Geochemistry and Geochronology of I-Type Granitoid Rocks in the Northeastern Central Iran Plate Abolfazl Soltani University of Wollongong
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2000 Geochemistry and geochronology of I-type granitoid rocks in the northeastern central Iran plate Abolfazl Soltani University of Wollongong
Recommended Citation Soltani, Abolfazl, Geochemistry and geochronology of I-type granitoid rocks in the northeastern central Iran plate, Doctor of Philosophy thesis, School of Geosciences, University of Wollongong, 2000. http://ro.uow.edu.au/theses/1970
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GEOCHEMISTRY AND GEOCHRONOLOGY OF I-TYPE GRANITOID ROCKS IN THE NORTHEASTERN CENTRAL IRAN PLATE
A thesis submitted in fulfilment of the requirements for the award of the degree of
DOCTOR OF PHILOSOPHY
from
UNIVERSITY OF WOLLONGONG,
by
ABOLFAZL SOLTANI BSc Mashhad, MSc (Hons) Wollongong
SCHOOL OF GEOSCIENCES 2000 I dedicate this thesis to my wife Zahra and my children Somayeh and Farzad This work has not been submitted for a higher degree at any other university or institution and, unless otherwise acknowledged, is the author's original research.
Abolfazl Soltani
31/03/2000 i
ABSTRACT
The Taknar and Sabzevar Zones in the northeastern Central Iran Plate (CIP) encompass a large variety of volcanic and plutonic rocks. The plutonic rocks are mainly I-type granitoids, ranging in age from Late Jurassic to Middle-Late Eocene. Among the plutonic rocks, granodiorite and granite are the most abundant rock types. This dissertation summarises the results of a detailed petrographic, geochemical and isotopic study of granitoid rocks from three areas of northeastern CEP, comprising the Kashmar and Bornavard granitoids in the
Taknar Zoiie, and the Kuh Mish intrusions in the Sabzevar Zone. The Kashmar and
Bornavard granitoids are generally high in Na20, total Fe as Fe203, Mn, Ba, Zr and Sr, and low in Ti02, P205, Rb, Nb, Cr, Ni and Sn contents. On Harker plots, they show regular trends for most major and trace element concentrations. They are characterised by steep negative slopes for LREE, flat to slightly negative gradients for HREE and moderate to strongly negative anomalies for Eu, features attributed to fractional crystallisation. However, in the Kashmar granitoid, restite separation and fractional crystallisation may be responsible for compositional variations. Mineralogical and chemical data suggest that the Kashmar and
Bornavard granitoids have formed from low temperature I-type magmas and can be assigned to a 'simple suite' of White et al. (2000). In the Kashmar granitoid, initial 87Sr/86Sr
(0.70471-0.70569) and 6Nd (-0.70 to -1.86) values are low and exhibit a restricted range, indicating a homogeneous lower crustal protolith. In the Bornavard granitoid, however,
87 86 initial Sr/ Sr (0.70757-0.75008) and eNd values (-1.41 to -5.20) exhibit a large range, suggesting that magmas were extensively contaminated with older continental crust or they were derived from partial melting of older felsic rocks of the continental crust. ii
The Kuh Mish intrusions are compositionally diverse, ranging from gabbro to quartz monzodiorite, but are dominated by granodiorite. They have low abundances of alkalis,
LFSE, HFSE and LREE relative to the Kashmar and Bornavard granitoids. They are also the most isotopically primitive plutonic rocks in northeastern CJP, typically having initial (at
87 86 42.8 Ma) Sr/ Sr of 0.70386-0.70475 and sNd values of +8.02 to +6.30 that indicate a mantle source. In these aspects, they are similar to the tonalitic association in the American
Cordillera and, in particular, to the western Peninsular Ranges Batholith.
The granitoid rocks of the northeastern CJJP show characteristics of magmas that originated in a subduction-related environment. The Rb/Sr ages of biotite-whole rock pairs from granitoids of northeastern CIP are consistent with the timing of subduction of the Neo-
Tethys Oceanic crust beneath the CIP. In particular, Sr-Nd isotopic data show that in the northeastern CIP, Middle-Late Eocene granitoids are isotopically less evolved or have primitive features compared with Late Jurassic/Early Cretaceous granitoids. It seems that voluminous injection of basaltic and andesitic magmas derived from subduction of the oceanic crust resulted in a complete change in the genesis of magmas in the northeastern
CIP. Using tectonic discrimination diagrams, the Kashmar and Bornavard granitoids typically plot in the 'volcanic arc and syn-collisional' granite field. However, the Kuh Mish intrusions are strongly depleted in Rb, Nb and Y contents suggesting that they may have emplaced in an island arc environment. iii
ACKNOWLEDGEMENTS
This thesis could not have completed without the help and assistance of many people who contributed their time, enthusiasm and support towards the intellectual and practical understanding of granitoids.
My supervisor Dr Paul F. Carr is especially thanked for giving me the opportunity to complete a doctoral study in one of the most scenic field sites in the world. He also provided valuable insight and fruitful discussion through all phases and aspects of the present study.
His patience, during my field and laboratory work and his critical reading of chapter drafts, is hereby acknowledged.
Special thanks are due to my family for their constant support during my studies, often under difficult and trying circumstances, and for their love and faith in me.
The various Heads of the Department of Geology/School of Geosciences (A/Prof. Tony J.
Wright, A/Prof. Brian G. Jones, Prof. Allan R. Chivas and A/Prof. Ted Bryant) are thanked for making available the facilities of the Department/School. I am indebted to A/Prof. Brian
G. Jones and Mr Aivars M. Depers for encouragement and reading of many drafts of my thesis chapters and bringing to my attention errors and omissions, all beyond the call of duty. Prof. Bruce W. Chappell from the Australian National University is sincerely thanked for giving me his advice and forwarding copies of his publications. I would like to thank Mr.
David Carrie for his help in the preparation of polished thin sections. Thanks are also due to
Dr Norm Pearson from Macquarie University for his assistance with the electron microprobe iv analyses. I wish to express my sincere appreciation to Dr Masoud Doroudian and Mr Ali A.
Shojaei for proof reading my thesis references and linking these with the text. Mr Ali A.
Faramarziah is thanked for help in final printing of the thesis.
The field aspects of the present study occupied a large proportion of my time and numerous people need to be thanked for their assistance, not only in active muscle work, but also camaraderie, often during inclement weather. I am indebted to Mr Mahmoud Refabi for his continued and enthusiastic help with all aspects of field sampling, as well as logistics. I am grateful to the Director of the Geological Survey of Mashhad, Mr Jafar Taheri for allowing me to use unpublished data from the Kashmar 1:100 000 Geological Sheet. He also provided assistance and informative discussions during field trips. The Governor's Office of Kashmar provided suitable vehicles for field trips. I am grateful to fellow postgraduate students of the
School of Geosciences, particularly Musa Arhoma, Jenny Atchison, Sue Murray, Mark
Dickson, Alex Golab, Daniel Palamara and Simon Clarke for providing amusement and social relief as well as different academic insights into research, both relevant and irrelevant to granitoids.
This work was supported by a scholarship from the Ministry of Culture and Higher
Education of the Islamic Republic of Iran. Without this assistance, this study would not have been possible. V
TABLE OF CONTENTS
ABSTRACT i ACKNOWLEDGEMENTS iii TABLE OF CONTENTS v LIST Of FIGURES x LIST OF TABLES xvi
CHAPTER 1 INTRODUCTION 1.1 INTRODUCTION 1 1.2 GENERAL GEOLOGY 2 1.3 AIMS 4 1.4 PREVIOUS WORK 4 1.5 IGNEOUS ROCK NOMENCLATURE 5 1.6 LAYOUT OF THE THESIS 5
CHAPTER 2 REGIONAL GEOLOGY OF IRANIAN GRANITOIDS 2.1 INTRODUCTION 7 2.2 CENTRAL IRAN PLATE (CIP) 8 2.3 SANANDAJ-SIRJAN MET AMORPHIC ZONE (S-SMZ) 10 2.4 URUMIEH-DOKHTAR VOLCANIC BELT (U-DVB) 12 2.5 GEOLOGICAL SETTING OF IRANIAN GRANITOIDS 12 2.5.1 REGIONAL-AUREOLE GRANITOIDS 12 2.5.1.1 Chapedony Complex 13 2.5.1.2 Doran Granite 13 2.5.2 CONTACT-AUREOLE GRANITOIDS 14 2.5.2.1 Mashhad Granite 14 2.5.2.2 Shahkuh Granite 16 2.5.2.3 Shirkuh Batholith 17 2.5.2.4 Muteh Granite 17 2.5.2.5 Hamadan Batholith 17 VI
2.5.3 SUB VOLCANIC GRANITOIDS 19 2.5.3.1 Natanz Intrusive Complex 19 2.5.3.2 Karkas and Jebal-e-Barez Intrusions 20 2.5.4 SUMMARY 20
CHAPTER 3 GEOLOGICAL SETTING AND GEOCHRONOLOGY 3.1 GEOLOGICAL SETTING 23 3.1.1 TAKNAR ZONE 23 3.1.2 SABZEVAR ZONE 24 3.2 MAJOR FAULT SYSTEMS 25 3.2.1 DORUNEH FAULT 25 3.2.2 RIVASH FAULT 26 3.3 GEOCHRONOLOGY 26 3.3.1 KASHMAR GRANITOID 26 3.3.1.1 Rb/Sr Age Dating 28 3.3.1.2 IsotopicData 28 3.3.2 BORNAVARD GRANITOID 31 3.3.2.1 IsotopicData 32 3.3.2.2 Age Discussion on the Bornavard Granitoid 35
CHAPTER 4 PETROGRAPHY AND MINERAL CHEMISTRY 4.1 PETROGRAPHY OF KASHMAR GRANITOID 38 4.1.1 TONALITE 38 4.1.2 GRANODIORITE 39 4.1.3 GRANITE 39 4.1.4 ALKALI FELDSPAR GRANITE 40 4.2 MINERAL CHEMISTRY OF KASHMAR GRANITOID 41 4.2.1 PLAGIOCLASE 41 4.2.2 AMPHD30LE 43 4.2.3 BIOTITE 47 4.2.4 Fe-Ti OXIDES 52 4.2.5 K-FELDSPAR 54 4.2.6 QUARTZ 56. Vll
4.2.7 ACCESSORY MINERALS 57 4.2.8 ALTERATION PRODUCTS 59 3 PETROGRAPHY OF BORNAVARD GRANITOID 60 4.3.1 TONALITE 60 4.3.2 GRANODIORITE 60 4.3.3 GRANITE 61 4 MINERAL CHEMISTRY OF BORNAVARD GRANITOID 62 4.4.1 PLAGIOCLASE 62 4.4.2 K-FELDSPAR 63 4.4.3 AMPHIBOLE 64 4.4.4 BIOTITE 67 4.4.5 ACCESSORY MINERALS 69 4.4.6 ALTERATION PRODUCTS 72 5 TAKNAR RHYOLTTE 73 4.5.1 PETROGRAPHY AND MINERAL CHEMISTRY 74 6 KUH MISH INTRUSIONS 76 4.6.1 GABBRO 77 4.6.1.1 Mineralogy of Gabbro 77 4.6.2 QUARTZ MONZODIORITE 78 4.6.2.1 Mineralogy of Quartz Monzodiorite 78 4.6.3 GRANODIORITE 79 4.6.3.1 Mineralogy of Granodiorite 80 7 SUMMARY 83
CHAPTER 5 WHOLE ROCK GEOCHEMISTRY 1 INTRODUCTION 85 2 KASHMAR GRANITOID 85 5.2.1 MAJOR ELEMENTS 85 5.2.2 SUMMARY OF MAJOR ELEMENTS 90 5.2.3 INCOMPATIBLE ELEMENTS 91 5.2.3.1 Low Field Strength Elements (LFSE) 92 5.2.3.2 High Field Strength Elements (HFSE) 95 viii
5.2.3.3 Rare Earth Elements (REE) 96 5.2.4 COMPATIBLE ELEMENTS 97 5.2.5 Sr AND Nd ISOTOPES 98 5.3 BORNAVARD GRANITOID 100 5.3.1 MAJOR ELEMENTS 101 5.3.2 INCOMPATIBLE ELEMENTS 101 5.3.2.1 Low Field Strength Elements (LFSE) 103 5.3.2.2 High Field Strength Elements (HFSE) 105 5.3.2.3 Rare Earth Elements (REE) 106 5.3.3 COMPATIBLE ELEMENTS 107 5.3.4 Sr AND Nd ISOTOPES 108 5.4 TAKNAR RHYOLITE 110 5.4.1 MAJOR AND TRACE ELEMENTS 110 5.4.2 Sr AND Nd ISOTOPES 112 5.5 KUH MISH INTRUSIONS 113 5.5.1 MAJOR ELEMENTS 113 5.5.2 INCOMPATIBLE ELEMENTS 114 5.5.3 RARE EARTH ELEMENTS (REE) 115 5.5.4 COMPATIBLE ELEMENTS 116 5.5.5 Sr AND Nd ISOTOPES 117
CHAPTER 6 GENETIC CLASSIFICATION AND COMPARISON WITH OTHER GRANITOIDS 6.1 INTRODUCTION 119 6.2 FIELD AND PETROGRAPHIC EVIDENCE 120 6.3 MPNERALOGICAL EVIDENCE 120 6.4 EVIDENCE FOR RESTTTE 122 6.5 CHEMICAL COMPOSITIONS 124 6.6 ALUMINUM SATURATION INDEX (AST) 126 6.7 Sr AND Nd ISOTOPES 132 6.8 ALLOCATION OF GRANITOIDS TO SUITE 133 6.9 HIGH- AND LOW-TEMPERATURE I-TYPE GRANITES 135 6.10 COMPARISON WITH OTHER GRANITOID TYPES 137 ix
6.10.1 COMPARISON WITH S-TYPE GRANITES 137 6.10.2 COMPARISON WITH A-TYPE GRANITES 138 6.10.3 COMPARISON WITH I-TYPE GRANITES 140
CHAPTER 7 PETROGENESIS AND TECTONIC SETTING 7.1 PETROGENESIS 143 7.1.1 PRODUCTION OF I-TYPE GRANITE SOURCE ROCKS 143 7.1.2 PRODUCTION OF I-TYPE GRANITES BY PARTIAL MELTING WITHIN THE CRUST 144 7.1.3 FRACTIONAL CRYSTALLISATION IN LOW-TEMPERATURE I-TYPE GRANITES 146 7.1.4 RESTHE FRACTIONATION 147 7.1.5 'MINIMUM-MELT COMPOSITIONS 150 7.1.6 Sr AND Nd ISOTOPES 151 7.1.7 LFSE ENRICHMENT AND HFSE DEPLETION 153 7.1.8 A MODEL FOR EVOLUTION OF MAGMAS IN NORTHEASTERN CIP 155 7.2 TECTONIC SETTINGS 158 7.2.1 ANOROGENIC GRANITES 159 7.2.2 OROGENIC GRANITES 160 7.2.2.1 Island Arc Granites 160 7.2.2.2 Magmatic Arcs of Continental Margines 161 CHAPTER 8 CONCLUSIONS 8.1 CONCLUSIONS 164
REFERENCES 174
APPENDIX 1 240 ANALYTICAL METHODS
APPENDIX 2 246 MODAL MINERALOGY AND C.I.P.W. NORMS
APPENDIX 3 252 ELECTRON MICROPROBE ANALYSES X
APPENDIX 4 290 WHOLE-ROCK GEOCHEMICAL DATA
LIST OF FIGURES
CHAPTER 1 Figure 1.1 Generalized tectonic map of Iran, based on the geological maps of Ruttner and Stocklin (1967) and Alavi (1991). 200
Figure 1.2 Location of the area studied. 201
CHAPTER 2 Figure 2.1 Tectonic subdivision of the Sanandaj-Sirjan Metamorphic Zone (after Mohajjel, 1997). 202
Figure 2.2 The Hamadan Batholith, a typical contact-aureole granitoid in the Sanandaj- Sirjan Metamorphic Zone, Iran (after Mohajjel, 1997). 203
Figure 2.3 Representative regional-aureole granites of Iran. 204
Figure 2.4 Representative contact-aureole granites of Iran. 205
Figure 2.5 Representative subvolcanic granites of Iran. 206
CHAPTER 3 Figure 3.1 Generalized geological map of east of the Taknar Zone, northeastern Central Iran Plate (after Valipour, 1992). 207
Figure 3.2 Geological map of the Kashmar area. 208 xi
Figure 3.3 Rb/Sr whole rock isochron diagram for granodiorite, granite and alkali feldspar granite from the Kashmar granitoid. MSWD = Mean Squares of Weighted Deviates, used as a measure of the goodness offit o f the isochron. 209
Figure 3.4 Geological map of the Bornavard area. 210
CHAPTER 4 Figure 4.1 Modal compositions (quartz, K-feldspar, plagioclase) of the Kashmar granitoid. Fields are based on classification of igneous rocks, proposed by Streckeisen (1976) and Le Bas and Streckeisen, 1991). 211
Figure 4.2 Plagioclase composition from the Kashmar granitoid. 211
Figure 4.3 Diagram of Ti versus total Al for magnesio hornblende from the Kashmar granitoid, with pressure contours determined according to Johnson and Rutherford (1989a). 212 Figure 4.4 Composition of biotite crystals from the Kashmar granitoid. The boundary between phlogopite and annite is proposed at Mg/(Mg + Fe) = 0.7 (Gribble, 1988). 212 Figure 4.5 Plot of MgO versus AI2O3 from biotite of the Kashmar granitoid. Discriminant lines are after Abdel-Rahman (1994). A = Anorogenic, P = Peraluminous and C = Calcalkaline orogenic. 213
Figure 4.6 Relationship between Mg/(Mg + Fe) and whole rock Si02 contents for biotite from the Kashmar granitoid. 213
Figure 4.7 Diagram showing negative correlation between Mg/(Mg + Fe) and total Fe (a.f.u.) from biotite of the Kashmar granitoid. 214
Figure 4.8 Diagram showing negative correlation between Mg/(Mg + Fe) and Ti (a.f.u.) in biotite from the Kashmar granitoid. 214 xii
Figure 4.9 Comparison of Mg/(Mg + Fe) between coexisting hornblende and biotite from the Kashmar granitoid. 215
Figure 4.10 Composition of K-feldspar crystals from the Kashmar granitoid. 215
Figure 4.11 Modal compositions (quartz, K-feldspar, plagioclase) of the Bornavard granitoid. Fields are based on classification of igneous rocks, proposed by the IUGS Subcommission on the systematics of igneous rocks (Streckeisen, 1976; Le Bas and Streckeisen, 1991). 216
Figure 4.12 Anorthite-Albite-Orthoclase triangle plot for plagioclase composition from the Bornavard granitoid. 216
Figure 4.13 Anorthite-Albite-Orthoclase triangle diagram showing the composition of K- feldspar crystals in granite from the Bornavard granitoid. 217
Figure 4.14 Diagram of Ti (a.f.u.) versus total Al (a.f,u.) for magnesio hornblende from the Bornavard granitoid, with pressure contours determined according to Johnson and Rutherford (1989a). 217
Figure 4.15 Composition of biotite crystals from the Bornavard granitoid. 218
Figure 4.16 Negative correlation between Mg/(Mg + Fe) and total Fe (a.f.u.) for biotite crystals from the Bornavard granitoid. 218
Figure 4.17 Total alkali contents versus Si02 (wt%) classification (TAS) for the Taknar Rhyolite (fields after Le Maitre, 1989 and Le Bas and Streckeisen, 1991). 218
Figure 4.18 Simplified geological map of the Kuh Mish area. 219
Figure 4.19 Modal compositions (quartz, K-feldspar, plagioclase) of the Kuh Mish intrusions. Fields are based on classification of igneous rocks, proposed by the IUGS xiii
Subcomrnission on the systematics of igneous rocks (Streckeisen, 1976; Le Bas and Streckeisen, 1991). 220
Figure 4.20 Composition of clinopyroxene in gabbro from the Kuh Mish intrusions. 220 Figure 4.21 Composition of plagioclase crystals in gabbro and granodiorite from the Kuh Mish intrusions. 220
CHAPTER 5 Figure 5.1 Harker diagrams for the Kashmar granitoid (oxides, wt%). Symbols: Tonalite (+), Granodiorite (x), Granite (*) and Alkali feldspar granite (a). 221
Figure 5.2 Multi-element patterns (spider diagrams) for the Kashmar granitoid. The normalised values are from Mc Donough et al. (1991). Symbols as Figure 5.1. 222
Figure 5.3 Harker diagrams for trace elements abundances of the Kashmar granitoid (oxides, wt% and traces, ppm). Symbols as Figure 5.1. 223
Figure 5.4 Rare earth element patterns for granodiorite, granite and alkali feldspar granite from the Kashmar granitoid. The normalised values are from Taylor and Mc Lennan (1985). Symbols as Figure 5.1. 224
Figure 5.5 Harker diagrams for the Bornavard granitoid (oxides, wt%). Symbols: Tonalite (+), Granodiorite (x) and Granite (*). 224
Figure 5.6 Multi-element patterns (spider diagrams) for the Bornavard granitoid. The normalised values are from Mc Donough et al. (1991). Symbols as Figure 5.5. 225
Figure 5.7 Harker diagrams for trace elements abundances of the Bornavard granitoid (oxides, wt% and traces, ppm). Symbols as Figure 5.5. 226 xiv
Figure 5.8 Rare earth element patterns for tonalite, granodiorite and granite from the Bornavard granitoid. The normalised values are from Taylor and Mc Lennan (1985). Symbols as Figure 5.5. 227
Figure 5.9 Plot of initial 87Sr/86Sr versus Si02 (wt%) for igneous rocks of the northeastern CIP. Symbols: Kashmar (+), Bornavard (x), Taknar (*) and Kuh Mish (•). 227
Figure 5.10 Multi-element patterns (spider diagrams) for the Taknar Rhyolite, northeastern CIP. The normalised values are from Mc Donough et al. (1991). 228
Figure 5.11 Rare earth element pattern for the Taknar Rhyolite, northeastern CIP. The normalised values are from Taylor and Mc Lennan (1985). 228
Figure 5.12 Harker diagrams for Gabbro (+), quartz monzodiorite (x) and granodiorite (*) from the Kuh Mish area, northeastern CIP (oxides, wt% and traces ppm). 229
Figure 5.13 Multi-element patterns (spider diagrams) for gabbro (+), quartz monzodiorite (x) and granodiorite (*) from the Kuh Mish area, northeastern CIP. The normalised values are from Mc Donough et al. (1991). 230
Figure 5.14 Rare earth element patterns for gabbro (+) and granodiorite (*) from the Kuh Mish area, northeastern CIP. The normalised values are from Taylor and Mc Lennan (1985). 230
CHAPTER 6
Figure 6.1 Plot of major elements against Si02 contents for plutonic rocks of the northeastern CIP. (oxides, wt%). Symbols: Kashmar granitoid (+), Bornavard granitoid (x) and Kuh Mish intrusions (D). 231
Figure 6.2 Plot of trace elements abundances against Si02 contents for plutonic rocks of the northeastern CIP. (oxides wt%, traces, ppm, symbols as Figure 6.1). 231 XV
Figure 6.3 (a) Aluminum Saturation Index (AST) and (b) Si02 contents versus ASI values for igneous rocks of the northeastern CIP. The boundary between metaluminous and peraluminous at ASI =1.1 proposed by Chappell and White (1974) because they recognised more generally that very felsic I-type granites may be weakly peraluminous (Chappell, 1998b). Symbols: Kashmar granitoid (+), Bornavard granitoid (x), Kuh Mish intrusions (•) and Taknar Rhyolite (*). 232
Figure 6.4 Ternary plot of normative Q-Ab-Or for granite from the Bornavard granitoid. The curves for water-saturated liquids in equilibrium with quartz and K-feldspar at 0.5 and 3.0 kb are from Tuttle and Bowen (1958). The position of 'minimum-melt' compositions of Tuttle and Bowen (1958) are shown by a cross (+) on each curve. 232
Figure 6.5 Histograms of ASI frequency for igneous rocks of the northeastern CJP. Numbers in the right side of the histograms indicate 1 = Kashmar granitoid, 2 = Bornavard granitoid, 3 = Taknar Rhyolite and 4 = Kuh Mish intrusions. 233
CHAPTER 7 Figure 7.1 ACF diagram for plutonic rocks of the northeastern CIP. Plagioclase-(FeO + MgO) bine defines ASI (aluminum saturation index) = 1, which divides peraluminous and metaluminous granitoids (Chappell and White, 1992). 233
Figure 7.2 Harker diagram for total Fe as Fe203 in the Kashmar granitoid. The diagram represents the partial melting of the source rock (S) to produce restite (R) and a liquid as 'minimum-melt' (M) composition (Chappell et al., 1987). The magma at its source consists of (R+M) and varying degrees of separation of R from M generated a range of magma and rock compositions, illustrated by granodiorite (x), granite (*) and alkali feldspar granite (D). 234
87 86 Figure 7.3 Diagram showing initial Sr/ Sr versus eNd values for Kashmar granitoid (+), Bornavard granitoid (x) and Kuh Mish intrusions (•). The boundary between mantle and crust (sNd = 0) is from Rollinson (1993). 234 XVI
Figure 7.4 The Nb-Y discrimination diagram for granitoid rocks of the northeastern CIP. The fields (Pearce et al, 1984; Forster etal, 1997) show volcanic-arc granites (VAG), syn- collisional granites (Syn-COLG), within-plate granites (WPG) and ocean-ridge granites (ORG). The broken line is thefield boundary for ORG from anomalous ridges. Symbols: Kashmar granitoid (+), Bornavard granitoid (x) and Kuh Mish intrusions (D). 235
Figure 7.5 The Rb versus (Y + Nb) discrimination diagram for granitoid rocks of the northeastern CIP. The fields and symbols are as Figure 7.4. 235
LIST OF TABLES
Table 2.1 Isotopic age data for some Iranian granitoids and volcanic rocks. 236
Table 3.1 Rb/Sr isotopic age data for Kashmar and Bornavard granitoids, northeastern CIP, Iran. 237
Table 5.1 Sr and Nd isotopic data for Kashmar granitoid, Bornavard granitoid, Taknar Rhyolite and Kuh Mish intrusions. 238
Table 6.1 Compilation of mean whole rock major and trace element data for I-, S- and A- type granitoids. 239 1
CHAPTER 1
INTRODUCTION
1.1 INTRODUCTION
The concept that the chemical characteristics of many igneous rocks reflect the composition of their source regions is widely accepted. This concept has been used in many studies of the petrogenesis of volcanic and plutonic rocks of diverse compositions and sources. Interpretation of the chemical variations within granitoid suites and their relationship to petrogenesis and source-rock compositions is still controversial. For example, two contrasted chemical types of granitoids (I- and S-type) from the Lachlan
Fold Belt (LFB) of eastern Australia imply melting of infra- and supra-crustal protoliths, respectively and chemical variation within related granites is considered to be controlled by processes such as fractional crystallisation and restite unmixing (White and Chappell,
1977; Chappell et al, 1987; Chappell, 1996a,b; Chappell, 1998a,b). hi contrast, Collins
(1996, 1998) and Keay et al. (1997) suggested that I- and S-type granites of the LFB are products of large scale three-component (Ordovician turbidites, Cambrian greenstones and depleted mantle) mixing, rather than unique products from different sources. But some low-temperature granites show compositional, petrographic, zircon age inheritance, and other features, which cannot be accounted for satisfactorily by the classical models of petrogenesis. The restite model is account for these features and recognises that unmelted but magmatically equilibrated source material (restite) may be entrained in a partial melt, together comprising magma (Chappell et al, 2000). In the broader context, the source and tectonic setting are intimately inter-related in the generative processes, so it is likely that the composition of granites characterizes their tectonic environment (Pitcher, 1993). 2
Plutonic rocks of diverse compositions crop out in many parts of Iran (Berberian, 1981;
Haghipour and Aghanabati, 1989) particularly in the Central Iran Plate (CIP) and the
Sanandaj-Sirjan Metamorphic Zone (S-SMZ). m Iran (Fig. 1.1), the emplacement of granitoid rocks occurred mainly during the Mesozoic (Jurassic and Cretaceous) and
Tertiary (Oligo-Miocene), but a few intrusions in the CIP have previously been assigned a
Precambrian age (Berberian, 1981; Emami et al, 1993). According to new isotopic data
(Section 3.3.2.1), a Precambrian age for some Iranian granitoids is unlikely. The results of preliminary investigations of Iranian granitoids are scattered among various publications, but no comprehensive study of any of these rocks has been undertaken previously.
1.2 GENERAL GEOLOGY
The present study is concerned with granitoid occurrences in the northeastern part of the
CIP between the cities of Kashmar, Bardaskan, Sabzevar and Neyshabur (Fig. 1.2). The granitoids occur in two geological zones, the Taknar Zone between the Doruneh and
Rivash faults (Lindenberg and Jacobshagen, 1983), and the Sabzevar Zone (Pilger, 1971) that occurs on the northern side of Rivash Fault. These granitoids are further subdivided into the Kashmar and Bornavard granitoids of the Taknar Zone, and the Kuh Mish intrusions of the Sabzevar Zone.
The Kashmar granitoid is the largest of these three granitoids and occurs on the northern side of the east-west trending Doruneh Fault, north of the city of Kashmar (Fig. 1.2). This granitoid is approximately 50 km long and 7 km wide, and is intruded into genetically related andesitic lava and pyroclastic rocks of Eocene age (Bernhardt, 1983). Several plutons comprising mainly tonalite, granodiorite, granite and alkali feldspar granite have 3
been recognised in this granitoid mass (Eftekhar-Nezhad, 1976; Behroozi, 1987). Before the present study, no isotopic or geochemical data have been reported for the Kashmar granitoid.
The Bornavard granitoid occurs -20 km northwest of the city of Bardaskan (Fig. 1.2). The country rocks in this region consist of metavolcanic units comprising mainly rhyolitic lavas and tuffs known as the Taknar Rhyolite (Stocklin and Setudehnia, 1991), which is considered to be Precambrian in age and the oldest unit in the Taknar Zone
(Razzaghmanesh, 1968). In the Bornavard area, the Taknar Rhyolite has been intruded by two major intrusive phases comprising granite and granodiorite. Before the current study, the granite and granodiorite of this region were considered to be representatives of
Precambrian and Tertiary granites of Iran, respectively (e.g., Forster, 1968;
Razzaghmanesh, 1968; Eftekhar-Nezhad, 1976). The current study presents new isotopic ages for rocks of the Bornavard-granitoid (Section 3.3.2.1).
The Kuh Mish intrusions crop out around the prominent mountain of Kuh Mish in the western part of the Sabzevar Zone (Fig. 1.2). The Kuh Mish intrusions include a large variety of rock types, ranging from gabbro through quartz monzodiorite to granodiorite.
Granodiorite occurs in three localities including Kuh Mish, Darin and Namin. The largest granodiorite pluton occurs in the Kuh Mish locality. This pluton intrudes volcano- sedimentary rocks of Eocene age, and a large body of quartz monzodiorite intrudes the granodiorite pluton. The northern part of the granodiorite pluton is cut by several parallel dykes which are quartz monzodiorite in composition. Gabbro, the last intrusive phase, occurs only in the central part of the quartz monzodiorite. In the Darin region, the granodiorite intrudes volcano-sedimentary rocks of Eocene age, whereas in the Namin 4
region, the contact between granodiorite and country rocks is covered by Quaternary deposits. Based on these stratigraphic relationships, the Kuh Mish intrusions are known to be Middle-Late Eocene in age (Eftekhar-Nezhad, 1976; Lindenberg et al, 1983).
1.3 AIMS
The major part of this dissertation is concerned with a detailed petrographic, geochemical and isotopic investigation of granitoids that occur in the northeastern part of the CIP.
Development of a model for the petrogenesis of these granitoids requires the following:
• isotopic age dating of appropriate granitoids;
• documentation and interpretation of the petrography, mineral chemistry and total rock geochemistry of the granitoids;
• documentation and interpretation of the isotopic compositions of various rock types;
• recognition and classification of the granitoid types on the basis of source regions;
• determination of the relationship between the granitoids and associated igneous rocks;
• comparison of granitoids on a local and global scale; and
• determining the tectonic environment in which granitoids have been emplaced.
1.4 PREVIOUS WORK
Previous knowledge of much of the area was summarised on the first comprehensive
1:250 000 geological map of Kashmar compiled by Eftekhar-Nezhad (1976). This map served as a valuable base map for the western parts of the study area, while the 1:100 000 geological map of Feyz-Abad compiled by Behroozi (1987) was used as a base map for the eastern part of the area.
Previous investigations have concentrated on Cu, Fe and Zn deposits in the Taknar Zone 5
(Razzaghmanesh, 1968; Muller and Walter, 1983), together with studies of the stratigraphy and structural evolution of the Sabzevar Zone (Lensch et al, 1980;
Lindenberg et al, 1983; Lindenberg and Jacobshagen, 1983). Publications directly related to the granitoids are scarce, but brief descriptions of petrography have been presented by
Homam (1992), Sepahi (1992) and Valipour (1992).
1.5 IGNEOUS ROCK NOMENCLATURE
Petrographic names are based on modal data utilizing the IUGS classification
(Streckeisen, 1976; Le Bas and Streckeisen, 1991) for plutonic rocks; and the TAS classification (total alkalis vs silica) as proposed by Le Maitre (1989) and Le Bas and
Streckeisen (1991) for volcanic rocks. The term 'granitoid' is used in the general sense for plutonic rocks ranging in composition from tonalite to alkali feldspar granite with quartz contents between 20 and 60% by volume of the rock.
L6 LAYOUT OF THE THESIS
This thesis has been organised into eight chapters. In Chapter 2 the classification and regional geology of the major Iranian granitoids occurring in the CIP, S-SMZ and
Urumiyeh-Dokhtar Volcanic Belt (U-DVB) are presented. In Chapter 3 the geological setting of the Taknar and Sabzevar Zones, in which rocks of the present study have been emplaced, and their geochronology are discussed. Chapter 4 is devoted to petrography and mineral chemistry of the granitoids and the Taknar Rhyolite Chapter 5 contains whole rock geochemistry, while Chapter 6 discusses the genetic classification (e.g., I- and S-type). It also compares the rocks studied to other granitoids on a local and global scale, based on several criteria, including chemical and mineralogical properties, aluminium saturation index (ASI) and Sr-Nd isotopic compositions. Using a genetic 6
classification, the rocks studied from the Taknar Zone are allocated into a 'simple suite' while those of the Sabzevar Zone are assigned into a different suite. Chapter 7 is concerned with petrogenesis and tectonic setting of the granitoids, whereas Chapter 8 presents the major conclusions from the present study. 7
CHAPTER 2
REGIONAL GEOLOGY OF IRANIAN GRANITOIDS
2.1 INTRODUCTION
Iranian granitoids occur mainly in the CIP and the S-SMZ. Stocklin (1972) originally divided these granitoids into five groups, but later, Stocklin and Nabavi (1973) classified them into three groups comprising Precambrian, Mesozoic and Tertiary granitoids. Haghipour and Aghanabati (1989) also adopted this three-fold subdivision on the geological map of Iran, but Berberian (1981) recognised eight plutonic episodes.
The three-fold subdivision of Iranian granitoids was based on similarities in petrography or stratigraphic relationships (Stocklin and Setudehnia, 1991). Recent geological investigations and limited isotopic data (Tables 2.1 and 3.1) suggest that the
Precambrian ages reported for some Iranian granitoids are unlikely (Otroudi, 1987;
Sepahi, 1992; Emami etal, 1993; Noorbehesht and Sharifi, 1997).
Relationships between granitoids and the associated country rocks, in the Lachlan Fold
Belt of eastern Australia, have provided the basis for subdivision into regional-aureole, contact-aureole and subvolcanic types (White et al, 1974) that has significant implications for depth of emplacement and origin of these granitoids (White and
Chappell, 1988; Chappell and White, 1992; Collins, 1996). Iranian granitoids are associated with either Precambrian and Mesozoic metamorphic rocks (Darvichzadeh,
1992), or with Tertiary volcanic rocks, especially in the CIP and its western and southwestern limits known as the Urumiyeh-Dokhtar Volcanic Belt (U-DVB; Kazmin et 8
al, 1986b; Darvichzadeh, 1992; Moradian, 1997). Using the same approach as adopted by White et al. (1974) for subdivision of the granitoids in the Lachlan Fold
Belt, this chapter reviews the available data from geological maps and other publications, together with new observations to subdivide Iranian granitoids into regional-aureole, contact-aureole and subvolcanic types. As plutonic activity in Iran is restricted to the CIP, S-SMZ and U-DVB, a brief review of geology of these three major geological zones is also presented.
2.2 CENTRAL IRAN PLATE (CIP)
The term Central Iran Plate or CIP is applied to an approximately triangular shaped area limited by the active marginal basins of the East Iran Belt to the east, the Alborz Belt to the north and the S-SMZ to the southwest (Fig. 1.1; Stocklin, 1968). During the
Precambrian and Palaeozoic, the CIP was part of the Arabian Plate and was separated from the Eurasia Plate by the Hercynian Ocean (Berberian and King, 1981). The CIP was a stable platform during the Palaeozoic, but tectonic activity in the Late Triassic produced a series of horsts and grabens between major faults (Stocklin, 1968; Hamedi,
1995).
Late Palaeozoic rifting in the Arabia-Iran platform, along the present line of the Main
Zagros Thrust Line, initiated separation and northward movement of the CIP and opened the Neo-Tethys Ocean in the south. The Neo-Tethys Ocean in this region started to close in the southwest towards the end of the Cretaceous and the northern parts of this ocean were almost entirely closed by the Eocene. Subduction of the Neo-Tethys Oceanic crust beneath the southern margins of the CJP produced an Andean-type magmatic arc during the Mesozoic and possibly Early Tertiary (Berberian, 1981). 9
Paleomagnetic data indicate that, during the Jurassic to Late Cretaceous, the CJP moved northwards accompanied by a counterclockwise rotation of about 100° towards the southern rim of Eurasia (Davoudzadeh et al, 1981; Schmidt and Soffel, 1983).
During the Late Cretaceous, the CJP converged with the Turan Plate. Additional northwards movement during the Paleocene-Eocene resulted in the collision of the CIP with the southern rim of Eurasia (Turan Plate) followed by the emplacement of large volumes of felsic magmas into the CIP (Soffel and Forster, 1983). In the Late Eocene-
Oligocene a limited tensional episode produced long fractures in the CIP, followed by volcanic activity (mainly andesite) and emplacement of subvolcanic granitoids along these fractures (Dercourt et al, 1986).
An Early Cimmerian orogenic event is evident in many parts of Iran, as indicated by changes in depositional environment from shallow continental to open marine during the Late Triassic (Kazmin et al, 1986a; Darvichzadeh, 1992). The Late Cimmerian orogenic event is responsible for subduction of the Neo-Tethys Oceanic crust beneath the active continental margins of the CIP. The latter event was accompanied by complex deformation in the S-SMZ (Mohajjel, 1997, 1998), folding, igneous activity and metamorphism in the CIP, and the development of unconformities in the Zagros, Alborz and Kopeh Dagh Fold Belts (Aghanabati, 1993). Granitic rocks, related to the Late
Cimmerian orogenic event, were emplaced in the CIP during the Late Jurassic to Early
Cretaceous (Aghanabati, 1993).
According to Berberian and Berberian (1981), after the Late Cretaceous orogenic activity, very large volumes of dacitic, andesitic and basaltic lavas, with tuffaceous and 10
other clastic sediments, were formed during the Eocene in the CIP and Alborz Fold
Belt. During the Late Eocene-Oligocene, Oligo-Miocene and Pliocene epochs, these rocks were cut by several intrusive bodies that are mainly coarse- to medium-grained biotite granite, hornblende biotite granodiorite, monzonite and diorite. The Tertiary plutonism in the CIP was not restricted primarily to plate margins and the bulk of the magmatism appears to have occurred within continental margins (Berberian and
Berberian, 1981).
2.3 SANANDAJ-SIRJAN METAMORPHIC ZONE (S-SMZ)
The Zagros Orogen (Alavi, 1994) developed from continental separation and subsequent collision between the Arabian platform and the CJP. It is part of the Tethyan orogenic collage that developed between Eurasia and dispersed fragments of Gondwana (Sengor,
1984). The Zagros Orogen consists of the U-DVB, S-SMZ and the Zagros Fold-Thrust
Belt (Alavi, 1991, 1994; Mohajjel, 1997).
The S-SMZ is located southwest of the U-DVB and is characterized by metamorphic and complexly deformed rocks, associated with abundant deformed and undeformed plutons, in addition to widespread Mesozoic volcanic rocks. It has a length of 1500 km, from the northwest to southeast Iran (Fig. 1.1), and a width up to 200 km. The S-SMZ is considered to be the most tectonically active zone in Iran. The rocks in this zone are mostly of Mesozoic age. Palaeozoic rocks are rarely exposed in the northwestern part of the S-SMZ, whereas they commonly occur in the southeastern part (Berberian, 1995;
Sabzehei and Eshraghi, 1995). In the southeastern part of the S-SMZ, deformation and metamorphism have been attributed to a number of orogenic episodes by different authors (e.g., Sengor, 1990). The major deformation and metamorphic events that 11
affected the S-SMZ are associated with the opening and closing of the Neo-Tethys
Ocean during the Mesozoic (Alavi, 1994). Based on the structural zonation recognized in Palaeozoic and Mesozoic rock assemblages, Mohajjel (1997) recently subdivided the
S-SMZ into several elongate sub-zones (Fig. 2.1). From southwest to the northeast these sub-zones are: (1) radiolarite sub-zone; (2) Bistoon sub-zone; (3) ophiolite sub-zone; (4) marginal sub-zone; and (5) complexly deformed sub-zone. Plutonic rocks of the S-SMZ are mainly granite and all occur in the complexly deformed sub-zone.
One of the distinctive features of the S-SMZ is that it contains Mesozoic-Cenozoic plutonic rocks that do not occur in the Zagros Fold-Thrust Belt. In contrast, most plutonic rocks in the CIP are of Cainozoic age.
The plutonic rocks of the S-SMZ are divided into two groups comprising plutons of
Late Jurassic age and plutons of Late Cretaceous-Palaeocene age. Late Jurassic plutonic rocks are less abundant than the younger plutonic rocks. In the southeastern complexly deformed sub-zone, plutonic rocks of Triassic age have been reported (Davoudzadeh and Weber-Diefenbach, 1987), but more work is required to establish the true age of these rocks. Most of the Late Cretaceous-Palaeocene plutonic rocks occur in the northwestern complexly deformed sub-zone and range in composition from gabbro to granite. The main granitic plutons are at Hamadan, Borujerd, Astaneh, Aligudarz, Boin-
Miandasht and Hasan-Robat (Figs 2.1 and 2.2). All these plutons have elliptical outcrop patterns that are elongate in a northwest-southeast direction. These plutons are known as
Hamadan Batholith. Isotopic ages of the Hamadan Batholith will be discussed in detail in Section 2.5.2.5. 12
2.4 URUMTYEH-DOKHTAR VOLCANIC BELT (U-DVB)
The U-DVB consists of a distinctive, thick sequence (up to 4 km) of volcanic and subvolcanic rocks of Eocene-Quaternary age that extends along the entire northern section of the S-SMZ as a 50 km wide zone (Berberian and Berberian, 1981). Plutonic outcrops within the U-DVB comprise a wide variety of lithologies including granite, granodiorite, diorite and gabbro, as well as widely distributed basaltic, trachybasaltic
(locally shoshonitic), andesitic, dacitic, trachytic lavas and pyroclastic units. Recently, based on geochemical and isotopic evidence, Moradian (1997) subdivided the U-DVB into three parts: (1) comprising: the "Urumiyeh-Nain" part in the northwest, (2) the
"Nain-Baft" part in the centre, and (3) the "Baft-Dokhtar" part in the southeast. The southeastern part of the U-DVB is still active and is associated with the ongoing subduction of Indian Ocean crust (White and Rose, 1979; McCall and Kidd, 1982;
McCall, 1985).
The genesis of the U-DVB has been controversial with several major tectono-magmatic models being proposed. The most popular model, however, involves Andean-type subduction of the Neo-Tethyan oceanic crust beneath the CIP during the Tertiary
(Berberian et al, 1982; Alavi, 1994; Moradian, 1997). In the U-DVB, the peak magmatic activity occurred during the Eocene (Alavi, 1994).
2.5 GEOLOGICAL SETTING OF IRANIAN GRANITOIDS
2.5.1 REGIONAL-AUREOLE GRANITOIDS
Regional-aureole granitoids of Iran intrude mainly Late Precambrian rocks that were folded, metamorphosed and uplifted during the Late Precambrian Katangan Orogeny
(Stocklin, 1968). The major representatives of this group occur in the CJP and are 13
known as the Chapedony Complex and Doran Granite (Fig. 2.3).
2.5.1.1 Chapedony Complex
The Chapedony Complex comprises the oldest of the Precambrian metamorphic rocks of Iran (Darvichzadeh, 1992) and occurs in the eastern part of the CJP (Fig. 2.3). It consists of ribboned gneiss, schist, migmatite, granite, granodiorite, quartz diorite, amphibolite and marble (Stocklin and Setudehnia, 1991). The granite and granodiorite consist mainly of deformed alkali feldspar, plagioclase, quartz and biotite, and small amounts of clinopyroxene and hornblende (Berberian, 1981). Gradational contacts between granite and ribboned gneiss indicate that emplacement was synchronous with metamorphism. The Rb/Sr isochrons formed by two suites of whole rock samples from the Chapedony Complex (Table 2.1) suggest ages of 541 and 550 Ma (Haghipour,
1978), but these dates may reflect a younger, high grade metamorphic imprint
(Berberian and Berberian, 1981).
2.5.1.2 Doran Granite
The Doran Granite (Fig. 2.3) is an equigranular to slightly porphyritic, white to pinkish coloured alkaline granite, with high contents of K-feldspar (mainly perthitic microcline) and quartz, but with low contents of plagioclase, muscovite, biotite, titanite, apatite, zircon and Fe-Ti oxides (Stocklin and Eftekhar-Nezhad, 1969). The granite has a slightly gneissic texture, intrudes regional metamorphic rocks (Kahar Phyllite) of
Precambrian age and is reportedly unconformably overlain by Late Precambrian dolomite (Stocklin et al, 1964). It contains abundant xenoliths derived from the Kahar
Phyllite. Although cited as a typical Precambrian granite (Stocklin et al, 1964), Rb/Sr dating of a biotite-whole rock pair (Table 2.1) gave an age of 175±5Ma (Crawford, 14
1977), which suggests that the Precambrian age is suspect.
Another exposure of regional-aureole granitoids is found in the Moghanlu area (Alavi et aL, 1982), approximately 30 km to the west of the Doran Granite (Fig. 2.3). In the
Moghanlu area, a granitic pluton intrudes green tuffaceous slates and encircles a group of high grade metamorphic rocks that are mainly augen gneisses, amphibolite and biotite schist (Stocklin and Eftekhar-Nezhad, 1969; Valizadeh and Esmaeili, 1994). The granitic pluton is enriched in biotite towards the contact with the high grade metamorphic rocks. Although no field evidence for gradual changes between granite and gneisses has been observed, geochemical data (Valizadeh and Esmaeili, 1994) indicate that the granite probably derived from partial melting of the gneiss that may have occurred at depth.
2.5.2 CONTACT-AUREOLE GRANITOIDS
In the S-SMZ, plutons have been emplaced both during and after tectonic activity.
Deformed syn-tectonic plutons are characterised by narrow contact aureoles, but post- tectonic plutons have wide aureoles or faulted margins (Soheili et al, 1992; Mohajjel,
1997). The major contact-aureole granitoids of Iran are the Muteh Granite and Hamadan
Batholith in the S-SMZ, together with the Mashhad, Shahkuh and Shirkuh Granites in the CJP (Fig. 2.4).
2.5.2.1 Mashhad Granite
The Mashhad Granite occurs in the northeastern part of Iran (Fig. 2.4) and it is surrounded by a remarkable contact-aureole up to 200 m wide (Esmaeili et al, 1998b), superimposed on regional metamorphic rocks which are of Late Permian to Early 15
Jurassic age (Aghanabati, 1986). The regional metamorphic rocks are pelitic, psammitic, calcareous and mafic in nature, and have been affected by low- to medium- grade metamorphism. The polymetamorphic rocks in the aureole around the Mashhad
Granite comprise almandine- and andalusite-rich pelite, garnet and biotite schist, quartzite, calc-silicate rocks and amphibolite, whereas the regionally metamorphosed rocks comprise almandine-, staurolite- and chloritoid-bearing pelite, andalusite-bearing pelite, quartzite, carbonate rocks, amphibolite and serpentinite (Holzer and
Momenzadeh, 1969; Majidi and Alavi, 1972; Aghanabati, 1986; Iranmanesh and
Sethna, 1998).
The Mashhad Granite was emplaced in three major intrusive phases (Majidi and Alavi,
1972). The first phase is characterised by porphyritic granite, biotite granite, granodiorite and hornblende-biotite tonalite, and occurs in the southern part of the
Mashhad Granite (Holzer and Momenzadeh, 1969; Majidi, 1978). In the first phase biotite and hornblende occur in some parts of the granite and an overprinted schistosity continues into the country rocks. The granite also contains numerous lens-shaped xenoliths of the same composition as the country rocks. These features were used to infer a syn-tectonic intrusion for the Mashhad Granite (Majidi and Alavi, 1972). The second phase produced leucogranite that consists mainly of quartz, K-feldspar, muscovite and rare biotite, together with accessory garnet, tourmaline, apatite, zircon, rutile and Fe-Ti oxides. The leucogranite intrudes the first phase and crops out in the central exposures of the Mashhad Granite. The third phase, represented by several aplitic, pegmatitic and pneumatolytic veins, intrudes both the earlier phases (Majidi and
Alavi, 1971) and contains quartz, K-feldspar, muscovite and minor amounts of biotite.
Chemical analyses of different intrusive phases of the Mashhad Granite indicate 16
metaluminous I-type features for hornblende-bearing granite that is considered as the first phase, and peraluminous S-type features for biotite/muscovite granite and pegmatite that are referred to the second and third phases (Iranmanesh and Sethna,
1998).
The Mashhad Granite is overlain by Cretaceous conglomerate, arkose and limestone about 40 km southeast of Mashhad. Holzer and Momenzadeh (1969), however, reported that granite intruded Early Cretaceous rocks to the south of Mashhad city. The K-Ar dates of 146-120 Ma for the Mashhad Granite (Alberti et al, 1973) and the stratigraphic relationships indicate high level multiple intrusions of granitic magma from Late
Jurassic to Early Cretaceous. A K-Ar age of 146+3 Ma (Table 2.1) for quartz diorite at
Mashhad (Alberti et al, 1973) indicates that the more mafic and peripheral phases of the
Mashhad Granite were probably the earliest intrusions in the composite granitic mass
(Iranmanesh and Sethna, 1998).
2.5.2.2 Shahkuh Granite
The Shahkuh Granite (590 km2) occurs in eastern Iran (Fig. 2.4) and intrudes Early
Jurassic shale. It is overlain by Orbitolina limestone of Late Jurassic-Early Cretaceous age (Darvichzadeh, 1992). The northern part of the Shahkuh Granite produced slight metamorphism, but the southern part is surrounded by a well-developed contact-aureole, containing andalusite, cordierite and biotite hornfels, and occasional copper mineralisation (Mobasher, 1992). This granite is petrographically similar to other
Middle-Late Jurassic granites (Aghanabati, 1993) occurring in the northern and northeastern parts of the CIP, including the Airakan Granite (165±8 Ma) and Torbat-e-
Jam Granite (153±5 Ma). 17
2.5.2.3 Shirkuh Granite
The Shirkuh Granite (>1000 km2) occurs in the CIP (Fig. 2.4). It is a peraluminous S- type calcalkaline granite, characterized by the presence of accessory cordierite, garnet, graphite, andalusite, silhmanife, ilmenite, zircon and apatite (Arnini and Kalantari,
1997). The Shirkuh Granite has a thermal metamorphic contact with Triassic rocks
(Darvichzadeh, 1992; Arnini and Kalantari, 1997) and it is unconformably overlain by
Late Jurassic conglomerate containing pebbles from the underlying granite (Aghanabati,
1993; Kh-Tehrani and Vaziri, 1993). K/Ar dating of K-feldspar grains from Shirkuh
Granite (Table 2.1) gave ages of 186 and 159 Ma (Reyre and Mohafez, 1972). The younger age (159 Ma) is slightly anomalous on the basis of the stratigraphic relationships. The younger age from low temperature feldspar probably reflects loss of radiogenic argon (e.g., Harrison and McDougall, 1981; McDougall and Harrison, 1988).
2.5.2.4 Muteh Granite
The calcalkaline Muteh Granite occurs in the S-SMZ (Fig. 2.4) and is rich in quartz and
K-feldspar, but poor in ferromagnesian minerals (Berberian, 1981). This granite was considered to be Precambrian in age by Theileh et al. (1968), but thermal metamorphism of Early Jurassic rocks indicates that the intrusion is Late Mesozoic or younger in age (Valizadeh, 1992). The petrography and whole rock geochemistry of the
Muteh Granite are very similar to Cretaceous granitoids in the S-SMZ (Otroudi, 1987;
Valizadeh and Ghasemi, 1993; Noorbehesht and Sharifi, 1997).
2.5.2.5 Hamadan Batholith
The Hamadan Batholith is typical of contact-aureole granites in Iran. It intruded the 18
Jurassic Hamadan Phyllite in the S-SMZ (Amidi and Majidi, 1977) and comprises major plutons in the Hamadan, Borujerd and Aligudarz areas (Fig. 2.2). These plutons are surrounded by wide aureoles (up to 5 km) of hornfels and other contact metamorphic rocks, containing cordierite, quartz and graphite, and minor andalusite, staurolite, garnet and biotite (Mohajjel, 1992, 1997). The contact-aureoles are superimposed on various grades of regional metamorphic rocks, including slate, phyllite, biotite schist, amphibolite and gneiss (Darvichzadeh, 1992; Husseini-Doost, 1997). The granite to granodiorite pluton in the Borujerd area intrudes the Hamadan Phyllite; a granite sample from this pluton has been dated by the K/Ar method at 100 Ma (Farhadian, 1991;
Mohajjel, 1997).
Rb/Sr and K/Ar isotopic data from Hamadan Batholith (Table 2.1) gave ages of 104±3 and 82.8±3 Ma respectively for muscovite from pegmatite, 88.5 and 89.6±3 Ma respectively for biotite from the gabbro, and 68±2 and 63.8±2.5 Ma respectively for biotite from the granodiorite (Valizadeh and Cantagrel, 1975). These ages indicate that the Hamadan Batholith emplace*d during the Late Cretaceous times. For gabbro and granodiorite, there is a good agreement between the ages obtained by Rb/Sr and K/Ar methods on biotite-whole rock and biotite, respectively. However, there is significant difference in Rb/Sr and K/Ar ages for the pegmatite. This difference may be the result of loss of radiogenic argon, because there is some evidence that this part of the S-SMZ has been affected by a high thermal gradient as a result of intrusion of several plutons during the Late Cretaceous-Palaeogene times (Section 2.3). This plutonism has been related to closure of the Neo-Tethys Ocean during the Late Mesozoic by many authors (e.g.,
Dercourt etal, 1986; Alavi, 1994). 19
2.5.3 SUBVOLCANIC GRANITOIDS
Subvolcanic granitoids of Iran occur in the CIP and U-DVB. Geochemical data are lacking for most of the subvolcanic granitoids of Iran, particularly for those that occur in the northeastern CIP. Accordingly, the genesis of the intrusive rocks in this part of the country is controversial. Representative subvolcanic granitoids (Fig. 2.5) from the U-
DVB occur in the Natanz, Karkas and Jebal-e-Barez areas (Berberian and Berberian,
1981).
2.5.3.1 Natanz Intrusive Complex
The Natanz Intrusive Complex (Fig. 2.5) comprises calcalkaline rocks, including gabbro, diorite, quartz diorite, quartz monzonite, granite and granodiorite, that intrude folded
Eocene volcanic rocks (Berberian, 1981). The complex is surrounded by a low-grade hornfelsic aureole up to 1.5 km wide (Amidi, 1977). Emplacement began with mafic intrusions and ended with granite and granodiorite. The latter intrusions contain several xenoliths of host volcanic units, as well as early mafic intrusions. Hornblende, biotite and magnetite are the most common ferromagnesian minerals of the Natanz intrusive complex.
Rb/Sr dating of biotite-whole rock pairs (Table 2.1) gave ages of 33.5+1.2 Ma and
25.5±0.5 Ma for gabbro and granite respectively. The corresponding initial Sr/ Sr values are 0.70524 and 0.70573 for gabbro and granite, respectively (Berberian, 1981).
These values indicate that the gabbro and granodiorite originated from a heterogeneous mantle source (e.g., Faure, 1986) or melting of infra crustal source rocks (e.g., Brownlow,
1996). However, Emami and Khalatbari (1997) showed evidence of wall rock assimilation for granitic rocks of the Natanz Intrusive Complex. 20
2.53.2 Karkas and Jebal-e-Barez Intrusions
The Karkas and Jebal-e-Barez intrusions (Fig. 2.5) form part of the U-DVB that is parallel to the Zagros-Central Iranian convergent plate boundary (Berberian et al,
1982). According to Berberian (1981), the last major intrusive episode in Iran occurred during the Oligocene to Early Miocene and is mainly developed in the Karkas and Jebal- e-Barez areas. The medium-grained intrusions are composed mainly of gabbroic to granitic rocks that intrude folded sedimentary and volcanic units of Eocene to Miocene age (Berberian, 1981). The granitic rocks contain abundant magnetite grains and show microgranophyric intergrowths. The volcanic units are mostly andesitic lavas. The
Karkas and Jebal-e-Barez intrusions have produced a thermal aureole in the surrounding andesitic lavas. As occurs in the Natanz Intrusive Complex, felsic intrusive rocks cut the mafic intrusions showing their relatively younger age. Whole rock Rb/Sr isochron of the granodiorite from the Karkas intrusions gave an age of 78 Ma (Reyre and Mohafez,
1972), which is definitely not the age of intrusion since the granodiorite cuts the folded
Eocene volcanic units. Conversely, K/Ar dating of biotite gave ages of 38-33 Ma
(Table 2.1) for the Karkas Granodiorite (Reyre and Mohafez, 1972) and an age of
24+0.1 Ma for the Jebal-e-Barez Granite (Conrad et al, 1977). These ages are consistent with the stratigraphic position of the Karkas and Jebal-e-Barez intrusions.
2.6 SUMMARY
The regional-aureole granitoids of Iran occur in the CIP and intrude high-grade metamorphic rocks of Precambrian age. The emplacement of such granitoids from other parts of the world is considered synchronous with metamorphism (e.g., Wickham, 1986;
Collins et al, 1991; Collins and Vernon, 1991). They may have developed by melting of 21
the older high-grade regional metamorphic rocks (White and Chappell, 1988).
The contact-aureole granites of Iran occur as large batholiths in the CJP and the S-SMZ.
Most of them are Late Mesozoic in age. According to mineralogical features, both I-type metduminous and S-type peraluminous are present. But the S-type granites are dominated and commonly attributed to those plutons showing wide contact-aureoles.
Such S-type granites contain peraluminous minerals and xenoliths that can be matched with the surrounding rocks (Valizadeh, 1992; Husseini-Doost, 1997). In the S-SMZ, deformed and undeformed plutons are mostly circular to elliptical in plan (Fig. 2.2), with the major axes of the ellipses oriented parallel to the northwest-southeast trend of the S-SMZ (Soheili et al, 1992), but undeformed plutons are large and surrounded by pronounced hornfelsic aureoles (Mohajjel, 1997). Preliminary data from some S-type granites of Iran, such as the Hamadan and Shirkuh Granites (Valizadeh, 1992; Amini and Kalantari, 1997; Noorbehesht and Sharifi, 1997), suggest that they are post-tectonic intrusions and a supracrustal protolith is favoured for their generation.
Tertiary granites are mostly subvolcanic intrusions and occur throughout Iran, except in the Zagros and Kopeh Dagh Fold Belts. Contacts with the host volcanic rocks are sometimes low-grade hornfelsic aureole, but commonly steeply deepened or faulted.
These granites are medium-grained and sometimes show granophyric intergrowths and porphyritic textures. Their mineral assemblages indicate that they can be classified as magnetite-series granites (I-type) that tend to occur along the continental margins (e.g.
Ishihara, 1998). They contain basaltic enclaves and have similar minerals to granites from elsewhere that are derived from igneous source regions. Most of them occur along major fault systems, suggesting possibly an extensional structural setting (e.g., 22
Darvichzadeh, 1992). 23
CHAPTER 3
GEOLOGICAL SETTING AND GEOCHRONOLOGY
3.1 GEOLOGICAL SETTING
The granitoids that form the basis of the present study crop out in two distinct geological zones comprising the Taknar Zone, with the Kashmar and Bornavard granitoids, and the
Sabzevar Zone that contains the Kuh Mish intrusions (Fig. 1.2).
3.1.1 TAKNAR ZONE
The Taknar Zone (Lindenberg and Jacobshagen, 1983) is a wedge-shaped region between two major fault systems, the Rivash Fault to the north and the Doruneh Fault to the south
(Fig. 1.2). The southern part of the Taknar Zone extends to the east at least to the boundary of Iran and Afghanistan. The basement of the Taknar Zone consists of metavolcanic rocks
(andesitic lava and tuff) of probably Precambrian age (Eftekhar-Nezhad, 1976). The
Bornavard granitoid intrudes the metavolcanic rocks (Fig. 3.1). In the Taknar Zone, Late
Palaeozoic rocks occur as thin sequences and are scattered around the most western and southern parts of the zone (Muller and Walter, 1983). In the southern parts of the zone, Late
Palaeozoic - Early Mesozoic dolomite is unconformably overlain by Late Cretaceous limestone (Fig. 3.1). The Palaeozoic and Mesozoic sedimentary rocks of the Taknar Zone cannot be correlated with rocks of similar age in the Sabzevar Zone (Sepahi, 1992). The eastern part of the Taknar Zone is widely covered by volcano-sedimentary rocks of Eocene age (Fig. 1.2) and is intruded by the Kashmar granitoid. 24
The metavolcanic rocks of the Taknar Zone are characterised by a thick sequence (about
2000 m) of essentially andesitic lava and tuff, with intercalations of sandstone and dolomite. The contact between the metavolcanic unit and the underlying rocks is not clear.
The metavolcanic rocks are lithologically subdivided into five groups (Muller and Walter,
1983) comprising:
(a) light coloured rhyolite containing quartz phenocrysts in a fine-grained groundmass, and
showing a distinctive flow texture;
(b) dark-grey to black coloured rhyolite containing phenocrysts of quartz and K-feldspar in
a very fine-grained groundmass;
(c) green-grey coloured rhyolite contains large phenocrysts of quartz in a fine-grained
groundmass. This rock is sometimes intercalated with metarhyodacite;
(d) fine-grained uniform tuff, grey to dark green colour and composed of quartz and
feldspar; and
(e) light-green coloured banded tuff containing various amounts of quartz grains.
3.1.2 SABZEVAR ZONE
The Sabzevar Zone occurs between the southern border of the Alborz Belt and the northern
border of the Rivash Fault. This zone is subdivided into four geological units (Lindenberg
et al, 1983) comprising:
(a) andesitic and basaltic rocks (Cretaceous in age);
(b) ophiolitic melanges (Early to Late Palaeogene in age) along the southern and northern
parts of the zone; 25
(c) volcano-sedimentary rocks (Eocene in age); and
(d) Kuh Mish intrusions (Middle-Late Eocene in age).
The andesitic and basaltic rocks are interbedded with marine sedimentary strata comprising thin-bedded chert, radiolarite, marl and limestone. These rocks are usually rich in
Globotruncanas indicating a Late Cretaceous age. The ophiolitic melanges consist of a large variety of ultramafic and typical flysch-like deposits, occurring in the vicinity of the
Late Cretaceous rocks. Based on their fossil contents, the age of the ophiolitic melanges ranges from Early to Late Palaeogene (Lindenberg and Jacobshagen, 1983). The volcano- sedimentary rocks occur in the central parts of the zone (Eftekhar-Nezhad, 1976). The volcanic units are generally andesitic lavas and tuffs, whereas the sedimentary rocks include conglomerate, sandstone, limestone, marl and evaporite. The marl and limestone are
Nummulites-bear'mg indicating a Late Palaeocene to Middle Eocene age (Lindenberg et al,
1983). The Kuh Mish intrusions occur around the prominent mountain of Kuh Mish where they intrude the Late Cretaceous-Early Tertiary andesitic and basaltic rocks. The intrusions are mainly gabbro, quartz monzodiorite and granodiorite. The current study presents the first petrographic and chemical information for the Kuh Mish intrusions.
3.2 MAJOR FAULT SYSTEMS
3.2.1 DORUNEH FAULT
The presently active Doruneh Fault is east-west trending, with a slight convexity to the north, and merges with the Rivash Fault to the west of the Taknar Zone (Fig. 1.2;
Lindenberg et al, 1983; Lindenberg and Jacobshagen, 1983). In a broad sense, the Doruneh 26
Faultfits a model of a composite fracture zone with a general trend towards sinistral and strike-slip movement (Jackson et al, 1995).
3.2.2 RIVASH FAULT
The Rivash Fault separates the northern border of the Taknar Zone from the southern border of the Sabzevar Zone (Fig. 1.2). Major lateral movements have occurred along the
Rivash Fault. The Cretaceous andesitic and basaltic rocks of the Sabzevar Zone never occur in the Taknar Zone and Tertiary facies show abrupt changes across the Rivash Fault. This can be related to the lateral movements of the fault, in combination with a continuous uplift of the Taknar Zone after the Middle Eocene (Davoudzadeh et al, 1981; Lindenberg and
Jacobshagen, 1983).
3.3 GEOCHRONOLOGY
3.3.1 KASHMAR GRANITOID
The Kashmar granitoid (-200 km2) occurs throughout the northern parts of the Kashmar area (Fig. 3.2) located to the east of the Taknar Zone (Fig. 1.2). It is bordered by the
Doruneh Fault to the south and the Rivash Fault to the north. It forms the central part of the
'north Doruneh Fault magmatic belt' that runs from the northern to eastern parts of the CIP for a length of -300 km with a width of -10 km (Emami et al, 1993). In this belt, volcanic and plutonic rocks occur in an arc with a convexity to the north, defining the margins of the
CJP (Taheri, 1999). The Kashmar granitoid is partly overlain in the south by a Neogene sandstone -100-150 m thick. Where observed, contact of the granitoid with the surrounding 27
rocks is either faulted or narrow homblende-hornfelsic rims have developed in the Eocene volcanic rocks.
The Kashmar granitoid is composed of four major plutons including tonalite, granodiorite, granite, and alkali feldspar granite. Among these, granodiorite and granite plutons constitute approximately 90% of the granitoid exposure. They are mainly light grey to dark brown in colour and occupy much of the northern outcrop area, although in places they are the marginal constituents of the southern areas. Contacts between different plutons are sharp and mostly faulted. Along the faulted contacts, particularly between granite and granodiorite, several alteration products have been developed. These plutons have been cut by aplitic and dacitic dykes. The alkali feldspar granite is the most homogeneous pluton and is distinguished in the field by a light pink-cream colour. It occupies much of the southern margin of the Kashmar granitoid. Rocks of this alkali feldspar pluton are medium-grained.
The homogeneous interlocking quartz-feldspar textures and the presence of nearly aphyritic volcanic rocks at the same structural level indicate that these plutons were emplaced at very high levels in the crust and solidified contemporaneously (e.g., Kistler and Swanson, 1981;
Turner et al, 1992). In general, contacts between different plutons, together with a longitudinal shape of the igneous assemblage, are consistent with the east-west trend of the
Doruneh Fault to the south. This implies that the emplacement of these plutons is most likely related to the activity of the Doruneh Fault. 28
3.3.1.1 Rb/Sr Age Dating
Due to the widespread occurrence of fresh biotite in the rocks studied, the Rb/Sr method on biotite-whole rock pairs was selected for isotopic age determination. Ages are calculated using a two point isochron method in which low concentrations of Rb in the whole rock and very high concentrations of Rb in biotite are used as a control on the accuracy of the initial ratio. The significance of the Rb/Sr method results from the fact that Rb and Sr have a very close geochemical relationship to K and Ca, respectively (Geyh and Schleicher, 1990).
These elements are important in magmatic processes, as they can provide important information, particularly for the petrology of granites. For example, Sr substitutes predominantly in early phase minerals such as apatite and plagioclase, but Rb becomes enriched in residual melts. This leads to a large variability in the Rb/Sr ratio during differentiation and, therefore, provides ideal conditions for the isochron method (Faure,
1986; Geyh and Schleicher, 1990; Brownlow, 1996). The Rb/Sr data are summarised in
Table 3.1. Sample preparation and analytical methods are presented in Appendix 1. Results presented in this thesis are the first Rb/Sr isotopic ages recorded for the Kashmar and
Bornavard granitoids.
3.3.1.2 IsotopicData
On the basis of stratigraphy and comparison with subvolcanic intrusions occurring in the
CP, the Kashmar granitoid was mapped as Tertiary in age by Eftekhar-Nezhad (1976) and
Taheri (1999). The only isotopic age data for the host volcanic rocks to the north of
Kashmar city have been reported by Bernhardt (1983), giving ages of 57.2±3.7 and 29
43.7±1.7 Ma for K/Ar dates on hornblende and biotite, respectively. As hornblende has the capacity to be more retentive with respect to 40Ar than biotite (Faure, 1986; McDougall and
Harrison, 1988), the age of the hornblende is interpreted as the emplacement age of the volcanic rocks, whereas the younger age for biotite may be the result of loss of 40Ar due to the emplacement of a subvolcanic granitoid. This biotite age of 43.7±1.7 Ma is very similar to the ages obtained by the Rb/Sr dating method on several biotite-whole rock pairs from different plutons of the Kashmar granitoid (Table 3.1).
Also, Tertiary volcanic rocks occur extensively in the Gonabad and Bejestan areas, located to the southern parts of the Kashmar granitoid. Bina et al (1986) obtained whole rock K/Ar ages of 61±2 and 54±2Ma for andesitic rocks from the Gonabad and Bejestan areas, respectively (Table 2.1). The age of 57.2±3.7Ma (Bernhardt, 1983) for hornblende from volcanic rocks of the Kashmar area is within the 2o analytical uncertainty of the isotopic ages for andesitic rocks from the Gonabad and Bejestan areas. These isotopic data indicate
Early Eocene volcanic activity in the northeastern CIP.
In the present study, eight samples from three major plutons (granodiorite, granite and alkali feldspar granite) of the Kashmar granitoid were selected for Rb/Sr dating of biotite- whole rock pairs. Biotite was separated from four samples including one alkali feldspar granite, one granodiorite and two granites. The biotites from all dated rocks have been analysed by electron microprobe (Sections 4.2.3 and 4.4.4). 30
Rb/Sr data for the whole rock and biotite separates, together with calculated ages for the
Kashmar granitoid, are presented in Table 3.1. The four isotopic ages for different plutons of the Kashmar granitoid range from 43.5±0.4 to 42.4+0.4 Ma, indicating Middle-Late
Eocene plutonism in northeastern CIP. These ages are essentially indistinguishable with ages differences of - 1 Ma being negligible for extensive magmatic suites (e.g., Ewart et al, 1992; Nakajima, 1996).
The 87Rb/86Sr and 87Sr/86Sr values for whole rock (six samples) and biotite (four samples) of the Kashmar granitoid define an isochron (Fig. 3.3), with MSWD (mean square weighted deviate) value of 5.3. Two whole rock samples from alkali feldspar granite are not plotted in Figure 3.3 because they produce high MSWD. The slope of the isochron is clearly constrained by biotite samples because they are significantly higher in Rb/ Sr and
87Sr/86Sr values. This isochron yields an age of 42.8±0.2 Ma and an initial 87Sr/86Sr value of
0.70548±0.00003. The age and initial 87Sr/86Sr value given by the isochron are within the limited range of age and initial 87Sr/86Sr values obtained only by biotite-whole rock pairs
(Table 3.1). For each pluton similarity in initial 87Sr/86Sr values indicate that the total rock
Rb/Sr system was closed simultaneously, and that the Rb/Sr system has remained a closed system since the time of emplacement (e.g., Faure, 1986; Pollard et al, 1995).
The accuracy of the isotopic ages for biotite-whole rock pairs from the Kashmar granitoid is dependant upon the closed system behaviour of biotite (e.g., Ganguly and Ruiz, 1986;
Milner et al, 1993). Although Late Palaeogene volcanic activity has been reported in this area (Behroozi, 1987; Taheri, 1999), the Kashmar granitoid was not subjected to 31
disturbance by younger thermal events. The similarity in isotopic ages from biotite-whole rock pairs clearly records a short duration for the emplacement of different plutons of the
Kashmar granitoid in a subvolcanic environment.
hi view of the field data, the ages obtained support a close genetic connection between host- volcanic rocks and granitoid emplacement. The volcanic activity in the Kashmar area is believed to have started approximately 57.2±3.7 Ma (Table 2.1), continued through the
Early to Middle Eocene (Bernhardt, 1983) and was followed by intrusion of the Kashmar granitoid at 42.8+0.5 Ma, suggesting that magmatic activity occurred over an interval of
-15 million years. Such a duration for intense magmatic activity is not unusual (e.g., Milner et al, 1993). The precise Rb/Sr ages determined for biotite-whole rock pairs (43.5-42.4 Ma) of the Kashmar granitoid, when taken together with the few reliable published isotopic ages for volcanic rocks (57-43 Ma), suggest that these subvolcanic plutons were emplaced after or contemporaneous with widespread volcanic extrusions (mostly andesite) during the
Middle Eocene. The east-west distribution (-300 km) of volcanic and plutonic rocks along the northern parts of the Doruneh Fault may be interpreted to be related to extensional magmatism resulting from the upwelling of large volumes of magma being focused along a structural discontinuity in the northeastern margins of the CJP.
3.3.2 BORNAVARD GRANITOID
The Bornavard granitoid (39 km2) includes the oldest intrusive rocks in the Taknar Zone
(Fig. 2.8). The granitoid occurs in the western part of the Taknar Zone (Razzaghmanesh,
1968; Muller and Walter, 1983). It comprises three distinct plutons mainly of tonalite, 32
granodiorite and granite composition. The tonalite and granodiorite are dark green in colour and occur in the central part of the granitoid. Aplitic, dioritic and doleritic dykes intrude tonalite and granodiorite. Hornblende and biotite are the most common ferromagnesian minerals occurring in tonalite and granodiorite, respectively. Xenoliths rich in biotite, plagioclase and quartz are common in granodiorite. Granite intrudes into the external margins of the granodiorite. It is light pink in colour and shows chilled margins and narrow thermal contacts to the metavolcanic units. Doleritic dykes intrude granite especially in the southern part. Hydrothermal alteration has produced quartz-chlorite viens that usually cut the granite near the major joint systems.
The age and the origin of the Bornavard granitoid have long been the subject of interest in
Iran, particularly given their association with metavolcanic units that have been compared with the Precambrian basement rocks in other parts of the CJP (Homam, 1992; Esmaeili et al, 1998a). Based on field observations, the Bornavard granitoid was emplaced during two major magmatic episodes. The early episode is characterised by small plutons of mainly dark tonalite and granodiorite, occurring in the inner parts of the granitoid. Contacts between these plutons are faulted. The later episode is characterised by emplacement of a larger granitic pluton in the outer parts. The present study confirms the younger age of the granitic pluton by Rb/Sr dating on biotite-whole rock pairs.
3.3.2.1 IsotopicData
Rb/Sr isotopic data on biotite-whole rock pairs clearly distinguish at least two different plutonic episodes for the Bornavard granitoid (Table 3.1). Two samples from the 33
granodiorite, representing the early intrusive episode, yield precise ages of 152.8±1.3
(Sample R15947) and 145.6±1.3 Ma (Sample R15946), indicating that the oldest plutonic activity in the Bornavard area occurred during the Late Jurassic (Tithonian; Harland et al,
1990). The younger age for Sample R15946 may be due to episodic emplacement of the granodiorite. The Rb and Sr concentrations in biotite from both samples are nearly similar and indicate that the biotite from younger sample is not altered. Because Sample R15946 has taken close to the late intrusive episode, it is possible that the isotopic system was disturbed by thermal effect.
The isotopic ages of the granodiorite (153.8 and 145.6 Ma) are similar to the age of some contact-aureole granitoids of the northeastern CJP. For example, the age of 145.6±1.3 Ma is indistinguishable from the age of 146±3 Ma from the Mashhad Granite (Table 2.1). hi particular, the age of 145.6 Ma, which corresponds to the boundary of the Jurassic-
Cretaceous (Harland et al, 1990), confirms granite magmatism is related to the Mid-
Cimmerian Orogeny, recognised in the CIP by Alavi-Naini (1992) and Aghanabati (1993).
This orogeny is recognised in Iran by a lack of sedimentary deposits during the Early
Cretaceous (Neocomian) in the Alborz Belt and an angular unconformity between Jurassic and Cretaceous sedimentary rocks in the CJP (Darvichzadeh, 1992).
Rb/Sr data for two pairs of biotite-whole rock from the granitic pluton (later episode) yielded ages of 123.8+1 (Sample R15938) and 111.8±1.1 Ma (Sample R15941), indicating that plutonic activity in the Bornavard area continued episodically until the Early to Late
Aptian (Harland et al, 1990). There is significant difference in Rb/Sr ages of granite 34
samples. Both samples have similar whole rock Sr contents (39 ppm). Several biotite grains from each sample were analysed by electron microprobe and they are very similar in composition (Section 4.4.4). The Rb content of younger sample (R15941, Rb = 128 ppm) is higher than Rb content of older sample (R15938, Rb = 95 ppm) and there is no indication for hydrothermal alteration. Therefore, younger isotopic age of sample R15941 may be related to episodic emplacement of the late intrusive episode.
Assuming that the age of 123.8 Ma is the age of emplacement of the granite pluton, a time interval of -22 million years between two major magmatic episodes for the Bornavard granitoid is suggested. This time interval is significant, but it seems to be common for emplacement of granites along continental margins. Nakajima et al. (1990) and Nakajima
(1994, 1996) examined the Rb-Sr biotite-whole rock isochron ages of the Cretaceous I-type granites of southwest Japan and showed episodic emplacement from 100 Ma to 70 Ma, with periods of 5-10 million years between emplacement and cooling. They suggested different magma sources and showed an increase in initial 87Sr/86Sr ratios towards younger magmatic events. Similarly, in the Bornavard granitoid, a -22 million years age difference between emplacement of the granodiorite and granite plutons resulted in a complete change in the source of magma generation. This is evident from the extremely high initial Sr/ Sr values
(0.73978-0.75008) and the very felsic composition of granite (Si02 = 74.84-76.04 wt%), compared with the granodiorite. 35
3.3.2.2 Age Discussion on the Bornavard Granitoid
On the geological map of the Kashmar area (1:250 000; Eftekhar-Nezhad, 1976), the central part of the Bornavard granitoid, which is characterised by tonalitic and granodioritic rocks, has been labeled as Tertiary in age. It appears that this age was based on a petrographic comparison with other Tertiary granites in the Taknar Zone. The present level of exposure of the Bornavard granitoid precludes determination of any stratigraphic relationship to
Tertiary rocks. Also, Tertiary granites do not crop out in this part of the Taknar Zone. The isotopic ages on biotite-whole rock pairs in the present study indicate that the central part of the Bornavard granitoid formed in the Late Jurassic (152.8-145.6 Ma).
The above mentioned map also indicates a Doran-type (Precambrian) granite occurring in the external parts of the Bornavard granitoid and intruding the metavolcanic rocks. The granite and metavolcanic rocks were compared with regional-aureole granites of the CIP
(e.g., Doran and Muteh granites; Razzaghmanesh, 1968). This comparison was based on petrography and minor stratigraphic relationships (Huckriede et al, 1962; Stocklin and
Setudehnia, 1991) and it has now been shown to be incorrect. A Precambrian age for the
Doran Granite itself is suspect because using the Rb/Sr method on a biotite-whole rock pair from the granite yielded an age of 175+5 Ma (Crawford, 1977). The recent magmatic map of Iran, compiled by Emami et al. (1993), correctly assigned the bulk of the Bornavard granitoid as Mesozoic calcalkaline plutonic rocks, but without isotopic evidence.
Granites related to Late Jurassic and Early Cretaceous plutonism, particularly in the north and northeastern CIP, have been recognised (e.g., Darvichzadeh, 1992; Aghanabati, 1993). 36
For example, the Airakan Granite occurs in the southwestern limit of the Bornavard granitoid. Rb/Sr dating on whole rock samples from the Airakan Granite yielded an age of
165±8Ma (Table 2.1) that is close to the oldest age of the Bornavard granitoid.
Furthermore, Aghanabati (1993) reported an age of 153±5 Ma for the Torbat-e-Jam Granite that occurs in the easternmost part of the Taknar Zone. This age is similar to the older age of granodiorite (152.8 Ma) from the Bornavard granitoid (Table 3.1). Additionally, the
K/Ar ages of 146±3 and 120±3 Ma (Table 2.1) for biotites from the Mashhad Granite that occurs in the northeast of the CIP supports the Late Jurassic Early Cretaceous plutonism in this part of Iran. The K/Ar ages of biotite from the Mashhad Granite are very similar to the
Rb/Sr biotite-whole rock ages of the granodiorite (145.6 Ma) and granite (123.8 Ma) from the Bornavard granitoid.
In addition, Crawford (1977) reported a low-grade metamorphic event that occurred between 250 and 190 Ma for the Taknar Rhyolite, based on Rb/Sr dating of whole rock samples. Muller and Walter (1983) confirmed that metamorphism of some pelitic rocks of the Taknar Zone occurred during this interval. These authors clearly reject the Precambrian age for metamorphism in the Taknar Zone. Assuming that plutons of the Bornavard granitoid were emplaced before metamorphism of the Taknar Zone, they would have been isotopically homogenised and similar in ages. But isotopic data from the Bornavard granitoid (Table 3.1) distinguish a series of age ranging from Late Jurassic to Early
Cretaceous, all are significantly younger than the age of metamorphism of the Taknar Zone.
Microscopic examinations of the rocks from the Bornavard granitoid do not show any evidence of metamorphism. Also, the Bornavard granitoid shows a thermal contact to the 37
metavolcanic rocks. These features indicate that comparison of the Bornavard granitoid with the Precambrian regional-aureole granitoids of the CIP is meaningless. The Rb/Sr isotopic ages of Late Jurassic and Early Cretaceous, respectively for the early and late episodic rocks of the Bornavard granitoid are considered to be the best estimate of the timing of the most intense plutonic activity in the Middle East (Laws and Wilson, 1997).
This plutonic activity is related to the Middle to Late Cimmerian Orogeny, recognised in
Iran by intrusion of several granitoid bodies in the CIP and the S-SMZ (Alavi-Naini, 1992;
Darvichzadeh, 1992; Emami etal, 1993). 38
CHAPTER 4 PETROGRAPHY AND MINERAL CHEMISTRY
4.1 PETROGRAPHY OF KASHMAR GRANITOID
Analytical methods and modal mineralogy for representative samples of four plutons from the Kashmar granitoid are listed in Appendices 1 and 2 respectively. Modal data for the Kashmar granitoid indicate that tonalite, granodiorite, granite and alkali feldspar granite are the only plutonic rocks that occur in the granitoid mass (Fig. 4.1).
Representative petrographic features for each pluton are outlined in the following sections.
4.1.1 TONALITE
Tonalite comprises the smallest pluton occurring in the north of the Kashmar granitoid.
It is green to grey in colour, and intrudes into andesitic lavas of Eocene age. Also, it has sharp contacts with granodiorite and granite plutons. The tonalite is a medium-grained rock and contains a lower modal content of quartz compared with granite and alkali feldspar granite (Appendix 2.1). Microgranular enclaves are common in tonalite. Most of them range between 3 and 4 cm in size that could be related to the rheological properties of the magma; that is, enclaves greater than the size distribution observed would have mostly been left behind and separated during upward movement of the host magma. The modal content of total ferromagnesian minerals in the tonalite is up to
37.4% (Sample R15911), which is higher than in other plutons of the Kashmar granitoid. Plagioclase (44.2-58% modal), quartz (12.8-18% modal), amphibole (8.2-
20.4% modal) and biotite (up to 17.8% modal) are major mineral components of the 39
tonalite (Appendix 2.1). Plagioclase crystals are commonly euhedral with normal zoning and the rock has a hypidiomorphic granular texture, although some marginal samples are slightly glomeroporphyritic. Amphibole grains occur either as prismatic crystals or as irregular grains. They have been variably converted to biotite. Sometimes, amphibole grains form clusters together with magnetite and biotite (e.g., Sample R15912).
4.1.2 GRANODIORITE
The granodiorite is grey to black in colour and essentially medium-grained (2-3 mm), with a few plagioclase crystals up to 5 mm long. Plagioclase, quartz, amphibole, biotite and K-feldspar are the major mineral components of the granodiorite. Plagioclase constitutes an average of 45% by volume of the rock, ranging from 36.8% to 60.4%
(Appendix 2.1). Quartz and K-feldspar are interstitial, or occur as interlocking anhedral grains. Hypidiomorphic granular textures are the most commonly observed textures in thin sections, although some samples show allotriomorphic granular and micrographic textures. Dark microgranular enclaves up to 30 cm across, with igneous textures, occur in the granodiorite. Most enclaves are spherical in shape and 3-4 cm in size. They are rich in amphibole, plagioclase and apatite and seem to be mineralogically related to the host granodiorite. They show some characteristics of common inclusions found in I-type granites and they may be restite. The enclaves taken from marginal outcrops display porphyritic textures and a similar mineralogy to the host-volcanic rocks. Such enclaves are interpreted as accidental inclusions of wall-rock.
4.1.3 GRANITE
Granite is the most abundant rock-type in the Kashmar granitoid. The granite is medium-grained (3-4 mm) and equigranular, with fresh exposures being light grey to 40
white in colour. It consists of plagioclase, K-feldspar, quartz, amphibole, and biotite as major minerals. Modal data show that the content of a particular mineral is more variable in the granite, compared with other plutons of the Kashmar granitoid. For example, plagioclase ranges from 19.8 to 39.6% by volume of the rock. When biotite and amphibole coexist, the latter mineral always occurs in low amounts (Appendix 2.1), but some samples from the granite are completely lacking in biotite or amphibole. Such variation in mineral content may be a result of fractional crystallisation in the granitic pluton. K-feldspar is sometimes microperthitic, although commonly it occurs in microgranophyric intergrowths particularly when hornblende and biotite are absent. A noteworthy characteristic of the granite is the magmatic reaction in which amphibole is converted to biotite. Rare clinopyroxene cores are found in the hornblende crystals of the granite (Sample R15903).
4.1.4 ALKALI FELDSPAR GRANITE
This is a medium-grained rock (minerals 2-3 mm in size) which, in hand specimen, is characterised by a light cream to pinkish colour. A low content of mafic silicates and lack of amphibole differentiate the alkali feldspar granite from other plutons of the
Kashmar granitoid. The essential modal constituents are microperthite (53-63%), quartz
(28-37%) and sporadic crystals of biotite (up to 4.6%; Appendix 2.1). Biotite is fresh, blue-green in colour and medium-grained. Accessory apatite occurs as tiny slender prisms within biotite that may represent a 'primary restite phase' that was incorporated into the biotite crystals when pyroxene reacted with the magma (Chappell et al, 1987).
Magnetite, minor ilmenite, zircon, titanite and allanite commonly accompany the biotite.
Such accessory minerals commonly occur in I-type granites. In particular, allanite may 41
contain appreciable ferric iron (Hine et al, 1978) and indicates high f02 that is characteristic of many I-type granites.
4.2 MINERAL CHEMISTRY OF KASHMAR GRANITOID 4.2.1 PLAGIOCLASE
Plagioclase is usually the most abundant mineral that occurs in the tonalite, granodiorite and granite, where it commonly occurs as subhedral to euhedral crystals. Plagioclase crystals mostly range between 3 and 5 mm in size, with decreasing size and abundance towards more felsic variants. They sometimes contain inclusions of hornblende, Fe-Ti oxide and long needles, but commonly prismatic crystals of apatite (e.g., R15912,
R15958). In some samples of tonalite, plagioclase occurs as larger euhedral grains that are surrounded by fine-grained, lath-shaped crystals of the same mineral. The lath- shaped plagioclase crystals have strongly resorbed boundaries (e.g., Sample R15912).
Polysynthetic twinning is the most common feature of plagioclase grains and they are often twinned after the albite and pericline laws. Plagioclase crystals from the granodiorite and granite typically show normal zoning, which is attributed to normal magmatic fractionation.
Electron microprobe analyses of plagioclase crystals from the Kashmar granitoid are shown in Appendix 3.1. The composition of plagioclase ranges from A1148-18 in the
granodiorite, An50-i5 in the granite and An25-i6 in the alkali feldspar granite. The most calcic plagioclase core is An5o that occurs in the granite (R15910). All the compositional data obtained for the plagioclase cores and rims are plotted in Figure 4.2. Chemical data show that most of plagioclase cores and rims are Anso-30 and AH30-18, respectively. The higher anorthite content of crystal cores is consistent with normal zoning that is 42
observed in thin sections. In normally-zoned grains, the maximum difference between the composition of core and rim is 26 mole% anorthite (Appendix 3.1). There is a good similarity in anorthite content of plagioclase from the granodiorite and granite. Most plagioclase grains from the alkali feldspar, granite (Samples R15914 and R15900) are typically homogeneous and low in anorthite content ( Sometimes plagioclase cores have embayed margins, or are rounded (Sample R15908), indicating partial resorption. The compositional range of such cores is low (e.g., An34_ 39). During fractional crystallisation, if plagioclase precipitates from a granitic melt, it is not uniform in composition (Chappell et al, 1987; Hall, 1987) because plagioclase is a very difficult mineral to re-equilibrate with the melt, especially at temperatures less than ~1000°C (Johannes, 1978; Chappell, 1996b). The uniform calcic plagioclase cores are interpreted as representing restite by some workers (White and Chappell, 1977; Chappell etal, 1987, 1988,1999; Champion, 1991; Chappell, 1996b). If the plagioclase system were in perfect equilibrium, crystals would react continuously with the melt to produce unzoned plagioclase. The fact that normally zoned crystals are common shows that the kinetics for plagioclase equilibration are slow and virtually all igneous plagioclases are zoned (Shelley, 1993). Also, several possibilities can prevent equilibrium crystallisation. For example, intracrystalline diffusion in plagioclase is essentially slow because of the incorporation of high charge elements with large ionic radius (Hess, 1989; Deer et al, 1992; Shelley, 1993). Some authors believe that crustal contamination (Wilson, 1989; Mason, 1996), high viscosity of magma (Ragland, 1989) or change in water pressure and temperature (Loomis, 1982; Mason, 1985; Holtz et al, 1995; Singer, 1995) lead to the development of zoned crystals. Because the Kashmar 43 granitoid emplaced in a subvolcanic environment it is highly probable that temperatures dropped too fast to exchange reactions to produce unzoned plagioclase grains. In general, plagioclase grains are low in Ti02 (<0.06 wt%) and total Fe as FeO (<0.33wt%) but high in Si02 (up to 64.8 wt%) and A1203 (up to 29.4 wt%; Appendix 3.1). Fe is always less than 0.1 a.f.u. in the structural formula of each plagioclase grain. Because Fe3+ replaces Al3+ and Fe2+ replaces Ca2+ (e.g., Deer et al, 1992), high AI2O3 contents of plagioclase grains from the Kashmar granitoid may be related to lower activity of iron during plagioclase crystallisation or low iron content in the source. The plagioclase crystals normally contain some orthoclase (KAlSisOg) in solid solution, varying up to Ors.s from andesine to oligoclase and tending to increase towards the more Na-rich plagioclase rims (e.g., Sample R15958, Ab78.s). This is expected from normal magmatic reactions that observed in most granite and granodiorite samples. 4.2.2 AMPHIBOLE Amphibole is the major mafic mineral in the tonalite ranging from 8.2% to 20.4% by volume of the rock (Appendix 2.1). It forms rarely euhedral but commonly subhedral crystals with a pleochroic scheme (X and Y = pale green to green, Z = straw-yellow), and a maximum average grain size of 1.4 mm. It is associated with biotite, titanite and Fe-Ti oxide. Both amphibole and biotite decrease in abundance towards the more silica- rich variants. Naney (1983) and Hogan and Gilbert (1995) showed that for magmas crystallising even at low H20 contents and low confining pressures (<3.5 wt% at 200 MPa), amphibole reacts with the melt to form biotite. Such a reaction is evidenced by the fresh fine-grained biotite that replaces amphibole around rims and along the 44 cleavages (e.g., Sample R15904). In the reaction area magnetite and titanite are usually common, suggesting oxidising conditions (Mason, 1985; Hammarstrom and Zen, 1986). As indicated in several publications (e.g., Elliott et al, 1998), it is possible that amphibole replaces pyroxene with increasing H20 content in the magma; a natural consequence of crystallisation of early anhydrous phases. In the Kashmar granitoid, small clinopyroxene relicts within amphibole occur only in one sample of granite (R15903). It seems that the H20 content of magmas was high enough to change pyroxene completely to amphibole (Bateman and Chappell, 1979; Wilson, 1989). Therefore, most amphibole grains in the Kashmar granitoid can be presumed to result from normal magmatic reactions. Then, at higher H20 pressures amphibole reacted with the melt to produce biotite. This explains the occurrence of amphibole cores within biotite. Microprobe analyses along with structural formulae for amphibole are listed in Appendix 3.2. To determine the structural formulae of amphibole, maximum Fe3+ was estimated using the methodology of Robinson et al. (1982, pp. 6-10). Mw was calculated as the difference between 8.0 cations (full tetrahedral occupancy) and the number of Si cations. Following the recommendations of Leake (1978) and Deer et al (1997) all the Kashmar amphiboles are calcic amphiboles [Ca(M4) + Na(M4) >1.34; Na(M4) <0.67]. Their (Na+K)A and Ti are both always less than 0.50, indicating magnesio-hornblende which is typical of I-type granites (Czamanske et al, 1981). The hornblende grains encountered in the present study are chemically homogeneous. Hornblende crystals from different samples of granodiorite are similar in composition, whereas those from the granite are more varied in composition. This is consistent with 45 the observation that the granodiorite pluton is more homogeneous than the granite pluton. The chemistry of a crystallising hornblende is sensitive to pressure, temperature,/02 and FH20 (Rutherford et al, 1985; Johnson and Rutherford, 1989a,b; Blundy and Holland, 1990; Schmidt, 1992). Hornblendes from the Kashmar granitoid are characteristically low in A1203 (1.74 to 8.41 wt%, total Al <1.5 a.f.u.) and Ti02 (0.13 to 2.37 wt%, Ti <0.27 a.f.u.) which is indicative of low-temperature, high f02 crystallisation (Mason, 1978; Hammarstrom and Zen, 1986; Hollister et al, 1987). Octahedral Al (A1VI) is less than 0.1 a.f.u., and Fe3+ is much higher (up to 0.93 a.f.u.) than Al^ in the Kashmar hornblende. These features are typical of low-pressure calcic amphiboles (Leake, 1971; Mason, 1985; Green, 1992) from shallow-level intrusions (e.g., Wyborn, 1983; Hammarstrom and Zen, 1986). According to the geobarometer of Johnson and Rutherford (1989a), the total Al content of hornblende (mostly <1.3 a.f.u.) from Kashmar implies variable pressures up to a maximum of 3 kbar (Fig. 4.3). The range of indicated pressures record polybaric crystallisation during ascent through the crust before emplacement near the surface. The contents of FeO (13 to 18.93 wt%) and MgO (10.45 to 15.71 wt%) are relatively high in the analysed hornblendes, whereas the MgO/FeO is low (1.20 to 0.54), mostly <1, indicating hornblende crystallised from a felsic melt (Gribble, 1988). Variation between FeO and MgO contents in amphibole depends on f02 which can be assessed by Mg/(Mg + Fe2+) or FeO/(FeO + MgO) (Czamanske et al, 1981). In the Kashmar granitoid, the Mg/(Mg + Fe2+) for all analysed hornblendes, is high and ranging from 46 0.60 to 0.77 (Appendix 3.2), again suggesting low pressure, bigh/02 for crystallisation (e.g., Hammarstrom and Zen, 1986; Anderson and Smith, 1995). Hornblende from the granite (e.g., Sample R15910) shows the highest Mg/(Mg + Fe2+) (0.77) and the lowest FeO content (13.35 wt%) among all analyses listed in Appendix 3.2. Modal content of Fe-Ti oxide from this sample (3.2 vol %) is the highest for all granite samples from the granitoid. The above features may support higher /02 for hornblende crystallisation in granite. Because at higher /02 magnetite precipitates and lowers the activity of FeO in the melt, consequently hornblende crystallises with a low FeO/MgO contents but high Mg/(Mg + Fe2+) values (Mason, 1978; Brownlow, 1996). Fractionation factors between amphibole and magma composition for MgO, FeO and Ti02 were determined by dividing each of the oxide concentrations in the mineral by that in the whole rock. The MgO fractionation factor for amphibole grains in igneous rocks ranges from approximately 2 for high temperature magmas (~1000°C) to 10 at lower temperatures (~800°C; Cawthorn, 1976). Except for Sample R15909 (granite), the calculated MgO fractionation factor for all magnesio-hornblendes from the Kashmar granitoid ranges from 5.79 to 7.05, indicating that crystallisation occurred at moderate to low temperatures. Cawthorn (1976) determined a similar temperature dependence for FeO as for MgO with values ranging from 2 to 6. The FeO fractionation factor for magnesio-hornblendes from the Kashmar granitoid ranges from 2.75 to 8.21, but most ratios are >3 being consistent with low temperature crystallisation. The Ti02 fractionation parameter is pressure-sensitive, decreasing from a value of 5-10 for extrusive rocks, to 1-5 for rocks crystallised at crustal pressures (Cawthorn, 1976). The 47 Ti02 fractionation factor for Kashmar hornblende mostly ranges from 1 to 4.65, indicating the hornblende grains crystallised at crustal pressures. Hornblende grains from the Kashmar granitoid are high in MnO contents (0.36- 0.69 wt%, 2-20 times greater than the whole rock values). The MnO contents of hornblende show inverse relationship to whole rock MnO contents. Similar MnO behaviour has been reported from subduction-related I-type granites in north Queensland, Australia (Champion, 1991). Collectively, all amphibole grains from the Kashmar granitoid show evidence of low temperature but high y02 and PH20 conditions. Magnesio-hornblendes from the Kashmar granitoid are compositionally similar to hornblende grains from the shallow-level calcalkaline plutons of eastern Peninsular Ranges Batholith (Clinkenbeard and Walawender, 1989) and from Pioneer Batholith, western North America (Hammarstrom and Zen, 1986; Johnson and Rutherford, 1989a). 4.2.3 BIOTITE Biotite occurs both as a reaction product after hornblende and as euhedral to subhedral grains up to 3 mm long, that could have entirely replaced hornblende. Biotite is commonly grouped into aggregates composed of several biotite grains (e.g., R15908) as well as other minerals (e.g., titanite, Fe-Ti oxide, apatite and zircon). These aggregates commonly occur along the margins of larger grains of hornblende, feldspar and quartz. Biotites from granite (e.g., R15914) and alkali feldspar granite (e.g., R15900) contain zircon inclusions. The inclusions commonly show pleochroic haloes. Biotite exhibits the distinctive pleochroic scheme from X = Y = light brown-chocolate, to Z = straw- coloured, which is typical of oxidised I-type granites (Chappell and White, 1992). 48 Biotite is a common mafic mineral in the Kashmar granitoid (<0.2-17.8% modal) and notably decreases in abundance as the rocks become more felsic. For example, the modal content of biotite from the alkali feldspar granite (Si02 >74 wt%) never exceeds 4.6% by volume of the rock. The absence of significant alteration is supported by the relatively high K20 contents (up to 10 wt%) in the biotite. SHghtly chloritised biotite is localised along obvious zones of fluid alteration defined by fractures whereas light green chlorite is less common in the alkali feldspar granite. Over 150 analyses of fresh biotite were carried out with the electron microprobe. Typical examples are listed in Appendix 3.3. Al™ is calculated as the difference between 8 (total tetrahedral occupancies) and the number of Si cations. Because Fe203 cannot be determined on the electron microprobe, all Fe has been assumed to be FeO in the calculation; and the amount of Fe3+ is considered to be negligible. Biotite grains from all samples except R15900 are homogeneous on the scale of a single thin section. In Figure 4.4 Mg/(Mg + Fe) is plotted versus Mw (a.f.u.) content of biotite. Only three analyses from the alkali feldspar granite (Sample R15900) show a phlogopite composition with Mg/(Mg + Fe) = 0.73-0.77, while the rest of analyses have Mg/(Mg + Fe) values between 0.45 and 0.63. Biotite grains from the granodiorite and granite are relatively titaniferous (Ti02 up to 5 wt%) and high in total Fe as FeO (mostly between 16 to 22 wt%). Some biotite analyses from the alkali feldspar granite (R15900) are lower in total Fe as FeO content (10-15 wt%) but higher in MgO content (up to 18.5 wt%). Such analyses are lower in Ti02 content (1-2 wt%), consistent with the lower Ti02 content of whole rock analyses from the alkali feldspar granite. This 49 difference is supported by microscopic evidence that pale red to green biotite (low Ti02) occurs in the alkali feldspar granite, whereas brown biotite (high Ti02) occurs in the granodiorite and granite. Also, thin sections show titanite and exolution lamellae of ilmenite in magnetite (confirmed by microprobe) in the alkali feldspar granite. Because the Ti content of the melt is buffered by the coexisting titanite or ilmenite (Wyborn, 1983; Harrison, 1990), a lower Ti02 content of biotite from the alkali feldspar granite may be related to fractionation of Ti-rich phases (titanite and ilmenite). Using biotite analyse from numerous localities around the world, Abdel-Rahman (1994; 1996) has shown that igneous biotites crystallising from alkaline (A), peraluminous (P) and calcalkaline (C) orogenic magmas are chemically distinct from one another. Because biotites from the Kashmar granitoid are moderately enriched in MgO content (average Mg/Mg + Fe = 0.57), all plot within the calcalkaline orogenic field (field C) on the AI2O3 versus MgO diagram (Fig. 4.5). In addition they are characterised by a low A1203 content (11.64-14.96 wt%) which is consistent with the absence of Al-rich minerals such as cordierite, garnet and sillimanite in the host-rocks (e.g., Wones, 1980). Al™ is always <2.5 a.f.u. and AT71 is negligible (mostly <0.1 a.f.u., or absent) in the Kashmar biotite. Chappell and White (1992) noted that biotite in granite always contains close to 2.5 atoms of Al™ per 24 (O, OH, F), and that biotite coexisting with hornblende contains negligible amounts of A1VI. In contrast, biotite that coexists with muscovite and/or other Al-rich minerals such as garnet and cordierite, contains appreciable amount of Al^ (-0.6 a.f.u), features attributed to an S-type source (Whalen and Chappell, 1988). According to this scenario, low levels of A1VI in the Kashmar biotite reflect the I-type source composition (Liu et al, 1989; Chappell and White, 50 1992). The overall low levels of Al in the Kashmar biotite are similar to those of the I- type granites reported from the Lucerne pluton, New Brunswick (Wones, 1980), Natanz complex, Iran (Berberian, 1981) and Lachlan Fold Belt, Australia (Wyborn, 1983; Whalen and Chappell, 1988; Turner et al, 1992). The most significant aspect of biotite chemistry is the relationship between Mg/(Mg + Fe) of biotite and the Si02 content of the host-rock (e.g., Czamanske et al, 1981; Whallen and Chappell, 1988; Bacon, 1992). In Figure 4.6, Si02 content in the host- rocks of biotite increases from 62.30 to 76.97 wt% and Mg/(Mg + Fe) ranges from 0.45 to 0.77. Typical positive correlation is observed when analyses of biotite from a single pluton are compared. For example Sample R15908 from granodiorite is low in whole rock Si02 (62.3 wt%) and its biotite grains are low in Mg/(MgO + FeO) (0.47-0.53). Sample R15915 from the same rock is higher in whole rock Si02 (66.4 wt%) and its biotite grains are higher in Mg/(Mg + Fe) (0.54-0.60). This trend is interpreted as indicating melt evolution to more f02, and is supported by presence of euhedral magnetite grains which accompany the biotite (e.g., Mason, 1978) or magnetite inclusions that occur in the biotite (e.g., Whallen and Chappell, 1988). During biotite crystallisation, iff02 is low, iron is preferentially incorporated into biotite (Elliott et al, 1998) but for the Kashmar granitoid iron incorporated into co-existing Fe-Ti oxides, therefore a high/02 was dominant during biotite crystallisation. There is, however, an excellent negative correlation between Fe (a.f.u.) and Mg/(Mg + Fe) values of biotite from the Kashmar granitoid (Fig. 4.7). Three biotite analyses from the alkali feldspar granite plot at the uppermost portion of the trend with Mg/(Mg + Fe) >0.70. These three analyses have a significantly lower content of Ti02, consistent with 51 their lower Al (Appendix 3.3). Such positive correlation between Ti and Al has been related to Al-Tschermak substitution (Deer et al, 1992, 1997; Harrison, 1990; Schneiderman, 1991). Biotites from the granodiorite and granite are more titaniferous than biotites from the alkali feldspar granite. This may be related to octahedral vacancy in which Ti can easily be accommodated in biotite (Schneiderman, 1991). An apparent negative correlation between Mg/(Mg + Fe) and Ti (Fig. 4.8) may confirm the coupling of Ti substitutions with an octahedral vacancy (e.g., Schneiderman, 1991). Based on the data listed in Appendices 3.2 and 3.3, the Mg(Mg + Fe) for hornblende ranges from 0.60 to 0.77 with an average of 0.65. Also, the Mg/(Mg + Fe) for co existing biotite ranges from 0.45 to 0.57 with an average of 0.51. The higher average Mg/(Mg + Fe) for hornblende (Fig. 4.9) is consistent with the high compositional potential of hornblende for generation of biotite under normal magmatic reactions. Also, a lower Mg/(Mg + Fe) in biotite is in agreement with an increase in/02 after hornblende crystallisation (e.g., Clinkenbeard and Walawender, 1989). The average contents of total Fe as FeO and MgO are -16 and 12.81 wt% for hornblende, respectively and 18 and 14.17 wt% for biotite, respectively. The presence of Fe-rich hornblende and Mg-rich biotite in the Kashmar granitoid may reflect the availability of oxygen. Since granitic melts are undersaturated in H20 and the PH20 determines f02 during hornblende crystallisation, due to the diffusive escape of H2 formed by the dissociation of H20, the remaining oxygen would increase f02 (Burkhard, 1991), leading to early crystallisation of Fe-rich hornblende and magnetite (Section 3.2.1.4). This in turn precludes the build up of Fe in the granitic melts and hence, Mg-rich biotite crystallises (Abdel-Rahman, 1994). 52 4.2.4 Fe-Ti OXIDES Fe-Ti oxide is ubiquitous in samples from the Kashmar granitoid and mostly occurs as fine- to medium-grained, subhedral to anhedral granular aggregates. Sometimes Fe-Ti oxide grains contain minor zircon inclusions. Small Fe-Ti oxide inclusions occurring in quartz and plagioclase rims may suggest alteration or later generation of Fe-Ti oxide than that contained in the biotite and hornblende. Fe-Ti oxide commonly replaces hornblende or biotite crystals. Fe-Ti oxide is modally significant, and represents an extensive range from 0.4 to 9.6% by volume of the rock (Appendix 2.1) with increasing abundance when the modal contents of hornblende and biotite increase. Most samples of the Kashmar granitoid contain >1% modal Fe-Ti oxides; this is a fundamental characteristic for distinction of I-type granites (Whalen, 1985; Whalen and Chappell, 1988). The greater opaque mineral abundance in I-type granites is a feature mainly attributable to the higher state of oxidation, particularly for those emplaced at shallow levels of the crust (Burnham and Ohmoto, 1980; London, 1990; Turner et al, 1992; Candela and Blevin, 1995). Microprobe analyses of Fe-Ti oxides are presented in Appendix 3.4. Magnetite, titanomagnetite and ilmenite are present. Among the aforementioned Fe-Ti oxides, magnetite is the predominant opaque mineral occurring through samples of the Kashmar granitoid. Magnetite grains are commonly surrounded by biotite; this is characteristic of I-type granites (Whalen and Chappell, 1988). Ilmenite and titanomagnetite grains are accompanied by magnetite, while magnetite grains can occur independently (e.g., Samples R15908 and R15915). Some magnetite grains are homogeneous (e.g., Samples R15908, R15914). In heterogeneous grains, titanomagnetite is commonly encountered in 53 magnetite cores (e.g., Samples R15918 and R15958). Magnetite grains are generally very low in Ti02 (usually <1 wt%) and have probably re-equilibrated at low temperatures but high,/02. The assemblage magnetite + titanite included in hornblende or biotite indicates they formed above the QFM buffer, a feature for oxidised I-type granites which is thought to be inherited from the source (Wones, 1989; Chappell and White, 1992). Ilmenite grains mostly occur as a 'later phase', forming narrow exolution lamellae (up to 50 microns wide) within magnetite. Ilmenite grains were detected only by electron microprobe. Alteration to rutile + hematite did not take place as Fe and Mn are not depleted. MnO is strongly enriched in ilmenite (4.53-5.27 wt%) relative to coexisting magnetite (MnO = 0.02-0.09 wt%), representing manganon-ilmenite. The Mn enrichment in the ilmenite and its coexistence with virtually homogeneous magnetite suggests low-temperature, high/02 re-equilibration of these phases, possibly with a late fluid phase (e.g., Harrison, 1988). Manganon-ilmenite from the granodiorite (Si02= 62.30 wt%) and alkali feldspar granite (Si02= 76.75 wt%) are compositionally different. Analysis of manganon-ilmenite from the former rock contains appreciable CaO (3.8 wt%) and is lower in FeO content (40 wt%), while manganon-ilmenite from the latter rock contains negligible CaO but higher FeO content (up to 47 wt%). MnO shows the highest concentration in manganon-ilmenite from the alkali feldspar granite (up to 5.23 wt%). These features support a distinct source for the alkali feldspar granite. Overall, ilmenite is not widespread through samples from the Kashmar granitoid. The occurrence of ilmenite in R15900 and R15908 is in accord with their lowest opaque mineral contents (0.8 and 1.0% modal, respectively), indicating a relatively lower y02 54 (e.g., Turner et al, 1992). Like elsewhere, ilmenite from I-type magnetite-granites (e.g., Lachlan Fold Belt, Australia; eastern Peninsular Ranges Batholith; coastal areas of Japanese Islands), ilmenites of the Kashmar granitoid have high MnO and FeO contents. These features, together with a dominance of magnetite grains (Appendix 3.4) rather than ilmenite in the Kashmar granitoid, indicate oxidised magnetite-series or I-type granite. According to Ishihara (1981, 1998) oxidised I-type granites typically occur in the back-arc extensional zone of subduction-related continental margins. Due to lower pressure along these extensional zones, oxidised magmas ascend through the crust without significant amount of crustal contamination (Ishihara, 1998). 4.2.5 K-FELDSPAR In the Kashmar granitoid, K-feldspar is essentially identical in samples containing Si02 >70 wt%. In the granite and alkali feldspar granite, K-feldspar grains occur as prominent coarse-grained, subhedral to anhedral crystals with a pale pink colour in hand specimen. This colour is a field criterion that reflects the higher oxidation states for I-type granites when compared with S-type granites (Chappell and White, 1992). K-feldspar and albite coexist in the granodiorite, granite and alkali feldspar granite. Some K-feldspar grains show alteration to clay minerals, but the majority of grains look light brown or dusty in transmitted light. K-feldspar usually occurs as microperthitic intergrowths. Microperthite commonly shows strings, veins and braided varieties. Microcline is recognised by typical development of cross-hatched twinning characteristic of combined albite and pericline twins, indicating low temperature feldspar. In the alkali feldspar granite K-feldspar typically displays Carlsbad twinning (e.g., Sample R15900 and R15919). These textures indicate slowly changing conditions in the plutonic 55 environment (Shelley, 1993). Variations in exolution are governed by magmatic water, which is the prime catalyst in aiding perthite coarsening (Parsons, 1978; Mason, 1985). Microprobe analyses of K-feldspar, and calculated end-member compositions, are given in Appendix 3.5 and plotted on Figure 4.10. In general, K-feldspar grains from the Kashmar granitoid have a composition of 0^4-94. Based on the feldspar nomenclature of Deer et al. (1992), the compositions of K-feldspar from different plutons of the Kashmar granitoid plot in the fields of perthite and perthitic orthoclase or microcline, indicating a hypersolvus condition of differentiation (Fig. 4.10). The composition of K- feldspar ranges from 0^3.9.92.6, 0^3.9-94.4 and Or The content of CaO is low in K-feldspar from the Kashmar granitoid (mostly An <1 mole% in solid solution). There is an inverse correlation between Na20 (4-0.7 wt%) and K20 (11-16 wt%) contents but the total content of alkalis (Na20 + K20) in the K- feldspar from different plutons is essentially constant (15 to ~16wt%). This may support a similarity in chemical composition of these plutons. The amount of Ti02 (<0.1 wt%) and FeO (<0.4 wt%) are low in all analysed K-feldspar grains. The contents of Si02 (64-66 wt%) and A1203 (18-19 wt%) are constant among K-feldspar analyses from different plutons. 56 Development of microperthitic intergrowths is particularly common in the alkali feldspar granite (R15900 and R15914). The perthite has a uniform narrow lamellae or bleb form. In a single perthitic K-feldspar grain, the lamellae are oriented, while within a thin section preferred orientation of lamellae is not evident. In perthitic grains, compositions with less than Or64 do not occur, suggesting initial plagioclase crystallisation depleted the melt in Ca (Fenn, 1986; Petersen and Lofgren, 1986). The remaining melt, which was enriched in K, crystallised at eutectic temperature then exsolution with the residual plagioclase resulted in composite growth (Hess, 1989). In addition, tectonic strain may have induced unmixing and influence the preferred orientation of perthitic lamellae (Shelley, 1993). Although there is no evidence of ductile deformation in the Kashmar granitoid, the production of perthite, bending of perthite and possibly inversion of perthite to patchy cross-hatched twinning is visible in the alkali feldspar granite (e.g., Sample R15900). Lobate myrmekite is usually situated on the boundary between plagioclase and K- feldspar, and projects into the K-feldspar in a clearly replacive manner. Apparently myrmekite formation was enhanced by late stage lime-bearing magmatic fluids which attack the margins of K-feldspar and release potassium (Shelley, 1993). The released potassium would be fixed in sericite or very fine-grained secondary muscovite that rarely occurs in samples containing myrmekite (e.g., R15914; R15915). 4.2.6 QUARTZ Subhedral and anhedral quartz grains range in size from 0.7 to 1.5 mm. They constitute more than 20% of the rock volume and commonly fill the interstices between the earlier formed minerals, especially in the tonalitic rocks. In the granodiorite and granite, quartz 57 shows corroded margins and undulose extinction. Occasionally it contains inclusions of apatite and other accessory minerals. Quartz has three modes of occurrence: as separate large anhedral grains, as irregular patches, and as microgranophyric intergrowths. In the granodiorite it forms dominantly in separate grains as a consequence of slow near-equilibrium growth at water pressures high enough to restrict solid solution in K-feldspar. However, in the granite, typical granophyric intergrowths developed. Development of microperthitic structure in the granite may be interpreted by a change in PH20. For example, a decrease in PH20 may result by losing vapour through fractures (raising of the liquidus/solidus curves) with a consequential increase in undercooling and rate of crystal growth. Under such conditions independent crystals do not develop, instead the simultaneous growth of quartz and K-feldspar is coupled to produce microgranophyric structure (Tuttle and Bowen, 1958; Philpotts, 1990). 4.2.7 ACCESSORY MINERALS Apatite and titanite are the most common accessory minerals in all rocks of the Kashmar granitoid. Apatite is not uniformly distributed. It occurs in two forms comprising small inclusions in early crystallised minerals (e.g., plagioclase), and as euhedral grains accompanying hornblende, magnetite (e.g., Sample R15922) and plagioclase (e.g., Sample R15911 and R15912). Biotite and plagioclase from the alkali feldspar granite contains apatite inclusions. Association of apatite with plagioclase and ferromagnesian phases suggests that apatite could be incorporated into early crystallised minerals. However, apatite is also the most abundant accessory mineral occurring in microgranular enclaves in the tonalite, granodiorite and granite. Since the microgranular 58 enclaves show a similar mineralogy to the host rocks, and the apatite has a similar morphology to those of the enclaves, it is more likely that apatite may be a restite phase (Chappell et al, 1987; Chen et al, 1990). Titanite commonly occurs as a late stage crystallisation product. Isolated euhedral grains of titanite typically occur in the alkali feldspar granite, while anhedral grains are dominant and occur in the tonalite, granodiorite and granite. Anhedral grains of titanite commonly replace hornblende, Fe-Ti oxides and biotite crystals. Secondary titanite forms narrow rims or blebs around Fe-Ti oxides and hornblende. In the alkali feldspar granite, titanite and ilmenite occur together in the same rock, but the ilmenite is an early crystallising phase as titanite replaces ilmenite. The change from ilmenite to titanite as the titaniferous mineral in the Kashmar granitoid corresponds to an increase mf02 (e.g., Wyborn, 1983; Whalen and Chappell, 1988). Zircon is an ubiquitous accessory mineral in all rock types of the Kashmar granitoid. It is distinctly euhedral and mostly occurs in association with hornblende, biotite and magnetite grains. In particular, it is distributed as inclusions in hornblende and magnetite but is less common in biotite (R15900). When zircon occurs as inclusions in biotite, it shows typical pleochroic haloes (e.g., Sample R15914) indicating radioactive emanations from the inclusion. Euhedral zircon grains from the Kashmar granitoid are unique in shape, and growth zones are not observed. Euhedral isotropic rims are broader at the ends of the crystals and are probably metamict. At high temperatures, when the granitic melt is Zr-saturated, zirconium as a trace element can be accommodated in pyroxene and amphibole (Chen et al., 1990). But 59 zircon grains in the Kashmar granitoid mostly accompany amphibole, indicating they were possibly present in the melt prior to the conversion of pyroxene to amphibole. In this case, zircon grains in the Kashmar granitoid may be residual from the melting of the source (restite). The alternative possible explanation is that the zircon grains resulted from accidental contamination, but other mineralogical and chemical evidence is not consistent with contamination. If zircon grains are inherited from the source, the melts that produced rocks of the Kashmar granitoid would have been saturated in zircon and low in temperature (e.g., Williams, 1992; Chappell et al, 1998). Calculated zircon saturation temperatures (Watson and Harrison, 1983) for the most mafic zircon- saturated I-type granites of the Cobargo suite and Inlet pluton from the Lachlan Fold Belt, Australia, represent maximum temperatures of 740°C to 762°C at which zircon grains could have been present in their melts. According to Chappell et al (1998) such melts form from low-temperature origin. Chemical characteristics of most minerals from the Kashmar granitoid, particularly magnesio-hornblende and feldspars, show that crystallisation occurred at low temperatures, therefore zircon grains would likely be present in the melts that produced rocks of the Kashmar granitoid. 4.2.8 ALTERATION PRODUCTS Plutons of the Kashmar granitoid have been affected by late stage, low temperature subsolidus alteration which occurs extensively along the faults, major joints and contacts. The strongly altered zones occur in the south and southwestern parts of the Kashmar granitoid, particularly to the north of the city of Kashmar. The main mineral assemblages developed are pyrite, sericite, epidote, minor carbonate and titanite. Chlorite alteration is variably present with local pseudomorphs after hornblende and 60 biotite. Replacement by fine-grained accessory muscovite is also evident, the muscovite occurring around the margins and along the biotite cleavages. 4.3 PETROGRAPHY OF BORNAVARD GRANITOID The early magmatic episode in the Bornavard area is characterised by the occurrence of tonalite and granodiorite. Granite represents the late stage intrusive episode which occurs extensively in the Bornavard area. General geology and isotopic ages of these intrusive episodes have been discussed in detail in Sections 3.3.2 and 3.3.2.1. Modal data for intrusive rocks of the Bornavard area are illustrated in Figure 4.11. 4.3.1 TONALITE The tonalite is grey to dark green in colour and occurs in the central part of the Bornavard granitoid. It has a sharp or faulted contact with the granodiorite. Subhedral amphibole grains, anhedral quartz and euhedral to subhedral plagioclase crystals (up to 3 mm long) form the major mineral components of the tonalite (Appendix 2.2). Contents of K-feldspar (3% modal) and plagioclase (40-41% modal) do not vary significantly but quartz ranges from 9.2 to 23.6% modal in samples from the tonalite. This increase in quartz content correlates with a decrease in amphibole (from 45.8 to 29.8% by volume of the rock) which is the only ferromagnesian mineral in samples taken from the tonalite. 4.3.2 GRANODIORITE Granodiorite occurs in the inner parts of the Bornavard granitoid. It is always in sharp contact with the granite and tonalite. Quartz, plagioclase, K-feldspar, biotite and amphibole are major mineral components of the granodiorite. Quartz always shows 61 undulose extinction, indicating tectonic deformation. It contains several inclusions such as apatite, Fe-Ti oxide and titanite (Sample R15947). Dark xenoliths (up to 10 cm across) rich in biotite are common in the granodiorite. In hand specimen, the granodiorite is distinguished from the tonalite by the presence of light milky to pink K- feldspar grains that are surrounded by dark-green, medium-grained accumulations of biotite (e.g., Sample R15946). In thin sections, some K-feldspar grains from the granodiorite are characterised by cross-hatched twinning indicative of microcline (e.g., Sample R15943). The modal content of amphibole in the granodiorite (<0.2-ll% by volume of the rock; Appendix 2.2) is considerably lower than in the tonalite (29.8- 45.8% modal), whereas the modal content of biotite from the granodiorite is high (up to 29.2%) as biotite commonly replaces amphibole in granodiorite (R15953). Quartz and K-feldspar increase in both content and grain size from the tonalite to granodiorite. They display microgranophyric texture in the granodiorite. 4.3.3 GRANITE Granite is coarse- to medium-grained, pinkish in colour and is the late intrusive episode occupying much of the northern and southern parts of the Bornavard granitoid. It mostly occurs in sharp contact around the outer margins of the tonalite and granodiorite plutons. At external contacts, the granite intrudes into very low-grade metavolcanic rocks of the Taknar Zone. Modal data for the granite are shown in Appendix 2.2 and Figure 4.11. Except for one sample (R15939), the granite exhibits a relatively similar modal mineral content that is consistent with the homogeneous features of this pluton. All samples from granite lack amphibole and the content of biotite (<0.2 to 8.4 modal%) is lower than granodiorite 62 and tonalite, consistent with the very felsic nature of granitic pluton. Quartz and K- feldspar occur in typical microgranophyric and graphic textures. Both vermicular and cuneiform types of microgranophyric textures occur, especially in samples having negligible amount of biotite (e.g., R15939). The presence of microgranophyric intergrowths implies rapid undercooling and freezing under low PH20 and low confining pressures (Pitcher, 1993; Shelley, 1993). A perthitic structure is not observed in the K-feldspar of the Bornavard granitoid, as it was for the Kashmar granitoid. This could possibly indicate that the post-crystallisation cooling process was not appropriate in the case of the Bornavard granitoid to break down the feldspar structure to form the perthitic K-feldspar, or the composition was not appropriate. Mineralogical differences between the tonalite, granodiorite and granite suggest that the different plutons of the Bornavard granitoid may not be genetically related. 4.4 MINERAL CHEMISTRY OF BORNAVARD GRANITOID 4.4.1 PLAGIOCLASE Plagioclase forms euhedral to subhedral crystals in the tonalite and granodiorite, but usually anhedral crystals in the granite. It contains common polysynthetic and pericline twins (e.g., Sample R15943). In the granite, plagioclase shows strong resorption along crystal borders (e.g., Sample R15941) and only small anhedral plagioclase grains occur in some samples. Several inclusions such as biotite, Fe-Ti oxides, apatite, and some alteration products including sericite, epidote and muscovite are common in plagioclase crystals from the granodiorite and granite. Most of these inclusions, particularly aluminosilieates, cover a large area in the core of the plagioclase crystals. Normal compositional zoning is common in plagioclase grains from the tonalite and 63 granodiorite. However, owing to changes of temperature (and depth), or composition, the early formed plagioclase crystals have partially resorbed borders. Microprobe analyses of plagioclase are presented in Appendix 3.6 in which the rock series is arranged according to increasing whole rock Si02 contents. The anorthite content of plagioclases in the Bornavard granitoid decreases systematically from mafic to felsic varieties, from A1143.5 to Ani.8. Homogeneous grains predominantly occur in the granite, although variation in the composition of core and rim in plagioclase from the tonalite and granodiorite is usually less than Anio A wider range in plagioclase composition occurs in a normal-zoned grain with a core of An38 and rim of Anig from the granodiorite (Sample R15947). For most plagioclase crystals there is an increase in Na20 and a decrease in CaO with increasing Si02 content from core to rim. Variation in plagioclase composition is relatively wider in the tonalite (Ani3_43) and granodiorite (An2-38) compared with plagioclase in the granite (Ani,8-24)- The plot of plagioclase composition (Fig. 4.12) from all rock types of the Bornavard granitoid shows a range from albite to andesine that is slightly wider than the compositional range of plagioclase from the Kashmar granitoid (Fig. 4.2). 4.4.2 K-FELDSPAR K-feldspar is anhedral, minor and interstitial in the tonalite and granodiorite but it is the most abundant phase in the granite (31.6-52.0% modal). It is partially twinned after Carlsbad or Pericline laws, but commonly shows both in the well-known cross-hatched pattern representing microcline. The presence of abundant microcline in granite suggests equilibrium at a very low temperature in the plutonic environment (Gribble, 1988). This feature is consistent with the quartzofeldspathic nature of the granite. Also, K-feldspar 64 occurs in perfect intergrowths with undulose quartz, producing a typical microgranophyric texture, particularly in samples from the granite (Sample R15939). Some K-feldspar grains are slightly perthitic but development of such a microstructure is not as common as it was observed in the Kashmar granitoid. Microprobe analyses of K-feldspar from the Bornavard granitoid are presented in Appendix 3.7. K-feldspar from the granodiorite and granite have compositions of Or34_83 and Or?3-97 respectively. The small variation in composition of K-feldspar from the granite (Fig. 4.13) is consistent with the homogeneous nature of microcline that shows unique cross-hatched twinning in thin sections. All K-feldspar grains from the Bornavard granitoid are high in Si02 (64.18-66.74 wt%) and A1203 (18.32-19.79 wt%) but low in FeO (0.00-0.07 wt%) and CaO (0.00-0.66 wt%) contents (Appendix 3.7). The amount of Ti02 (0.00-0.03 wt%) and MnO (0.00-0.03 wt%) in the all samples from the granodiorite and granite are very low. All the above major oxides, particularly the contents of Si02 and A1203, are very similar to those of the K-feldspar grains from the Kashmar granitoid. 4.4.3 AMPHIBOLE Amphibole is the only abundant mafic mineral in the tonalite (29.8-45.8% modal) but it is the second most abundant mafic mineral after biotite in the granodiorite and ranges from <0.2 to 11% by volume of this rock (Appendix 2.2). The pleochroic scheme for amphibole is X = Y = pale green to bluish, and Z = pale yellow. Amphibole grains are fresh, occurring as euhedral to subhedral crystals in the granodiorite, but interstitial in the tonalite as a result of rapid cooling of this early intrusion. Amphibole grains from 65 the tonalite and granodiorite have been partially replaced by titanite and biotite, respectively (e.g., R15945; R15953). Chemical data and formulae listed in Appendix 3.8 show that the amphibole grains from the Bornavard granitoid are relatively high in Mg/(Mg + Fe) (0.52-70) and Si contents (6.65-7.78 a.f.u.), characteristic of magnesio-hornblende (Leake, 1978: Deer et al, 1997). Only one rim analysis (R15953) is ferro-Tschermakitic hornblende (Mg/(Mg + Fe) <0.50; Si <6.5 a.f.u.) due to its lowest MgO and Si02 contents. All hornblende analysed are low in Ti (<0.2 a.f.u); their (Na + K)A <0.50 but [Ca(M4) + Na(M4)] >1.34 is characteristic of calcic amphibole (Deer et al, 1997). In comparison with magnesio- hornblende from the Kashmar granitoid, most amphibole analyses from the Bornavard granitoid are lower in MgO/FeO (<0.75), possibly indicating higher f02 for the Bornavard magmas. Except for the ferro-Tschermakitic hornblende rim (R15953), the rest of analyses always contain total Al <2 a.f.u. and Fe3+ is, in most cases, greater than Al^, which is characteristic of calcic amphiboles from shallow-level metaluminous granitoids (e.g., Hammarstrom and Zen, 1986). The CaO contents of hornblende crystals do not show significant variation from the tonalite to granodiorite, but hornblendes become more Si02-rich in the granodiorite (Sample R15943), possibly because they crystallised from a melt higher in Si02 content. This accords with the observation of Cawthorn (1976) that the CaO content of hornblende appears to be insensitive to the CaO content of the magma, while Si02 content is sensitive (Hammarstrom and Zen, 1986). Fractionation factors between amphibole and magma composition for MgO, FeO and Ti02 were determined for hornblende analyses from the Bornavard granitoid (Appendix 3.8). The calculated MgO 66 and FeO fractionation parameters show that both values are mostly >5, suggesting low temperature for hornblende crystallisation. The Ti02 fractionation factor is always less than 2.8, indicating crustal pressure during crystallisation of the Bornavard hornblende. This is in accord with the geobarometer of Johnson and Rutherford (1989a) that suggests low Al content of hornblende (total Al <1.5 a.f.u.) is indicative of pressures less than 3 kbar (Fig. 4.14). A significant difference is observed in the composition of hornblende crystals from different samples of the granodiorite (Samples R15953 and R15943). Most of the hornblende analyses from Sample R15953 are higher in Fe/(Fe + Mg) (0.52-0.62) and 3+ Fe (0.458-1.034 a.f.u.) but lower in Mg/(Mg + Fe) (0.43-0.59) and Si02 contents (41.46-45.15 wt%). Also, hornblende grains in this sample are higher in total Al (1.184- 2.188 a.f.u.), Ti (0.037-0.181 a.f.u.) and Na (0.249-0.381 a.f.u.) compared with hornblendes from Sample R15943. Experimental studies on amphibole compositions as a function of temperature, pressure and f02 have been carried out by several workers (e.g., Hammarstrom and Zen, 1986; Schmidt, 1992; Anderson and Smith, 1995). Most workers report similar results which indicate that, with increasing temperature, Si decreases and Al, Ti and Na increase, while with increasing y02 the reverse is the case. Thus in a magma where f02 increases from early to late stages, early high-temperature amphiboles crystallising at low f02 will be considerably enriched in Al, Ti and Na compared to late stage low-temperature amphiboles crystallising at higher f02. According to this scenario differences in hornblende compositions from the granodiorite may be related to early and late crystallisation of this mineral under different oxidation states. 67 4.4.4 BIOTITE Biotite occurs as medium- to fine-grained (<4 mm in grain size), subhedral to anhedral crystals in the granodiorite and granite. It is fresh, sometimes bent, and strongly pleochroic (from X = Y = dark brown to green and Z = straw coloured) typical of biotite in oxidised I-type granites (Chappell and White, 1992). In the granodiorite, biotite is red to green in colour; it sometimes replaces hornblende and forms the greatest mafic component (up to 29.2 modal%; Appendix 2.2). While in the granite, biotite is less abundant, it is green in colour and occurs as the only ferromagnesian mineral (up to 8.4 modal%). Biotite in the granite has been partially replaced by Fe-Ti oxide (Sample R15941). In particular, biotite becomes less common in the granite when granophyric intergrowths are dominant. Such samples are strongly enriched in total-rock Si02 content (e.g., R15939). Microprobe data for biotites from the Bornavard granitoid are listed in Appendix 3.9 and plotted in Figures 4.15 and 4.16. Biotite analyses from the Bornavard granitoid are lower in Mg/(Mg + Fe) (0.13-0.50) and higher in A1203 contents (15-17 wt%) compared with biotites from the Kashmar granitoid. The composition of biotite in the granodiorite is different from the composition of biotite in the granite (Appendix 3.9). A negative correlation is observed between Mg/(Mg + Fe) and total Fe as FeO in the granodiorite (Fig. 4.16). The average Mg/(Mg + Fe) for biotite in the granodiorite is 0.38 which is significantly higher than for biotite in the granite (Mg/Mg + Fe = 0.15, on average). Low Mg/(Mg + Fe) in biotite from he granite is consistent with the strong enrichment of granite in whole rock Si02 content (74-76 wt%). Biotite crystals from the granodiorite and granite are homogeneous in the composition of core and rim. However, biotite in the granodiorite shows appreciable variation in the composition of some major oxides 68 (e.g., MgO = 5.76 to 11.49wt%). In contrast, biotites from different samples of the granitic pluton are very similar in chemical composition (e.g., MgO = 2.70 to 3.59 wt%). A relatively large variation is observed in the content of total Fe as FeO (17.58 to 27.66 wt%) from biotite in the granodiorite, which is common in most I-type granites (Whalen and Chappell, 1988). Biotite crystals in the granite are strongly enriched in total Fe as FeO content (28.66-32.25 wt%) with an average FeO/MgO = 10.27. This value is much higher than the average FeO/MgO (3.13) values from biotite in the granodiorite. It is also significantly higher than average FeO/MgO (1.45) from biotites in the Kashmar granitoid. Because high/02 conditions allow partition of Fe into oxide rather than biotite (Burkhard, 1991), the Fe enrichment of biotite in granite from the Bornavard area suggests crystallisation at relatively low f02 conditions. However, differences in chemical composition of biotite in the granodiorite and granite suggest that the granite and granodiorite plutons originated from different source compositions. In general, the Bornavard biotites are low in Ti02 (<2.5 wt%). The Ti content of biotite is mainly controlled by temperature and liquid composition; it appears to be particularly insensitive to f02 (Wyborn, 1983). Compared with titaniferous biotite from the Kashmar granitoid, the low Ti02 content of Bornavard biotite implies relatively lower temperature. Also, low Ti02 content in biotite indicates possibly Ti depletion occurred at the source (e.g., Harrison, 1990). Because ilmenite and titanite commonly occur in the tonalite and granodiorite, early fractionation of Ti-rich phases, together with low temperature, may be responsible for the low Ti02 content of biotite in the Bornavard granitoid. 69 According to the statistics on mica compositions for S- and I-type granites in southeast China (Liu et al, 1989), the amount of AT71 in biotite is more than 0.5 for S-type granites and less than 0.5 for I-type granites. The biotite from the Bornavard granitoid has a relatively high amount of A1VI in the octahedral layer, usually more than 0.5, ranging from 0.43 to 0.89, averaging 0.59 (n = 24). This average is lower than 0.6 a.f.u, a value for biotite occurring with muscovite (Chappell and White, 1992). The high amount of A1VI in the Bornavard biotite is not consistent with petrographic and mineralogical evidence that suggest an I-type source for the Bornavard granitoid. It is in agreement with the high FeO and low MgO contents of these biotites that indicate, in addition to the substitution of Fe for Mg, the substitution of kf1 for Mg has occurred (e.g., Liu et al, 1989). Biotite in the Bornavard granitoid exhibits a great range in total Fe as FeO content that is characteristic of biotite from I-type granites (Whalen and Chappell, 1988). Because biotite in the granite coexists with secondary muscovite, high A1VI in the Bornavard biotite may be the result of subsolidus alteration (e.g., Harrison, 1990). 4.4.5 ACCESSORY MINERALS Fe-Ti oxide, titanite, allanite, apatite and zircon are accessory minerals occurring in rocks of the Bornavard granitoid. They mostly accompany hornblende and biotite. The Fe-Ti oxides and titanite are products of magmatic reactions but allanite may be the result of hydrothermal alteration of biotite in granite samples. Ilmenite is the common Fe-Ti oxide, coexisting with hornblende (R15945; R15953) in the tonalite and granodiorite that are the early intrusive rocks of the Bornavard granitoid. Some ilmenite grains have rounded edges, and ilmenite alteration is evidenced by narrow reaction rims of fine-grained titanite along all rims, fractures and cracks. Such altered ilmenite can 70 result from magmatic evolution that is common in I-type granites, and indicates relatively oxidising conditions (Whalen and Chappell, 1988; Petrik and Broska, 1994). Microprobe data for Fe-Ti oxides from the Bornavard granitoid are listed in Appendix 3.10. Homogeneous ilmenite (predominant) and magnetite occur in the tonalite and granodiorite, while magnetite is the only Fe-Ti oxide typically occurring in the granite. Some of the magnetite grains in the granite are heterogeneous, having titanomagnetite (Ti02 up to 7.47-9.30 wt%) in the (R15938; R15941). The ilmenite is slightly high in MnO content (2.40-3.18 wt%) but very low in AJ2O3, MgO and other major oxides. The content of Ti02 (up to 52.45 wt%) is much higher than FeO (up to 44.66 wt%) in ilmenite. With increasing content of Si02 in the tonalite and granodiorite (from 58.09 to 71.32 wt%), significant variation in MnO and total Fe as FeO contents of ilmenite is not observed. This indicates that the composition of ilmenite is independent of variation in silica content of magma, but is related to the/02 (e.g., Czamanske et al, 1981; Petrik and Broska, 1994). Compared with ilmenite grains from the Kashmar granitoid (Appendix 3.4), ilmenite from the Bornavard granitoid is slightly higher in Ti02 and lower in MnO and FeO contents. These features may suggest slightly lower f02 conditions for ilmenite of the Bornavard granitoid. However, ilmenite from both areas are compositionally similar to ilmenite analyses from I-type granites of Central Chugoku, Japan (Czamanske et al, 1981) and the Lachlan Fold Belt, Australia (e.g., Wyborn, 1983; Whalen and Chappell, 1988). 71 Magnetite is absent in the tonalite, it coexists with ilmenite in the granodiorite, whereas it occurs without ilmenite in the granite. After biotite, magnetite is the second most common mafic mineral occurring in the granite but both minerals are less abundant than in the tonalite and granodiorite. The absence of ilmenite in granite is consistent with the higher total Fe as FeO contents and lower Mg/(Mg + Fe) values of biotite in granite samples (Fig. 4.16); all are indications of progressive increase mf02. Therefore, the nature of the source compositions, together with differences in f02 and PH20, are essential to explain compositional variation of Fe-Ti oxide minerals for different plutons of the Bornavard granitoid. In the granite, the presence of titanomagnetite cores and magnetite rims (R15938; R15941) indicates Ti depletion occurred in magnetite grains. Such grains are commonly accompanied by biotite and titanite. The presence of reaction rims indicate that the loss of Ti in magnetite occurred, either through exchange with biotite or by the formation of titanite from magnetite by oxidising deuteric fluids (e.g., Petrik and Broska, 1994). Lowering the Ti02 content of magnetite at high oxidation state, particularly in I-type granites, may favour the presence of allanite (Hine et al., 1978; Petrik and Broska, 1994) which is observed only in samples from the granite. Allanite in the granite (Appendix 3.10; Sample R15940) is subhedral to anhedral, homogeneous, reddish in colour, slightly anisotropic, low in birefringence, and has anastomosing cracks that extend to the boundary of adjacent minerals; an indication of radioactivity. Allanite grains are high in Ti02 and total Fe as FeO (up to 19 wt%). High FeO content of allanite is consistent with coexisting homogeneous magnetite. It is suggested that the entry of the relatively large amount of Fe and Ti into the allanite 72 structure is favoured by lower pressures (Deer et al, 1997). The sum of the major oxides in allanite is low (70.13 wt%) because REE oxides were not analysed. Allanite and monazite are included in contrasting mineralogies of both I- and S-type granites by Chappell and White (1974) and later classifications (Hine et al, 1978; Petrik and Broska, 1994). The presence of allanite in the Bornavard granite is considered to be a strong indicator of the oxidised I-type feature (Sawka et al, 1986; Sawka and Chappell, 1988). Burnham and Ohmoto (1980) and Wones (1980, 1989) suggested that magnetite- granites formed at higher relative f02 than ilmenite-granites, with the QFM buffer as the fundamental division between the two. Chappell and White (1992) interpreted thatjG2 is an inherited feature from the source for I- and S-type granites. In the early intrusive rocks of the Bornavard granitoid, the assemblage titanite + ilmenite + magnetite + quartz occurs with hornblende grains (e.g., Sample R15953). The hornblende grains are relatively high in Mg/(Mg + Fe) (0.52-0.70), therefore, a reduced condition such as for S-type granites is not implied (Wones, 1989). Accordingly, the occurrence of allanite + magnetite + biotite instead of ilmenite + hornblende assemblage in the late intrusive rocks (granite) implies higher J02 (e.g., Petrik and Broska, 1994). 4.4.6 ALTERATION PRODUCTS In the Bornavard granitoid, common alteration products comprise sericite (up to 4.4% by volume of the rock), epidote, chlorite and calcite. Some plagioclase grains are intensely sericitised, particularly in samples from the granite. The biotite is replaced by magnetite and sericite. Epidotisation of biotite and K-feldspar is common in the granodiorite and granite. Secondary muscovite is fine-grained, commonly occurs 73 as interstitial crystals, as well as in small dispersed flakes within plagioclase. Secondary muscovite becomes euhedral and more abundant when magnetite has a higher abundance (e.g., R15938) and granophyric intergrowths are well developed (e.g., R15955). The presence of small flakes of muscovite within plagioclase may suggest leaching of K and Si from the plagioclase at temperatures below the granite liquidus. Epidote is a constituent of the late kinematic granites; its abundance is related to the calcium content of the host granite (Deer et al, 1997). The amount of epidote increases as the calcium content of the host rock decreases. In some samples from the granite, the abundance of epidote and secondary muscovite resulted in lowering the whole rock CaO content, consequently those samples show a slightly higher alurriinium saturation index (e.g., Sample R15955). This scenario will be discussed in detail in Chapter 6. 4.5 TAKNAR RHYOLITE The Taknar Rhyolite is part of the metavolcanic rocks of the Taknar Zone (Razzaghmanesh, 1968). The Taknar Rhyolite is partly intruded by the Bornavard granitoid, indicating a relatively younger age for the granitoid. The metavolcanic rocks of the Taknar Zone are petrographically compared with the Precambrian basement sequences in Iran (Crawford, 1977; Berberian and King, 1981; Hamedi, 1995). The Taknar Rhyolite is widely distributed around the southwestern and northeastern parts of the Bornavard granitoid (Fig. 3.4). The rhyolite is a light grey to slightly green in hand specimen. K-feldspar, quartz, plagioclase, magnetite and rare biotite commonly occur in samples from the Taknar Rhyolite. Phenocrysts of the first three minerals are set in a holocrystalline fine-grained groundmass. The Taknar Rhyolite has been affected by very low-grade metamorphism and hydrothermal alteration (Crawford, 1977; Muller and 74 Walter, 1983). The low-grade metamorphism in the rhyolite is distinguished by slight recrystallisation of quartz, particularly in the groundmass (e.g., Sample R15948). Common alteration products include sericite and chlorite. 4.5.1 PETROGRAPHY AND MINERAL CHEMISTRY Petrographic data and chemical classification for the Taknar Rhyolite are presented in Appendix 2.3 and Figure 4.17, respectively. Five samples were chemically analysed from the Taknar Rhyolite. Based on Si02 vs (Na20 + K20) contents (Le Maitre, 1989; Le Bas and Streckeisen, 1991), the rhyolites fall within the alkali rhyolite field (Fig. 4.17). Only one sample (R15949) is lower in total alkali contents (Na20 + K20 = 3.69 wt%) due to the extensive sericitisation of K-feldspar in the groundmass. Microprobe analyses show that plagioclase phenocrysts in the Taknar Rhyolite have an albite composition (Ab99, Appendix 3.11). The plagioclase phenocrysts occur mostly as subhedral grains up to 3.5 mm in size and some have common pericline twinning. Some plagioclase phenocrysts have been partially altered to sericite. K-feldspar occurs both in the groundmass and as small phenocrysts having simple twinning. After quartz, it is the second most common mineral component of the groundmass. The composition of K-feldspar from the Taknar Rhyolite (Appendix 3.11) indicates homogeneous sanidine (Ofyi.z) as the phenocrystic phase and sanidine- anorthoclase (Or22.8o.6) in the groundmass. The contents of Ti02 (<0.04 wt%), MgO (<0.02 wt%), MnO (<0.05 wt%) and FeO (<0.3 wt%) in the all feldspars analysed from the Taknar Rhyolite are very low (Appendix 3.11). 7 b Biotite with Mg/(Mg + Fe) = 0.66-0.68, magnetite and titanomagnetite are the only ferromagnesian minerals occurring in the Taknar Rhyolite. Biotite is homogeneous, relatively low in K20 content (7-8 wt%) that may be the result of partial alteration of biotite to chlorite. The chemical data that are summarised in Appendices 3.9 and 3.11 show that the compositions of biotite from the Bornavard granitoid and Taknar Rhyolite are different. In particular, biotite from Taknar Rhyolite is low in Ti02, FeO/MgO and K20, but high in Si02, AI2O3 and MgO contents. Such differences in chemical composition of minerals suggest different sources for the Taknar Rhyolite and the Bornavard granitoid. Fe-Ti oxide is abundant (up to 8.2% modal) in the Taknar Rhyolite. Equant fine-grained Fe-Ti oxide grains are scattered through the groundmass (e.g., R15950, R15951). These grains are anhedral and commonly accompany alteration products. Analyses of Fe-Ti oxide from the Taknar Rhyolite (Appendix 3.11) indicate a heterogeneous composition, with titanomagnetite cores and magnetite rims. The core composition is strongly enriched in Ti02 (18.6 wt%) and MnO (4.5 wt%) contents, while the rim composition is depleted in these oxides. This trend indicates progressive oxidation during Fe-Ti oxide crystallisation. Overall, the composition of Fe-Ti oxide from the Taknar Rhyolite is different from those of in the Bornavard granitoid. Quartz is a minor phenocrystic phase, it occurs both as rounded or anhedral phenocrysts and in the groundmass. It is embayed and highly strained, as manifest by strong undulose extinction to extensive subgrain development, and most often shows recrystallisation with local granulation (e.g., R15951). It sometimes occurs as microgranophyric intergrowths. Secondary muscovite, biotite and Fe-Ti oxides are other 76 constituents occurring in the groundmass. Zircon, titanite and apatite are the accessory mineral assemblage occurring in the Taknar Rhyolite. In the Taknar Rhyolite sericitisation of feldspars is mostly accompanied by chloritisation of biotite (e.g., Sample R15949). Sericitisation requires the addition of water-rich fluids and K+ (Shelley, 1993). The source of K+ can be found in the chloritisation of biotite + which is evidenced from the low K20 content of the biotite (Appendix 3.11); the K released reacts with plagioclase to free Ca2+ (Deer et al., 1992). Due to the relative immobility of Si and Al (Rollinson, 1993), the host-rock became relatively poor in CaO content (CaO = 0.13-0.35 wt%) with a consequential increase in the calculated Aluminium Saturation Indices (ASI), showing a strong peraluminous character (ASI > 1.1) for the Taknar Rhyolite. These criteria will be discussed in detail in Chapter 6. 4.6 KUH MISH INTRUSIONS The Kuh Mish intrusions (Fig. 4.18) are Middle-Late Eocene in age (Sahandi, 1989) and cover a total area of approximately 84 km2. They intrude into Late Cretaceous andesitic and basaltic rocks, as well as into volcano-sedimentary rocks of Early Eocene age (Eftekhar-Nezhad, 1976). These intrusions are mainly granodiorite in composition. In the Kuh Mish area, the granodiorite has been intruded by medium-grained rocks of mostly gabbroic and quartz monzodioritic composition. Two isolated granodiorite plutons occur in the Darin and Namin regions, both located in the northwestern parts of the Kuh Mish area (Fig. 4.18). Representative modal analyses for the Kuh Mish intrusions are shown in Appendix 2.4 and Figure 4.19. 77 4.6.1 GABBRO Gabbro (Si02 = 45.75 wt%) and quartz monzodiorite (Si02 = 51.85-63.93 wt%) comprise the only mafic rocks analysed from the Kuh Mish intrusions, m the Kuh Mish area, gabbro intrudes in dyke-form in the central part of the quartz monzodiorite; while the latter intrusion crops out widely around, and into the granodiorite pluton. Quartz monzodiorite is the finest grained rock (minerals, 2-3 mm in size) among these intrusions. Its contact with the surrounding rocks, as well as into the granodiorite, is clearly intrusive. 4.6.1.1 Mineralogy of Gabbro Clinopyroxene grains form 44.6% by volume of the gabbro (Appendix 2.4). They are simply-twinned, yellowish orange in colour, and show exsolution lamellae of hedenbergite. Sometimes clinopyroxene crystals contain pale yellowish cores of relict olivine, which are altered. Microprobe data (Appendix 3.12) show that diopside (E1145- 47-W044-49-FS6-7) is the dominant clinopyroxene present in the gabbro (Fig. 4.20). Only one rim is slightly lower in CaO (14.5 wt%) and higher in MgO contents (18.2 wt%), representing augite in composition (En57.7-Wo32.9-Fs9.s), which is a common constituent of gabbros. However, the diopside grains are optically homogeneous. They are depleted in Ti02 content. According to Deer et al (1992), formation of augite as a low-calcium clinopyroxene in the rim composition or in the groundmass is due to rapid cooling. Plagioclase grains form 53.6% by volume of the gabbro (Appendix 2.4). Euhedral crystals of plagioclase together with diopside grains display a hypidiomorphic granular texture in the gabbro. The composition of the plagioclase in gabbro ranges from Ango to 78 An98.8 (Appendix 3.13), representing the most calcic plagioclase grains that were analysed in the present study (Fig. 4.21). The plagioclase grains are homogeneous, high in Al203 (up to 35.8 wt%) and low in Si02 (up to 45 wt%) contents. They are compositionalry different from plagioclase crystals of the Kashmar and Bornavard granitoids, representing a distinct source for gabbro from the Kuh Mish intrusions. Minor mineral constituents in the gabbro include alteration products such as Fe-Ti oxide, sericite, tremolite, actinolite and epidote. Microscopic evidence shows possibly two stages of alteration for the gabbro. At higher temperatures plagioclase changed to epidote, as well as pyroxene changed to fine-grained fibrous tremolite and actinolite. Then at lower temperatures, probably with an enrichment in H20 content (magmatic or meteoritic), sericitisation developed and became dominant as the last secondary mineral phase in gabbro (Sample Rl5929). 4.6.2 QUARTZ MONZODIORITE The quartz monzodiorite is grey to green in colour, hi the Kuh Mish locality, quartz monzodiorite partly intrudes into the granodiorite in east-west trending parallel dykes. The quartz monzodiorite seems to be a very high level intrusion as evidenced from its medium-grained minerals, particularly near its contact with the granodiorite. Towards the contact with the granodiorite, chlorite alteration has locally been developed. There is no evidence for xenoliths derived from the granodiorite or the adjacent volcanic rocks. 4.6.2.1 Mineralogy of Quartz monzodiorite The modal mineral assemblage for the quartz monzodiorite is shown in Appendix 2.4 and Figure 4.19. Plagioclase (36-50 wt%), amphibole (10-33.6 wt%), K-feldspar (9- /y 19 wt%) and quartz (9-14.6 wt%) are the major mineral components in the quartz monzodiorite. Minor diopside occurs as anhedral crystals, most of them have been replaced by amphibole due to normal magmatic reactions. In some hornblende crystals diopside relicts are observed. Biotite is absent since after amphibole crystallisation, or possibly during its evolution, the melt encountered hydrothermal fluids that caused Fe- Ti oxide, chlorite, epidote and titanite to commonly replace amphibole grains. Apatite is present as prisms up to 0.8 mm long embedded in late crystallising minerals and as intergranular crystals. Plagioclase occurs mostly as laths (up to 3 mm long) in samples taken near the contacts. In other places, it is euhedral and normally zoned, or anhedral with strongly resorbed boundaries. Alteration of the plagioclase cores is ubiquitous, varying in intensity from slight to extensive sericitisation. K-feldspar occurs as anhedral grains or interstitial fillings. In some samples (e.g., R15956) it is subhedral and shows simple twinning. Amphibole and Fe-Ti oxide are the most common ferromagnesian phases in the quartz monzodiorite. In general, amphibole and Fe-Ti oxides decrease as the host-rock Si02 content increases. Amphibole is anhedral in fine-grained samples with a lower Si02 content, but subhedral in other samples. The amphibole is slightly pleochroic from X = Y = pale green, to Z = light-yellow. Actinolite and tremolite after amphibole are common in quartz monzodiorite. 4.6.3 GRANODIORITE Three separated homogeneous plutons of granodiorite composition occur in the Kuh Mish (28 km2), Darin (21 km2) and Namin (5 km2) localities (Fig. 4.18). Rocks of the Kuh Mish and Darin plutons are mineralogically similar, which may imply similarity of the source compositions. They are white on fresh surfaces, low in mafic minerals and lacking in biotite. In the Kuh Mish and Darin plutons, amphibole shows partial 80 alteration to chlorite. In the Kuh Mish pluton the intensity of alteration in some amphibole grains resulted in complete pseudomorphism of the amphibole. Samples from the Namin pluton are black to grey in colour due to the presence of abundant amphibole, biotite and Fe-Ti oxides. Granodiorite from the Namin pluton (Sample R15926) contains a total of 23.4% modal contents of biotite, amphibole and Fe-Ti oxide. Petrographic data for eight granodiorite samples from these plutons are summarised in Appendix 2.4 and Figure 4.19. 4.6.3.1 Mineralogy of Granodiorite Plagioclase is the dominant mineral in the granodiorite samples from Kuh Mish, Darin and Namin plutons; it is typically euhedral and commonly shows polysynthetic and pericline twinnings. Most plagioclase grains are normally zoned, but some have resorbed rims. The grains are variably altered to epidote and sericite; the latter has partly proceeded to form small crystals of secondary muscovite. Chemical data for plagioclase crystals in granodiorite from the Namin (R15926) and Darin (R15927) plutons are shown in Appendix 3.13. The composition of plagioclase ranges from An26-57 and Ann-54 in the Namin and Darin plutons, respectively. However, most plagioclase grains from Namin and Darin plutons are compositionally overlapped, that indicates these plutons may be genetically related. Similar to plagioclase composition from the Kashmar granitoid (Fig. 4.2), most plagioclase grains from the Namin and Darin plutons represent an intermediate composition (Fig. 4.21) which is typically recognised in low-temperature granitoids. The presence of some normally- zoned plagioclase crystals in these plutons suggests that mineralogical equilibration was much slower than the rate of crystallisation. This is characteristic of low-temperature 81 systems in which plagioclase takes a very long time to equilibrate with the melt because that would involve a change in the Al/Si (e.g., Pearce and Kolisnik, 1990; Shelley, 1993). Subhedral amphibole typically occurs in the Namin and Darin plutons. Amphibole grains are up to 2.5 mm long and form up to 14.8% by volume of the rock. The pleochroic scheme is generally from X = Y = pale green/brown to Z = light yellowish. In Namin pluton, amphibole grains are partially replaced by fine-grained biotite, chlorite, irregular grains of Fe-Ti oxide and titanite. Replacement is common along the cleavages or grain boundaries. Actinolite occurs where moderate to complete pseudomorphs of amphibole are visible. Microprobe analyses of amphibole grains from the Namin and Darin plutons are listed in Appendix 3.14, For all analyses Ca(M4) is more than 1.5 a.f.u. and Na(M4) is less than 0.50 a.f.u., indicating calcic amphiboles. They are high in Mg/(Mg + Fe) (>0.50) and Si content (Si = ~7 a.f.u.), but their (Na + K)A and Ti are both less than 0.50 a.f.u. These characteristics represent typical magnesio-hornblende (Deer et al, 1997) for the Namin and Darin plutons. Magnesio-hornblendes are high in MgO/FeO, in contrast the Fe/(Fe + Mg) is commonly less than 0.50 indicating magmatic amphibole (Harrison et al, 1990). These distinctive properties may be a result of a low FeO content in the source or a lower oxidation state for amphiboles from the Namin and Darin plutons. Magnesio-hornblendes from the Namin and Darin plutons have a higher Mg/(Mg + Fe) (up to 0.83) compared with analyses of magnesio-hornblende from the Kashmar and Bornavard granitoids (Appendices 3.2 and 3.8). The high Mg/(Mg + Fe) of magnesio- 82 hornblende from the Namin and Darin plutons suggests that these plutons originated from sources being more mafic in nature. They are low in total Al (<1.5 a.f.u.) indicating a low pressure in the source. Fractionation factors for MgO (5-10), FeO (2-5) and Ti02 (1-4) for most magnesio-hornblende analyses imply low-temperature and relatively low crustal pressures, as suggested for the Kashmar and Bornavard granitoids. Biotite is fine-grained, it occurs in the Namin pluton (R15926), while other plutons of the Kuh Mish area lack this mineral because alteration, together with a high state of oxidation, were dominant during or after amphibole crystallisation. Under these conditions, possibly hornblende would have directly transfered to opaque minerals and chlorite instead of being replaced by biotite (e.g., Shelley, 1993). Analyses of biotite grains from the Namin pluton are high in MgO (up to 13.6 wt%) and low in total Fe as FeO (19-21 wt%) contents, with a high Mg/(Mg + Fe) (0.50-0.54), which is in accord with the high Mg/(Mg + Fe) of coexisting magnesio-hornblendes. Biotite crystals from the Namin pluton (Appendix 3.15) are compositionally different from biotites in the Bornavard granitoid (Appendix 3.9), while they are similar in composition to those of the Kashmar granitoid (Appendix 3.3). Analyses of Fe-Ti oxide from the Darin pluton represent heterogeneous grains, including magnetite and titanomagnetite compositions (Appendix 3.15). Titanomagnetite grains contain notable Si02 (up to 5.64 wt%) and CaO (up to 4.58 wt%) contents but they are low in Ti02 content (up to 5.67 wt%). One rim is depleted in these components, being magnetite in composition, that possibly means an increase in f02 during crystallisation of Fe-Ti oxides encouraged precipitation of magnetite. 83 4.7 SUMMARY Apparently, all ferromagnesian minerals in magmas of the northeastern CIP crystallised at an initially high f02 and PH20 that possibly increased during the final stages of crystallisation in the relevant pluton. The coexisting biotite, magnesio-hornblende and magnetite suggests a direct relationship between f02 and PH20 (Mason, 1978; Burkhard, 1991), the two fundamental criteria important in determining the characteristics and generation of mineralised I- and S-type granites (Blevin and Chappell, 1992; Blevin et al, 1996). The availability of H20 and equilibrium conditions (e.g., Burkhard, 1991) are responsible for homogeneous crystallisation of hornblende and biotite grains in different plutons of the northeastern CJP. When biotite and magnesio-hornblende coexist, the former replaces the latter, indicating normal magmatic reactions. Biotite crystals are typically pleochroic from dark brown to straw-coloured, and they are high in FeO/MgO representing oxidised I-type magmas. The composition of biotite in the Bornavard granitoid is multi-modal, representing three different sources for episodic intrusions of the Bornavard granitoid. Biotite and magnesio-hornblende from the Kuh Mish intrusions are high in Mg/(Mg + Fe) values that implies magmas of the Kuh Mish intrusions originated from sources being more mafic in nature. Except for two analyses, structural formulae of the analysed amphibole from northeastern CJP show a total Al <1.5 a.f.u. that is characteristic of shallow-level intrusions (Hammarstrom and Zen, 1986; Hollister et al, 1987). Fe-Ti oxide occurs as magnetite, titanomagnetite and ilmenite grains, but magnetite is the dominant opaque phase, indicating high f02. The Fe-Ti oxides are usually accompanied by biotite, titanite and amphibole. Where titanomagnetite occurs, it is b4 commonly encountered in magnetite cores. In the Bornavard granitoid, ilmenite and amphibole specifically are found in rocks of the early intrusive episode, while magnetite and biotite are present as the only ferromagnesian phases in rocks of the late intrusive episode. Plagioclase crystals from the Kashmar and Bornavard granitoids are commonly lower than An5o in composition. For the Kuh Mish intrusions plagioclases in the Namin and Darin plutons are compositionally lower than An57. In each area of this study, plagioclase from a certain pluton shows an evolutionary trend towards enrichment in albite content by progressive increase in whole rock Si02 content, that is characteristic of fractional crystallisation. Plagioclase crystals are homogeneous or variably normally- zoned which is a response to the nature of plagioclase crystallisation during changes in temperature and fluctuation in PH20. Analyses of feldspar grains from each pluton record evidence of low-temperature crystallisation. CHAPTER 5 WHOLE ROCK GEOCHEMISTRY 5.1 INTRODUCTION In the present study, 60 representative samples of igneous rocks from the Kashmar, Bornavard and Kuh Mish areas have been analysed for major and trace elements (Appendices 4.1 to 4.4). Major elements were determined by X-ray fluorescence spectrometry (XRF), while trace elements were analysed by a combination of XRF and instrumental neutron activation analysis (INAA). Tantalum, Co and W are not reported because of contamination of the sample by crushing in a tungsten-carbide mill (Tema). The loss on ignition (LOT) (Appendix 1) for all samples in the present study is less than 3.5 wt%, with most less than 1 wt%, supporting the contention that samples are fresh and represent magmatic compositions. The CIPW normative mineralogy (Appendix 2) was calculated using the computer program Geochemical Data Analysis (GDA) of Sheraton and Simons (1991). Since most of the I-type granites (e.g., Lachlan Fold Belt) maintain a ratio of 2:1 for FeO/Fe203, this ratio was assumed for calculation of CIPW normative mineralogy. 5.2 KASHMAR GRANITOID 5.2.1 MAJOR ELEMENTS Based on the silica content, all samples from the tonalite (Si02 = 54.18-59.79 wt%) and one sample (R15908) from the granodiorite are mafic in composition, whereas other samples are felsic in composition (Si02 > 63 wt%). The widest range in silica content is observed in the granite (Si02 = 63.42-71.81 wt%) but other plutons of the Kashmar granitoid are relatively uniform in composition and show small variation (-5 wt%) in Si02 contents, which is characteristic of I-type granites (Chappell and Stephens, 1988). Wide variation in Si02 content is usually interpreted as fractional crystallisation (Azevedo and Nolan, 1998). This process is observed in the granitic pluton by a general decrease in plagioclase and increase in K-feldspar contents towards more silica variants (Appendix 2.1). Overall, the Si02 content in rocks of the Kashmar granitoid ranges from 54.18 to 77.06 wt% with an average of 66.80 wt% (29 analyses). In the classification of I-type granites, this average content of Si02 would typically be I-(granodioritic) type (e.g., Hill et al, 1988, 1992; Pharaoh et al, 1993; Waight et al, 1998; Ajaji et al, 1998). Ti02 is typically low, mainly <0.5 wt% and all <0.92 wt%. The highest content of Ti02 (0.91 wt%) occurs in the tonalite, whilst the lowest content is in the alkali feldspar granite (Fig. 5.1). High Ti02 content in the tonalite is consistent with its higher modal contents of biotite, amphibole and Fe-Ti oxides. In the alkali feldspar granite, biotite and ilmenite are the only Ti-rich minerals. Although they occur in small amounts, microprobe analyses (Appendices 3.2 to 3.4) show that biotite (Ti02 up to 4.90 wt%) and ilmenite (Ti02 up to 53.54 wt%) are richer in Ti02 content than the hornblende (Ti02 up to 2.37 wt%). It can be concluded that the lower Ti02 content of the alkali feldspar granite may be related to fractionation of biotite and ilmenite. The content of A1203 is low in the alkali feldspar granite (11.70-12.73 wt%) but high in the tonalite (16.04-17.04 wt%), granodiorite (15.22-16.20 wt%) and granite (13.61- 15.89 wt%) as they contain high plagioclase contents. Referring to Appendices 3.1 and 3.2, both plagioclase and hornblende are enriched in A1203 contents (up to 27.72 wt% S7 and up to 8.41 wt%, respectively), thus explaining the higher Al203 content of the granodiorite and granite. The same explanation can be attributed to the tonalite that contains higher modal contents of plagioclase and hornblende. The content of Al203 in the granite is reasonably wide because some samples from this pluton are relatively enriched in Al203 by the presence of plagioclase, hornblende and biotite, whereas others are low in such components and high in K-feldspar and quartz contents. This resulted in a good linear trend on the plot of Al203 versus Si02 that is observed from the granite samples (Fig. 5.1). Total Fe as Fe203 content is high (mostly between 3.5 to 9.5 wt%) and is consistent with high/02 conditions, suggested by chemical data for ferromagnesian minerals (Chapter 4). The content of total iron as Fe203 shows a negative correlation with Si02 (Fig. 5.1) that is related to the decreasing modal content of Fe-Ti oxide and other ferromagnesian minerals from the tonalite to the alkali feldspar granite. The content of MnO is low and decreases from 0.14 wt% in the tonalite to 0.01 wt% in the alkali feldspar granite (Appendix 4.1). A slight scattering is observed in the granite analyses (Fig. 5.1) that is the result of differences in modal contents of Fe-Ti oxides, hornblende and biotite, common minerals that accommodate variable Mn in their crystal structures (Deer et al, 1997). Since Mn+2 has an ionic radius (0.80 A0) near to Fe+2 (0.76 A0), it can substitute for Fe+2 in ferromagnesian minerals (Harrison, 1990). This substitution seems to be higher in ilmenite (MnO up to 5.27 wt%) than in hornblende (MnO up to 0.88 wt%). Therefore, a low content of MnO, particularly in samples from the alkali feldspar granite, is related to fractionation of ilmenite. MgO is relatively low and mostly ranges from 2.97 to 0.43 wt%. It shows an excellent negative correlation with Si02 throughout different samples of the Kashmar granitoid. The highest MgO content (4.27 wt%) is observed in Sample R15924 (tonalite) as a result of its high modal content of hornblende (20.4% by volume of the rock). For the granodiorite and granite, a lower content of MgO is consistent with their lower modal contents of hornblende. The lowest content of MgO (0.09-0.30 wt%) is observed in the alkali feldspar granite because this rock lacks hornblende. Therefore, variation in the contents of MgO from different rock types of the Kashmar granitoid is mainly related to differences in modal contents of hornblende. CaO shows the same compositional trend as for MgO. The content of CaO decreases from 7.6 to 5 wt% in the tonalite, to 4.5 to 2 wt% in the granodiorite and 4.7 to 1.3 wt% in the granite, with <1 wt% in the alkali feldspar granite. For each pluton, the decrease in CaO content from rocks of low Si02 to high Si02 content would be the result of crystal fractionation of plagioclase and calcic amphiboles. Two samples of tonalite (R15911 and R15924) at the same silica level (Si02 = -54 wt%) plot away from the general trend, however, one of them (R15924) which has a lower CaO content (4.18 wt%), contains the highest modal content of plagioclase (51.6%). The contrast in behaviour of CaO in Sample R15924 may be related to secondary processes owing to mobile nature of CaO. Na20 is high and is dominated by values between 3 and 5 wt%. The high content of Na20 means that Na has not been removed from the source rocks, therefore, these plutons are genetically I-type (e.g., Kleemann and Twist, 1989; Chappell and White, 1992; Pitcher, 1993; Raymond, 1995). The content of Na20 is never less than 2.6 wt% and is only slightly <3 wt% in three samples from the alkali feldspar granite (R15916, R15900, R15914). A relatively low content of Na20 in the alkali feldspar granite is in accord with its lower content of plagioclase and absence of hornblende, whereas the presence of abundant K-feldspar enriched the rock in K20 content (up to 5.88 wt%). This resulted in a higher K20/Na20 in the alkali feldspar granite, although this ratio is mostly low in other rock types (Appendix 4.1). Overall, the low K20/Na20 ratio supports an I-type granodioritic composition and generation by partial fusion of pre existing igneous rocks in the crust (e.g., Sun and Chen, 1992; Roberts and Clemens, 1993; Ajaji et al, 1998). This feature indicates that similar to I-type granites from elsewhere (Gromet and Silver, 1987: Wyborn et al, 1992; Wyborn, 1998), most plutons of the Kashmar granitoid were derived from plagioclase-dominated sources. The content of K20 is variable (1.36 to 5.88 wt%), but typically high (mostly >2.5 wt%), characteristic of subduction-related granites in continental margin settings (Grigoriev and Pshenichny, 1998; Rottura et al, 1998). K20 shows a positive correlation with Si02 but an inverse correlation with all other major oxides. This behaviour is normal and occurs when granitoid magmas fractionate and become enriched in alkalis, Si02 and a water-rich vapour phase (Raymond, 1995). Such processes are supported by the occurrence of aplitic veins and dykes that microscopically show allotriomorphic granular textures. Rapid chilling or depletion of certain chemical species in the melt, as well as other processes, may initiate the development of allotriomorphic granular texture. The content of P205 is low (<0.4 wt%). One analysis from the tonalite (R15912) has the highest content of P205 (0.33 wt%) as a result of the presence of abundant apatite 9TJ crystals. In the granodiorite and granite, with increasing Si02 contents the content of P2Os decreases in a regular fashion (Fig. 5.1). This behaviour is consistent with a decrease in apatite and microgranular enclaves towards higher Si02 contents in these plutons. The regular pattern of decreasing P205 with increasing Si02 is characteristic of low-temperature I-type granites (Chappell et al, 1998). The lowest content of P2Os (0.01-0.03 wt%) is observed in the alkali feldspar granite, which is dominated by quartz and K-feldspar, consistent with the lower solubility of P in more felsic and lower temperature granite melts (Harrison and Watson, 1984; Chappell et al, 1998). Typical examples of low-temperature I-type granites reported from Cobargo and Inlet plutons (Lachlan Fold Belt, southeastern Australia) that show regular variation of P2Os from mafic to felsic rocks is a consequence of crystallisation from a 'minimum-melt' composition (Chappell et al, 1998). In the Kashmar granitoid, the behaviour of P205 and the quartzofeldspathic nature of the alkali feldspar granite indicate that plutons of the Kashmar granitoid are I-type and would have crystallised at low temperature. 5.2.2 SUMMARY OF MAJOR ELEMENTS Major element data from mafic and felsic rocks of the Kashmar granitoid show several common features including high A1203, Na20, K20, total Fe as Fe203, and low CaO, MnO, P205, Ti02 and MgO contents (Appendix 4.1). Harker diagrams (Fig. 5.1) show a continuum of compositions with linear trends from the tonalite to the alkali feldspar granite. All tonalite samples are low in Si02 content and define a mafic composition. But the main plutonic bodies of the Kashmar granitoid are felsic in composition. In the granodiorite and granite, with increasing Si02, hornblende disappears, plagioclase, biotite and Fe-Ti oxides decrease to a smaller extent but K-feldspar and quartz become dominant. These continuous and regular trends for major oxides favour the concept of crystal fractionation (e.g., Sewell et al, 1992; Mason, 1996; Grigoriev and Pshenichny, 1998). The absence of microgranular enclaves and sparse occurrence of apatite in the most felsic compositions (Si02 >74 wt%), may suggest restite fractionation. All samples from the alkali feldspar granite are strongly enriched in silica and K-feldspar. They are low in Al203, MgO, CaO and total Fe as Fe203 contents. In the alkali feldspar granite, biotite is the only hydrous mineral and occurs in low amount (<0.2-4.6 modal%). The composition of the alkali feldspar granite is similar to the felsic haplogranites of the LFB (cf. Chappell, 1998b). The haplogranites dominantly can be formed initially as primary melts that separated from restite, and less often by the fractionation of mafic high temperature melts (Chappell et al, 2000). But all mineralogical features of the Kashmar granitoid indicate crystallisation from low temperature magma. For each pluton of the Kashmar granitoid, the absence of inflection in the P205 data and the trend of decreasing P2Os with increasing Si02 support crystallisation from a relatively low temperature magma. Therefore, fractional crystallisation is not the only possible process for variation in the chemical composition of the Kashmar granitoid. 5.2.3 INCOMPATD3LE ELEMENTS As the primordial mantle is the ultimate source of all I-type magmas, incompatible trace element contents in Tertiary rocks from the Kashmar granitoid were normalised to primordial mantle, using the normalising values of McDonough et al. (1991). To assess the degree of chemical fractionation of the granitoid from this original common source, samples are plotted in spider diagrams (Fig. 5.2). The patterns for each pluton are almost parallel but in the tonalite, granodiorite and granite, some incompatible elements such as Ba, Th and U are slightly scattered due to differences in the contents of hornblende, biotite and K-feldspar. In general, the spider diagrams for different plutons of the "TZ Kashmar granitoid are similar, implying their close genetic relationship. Slight depletion is observed only in Ti and P in samples from the alkali feldspar granite. This could reflect apatite and ilmenite fractionation. Other trace elements, particularly Rb, Ba, Th, K and La, are strongly enriched in different samples from the Kashmar granitoid. The tonalite, granodiorite and granite show an overall relative enrichment with increasing incompatibility from right (Na) towards left (Pb) on the spidergram patterns, with distinct negative anomalies for Ba, Nb, P and Ti. For Sr, a negative anomaly is observed only in samples from the alkali feldspar granite. This may indicate that fractionation of plagioclase depleted the magma in Sr. All plutons show distinct Nb and Ti anomalies when normalised to primitive mantle compositions, which is a geochemical indication for involvement of a subduction-type environment (Hill et al, 1992; Whalen et al, 1996; Ajaji et al, 1998; Soesoo, 2000). Compared to some I-type granites that show direct mantle contribution (e.g., Kent, 1994), the relatively moderate to high concentration of various incompatible elements such as Ba, Rb, Sr, Th, Zr, Y, La, Ce, Nd, Sc and V (Appendix 4.1) indicate an infracrustal source (e.g., Brownlow, 1996). In particular, most of the low field strength elements (LFSE) have higher concentrations due to the abundance of various silicate minerals such as amphibole, biotite, feldspars and accessory minerals (Appendix 2). 5.2.3.1 Low Field Strength Elements (LFSE) The content of Ba is high but variable (140 to 745 ppm), with most samples containing -500 ppm Ba. High content of Ba is consistent with the presence of abundant biotite and K-feldspar in the analysed rocks (e.g., Sample R15906). Jn general, with increasing Si02, a convex curvature for Ba is observed in Figure 5.3. With increasing Si02 contents, samples from the tonalite, granodiorite and granite show an increase in Ba content. This may be related to fractionation of plagioclase and hornblende, particularly Pas/L for plagioclase because it has a mineral partition coefficient for Ba (DBa ) that is substantially less than one (Blundy and Wood, 1991; Chappell et al, 1998). With increasing Si02, a negative trend is observed for Ba contents from the alkali feldspar granite (Fig. 5.3). This rock contains abundant K-feldspar and a small amount of biotite. Because biotite and K-feldspar have high partition coefficients for Ba, their crystallisation can significantly lower the Ba content of the melt. For example the lowest Ba content (140 ppm) is observed in Sample R15914 from the alkali feldspar granite. This sample is low in biotite (2.2% modal) but extremely high in Si02 (76.97 wt%) and modal content of K-feldspar (56.2%). Therefore, a large fraction of Ba was partitioned into K-feldspar crystals. Precipitation of K-feldspar crystals from the melt of the alkali feldspar granite resulted in Ba content falling sharply in the remaining melt, consistent with fractional crystallisation processes (e.g., Wybom et al, 1987; Chappell et al, 1998; Wyborn, 1998). The content of Rb is low and mostly <150 ppm (Appendix 4.1). For all samples from the tonalite, granodiorite and granite, Rb/Sr is less than one, whereas for the alkali feldspar granite Rb/Sr is high (1.14-5.75). Rb does not correlate with Si02 content, possibly reflecting on variability of biotite and K-feldspar content in different rock types. The low content of Rb in the granite samples may reflect the source composition (Faure, 1986; Grigoriev and Pshenichny, 1998) or fractional crystallisation of biotite and K-feldspar (Chappell, 1996b). In the case of the Kashmar granitoid, fractional crystallisation is more likely, because rocks with lower Rb content have a higher modal content of biotite. For instance, the tonalite with Rb = 45-72 ppm and the granodiorite with Rb = 56-88 ppm, contain much more biotite than the alkali feldspar granite which has a Rb content mostly between 120 and 207 ppm. However, the latter rock is not significantly enriched in Rb content, because it contains abundant K-feldspar. Although Rb and K are similar in chemical properties and ionic radii, biotite has a substantially higher mineral partition coefficient for Rb than K-feldspar (Ewart and Griffin, 1994). This feature is consistent with the lower Rb content in the tonalite and granodiorite that contain appreciable amounts of biotite. Therefore, the low content of Rb in rocks of the Kashmar granitoid is related to fractionation of biotite and K-feldspar. The content of Sr ranges from 367 to 36 ppm, and shows a typical negative correlation with increasing Si02 contents (Fig. 5.3). Sr is moderately high in the tonalite (316- 367 ppm), granodiorite (256-342 ppm) and granite (188-315 ppm; Appendix 4.1). Concentration of Sr is relatively low (36-72 ppm) only in samples from the alkali feldspar granite, consistent with the lower plagioclase abundances of this rock. Since Sr is partitioned into and retained by plagioclase, higher modal abundances of plagioclase in the tonalite, granodiorite and granite may be correlated with higher concentration of Sr in these rocks. Although fractionation of plagioclase may lower the Sr contents of the melt, the high Na20 and Sr contents observed in the Kashmar granitoid, is considered to be dominantly a primary feature of the source rocks. Also, high Na20 and Sr contents precludes significant alteration in the Kashmar granitoid (e.g., Pollard et al, 1995). Concentration of Sr in rocks with Si02 >69 wt% from the Kashmar granitoid are similar to Sr contents of highly fractionated I-type granites (at the given Si02 content) from western Tasmania (Sawka et al, 1990). The content of Th increases from mafic towards felsic compositions (Fig. 5.3). Among different plutons of the Kashmar granitoid, the highest concentration of Th (31 ppm) is observed in the alkali feldspar granite, which contain Si02 >74 wt%. Although, Th mostly accommodates in the structure of the ferromagnesian minerals, but all samples from the alkali feldspar granite lack in calcic amphibole and are very low in biotite. The presence of zircon, titanite and apatite may explain high Th content of the alkali feldspar granite. Some authors believe that the enrichment in Th is typical of that expected for an incompatible LFSE in a fractional crystallisation system (e.g., Wyborn, 1983, Chappell, 1998b). 5.2.3.2 High Field Strength Elements (HFSE) The Y content of the Kashmar granitoid is relatively high and dominated by values between 20 and 32 ppm. On a Harker plot (Fig. 5.3), samples from the tonalite show an increase in Y content with increasing Si02 content from 55 to 60 wt%. For other rock types the Y contents show scattering with increasing Si02 contents. This behaviour of Y content is due to its substitution into many different minerals (Belolipetskii and Voloshin, 1996; Larsen, 1996). The relatively high and scattered Y content in the Kashmar granitoid is mainly due to the presence of various quantities of calcic amphibole, apatite, titanite and K-feldspar. It seems that calcic amphibole is critical because trivalent Y replaces Ca2+ in calcic amphiboles probably by coupled substitution with Na or K (e.g., Wyborn, 1983). This explanation is plausible for most samples of the granodiorite and granite that contain appreciable apatite, calcic amphibole and have relatively high P205 contents. But the content of Y is high in the most felsic samples (Si02 >74 wt%) from the alkali feldspar granite. The other common minerals that can readily accommodate Y are biotite and zircon (Ewart and Griffin, 1994); both minerals are common in the alkali feldspar granite. Among the HFSE, Zr has the highest concentration and ranges in value between 84 and 250 ppm; most analyses contain Zr >150 ppm (Appendix 4.1). High concentration of Zr is consistent with high Na20 contents of these rocks (e.g., Deer et al., 1992) and indicates the absence of significant alteration (Rubin et al, 1993). Figure 5.3 shows that the content of Zr is variable in the tonalite. This may be related to the different abundances of amphibole and zircon grains (Chen et al, 1990). From the granodiorite towards the alkali feldspar granite Zr content generally decreases with increasing Si02 content. This is related to a decrease in modal abundances of zircon and ferromagnesian minerals. Because, in the Kashmar granitoid, zircon is mostly accompanied by calcic hornblende which indicates a low temperature crystallisation, the negative trend for Zr versus Si02 may suggest zircon crystals would be present in the melt (e.g., restite). If the melts that produced the different plutons of the Kashmar granitoid were high temperature, they would be Zr-undersaturated. Hence by increasing Si02, the content of Zr would increase in the melt until zircon started to crystallise, then Zr would decrease in abundance. This behaviour is typical of high-temperature I-type granites (King et al., 1997; Chappell et al., 1998) but does not accord with the Kashmar granitoid. 5.2.3.3 Rare Earth Elements (REE) The REE concentrations were normalised to standard CI chondrite using normalising values of Taylor and McLennan (1985) and plotted in Figure 5.4 These rocks are characterised by enrichment in LREE relative to HREE, with Lau values between 38.15 and 87.19. The EREE ranges from 83.74 to 94.32 ppm in the granodiorite, 80.55 to 106.73 ppm in the granite and 132.06 to 143.28 ppm in the alkali feldspar granite (Appendix 4.1). For most samples, the increase in XREE is slightly greater in LREE than in HREE, so that LaN/YbN (5.60-8.71) increases with increasing whole rock Si02 y / contents (Appendix 4.1), All samples display steep negative slopes for LREE, moderate to strongly negative anomalies for Eu (Eu/Eu* = 0.88-0.18), and flat to slightly negative gradients for HREE. The steep LREE and flat HREE, with slight depletion trend, imply fractionation of amphibole, plagioclase and accessory minerals, and indicate garnet-free residuum in the source (Henderson, 1984; Skjerlie, 1992; Green, 1994); characteristics of I-type magmas (King et al, 1997). Furthermore, amphibole fractionation is suggested because the content of Sc drops with increasing Si02 (Fig, 5.3), and the most REE-rich samples lack amphibole. Also, the flat HREE gradient is consistent with the relatively high concentration of Y (Fig. 5.3) and negative Eu anomalies, all features of granites that have experienced fractional crystallisation (Allen et al, 1997). A strong similarity of REE patterns and incompatible element ratios between the granodiorite and granite support a genetic link between them. The main difference between REE patterns is that the lowest level of ZREE and Eu anomaly is observed in the granodiorite. Both values are intermediate in the granite. The alkali feldspar granite has the highest SREE (132- 143 ppm) with a pronounced Eu anomaly (Eu/Eu* = 0.18-0.20). Such differences suggest a higher degree of fractionation for the alkali feldspar granite (e.g., Sun and Chen, 1992). Petrographic observations demonstrate that the abundances of the REE are correlated with the abundances of accessory minerals such as titanite, apatite, and zircon. This is compatible with the fact that titanite alone may hold up to 90% of the REE in felsic rocks (Skjerlie, 1992; Peacock et al, 1994). 5.2.4 COMPATIBLE ELEMENTS Nickel and Cr have very low concentrations (<6 ppm and <22 ppm, respectively). Even the most mafic samples analysed from the tonalite are depleted in these elements. Ni and Cr contents in the granites are highly dependent on hornblende and Fe-Ti oxide contents (Soesoo, 2000). Therefore low contents of Cr and Ni in the Kashmar granitoid suggests that a large amount of ferromagnesian minerals may have been fractionated or the source was low in such compatible elements (e.g., Tate et al, 1999). The content of Sc is <15 ppm; it becomes extremely depleted in samples with Si02 >70 wt% (Fig. 5.3), which supports fractionation of amphibole and Fe-Ti oxides. Vanadium content shows a wide range (200-4 ppm). With increasing silica, the content of V decreases with a regular trend (Fig. 5.3). This is to be expected because V substitutes for Fe3+ in mafic silicates, especially in amphibole and Fe-Ti oxides. Similar V behaviour has been reported from I-type granites of the New England Batholith of eastern Australia (Bryant et al, 1997). Among the compatible elements, Mn has the highest concentration particularly in the mafic rocks (up to 1070 ppm), and decreases in a linear trend through the felsic rocks (Fig. 5.3). The higher Mn content of mafic rocks indicates that Mn is partitioned into ferromagnesian minerals. Concentrations of Mn from different rock types in the Kashmar granitoid are similar to Mn contents of I-type granites that have been generated in subduction-related environment (e.g., Silver and Chappell, 1988). The behaviour of Ga versus Si02 (Fig. 5.3) is very similar to Sr versus Si02 contents and represents a negative trend from mafic to felsic compositions. This trend is typical of that expected for a compatible element, since Ga3+ substitutes Al3+ in plagioclase. Collectively, all transition metals in the different plutons of the Kashmar granitoid show increases in compatibility with increasing Si02 contents and supports fractional crystallisation process. 5.2.5 Sr AND Nd ISOTOPES Isotopic data from different plutons of the Kashmar granitoid (Table 5.1) show low 87 86 initial Sr/ Sr (0.70471-0.70569) and eNd values (-0.70 to -1.86). The eNd values show a narrow range, and for most samples these values are indistinguishable within the limits of analytical uncertainty. For each pluton differences between initial 87Sr/86Sr values are small but are significantly larger than 2a analytical uncertainties. Consequently, a model of simple fractional crystallisation of any isotopically uniform melt for the generation of on a/: each pluton is unlikely. Low initial Sr/ Sr and slightly negative sm values suggest a typical infra-crustal source for the origin of the Kashmar granitoid (e.g., Rapela et al, 1992). The initial 87Sr/86Sr ratio is relatively higher in the granodiorite (0.70552-0.70569) than in the granite (0.70516-0.70550). Slight Sr isotopic heterogeneities are apparent for each pluton but this is more distinct for samples from the granite. Considering the 2o analytical uncertainty (±0.00005), slight overlap is observed between initial 87Sr/86Sr ratios from the granodiorite and granite. This overlap suggests similarity in the source composition of the granodiorite and granite and supports the similarity in REE patterns and incompatible element ratios between them. In the granite, except for Sample R15909, other samples show an increase in initial 87Sr/86Sr ratio with increasing whole rock Si02 content. This feature would be the result of fractional crystallisation (e.g., Soesoo, 2000). Some authors (Graham and Hackett, 1987; Price et al, 1999) pointed out that a broad positive correlation between Si02 abundances and Sr/ Sr isotopic ratios is a feature attributed to assimilation of a crustal component during crystal fractionation. But granite from the Kashmar granitoid shows a very limited range in 87 86 Sr/ Sr isotopic ratios and eNd values (-0.70 to -1.58) that suggest assimilation of older crust is unlikely (cf. Bryant et al, 1997). This implies that the magma chemistry was effectively controlled by fractional crystallisation as evidenced from behaviour of most of the major and trace element data. In the Kashmar granitoid, the lowest initial 87Sr/86Sr ratio (0.70471-70478) is observed in the alkali feldspar granite. This characteristic suggests a different source for the alkali feldspar granite. However, this rock is quartzofeldspathic with a strongly enriched in 87 86 Si02 content (74.63-77.06 wt%). The relatively low initial Sr/ Sr in the alkali feldspar granite may be the result of a low 87Rb/86Sr in the quartzofeldspathic source at the time of melting (e.g., Chappell et al, 1999). Although a different source is inferred, the source rocks do not impose the sole control on the Sr isotopic signature of granitic magmas (Gray, 1990; Barth et al, 1993; Nelson et al, 1993). For example, if melting R*7 Rri proceeds faster than the Sr diffusion rate in crystals, then a hquid with Sr/ Sr value that is lower than the source can be produced (Pichavant et al, 1996). Similar 87 86 contrasting behaviour between Si02 and initial Sr/ Sr ratio may be observed when compositional variation of granites is mostly controlled by restite (Chappell et al, 1999, 2000). 5.3 BORNAVARD GRANITOID The whole rock samples analysed from the Bornavard granitoid comprise two tonalites, four granodiorites and seven granites. To understand relationships between the composition of volcanic and plutonic rocks of the Bornavard area, five samples from the Taknar Rhyolite have been chemically analysed. The whole rock geochemical data for the Bornavard granitoid and the Taknar Rhyolite are presented in Appendices 4.2 and 4.3, respectively. 5.3.1 MAJOR ELEMENTS Tonalite is the only mafic rock that occurs in the central part of the Bornavard granitoid. The tonalite (Si02 = 48.64-58.09 wt%) and granodiorite (Si02 = 63.18-71.32 wt%) show a relatively wide range in silica contents, whereas the granite is more homogeneous and shows a very narrow range in Si02 contents (74.84-76.04 wt%). The distinctive geochemical features of these rocks, particularly the granodiorite and granite, are low (<1 wt%) Ti02, MnO, P205 and high A1203, total Fe as Fe203 and Na20 contents. The content of Na20 is mostly >3.5 wt% which suggests that the source rocks have not undergone hydrothermal or meteoric alteration, and hence they would be genetically I-type (e.g., Sun and Chen, 1992; Raymond, 1995; Ferre et al., 1998). The content of K20 is high (>3.5 wt%) in the granite due to the presence of large quantities of K-feldspar (Appendix 2.2). MgO and CaO are variable but, relative to the granodiorite, they are higher in the tonalite, reflecting the higher modal proportions of calcic amphibole and plagioclase in the tonalite (Appendix 2.2). On Harker plots (Fig. 5.5) the tonalite and granodiorite show negative correlation for Ti02, total Fe as Fe203, MnO, MgO and CaO contents. This feature may suggest that fractional crystallisation was involved in the generation of the tonalite and granodiorite plutons (e.g., Jung et al, 1998). In samples taken from the granite, significant variation in the composition of major oxides was not observed. This is consistent with small variation in mineralogy and the homogeneous nature of the granitic pluton. 5.3.2 INCOMPATIBLE ELEMENTS Primitive mantle-normalised elemental contents for the Bornavard granitoid are shown in Figure 5.6. These rocks show an overall relative enrichment with increasing incompatibility. The most striking feature of these data is that all plutons show marked negative anomalies for Ba, Nb, Sr, P and Ti. Compared with primordial mantle, Ti and P depletion occurs only in the granite samples, consistent with the low modal content of apatite, biotite and Fe-Ti oxides. Also, Sr shows a stronger negative anomaly in the granite compared with the granodiorite. This can be explained by lower modal abundances of plagioclase in the granite (Appendix 2.2). The behaviour of incompatible elements for most samples from the Bornavard granitoid is generally similar to those of the Kashmar granitoid (Figs 5.2 and 5.6). This is related to the similarity in mineralogy for different plutons from both granitoids. In particular, enrichment in Rb, Ba, Th, U, K, La, Ce and Nd elements is stronger in the granite and the alkali feldspar granite from the Bornavard and Kashmar granitoids, respectively. Both rock types are quartzofeldspathic (Si02 = 74-77 wt%) and enrichment in Rb, Ba and K may have resulted from high modal abundances of K-feldspar and common accessory minerals such as zircon, titanite and allanite. In the Bornavard granitoid, most incompatible elements (e.g., Rb, Ba, Th, U and K) are higher in the granite than in the tonalite and granodiorite. However, for most analyses there is a considerable similarity in the behaviour of less incompatible elements (Nd, Zr, Y and Na). Low contents of incompatible elements in the tonalite and granodiorite may be explained by the presence of appreciable amount of amphibole and biotite. It seems that biotite is favoured by Rb, Ba and K, while amphibole is favoured by Th and U (Wyborn, 1983), thus explaining the lower content of the above incompatible elements in the tonalite and granodiorite. The relative enrichment with increasing incompatibility, together with negative anomalies particularly for Nb and Ti, represent characteristics of I-type granites from subduction-related environment (e.g., Ferret al, 1998). 5.3.2.1 Low Field Strength Elements (LFSE) The content of Ba is low in the tonalite (70-85 ppm) and shows a wide range in the granodiorite (55-810 ppm). It is notably high in the granite (580-890 ppm; Appendix 4.2). It seems that variation in Ba content of the tonalite and granodiorite is mainly controlled by biotite. Thus the low content of Ba in the tonalite is consistent with the absence of biotite (Appendix 2.2). Ba variation in the granodiorite (Fig. 5.7) illustrates an inflexion point at Si02 = -69 wt% that is characteristic of granites in which the compositional variation resulted from fractional crystallisation (e.g. Chappell, 1998b). From 63 wt% Si02 up to -69 wt Si02, Ba increases in abundance, from 535 ppm to 810 ppm (Appendix 4.2). Samples with Si02 below 69 wt% (R15946 and R15947) are high in modal abundances of biotite (29.2 to 12.4%). With increasing Si02 from -69 wt% to 71.32 wt%, the abundance of biotite decreases to a negligible amount and Ba decreases to its lowest content (55 ppm). Therefore variation in Ba content in the granodiorite is mainly related to fractionation of biotite. The Ba variation in the granite is different from Ba variation in the granodiorite and tonalite. In the granite, Si02 is relatively constant and the high concentration of Ba is related to the presence of abundant K-feldspar and small amount of biotite. These minera]/L minerals have high partition coefficients for Ba (£>Ba - ~7). According to Chappell (1996b) and Chappell et al. (1998) in the most felsic compositions, the mineraiyL magnitude of DBa substantially increases to higher values than for the most mafic composition. This explains the higher Ba content in the granite, compared with the tonalite, the former rock is strongly felsic (Si02 = 74.8-76 wt%) whereas the latter is strongly mafic (SiQ2 = 48-58 wt%) in composition. Since the granite shows features of high-K systems (K20 = 3.65-4.35 wt%), the Ba contents are higher in that pluton, indicating K-feldspar was saturated in that melt because of the higher overall K content of the melt (e.g. Chappell et al, 1998). The content of Rb is very low in the tonalite (6-12 ppm). The behaviour of Rb in the granodiorite is noteworthy because it rapidly decreases from moderate levels (129 ppm) to less than 10 ppm with increasing Si02 contents (Fig. 5.7). Such rapid depletion, coupled with decreasing K20 and Ba contents (Appendix 4.2), can be explained by fractional crystallisation of biotite (e.g., Champion, 1991), a conclusion consistent with the large changes in the incompatible elements in the granodiorite (Fig. 5.6). In the granite Si02 has a restricted range (74.8-76 wt%), but Rb changes from 54 to 128 ppm, resulting in a wide range of Rb/Sr (1.13-3.28) in this felsic rock. However, variation in Rb content in the granite is independent of modal abundances of biotite and K-feldspar. This is a common feature in quartzofeldspathic rocks, as stated by Chappell (1996b), in felsic I-type granites those trace elements that occur in feldspars (e.g., Rb, Sr, Ba) may vary widely in abundance. In general, plutons of the Bornavard granitoid have slightly lower contents of Sr (168- 38 ppm) compared with plutons of the Kashmar granitoid. This may be related to lower abundances of plagioclase in the Bornavard granitoid that reflects the initial source composition was different. In the granodiorite, removal of Sr may be the result of fractional crystallisation of plagioclase. Similar to the Kashmar granitoid, the Rb/Sr ratio is low (0.04-0.89) in rocks with Si02 <74 wt% and high (1.13-3.28) in rocks with Si02 >74 wt% (Appendices 4.1 and 4.2). In the granite, the content of Sr does not show any regular variation with Si02 content (Fig. 5.7). J.UJ 5.3.2.2 High Field Strength Elements (HFSE) The Zr contents in the granite samples is relatively constant (226-240 ppm). In the tonalite samples zircon was not observed in thin sections but concentration of Zr ranges from 48 to 208 ppm that may reflect variation in modal abundances of hornblende (e.g. Chen et al, 1990). In the granodiorite samples, Zr variation is very similar to that of Ba variation (Fig. 5.7). Below about 69 wt% Si02, zircon is not observed in the granodiorite, possibly because the melt was undersaturated in zircon (e.g., Chappell et al, 1998). Again, beyond that point, concentrations of Zr decrease in a regular fashion from 448 to 142 ppm. This is consistent with the presence of zircon crystals in all samples with Si02 >69 wt% from the granodiorite. This means that after the inflexion point the granodiorite melt was saturated in zircon and crystallisation of zircon from that melt resulted in a decrease in Zr abundance. Ba concentrations show similar trends to Zr, and Rb concentrations fall sharply (Fig. 5.7). All are characteristics of fractional crystallisation occurring in I-type melts (Chappell, 1996a; Chappell et al, 1987,1998). Uranium is low and shows a similar increasing trend to Th (Fig. 5.7). The Th/U ratio is mostly around 4 to 8 and shows a weak increasing trend from mafic to felsic rocks. However, there is a large scatter in Th/U, probably because of large variations in the content of accessory phases such as titanite, apatite, zircon and allanite. The contents of U and Th are higher in the granite samples because they contain abundant accessory phases including allanite. Yttrium ranges between 19 and 69 ppm for different rock types from the Bornavard granitoid. For most samples the content of Y is high (>38 ppm) as a result of the 1UD presence of a variety of accessory minerals such as titanite and zircon. There is an apparent jump in Y values at about 69 wt% Si02 for the granodiorite, probably corresponding to the appearance of zircon grains and the presence of abundant biotite as another important Y-bearing phase (e.g., Ewart and Griffin, 1994; Green, 1994). Except for one sample (R15946) from the granodiorite, other samples of this rock show a negative trend for Y with increasing Si02 that is a general feature of LREE in low- temperature I-type granites (Chappell et al, 1998). The content of Y is relatively high and shows small variation in the granite (48-59 ppm). This is consistent with a lack of significant variation in the abundances of accessory phases particularly zircon grains in the granite samples. 5.3.2.3 Rare Earth Elements (REE) The REE concentrations in one tonalite, four granodiorite and three granite samples are plotted relative to standard CI chondrite (Fig. 5.8). The EREE content ranges from 109 to 214 ppm in samples from the Bornavard granitoid (Appendix 4.2). The chondrite- normalised REE patterns display variable enrichment in LREE and MREE, with LaN/YbN values mainly ranging between 3.38 and 5.76. There is a good consistency in the REE patterns from the granite samples, reflecting the homogeneous nature of this pluton. In the tonalite and granodiorite samples, enrichment is chiefly controlled by presence of hornblende and biotite crystals, whereas in the granite accessory phases such as apatite, titanite, zircon and allanite strongly influenced the REE patterns towards extreme enrichment. One sample from the granodiorite (R15953) with EREE = 120 ppm, shows significant enrichment in HREE, possibly due to the presence of abundant titanite grains (2.4% modal). Except for Sample R15947 (Eu/Eu* = 1.1) from the granodiorite, other analyses display moderate to strong negative Eu anomalies 107 (Eu/Eu* = 0.37-0.97) indicating fractionation of plagioclase (e.g., Henderson, 1984; Villaseca et al, 1998). The chondrite-normalised patterns for the granite are characterised by strongly enriched LREE and MREE (LaN = 86-125), pronounced Eu anomalies (Eu/Eu* = 0.37-0.45) and the highest in IREE contents (177-214 ppm). The stronger negative Eu anomaly in the granite is consistent with the low Sr content of this rock (Appendix 4.2). The tonalite and most samples from the granodiorite are variably depleted in HREE because of differences in amounts of hornblende and accessory phases. Collectively, the REE patterns exhibit a broad spectrum of compositions that imply rocks of the Bornavard granitoid are not genetically related. 5.3.3 COMPATIBLE ELEMENTS Most compatible elements, including Sc, V, Cr, Mn, Ni, Cu and Zn, decrease in abundance from mafic to felsic plutons of the Bornavard granitoid (Appendix 4.2). This behaviour is normal because compatible elements such as Ni, V and Cr preferentially partition into mafic minerals (e.g., hornblende and biotite) particularly at higher temperatures (Rollinson, 1993). Among all trace element data Mn has the highest concentration in the tonalite (up to 1300 ppm) reflecting its high contents of ilmenite (Appendix 3.10) and hornblende (Appendix 2.2). On Harker plots, Mn and V display negative correlation with increasing Si02 contents (Fig. 5.7). Concentrations of Zn in the granodiorite show an inflexion at Si02 = -69 wt% that is similar to the behaviour of Ba and Zr elements versus Si02 contents (Sections 5.3.2.1 and 5.3.2.2), emphasising fractional crystallisation in the granodiorite melt. The content of V rapidly decreases from 240 ppm in the tonalite to <2 ppm in the granite. The different behaviour of V is because of its strong partitioning into magnetite, consistent with the composition of Fe- Ti oxides, being magnetite and titanomagnetite in granite samples (Appendix 3.10). lUt! However, the granite samples are extremely depleted in most of the compatible elements suggesting that the source of the granite was strongly fractionated. Sc is nearly constant among samples from each pluton. Concentration of Sc in the granodiorite (16.4- 17.0 ppm) is higher than in the granite (11.5-11.8 ppm). This is expected because the granodiorite contains hornblende and Sc accommodates into hornblende (Wyborn, 1983). 5.3.4 Sr AND Nd ISOTOPES For whole rock analyses, the initial ratios were calculated at 149.2 and 117.8 Ma for the tonalite and granodiotite, respectively. These ages are the average ages obtained at the present study by Rb-Sr dating on biotite-whole rock pairs for the first and the second intrusive episodes of the Bornavard granitoid (Section 3.3.2.1). The Bornavard granitoid R7 Q^ exhibits a broad spectrum of isotopic characteristics with initial Sr/ Sr ranging from 0.70757 to 0.75008 values and the sNd ranging from -1.41 to -5.20 values (Table 5.1). The range of initial isotopic ratios and the &m values are significantly larger than the estimated 2c analytical uncertainties which equate to ±0.00005 and ±0.5 for initial R7 Rfi Sr/ Sr and eNd values, respectively (Appendix 1). An important feature of these data is the high initial 87Sr/86Sr and low ENd values, with Sr and Nd-isotope heterogeneity within and between individual plutons. The lowest initial 87Sr/86Sr (0.70798) was observed for the granodiorite but it is significantly higher than values are proposed for mantle-like sources (cf. Soesoo, 2000). The granodiorite 87 86 has wide range in initial Sr/ Sr (0.70757-0.72153) and eNd values (-1.41 to -4.29). The 87 86 tonalite has an initial Sr/ Sr of 0.70820 and sNd value of -4.16, both values are within 109 "^^^ the range of isotopic values from the granodiorite, indicating similarity in isotopic characteristics of the early intrusive rocks of the Bornavard granitoid. The initial 87Sr/86Sr values of granite are significantly high and wide in range (0.73622-0.75008). The wide range in initial 87Sr/86Sr ratios of the granite is in contrast with its uniform chemical composition and a limited range of em values (-4.50 to -5.20) from different samples of the granite. Differences between initial 87Sr/86Sr values of the early and late intrusive episodes is the result of a -22 Ma age different between emplacement of the granodiorite and granite. During this time interval, the source of magmas changed towards strongly felsic nature. The origin of very high values of initial 87Sr/86Sr in granites has been attributed to a variety of processes including: (1) derivation from metasedimentary sources (Sylvester, 1998); (2) protracted crystal fractionation (McCarthy and Cawthorn, 1978); (3) hydrothermal exchange of Sr with surrounding country rock (Richardson et al, 1990; Cuney et al, 1992); and (4) assimilation of older crustal components (e.g., Graham and Hackett, 1987; Hess, 1989; Mason et al, 1996; Moghazi et al, 1998; Price et al, 1999). In the Bornavard granitoid, metasedimentary enclaves are absent and mineralogical 87 Rfi properties are not consistent with S-type sources. In plots of Si02 against initial Sr/ Sr (Fig. 5.9),* rocks of the Bornavard granitoid show two trends. By increasing Si02 content, the initial 87Sr/86Sr values in the granodiorite decrease. This is in contrast with the operation of fractional crystallisation. A decrease in the initial 87Sr/86Sr may be a response to restite separation (e.g., Chappell et al, 1999). This process may be supported by presence of dark xenoliths (Section 4.3.2), possibly being restite in the 87 86 granodiorite. In the granite, the initial Sr/ Sr ratio increases while Si02 is relatively constant. Due to the lower Sr contents of the granite samples (38-54 ppm) and presence 110 of allanite and notable amount of sericite in this rock, hydrothermal alteration may be responsible for the generation of high initial 87Sr/86Sr values. But the effect of hydrothermal fluids is not consistent with the low variation in 8Nd values (-4.5 to -5.20) in the granite (e.g., Darbyshire and Shepherd, 1994; Darbyshire and Sewell, 1997). Therefore, significant enrichment in initial 87Sr/86Sr values of the Bornavard granitoid may be the result of extensive contamination of magmas with radiogenic Sr derived from old felsic rocks of the continental crust, or the magmas were produced by partial melting of old felsic rocks (e.g., Sewell and Campbell, 1997). 5.4 TAKNAR RHYOLITE 5.4.1 MAJOR AND TRACE ELEMENTS Chemical data for five samples from the Taknar Rhyolite are presented in Appendix 4.3. The high content of Si02, ranging from 75.75 to 77.90 wt%, is a typical feature of continental arc rhyolites (Raymond, 1995). The content of A1203 shows a limited range from 10.9 to 12.6 wt%. The molar proportion of A1203 is higher than total alkalies, consistent with the presence of appreciable normative C (0.51-6.80%) for most of the rhyolite samples (Appendix 2.3). The aluminium saturation index (ASI) ranges from 1.04 to 2.46 and is characteristic of weakly to strongly peraluminous rhyolites (e.g., Reece et al, 1990; Feldstein et al, 1994). However, the peraluminous characteristic may not be an intrinsic feature of the Taknar Rhyolite because primary aluminous minerals are absent. The development of secondary muscovite may have resulted in the higher ASI values (Appendix 4.3). Secondary muscovite commonly replaces K-feldspar in the groundmass. The strongly peraluminous feature is observed only in Sample R15949 because K-feldspar in the groundmass is extensively sericitised. The Taknar Rhyolite is low in Ti02 (0.13-0.16 wt%), MnO (0.00-0.06 wt%), MgO (0.22-0.64 wt%), Ill CaO (0.13-0.35 wt%) and P205 contents (0.03-0.06 wt%). The content of K20 is high and ranges from 3.31 to 5.35 wt%. The highest K20 content is observed in Sample R15952 because it contains higher abundance of sanidine phenocrysts. i Spidergram patterns (Fig. 5.10) show similar trends for most of the incompatible elements from different samples of the Taknar Rhyolite. Depletions in Sr, P and particularly Ti are pronounced but other incompatible elements show strong enrichment in the rhyolite compared to primordial mantle. Significant negative anomalies for Sr, P and Ti elements indicate they are compatible in the residual mineral assemblage, which may be inferred to include plagioclase, apatite and titanomagnetite. One sample (R15949) with the highest Si02 content (-78 wt%) is extremely depleted in Na (Na20 = 0.37 wt%) and Sr (12 ppm) possibly due to lack of plagioclase phenocrysts. Some light rare earth elements such as La and Ce generally increase with increasing Si02 content, but also show quite an amount of scatter. This is perhaps caused by different amounts of accessory phases. The Taknar Rhyolite is high in Ba content (545-1020 ppm). This feature may be attributed to greater incorporation of Ba into K-feldspar that is the major component of the groundmass. Other LFSE such as Rb (76-130 ppm), Sr (12-93 ppm), Pb (6-12 ppm) and Th (17-23 ppm) are have low abundances in the Taknar Rhyolite (Appendix 4.2). Also, most of the HFSE, including Zr, Y, REEs and Hf, show lower concentrations in the rhyolite samples because biotite is negligible, accessory phases occur in low abundances and particularly titanite and allanite are absent. The concentrations of most of the incompatible elements and the first transition metals such as V, Cr, Ni, Cu and Sn are lower in the Taknar Rhyolite, compared with tonalite and granodiorite from the 112 Bornavard granitoid. This supports the lack of relationship between the Taknar Rhyolite and the Bornavard granitoid. The REE pattern from the Taknar Rhyolite (Fig. 5.11) is fractionated and enriched in LREE (LaN = 77.66), has a strong negative Eu anomaly (Eu/Eu* = 0.25) and flat HREE with LaN/YbN = 4.01. The abundances of REE in the Taknar Rhyolite (EREE = 84.48 ppm) is lower than REE abundances in the different plutons of the Bornavard granitoid (EREE = 109.41-214.05 ppm). Also, the Eu anomaly is much stronger in the Taknar Rhyolite. Overall, differences in mineralogy (Section 4.5.1) and concentrations of REE and incompatible elements from the Taknar Rhyolite preclude any relationship with the Bornavard granitoid. Such differences are consistent with the older age of the Taknar Rhyolite that forms part of the country rock of the Bornavard granitoid. 5.4.2 Sr AND Nd ISOTOPES 87 86 In the Taknar Rhyolite the initial (at 190 Ma) Sr/ Sr is high (0.72378) and the eNd value is low (-4.1; Table 5.1). The high initial 87Sr/86Sr of the rhyolite is consistent with the silica-rich nature of this rock and indicates that derivation of Taknar Rhyolite from the mantle would be precluded (e.g. Lightfoot et al, 1987). The initial (at 190 Ma) 87Sr/86Sr value of the Taknar Rhyolite is significantly lower than the initial 87Sr/86Sr of the granite from the Bornavard granitoid. This supports the lack of any genetic relationship between the granite and rhyolite in the Bornavard area. Because the Taknar Rhyolite has undergone hydrothermal alteration and low-grade metamorphism (Section 3.3.2.2), its isotopic signatures would have modified by these processes. Enrichment in initial 87Sr/86Sr and depletion in HFSE in the Taknar Rhyolite would appear to be the result of hydrothermal alteration (e.g., Pollard et al, 1995). In 1 particular, the Zr content of the Taknar Rhyolite is low (124-144 ppm) because this element is highly mobile in hydrothermal systems (Rubin et al, 1993). 5.5 KUH MISH INTRUSIONS Based on the Le Bas and Streckeisen (1991) scheme, the analysed samples from the Kuh Mish intrusions comprise one gabbro, four quartz monzodiorite and eight granodiorite samples. The gabbro is uniform in composition (Si02 = 45.75 wt%) but the quartz monzodiorite shows a relatively wide range in chemical data (Si02 = 51.85-60.71 wt%). The granodiorite occurs in three localities including the Kuh Mish, Darin and Namin areas. In the Kuh Mish area, granodiorite has been intruded by a quartz monzodiorite which shows sharp contacts. The Kuh Mish granodiorite ranges in composition from Si02 = 70.66 to 75.96 wt%. Granodiorite from the Darin (Si02 = 71.58 wt%) and Namin (Si02 = 63.93 wt%) plutons are similar in mineralogy, except the former pluton lacks biotite (Appendix 2.4). Geochemical data for the Kuh Mish intrusions are summarised in Appendix 4.4. 5.5.1 MAJOR ELEMENTS The gabbro is strongly enriched in A1203 (17 wt%), MgO (11.8 wt%) and CaO (16.6 wt%), because it contains mainly anorthite and diopside minerals. Except for total Fe as Fe203 (4.19 wt%), contents of other major oxides from the gabbro are low (<1 wt%), consistent with the most primitive nature of the gabbro. The Kuh Mish intrusions are low in normative C (0.00-1.69) and high in Na20 (mostly between 3 and 5 wt%). The content of K20 is low (<2 wt%) and variable because of different modal abundances of K-feldspar and biotite. Low content of normative C and high content of 114 Na20 are consistent with mineralogical features of the Kuh Mish intrusions that suggest I-type source. For all analyses of the Kuh Mish intrusions K20/Na20 is extremely low (mostly <0.5) which may indicate the absence of significant involvement of the continental crust in generation of these intrusions (e.g., Roberts and Clemens, 1993). On Harker diagrams, from the gabbro to granodiorite, contents of A1203, MgO and CaO decrease with increasing Si02 content (Fig. 5.12), consistent with an increase in Rb/Sr ratios from 0.01 to 0.63 (Appendix 4.4). Also, from the quartz monzodiorite to granodiorite, contents of total Fe as Fe203 and MnO typically decrease (Fig. 5.12). These features suggest that the Kuh Mish intrusions may form a differentiation series. In the granodiorite, Ti02 and P2Os decrease regularly with increasing Si02 (Appendix 4.4) because towards more felsic compositions, hornblende and biotite disappear and plagioclase decreases to a lower content. With the exception of the Namin Granodiorite (Sample R15926), all samples from Darin and Kuh Mish Granodiorites show a narrow range of composition for most of the major and trace element data. This feature suggests that these plutons have a close genetic relationship. 5.5.2 INCOMPATIBLE ELEMENTS In the gabbro most of the incompatible elements, including Ba, Rb, Pb, Th, U, Zr, Nb, Y and REEs, are lower than in the quartz monzodiorite and granodiorite. The abundances of most of these elements in the gabbro are close to the primordial mantle (Fig. 5.13). A slight positive Sr anomaly in the gabbro is possibly related to the crystallisation of calcic plag/L plagioclase (A1190-99) at low temperatures, a factor contributing to an increase in DSr values of plagioclase relative to the melt (Chappell, 1996b). Among the HFSE, the 115 content of Sc is high (41 ppm) in the gabbro due to the presence of abundant clinopyroxene grains. The spidergram patterns for the quartz monzodiorite and granodiorite show enrichment in most of the incompatible elements, particularly for some LFSE such as Rb, Ba, Th and K. hi the granodiorite, enrichment in some incompatible elements (e.g., Zr, Sr, K, Rb, La and U) is stronger than for the quartz monzodiorite. This is related to the presence of biotite, zircon and apatite in the granodiorite samples. Other trace elements in granodiorite and quartz monzodiorite show some similarities in concentrations. All samples analysed from the Kuh Mish intrusions are very low in Nb (<2 ppm) and Ti02 contents (mostly <0.56 wt%). When normalised to primitive mantle composition, the granodiorite shows distinct Nb and Ti anomalies, which is a geochemical indication for involvement of a subduction-type environment (e.g., Soesoo, 2000). This is also in agreement with recent tectonic models for the Sabzevar Zone proposing that formation of the Eocene ophiolites occurred in an island arc related environment (Ghazi and Hassanipak, 1999). 5.5.3 RARE EARTH ELEMENTS (REE) The REE concentrations from the Kuh Mish intrusions represent two distinct patterns (Fig. 5.14). In the gabbro, EREE is very low (3.93 ppm), LREE are strongly depleted and are lower than HREE (LaN/YbN = 0.16), the negative Eu anomaly is small (Eu/Eu* = 0.7) and the HREE display a convex downward trend becoming flat at Yb and Lu. None of the chondrite-normalised REE abundances approach 4x chondritic values and the REE signature is consistent with low concentrations of other incompatible trace elements in the gabbro. Most of the REE abundances in the gabbro are comparable to the concentrations of REE in the mantle peridotite (harzburgite) from Tethyan ophiolite belts in Iran, which have island arc affinities in the Sabzevar Zone (Lensch and 116 Davoudzadeh, 1982; Frey, 1984; Ghazi and Hassanipak, 1999). The REE concentrations for two granodiorites from the Namin (Sample R15926) and Darin (Sample R15927) plutons are very similar (LaN = 16 and 18 ppm, EREE = 38 and 40 ppm, respectively). The chondrite-normalised patterns for both plutons (Fig. 5.14) show slight fractionation of LREE (LaN/YbN = 1.54-2.18), flat MREE to HREE and small negative Eu anomalies (Eu/Eu* = 0.8-0.9). These features indicate that the Namin and Darin plutons evolved under similar magmatic conditions and were derived from similar source compositions. 5.5.4 COMPATD3LE ELEMENTS In the gabbro, the contents of Cr (805 ppm) and Ni (242 ppm) are extremely high and are good indicators for derivation of gabbroic magma from a peridotite mantle source (e.g., Wilson, 1989). The high content of Cr and Ni is in agreement with the REE pattern (Fig. 5.14) and negligible quantities of HFSE, particularly Zr (2 ppm) and Hf (0.1 ppm), which are classic incompatible elements, not readily substituted in major mantle phases. Samples from the quartz monzodiorite and granodiorite have variable low Cr and Ni contents but high in V (up to 316 ppm), Mn (up to 1490 ppm) and Zn contents (up to 90 ppm). A few anomalous samples from the quartz monzodiorite have slightly higher Cr and Ni contents (up to 62 ppm of Cr and 26 ppm Ni). On Harker diagrams, V, Mn and Zn elements decrease in concentration from the quartz monzodiorite to the granodiorite (Fig. 5.12). This behaviour is related to differences in the contents of hornblende, titanomagnetite and biotite. All these minerals variably decrease towards the higher Si02 contents. 5.5.5 Sr AND Nd ISOTOPES The Kuh Mish intrusions are isotopically different from the Bornavard granitoid (Fig. 5.9). These intrusions have a remarkably low initial 87Sr/86Sr (Table 5.1) and high positive 8Nd values. Low initial 87Sr/86Sr values of the Kuh Mish intrusions are somewhat sirnilar to the Kashmar granitoid (Fig. 5.9) but the positive end values are not matched. The gabbro has the lowest initial (at 42.8 Ma) 87Sr/86Sr (0.70386) and the highest 8Nd value (+8.02). These values are consistent with strong depletion of incompatible elements (particularly LREE) and enrichment of Sc, Cr and Ni in the gabbro. All the above features support the assumption that the gabbro was derived from a primitive mantle peridotite source. The initial (at 42.8 Ma) 87Sr/86Sr values for the Namin and Darin Granodiorites are 0.70388 and 0.70475, respectively. The corresponding £Nd values are +6.73 and +6.30, respectively. Sr-Nd isotopic data for the Kuh Mish intrusions suggest that the gabbro and granodiorites are genetically related. The initial (at 42.8 Ma) 87Sr/86Sr ratios for the gabbro (0.70386) and the Namin Granodiorite (0.70388) are indistinguishable within experimental error (2a = ±0.00005). R7 Rfi The Darin Granodiorite does not show a distinctly more radiogenic initial Sr/ Sr value than the Namin Granodiorite, suggesting that the Darin Granodiorite derived from the same parent as the gabbro. This is supported by similarity in eNd values in the Darin (+6.73) and Namin (+6.30) Granodiorites. The Kuh Mish intrusions do not show 87 86 significant positive correlations between Si02 and initial Sr/ Sr values (Fig. 5.9), nor between Si02 and eNd values. This indicates that magma mixing is not the case and these rocks may be derived from a single parent through simple fractional crystallisation (e.g. Soesso, 2000). The possibility of simple fractional crystallisation is supported by: (1) reasonable fractionation trends in some major and trace element abundances (Fig. 5.12); 118 (2) similar Sr-Nd isotopic systematics of the gabbro and granodiorites; and (3) a strong similarity of REE patterns and incompatible element ratios between the Namin and Darin Granodiorites (Figs. 5.13 and 5.14). Sometimes, linear trends of chemical elements can be produced by magma mixing of, for instance, mantle derived and crustal magmas (Collins, 1996, 1998). However, the isotopic similarities between the Kuh Mish intrusions preclude such mixing. The only possible explanation for the slightly less isotopically primitive feature of the Darin Granodiorite (initial 87Sr/86Sr = 0.70475 and ENd = +6.30) is that it may be the effect of very limited local crustal contamination or fractional crystallisation. It is interesting to note that the isotopically most primitive compositions in the Kuh Mish intrusions show similar characteristics to the Moruya I-type granite suite in central eastern part of the Lachlan Fold Belt (Keay et al, 1997). The Moruya suite contains a gabbroic-granitic component. The gabbro sample from the Kuh Mish intrusions and the on or microdiorite sample (BN11) from the Moruya suite are similar in Sr/ Sr (0.70388), and show ENd values of +8.02 and +8, respectively. The geological evolution of the Moruya suite is controversial. It is recognised as a type example of restite-controlled chemical variation by White and Chappell (1977) and Chappell et al (2000), while Keay et al. (1997) believed that the parent magma was derived as a mixture of mantle and crustal melts, followed by fractional crystallisation (Collins, 1996, 1998). But for the evolution of the Kuh Mish intrusions, restite or mixing models are not involved, because microgranular enclaves and isotopically evolved rocks are absent. Therefore, fractional crystallisation of a common mantle-derived magma is more likely responsible for generation of the Kuh Mish intrusions (e.g., Soesso, 2000). CHAPTER 6 GENETIC CLASSIFICATION AND COMPARISON WITH OTHER GRANITOIDS 6.1 INTRODUCTION Granitoid rocks have been subdivided on the basis of a variety of criteria by many authors (e.g., Ishihara, 1981; Petrik and Broska, 1994; Forster et al, 1997). A fundamental subdivision into I- and S-types was proposed by Chappell and White (1974) on the basis of petrographic and geochemical characteristics of the LFB granitoids, southeastern Australia. Later contributions (Chappell and White, 1984,1992; Chappell et al, 1998) confirmed the I- and S-types subdivision and proposed that I-type granites occur as two distinct groups, high- and low-temperature, based on the absence or presence, respectively, of inherited zircons. The high-temperature I-type granites are the most primitive and form by the partial melting of mafic rocks in the deep crust, or perhaps in modified mantle (Chappell et al, 1998), whereas the dominant group of low- temperature granites have been formed by partial melting of quartzofeldspathic crust at low magmatic temperatures (Chappell et al, 2000). The characteristics of the I- and S- type granites of the LFB have been widely used for the petrogenetic distinction of many granites around the world (e.g., Polard et al, 1995; Harris et al, 1997; Encarnacion and Mukasa, 1997; Sylvester, 1998; Moghazi, 1999). In this chapter, the most important characteristics of granites from the northeastern CIP are discussed to assign them according to the I- and S-type scheme. 6.2 FIELD AND PETROGRAPHIC EVIDENCE Granitiod rocks from the Kashmar, Bornavard and Kuh Mish areas are part of the Ururniyeh-Dokhtar Volcanic Belt that is a major Cretaceous to Recent geological structure in the CIP. In this belt, andesite is the most common igneous rock that has been generated by subduction of the Tethyan Oceanic crust beneath the CIP (Hassanzadeh, 1993; Alavi, 1994; Moradian, 1997). At most places in this belt, plutonic and volcanic rocks are broadly similar in mineralogy since the igneous protoliths are probably similar. Granitic plutons of the northeastern CJP (present study) commonly have mineral assemblages of quartz + plagioclase + K-feldspar + biotite ± hornblende. In the terminology of Chappell and White (1992), this uniform mineral assemblage is a feature of metaluminous I-type sources. When present, xenoliths in rocks from the current study are commonly hornblende-bearing microgranular varieties (Section 4.1.2). The minerals present in the xenoliths match those of the enclosing rocks (e.g., Sample Rl 5912). 6.3 MINERALOGICAL EVIDENCE Granitoid rocks of the northeastern CJP lack Al-rich minerals such as cordierite, andalusite, sillimanite and garnet, Muscovite is rare, but it is fine-grained secondary and formed by subsolidus alteration. It sometimes occurs in granite from the Bornavard granitoid. The K-feldspar crystals are frequently pale pink in colour, and less commonly white, reflecting high/02; characteristic of I-type granites (Chappell and White, 1992). The mafic minerals are magnesio-hornblende, biotite and Fe-Ti oxides. In the Kashmar and Kuh Mish areas, biotite coexists with hornblende and contains lower total Fe as FeO content than biotite from the Bornavard granitoid (Appendices 3.3; 3.9 and 3.15). The presence of two OH-rich phases indicates a relatively high H20 content in the magma (Peacock et al, 1994), consistent with the more homogeneous nature of biotite grains from the Kashmar and Kuh Mish areas. Burkhard (1991) stated that granitic melt will cross the saturation point of H20 during cooling at a temperature that depends on the amount of H20 dissolved in the melt. Once this point is crossed, an increasing amount of H20 will be set free during further cooling and PH20 and/02 will increase. Increase m.j02 allows the crystallisation of magnetite, i.e. partitioning of Fe into the oxide rather than into biotite. Therefore, lower total Fe as FeO in biotite is attributed to higher PH20 andy02 of the Kashmar and Kuh Mish magmas, compared with the Bornavard magmas. Biotite grains from the Kashmar and Kuh Mish areas are high in Ti02 content (up to 4.90 wt%) and contain negligible amounts of A1VI (mostly <0.1 a.f.u.). This is characteristic of biotite coexisting with hornblende in I-type granites (Chappell and White, 1992). In the Bornavard granitoid, where biotite coexists with secondary muscovite, it contains appreciable A1VI (0.43-0.89 a.f.u.) in the structural formulae and has a lower Ti02 content. This is not characteristic of the source because of subsolidus alteration. Biotite from the Bornavard area sometimes coexists with allanite and shows a wide range in Fe/(Fe + Mg), both are typical I-type features (Whalen and Chappell, 1988; Petrik and Broska, 1994). Biotite in all granitoid rocks from the northeastern CIP is, however, typically high in FeO/MgO and shows the distinctive pleochroic scheme of oxidised I-type granites (X = Y = reddish to dark brown, Z = straw-coloured). When present, ilmenite commonly occurs with hornblende. Ilmenite without magnetite is less common in granitoids of the northeastern CIP. Microprobe data show that ilmenite and titanomagnetite preferentially occur as cores of Fe-Ti oxide grains while magnetite occurs in the rim or as single grains (Appendices 3.10 and 3.15). These features indicate evolution towards higher f02 (Whalen and Chappell, 1988; Blevin and Chappell, 1995). Titanite usually replaces hornblende and biotite, however large euhedral titanite grains seem to be primary (e.g., Sample R15900). Overall, the chemical characteristics of the ferromanesian minerals and presence of accessory phases such as titanite, apatite and allanite suggest that the granitoids of the northeastern CJP are I-type (e.g., Chappell and White, 1992; Pollard et al, 1995). 6.4 EVIDENCE FOR RESTITE Microgranular enclaves are common in the Kashmar granitoid, particularly in the tonalite, granodiorite and granite. They are mostly dark grey to black in colour and occur as angular or spherical shapes 3-4 cm in diameter. The enclaves are mainly more mafic and finer grained than the host rocks. They have an igneous texture. They show solid state reactions on their margins and become less common and smaller in size as the host-rock increases in Si02 content. According to the modal analyses, the enclaves are microtonalites rich in plagioclase, biotite and hornblende. Magnetite (5.4%), apatite (0.8%), titanite (0.2%) and zircon (<0.2%) are the only accessory minerals occurring in the microgranular enclaves as well as in the host-rocks. The apatite occurs as tiny needles or prisms concentrated in plagioclase crystals, and less commonly in biotite and hornblende. These occurrences suggest that the apatite is restite, crystallised at low- temperature I-type melts (Chappell et al, 1987). In contrast, apatite occurs as large squat prisms where it has precipitated from high-temperature I-type melts as in the Toulumne Intrusive Series, California (Bateman and Chappell, 1979; Beams, 1980). One sample (R15912) of microtonalite enclave taken from granodiorite has been chemically analysed. This enclave is high in Na20 (4.33 wt%), CaO (5.13 wt%), Ba (455 ppm) and Sr (367 ppm) and low in Rb (45 ppm) contents that suggests an I-type source (e.g., White et al, 1999). The abundances of most major and trace elements in the enclave are J.ZO similar to those in tonalite and granodiorite (Appendix 4.1). The chemical composition of the enclave reflects an ASI = 0.87 that is metaluminous. Many granitic rocks from elsewhere contain darker coloured enclaves, typically more mafic in composition (usually tonalitic to dioritic) than their hosts. For example in some granitic suites of the LFB the enclaves reflect mechanical interaction (mingling) and partial chemical hybridisation (mixing) of basaltic to andesitic magmas with host granitoids, and suggest the involvement of mafic mantle-derived magmas in granite petrogenesis (Gray, 1984, 1990; Keay et al, 1997; Collins, 1998). Variation in isotopic signatures (Sr, Nd and Pb) of such granitic suites is compatible with an origin involving contrasted crustal and mantle source components (Collins, 1996; Keay et al, 1997). In two-component mixing model, mafic rocks and enclaves have primitive, mantle-like initial 87Sr/86Sr and ENd values, however in some cases the isotopic mixing arrays, do not match the predicted trace element mixture (Soesoo and Nicholls, 1999; Soesoo, 2000). In the Kashmar granitoid, variation in initial 87Sr/86Sr is low (0.70471-0.70569) and all ENd values are negative and show a limited range (-0.70 to -1.86). These data suggest that possibly lower crust is the source for the Kashmar granitoid. Therefore, origin of the Kashmar granitoid by two-component mixing model is unlikely, and microgranular enclaves are not fragments of solidified mantle-derived magmas. This is supported by enrichment in most of the incompatible elements and similarity in rock/primordial mantle normalised patterns and the REE abundances in different rock types of the Kashmar granitoid. Assuming that compositional variation of the Kashmar granitoid was solely resulted by fractional crystallisation of a mafic melt (e.g., Wyborn et al, 1987; Soesoo, 2000), then 124 the parental magma would be completely or largely molten, high in temperature and free of restite (e.g., Chappell et al, 1998). At high temperature, crystals of zircon are not initially present in the melt because the melt is undersaturated in zircon (King et al, 1997; Chappell et al, 2000). But in the Kashmar granitoid, zircon grains mostly coexist with hornblende or biotite and commonly occur in mafic and felsic rocks as well as in microgranular enclaves. Most zircon grains are similar in morphology and size, some have a very narrow black rim. In each pluton of the Kashmar granitoid, inflexion in Zr versus Si02 contents is not observed that may imply zircon was not homogenised in the magma, therefore the melt was low in temperature and zircon may be a restite phase (e.g., Chappell, 1998c). In addition, some uniform plagioclase cores that show discontinuity in their outer rims (e.g., Samples R15908 and R15918) may be interpreted as restite phase (e.g., Chappell, 1996b). Also, in the Kashmar granitoid, most major and trace element data show negative correlation with Si02 contents, particularly decrease in the contents of P2Os suggests lower solubility of P in more felsic and lower temperature granite melts (Harrison and Watson, 1984). These characteristics are invoked to explain compositional variations of the Kashmar granitoid may be generated by fractional crystallisation and removal of restite from a low-temperature melt compositions (e.g., Chappell et al, 1987). Therefore, microgranular enclaves of the Kashmar granitoid may be restite. 6.5 CHEMICAL COMPOSITIONS The granitoid rocks of northeastern CJP are generally high in Na20, K20, CaO, Ba and Sr. These elements are distinct because they occur in feldspars. High content of these elements suggest I-type characteristics. According to White and Chappell (1977) and Chappell and White (1992) high total alkalis (Na20 + K20) relative to A1203 is 125 characteristic of granites derived from source rocks that would not have undergone weathering or alteration processes. These characteristics are retained during production of I-type granites. In northeastern CIP, granitoid rocks are low in Rb content but Cr, Ni, Pb and Sn are very low that suggest the source rocks were previously fractionated (Chappell and White, 1992). Within and between each rock type, particularly in the Kashmar and Bornavard granitoids, with increasing Si02, there are marked increases in the contents of K20, Rb, Th, EREE, and the magnitude of the Eu anomaly, while the transition metals and Sr decrease in abundance. These features may be attributed to fractional crystallisation in I- type granites. The process of fractional crystallisation is supported by observation of inflexion point on Harker diagrams for Ba, Zr and Zn elements from granodiorite of the Bornavard granitoid (Sections 5.3.2.2 and 5.3.3). Chappell and White (1992) point out distinctive chemical contrasts between I- and S- type granites of the LFB that are manifested during fractionation. For example, P205 decreases in I-types and increases in S-type granites, Th and Y increase in I-types but remaining fairly constant in the S-type granites. The contents of La and Ce do not change in I-type but decrease markedly in S-type granites. According to Harker plots (Fig. 6.1) for all felsic rocks (Si02 >63 wt%) in the Kashmar and Bornavard granitoids, P205, A1203 and total Fe as Fe203 markedly decrease towards higher silica contents. Decrease in P205 means that P was saturated in the melt and behaved compatibility during fractional crystallisation, implying that the the Kashmar and Bornavard granitoids are I-type (e.g., Pollard et al, 1995; Chappell, 1998b). The concentrations of some incompatible elements such as Ce, Th, La, and Y are illustrated in Figures 6.2. The concentration of Th in rocks of the Kashmar and Bornavard granitoids shows a typical positive correlation with increasing Si02 contents, again supporting the I-type characteristics (Chappell, 1998b). In the Kuh Mish intrusions, Ce, Th, La and Y are relatively constant and strongly depleted. This feature is consistent with primitive isotopic signature of the Kuh Mish intrusions. The concentrations of the above HFSE in the Kashmar and Bornavard granitoids are high and variable. However, above 74 wt% Si02 contents corresponding to crystallisation of alkali feldspar granite and granite, respectively from the Kashmar and Bornavard granitoids, La, Ce and Y increase dramatically. Because alkali feldspar granite and granite are the most quartzofeldspathic rocks of the present study, the strongly incompatible behaviour of HFSE suggests no Y- bearing accessory phases (e.g., fluorite) were being fractionated (King et al„ 1997). Such behaviour in REE concentrations implies I-type characteristics (Chappell, 1998b). 6.6 ALUMINUM SATURATION INDEX (ASI) The aluminum saturation index (ASI = molecular Al203/[CaO + Na20 + K20]) is one of the most useful chemical discriminant between peraluminous and metaluminous granitoid rocks (Zen, 1988; Barbarin, 1996; Chappell, 1998b). The S-type granites are always saturated in Al (ASI >1) so they are peraluminous. Since a degree of Al- oversaturation is an intrinsic property of most felsic granite melts, I-type granites may be either metaluminous or weakly peraluminous (e.g., Nakajima, 1996). In particular the most felsic I-type granites are dominated by quartz and feldspars (haplogranite, ASI = 1), hence their ASI converges to values close to one (Chappell, 1998b). In the present study, the boundary between metaluminous and peraluminous terms is not exactly defined at ASI = 1 because the CaO value is not corrected for apatite. Due to the absence of minerals being more peraluminous than biotite, rocks with ASI = 1 to -1.1 127 are known as 'weakly peraluminous I-type' (e.g., Chappell, 1984; Miller, 1985; Nakajima, 1996; Chappell, 1998b). Calculated ASI values for rocks of the northeastern CJP are shown in Appendices 4.1 to 4.4. Figure 6.3a shows that most analyses are metaluminous. The peraluminous feature is attributed to most Samples from the Taknar Rhyolite and few samples from granitoid rocks. For all analyses the molar proportion of total alkalis (Na20 + K20) is less than A12Q3 (in moles) that indicates that these rocks are not A-type (cf. Wormald and Price, 1988; Whalen et al, 1996). With increasing Si02, the ASI values generally increase but some scattering is observed (Fig. 6.3b). The positive correlation between ASI and Si02 is consistent with I-type metaluminous features (Chappell, 1998b). The scattering results from different modal contents of biotite, the only peraluminous mineral that occurs in these granitoid rocks. In the Kashmar granitoid, about two thirds of samples have ASI <1. With the exception of sample R15906, others stay at ASI values close to the unity (Appendix 4.1). The average ASI value for 29 analyses of the Kashmar granitoid is 0.97 that suggests typical metaluminous I-type characteristic. This feature is consistent with mineralogy (Section 4.2) and low initial 87Sr/86Sr values reported for the Kashmar granitoid (Section 5.2.5). In particular, the metaluminous nature of the Kashmar granitoid is reflected by the appearance of CIPW normative diopside (Di) that ranges from 0.0 to 8.4%. In most samples of the Kashmar granitoid normative corundum (Q ranges between 0.00 and 0.82% (Appendix 2.1). Some samples contain normative hypersthene (Hy) (average of 7.1% on 29 analyses). Only one sample from granite (R15906) is weakly peraluminous (ASI =1.15) and contains -2.5 % normative C. This sample is i!r IZS quartzofeldspathic with high Na20 (4.68 wt%) and low CaO (1.33 wt%) contents but its feldspars are commonly serialised. It seems that sericite increases normative C because except biotite other aluminous minerals are not present in this sample. Samples from the alkali feldspar granite have ASI = 1 to 1.05. These values are slightly higher than ASI values for granite. The higher ASI values of alkali feldspar granite, is consistent with the presence of quartz and K-feldspar as the major mineral components of this rock. The Si02 contents of alkali feldspar granite range from 74 to 77 wt%. It also contains minor amount of fresh euhedral biotite, and approximately equal amounts of normative quartz (0, albite (Ab) and orthoclase (Or). All the above features are similar to the composition of low temperature hydrous silicate melt in equilibrium with quartz and feldspar; i.e. haplogranite composition. According to Chappell (1998b), the composition of haplogranites is usually weakly peraluminous and produces an overlap between ASI values of I- and S-type granites. To a first approximation, haplogranites have compositions that are independent of their precursor materials and separation into I- or S-type, is often difficult. But in the Kashmar granitoid, the alkali feldspar granite is part of a suite that extends to more mafic compositions with ASI values less than one. This is conformed by similarity in initial 87Sr/86Sr and ENd values of alkali feldspar granite and other plutons of the Kashmar granitoid. Therefore, weakly peraluminous characteristic of alkali feldspar granite is different with those haplogranites derived from S-type origin. The ASI values for Bornavard granitoid range from 0.64 to 1.11 (Appendix 4.2) with an average of 0.95 that is very similar to the average ASI values of the Kashmar granitoid. With the exception of Sample R15955 (normative C = 1.30%), other samples from the Bornavard granitoid (Appendix 2.2) are low in normative C (0.00-0.77%), emphasising the metaluminous feature. Samples from tonalite are strongly metaluminous (ASI = 0.64-0.67) as they contain modal abundances of hornblende. Most samples from the granodiorite are weakly metaluminous (ASI = 0.84-0.98). But the ASI values for granite samples range from 1.04 to 1.11, indicating weakly peraluminous characteristic. This is due to the appearance of secondary muscovite in most of the granite samples (Appendix 2.2). For example the highest ASI value (1.11) for granite is observed in Sample R15955 because it contains the highest modal content of secondary muscovite (9.0% modal). The granite from Bornavard granitoid is a very felsic rock (Si02 = 74- 76wt%) and high in Ba contents, varying between 580 and 890 ppm that may be resulted from fractional crystallisation. The granite compositionally plots close to the Tuttle and Bowen (1958) minimum-temperature melt composition (Fig. 6.4), which shows that the granite represents liquid composition, resulted from the dominant role of fractional crystallisation (e.g., Chappell, 1998b). Water is one component that would be concentrated by such extreme fractionation. In Figure 6.4, the curves for water-saturated liquids in equilibrium with quartz and K-feldspar at confining pressures of 0.5 and 3.0 kb are also shown (Tuttle and Bowen, 1958). The granite samples are clustered near the isobaric line of 0.5 kb, consistent with the general lack of primary muscovite. In the system KAlSi308-Si02-H20, muscovite crystallisation reflects PH20 exceeding 5 kb (Sun and Chen, 1992). Therefore, low PH20 of the granite (0.5 kb) suggests that muscovite in the granite samples is not primary. Hence, the weakly peraluminous feature of the granite resulted from its quartzofeldspathic nature. Also, the granite contains biotite, titanite, magnetite and allanite, the last three minerals are typical accessories in I-type granites (Pollard et al, 1995). 13U For most samples of the Taknar Rhyolite, the ASI ratio is higher than 1.1 and shows peraluminous to strongly peraluminous features (Appendix 4.3). This characteristic is different with the granite form the Bornavard granitoid, consistent with differences in age, mineralogy and concentrations of most incompatible elements (Section 5.4.1), all indications that the granite and rhyolite are not genetically related. The extremely high value of the ASI (2.46) in Taknar Rhyolite is related to Sample R15949 that is very high in normative C (6.8%). This sample is depleted in whole rock Na20 content (0.37 wt%) as all feldspars in the groundmass are strongly sericitised that is a major contributor to the very high C content. It must be noted that the ASI values are very sensitive to the effect of hydrothermal alteration. This process can lead to the destruction of feldspars and the mobilisation of Ca, Na and K, with consequence increase in ASI value. As the rhyolite samples are quartzofeldspathic in nature, it is expected to show ASI values close to one (e.g., Chappell, 1998b), but hydrothermal alteration may be responsible for peraluminous feature of the Taknar Rhyolite. The Kuh Mish intrusions mostly range from strongly metaluminous to weakly peraluminous (ASI = 0.55 to 1.13), typical of I-type granites. This feature is consistent 87 86 with low initial Sr/ Sr (0.70386-0.70475) and positive £Nd values (+6.30 to +8.02) of these rocks. Only two samples of granodiorite, R15935 and R15931 are slightly peraluminous (ASI = 1.13 and 1.25). They contain normative C of 1.69 and 0.38%, respectively. Both samples have higher Na20 contents (5.63 and 4.05 wt%, respectively) than strongly metaluminous rocks of the Kuh Mish intrusions. They are felsic and dominated by quartz and feldspars. They do not contain a mineral more aluminous than biotite. Their normative C is not unusual because felsic I-type granites may contain more than 1% C and ASI value >1.1 (Nakajima, 1996: Chappell et al, 1998; Chappell, 1998b). Collectively, on the ASI frequency histogram (Fig. 6.5), igneous rocks of the northeastern CJP have a distribution between strongly metaluminous to weakly peraluminous values that is a very similar pattern to that shown by 1025 analyses of I- type granites of the LFB, Australia (Chappell, 1998b). The Kashmar and Bornavard granitoids show similar distributions in ASI values (Fig. 6.5). The average ASI values for the Kashmar and Bornavard granitoids are 0.97 and 0.95, respectively (Table 6.1). Both distributions are approximately symmetrical and centred at an ASI equal to one, with a range between 0.8 and 1.2. This implies that the Kashmar and Bornavard granitoids are weakly metaluminous to weakly peraluminous and this is characteristic of low-temperature I-type granites (Chappell, 1998b). The mean ASI value for 59 analyses of igneous rocks of the northeastern CJP is 0.97 and the median value is 0.99 with 90% of the values being greater than ASI = 0.80, so that the studied rocks are not cumulates. The ASI values confirm that most of the analyses, particularly those of the weakly peraluminous, have a composition of hydrous silicate melt in equilibrium with quartz and feldspar. Most of the weakly peraluminous rocks are strongly felsic (Si02 >74 wt%) and lacking in microgranular enclaves. They may have been formed at low temperature when only the felsic components of the source rocks were fused, or by fractional crystallisation (e.g. Chappell and White, 1992; Chappell, 1998b). PP; 132 6.7 Sr AND Nd ISOTOPES Isotopic ratios are potentially useful indicators of granite classification because granites compositionally image their source rocks in the deep crust (Chappell, 1994; Chappell et al, 1998). For I-type granites the range in initial 87Sr/86Sr is from 0.703 to 0.712 and for eNd from + 8.4 to - 7.2 (e.g., Chappell and White, 1992; Keay et al, 1997). For S-type granites in the LFB, the corresponding values are 0.708 to 0.720 and -5.8 to -9.2 (Chappell and White, 1992; Sun and Chen, 1992; Raymond, 1995). 87 86 In the Kashmar granitoid initial Sr/ Sr (0.70471 to 0.70569) and £Nd values (-0.70 to -1.86) are typically within the range of Sr-Nd isotopic signature of I-type granites. For 87 86 Kuh Mish intrusions both initial Sr/ Sr (0.70386 to 0.70475) and eNd values (+6.30 to R7 Rfi +8.02) are not evolved. These initial Sr/ Sr values are the lowest extreme of the I-type range and the ENd values are similar to more primitive granitoids (e.g., Keay et al, 1997). Also, the initial 87Sr/86Sr values of the Kuh Mish intrusions are very similar to Mt Buller I-type suite in the southeastern LFB (Soesoo, 2000) that shows granites and mafic rocks are likely to have formed by fractional crystallisation from a common mantle-derived magma. In the Bornavard granitoid, the initial 87Sr/86Sr values are variably high. For example tonalite and granodiorite have initial 87Sr/86Sr (0.70757 to 0.72153) that fall within the range of values shown in S-type granites. The granite shows distinctively more radiogenic initial 87Sr/86Sr values, 0.73622-0.75008. But the eNd values (-1.41 to -5.20) from all plutons of the Bornavard granitoid are within the range of I-type granites. Except for initial ratios, other petrographic and geochemical data for the Bornavard granitoid confirm I-type characteristics. As mentioned in Section 5.3.4, the initial 87Sr/86Sr values of the Bornavard granitoid would be the result of contribution of older continental crust in the source. Such isotopic values suggest that the Bornavard granitoid may be an I-type contaminated granitoid (e.g., Ague and Brimhall, 1988b; Yui et al, 1996). 6.8 ALLOCATION OF GRANITOIDS TO SUITE An important feature of the study of granitoid rocks in Australia has been the recognition of sets of plutons that can be often grouped into suites. Such grouping is based on the shared similarities of field, petrographic and compositional data (Griffin et al, 1978; Hine et al, 1978; Chappell, 1984; Chappell and White, 1992; Blevin and Chappell, 1995; White et al, 2000). Granites from a particular suite may be uniform or varied in chemical composition but they should have distinctive properties, reflecting similar features of their source rocks. The allocation of granites to suites is fundamental to understanding their petrogenetic concepts (e.g., Whitten, 1991). Granites, which might be grouped together in suites except for small compositional differences that preclude such a precise grouping, are placed in supersuites (White et al, 2000). In the Kashmar granitoid, rocks of different plutons are fairly equigranular, homogeneous in appearance and similar in age (-42.5 Ma). Hornblende, biotite, magnetite and titanite are common minerals. The abundance of most major and trace elements (e.g., P205, total Fe as Fe203, Sr and V) are well correlated (linear with smooth trends). They are high in Na20 (mostly between 3 to 5 wt%), Mn (up to 1070 ppm), Ba (-500 ppm, on average), and low in Ti02 (all <1 wt%), Rb (all <210 ppm), Cr (mostly <10 ppm), Ni (-2 ppm) and Sn (all <5 ppm). These plutons are enriched in LREE, flat 87 86 in HREE and low in initial Sr/ Sr (0.704 to 0.705) and the ENd values (-0.70 to -1.86). The small but significant isotopic differences suggest that each pluton represents independent evolution from similar, but not identical parent magmas (e.g., Rapela and Pankhurst, 1996). The above similarities and common metaluminous I-type characteristic (ASI = 0.81 to 1.05) precisely suggest a 'simple suite' for the Kashmar granitoid. Using modern nomenclature of granites, 'simple suite' corresponds fairly closely to the low-temperature granites of Chappell et al (1998). In the Bornavard granitoid, rocks of different plutons are metaluminous (ASI <1) or weakly peraluminous (ASI = 1 to 1.1) that is similar to the Kashmar granitoid. Typical I- type features include the occurrence of hornblende, biotite, magnetite, titanite and allanite (e.g., Ague and Brimhall, 1988a; Petrik and Broska, 1994). The weakly peraluminous feature is only attributed to the granite that is lacking in hornblende. In general, rocks of the Bornavard granitoid are high in Na20, Ba, Zr, Mn, and low in Ti02, Rb, Pb, U, Nb, and most transition metals (Appendix 4.2). These rocks share in most petrographic and chemical characteristics with the Kashmar granitoid (Figs 6.1 and 6.2). Both granitoids show similar patterns for incompatible elements. The content of Rb is lower than Sr for all rock types with Si02 <74 wt% whereas Rb is higher than Sr in rocks with Si02 >74 wt%. Isotopic data for the Kashmar and Bornavard granitoids suggest crustal source. The compositions of granite from the Bornavard granitoid and alkali feldspar granite from the Kashmar granitoid represent typical examples of low- temperature I-type melts. The above similarities suggest that the Kashmar and Bornavard granitoids can be assigned as a 'simple suite' in the Taknar Zone. However, the Bornavard granitoid is Late Jurassic Early Cretaceous in age and the Kashmar granitoid is Middle-Late Eocene that is significantly younger. But age is not used as part of the recognition of a suite (White et al, 2000). Isotopic data show that rocks of the Bornavard granitoid are more evolved than the Kashmar granitoid but the isotopic composition of a granite suite may be varied (e.g., Bullenbalong suite, Chappell et al, 5 1999). Because the Kashmar granitoid represents typical 'simple suite' and is compositionally similar to the average composition of the Bornavard granitoid (Table 6.1), the name of the Kashmar suite is introduced as a fundamental lithological concept for granite petrology in the Taknar Zone. The Kuh Mish intrusions show some differences with the Kashmar and Bornavard granitoids (Fig. 6.2). The Kuh Mish intrusions are lower in most of the incompatible elements particularly Ba, Rb, Sr, Pb, Th, U and REE (Appendix 4.4). They are high in some transition metals such as V, Cr, Mn and Ni. Isotopic data show that mafic and felsic rocks of the Kuh Mish intrusions have primitive, mantle-like initial 87Sr/86Sr and End values (Table 5.1). Differences between the composition of rocks of the Kuh Mish intrusions, with the Kashmar and Bornavard granitoids preclude allocation of all the igneous rocks of the northeastern CJP into a 'simple suite'. The Kuh Mish intrusions form a magmatic suite that magmas are mantle-related. Whereas, the Kashmar and Bornavard granitoids form a magmatic suite that magmas are crastal-related (indirect- mantle). 6.9 HIGH- AND LOW-TEMPERATURE I-TYPE GRANITES Recently, new and conclusive information provided by Chappell (1998a,b) and Chappell et al (2000) has led a fundamental subdivision of I-type granites into two groups, formed at 'high- and low-temperatures'. The subdivision was deduced specifically from the criteria of zircon age inheritance and the abundance of Zr and its pattern of variation in Harker plots. The high-temperature I-type granites formed by partial melting of mafic source rocks at completely or largely molten state while the low-temperature I-type 13b granites formed by partial melting of quartzofeldspathic rocks such as older tonalites (Chappell etal, 1998). There are some indications that the Kashmar and Bornavard granitoids have been formed at low magmatic temperatures. Although, contacts with the country rocks are often steeply inclined but the absence of significant thermal aureole around plutons indicates that magmas were possibly low in temperature. Variations in concentrations of Ba and Zr for most plutons indicate that they did not form progressively as cumulates from an originally liquid or largely liquid magma (e.g., Chappell, 1996b). This is supported by presence of microgranular enclaves that were interpreted as restite in the Kashmar granitoid. The presence of zircon grains in mafic and felsic plutons of the Kashmar and Bornavard granitoids suggests that zircon possibly presented in the magmas, hence temperature was low (e.g., King et al, 1997). But in the Kuh Mish intrusions, zircon is absent, possibly because zircon was not in the source or magmas were at higher temperature. Alkali feldspar granite from the Kashmar granitoid and granite from the Bornavard granitoid are dominated by quartz and feldspars. These rocks contain microcline that is typical low-temperature K-feldspar in plutonic environments. According to Chappell (1998b), such quartzofeldspathic granites may be produced by partial melting of the crust at low magmatic temperature or by fractional crystallisation from mantle-derived magmas. The latter alternative is not indicated in the Kashmar and Bornavard granitoids, because they are not associated with cumulates to suggest fractionation of mantle-derived magma and also isotopic data show crustal sources. In the Kashmar and Bornavard granitoids, the content of P205 is low (all <0.34 wt%) and decreases to negligible amount in granite and alkali feldspar granite, suggesting low solubility of P in low-temperature I-type granites (Harrison and Watson, 1984). The normal distribution of the ASI values with a mean value close to the unity (0.97) and low deviation (o = 0.1155) indicate that the Kashmar and Bornavard granitoids are generally neither strongly metaluminous nor strongly peraluminous. Hence, they are not cumulates and they have not been crystallised from completely molten magmas. Whereas in high-temperature I-type granites (e.g., Boggy Plain and Marulan suites of the LFB), the ASI values are composite with a wide range of values (0.18 to 1.07), relative to low-temperature I-type granites (Chappell and White, 1992; Chappell, 1998b). The above features are consistent with the presence of calcic hornblende with low contents of A1203 and Ti02 (e.g., Section 4.2.2) that are characteristics of amphibole from low-temperature magmas at high^02 (Mason, 1978; Hammarstrom and Zen, 1986; Hollister^ al, 1987). 6.10 COMPARISON WITH OTHER GRANITOID TYPES In Table 6.1 the average geochemical data from igneous rocks of the northeastern CJP are compared with the averages of the three major granite types (I-, S- and A-type). The main reason for comparison with the Mashhad Granite is because the last two phases of the Mashhad granite which are biotite/muscovite granites and pegmatites, respectively, are known as the only peraluminous S-type granites (Section 2.5.2.1) that reported from northeastern CJP (Iranmanesh and Sethna, 1998). Comparison with A-type granites is because Esmaeili et al (1998a) stated that the Bornavard granitoid is A-type in character. 6.10.1 COMPARISON WITH S-TYPE GRANITES Compared with averages of 18 analyses of the S-type granites from Mashhad, Iran and 704 analyses of S-type granites from the LFB (Table 6.1), the most obvious uts .. compositional differences are the higher Sr, CaO, MgO and total Fe as Fe203 and lower Pb and Rb contents of KBTK (Kashmar, Bornavard, Taknar, Kuh Mish). In particular, rocks of the KBTK are significantly higher in Na20 contents than S-type granites of the LFB. These features are fundamental and indicate that the sources of magmas in northeastern CJP have not been previously weathered, hence they are I-type. In contrast, S-type granites are low in Na, Ca, Mg and Sr but high in K20 and Rb because prior weathering of their supracrustal source, particularly involves the loss of key elements, Na, Ca and Sr in solution (Chappell, 1996b; Chappell et al, 1998; 2000). Both Mashhad and the LFB S-type granites are higher in Si02 contents than granitoids of the northeasten CJP. The reason is because S-type granites are in many ways analogous to the low-temperature I-type granites that were generated initially by partial melting of quartzofeldspathic rocks in the crust (Chappell et al., 2000). The content of total Fe as Fe203 is significantly higher in granitoids of northeastern CJP, compared with S-type granites from Mashhad and the LFB. This feature is consistent with modal abundances of magnetite, and indicates that I-type granites of northeastern CJP axe high mf02, while S-type granites are low in f02 because they are mostly ilmenite-bearing and contain graphite in their source rocks (Chappell et al, 1998). 6.10.2 COMPARISON WITH A-TYPE GRANITES Using XRF data, Esmaeili et al. (1998a) stated that the granite from Bornavard granitoid is peraluminous, sub-alkaline, lacking in OH-bearing minerals, and formed in anorogenic environment, hence it is A-type granite. Such arguments for this granite are in contrast with the petrographic and geochemical data that were obtained by the present study. According to modal petrography (Appendix 2.2), quartz, K-feldspar and plagioclase are major mineral components of the granite. In most granite samples, minor 139 biotite is conspicuous and may form up to 8.4% by volume of the rock (Appendix 2.2). Biotite is the only primary aluminous mineral in the granite samples. The biotite grains contain Mg/(Mg + Fe) = 0.13 to 0.19 and shows typical pleochroic scheme of oxidised I-type granites (Section 4.4.4). Whereas in A-type granites, annite (Fe/Fe+Mg = -0.00), Fe-rich amphibole, and fluorite are common (Collins et al, 1982; King et al, 1997). Analyses of Fe-Ti oxides in granite represent magnetite, and only two core composition as titanomagnetite (Appendix 3.10), indicating oxidised conditions, while the A-type granites are relatively reduced as they have primary ilmenite and Fe-rich amphiboles (Wones, 1989; Weaver et al, 1992; Landenberger and Collins, 1996: King et al, 1997). Such differences preclude A-type feature for the granite of the Bornavard granitoid. The ASI values for most granite samples are close to the unity indicating sub-aluminous (e.g., Hess, 1989), reflected from quartzofeldspathic nature of the granite. Only one sample (R15955) is weakly peraluminous (ASI = 1.11) because it contains 9% modal abundance of secondary muscovite. Therefore, peraluminous feature that proposed by Esmaeili et al. (1998a) is not valid. Most of the A-type granites from elsewhere are peralkaline (Wormald and Price, 1988; Sylvester, 1989; Eby, 1990; Poitrasson et al, 1995; Whalen et al, 1996), but for all analyses of the Bornavard granitoid, the molar percent of A1203 is less than (Na20 + K20). For this reason, the Bornavard granitoid is compared with 43 analyses of ahiminous A-type granites of the LFB (Table 6.1). The LFB A-type granites are characterised by higher K20 and lower MgO, CaO, Fe203 and A1203 contents compared ,+ h the Bornavard granitoid (Table 6.1). Compositional similarities include high Na20, Y, REE and low MnO and P205 contents. Because A-type granites overlap in some major element compositions to very felsic I- type granites (Chappell and White, 1992), distinctive differences between the Bornavard granitoid and A-type granites become clear when trace elements are considered. The obvious feature of the LFB A-type granites is much higher abundances of HFSE (Zr, Nb, Y, La, Ce), LFSE (Rb, Pb, Th) and transition metals (Zn and Ga) (Table 6.1). The higher abundances of HFSE suggest that A-type granites are higher in temperature. This is consistent with experimental studies that show partition coefficient of HFSE particularly Zr increases with increasing temperature (King et al, 1997; Chappell et al, 1998). But lower contents of HFSE particularly Zr indicate that the Bornavard granitoid may be formed at lower temperature than the A-type granites. The ternary plot of normative Q-Ab-Or (Fig 6.4) for the granite samples supports low temperature crystallisation of the granite magma. The lower concentrations of La, Ce and Y in the Bornavard granitoid suggests that no Y-bearing minerals particularly fluorite was crystallised, whereas fluorite commonly occurs in A-type granites (Clemens et al, 1986; Eby, 1990; Landenberger and Collins, 1996). The above characteristics suggest that the presence of A-type granite in the Bornavard granitoid is unlikely. 6.10.3 COMPARISON WITH I-TYPE GRANITES Compared with subvolcanic I-type intrusions from Natanz, Iran (Table 6.1), the Kashmar and Bornavard granitoids are lower in average content of A1203, total Fe as Fe203, MgO and CaO but higher in Si02 and K20 contents. Differences in chemical composition are consistent with isotopic data that suggest magmas of the Kashmar and Bornavard granitoids are more evolved and derived from igneous rocks of the crust (Sections 5.2.5 and 5.3.4), while magmas of the Natanz intrusions are less evolved and derived from mantel source (Berberian, 1981). The Kuh Mish intrusions are similar to 141 the Natanz intrusions in some major element data such as total iron, MgO, CaO and Na20 contents. This similarity is consistent with primitive isotopic features of the Kuh Mish (Section 5.5.5) and Natanz intrusions. The average content of Na20 from igneous rocks of the northeastern CIP is higher than average Na20 content of 1074 analyses of Early Palaeozoic I-type granites of the LFB (Table 6.1). The lower Na20 contents of igneous rocks of the LFB may be related to different tectonic settings (Chappell, 1994) as the bulk of the I-type granites of the LFB formed at low magmatic temperatures and involved the partial melting of older quartzofeldspathic crust (Chappell et al, 2000). The higher Na20 content of igneous rocks of the northeastern CJP is comparable with average Na20 contents of I-type granites from different magmatic provinces of the Peninsular Ranges Batholith (PRB, Table 6.1). Chemical and isotopic properties of I-type granites of the PRB indicate a subduction-related environment (Silver and Chappell, 1988). In the present study, the Kuh Mish intrusions are typically low in total alkalis and high in CaO, Cr and Ni contents. These features are very similar to I-type granites of the WPRB and indicate possibly involvement of subduction in generation of the Kuh Mish intrusions. This may be supported by large volume of andesitic to basaltic rocks, and an extensive ophiolitic belt with island arc character that occurs to the northern parts of the Kuh Mish intrusions (Ghazi and Hassanipak, 1999). Similar to the average chemical composition of I-type granites of the LFB, the Kashmar and Bornavard granitoids are high in average contents of Si02 (66.80 and 69.84 wt%), Ba (494 and 516 ppm), Sr (249 and 90 ppm) and Zr (175 and 234 ppm, respectively). Also, the Kashmar and Bornavard granitoids are high in most of the REE contents such as Nb, Y, La and Ce (Table 6.1). These elements have the potential to discriminate between I-type granites formed at different temperatures. Some REEs (e.g. Ce and Y) behave compatibility in low-temperature I-type magmas (Chappell et al, 1998). Therefore, high content of the above REE in the Kashmar and Bornavard granitoids indicates I-type low-temperature. 143 CHAPTER 7 PETROGENESIS AND TECTONIC SETTING 7.1 PETROGENESIS 7.1.1 PRODUCTION OF I-TYPE GRANITE SOURCE ROCKS At an earlier stage of evolution, all of the Earth's crust was produced by partial melting of the mantle, with the S-type granites being derived after that more primitive crust was modified by weathering processes. I-type granites are produced either directly by fractional crystallisation of mantle-derived liquids or by direct partial melting of mantle- derived source rocks of the crust, therefore the source materials for 1-type granites will be broadly basaltic to andesitic in composition (Chappell and White, 1992; Pitcher, 1993). The high-temperature I-type granites are the most primitive and form by the partial melting of mafic rocks in the deep crust, or perhaps in modified mantle (Chappell et al, 1998). They are dominantly high-Ca tonalitic to low-K granodioritic rocks and occur in younger subduction-related continental margins. They characterise the Cordilleran I-type granites of Pitcher (1993). The 'low-temperature' I-type granites form by the partial melting of older quartzofeldspathic igneous rocks such as older tonalites, to produce magmas that comprise varying proportions of low-temperature felsic melts and restite. These rocks are mostly granodiorite and granite that typically occur in continental marginal arcs (e.g., eastern province of the LFB, Idaho Batholith, western USA and the Newer Granites of the Scottish Caledonides). They are generally more potassic than Cordilleran I-type granites since they represent a further stage of fractionation away 11<± from mantle-derived materials. Such low-temperature I-type granites are known as Caledonian I-type granite of Pitcher (1993), however they do occur in Cordilleran fold belts (Chappell and Stephens, 1988). In the present study, granodiorite and granite are the main rock types of the Kashmar and Bornavard granitoids. Chemical data show that the average contents of K20 in rocks of the Kashmar and Bornavard granitoids are higher than average K20 contents of Cordilleran I-type granites (e.g. PRB, Table 6.1). Isotopic data show that older igneous rocks of the crust that are virtually quartzofeldspathic in composition would likely be the source rocks for I-type granites of the Kashmar and Bornavard granitoids. These characteristics suggest that the Kashmar and Bornavard granitoids are similar to the Caledonian I-type granites. In contrast, the Kuh Mish intrusions are high in CaO (5.56 wt% on average) and low in K20 contents (0.92 wt% on average), features attributed to Cordilleran I-type granites. This similarity is consistent with low initial 87 86 Sr/ Sr and high £Nd values of the Kuh Mish intrusions. Primitive isotopic characteristics of the Kuh Mish intrusions indicate that mantle is the only possible source for generation of I-type granites in the Sabzevar Zone. 7.1.2 PRODUCTION OF I-TYPE GRANITES BY PARTIAL MELTING WITHIN THE CRUST For both I- and S-type granites, if melting of the source rocks occurs at minimum temperatures, then its composition will be slightly peraluminous, and is known as 'minimum-melt' composition (Chappell and White, 1992). In this case, the magma or 'minimum-melt' plus restite will initially have the same composition as that of the crustal source material. As the mantle is the ultimate source of all I-type granites, with 145 fractional crystallisation and/or restite separation, a range of compositions between 'minimum-melt' and mantle-derived material are produced (Chappell et al, 1998, 1999). Using an ACF diagram (Fig. 7.1), molecular percent of major oxides from plutonic rocks of the northeastern CJP, fall within the field of the metaluminous granites containing amphibole and biotite, implying source rock materials of mafic to felsic igneous composition (e.g. Chappell, 1984; Chappell and Stephens, 1988; White and Chappell, 1988). The join of plagioclase to the (FeO + MgO) apex defines ASI = 1, which divides metaluminous from peraluminous compositions. Most samples from the Kuh Mish intrusions fall within the central parts of the ACF diagram, where Chappell and White (1992) considered as 'mantle-derived source composition for I-type granites. This is consistent with low K20 and high CaO contents of the Kuh Mish intrusions and 87 86 supported by very low initial Sr/ Sr and very high £Nd values (Section 5.5.5). For the Kashmar and Bornavard granitoids, a series of rock compositions ranging from mafic to strongly felsic occurs and may be interpreted as having been generated by a combination of restite separation and fractional crystallisation processes (Sections 5.2.2, 5.3.1 and 6.4). All rock types from the Kashmar and Bornavard granitoids are low in Rb (mostly < 150 ppm) implying that the sources have been previously fractionated (Chappell and White, 1992). Also, the average Si02 contents from Kashmar and Bornavard granitoids are 66.80 and 69.84 wt%, respectively (Table 6.1). These values are higher than chemical composition of the continental crust (Si02 = 66 wt%) estimated by Taylor and McLennan (1985). The above features are consistent with isotopic data (Table 5.1) that suggest magmas of the Kashmar and Bornavard granitoids originated from partial melting of igneous rocks in the crust (e.g., Krogstad and Walker, 1997). 7.1.3 FRACTIONAL CRYSTALLISATION IN LOW-TEMPERATURE I-TYPE GRANITES All granite's are, in a sense, fractionated rocks, especially 'high-temperature' I-type granites because they formed originally from a magma that was completely or largely molten (Chappell et al, 1998). In this case, a series of more mafic compositions will be generated, but felsic compositions corresponding to 'minimum-melt' will only result when all restite has been removed and further crystal fractionation continues (e.g., Boggy Plain Zoned Pluton, Australia and Tuolumne Intrrusive Series, California). Low- temperature I-type granites produced by continuing fractional crystallisation of quartz and feldspars from felsic granite melt. Under such conditions the major element compositions of the melts change very little and hence major element abundances in the granites are all very similar, being determined by the equihbrium between quartz, feldspars and melt at low magmatic temperatures. However, trace element contents can range widely, and in some cases in a different way with high-temperature I-type granites (Chappell etal, 1998). In the present study, fractional crystallisation at low magmatic temperature is attributed to the Kashmar and Bornavard granitoids. For example, granite and granodiorite from the Kashmar granitoid do not show significant variation in major element composition 87 86 (Appendix 4.1). Also, the initial Sr/ Sr and ENd values of these plutons are very limited in range (Table 5.1). For each pluton, with increasing Si02 contents, little variation is observed in concentration of most of the trace elements such as Ba, Rb, Sr and Zr. In the granite, with increasing Si02 from 63.42 to 71.81 wt%, the content of Sr decreases from 315 to 188 ppm (Appendix 4.1). In the granodiorite, Si02 increases from 62.30 to 67.47 wt%, while Sr decreases from 342 to 282 ppm (Appendix 4.1). The granodiorite and granite contain variable amount of microgranular enclaves or restite (Section 6.4). It seems that magmas containing significant amount of restite can not readily undergo fractional crystallisation because relatively little liquid is available (Wyborn, 1983; Wyborn et al., 1987). This may be supported by low content of Rb in tonalite (45-72 ppm) and granodiorite (56-88 ppm) from the Kashmar granitoid (e.g., Azevedo and Nolan, 1998). The role of fractional crystallisation is well observed in producing variation in chemical composition of granodiorite from the Bornavard granitoid. In the granodiorite, with increasing Si02 contents from 63.18 to 71.32 wt% (Appendix 4.2), wide range is observed in concentrations of Ba (535-55 ppm), Rb (129-8 ppm) and Zr (272-142 ppm). On Harker plots (Fig. 5.7), Ba and Zr show inflexion at Si02 = 68.85 wt% that supports the process of fractional crystallisation (Section 5.3.2.1). For most analyses from the granodiorite, the elements Ce and Y decrease with increasing Si02 contents (Appendix 4.2). This behaviour is consistent with fractional crystallisation of low- temperature I-type granites (Chappell et al, 1998). 7.1.4 RESTITE FRACTIONATION Linear variations in chemical composition of granitoids (e.g., LFB) have been considered to result from magma mixing, restite unmixing and crystal fractionation (Collins, 1996, 1998; Chappell, 1978, 1994, 1996a,b; Chappell et al, 1999). In the present study, linear variation is well developed for the Kashmar granitoid, but direct field or petrographic evidence for magma mixing, even on small scale, is absent. In addition, major problem with the magma mixing is very httle occurrence of mafic rocks 148 87 86 (e.g., tonalite) and limited range in initial Sr/ Sr (0.70471-0.70569) and ENd values (- 1.86 to -0.70) of the Kashmar granitoid. Similarity of chemical composition (e.g., Ba, Rb, Sr) of the microgranular enclaves with the host tonalite, granodiorite and granite (Section 6.4) indicates that the enclaves are not parts of different source (e.g., Chappell, 1996a). If the enclaves were produced by fractional crystallisation of the host granite magmas (e.g., cumulates), then enclaves should be relatively higher in Cr, Ni and Sr and lower in Ba and Rb abundances compared with the host rocks (e.g., Chen et al, 1990). This behaviour is not observed in the present study. Magma mixing requires very large scale mixing of very fluid, high-temperature mafic melt with much less fluid and low- temperature felsic melt (Chappell, 1996a; Passmore and Sivell, 1998). Assuming that pure combination of two components (mantle and crust) produced a homogenised magma, then microgranular enclaves would be absent or uniformly distributed. But in the Kashmar granitoid, the microgranular enclaves reduce in size towards higher silica contents and disappear at Si02 >74 wt% (e.g., alkali feldspar granite). In the case of 87 86 well-mixed magmas, initial Sr/ Sr and sNd values most often fall within the mantle values (Castro et al, 1991). For example, Pankhurst et al. (1988) reported isotopic evidence relating to the petrogenesis of the Andean granitoids. Their values are nearly homogeneous within a given pluton and are consistent with a mixing model involving 87 86 variable proportions of mantle-derived magmas ( Sr/ Sr <0.704 and £Nd >0). In the 87 86 Kashmar granitoid each pluton is homogeneous in the initial Sr/ Sr and eNd values but all values indicate crustal origin. Therefore, linear variations in chemical composition of the Kashmar granitoid can not be explained by magma mixing. Since mineralogical and chemical data of the Kashmar granitoid show characteristics of low temperature I-type granites (e.g., Section 6.9), fractional crystallisation is not the 149 only possible process for variation in chemical compositions (Section 7.1.3). Therefore, the observed linear variations are interpreted to reflect restite separation. In this, case the primary melt composition can not be determined precisely because it has been modified by removal of restite and fractional crystallisation (Chappell et al, 1987). The composition of the source rocks for the Kashmar granitoid can be estimated by plotting a refractory component such as total Fe as Fe203 versus Si02 contents (Fig 7.2). The compositions of rocks of the Kashmar granitoid show a well linear trend. The composition of one microgranular enclave (R15912) plots very close to the trend line. For other samples, when selecting granite for chemical analysis, inclusions of all types were discarded. As the most felsic rock containing restite, is granite (Si02 <74 wt%, Appendix 4.1), then the composition of restite free rock (alkali feldspar granite) would be between Si02 = 74 to 77 wt%, that is a composition nearly similar to 'minimum- temperature' melts. The most abundant microgranular enclaves (restite) occur in tonalite that is low in Rb, indicating little fractionation. The enclaves are mineralogically and compositionally similar to tonalite that is the only mafic rock of the Kashmar granitoid. Therefore, according to the restite model (Fig. 7.2), the composition of the source rocks (restite + melt) would lie between restite (Si02 = -60 wt%) and 'minimum-melt' (Si02 = 74-77 wt%). The limited range in initial 87Sr/86Sr and ENd values of the Kashmar granitoid is usual condition for the restite model, indicating a relatively homogenisation during partial melting of the source rocks. In contrast magma mixing occurs in an open system and produces wide range in isotopic values (Collins, 1998). The Si02 content of restite is similar to the average composition of andesitic rocks that are complete equivalence in chemical composition with respective plutonic rocks of the Peninsular Ranges Batholith (cf. Hess, 1989; Wilson, 1989). This is consistent with the range of initial 87Sr/86Sr values of the Kashmar granitoid that is completely within the range of 150 initial 8/Sr/80Sr values of andesites (0.7046-0.7063) from the continental margins (Hall, 1987). 7.1.5 'MMMUM-MELT' COMPOSITIONS Analyses of alkali feldspar granite and granite, respectively from the Kashmar and Bornavard granitoids show the limited range in major element compositions (Appendices 4.1 and 4.2). These rocks are very siliceous (Si02 = 74-77 wt%). They contain low A1203 (<13 wt%), and they have very low contents of MgO (<0.40 wt%), CaO (<1 wt%) and the transition elements (except for Mn). Na20 and K20 have high and relatively constant abundances. These rocks are dominated by four metals, Si, Al, Na and K, which do not vary greatly in amount, and the variation in abundances of the normative minerals Q, Ab and Or, that incorporate those four elements are likewise restricted (Appendices 2.1 and 2.2). The alkali feldspar granite is typical example of extreme evolution of felsic I-type granite magmas. As mentioned earlier (Section 7.1.4), it may result from a combination of restite separation and fractional crystallisation processes. The granite from the Bornavard granitoid is free of microgranular enclaves, low in P205 (<0.04wt%), Rb (54-128 ppm) and Sr (38-54 ppm) but high in Ba contents (580- 890 ppm). Low content of P205 indicates low solubility of P that is characteristic of low- temperature I-type granites, consistent with abundant microcline and microperthitic K- feldspar in the granite. The I-type characteristic is confirmed by low A1203 content and presence of valuable modal abundance of biotite that is the only ferromagnesian mineral in the granite (Appendix 2.2). Major and trace element data from the granite do not show significant linear variation (Figs 5.5 and 5.7). This is consistent with homogeneous nature of the granite pluton and indicates that fractional crystallisation possibly was not operated 151 or it occurred below the present level of exposure. However, the granite is strongly enriched in LREE and initial 87Sr/86Sr (0.73978-0.75008) but low in concentration of transition metals that indicate earlier fractional crystallisation of the source rocks. Based on the chemical composition, the most favoured and significant process for generation of the granite, is selective fusion of only the quartzofeldspathic components of the crust and leaving the mafic components as a solid residue, which may have disengaged from the melt at or near the source. Evidence supporting the likelihood of such origin for the granite include textural relations (e.g.,; corroded or partly melted plagioclase and quartz crystals); absence of high-temperature mafic rocks; elevated initial 87Sr/86Sr and negative ENd values (e.g., Raymond, 1995). But on the ternary plot of normative Q-Ab-Or (Fig 6.4), the granite samples do not fall exactly in the position of the 'minimum-melt' of Tuttle and Bowen (1958). 7.1.6 Sr AND Nd ISOTOPES The regional variation in initial 87Sr/86Sr of granitoid rocks from the northeastern CJP is wide and range from 0.70471 to 0.70569 for the Kashmar granitoid, 0.70757 to 0.75008 for the Bornavard granitoid and 0.70386 to 0.70475 for the Kuh Mish intrusions (Table 5.1). The Kashmar and Bornavard granitoids are characterised by low sNd values (Fig. 7.3) ranging from -0.70 to -5.20 that indicate crustal sources. In the Bornavard 87 86 granitoid, granodiorite shows wider range in initial Sr/ Sr and ENd values compared with the granite. The wider range in Sr-Nd isotopic compositions may be related to heterogeneity of the source or fractional crystallisation of accessory minerals such as apatite, titanite, and zircon. These minerals are the main carriers for the REE and may have prevented equilibration for the Nd isotopes (e.g., Holden et al, 1987; Dias and Leterrier, 1994). In the Bornavard granitoid, granite is strongly higher in initial Sr/ Sr 152 than the granodiorite. This is consistent with the quartzofeldspathic nature of the granite (Section 7.1.5). The Sr-Nd isotopic values indicate that the plutons of the Kashmar granitoid have incorporated an isotopically less evolved component that may represent either young isotopically lower continental crust or enriched mantle magmas with juvenile sources (e.g., Rapela et al, 1992: Sewell et al, 1992; Bryant et al, 1997). The sources incorporated little or no Nd from the country rocks. The lower continental crust is characterised by low 87Sr/86Sr values that are not greatly different with modern enriched mantle values (e.g., Rollinson, 1993). This means that modern granites derived from the lower crust and those derived from the mantle will have very similar initial 87Sr/86Sr values. However, most of the granitic magmas are unlikely to have been derived directly from the mantle, but the relative contributions of components from the crust and mantle produce the range of isotopic compositions for I-type granites (Williams, 1998; Wyborn et al, 1998). Sr-Nd isotopic data for the Kuh Mish intrusions suggest that gabbro and the granodiorite from Darin and Namin plutons are genetically related. Variation in Sr-Nd isotopic values of the Kuh Mish intrusions is low and all values typically show characteristics of mantle-derived magmas (Fig. 7.3). The strong positive sNd values, low concentrations of incompatible elements (Section 5.5.2) and Nb and Ti anomalies may suggest a volcanic arc environment for the Kuh Mish intrusions (e.g., Allen et al, 1997; 1998; Price et al, 1999). The Kuh Mish intrusions are the most primitive I-type rocks of the northeastern CJP (Table 5.1). It is possible that the CJP is relatively more mafic and thinner in the 87 86 Sabzevar Zone than in the Taknar Zone. However, low initial Sr/ Sr and high £Nd values for I-type granites from elsewhere have been attributed to thinner and younger 153 age of the continental crust by some authors (e.g., DePaolo, 1988; Soler and Rotach- Toulhoat, 1990; Grigoriev and Pshenichny, 1998). In general, the initial 87Sr/86Sr values of the Kashmar granitoid and Kuh Mish intrusions are low and similar to initial Sr/ Sr values of the Oligocene to Miocene metaluminous I-type granites from the central and western American Cordillera, respectively (Silver and Chappell, 1988; Hess, 1989; Soler and Rotach-Toulhoat, 1990). Similarity in isotopic composition with young granitoids is in accord with the Middle-Late Eocene age, obtained at the present study, for different plutons of the Kashmar granitoid (Section 3.3.1.2). Also, the initial 87Sr/86Sr values of the Kashmar granitoid (0.704- 0.705) are similar to the values reported from I-type granites of the New England Fold Belt, eastern Australia (Shaw and Flood, 1981, 1993; Hansel et al, 1985; Bryant et al, 1997). Such I-type granites with low initial 87Sr/86Sr values would have been developed according to the processes of subduction-related magmatism (Chappell, 1994). 7.1.7 LFSE ENRICHMENT AND HFSE DEPLETION On rock/primordial mantle-normalised spider diagrams (Figs 5.2, 5.6 and 5.13), representative compositions from different plutons of the northeastern CIP show compositional similarities with those igneous rocks that occur in active continental margins (e.g., Sewell and Campbell, 1997). In particular, the spider diagrams show enrichment in most LFSE (e.g., K, Rb, Ba, Th) and LREE but relative depletion in HFSE (e.g., Nb, Ti, U, Hi), compared with primordial mantle-normalised values. Enrichment in LFSE and depletion in HFSE are stronger for the Kashmar and Bornavard granitoids, compared with the Kuh Mish intrusions. This is characteristic of subduction-related magmas and involvement of the continental crust in magma genesis 154 (Rottura et al, 1998; Feldstein and Lange, 1999), consistent with negative sNd values of the Kashmar and Bornavard granitoids. All plutons of the Kashmar and Bornavard granitoids display a distinct negative anomaly in Ba with respect to the adjacent Rb and Th. This behaviour of the trace elements is common in granites that occur in continental regions (e.g., Pearce et al, 1984). The negative anomaly in Ba is not conspicuous for the Kuh Mish intrusions (Fig. 5.13) indicating possibly an island arc environment (e.g., Price et al, 1999). The enrichment of Rb and Th relative to Ba cannot result from hydrothermal overprint, since Ba and Rb show comparable mobility and Th is considered to be immobile in aqueous fluids (Rottura et al, 1998). Many authors believe that LFSE enrichment and HFSE depletion are intrinsic features of the mantle wedge and a consequence of the immobility of most HFSE in aqueous fluids derived from dehydration of the subducted oceanic crust (McCulloch, 1993; Pearce and Peate, 1995; Price et al, 1999). There are numerous theories for the origin of HFSE depletion in arc magmas, but no consensus of opinion. In magmas of the northeastern CJP, variation in LFSE and HFSE probably results from differences in degree of partial melting and amount of LFSE enrichment of the source rocks and extent of fractional crystallisation. For example, the Kuh Mish intrusions originated from mantle source and are low in most of the incompatible elements. They show small variation between HFSE and LFSE on spider diagrams (Fig. 5.13). Whereas, the Kashmar and Bornavard granitoids, originated from crustal sources. They are high in most of the incompatible elements. They show large variation between LFSE and HFSE (Figs. 5.2 and 5.6). In the Kashmar and Bornavard granitoids, variation between LFSE 155 and HFSE are similar and becomes larger in plutons that extended fractional crystallisation. Examples are alkali feldspar granite and granite from Kashmar and Bornavard, respectively. Both the Kashmar and Bornavard granitoids generated from low-temperature I-type magmas that may suggest similar heat flow beneath the two granitoids. They occur in the same crustal structure that is probably a thick continental crust where regional heat flow is significantly low (e.g., Price et al, 1999). According to Chappell et al. (2000), low-temperature I-type granites mostly formed through partial melting of pre-existing quartzofeldspathic igneous rocks. This is in agreement with similar distribution of the ASI values that converge towards unity in the Kashmar and Bornavard granitoids. Therefore, similarity in variation of the LFSE and HFSE between the Kashmar and Bornavard granitoids may be the result of similar degree of partial melting of the crust. 7.1.8 A MODEL FOR EVOLUTION OF MAGMAS IN NORTHEASTERN CIP Rb/Sr ages of biotite-whole rock pairs for I-type granites of the northeastern CIP (Section 3.3) are consistent with the timing of magmatism associated with Late Mesozoic to Tertiary subduction of the Neo-Tethys Oceanic crust beneath the CIP (e.g., Kazmin et al, 1986a). In the present study, Mesozoic magmatism is reflected in Late Jurassic and Early Cretaceous ages, obtained respectively for granodiorite (153-145 Ma) and granite (124-112 Ma) from the Bornavard granitoid. The Late Jurassic and Early Cretaceous ages indicate the oldest plutonic activity in the Taknar Zone, and are consistent with widespread Jurassic-Cretaceous magmatism in the Middle East (e.g., Laws and Wilson, 1997). 156 Tertiary magmatism related to subduction of the Neo-Tethys produced extrusion of calcalkaline basaltic to andesitic rocks and intrusion of subvolcanic metaluminous granitoids in Iran and adjacent countries (Wilson, 1989). In the northeastern CJP, Tertiary magmatism is reflected in Middle-Late Eocene ages (43-42 Ma) of the Kashmar granitoid (Section 3.3.1.2). The isotopic ages indicate a significant time interval (-70 Ma) between the emplacement of the Bornavard granitoid and the Kashmar granitoid. However, during this time interval magmatism occurred in other parts of Iran (Darvichzadeh, 1992). For example, in the S-SMZ, most of the contact-aureole granites intruded before Eocene times (Mohajjel, 1997). In the CJP, igneous rocks of Tertiary age are widespread and the peak of volcanic activity is related to Eocene times (Darvichzadeh, 1992; Moradian, 1997). The isotopic data (Section 3.3) show that in the Taknar Zone, Middle-Late Eocene rocks are less isotopically evolved than Late Jurassic Early Cretaceous rocks. Because the Kashmar granitoid (Middle-Late Eocene) is less isotopically evolved than the Bornavard granitoid (Late Jurassic-Early Cretaceous), it is not concluded that Middle-Late Eocene rocks have intruded outboard of a significant thickness of continental crust. The Kashmar and Bornavard granitoids occur in the same geological zone and within -25 km at the present level of exposure. Explanation for a trend toward more isotopically primitive magmatism in younger rocks from elsewhere has been attributed to various models (Allen et al, 1997, 1998). Two of these models include 'plumbing' and steepness of the subduction zone and 'basification' of the continental lower crust (Pankhurst et al, 1988). In bran, a steep angle of subduction was first proposed by Berberian (1981) who believed that during the Middle Tertiary, the subduction zone steepened due to the collision of the Arabian plate and CJP. She also argued that 157 differences in isotopic compositions of Tertiary magmas derived from subduction of the Neo-Tethys beneath the CIP reflect a low-dip angle of subduction near the trench and a high-dip angle beyond the arc magmatic chain. According to Pankhurst et al. (1988), increase in subduction angle results a hotter melting, while low-temperature characteristic of the Kashmar and Bornavard granitoids indicate similar heat flow beneath the continental crust (Section 7.1.7). Therefore, the steepened aiigle of subduction can not explain lower isotopic values of the Kashmar granitoid, compared with the Bornavard granitoid. The second model relies on 'basification' which means that continental lower crust becomes more mafic through time by repeated injection of basalt (Asmerom et al, 1991; Farmer et al, 1991). This model has been used for at least two locations (Nevada and Arizona) in the Cordilleran Interior of the western United States. It is in agreement with isotopic data and some field relations of the granitoids occurring in northeastern CIP. For example, the Bornavard granitoid is Late Jurassic to Early Cretaceous in age and very high in initial 87Sr/86Sr. The Kashmar granitoid is Middle-Late Eocene in age, low in initial 87Sr/86Sr and associated with andesite. The Kuh Mish intrusions are associated with basaltic to andesitic rocks of Eocene age and have primitive mantle-like 87 86 initial Sr/ Sr and eNd values (Table 5.1). According to the above explanations, 'basification' of the continental lower crust with time is a more likely model to explain the magmatic evolution at the northeastern margin of the CIP. This model is supported by the occurrence of large volume of mainly intermediate magmas extruded in the U- DVB and the CIP through the entire Eocene and Oligocene (Haghipour and Aghanabati, 1989). The Eocene volcanic and plutonic rocks have been intruded by several gabbroic to dioritic bodies of Oligocene in age (Darvichzadeh, 1992). 158 7.2 TECTONIC SETTING Iran is located in the middle of the extensive Alpine-Himalayan orogenic system, produced by collision between the Eurasian plate to the north, and Afro-Arabian plate to the southwest (Bina et al, 1986). Along the entire active Eurasian continental margin of the Neo-Tethys, a large number of occurrences of calcalkaline rocks in Romania (Mason et al, 1996), the southern Alps (Rottura et al, 1998), the northern Arabian Plate (Laws and Wilson, 1997), Turkey (Wilson et al, 1997) and Pakistan (Petterson et al, 1993) have been investigated and the current study fills the missing link to the database of Tethyan I-type granitoids in northeastern CIP. According to many Iranian geologists the U-DVB is parallel to the Zagros Fold-thrust Belt, ophiolite-melange belt and linear metamorphic belt of the S-SMZ that extends from the Turkey to southeast of Iran (e.g., Moradian, 1997; Ghazi and Hassanipak, 1999). The U-DVB could have resulted from subduction of Neo-Tethys oceanic crust beneath the CJP during Mesozoic and Cainozoic (Alavi and Mahdavi, 1994; Berberian, 1995). In Iran, calcalkaline magmatism related to this subduction has occurred during five stages comprising Early Jurassic, Middle Jurassic-Early Cretaceous, Late Cretaceous, Paleogene and Late Miocene- Quaternary (Kazmin et al, 1986b). As closure of the Neo-Tethys proceeded, island arcs gave way to marginal continental volcanic belts with widespread development of high-K calcalkaline to shoshonitic series (Kazmin et al, 1986b; Moradian, 1997). hi Iran the intensity of calcalkaline magmatism is more pronounced in Jurassic-Cretaceous and Eocene times. In northeastern CJP, Late Jurassic Early Cretaceous and Middle-Eocene magmatism is recognised by occurrence of metaluminous I-type granitoids of the Kashmar and Bornavard, respectively. 159 7.2.1 ANOROGENIC GRANITES A chemical-tectonic approach to granite classification has been taken by some researchers, where chemical parameters are taken as indicative of the tectonic environment in which the granite formed (e.g., Pearce et al, 1984, Forster et al, 1997). Others have inferred magma source compositions from chemical and mineralogical features, for example, the I- and S-type classification of Chappell and White (1992). Anorogenic granites may occur in different tectonic settings but they are mainly related to rift, continental epirogenic and uplift stages. Among the anorogenic granites, 'A-type' is the most recognised granite that is referred to alkaline rocks (high K20 + Na20). Some A-type granites are metaluminous to weakly peraluminous that may be produced by high-temperature partial melting of a felsic infracrustal source (e.g., King et al, 1997). In the northeastern CJP anorogenic environment reported only for granite of the Bornavard granitoid by Esmaeili et al. (1998a). For several reasons (Sections 6.10.2), petrographic and geochemical data of the granite from the Bornavard granitoid are not similar with characteristics of the A-type granites. Esmaeili et al (1998a) believed that this granite emplaced in a rift-related tectonic setting. But rift-related granites may have a predominantly mantle source (Pitcher, 1993) or may be evolved by magma mixing to generate M-type (hybrid) granitoids (Castro et al, 1991). The rift-related magmas are 87 86 typically low in initial Sr/ Sr (<0.704) and have positive £Nd values (e.g. Wilson, 1989). In contrast, the granite samples of the Bornavard granitoid are significantly high 87 86 in initial Sr/ Sr and show negative sNd values (Table 5.1). Assuming that the granite of the Bornavard granitoid generated by lower crustal derived A-type magmas, it would be drier and hotter than I-type granites (e.g., Landenberger and Collins, 1996). But the 160 granite samples of the Bornavard granitoid plot close to the 'minimum-temperature' melt composition of Tuttle and Bowen (1958) and indicate a water-saturated melt at PH20 = -0.5 kb (Fig. 6.4). Also, the granite samples are low in the content of P205 that supports the low temperature characteristic of the felsic melt (Harrison and Watson, 1984; Chappell et al., 1998). Therefore hot and dry conditions for the granite of the Bornavard granitoid are not evidenced, hence A-type feature and rift-related environment are unlikely. 7.2.2 OROGENIC GRANITES Orogeny is characterised by plutonism, deformation and metamorphism (Lutgens and Tarbuck, 1992). In the CJP and the U-DVB of Iran, orogeny is characterised by extensive volcanic-plutonic associations and deformation during Late Mesozoic and Tertiary times (Darvichzadeh, 1992). The volcanic-plutonic associations are calcalkaline and alkaline in composition and occur through the entire southern parts of the Alp- Himalaya Belt (Kazmin et al, 1986a; Bina et al, 1986). In the northeastern CIP, the volcanic-plutonic rocks are calcalkaline in composition. The plutonic rocks of the northeastern CJP are I-type metaluminous granitoids. 7.2.2.1 Island Arc Granites The Kuh Mish intrusions show characteristics of the most primitive magmas in the northeastern CJP. They occur in the south of the extensive Late Cretaceous ophiolites of the Sabzevar Zone. The only published chemical data for the ophiolites of the Sabzevar Zone indicate calcalkaline character with trench-type, ridge-type and coloured melange affinities (Lensch and Davoudzadeh, 1982). The ophiolites of the Sabzevar Zone are 87 86 low in Ti02 and K20 contents. The initial Sr/ Sr values of the trench-type ophiolites 161 of the Sabzevar Zone are very close to values found in gabbro and granodiorite samples of the Kuh Mish intrusions. This may indicate that the Kuh Mish intrusions formed near an arc-trench system or even as a part of this system. Because rocks of the island arcs are usually high-temperature, at least plagioclase is cumulative, and hence their REE patterns mostly show pronounced positive Eu anomalies (Tate et al, 1999), while the REE patterns of the Kuh Mish intrusions show slightly negative Eu anomalies (Fig. 5.14). The Kuh Mish intrusions represent an average of 64.68 wt% Si02 (Table 6.1) that is significantly higher than that of the M- type rocks of the primitive island arcs (e.g., Bismark Volcanic Arc, mean Si02 = 60.13 wt%) (Hill et al, 1992). The average of 64.68 wt% Si02 from the Kuh Mish intrusions is very similar to the average of 64.63 wt% Si02 from 323 analyses of tonalitic I-type granites of the Cordillera (PRB, Table 6.1) that are excellent representatives of the active continental margins. 7.2.2.2 Magmatic Arcs of Continental Margins There is some evidence that granitoid rocks of the northeastern CJP originated in continental arc environment. For example the Kashmar and Bornavard granitoids together with the extensive Tertiary volcanic rocks occur in a belt parallel to the margins of the CJP. hi this belt, igneous rocks have typical arc shape and continuously extend towards the eastern boundary of Iran with a total length of about 300 km (Haghipour and Aghanabati, 1989). In northeastern CJP, granodiorite and granite are the most abundant plutonic rocks and. show characteristics of Cordilleran I-type granites. The associated volcanic rocks are mostly andesite, dacite and rhyolite, suggestive of the derivation of the more felsic magmas by fractional crystallisation of olivine, pyroxene, 162 plagioclase, Fe-Ti oxides and amphibole mineral assemblages from basaltic parental magmas (e.g., Wilson, 1989). The Si02 content is high and mostly ranges from 62 to 76 wt% with an average of 67.80 wt% for all analyses (Table 6.1). High content of Si02 is consistent with low-temperature characteristics of magmas generated in continental regions (Hess, 1989). The weakly metaluminous to weakly peraluminous features of the most plutons in the northeastern CJP support low-temperature characteristic, whereas granites from the island arcs are high-temperature and predominantly strongly metaluminous (Maniar and Piccoli, 1989). The Nb-Y diagram (Fig. 7.4) facilitates &scrimination of the tectonic environment of 'granitic magmas' (Pearce et al, 1984; Forster et al, 1997). The plutonic rocks of the northeastern CJP plot in the 'volcanic arc and syn-collisional' (VAG + syn-COLG) granite field because of their relatively low content of Y and Nb. The Bornavard granitoid is slightly higher in Y content, and plot at the edge of the within-plate (WP) field. A very similar distribution is observed in Figure 7.5, the Rb vs (Y + Nb), which confirms the arc affinities of the granitoid rocks of the northeastern CJP, although this diagram is less reliably interpreted because of the possibility of Rb mobility. The position of the Kuh Mish intrusions in the lowest part of the VAG + syn-COLG field is typical of immature arcs of the continental margin (e.g., Forster et al, 1997) that have relatively low degrees of the crustal influence. This is consistent with positive ENd values of the Kuh Mish intrusions. The Kuh Mish intrusions are relatively higher in Cr and Ni compared with the Kashmar and Bornavard granitoids and have similar situation in Rb vs. (Rb + Nb) plot as for the Cordilleran I-type granites (e.g., Forster et al, 1997). The Kashmar and Bornavard granitoids have negative £Nd values (Table 5.1), supporting the I-type characteristic of granites from continental margins (Hess, 1989). The position of 163 the most samples from the Kashmar and Bornavard granitoids in the upper most part of the VAG field is typically representative of arc rocks from an ancient continental margin. The Kashmar and Bornavard granitoids are strongly enriched in highly incompatible trace elements than the Kuh Mish intrusions, suggesting that the continental crust has had an important role in their magma genesis. Some samples from the Kashmar ganitoid are higher in Rb (-200 ppm) and plot very close to the Syn-COLG boundary (Fig. 7.5) but there is not overlap with the Syn-COLG field that indicates no incorporation of pelitic rocks in the melting process, consisting with typical I-type characteristics of the Kashmar granitoid. Collectively, plutonic rocks of the northeastern CIP occupy a similar area on the Rb vs (Y + Nb) diagram as the Cretaceous granitoids of eastern Eurasian margin (e.g., Nakajima, 1996). 164 CHAPTER 8 CONCLUSIONS 8.1 CONCLUSIONS A majority of the Iranian granites are either surrounded by contact-aureoles and are therefore contact-aureole types or intrude into related volcanic rocks and are therefore subvolcanic types. The contact-aureole granites of Iran are mostly Jurassic to Cretaceous in age and typically occur in the S-SMZ. Both contact-aureole and subvolcanic types are common in the CJP and the U-DVB. The subvolcanic granites of the CJP and U-DVB are associated with basalt and, most commonly, andesite. The volcanic-plutonic associations of the CJP and the U-DVB represent extensive Tertiary magmatism in Iran. The current study presents the first Rb/Sr isotopic data on biotite-whole rock pairs for the Kashmar and Bornavard granitoids (Table 3.1). Rb/Sr isotopic data on biotite-whole rock pairs clearly distinguish at least two different plutonic episodes for the Bornavard granitoid. Granodiorite pluton, the major representative of the early intrusive episode, yielded ages of 152.8±1.3 and 145.6±1.3Ma, and indicates that the oldest plutonic activity in the Taknar Zone occurred during the Late Jurassic. Granite pluton, representative of the late intrusive episode, yielded ages of 123.8±1 and 111.8±1.1 Ma, that indicate Early Cretaceous plutonic activity. The Late Jurassic and Early Cretaceous ages of the Bornavard granitoid are related to the Middle to Late Cimmerian Orogeny, recognised in Iran by intrusion of several contact-aureole granites in the S-SMZ and CIP. The Rb/Sr ages of the Bornavard granitoid are similar to isotopic ages for several contact-aureole granites that occur in northeastern CJP; for example, K/Ar ages of 165 146±3 and 120±3Ma for biotites from the Mashhad Granite (Table 2.1). Also, Rb/Sr ages of 152.8±1.3 and 111.8±1.1 Ma of the Bornavard granitoid are in agreement with K/Ar ages for biotite from the Torbat-e-Jam and Airakan Granites (153+5 and 113±8Ma, respectively), all indicative of Late Jurassic/Early Cretaceous plutonic activity in the northeastern CJP. Furthermore, the Late Jurassic and Early Cretaceous ages of the contact-aureole granites of the northeastern CJP are consistent with the time of extensive plutonic activity in the Middle East (e.g., Laws and Wilson, 1997). Before the present study, the granites and granodiorites of the Bornavard granitoid were compared with Precambrian and Tertiary granites of Iran, respectively (Section 3.3.2.2). These comparisons were based solely on petrography and limited stratigraphic relationships. The Rb/Sr data from the present study indicate that the earlier assumptions are incorrect. The granite pluton from the Bornavard granitoid intrudes the Taknar Rhyolite. The Taknar Rhyolite has been affected by hydrothermal alteration and low-grade regional metamorphism that occurred between 250 and 190 Ma (Crawford, 1977; Muller and Walter, 1983). The hydrothermal alteration is confirmed by a low content of K20 in biotite and high values of ASI due to the presence of abundant sericite in rhyolite samples. However, low content of Zr and total REE in samples from the Taknar Rhyolite indicate strongly fractionation. According to Rb/Sr dating on biotite-whole rock pairs, the oldest age of the Bornavard granitoid is 152.8±1.3 Ma (Table 3.1), which is significantly younger than the age of regional metamorphism of the Taknar Rhyolite. Differences in age, mineralogy (Section 4.5.1) and chemical data (Section 5.4.1) between the Taknar Rhyolite and the Bornavard granitoid indicate that they are not genetically related. 166 Rb/Sr isotopic data on biotite-whole rock pairs from different plutons of the Kashmar granitoid (Table 3.1) suggest Middle-Late Eocene plutonic activity in the northeastern CIP. The isotopic ages of the Kashmar granitoid show a narrow range from 43.5±0.4 to 42.4±0.4Ma that indicates typical synchronous plutonic activity in the northeastern CIP. These ages are very similar to a K/Ar age (43.7±1.7 Ma) determined from biotite in the host volcanic rocks and reported by Bernhardt (1983). The Middle-Late Eocene age of the Kashmar granitoid is consistent with the peak of the calcalkaline magmatism that occurred in the CJP during the Eocene (e.g., Darvichzadeh, 1992). Microprobe analyses show that in the Kashmar and Bornavard granitoids, the plagioclase grains have compositions between Anso-u and Ari44-2, respectively. In the Kashmar granitoid most plagioclase grains are typically normally-zoned and the compositions of most cores and rims are An50.3o and An3o-i8, respectively. There is some similarity in anorthite content of plagioclase crystals from granodiorite (An^-is) and granite (An50-i4) plutons of the Kashmar granitoid, indicating that they are genetically related. In the Kuh Mish intrusions, plagioclase grains from gabbro show a composition of An90-99 and these are the most calcic plagioclase compositions observed in the present study. Analyses of ferromagnesian minerals show that magnesio-hornblende, biotite, ilmenite, titanomagnetite and magnetite are the only mafic minerals occurring in the granitoid rocks of the northeastern CIP. Magnesio-hornblendes are homogeneous, low in Ti02, lv A1203 (Al <1 a.f.u.) and high in FeO and MgO contents that indicate crystallisation occurred at a low temperature and high f02 (e.g., Hammarstrom and Zen, 1986; 167 Anderson and Smith, 1995). The magnesio-hornblendes have been variably converted to biotite which is indicative of normal magmatic reactions. When magnesio-hornblende co-exists with biotite, the biotite is lower in Mg/(Mg + Fe), because it contains a higher total Fe as FeO content than magnesio-hornblende, and suggests an increase in/02 after amphibole crystallisation. When magnesio-hornblende is absent, biotite is significantly lower in Mg/(Mg + Fe) and shows the typical pleochroic scheme for biotite in oxidised I-type granites (e.g., biotite in granite from the Bornavard granitoid). Fractionation factors for MgO, FeO and Ti02 for magnesio-hornblende suggest that crystallisation occurred at moderate to low temperatures, but under variable crustal pressures up to a maximum of 3 kbar (Figs 4.3 and 4.14). According to the chemical classification of micas (Gribble, 1988), only three analyses of mica in alkali feldspar granite from the Kashmar granitoid are phlogopite in composition, whereas other analyses of mica from plutonic rocks of the northeastern CJP have Mg/(Mg + Fe) values ranging from 0.63 to 0.13 indicating biotite composition. Analysed biotites from granitoid rocks of northeastern CJP are homogeneous and titaniferous (Ti02 = 1.20-4.90 wt%). They are high in K20 content which indicates that the biotites are fresh. The biotite grains from the Kashmar granitoid and Kuh Mish intrusions are low in A1VI (all <0.5 a.f.u.), a feature typical of biotites from I-type granites. In the Bornavard granitoid, however, biotite is slightly higher in A1VI (mostly >0.5 a.f.u.), possibly because it co-exists with small amount of secondary muscovite that can be the result of subsolidus alteration (e.g., Harrison, 1990). Microprobe analyses of Fe-Ti oxides in granitoid rocks of northeastern CJP suggest that magnetite and titanomagnetite are the most common Fe-Ti oxides that accompany 1-68 hornblende and biotite. In heterogeneous grains, titanomagnetite is usually encountered in magnetite cores, emphasising increasingjfG2 after hornblende crystallisation. Ilmenite without magnetite occurs only in tonalite and some samples of granodiorite from the Bornavard granitoid. The ilmenite grains in the tonalite and granodiorite have rounded edges and show alteration to fine-grained titanite, features attributed to evolution of I- type magmas towards higher/02 (Petrik and Broska, 1994) and consistent with lack of any S-type characteristics in the tonalite and granodiorite. Chemical data show that the granitoid rocks of northeastern CJP are generally high in Na20, total Fe as Fe203, Ba, Sr, Mn and V, and low in P205, Ti02, Rb, Pb, Sn, Cu, Ni and Cr contents. High contents of Na20, Ba and Sr indicate that the source rocks were not previously weathered (Chappell and White, 1992; Chappell et al, 2000). This characteristic is consistent with mineralogical and chemical data that suggest an I-type source for the granitoid rocks of northeastern CJP. However, the low contents of Rb and most transition metals indicate that the igneous sources were previously fractionated or possibly low in these elements. The high content of total Fe as Fe203 is an intrinsic feature of the I-type magmas (Chappell et al, 2000) and indicates that magmas of the northeastern CIP evolved at high degrees of an oxidation condition. This is consistent with higher modal abundances of magnetite compared with other Fe-Ti oxides in most rock types of the present study. The association of magnetite and titanite supports high degrees of f02 in I-type granites. The ASI values are mostly less than one (Fig. 6.5), with averages of 0.97 for Kashmar, 0.95 for Bornavard and 0.92 for Kuh Mish, emphasising the metaluminous nature of I-type granitoids of the northeastern CJP. 169 • The Kashmar and Bornavard granitoids form a major plutonic suite of the Taknar Zone. They are similar in many petrological and some isotopic characteristics. Granite and granodiorite are the most abundant rock types of the Kashmar and Bornavard granitoids. In both granitoids, particularly for felsic rocks (Si02 >63 wt%), concentrations of most major and trace elements show regular trends towards higher Si02 contents (Figs 6.1 and 6.2). They show a similar distribution on histograms of ASI frequency (Fig. 6.5). Microprobe data confirm that magnesio-hornblende, biotite and Fe-Ti oxides are the only ferromagnesian minerals of the Kashmar and Bornavard granitoids. Plagioclase, K- feldspar, hornblende and biotite show characteristics of minerals crystallising at low temperature and high/02 under normal magmatic conditions. These characteristics are invoked to assign the Kashmar and Bornavard granitoids into a 'simple suite' that corresponds to low temperature I-type granites of Chappell et al. (1998). The Kuh Mish intrusions are isotopically and chemically different from the Kashmar and Bornavard granitoids. They are lower in Ba, Rb, Sr, Pb, Th, U and REE abundances, but higher in Cr, Ni and V contents. These characteristics are consistent with low initial 87Sr/86Sr and high £Nd values of the Kuh Mish intrusions. The absence of microgranular enclaves and zircon grains in samples from the Kuh Mish intrusions suggest that magmas of the Kuh Mish intrusions were higher in temperature than the Kashmar and Bornavard magmas. The Kuh Mish intrusions, therefore, define a different magmatic suite in the northeastern CIP. In the Bornavard granitoid, the role of fractional crystallisation is indicated in compositional variation of granodiorite samples (Sections 5.3.2.2 and 7.1.3). On Harker plots (Figs 5.5 and 5.7), with increasing Si02 contents in the granodiorite, a negative 170 correlation occurs for Ti02, total Fe as Fe203, MnO, MgO, CaO, Ba, Rb, Zr, Ce and Y contents. In particular, Ba, Zr and Zn show an inflexion at Si02 = -69 wt%, indicative of fractional crystallisation. Furthermore, a negative correlation of Si02 with REE, such as Ce and Y, indicate typical fractional crystallisation of low temperature I-type granites (e.g., Chappell et al, 1998). In the Bornavard granitoid, the youngest pluton is granite of Early Cretaceous age. The granite samples are quartzofeldspathic (Si02 = 74.84-76.04 wt%), similar in mineralogy and the contents of most major and trace elements. On the ternary plot of normative Q- Ab-Or (Fig. 6.4), the granite samples are clustered near the center of the triangle but do not fall precisely on the position of the 'minimum-temperature' melt of Tuttle and Bowen (1958). In the case of the 'minimum-melt' composition, due to the homogenisation of the partial melt, all samples of a pluton would be expected to be similar in initial 87Sr/86Sr values. The granite samples, however, have high and varied initial 87Sr/86Sr values (0.73622-0.75008), indicating that the granite is not a 'minimum- melt'. These features indicate that fractionation in a magma, derived from older continental crust, is more likely a process in generation of the granite from Bornavard granitoid. High contents of Na20 and a limited range in ENd values (-4.5 and -5.20) of the granite samples reduce the possibility of hydrothermal alteration for the generation of high initial 87Sr/86Sr values. In the Kashmar granitoid, with increasing Si02 contents in tonalite samples, significant compositional variation for most major and trace elements does not occur. The tonalite is rich in microgranular enclaves. It seems that the presence of a considerable amount of microgranular enclaves as restite (Section 6.4) resulted in the limited fractionation of 171 the tonalitic magma (e.g., Wyborn et al, 1987). In granodiorite and granite with increasing silica, the microgranular enclaves reduce in size and abundance. The enclaves are absent from alkali feldspar granite. On Harker plots, the granodiorite and granite show typical linear trends for most major and trace element contents. As mineralogical and chemical features of plutons of the Kashmar granitoid show characteristics of low temperature I-type magmas, fractional crystallisation is not the only possible process for generation of the linear trends. According to Figure 7.2, compositional variation of the Kashmar granitoid may be explained by restite fractionation (Section 7.1.4). This model explains the limited range of the initial Sr/ Sr and ENd values of the Kashmar granitoids. 87 86 The Kashmar granitoid has low initial Sr/ Sr (0.70471-0.70569) and sNd (-0.70 to - 1.86) values (Table 5.1 and Fig. 7.3). These values are consistent with Sr-Nd isotopic features of I-type granites originating from infra-crustal source rocks. The age (43.5- 42.4 Ma), initial 87Sr/86Sr and ENd values of the Kashmar granitoid show a very limited range, indicating typical synchronous plutonism and the generation of magmas from compositionally similar source rocks. Variation in chemical composition of these magmas resulted from fractional crystallisation and restite separation. The Bornavard granitoid exhibits a broad spectrum of isotopic characteristics with high 87 86 initial Sr/ Sr values ranging from 0.70757 to 0.75008 and low negative £Nd values ranging from -1.41 to -5.20 (Table 5.1). These data indicate that magmas of the Bornavard granitoid were generated from isotopically-evolved continental crust. Significant enrichment in initial 87Sr/86Sr values may indicate that magmas of the R7 Bornavard granitoid were extensively contaminated with radiogenic Sr ( Sr) derived 172 from older felsic rocks of the continental crust or that the magmas were produced by partial melting of old felsic rocks of the continental crust. The Kuh Mish intrusions are very low in initial (at 42.8 Ma) 87Sr/86Sr values (0.70386- 0.70475) and very high in £Nd values (+8.02 to +6.30). The primitive isotopic signatures of the Kuh Mish intrusions indicate that mantle is the only source for their generation. Similarity in Sr-Nd isotopic data between the Kuh Mish intrusions indicates that they are genetically related. However, gabbro is significantly higher in Cr, Ni and Sc contents, and lower in most incompatible elements than granodiorite samples, suggesting that chemical differences may be the result of fractional crystallisation of mantle-derived magma. On rock/primordial mantle-normalised spider diagrams (Figs 5.2, 5.6 and 5.13), I-type granitoids of northeastern CJP show enrichment in most LFSE and LREE but relative depletion in HFSE, compared with primordial mantle-normalised values. This is characteristic of subduction-related magmas, consistent with subduction of the Neo- Tethys Oceanic crust beneath the CJP during the Late Mesozoic and Tertiary. Enrichment in HFSE and depletion in LFSE (particularly Nb, Ti, U and Hf) contents for the Kashmar and Bornavard granitoids, compared with the Kuh Mish intrusions, are consistent with crustal sources of the Kashmar and Bornavard granitoids. Distinct negative anomalies in Ba with respect to the adjacent Rb and Th indicate that the Kashmar and Bornavard granitoids were emplaced in continental margins (e.g., Sewell and Campbell, 1997). Conversely, primitive isotopic features and small variations between LFSE and HFSE in the Kuh Mish intrusions indicate the possibly an island arc environment (e.g., Price et al, 1999). 173 Sr-Nd isotopic data show that granitoids of the Taknar Zone originated from crustal igneous rocks, but Middle-Late Eocene granitoids are less isotopically-evolved than Late Jurassic and Early Cretaceous granitoids. In the Sabzevar Zone, Middle-Late Eocene rocks are the most isotopically primitive rocks of the northeastern CJP. Explanation for a trend towards more isotopically primitive magmatism in younger magmas is in agreement with a 'basification' model (Section 7.1.8). This model suggests that the continental crust in northeastern CIP became more mafic with time by repeated injection of basaltic and andesitic magmas derived from subduction of the Neo-Tethys Oceanic crust beneath the CIP. The granitoid rocks of northeastern CIP are low in the contents of Rb, Y and Nb. 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Annual Review of Earth and Planetary Sciences 16, 21-51. 200 Figure 1.1 Generalized tectonic map of Iran, based on the geological maps of Ruttner and Stocklin (1967) and Alavi (1991). 2U1 57" 30' 58" 00' 58" 30' 59" 00' Namin Neyshabur Darin o o^vVolcano-sedimentary rocks (Eocene) N Granitoid rocks (Mesozoic & Tertiary) 1 v v 1 Volcano-sedimentary rocks (Cretaceous) J Metavolcanic rocks (?Precambrian) I 1 Quaternary 4£ Area studied 500 km - 25' Figure 1.2 Location of the area studied (after Emami et al, 1993). 203 1 1 •=' S3 in r c rt> B tr 3 ro" P 3" ce O a o 3 a. rc 3 a a. 3 o00o tag 03 rt> trt ti* >-" S'*I f£ 3 P <£ o rp a o "-J CD a N)I iooCD •Hcr cr o CD N wo o a p H n BP o o tr p i a P a. so CJ P I. tu trJ O o P r-t- CD tr VO IVO I -Jcf ex. CD cr £: I' EL o o P o o r-i- o •—» cCrD CD oo P CJ a. ,p r O O .4 204 ' Figure 2.3 Representative regional-aureole granites of Iran. TURKMENIYA ARABIAN PLATFORM Figure 2.4 Representative contact-aureole granites of Iran. 206 Figure 2.5 Representative subvolcanic granites of Iran. 207 57" 45' 57" 50' V V V V v" V V V V 35*25' *T 35° 20' Qt Quaternary deposits (alluvium fans and terraces) Limestone \ (Late Cretaceous) V V V Andesite and basaltic rocks (Cretaceous) + + + Bornavard Granitoid f r r • Taknar Rhyolite ' r r r / / / Metavolcanic rocks .,/ / / WI'll Dolomite (Late Palaeozoic - Early Mesozoic) Sample Sites: 1 = R15948,2 = R15949, 3 = R15950,4 = R15951, 5 = R15952 Figure 3.1 Generalized geological map of east of the Taknar Zone, northeastern Central Iran Plate (after Valipour, 1992). 208 < o \ \ < s \ o < s \ < \ \ o o 1- n B E 8 ao n .* 8 » •i a 1 0 3 ^1 Sn" a s s 0 to n i a E o a + + o + + PP |tf oo >- CO o HHH || || S3 c e. vo vo jrj Jd a a K to WVO U> £ UlI- -I . " " vo ^D >-< WtO i-» © oCO —• 2. D. rH — n » 2. VO "g II a. 3. \0 W Ui J* h- ?a. Wi vo VO Ul VO VJMH VO • o *.o o 8 (0- - CO to ^ 00 - II w <* "^ u> & II II O II 8KK£S (0 M Ui £ M i, VO JX Iw ^ t- II 2 II II ^ w i*M • CLO W i Ui 3J »vo u.i o <- o so to . a 3" o- T ^ 00 W 8. 3 II 11 to II & 8 '-< ^ # hr- to —. vo W vj h- II to h- 11 n II 8 •3=Q s 2j p m s-1 ^2 8. (••3S.J >°° ro g J H o\ Ul in Age=42.8±0.2Ma Initial = 0.70548±0.00003 0.6 MSWD = 5.3 0.5 X 0 80 160 240 320 400 87Rb/86Sr Figure 3.3 Rb/Sr whole rock isochron diagram for granodiorite, granite and alkali feldspar granite from the Kashmar granitoid. MSWD = Mean Squares of Weighted Deviates, used as a measure of the goodness of fit of the isochron. 210 57" 45' 57" 50' 57" 55' 58° 00' ' '""'I 0 • fi v v v v v v v yn/ vvvvvvvvvvvvvvvvv 0 o o yj /jvvvvvv.vvvvvvv^ ^-n^^yf/j /- w « v V V V V V o SABZEVAR ZONE ^ -vvvvvvvvv/Jvv o ax £x \ y a/v v v v v v v /°\v v_ v v v v v^v V V V V V V V V V / / V ?//,i£iii: 35° 25' vvvvvvvvvv V V V V V V V V V V V vvvvvvvvvv V v V V V V v - 35° 20' o o o c> uncVi Fau« Bardaskau Anabad Babolhakain 'To Kashmar -»ji^. j£> •+ + >| Bornavard granitoid (Late Jurassic Early Cretaceous) I Qt I Quaternary deposits fj1 i\ Taknar Rhyolite o o | Volcano-sedimentary rocks (Eocene) 11 11\ Metavolcanic rocks I i ' j Limestone (Cretaceous) "1 Late Palaeozoic - Early Mesozoic sedimentary rocks V V | Andesile and basaltic rocks (Cretaceous) • Sample Sites: 1 = R15938, 2 = R15939, 3 = R15940,4 = R15941, 5 = R15942, 6 = R15943, 7 = R15944, 8 = R15945, 9 = R15946,10 = R15947,11 = R15953,12 = R15954,13 = R15955 Figure 3.4 Geological map of the Bornavard area (after Eftekhar-Nezhad, 1976). 211 QUARTZ + Tonalite x Granodiorite * Granite • Alkali feldspar granite K-FELDSPAR PLAGIOCLASE Figure 4.1 'Modal compositions (quartz, K-feldspar, plagioclase) of the Kashmar granitoid. Fields are based on classification of igneous rocks, proposed by the IUGS Subcommission on the systematics of igneous rocks (Streckeisen, 1976; Le Bas and Streckeisen, 1991). Albite 100 100 "A A A A A A A A A 100 90 80 70 60 50 40 30 20 in Anorthite Orthoclase Figure 4.2 Plagioclase composition from the Kashmar granitoid. 212 ~ 0.3^ 1 kbar 3 kbar 3 • «, 0.25 - o> 1 0.2- a E 0.15- (0 2 0.1- 55 1 0.05 - o 2 ^& ^» 'i 0VJ - i • 0 0.5 1 1.5 Total Al in amphibole (a.f.u.) Figure 4.3 Diagram of Ti versus total Al for magnesio hornblende from the Kashmar granitoid, with pressure contours determined according to Johnson and Rutherford (1989a). 2.7 -— -~: ~ ' " I 3 1* - j ; 2.3 - \ 3 Al" \ # ij 2.1 \ 1.9 - • 1 1.7 H —I — I 1 1 1 1 III 0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0,9 1 Annite Mg/(Mg+Fe) Phlogopite Figure 4.4 Composition of biotite crystals from the Kashmar granitoid. The boundary between phlogopite and annite is proposed at Mg/(Mg + Fe) = 0.7 (Gribble, 1988). 10 12 14 16 18 20 MgO (wt%) Figure 4.5 Plot of MgO versus A1203 from biotite of the Kashmar granitoid. Discriminant lines are after Abdel-Rahman (1994). A = Anorogenic, P = Peraluminous and C = Calcalkaline orogenic. f80 " I 60- • |* $ * • o o 50 - o JZ <* 4U "1 1 1 i 0.4 0.5 0.6 0.7 0.8 Mg/(Mg + Fe) Figure 4.6 Relationship between Mg/(Mg + Fe) and whole rock Si02 contents for biotite from the Kashmar granitoid. 214 4 -i 1 i "" 1 " i 0 ! 1 > t i 0.4 0.5 0.6 0.7 0.8 Mg/(Mg + Fe) Figure 4.7 Diagram showing negative correlation between Mg/(Mg + Fe) and total Fe (a.f.u.) from biotite of the Kashmar granitoid. Figure 4.8 Diagram showing negative correlation between Mg/(Mg + Fe) and Ti (a.f.u.) in biotite from the Kashmar granitoid. • Biotite Ll Hornblende 1 2 3 4 5 6 7 8 9 10 11 12 13 Number of analyses Figure 4.9 Comparison of Mg/(Mg + Fe) between coexisting hornblende and biotite from the Kashmar granitoid. Albite 100 \ A A A A A A A A A" 100 90 80 70 60 50 40 30 20 10 Anorthite Orthoclase Figure 4.10 Composition of K-feldspar crystals from the Kashmar granitoid. QUARTZ + Tonalite x Granodiorite * Granite K-FELDSPAR PLAGIOCLASE Figure 4.11 Modal compositions (quartz, K-feldspar., plagioclase) of the Bornavard granitoid. Fields are based on classification of igneous rocks, proposed by the IUGS Subcommission on the systematics of igneous rocks (Streckeisen, 1976; Le Bas and Streckeisen, 1991). Albite 100 100 A A A A A A A A 7\ 100 90 80 70 60 50 40 30 20 10 Anorthite Orthoclase Figure 4.12 Anorthite-Albite-Orthoclase triangle plot for plagioclase composition from the Bornavard granitoid. Albite 100 100 VX A A" A A A A A A" 100 90 80 70 60 50 40 30 20 10 Anorthite Orthoclase Figure 4.13 Anorthite-Albite-Orthoclase triangle diagram showing the composition of K- feldspar crystals in granite from the Bornavard granitoid. 0.25 - 1 kbar 3 kbar _ 0.2 - Jpl- 5 0.15 - r o.i- # # w 0.05 - •# 1 1 0 - ! I 0 0.5 1 1.5 Total Al (a.f.u.) Figure 4.14 Diagram of Ti (a.f.u.) versus total Al (a.f.u.) for magnesio hornblende from the Bornavard granitoid, with pressure contours determined according to Johnson and Rutherford (1989a). -a _ £ 2.5 - (0 1 c 1.3 > i t l i 0 0.2 0.4 0.6 0.8 1 Annite Mg/(Mg + Fe) Phlogopite Figure 4.15 Composition of biotite crystals from the Bornavard granitoid. Figure 4.16 Negative correlation between Mg/(Mg + Fe) and total Fe (a.f.u.) for biotite crystals from the Bornavard granitoid. 15 - 13 £ • *• o B + O 7 K < d I CO 2 I 5 t~- ill. ' - 1 - ' . I - I . I - ' • ' • ' • ' • ' ... I i I • I ' I i I'l , 1 73 77 '3V ' 4', ' '4A5 ' 'J49o S3 ' 57 ' 6» 6 5 6 9 7 3 I ' SiOz wl%. ULTRABASIC 45 BASIC 52 INTERMEDIATE "63 ACID Figure 4.17 Total alkali contents versus Si02 (wt%) classification (TAS) for the Taknar Rhyolite (fields after Le Maitre, 1989 and Le Bas and Streckeisen, 1991). 219 ' Qt Quaternary deposits |Q Oi Volcano-sedimentary rocks (Eocene); [-»• 4- ] Kuh Mish intrusions (Middle-Late Eocene) '' i ' I Limestone (Late Cretaceous) v v[ Andesite and basaltic rocks (Late Cretaceous-Early Tertiary) • Sample Sites: 1 = R15926, 2 = R15927, 3 = R15928, 4 = R15929, 5 = R15930, 6 = R15931, 7 = R15932, 8 = R15933, 9 = R15934, 10 = R15935, 11 = R15936, 12 = R15937, 13 = R15956 Figure 4.18 Simplified geological map of the Kuh Mish area (after Eftekhar-Nezhad, 1976 and Sahandi, 1989). * Gabbro x Quartz monzodiorite * Granodiorite K-FELP5FAR PLAGIOCLASE Figure 4.19 Modal compositions (quartz, K-feldspar, plagioclase) of the Kuh Mish intrusions. Fields are based on classification of igneous rocks, proposed by the IUGS Subcommission on the systematics of igneous rocks (Streckeisen, 1976; Le Bas and Streckeisen, 1991). Ca2Si20B (Wo) 100- 100 "A A A A A A—""A A A 7 100 90 80 70 60 50 40 30 20 10 Mg2S2D6 (En) Fe28iZ06 (Fs) Figure 4.20 Composition of clinopyroxene in gabbro from the Kuh Mish intrusions. Alblle O Gabbro • Granodiorite 100 "A A A A A A A A . A 100 90 80 70 60 50 • 40 30 20 10 Anorfhlle Orthoclase Figure 4.21 Composition of plagioclase crystals in gabbro and granodiorite from the Kuh Mish intrusions. 221 -i—i—i—r -i—i—r=jf—i—i—i—i—j—i—i—i—i—|—i—i—r -i—i—i—r-; •n 0-3 r 9, 0,2 I t * fo£## * 0.1 * * • O 4 * * + • 2 + ****m 5 4- -dr o X * 8 4 + ++ • * * .3 + fir%- O 6 + 4 + '^* U * * 2 r-ip| nryn O 4 + + + bO 2 + * #*** J I L ULJ L dni. ran i 2 ' ' i i ! i i i—i—I—i—;—i—L. iiit l i i i i M i i > • 0.12 t 0. 08 0.04 $ !^_ | "I I 1 1 1 1 1 I 1 1 M MlN' ^^ 8 ++ o X*# * E 0 3 V L 0 E 2 L CL \ u = 1 0 Rb Th K La Sr P Ti Na Figure 5.2 Multi-element patterns (spider diagrams) for the Kashmar granitoid. The normalised values are from Mc Donough et al (1991). Symbols as Figure 5.1. T—I—I—rr-]—i—i—i—i—|—i—i—i—r "J I I I I I I—I—1—1—]—I—I—I—1— t ++ X D 16 '.12 rfr ^c ]•, 30 JZ 20 * f • D • 10 + tW ?V "TP TJ^fffh, jfu •+ x 50 * * + + C 40 X iit N t 20 * ir * •a 800 X + X ^ ...... 7; 2 400 * # a + ' WUr 200 + + + > 1 00 * " i i i i I i i i i I i i i i I -i i i i i I * i i idniTO i i' * 1 2 * u * 00 8 4 * I I 1 I 4* 44- K 200 x* ^ + N + * B 100 * | I I M '| I -I I I | I I I I | I I I I | I I I t 30 + x * L? J x x >- .* < * •*.# * 20 X *"' ^x • * a a i i i, | ii ii ii | j i i i i | i i i i [ 11 i i 1 ii i i ; + *** ** 00 200 iD MIIMI ^ ® 200 D JQ * # D 100 xXy * a + + x *x 4-4-+- 111 MM X* x>«< 600 r * a O + X &.1V D m 400 + 4- 200 4- I • • • • l • • i • i i i t fi, ,i < • • i 55 60 65 . 70 75 Si02 (wt%) - Figure 5.3 Harker diagrams for trace elements abundances of the Kashmar granitoid (oxides, wt% and traces, ppm). Symbols as Figure 5.1. 1 00 CD 50 •D C O O 20 CJ 1 0 J* u o La Eu Tb Ho Lu Ce... .• Nd Sm Gd . Yb Figure 5.4 Rare earth element patterns for granodiorite, granite and alkali feldspar granite from the Kashmar granitoid. The normalised values are from Taylor and Mc Lennan (1985). Symbols as Figure 5.1. t;'l M I1 [ I i I l'| I I I I |l I I ! | I I II | I M I I I M O 2 * 8 + 9 4 x* x U v|Wfab 6 o + 4 if 2 44- 0.08 r x X '•¥ o + 4 r x X •¥• 0,8 h XX x £o 0,4 I'll I I—I—1_ I I I I I I I I I I I • I I I ' I I I I I I—I—I 50 55 60 65 70 75 Si02(wt%) Figure 5.5 Harker diagrams for the Bornavard granitoid (oxides, wt%). Symbols: Tonalite (+), Granodiorite (x) and Granite (*). 100 = 4-1 c 0 10 = 1 100 TJ 10 = L 1 CL 100 u 10 = 0 H. 1 E Rb Th K La Sr P Ti Na Figure 5.6 Multi-element patterns (spider diagrams) for the Bornavard granitoid. The normalised values are from Mc Donough et al (1991). Symbols as Figure 5.5. 1 i i i i | i i' r-i—i—I--I r 1 : ~ 1 1—1—1—1—1—r I'I 1 1 ffi* i—1—r _) 3 X * 2 x x 4- •4-4- 1 x 8 4- X 5 -£4-4- m .c 20 X f- 4- X 10 H^r- c 4- 120 X IS) XX X 80 'J 1 I4J -# 40 4- 1200 X* X 800 4#- 4-4-4 400 + yX _i 1 i_ 1 .1...1. 1 .1 ' ' ' 1 J I I l_J ' ' ' !'ii )W-4 • •' ••' 200 r 1 1 1 1 —I—T 1" I "I ITTI I I I I I I l"'l I 1 1 1 1 | I 1 1 1 u 100 LO 30 XX ^ 400 +x- 1*1 I iff 20 X N 200 4- X A- II!I •[••!• !• 60 -H- >- 40 f X 20 X -fr L. 160 £ + X 10 80 120 a: . 80 % 40 4=N- 4-£+ 4*H a 800 X CD 400 t 1 ,141 I. i 1 1 1 l.i 1 4 I'll I, ,1., (• I ,x (X [ I, I I.I I L 50 55 60 . 65 70 75 Si02 (wt%) Figure 5.7 Harker diagrams for trace elements abundances of the Bornavard granitoid (oxides, wt% and traces, ppm). Symbols as Figure 5.5. Q)- - 3 •o c o JZ : 2 O O o o - 1 Ce Nd Sm Gd Figure 5.8 Rare earth element patterns for tonalite, granodiorite and granite from the Bornavard granitoid. The normalised values are from Taylor and Mc Lennan (1985). Symbols as Figure 5.5. -r-i—i—1—|—i—i—i—i—[• i i i i [ i t i i | i i i i i.' ' ' » t ' ' r~r 0,75 « 0,74 X CO vo BO £ 0.73 73 # •iH X ••rH & 0 .72 X 0. 71 x + +n 4-4- 4^ •a dr .1 I I I I U_I—I—L I • , , • I • • • • I i i i i I i i r I 1 I I ' ••' 0. 70 0 55 60 65 70 75 Si02 (wt%) 87 86 Figure 5.9 Plot of initial Sr/ Sr versus Si02 (wt%) for igneous rocks of the northeastern CIP. Symbols: Kashmar (+), Bornavard (x), Taknar (*) and Kuh Mish (D). 100 r 30 - T5 O 10 r £ CL 3 - O 1 r o LY Rb Th K. La S'r P • Ti Na Figure 5.10 Multi-element patterns (spider diagrams) for the Taknar Rhyolite, northeastern CIP. The normalised values are from Mc Donough et al. (1991). i i i I i i i i i i i i i i 1 00 / v i i 50 i - C o JZ O CJ 20 - ^^_ ,i , i - u 10 o LY • i i i—i—i— 5 —i'— —i— Eu Tb Ho Lu La e N d S m G d Y b C Figure 5.11 Rare earth element pattern for the Taknar Rhyolite, northeastern CIP. The normalised values are from Taylor and Mc Lennan (1985). M I I | l)[ I I | I IJ I | I I I I |l I H | I I I I | I M I. 80 r X X X a 40 i * * j i 4^4-4 X X * 800 *3- * t-H- I I 11 T 300 X x > 200 * 100 ,** I I I I o 4 X *^## * X * 2 * X 16 s- 9 x X a x x :*. P- _1_ ) |* Vl*. I * r4- ++ i i r i r-r-r O 8 6fj X x X X 4 r * 4-x-ir \*\*\*ft\ 1 t 0.16 x x * 0 .08 7*- * ^>H-4- rm * o 8 x X 4 "M- **% ^ 1 6 X o * 14 X • .*r. i * t_l I I L • i i i i • i i i i i i i i i—L_J—i—i—i—i 3 50 55 -60 65 70 75 Si02(wt%) Figure 5.12 Harker diagrams for Gabbro (+), quartz monzodiorite (x) and granodiorite (*) from the Kuh Mish area, northeastern CJP (oxides, wt% and traces ppm). QJ C = 3 G TJ i_ 2 O £ Q. o 0 LY Rb Th K La Sr Figure 5.13 Multi-element patterns (spider diagrams) for gabbro (+), quartz monzodiorite (x) and granodiorite (*) from the Kuh Mish area, northeastern CIP. The normalised values are from Mc Donough et al. (1991). tu TJ C o JZ o CJ J* u o C£ Figure 5.14 Rare earth element patterns for gabbro (+) and granodiorite (*) from the Kuh Mish area, northeastern CIP. The normalised values are from Taylor and McLennan (1985). _p I I I I I I I I I I I I I I I I I I I I I I I i r * i ~ 8 r % 4 =• D J "- D s l"t.l ti^flWlEflrp^ *a o 10 n n*P x • u D 8 * X++D o + x 4 ^ffN--Hu#*•. .xflJJqQ ] o 1 6 x D • JCff""^''^ 3 14 X x^^cf 12 Mil ij?t 0,3 "+ +*& * • 0 r. 1X D "Q_i—i—i—I i_i i Ql i_ .,, i .... i ..... i .T^-Hl^^ j i_ 50 5•Fa 60. 65 70 75 Si02 (wt%) Figure 6.1 Plot of major elements against Si02 contents for plutonic rocks of the northeastern CIP. (oxides, wt%). Symbols: Kashmar granitoid (+), Bornavard granitoid (x) and Kuh Mish intrusions (D). 1 ' V I 'ill luy" i i i | i i i i ] i i i i ] i i i i | i i i r 800 C I 1 I 1 1 I R I I 1 Ml I 300 i- ' "dr D > 200 r x 7 x++ ° +* "4- 1 oo a fl^K 60 r HH-f x >- 40 I + + •n• + +,• +^=fo"^- -* 20 33- ^ 1 1 1 I 1 i,l I t 40 K+f +Xp%+#f x^^ +ff 20 a- 44^- i i [qi i9| i i i i ffifft | iii 20 + £ 4^ + ,+ , . HdfH-W=t: « 10 1 X tl l l*l CT P"ft[ [qqqq^g 80 tq no 03 ++ x +»c. o 40 i X +4^V^ x+X4- +* Tj i |X, | ,q Da J I I u in , ,p t , , •, , ipdlrm in • 50 55 60 65 70 75 Si02(wt%) Figure 6.2 Plot of trace elements abundances against Si02 contents for plutonic rocks of the northeastern CIP. (oxides wt%, traces, ppm, symbols as Figure 6.1). .1 | | I | | I i I | I. I I I \ 1 I I I | II I1 I ) I I M | .1 I I I | I I I I j 1 I I I | I I I I- a 0s ' 75 i P 70 X 65 2 60 (b) 55 50 a 6 .1 n 1111111 D i ji i M 1111 I 111 4+ 3.5 Et Peraluminous ¥ 3.0 a o 2.5 P X (a) a 2.0 # & r Metaluminous "•ft 1 ,5 O 1 ,0 L Peralkaline < 0 . 5 F. . i . I • . i . I I I I I I I I. I t I I I I I I I I,. . 11. i... • 111 • • i • • • •••i U 0.6 0.8 1.0 1 .2 1.4 ASI Figure 6.3 (a) Aluminum Saturation Index (ASI) and (b) Si02 contents versus ASI values for igneous rocks of the northeastern CJP. The boundary between metaluminous and peraluminous at ASI = 1.1 proposed by Chappell and White (1974) because they recognised more generally that very felsic I-type granites may be weakly peraluminous (Chappell, 1998b). Symbols: Kashmar granitoid (+), Bornavard granitoid (x), Kuh Mish intrusions (C) and Taknar Rhyolite (*). Q 0.5 kb 3.0 kb Ab ii u -S. Or Figure 6.4 Ternary plot of normative Q-Ab-Or for granite from the. Bornavard granitoid. The curves for water-saturated liquids in equilibrium with quartz and K-feldspar at 0.5 and 3.0 kb are from Tuttle and Bowen (1958). The position of •minimum-melt' compositions of Tuttle and Bowen (1958) are shown by a cross (+) on each curve. M I I | i i . i | i i i . | I I I I | I I I I | ) I I I | I M I I I I I I | I I { I [ | i | | 2 - _ 4 4- I I I I | I I I I Illllllll I I I I 1 I I I I 11 I I I I I 1 I I c 11 n [i 111 11 11 16 ± ll H IL 8 i i t i i i i i < i t i > i i f . i i i i i . / i i i t < i i i i i i < f i f i i i 11 11 » t i 0.6 0.8 1.0 1.2 1.4 Aluminum Saturation Index (ASI) Figure 6.5 Histograms of ASI frequency for igneous rocks of the northeastern CIP. Numbers in the right side of the histograms indicate 1 = Kashmar granitoid, 2 = Bornavard granitoid, 3 = Taknar Rhyolite and 4 = Kuh Mish intrusions. Igneous ACF (Molecular%) • Al203-K26-Na20 Plagioclase /Peraluminous CaO FeO+MgO Figure 7.1 ACF diagram for plutonic rocks of the northeastern CIP. Plagioclase-(FeO + MgO) line defines ASI (aluminum saturation index) = 1, which divides peraluminous and metaluminous granitoids (Chappell and White, 1992). •"•i—r ' ' 1 •1 J- 1 1 '1 . . 1 1 . . i -j—1—r- . . | 1 . 1 1 4- 8 - - + N 4- R o 6 - S(R + M) — - Is* - - - VI OS • ,. xX#xNf 7S rv (D - - P-, 4 - •5 - - * N* 2 - 1 1 1 1 1 ,i.. 1 1 1 11 1 . 1 1 1 t 1 _I 1....1 1,1,1 1 1 w 1, 55 60 . 65 •70 75 Si02(wt%) Figure 7.2 Harker diagram for total Fe as Fe203 in the Kashmar granitoid. The diagram represents the partial melting of the source rock (S) to produce restite (R) and a liquid as 'minimum-melt' (M) composition (Chappell et al., 1987). The magma at its source consists of (R+M) and varying degrees of separation of R from M generated a range of magma and rock compositions, illustrated by granodiorite (x), granite (*) and alkali feldspar granite (G). 10 • • mantle 1 , , 1 , 1 £Nd ° X • *x i * X X X x 1 crust 1 -10 1— , i 0.70 0.71 0.72 0.73 0.74 —,0.7 5 _0.7^ 6 Initial ^Sr/^Sr + Kashmar x Bornavard n Kuh Mish 87 86 Figure 7.3 Diagram showing initial Sr/ Sr versus eNd values for Kashmar granitoid (+), Bornavard granitoid (x) and Kuh Mish intrusions (D). The boundary between mantle and crust (SN 30 0 £ Q- 100 O. Z 30 VAG 4 Syn-COLG 10 t 3 : a • i a i i i i mi 100 1000 Y ppm Figure 7.4 The Nb-Y discrimination diagram for granitoid rocks of the northeastern CIP. The fields (Pearce et al, 1984; Forster et al, 1997) show volcanic-arc granites (VAG), syn-collisional granites (Syn-COLG), within-plate granites (WPG) and ocean-ridge granites (ORG). The broken line is the field boundary for ORG from anomalous ridges. Symbols: Kashmar granitoid (4), Bornavard granitoid (x) and Kuh Mish intrusions {•). 1 l I l l 11| 1—r—TT 1 1 M I I ll| I a • ' '' " a1 ' ' '' I I I I I I 111 10 - 100 1000 (Y 4- Nb) ppm Figure 7.5 The Rb versus (Y + Nb) discrimination diagram for granitoid rocks of the northeastern CIP. The fields and symbols are as Figure 7.4. Xable 2.1 Isotopic age data for some Iranian granitoids and volcanic rocks Locality Lithology Material dated Method Age (2a) Ref. Regional-Aureole Granitoids (CIP) Chapedony granite whole rock Rb/Sr . 541-550 (A) Doran granite biotite-whole rock Rb/Sr 175±5 (B) Contact-Aureole Granitoids (CIP) Mashhad cjuartz diorite biotite K/Ar 146±3 (Q granite biotite K/Ar 120±3 (C) Shirkuh granite K-feldspar K/Ar 186-159 (D) Airakan granite whole rock Rb/Sr 165±8 (D) granite biotite K/Ar 113±9 (D) ' Torbat-e-Jam granite biotite K/Ar 153±5 (E) Sangan granite biotite K/Ar 38.4+1 (F) Contact-Aureole Granitoids (SSMZ) Hamadan pegmatite muscovite-whole rock Rb/Sr 104±3 (G) pegmatite muscovite K/Ar 82.8±3 (G) gabbro biotite-whole rock Rb/Sr 88.5 (G) gabbro biotite K/Ar 89.6±3 (G) granodiorite biotite-whole rock Rb/Sr 68+2 (G) granodiorite biotite K/Ar 63.8±2.5 (G) Subvolcanic Granitoids (U-DVB) Natanz gabbro biotite-whole rock Rb/Sr 33.5±1.2 (H) granite biotite-whole rock Rb/Sr 25.5±0.5 (H) Karkas granodiorite whole rock Rb/Sr 78 (D) granodiorite biotite K/Ar 38-33 (D) Jebal-e-Barez granite biotite K/Ar 24+0.1 0) Volcanic Rocks (CIP) Kashmar dacite hornblende K/Ar 57.2±3.7 (J) dacite biotite K/Ar 43.7+1.7 (J) Gonabad andesite whole rock K/Ar 61±2 (K) Bejestan andesite whole rock K/Ar 54+2 CK) Volcanic Rocks (U-DVB) /t \ Shahrbabak trachyte albite Ar/Ar 37.5±1.4 (L) tephriphonolite K-feldspar Ar/Ar 22.8±3.2 (L) quartz monzonitei hornblende Ar/Ar 16.9±0.2 (L) dacite biotite K/Ar 16.4±1 (M) Aghda trachyandesite hornblende K/Ar 15.7+1 • (M) sanidine K/Ar 6.5+1 (M) -ikJiUU.JJLV*Islamic' 1Penin< WJJJijnL il a trachyte Zone; U-DVB = Abbreviations: CIP = Central Iraii Plate; S-SMZ = Sanandaj-Sirjan Metamorphic Urumiyeh-Dokhtar Volcanic Belt. 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For rocks of the current study, total Fe is assumed as Fe„0, Locality KASH BOR TAK KUH JOBTK NATAN MASH LFB LFB LFB PRB WPRB EPRB No. of samples 29 13 5 13 60 26 18 1074 704 43 323 174 149 Granite-type I-type I-type I-type I-type I-type I-type S-type I-type S-type A-type I-type I-type . I-type Si02 66.80 69.84 76.41 64.68 67.80 61.65 71.45 69.50 70.91 73.47 64.63 63.41 66.05 Ti02 0.47 0.44 0.14 0.36 0.41 0.69 0.08 0.41 0.44 0.30 0.65 0.64 0.65 A1203 14.89 13.11 12.01 14.34 14.15 16.37 16.39 14.21 14.00 12.88 15.94 15.70 16.23 FeA 3.94 3.84 2.21 5.03 4.01 2.07 0.35 1.01 0.52 0.90 1.20 1.55 0.78 FeO n.a. n.a. n.a. n.a. n.a. 3.40 0.44 2.22 2.59 1.63 3.19 3.60 2.71 MnO 0.06 0.05 0.03 0.08 0.06 0.12 0.07 0.07 0.06 0.06 0.08 0.09 0.06 MgO 1.45 1.99 I 0.4 3.13 1.84 3.00 0.24 1.38 1.24 0.30 2.15 2.71 1.49 CaO 2.97 2.72 1.02 5.56 3.31 6.05 1.50 3.07 1.88 1.06 5.10 5.70 4.41 Na20 3.81 3.95 3.10 3.46 3.71 3.44 3.73 3.16 2.51 3.50 3.62 3.45 3.82 K20 3.55 2.63 3.92 0.92 2.81 1.93 5.61 3.48 4.09 4.62 1.95 1.69 2.26 P2O5 0.12 0.08 0.04 0.07 0.09 0.08 0.12 0.11 0.15 0.07 0.13 0.10 0.15 ASI 0.97 0.95 1.12 0.92 0.97 0.87 1.10 0.985 1.179 n.a. n.a. n.a. n.a. Trace elements (ppm) Ba 494 516 802 140 448 n.a. 242 519 440 547 641 451 863 Rb 97 66 107 15 73 63 269 164 245 188 60 49 73 Sr 249 90 50 154 177 285 134 235 112 96 375 268 501 Pb 8 10 8 2 7 n.a. 44 19 27 27 10 8 12 Th 13 15 20 2 12 n.a. 6 20 19 24 7.2 6 8.6 U 3 2 3 1 2 n.a. 2 5 5 5 1.5 1.4 .1.7 Zr 175 234 134 57 159 100 25 150 157 322 139 130 150 Nb 10 10 9 2 8 10 19 11 13 26 6.7 5.1 8.6 Y 24 48 41 15 29 26 7 31 32 71 19 24 12 La 25 30 27 7 22 n.a. n.a. 31 27 55 16 12 20 Ce 50 67 58 9 45 n.a. n.a. 66 61 130 35 28 43 Sc 9 17 . 4 24 14 n.a. 2 13 11 11 14 19 8 V 71 48 4 117 70 n.a. 27 57 49 9 85 115 50 Cr 8 15 2 74 23 n.a. 26 20 30 2 47 67 24 Mn ,507 425 210 735 514 n.a. n.a. n.a. n.a. n.a. 587 704 451 Ni 2 11 2 26 9 n.a. 9 8 11 2 13 18 6 Cu 10 9 8 29 14 n.a. 6 9 9 5 10 16 3 Zn 32 42 110 38 J 42 n.a. 31 48 59 j 95 76 66 88 Sn <5 <5 <5 <5 <5 n.a. n.a. 6 10 8 n.a. n.a. n.a. Ga 15 16 15 12 15 n.a. 19 16 18 22 18.3 16.4 20.5 Abbreviations and the source of data: KASH = Kashmar granitoid; BOR = Bornavard granitoid; TAK = Taknar Rhyolite; KUH = Kuh Mish intrusions; KBTK = the average of 60 analyses in this study, from Kashmar, Bornavard, Taknar and Kuh Mish regions; NATAN = Natanz intrusive rocks that are subvolcanic and form part of the U-DVB in Iran (Berberian, 1981); MASH = Mashhad Granite that is the largest contact-aureole granite in northeastern Central Iran Plate (Iranmanesh and Sethna, 1998); LFB = Lachlan Fold Belt, Australia (Chappell and White, 1992); PRB = Peninsular Ranges Batholith, USA (Silver and Chappell, 1988); WPRB = Western Peninsular Ranges Batholith; EPRB = Eastern Peninsular Ranges Batholith. For easy compilation, the column data for KBTK (all analyses of this study) is bold. 240 APPENDIX 1 ANALYTICAL METHODS 1.1 SAMPLE COLLECTION With one of the main aims of the thesis being completion of the granite data-base for the Kashmar, Bornavard, and Kuh Mish areas, granite sampling were undertaken from several traverses across the granitoid bodies. To ensure something approaching representative sampling of all intrusions at least three samples were selected in more detail from each variety at different locality. All samples were microscopically studied and modally analysed; the general homogeneity of bodies reduced selected samples to 60 choices that govern both external and internal variability, and the validity of this approach. Granite samples were collected on field trip to Iran in 1994 by author and Dr. Paul F. Carr. The samples were taken by sledge hammer as to use only fresh, internal rock, and free of imperfections e.g. enclaves, viens, etc., and were of hand size. The samples were individually put into plastic bags, then wrapped in paper or cloth pieces, and transported in suitable boxes for shipment to Wollongong. 1.2 SAMPLE PREPARATION Samples were first cut by saw and carefully removed the weathered parts (although rare) and obtained rock slabs; then were broken down by hammer to small rock fragments. All rock fragments were washed and dryed in room temperature. Using a hydraulic rock 241 spliter (with tungsten-carbide plates), the rock fragments were split into small pieces roughly <1 cm3. This aggregate was milled in partial for 1 to 3 minutes by a tungsten- carbide mill (TEMA) embarked upon a shaking machine, to result at least 500 g of fine powder for each sample. Then the resulting powder was homogenised and bagged. A representative sample was taken from this powder used for analytical processes such as major oxides, trace and rare earth elements, and isotope determination. After preparation of each sample, the hydraulic spliter was brushed then cleaned up by acetone. The TEMA was washed with hot water, and followed by washing with acetone. Additionally, where the rocks were somewhat geochemically different from the previous sample (e.g. abundance of hydrous minerals), the TEMA was also cleaned by milling roughly 100 gm of clean sandbeach. 1.3 ANALYTICAL METHODS Variety of modern and suitable sample preparation (Saheurs and Wilson, 1993) and laboratory methods of analysis (Potts, 1993) were employed during this study for analysing of samples on whole rock powders, separated minerals and thin section, these being: - X-Ray fluorescence (XRF) for whole rock major and trace element analyses, - Instrumental Nuetron Activation Analysis (INAA) for trace and rare earth elements (REE), - Total volatile contents or "loss on ignition" (LOI) determination, - Isotope determinations (Rb, Sr, Nd) and - Wave lenght dispersive electron microprobe analysis of minerals. 242 1.3.1 WHOLE ROCK, MAJOR AND TRACE ELEMENT ANALYSES X-ray fluorescence (XRF) spectrometry was performed on the whole rock powders using both fused glass disks for major elements (using a SIEMENS SRS300 XRF) and pressed powder pellets for trace elements (using automated Phillips PW1450 spectrometer) at the Australian National University, following the methods of Norrish and Hutton (1969) and Norrish and Chappell (1977). Major and trace element abundances were determined for 60 representative samples. Other trace elements (REE, Sc, Cr, Sb, Hf, Th, and U) were determined on 20 representative samples by Instrumental Neutron Activation Analysis (INAA). Cobalt, tantalum and wolfram are not reported due to enrichment as a result of contamination during powder preparation using tungsten-carbide mill. 1.3.2 TOTAL VOLATILE DETERMINATION Total volatile or lose on ignition (LOI) were determined gravimetrically at the University of Wollongong. For each sample approximately 0.5 gm of powder was accurately weighed into a crucible, that had been previously washed, heated and cleaned by acetone. Crucible and sample were placed into a furnace at 1000°C for 12 hours. Then, samples took out from the furnace and cooled at the room temperature, then accurately weighed after five minutes cooling in a desiccator. The decrease in sample weight is due to loss of volatiles, and is therefore a direct measure of all volatiles + including H20 , H20", C02. The volatile contents reported as LOI which obtained from 100 x weight lost/weight of sample. 1.3.3 ISOTOPE DETERMINATION (Rb, Sr and Nd) After careful microscopic study on thin sections, 8 samples with the most freshest 243 biotite were selected for Rb/Sr age dating on biotite-whole rock pair. Samples were washed and dried in room temperature, then they were broken up into pieces with a size of 5-10 mm by a tungsten-carbide splitter. Samples were further ground to a grain size of less than 500 urn by tungsten-carbide ring mill. The powder was washed with tab water and slowly dried. The individual grain-size fractions were separated using a magnetic separator tuned up by electric current of 0.8 A and 20° inclination. Dry shaking (by hand) on white paper was very effective for the initial enrichment of small biotite flakes. A purity of >99% could be attained only by handpicking under microscope. To remove mineral inclusions (e.g. apatite), the hand-picked biotites were ground in an opal mortar, then washed by alcohol. As a final step, biotite grains were carefully controlled under microscope. At last all biotite fractions were washed by acetone and distilled water. For each sample, about 0.1 mg of purified biotite was dissolved in teflon capsule, mixed with a Rb/Sr spike, by using a mixture of HN03 and HF acids. Analytical processes followed by isotope dilution method. Rb and Sr were separated and concentrated for mass spectrometric analysis using standard ion-exchange procetures at School of Geosciences, University of Wollongong, by Dr. P. F. Carr. Also, a representative subset of 20 whole rock powders (0.1 mg) were analysed for Rb, Sr and Nd isotope ratios, at School of Geosciences, University of Wollongong and CSIRO, Sydney by Dr. P. F. Carr. Isotopic analyses undertaken on a VG 354 mass spectrometer in CSIRO, Sydney. Replicate analyses of SRM 987 gave 86Sr/88Sr = 0.710251 ± 28 (external precision at 2rj, n = 17) and the JM-Nd standard gave 146Nd/144Nd = 0.511111 ± 12 (external precision at 2a, n = 17). 87Sr/86Sr normalised to 86Sr/88Sr = 0.1194; 2a analytical uncertainty for 87Sr/86Sr is ± -0.00005.143Nd/144Nd normalised to 146Nd/144Nd = 0.7219; 244 2a analytical uncertainty for £Nd is ± -0.5 units. Ages, 87Rb/86Sr, 147Sm/144Nd, SN 1.42 xlO^a"1. The MSWD, 2a uncertainty in age and 87Sr/86Sr (initial) were calculated by standard percents of 0.8 and 0.008, respectively for X (87Rb/86Sr) and Y (87Sr/86Sr) axes. 1.3.4 MICROPROBE MINERAL ANALYSIS Mineral analyses reported in this work were undertaken on the CAMECA SX50 electron microprobe fitted with 5 wavelength dispersive spectrometers (WDS) at the School of Earth Sciences, Macquarie University. The WDS was operated using an accelerating voltage of 15 kV and a beam current of 20 nA with 10 pm beam size. The count times are 30 seconds on peak, 30 seconds on background for Si, Ti, Al, Cr, Fe, Mn, Mg, Ca, Ni, 10 seconds on peak, 10 seconds on background for Na and K. Typical natural mineral standards were run for each element program include Si, Al, Na- albite, Ti-rutile, Cr-chromite, Fe-hematite, Mn-spessartine, Mg-forsterite, Ca- wollastonite, K-orthoclase and synthetic mineral standard includes Ni-Ni-olivine. Correction proceture is based on the Cameca PAP method. Lower levels of detection (wt% oxide) are Si-0.04; Ti-0.03; Al-0.02; Cr-0.05; Fe-0.05; Mn-0.05; Mg-0.04; Ca- 0.03; Na-0.04; K-0.04; Ni-0.06. Precision (rsd at wt% in brackets) for elements include Si-0.21 (50 wt%); Ti-1.72 (1.0 wt%); Al-0.29 (20 wt%); Cr-3.46 (1 wt%); Fe-0.94 (10 wt%); Mn-3.18 (1 wt%); Mg- 0.46 (10 wt%); Ca-0.55 (10 wt%); Na-0.94 (10 wt%); K-0.96 (10 wt%); Ni-3.22 (1 245 wt%). Structural formulae for amphibole analyses were calculated on the basis of 23 oxygens (assumed anhydrous) with site allocation as suggested by Robinson et al. (1982, p.5-6). Ferric iron contents were estimated by utilizing assumptions of crystal-chemical limitations on cation substitution and total cation assumptions as outlined by Robinson et al. (1982, p.6-9). In this case the predominant option used was total cations exclusive of K, Na and Ca calculated to 13. This succeeded in successful atomic formulae for all analyses. Amphibole nomenclature follows the recommendations of Leake (1978) and Deer et al. (1997). About 20 analyses have been done on a certain mineral from every sample (including cores and rims). For each sample, representative analyses of a particular mineral were selected and listed in Appendix 3. 246 APPENDIX 2 MODAL MINERALOGY AND C. I. P. W. NORMS ABBREVIATIONS USED Rocks: A.F.G. = Alkali feldspar granite Grd. = Granodiorite Q.M.D. = Quartz monzodiorite For each sample, the modal mineral contents were determined by recognition of 500 spots under Research Microscope (Lietz Orthoplan), equiped by a digital point counter. C. I. P. W. Norms: Q = Quartz C = Corundum Or = Orthoclase Ab - Albite An = Anorthite Di = Diopside Hy = Hypersthene 01 = Olivine Mi = Montmorillonite 7Z = niite Ap = Apatite Diff. Index = Differentiation Index C. I. = Colour Index Norm Plag. = Normative Plagioclase Composition = 100 An/(An+Ab) 3 o o> =» "T > 5? ?* "2 z H^S^& ^OPL_ P* CD •- g*^ 3 O -•< D) o 3 3 .—* -i-o 5: SLCQ Q. CD ~> 3" CD O- O CD m CD ij S- 2. X o =i CO 3 ^- 8-® CD CD " 3 CO 4* 4s. fO 4s. ro co pS*>Sp-J.|N)J00(BJ.0!)OS A -*• -* •OO *S lo O 4*. *4s. Cd oCOo o*4>o -«CO• "-»• 4s. *0 4s. CO fe COO-i-iUO)CnOl4i>CSO!OOM p p O ^1 A fo io b b bo Ol CO M OI -J- -a 4*. _i. 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CO 4s CO 1 J A A A A ^ A -»• CDS 0l- 0)0)OOOO4-proC0W 4s O CO -s! 4s O O O p 4s S. <" b'^bb'^isbbbb'-tbbb-si fo fo fo fo is io b fo b i _avjOO- OC04sOOlOrO-40C04s. 7s 3" 2 CO 4s -a CO _ co ->• A A -* CO CO o -4 4s 4s O CD pproppppp O N 00 CD CO O poppppppco a-S-2 -si fo b-sicocbocoo-sjcooooc bo b fo b b o fofofoisiobbis4s oo co --I ro 252 APPENDIX 3 ELECTRON MICROPROBE ANALYSES ABBREVIATIONS USED Spot: R = Rim C = Core The number located on the left side of spot (e.g. 1-R or 1-C) indicates number of grain. The number located on theright side of spot (e.g. R-l or C-l) indicates number of analyses. Rocks: (3rd = Granodiorite A.F.G. = Alkali feldspar granite Tona. = Tonalite Minerals: Mag. = Magnetite Titmag. = Titanomagnetite Hmen. = Ilmenite Plag. = Plagioclase K-feld.= K-feldspar P = Phenocryst G = Groundmass The letters m and r, respectively define mineral and rock where they are subscripted. CHEMICAL CALCULATIONS For hornblende analyses Fe3+ estimated using 13CNK normalisation from Robinson (1982), and Mg* = Mg/(Mg + Fe2+). " For biotite analyses total Fe is assumed as FeO contents. For clinopyroxene analyses Fe* = total Fe + Mn (Deer et al., 1992). End members were calculated based on En = 100Mg/(Ca + Fe* + Mg); Wo = 100Ca/(Ca + Fe* + Mg) and Fs = 100Fe*/(Ca + Fe* + Mg). TABLE ORGANISATION In each table, rock types are listed according to decreasing in modal contents of plagioclase e.g. tonalite, granodiorite, Granite, etc. Also, for each rock-type, tables start with analyses containing the lowest Si02 content of the whole-rock and continue towards the highest Si02 content. 253 KASHMAR GRAMTOID: Appendix 3.1 Plagioclase (84 analyses) Appendix 3.2 Hornblende (22 analyses) Appendix 3.3 Biotite (24 analyses) Appendix 3.4 Fe-Ti oxides (22 analyses) Appendix 3.5 K-feldspar (3 6 analyses) BORNAVARD GRANITOID: Appendix 3.6 Plagioclase (44 analyses) Appendix 3.7 K-feldspar (12 analyses) Appendix 3.8 Hornblende (24 analyses) Appendix 3.9 Biotite (24 analyses) Appendix 3.10 Fe-Ti oxides (16 analyses), allanite (4 analyses) TAKNAR RHYOLITE: Appendix 3.11 Feldspars (8 analyses), Appendix 3.11 Biotite (2 analyses) and Appendix 3.11 Fe-Ti oxides (2 analyses) KUH MISH INTRUSIONS: Appendix 3.12 Clinopyroxene (6 analyses) Appendix 3.13 Plagioclase (21 analyses) Appendix 3.14 Hornblende (16 analyses) Appendix 3.15 Biotite (8 analyses) and Appendix 3.15 Fe-Ti oxides (4 analyses) 254 0££ £*?££>3* CO -J S3 to VO to ON 392 o oo •—« to O O Js O 1— to so O NO O ;IS IO U P> Ui Ul i—> CO p p t I—• i—» i—> "o isui vo oo o to O" vo b b i—> 0 00 00 "o H- ~0 © h— ~J O h- 00 NJtOUit-aONtOOUi 00 t04stOt-^0©©Ui 7* S3 © to ON O 4s to LO O 1— 4s O O S3 tu o © 00 O Ui to © -J HH O o o _ to i-> sl O W NO o o 4s ON Ul Ui S> na VO ON 00 ON '© '© '© '© © VO © VO to totooonaooooo tOi—• Jl 00 Ui s) IO Ul Ul OJ S3 NO S3 Ul 9 v W O H U O O © NO NO o © NO O ON O 00 ON O 00 o 0©4s©U>U«©4s NOi-aNoi—'tO©©tO NO "ON 4S H-tO©tOVO©©4s S>ONtOtO©tOtOON bJ WW 0N0N00tOtsJO4s4s t—» ON Q ON to to O S> Ul to © I—> Ul S3 to 4s 0 0 © © o © -o © -4 y» © vo 9 W P" 00 tO VO 0 1—1 Ul b 4s lo© ON tO Ui Si i—* CO is O ONS3 © 4s h-' to to to Ul O H-> H-• to OOONHaQOSJOtO Ul 4s ON © NO Ul Ui to to NO S> Ul S9 H-' vo 00 0 © to Ui © © vo©oN©-jy»©po 00 00 to 0 O 4s O Ul is "© is 4s to NO i—' VO 00 © SJ 00 NO tO 1—1 -4 H-» tO 4s 00 © 00 _p».(jj|_a4i.NOtOS)S> ON ON 4s 1— H-» OH U S3 -O. 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Mg* = Mg/(Mg + FeH, Sample No. R15908 R15908 R15908 R15908 R15908 R15908 R15910 R15910 Rock Name Grd. Grd. Grd. Grd. Grd. Grd. Granite Granite Rock Si02 62.30 62.30 62.30 62.30 62.30 62.30 63.42 63.42 Spot 1-R 1-C 2-R 2-C 3-R 3-C 1-C-l l-C-2 Si02 47.43 45.71 45.36 44.56 45.43 45.96 49.97 52.36 Ti02 1.32 1.75 1.84 2.37 1.71 1.71 0.63 0.13 AI203 6.51 7.17 7.39 8.41 7.20 7.11 4.43 2.35 MgO 12.35 11.17 12.30 12.95 11.48 12.04 15.05 13.92 CaO 10.81 11.21 11.20 11.22 11.03 11.10 11.70 11.90 MnO 0.69 0.68 0.39 0.24 0.59 0.40 0.39 0.60 FeO 16.71 17.56 16.31 14.54 17.09 16.77 13.35 15.39 Na20 1.16 1.43 1.48 1.92 1.25 1.33 0.86 0.42 K20 0.58 0.76 0.82 0.60 0.79 0.72 0.42 0.13 Total 97.56 97.44 97.09 96.81 96.57 97.14 96.83 97.20 Si 6.924 6.800 6.720 6.588 6.780 6.795 7.251 7.619 Ti 0.144 0.196 0.205 0.264 0.192 0.190 0.069 0.014 Al 1.120 1.257 1.290 1.465 1.267 1.239 0.756 0.403 Mg 2.687 2.476 2.716 2.854 2.554 2.653 3.254 3.018 Ca 1.690 1.787 1.778 1.777 1.764 1.758 1.819 1.855 Mn 0.085 0.085 0.049 0.031 0.074 0.050 0.048 0.074 Fe 2.039 2.185 2.021 1.798 2.134 2.073 1.620 1.873 Na 0.329 0.412 0.424 0.551 0.361 0.381 0.243 0.118 K 0.107 0.144 0.149 0.114 0.150 0.136 0.077 0.026 Total 15.125 15.342 15.352 15.442 15.276 15.275 15.137 15.000 A1IV 1.076 1.200 1.280 1.412 1.220 1.205 0.749 0.381 Al71 0.044 0.057 0.010 0.053 0.047 0.034 0.007 0.022 Fe3+ 0.930 0.623 0.729 0.612 0.748 0.758 0.650 0.475 Fe2+ 1.109 1.562 1.292 1.186 1.386 1.315 0.970 1.370 Ca(M3) 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 Fe(M4) 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.027 Ca(M4) 1.690 1.787 1.778 1.777 1.764 1.758 1.819 1.855 Na(M4) 0.310 0.213 0.222 0.223 0.236 0.242 0.181 0.118 Na(A) 0.019 0.199 0.202 0.328 0.125 0.139 0.062 0.000 K(A) 0.107 0.144 0.149 0.114 0.150 0.136 0.077 0.026 Mg* 0.71 0.61 0.68 0.71 0.65 0.67 0.77 0.68 Fe/(Fe+Mg) 0.43 0.47 0.43 0.39 0.46 0.44 0.33 0.38 MgO/FeO 0.74 0.64 0.75 0.89 0.67 0.72 1.13 0.90 MgO^MgO, 6.40 5.79 6.37 6.71 5.95 6.24 6.38 5.90 FeOJFeOr 3.49 3.67 3.41 3.04 3.57 3.50 2.75 3.17 Ti02m/Ti02r 2.10 2.78 2.92 3.76 2.71 2.71 1.07 0.22 262 Appendix 3.2 (Continued): Sample No. R15910 R15910 R15918 R15918 R15918 R15918 R15918 R15909 Rock Name Granite Granite Granite Granite Granite Granite Granite Granite Rock Si02 63.42 63.42 65.33 65.33 65.33 65.33 65.33 71.81 Spot 2-R 2-C 1-R-l l-R-2 1-C 2-R 2-C 1-R SiCb 48.45 49.20 46.13 47.43 45.74 46.58 45.85 47.18 Ti02 1.23 1.00 1.41 0.69 1.76 1.12 1.59 1.09 A1203 5.51 5.08 6.61 5.74 7.56 6.49 7.12 5.74 MgO 14.23 14.42 10.64 10.61 11.63 10.59 10.45 12.62 CaO 11.41 11.14 10.86 10.56 11.14 10.79 10.88 11.03 MnO 0.45 0.55 0.46 0.62 0.36 0.88 0.69 0.85 FeO 14.00 13.85 18.93 19.76 16.94 18.83 18.53 16.41 Na20 1.19 1.10 1.18 0.99 1.37 1.01 1.15 1.44 K20 0.54 0.50 0.69 0.51 0.76 0.69 0.81 0.57 Total 97.01 96.84 96.91 96.91 97.28 96.98 97.07 96.93 Si 7.064 7.147 6.888 7.041 6.774 6.933 6.843 6.968 Ti 0.135 0.109 0.158 0.076 0.195 0.126 0.179 0.121 Al 0.946 0.870 1.163 1.005 1.320 1.138 1.253 1.000 Mg 3.092 3.123 2.368 2.346 2.567 2.349 2.326 2.779 Ca 1.782 1.733 1.738 1.679 1.768 1.722 1.740 1.745 Mn 0.055 0.068 0.059 0.077 0.045 0.111 0.087 0.107 Fe 1.707 1.683 2.364 2.453 2.097 2.344 2.312 2.026 Na 0.335 0.309 0.343 0.284 0.392 0.292 0.333 0.410 K 0.101 0.093 0.131 0.096 0.144 0.130 0.154 0.110 Total 15.217 15.135 15.212 15.057 15.303 15.145 15.227 15.266 Af 0.936 0.853 1.112 0.959 1.226 1.067 1.157 1.000 Af" 0.010 0.017 0.051 0.046 0.094 0.071 0.096 0.000 Fe3+ 0.658 0.748 0.795 0.917 0.670 0.876 0.736 0.810 Fe2+ 1.049 0.935 1.569 1.499 1.427 1.468 1.576 1.216 Ca(M3) 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 Fe(M4) 0.000 0.000 0.000 0.037 0.000 0.000 0.000 0.000 Ca(M4) 1.782 1.733 1.737 1.679 1.768 1.722 1.740 1.745 Na(M4) 0.218 0.267 0.263 ' 0.284 0.232 0.278 0.260 0.255 Na(A) 0.117 0.042 0.080 0.000 0.160 0.014 0.073 0.155 K(A) 0.101 0.093 0.131 0.096 0.144 0.130 0.154 0.110 Mg* 0.75 0.75 0.60 0.61 0.64 0.62 0.60 0.70 Fe/(Fe+Mg) 0.36 0.35 0.50 0.51 0.45 0.50 0.50 0.42 MgO/FeO 1.02 1.04 0.56 0.54 0.69 0.56 0.56 0.77 6.43 7.05 6.42 6.33 18.04 MgOm/MgOf 6.03 6.11 6.45 4.21 4.68 4.61 8.21 FeOm/FeOr 2.89 2.86 4.71 4.92 3.59 2.29 3.24 4.19 Ti02n/Ti02r 2.09 1.70 2.88 1.40 263 Appendix 3.2 (Continued) Sample No. R15909 R15909 R15909 R15909 R15909 R15909 Rock Name Granite Granite Granite Granite Granite Granite Rock Si02 71.81 71.81 71.81 71.81 71.81 71.81 Spot 1-C 2-R 2-C 3-R-l 3-R-2 3-C Si02 48.97 47.59 48.27 52.09 53.05 53.05 Ti02 1.11 1.21 1.13 0.48 0.16 0.17 1.74 A1203 4.67 5.67 5.74 2.85 1.89 MgO 13.74 12.60 13.06 14.72 15.61 15.71 CaO 10.83 10.86 11.23 12.20 12.24 12.23 MnO 0.88 0.69 0.59 0.33 0.43 0.43 FeO 14.90 16.08 15.49 13.84 13.00 13.14 Na20 1.17 1.46 1.45 0.56 0.46 0.46 0.15 K20 0.51 0.61 0.52 0.25 0.15 Total 96.78 96.77 97.48 97.32 96.99 97.08 Si 7.155 7.030 7.066 7.557 7.671 7.664 Ti 0.122 0.134 0.123 0.052 0.018 0.018 Al 0.803 0.987 0.990 0.487 0.323 0.296 Mg 2.991 2.775 2.850 3.184 3.364 3.382 Ca 1.695 1.718 1.762 1.896 1.896 1.893 Mn 0.108 0.087 0.073 0.041 0.053 0.053 Fe 1.821 1.987 1.897 1.679 1.572 1.587 Na 0.331 0.418 0.413 0.158 0.128 0.130 K 0.095 0.115 0.098 0.046 0.028 0.027 Total 15.121 15.251 15.272 15.100 15.053 15.050 Alw 0.803 0.970 0.934 0.443 0.323 0.296 AF 0.000 0.017 0.056 0.044 0.000 0.000 Fe3+ 0.827 0.716 0.599 0.299 0.349 0.397 Fe2+ 0.994 1.271 1.298 1.380 1.223 1.190 Ca(M3) 0.000 0.000 0.000 0.000 0.000 0.000 Fe(M4) 0.000 0.000 0.000 0.000 0.000 0.000 Ca(M4) 1.695 1.718 1.762 1.896 1.896 1.893 Na(M4) 0.305 0.282 0.238 0.104 0.104 0.107 Na(A) 0.026 0.136 0.175 0.054 0.024 0.023 K(A) 0.095 0.115 0.098 0.046 0.028 0.027 Mg* 0.75 0.69 0.69 0.70 0.73 0.74 Fe/(Fe+Mg) 0.40 0.42 0.40 0.34 0.32 0.32 MgO/FeO 0.92 0.78 0.84 1.06 1.20 1.20 18.66 21.03 22.30 22.44 MgOm/MgOr 19.62 18.00 7.75 6.92 6.50 6.57 FeOm/FeOr 7.45 8.04 1.83 0.62 0.64 Ti02m/Ti02r 4.28 4.65 4.35 264 1 5£ *i1 H *? , ECfc \> tf. ffi >0 V P CO o 09 si < lo?olf^oo ° B €8 oFnT PV n- 9 00 00 &. + O § p CO • .co. tpj NO |_a na na tO £S92 I Ul © S3 © S3 © to © Ul C300©NO©S3tOtOON Ul Ul Lu "sl © 4s "© 00 © to Js. o^ ON ON '© S3 S3 S3 Ui 4s. rS3* jo P. 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Rock Si0 58.09 2 58.09 58.09 58.09 58.09 58.09 69.45 69.45 Spot 1-R 1-C 2-R 2-C 3-R 3-C 1-R 1-C 46.41 SiOs 48.71 45.93 46.34 45.78 47.36 44.45 43.89 Ti02 1.18 0.85 1.74 1.71 0.89 1.08 1.57 1.53 A1 0 6.34 2 3 4.75 7.09 7.21 7.13 5.46 6.85 6.92 MgO 10.76 12.15 11.91 12.55 10.20 11.39 9.26 9.25 11.57 CaO 11.72 11.12 11.22 11.80 11.52 10.42 10.49 MnO 0.43 0.38 0.24 0.27 0.40 0.43 0.46 0.54 FeO 18.53 17.34 16.59 15.71 19.30 18.16 20.38 20.95 Na20 0.74 0.59 1.39 1.25 0.78 0.66 0.88 1.28 K20 0.68 0.44 0.67 0.66 0.73 0.56 0.76 0.80 Total 96.64 96.93 96.68 96.92 97.01 96.62 95.03 95.65 Si 6.961 7.210 6.830 6.830 6.878 7.066 6.804 6.716 Ti 0.133 0.094 0.195 0.189 0.101 0.121 0.181 0.177 Al 1.121 0.828 1.243 1.254 1.262 0.961 1.236 1.247 Mg 2.406 2.680 2.639 2.758 2.285 2.533 2.112 2.109 Ca 1.860 1.858 1.771 1.772 1.900 1.841 1.708 1.719 Mn 0.054 0.047 0.030 0.034 0.050 0.054 0.060 0.070 Fe 2.325 2.146 2.063 1.936 2.424 2.265 2.609 2.680 Na 0.216 0.168 0.401 0.358 0.226 0.192 0.261 0.381 K 0.129 0.084 0.127 0.123 0.139 0.107 0.149 0.157 Total 15.205 15.115 15.299 15.254 15.265 15.140 15.120 15.256 1.039 0.790 1.170 1.170 1.122 0.934 1.196 1.247 w Al 0.082 0.038 0.073 0.084 0.140 0.027 0.040 0.000 34 Fe 0.630 0.622 0.637 0.681 0.615 0.684 0.946 0.993 *\2+ 1 Fe 1.695 1.524 1.426 1.255 1.809 1.581 1.645 1.687 Ca(M3) 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 Fe(M4) 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 Ca(M4) 1.860 1,858 1.771 1.772 1.900 1.841 1.708 1.719 Na(M4) 0.140 0.142 0.229 0.228 0.100 0.159 0.261 0.281 Na(A) 0.076 0.026 0.172 0.130 0.126 0.033 0.000 0.100 K(A) 0.129 0.084 0.127 0.123 0.139 0.107 0.149 0.157 Mg* 0.59 0.64 0.65 0.69 0.56 0.62 0.56 0.56 Fe/(Fe+Mg) 0.49 0.44 0.44 0.41 0.51 0.47 0.55 0.56 MgO/FeO 0.58 0.70 0.72 0.80 0.53 0.63 0.45 0.44 MgOm/MgOr 2.45 2.76 2.71 2.85 2.32 2.59 6.43 6.42 FeOm/FeOr 2.25 2.10 2.01 1.90 2.34 2.20 8.25 8.48 Ti02m/Ti02r 1.22 0.88 1.79 1.76 0.92 1.11 2.53 2.47 277 Appendix 3.8 (Continued): Sample No. R15953 R15953 R15953 R15953 R15953 R15953 R15943 R15943 Rock Name Grd. Grd. Grd. Grd. Grd. Grd. Grd. Grd. Rock Si02 69.45 69.45 69.45 69.45 69.45 69.45 71.32 71.32 Spot 2-R 2-C 3-R 3-C 4-R 4-C 1-R 1-C Si02 44.86 44.67 44.48 45.00 41.46 45.15 52.70 52.34 Ti02 1.52 1.53 1.53 1.56 0.31 0.32 0.01 0.14 A1A 7.13 6.81 6.68 6.64 11.98 8.19 1.19 2.98 MgO 9.75 9.45 9.90 10.22 7.01 9.87 12.62 14.49 CaO 10.86 10.27 10.58 10.62 11.67 11.77 12.50 12.23 MnO 0.44 0.48 0.33 0.24 0.41 0.37 0.16 0.31 FeO 19.54 21.03 18.87 19.67 20.70 17.82 17.99 14.36 Na20 0.85 1.00 1.17 1.22 1.18 0.93 0.13 0.35 K20 0.80 0.83 0.77 0.79 0.54 0.21 0.07 0.08 Total 95.75 96.07 94.31 95.96 95.26 94.63 97.37 97.28 Si 6.808 6.758 6.858 6.811 6.422 6.907 7.777 7.575 Ti 0.174 0.175 0.177 0.178 0.037 0.037 0.001 0.016 Al 1.275 1.213 1.214 1.184 2.188 1.476 0.207 0.508 Mg 2.206 2.131 2.275 2.307 1.618 2.251 2.776 3.125 Ca 1.765 1.664 1.748 1.722 1.936 1.930 1.977 1.896 Mn 0.057 0.061 0.043 0.031 0.053 0.049 0.020 0.039 Fe 2.480 2.661 2.433 2.490 2.682 2.280 2.220 1.738 Na 0.249 0.293 0.351 0.359 0.354 0.276 0.037 0.097 K 0.155 0.160 0.152 0.152 0.107 0.042 0.014 0.015 Total 15.169 15.115 15.251 15.234 15.397 15.248 15.029 15.009 Al™ 1.192 1.213 1.142 1.184 1.578 1.093 0.207 0.425 Al* 0.083 0.000 0.072 0.000 0.610 0.383 0.000 0.083 Fe3+ 0.827 1.034 0.717 0.881 0.563 0.458 0.230 0.404 Fe2+ 1.653 L627 1.716 1.609 2.119 1.822 1.990 1.334 Ca(M3) 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 Fe(M4) 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 Ca(M4) 1.765 1.664 1.748 1.722 1.936 1.930 1.977 1.896 Na(M4) 0.235 0.293 0.252 0.278 0.064 0.070 0.023 0.097 Na(A) 0.014 0.000 0.099 0.081 0.290 0.206 0.014 0.000 K(A) 0.155 0.160 0.152 0.152 0.107 0.042 0.014 0.015 Mg* 0.57 0.57 0.57 0.59 0.43 0.55 0.58 0.70 Fe/(Fe+Mg) 0.53 0.56 0.52 0.52 0.62 0.50 0.44 0.36 MgO/FeO 0.50 0.45 0.52 0.52 0.34 0.55 0.70 1.01 7.80 8.05 5.52 7.77 8.76 10.06 MgOm/MgOr 7.68 7.44 6.19 6.45 6.79 5.84 7.28 5.81 FeOm/FeOr 6.41 6.90 2.47 2.52 0.50 0.52 0.02 0.26 Ti02m/Ti02r 2.45 2.47 278 Appendix 3.8 (Continued): Sample No. R15943 R15943 R15943 R15943 R15943 R15943 R15943 R15943 Rock Name Grd. Grd. Grd. Grd. Grd. Grd. Grd. Grd. RockSi02 71.32 71.32 71.32 71.32 71.32 71.32 71.32 71.32 Spot 2-R 2-C 3-R 3-C 4-R 4-C 5-R 5-C Si02 47.57 49.35 44.30 48.67 49.92 51.28 46.33 50.38 Ti02 0.26 0.48 0.69 1.06 0.72 0.13 1.50 0.12 4.84 A1203 7.77 5.27 9.69 4.86 4.09 4.28 6.68 MgO 11.37 12.58 9.01 11.78 12.51 13.61 10.66 12.73 CaO 12.30 12.09 11.62 11.56 11.54 12.48 11.26 12.33 MnO 0.32 0.29 0.37 0.28 0.32 0.27 0.36 0.24 FeO 15.99 15.92 20.04 18.01 17.41 14.97 19.38 16.25 0.45 Na20 0.83 0.57 1.35 0.61 0.55 0.44 0.81 0.25 K20 0.23 0.35 0.39 0.53 0.41 0.20 0.83 Total 96.64 96.90 97.46 97.36 97.47 97.66 97.81 97.59 Si 7.066 7.272 6.639 7.185 7.312 7.451 6.865 7.364 Ti 0.030 0.054 0.078 0.117 0.080 0.015 0.167 0.014 Al 1.360 0.915 1.712 0.847 0.706 0.732 1.166 0.833 Mg 2.518 2.761 2.012 2.591 2.732 2.949 2.354 2.774 Ca 1.958 1.908 1.865 1.829 1.811 1.942 1.787 1.931 Mn 0.041 0.037 0.047 0.036 0.039 0.034 0.046 0.029 Fe 1.986 1.961 2.512 2.224 2.132 1.819 2.402 1.986 Na 0.238 0.163 0.392 0.175 0.158 0.125 0.233 0.129 K 0.044 0.066 0.075 0.101 0.076 0.038 0.157 0.046 Total 15.241 15.137 15.332 15.105 15.046 15.105 15.177 15.106 AF 0.934 0.728 1.361 0.815 0.688 0.549 1.135 0.636 Al* 0.426 0.187 0.351 0.032 0.018 0.183 0.031 0.197 Fe3+ 0.248 0.388 0.657 0.615 0.652 0.289 0.770 0.374 1.612 Fe2+ 1.738 1.573 1.855 1.609 1.480 1.530 1.632 Ca(M3) 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 Fe(M4) 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 Ca(M4) 1.958 1.908 1.865 1.829 1.811 1.942 1.787 1.931 Na(M4) 0.042 0.092 0.135 0.171 0.158 0.058 0.213 0.069 Na(A) 0.196 0.071 0.257 0.004 0.000 0.067 0.020 0.060 0.157 0.046 K(A) 0.044 0.066 0.075 0.101 0.076 0.038 0.66 0.59 0.63 Mg* 0.59 0.64 0.52 0.62 0.65 0.38 0.51 0.42 Fe/(Fe+Mg) 0.44 0.42 0.56 0.46 0.44 MgO/FeO 0.71 0.79 0.45 0.65 0.72 0.91 0.55 0.78 9.45 7.40 8.84 MgO /MgO 7.90 8.74 6.26 8.18 8.69 ffl r 7.85 6.58 6.47 6.45 8.11 7.29 7.05 6.06 FeO» ^-^m Jiy/FeO a* ^ ^a-r j 1.28 1.96 1.33 0.24 2.78 0.22 Ti02m/Ti02r 0.48 0.89 279 to 3 < si OQ OQ + O .CO. © pp Ui>a©10©©S3©S3©Ui its bo b "si '© "ON b '© ui ui is i— ui VO ON ON4stOtOS3SJpa©tOslON UJ4S>-S3©NOslUiNOUislUi © v-n Ul O S3 © © S3 O SJ © Ui is bo ON tO © 00 © © sl is 'sl 'na tO NO 4s UltOSJ-fstOONUitOtOONON ON00tOtOUiUiNOSJ4sHaON © H Ul t-H © NJ © © SJ © JO © Ul 21 7 44 9 Ul sl Ul Ul '© '© S3 00 ia la 00 O NO o sl pa SJ O NO ON tO tO to Ui pa 4a. © 4s ON OS Ul © H- Ul H- O JO O © JO © S3 © Ui 4s bo Ul ON tO tO '© '© SJ 'sl S3 Ha sl NO sl s!Ui©N0t0Ha004^ONUltO S3S3S3paONOtOSJOOtOSJ NO H, t—' © H Ul © JO © © S3 © tO © Ul tO NO © NO © © © ui bo s] 00 b© "uUil © ©© Ul is, 4S ia Ul 00 Ul '© 00 la '© NO NO Ul © H- 4sNOS300S3©4sVOS300sl O 4s NO tO Ui na UiONsltO©S3t04sUi©Ui r-^O N JO NO S3 l to 9 © to UlHa©tO©©HapjO© Ui NO © Ui © © ON ON na 4s p Ul to W 00 NO tO 'sl sl bo '© is © © b Ul "4S i-J Ul NO S3 O 4s to © © 4s ' 4s SJ to S3slS3©4s©SJONONsltO 4s O OS to Os Ul rP SJVDOONOOOOUlUIHaHaNO ON P-> ON JO 1 NO S3 H tO 1 to 9 Ulpa©tOO©HapjO©Ul NO O to o o sl p H- 4s IJ p Ul o to 4s 00 NO w tO ia sl NO © H © © "sl b 4S t-H Ul 4a. b O 00 tO '© Ui ON Ui sl SJ4sONOOtOOONOUisl 4s o ON 00 4^4SHaUi4s©ONS3NOpO© ON ©Ui00NO4s©NO©S3Ui00 SJ ON NO S3 h_. tO rH ""»*9 2 1 Ulpa©S3©pa*©JO©Ul S3 © © sl p Ha Ul Ui © to 4s NO 00 U) ia b NO b NO © © b b 4s to ui 00 00 ia sl 'tO "© S3 00 SO sl o p ON 4a. haUitOsl4s©00NDp-tO00 Ha©S3tOS3©UiOOOOtO 4s uttOt*H-S30*OS30Nt04s ON SJ ON Q JO S3 t-H tO 8. 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According to Deer et al. (1992) Fe* = total Fe + Mn; En = lOOMg / (Ca + Fe* + Mg), Wo = lOOCa / (Ca + Fe* + Mg) and Fs = lOOFe* / (Ca + Fe* + Mg). Oxides, wr%. Sample No. R15929 R15929 R15929 R15929 R15929 R15929 RockSi02 45.75 45.75 45.75 45.75 45.75 45.75 Spot 1-R-l l-R-2 1-C 2-R-l 2-R-2 2-C Si02 50.57 51.93 52.83 53.15 52.58 51.63 Ti02 0.48 0.35 0.26 0.11 0.31 0.26 A1203 6.99 3.23 2.45 1.84 2.93 3.31 Cr203 0.22 0.26 0.26 0.28 0.26 0.25 MgO 18.21 16.56 15.92 15.98 16.06 16.35 CaO 14.46 21.13 23.85 24.18 22.10 22.48 MnO 0.01 0.16 0.08 0.14 0.12 0.13 FeO 5.33 4.88 4.39 3.66 4.34 4.19 NiO 0.08 0.06 0.13 0.00 0.09 0.01 Na20 1.09 0.42 0.15 0.23 0.38 0.28 0.00 K20 0.02 0.06 0.00 0.00 0.00 Total 97.46 99.04 100.32 99.57 99.17 98.89 Si 1.865 1.918 1.933 1.954 1.937 1.911 Ti 0.013 0.010 0.007 0.003 0.009 0.007 Al 0.304 0.140 0.106 0.080 0.127 0.144 Cr 0.006 0.008 0.007 0.008 0.008 0.007 Mg 1.001 0.912 0.868 0.876 0.881 0.902 Ca 0.571 0.836 0.935 0.953 0.872 0.891 Mn 0.000 0.005 0.002 0.004 0.004 0.004 Fe 0.164 0.151 0.134 0.112 0.134 0.130 Ni 0.002 0.002 0.004 0.000 0.003 0.000 Na 0.078 0.030 0.011 0.016 0.027 0.020 K 0.001 0.003 0.000 0.000 0.000 0.000 Total 4.005 4.015 4.007 4.006 4.002 4.016 Fe* 0.164 0.156 0.136 0.116 0.138 0.134 En 57.66 47.90 44.77 45.04 46.59 46.81 Wo 32.89 43.91 48.22 49.00 46.11 46.24 Fs 9.45 8.19 7.01 5.96 7.30 6.95 285 9£pT £*?c?p.-p^ Hw:z;'Tift>Ht/ NO Ui 4s N0©O©4Ssl©00 © sl S3 NO © SJ b b to "©; to © sl to SJO4stOO4s©00 OsOONOONOOstrtSJ NO SJ to <1 © 00 © NO © OOO© '© 'sl SJ 4s © tO O sl is © is 4s © ON OJONOSJnaOO©Sl slS3NOHa©S3©tO NO S3 OJ sl © 00 00 t-n o © © © b to '© b to b si ui b to sl 4S Os©ONS300sl©S3 ND4s©S3tOONQ4s NO S3 © H 00 O © O O tO sl © 00 © SJ sl © 4s ON '© b b b "ND "oo ©i b OOUiS3N04s©NO Ha©©S34SUl©© sl S3 SJ t-H O S3 O H Ul © © S3 Ui © 00 S3 00 o © NO b '© © © oo ON Ul NO sl S3 © 4s NO © 00 to Ul 1—> © 00 Os S3 S3i 00 Ui 4s S3 S3 © 1-* Ul ©! 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PTo PT*%o a cfi.> SfiZS"""- og z "go CO • 71. 5 R1 5 Dar i pa Grd . ui 4s SJ NO S3 to © H— Ul <_> O © ON O po ^p u« to © © NO 00 ON © is ui b O © S3 © sl sl © Ha ©S34sOJNONO©Ha 00 NO b tOUi©sl00©paSJ NJ S3 4s 4s ONUIUIUI4*.UI©N0 sl 4s004sOsONslSJHa 4s Ul SJ NO S3 Ul 292 p- 00 © © © pa © H- ON © NO NO © Ul © O 00 4s ft Ul Ul © "ON to © NO to S3 p— © NO © '© NO "© "NO b 'to ia *4a. bo 00 OS ON OO © Ui 00 sl Ui pa SJ H 4s OJ Ul NO S3 © 00 ND SJ S3 Ui s] OJ S3 4s s] NO 4s 00 na NO H- OS © 00 © NO © to © © 4^ © H 00 © SO © S3 NO © sl g^9 2 SJ NO 00 b b to b to b b NO na © OJ © © Ha © Y Ul f" ui ON 4s Ui to ON sl © ON tO pa pa H-i PP O NO P- O 00 na 00 ' VD 4*-©00-pSp-©©Ha SJ sl Ul Js SJ NO S) Ul SJ Ui © o £292 © to Ul o © NO © ON © NO CO © Ui is b "NO OOaOOOSOOO '© © © "© 4s © "o to 73 Ul P* Ui © pa NO paH-tO©OJ00©tO S»00 VD NOON00tO4ssltO00 S3slOOS34sS3S3sl sj SJ sl ui to to NO S3 ' H Ui O © ON © 92 t-H NO NO © © S3 © . <^i © 00 ON © 1-H ~J H-> na 00 b b 4s b ui ubi is ia 'na bO IO Ha © "© OS P- Ul O 00 S3 ONSOUl4sSJNO©sl " ND pa 4i. 0 OJ sl Ul © H-" NO 4s © ON © to © ON SJ sl S3 sl ft JO rH tO Ul 4s ui SJ NO 1 © Ul to © © p JO p © NO NO © © t-n 00 © to S- Ul ON tO NO SJ Ul NO © JO © n- Ui © © £292Ul © tO ON NO O sl © sl Ul © 00 Ul NO '© is b 4s is "© Ul 00 ND OJ S3 4s b b ia '© 4s 'NO b oo5* u, P- -OHaVOHaOJslOOJ SJ sl © to ©ONOONOSOtOSJtO Ui4s4s4sO04sSJNO sl 4s 4s NO pa fO NO © pa © Ha Ul O O © Ul © © sl O £29Ul 2Ul © NO NO ND so b "ND b NO oo b b Ul P-» ON tO na Ul © sl V3 iono pp NO 00 NO ia S3 Ui 00 O tO tO pa SJ NJ NO tO OO 00tONO4sONsl©00 SOONtOUltO©HaHa sl w pa SJ Ui jo a 4s Ul S3 NO 2 9 © O n- © JO p © NO Ui © a 00 © 4s © Ul to 00iaHaOJ©ON©Ul b b i-J bbbo'©Hana©00 slh-»NOtOONUlSJpa ft £ p- ^ | ©S3S3Ui4*.©OUi SJ sl H- S3 00 SOS UlUlHa©S3ONtO00 sl 287 Appendix 3.14 Mineral chemistry and structural formulae of hornblende (23 oxygens) in granodiorite from Kuh Mish intrusions (oxides, wt%). Mg* = Mg/(Mg + Fe2+). Locality Namin Namin Namin Namin Namin Namin Namin Namin Sample No. R15926 R15926 R15926 R15926 R15926 R15926 R15926 R15926 Rock Si02 63.93 63.93 63.93 63.93 63.93 63.93 63.93 63.93 Spot 1-R 1-C-l l-C-2 2-R 2-C-l 2-C-2 3-R-l 3-R-2 Si02 46.63 45.75 48.15 44.36 46.81 48.02 46.62 47.98 Ti02 0.69 0.80 1.16 0.36 1.39 1.40 0.42 1.06 A1203 6.52 7.28 5.69 11.20 6.10 5.80 8.00 5.45 MgO 12.05 11.50 14.64 9.02 13.62 15.00 10.76 13.51 CaO 11.67 11.56 11.32 11.28 11.06 11.14 11.45 10.41 MnO 0.67 0.69 0.44 0.57 0.38 0.37 0.54 0.81 FeO 15.93 16.59 12.86 16.54 14.13 12.74 16.82 14.80 Na20 0.75 0.79 0.90 1.73 1.00 0.98 0.81 1.03 K20 0.49 0.53 0.30 0.72 0.31 0.30 0.61 0.35 Total 95.40 95.49 95.46 95.78 94.80 95.75 96.03 95.40 Si 6.994 6.872 7.059 6.738 6.956 6.996 6.972 7.056 Ti 0.078 0.091 0.129 0.041 0.155 0.153 0.047 0.117 Al 1.152 1.289 0.982 2.005 1.069 0.995 1.410 0.945 Mg 2.693 2.576 3.198 2.043 3.016 3.258 2.399 2.960 Ca 1.874 1.861 1.777 1.836 1.761 1.740 1.834 1.640 Mn 0.086 0.089 0.055 0.073 0.047 0.045 0.068 0.102 Fe 1.997 2.084 1.577 2.101 1.756 1.552 2.103 1.820 Na 0.219 0.231 0.256 0.509 0.290 0.277 0.235 0.293 K 0.094 0.100 0.055 0.140 0.059 0.056 0.118 0.067 Total 15.187 15.193 15.088 15.486 15.109 15.072 15.186 15.000 Al17 1.006 1.128 0.941 1.262 1.044 0.995 1.028 0.944 Af1 0.146 0.161 0.041 0.743 0.025 0.000 0.382 0.001 Fe3" 0.643 0.730 0.777 0.114 0.844 0.896 0.533 1.069 Fe2+ 1.354 1.354 0.800 1.987 0.912 0.656 1.570 0.684 Ca(M3) 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 Fe(M4) 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.067 Ca(M4) 1.874 1.861 1.777 1.836 1.761 1.740 1.834 1.640 Na(M4) 0.126 0.139 0.223 0.164 0.239 0.260 0.166 0.293 Na(A) 0.093 0.092 0.033 0.345 0.051 0.017 0.069 0.000 K(A) 0.094 0.100 0.055 0.140 0.059 0.056 0.118 0.067 Mg* 0.67 0.66 0.80 0.51 0.77 0.83 0.60 0.80 0.47 0.38 Fe/(Fe+Mg) 0.43 0.45 0.33 0.51 0.37 0.32 MgO/FeO 0.76 0.69 1.14 0.55 0.96 1.18 0.64 0.91 3.76 5.68 6.25 4.48 5.63 MgOm/MgOr 5.02 4.79 6.10 2.27 2.92 2.50 2.25 2.97 2.61 FeOm/FeOr 2.81 2.93 2.27 0.71 2.73 2.75 0.82 2.08 Ti02m/Ti02r 1.32 1.57 288 Appendix 3.14 (Continue d): Locality Namin Narnin Darin Darin Darin Darin Darin Darin Sample No. R15926 R15926 R15927 R15927 R15927 R15927 R15927 R15927 Rock Si0 63.93 2 63.93 71.58 71.58 71.58 71.58 71.58 71.58 Spot 3-C 4-C 1-R 1-C 2-R 2-C 3-R 3-C 47.65 Si02 48.96 46.92 47.17 48.09 45.97 47.48 47.32 Ti02 1.43 0.84 1.33 1.36 0.97 1.29 1.10 1.33 A1203 5.87 4.71 6.00 6.36 4.78 5.82 5.27 6.03 MgO 12.76 13.71 11.23 12.28 11.50 11.88 11.12 12.00 CaO 11.00 11.43 10.18 10.55 9.94 10.45 9.91 10.35 MnO 0.66 0.65 0.60 0.46 0.70 0.57 0.61 0.53 FeO 15.42 14.28 18.75 16.92 18.38 17.67 19.01 17.53 Na20 0.98 0.55 1.77 1.79 1.46 1.79 1.64 1.70 K20 0.34 0.27 0.33 0.30 0.45 0.31 0.41 0.34 Total 96.11 95.40 97.11 97.19 96.27 95.75 96.55 97.13 Si 7.030 7.227 6.937 6.926 7.140 6.889 7.054 6.958 Ti 0.158 0.093 0.148 0.151 0.109 0.145 0.122 0.147 Al 1.021 0.819 1.045 1.100 0.835 1.029 0.923 1.043 Mg 2.806 3.017 2.476 2.688 2.546 2.654 2.463 2.630 Ca 1.738 1.807 1.613 1.660 1.581 1.678 1.577 1.631 Mn 0.082 0.082 0.075 0.057 0.088 0.071 0.076 0.067 Fe 1.902 1.763 2.319 2.078 2.282 2.213 2.361 2.156 Na 0.281 0.158 0.506 0.509 0.420 0.520 0.472 0.484 K 0.065 0.051 0.063 0.056 0.086 0.060 0.077 0.064 Total 15.083 15.017 15.182 15.225 15.087 15.259 15.125 15.180 17 Al 0.970 0.773 1.045 1.074 0.835 1.029 0.923 1.042 71 Al 0.051 0.046 0.000 0.026 0.000 0.000 0.000 0.001 3+ Fe 0.783 0.716 0.990 0.861 0.999 0.965 1.024 0.935 2+ Fe 1.119 1.012 1.329 1.217 1.283 1.248 1.337 1.221 Ca(M3) 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.000 Fe(M4) 0.000 0.035 0.000 0.000 0.000 0.000 0.000 0.000 Ca(M4) 1.738 1.807 1.613 1.660 1.581 1.678 1.577 1.631 Na(M4) 0.262 0.158 0.387 0.340 0.419 0.322 0.423 0.369 Na(A) 0.019 0.000 0.119 0.169 0.001 0.198 0.049 0.115 K(A) 0.065 0.051 0.063 0.056 0.086 0.060 0.077 0.064 Mg* 0.71 0.74 0.65 0.69 0.66 0.68 0.65 0.68 Fe/(Fe+Mg) 0.40 0.37 0.48 0.44 0.47 0.45 0.49 0.45 MgO/FeO 0.83 0.96 0.60 0.73 0.63 0.67 0.58 0.68 MgOm/MgOr 5.32 5.71 9.85 10.77 10.09 10.42 9.75 10.53 FeOffl/FeOr 2.72 2.52 5.12 4.62 5.02 4.83 5.19 4.79 Ti02ffl/Ti02r 2.80 1.65 3.69 3.78 2.69 3.58 3.06 3.69 1 QQ ^ ^ Z £ ^ ^ ^ ^ JO oo t- 5 O O "» w to A n /"~N O & CO a6 bp o 50 0 o g o al co* g TO no o 8 f> + O 02; ' a 41 N> CJ o, Cr- • J-' P- th NO Ha Ha OJ JO © H- Ul na O tO © © tO © IO © Ul r w p, S3 00 © NO O © S3 Ui na .ja. 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NO o © SJ ONO©ONOstO©©ND4sON o g oo vo 5 ONO©tOONOONO©S3UlS3 TO ^ 290 APPENDIX 4 WHOLE ROCK GEOCHEMICAL DATA ABBREVIATIONS USED Grd. = Granodiorite A.F.G. = Alkali feldspar granite Qtz = Quartz ASI = Aluminum Saturation Index Trace elements with decimal places were obtained by instrumental neutron activation analyses (INAA) while the others were obtained by XRF. Appendix 4.1 Whole rock geochemical data from the Kashmar granitoid (oxides, wt% and traces, ppm). Sample No. R15911 R15924 R15907 R15912 R15908 R15959 R15902 R15925 R15904 R15915 R15901 R15917 Rock name Tonalite Tonalite Tonalite Tonalite Grd. Grd. Grd. Grd. Grd. Grd. Grd. Grd. Si02 54.18 54.35 59.01 59.79 62.30 63.36 64.00 64.28 66.35 66.41 66.76 67.47 Ti02 0.73 0.80 0.81 0.91 0.63 0.60 0.55 0.58 0.49 0.46 0.51 0.43 Al203 17.04 16.80 16.04 16.50 16.02 16.07 15.71 16.20 15.29 15.22 15.32 15.23 Fez03 7.69 9.50 7.63 6.76 5.33 4.98 4.28 4.52 3.87 3.90 3.93 3.75 MnO 0.11 0.10 0.1 0.14 0.09 0.07 0.07 0.06 0.06 0.06 0.03 0.05 MgO 2.97 4.27 2.76 2.11 1.93 1.81 1.44 1.50 1.32 1.41 1.80 1.28 CaO 7.58 4.18 4.26 5.13 4.51 3.16 3.98 3.50 3.06 3.23 2.01 3.04 NazO 3.48 4.81 4.15 4.33 3.81 4.06 3.84 4.62 3.96 3.89 5.33 3.82 K20 1.36 2.04 2.87 2.26 2.75 3.72 3.24 2.87 3.50 2.83 2.27 3.19 P205 0.17 0.20 0.21 0.33 0.18 0.19 0.16 0.18 0.14 0.13 0.15 0.12 LOI 3.46 2.60 1.62 0.93 1.03 2.22 1.46 1.19 0.97 1.62 1.15 1.48 Rest 0.15 0.15 0.19 0.17 0.17 0.19 0.19 0.18 0.19 0.16 0.13 0.16 Total 98.92 99.80 99.65 99.36 98.75 100.43 98.92 99.68 99.2 99.32 99.39 100.02 Ba 210 295 505 455 515 660 630 635 690 595 365 530 Rb 59 72 55 45 59 78 72 58 88 62 56 88 Sr 366 315 363 367 342 341 324 343 310 273 256 282 Pb 6 4 8 4 6 6 6 4 4 8 4 10 Th 4 3 11 5 5.30 10 14 10 12 10.80 10 10 U 1 2 2 1 0.90 2 3 3 3 3.10 3 2 Zr 98 84 144 156 198 216 222 216 210 180 192 172 Nb 6 4 8 10 8 12 12 10 10 8 10 8 Y 22 20 33 31 22 27 31 27 26 22 21 19 La 22 14 26 24 17.40 24 32 28 28 21.50 24 24 Ce 40 25 60 55 36.50 55 65 55 50 42.50 50 40 Nd - . - - 17.40 - - - - 17.60 - - Sm - - - - 4.00 - - - - 3.90 - - Eu . . . . 1.11 - - - - 0.87 - - Gd . . - - 3.60 - - - - 3.60 - - Tb . . - - 0.54 - - - - 0.62 - - Ho . . . - 0.75 - - - - 0.95 - - Yb . . - - 2.10 - - - - 2.40 - . - Lu . . • - 0.34 - - - - 0.38 - - Sc . . • - 13.00 - - - - 9.60 - - V 188 220 190 144 96 82 66 70 56 62 58 58 Cr 18 8 2 4 16.5j 6 6 6 6 10.2 6 8 Mn 925 815 825 1070 740j 510 560 470 535 480 210 375 Ni 2 6 <2 <2 4 <2 <2 <2 <2 <2 <2 2 Cu 60 <2 50 4 2 8 2 16 4 2 <2 2 Zn 50 50 34 56 38 40 32 24 26 34 18 30 Sn 5 5 <5 5 <5 <5 <5 <5 <5 <5 L_ <5 <5 Ga 18 19 18 18 18 17 17 17 16 15j 16 15 As 5 6 4 2 2 7 3 5 6 5j 2 6 Sb „ . . 0.8 - - - - 0.75 .- - Cs m _ . . 0.9 - - - - 2.6 - - Hf „ _ . - 4.8 - - - - 4.6 - - Total REE m _ . - 83.74 - - - - 94.32 • Eu/Eu* m _ . - 0.88 - - - 0.70 - - - -| 6.05 • • LaN/YbN . . • - 5.60 - 65.4 87.19 76.29 76.29 58.58 65.40 65.40 LaN 59.95 38.15 70.84 65.40 47.41 Rb/Sr 0.16 0.23 0.15 0.12 0.17 0.23 0.22 0.17 0.28 0.23 0.22 0.31 0.09 0.13 0.10 0.15 0.17 Rb/Ba 0.28 0.24 0.11 0.10 0.11 0.12 0.11 0.25 0.84 0.62 0.88 0.73 0.43 0.84 KzO/Na20 0.39 0.42 0.69 0.52 0.72 0.92 0.95 0.96 0.99 1.03 0.99 ASI 0.81 0.95 0.91 0.87 0.92 0.98 Appendix 4.1 (Continued): Sample No. R15910 R15903 R15918 R15957 R15958 R15906 R15923 R15921 R15922 R15905 R15913 R15909 Rock name Granite Granite Dranite Granite Granite Granite Granite Granite Granite Granite Granite Granite Si02 63.42 64.63 65.33 66.34 66.44 66.67 66.97 66.99 67.49 70.24 71.69 71.81 Ti02 0.59 0.56 0.49 0.46 0.45 0.54 0.43 0.45 0.41 0.34 0.27 0.26 Al 0 15.37 14.82 2 3 15.53 15.31 15.24 15.89 15.48 15.16 15.09 13.9 13.61 13.77 Fe 0 5.39 4.94 2 3 4.47 3.76 3.86 2.88 3.69 3.76 3.52 2.77 2.29 2.22 MnO 0.10 0.08 0.08 0.10 0.08 0.02 0.09 0.08 0.07 0.06 0.02 0.03 MgO 2.36 2.15 1.65 1.40 1.39 1.43 1.44 1.28 1.11 0.93 0.77 0.70 CaO 4.71 4.31 3.37 2.87 3.39 1.33 2.90 2.74 2.92 2.35 2.05 2.04 Na O a 3.34 3.15 3.70 3.95 3.83 4.68 3.93 3.66 3.56 3.25 3.06 3.10 K20 3.39 3.66 3.23 3.78 3.21 3.39 3.31 3.91 3.82 4.47 4.58 4.62 P205 0.17 0.14 0.14 0.12 0.12 0.12 0.13 0.13 0.10 0.08 0.06 0.06 LOI 1.37 0.94 1.61 1.66 1.77 2.26 1.44 1.47 0.92 0.83 1.13 2.04 Rest 0.17 0.17 0.17 0.18 0.16 0.18 0.17 0.16 0.18 0.16 0.14 0.15 Total 100.38 99.55 99.77 99.93 99.94 99.39 99.98 99.79 99.19 99.38 99.67 100.80 Ba 440 475 515 590 530 745 540 500 585 590 545 580 Rb 103 111 88 104 80 61 81 115 105 144 129 145 Sr 315 283 288 292 269 214 295 281 253 206 191 188 Pb 14 14 8 10 12 2 10 10 14 16 6 10 Th ' 10.70 12 10.10 10 10.9 12 10 11 13 17 20 17.6 U 1.80 2 2.30 2 1.5 5 2 2 2 4 4 4.6 Zr 180 176 174 206 170 214 204 192 250 178 140 148 Nb 8 8 8 10 8 12 8 10 10 10 10 8 Y 23 21 22 20 22 31 20 22 24 20 22 16 La 22.00 24 22.00 26 25.00 30 24 22 22 24 18 24.50 Ce 45.50 50 44.00 45 50.00 65 45 50 45 50 25 45.50 Nd 20.00 - 18.20 - 20.50 ------15.80 Sm 4.20 __-} 4.00 - 3.40 - - • - - . 3.30 Eu 0.93 - 0.87 - 0.94 - - • - • - - 0.57 Gd 3.80 - 3.70 - 3.00 - - . - - - 3.30 Tb 0.65 - 0.65 - 0.50 ------0.55 Ho 0.95 - 0.90 - 0.65 ------0.80 Yb 2.15 - 2.30 - 2.35 ------1.90 Lu 0.35 - 0.37 - 0.39 ------0.29 Sc 14.30 - 10.90 - 9.70 - - p - - - 4.70 V 108 102 78 62 58 64 62 56 52 44 32 30 Cr 20 22 12.40 6 9.00 6 10 6 10 6 6 5.10 Mn 775 658 635 805 665 . 175 725 645 615 470 200 280 Ni 4 6 2 <2 <2 <2 2 <2 <2 <2 <2 <2 Cu 28 30 20 8 12 4 2 4 6 16 2 4 Zn 62 50 32 48 46 16 62 38 46 38 12 22 Sn <5 <5 <5 <5 <5 <5 <5 <5 <5 <5 <5 <5 Ga 16 15 17 15 16 16 16 16 16 13 13 13 As 5 4 2 14 4 2 4 8 2 3 3 2 Sb 0.9 - 0.7 - 0.6 ------0.4 Cs 5.0 > 2.7 - 1.2 ------5.0 Hf 4.9 - 4.6 - 9.4 ------4.0 Total REE 80.55 - 96.99 - 106.73 ------96.51 Eu/Eu* 0.69 - 0.68 - 0.88 - - - - - 0.52 ------8.71 LaN/YbN 6.91 - 6.46 - 7.19 59.95 65.40 49.05 66.76 LaN 59.95 65.40 59.95 70.84 68.12 81.74 65.40 59.95 Rb/Sr 0.33 0.39 0.31 0.36 0.30 0.29 0.27 0.41 0.42 0.7 0.68 0.77 Rb/Ba 0.23 0.23 0.17 0.18 0.15 0.08 0.15 0.23 0.18 0.24 0.24 0.25 0.84 1.07 1.07 1.38 1.50 1.49 K20/Na20 1.01 1.16 0.87 0.96 0.84 0.72 ASI 0.87 0.87 0.99 0.97 0.96 1.15 1.01 1.00 0.98 0.96 0.99 0.99 Appendix 4.1 (Continued): Sample No. R15920 R15916 R15900 R15914 R15919 Rock name A.F.G. A.F.G. A.F.G. A.F.G. A.F.G. Si02 74.63 75.43 76.75 76.97 77.06 Ti02 0.17 0.24 0.19 0.15 0.17 Al203 12.73 12.31 11.79 11.70 12.52 Fe203 1.22 0.94 0.97 0.74 0.61 MnO 0.02 0.O1 0.01 0.01 0.01 MgO 0.17 0.15 0.30 0.14 0.09 CaO 0.45 0.87 0.51 0.43 0.48 Na20 3.75 2.61 2.85 2.63 5.26 K20 4.93 5.88 5.61 5.59 2.60 P2O5 0.01 0.02 0.02 0.01 0.03 LOI 1.14 1.64 0.74 1.76 1.70 Rest 0.13 0.13 0.09 0.08 0.09 Total 99.35 100.23 99.83 100.21 100.62 Ba 555 490 170 140 245 Rb 120 172 200 207 65 Sr 51 72 50 . 36 57 Pb 12 8 8 12 6 Th 16 19 30 31 17 U 2 3 5.6 3.4 2 Zr 164 168 134 112 172 Nb 14 10 14 10 14 Y 26 20 31 20 32 La 30 26 31.00 32.00 24 Ce 75 45 68.00 64.00 60 Nd - - 27.50 23.50 - Sm - - 5.00 4.30 - Eu - - 0.30 0.27 - Gd - - 5.10 3.60 - Tb - - 0.95 0.69 Ho - - 1.30 0.85 - Yb - - 3.60 2.50 - Lu - - 0.53 0.35 - Sc - - 4.30 3.30 - V 4 12 8 4 6 Cr <2 <2 2 1.2 <2 Mn 200 110 85 85 60 Ni <2 <2 <2 <2 <2 Cu <2 <2 <2 <2 <2 Zn 14 6 8 4 4 Sn <5 <5 <5 <5 <5 Ga 12 11 11 11 12 As 1 2 4 4 2 Sb - - 0,45 0.50 - Cs - - 3.30 4.40 - Hf - - 4.90 . 4.00 - Total REE - - 143.28 132.06 - Eu/Eu* - - 0.18 0.20 - - Laf/YbN - - 5.82 8.65 70.84 84.47 87.19 65.40 LaN 81.74 Rb/Sr 2.35 2.39 4.00 5.75 1.14 Rb/Ba 0.22 0.35 1.18 1.48 0.27 2.25 1.97 2.13 0.49 K20/Na20 1.31 ASI 1.03 1.01 1.01 1.05 1.01 Appendix 4.2 Whole rock geochemical data from the fornavard granitoid (oxides, wt% and traces, ppm). Sample No. R15944 R15945 R15946 R15947 R15953 R15943 Rock name Tonalite Tonalite Granodiorite Granodiorite Granodiorite Granodiorite SiOz 48.64 58.09 63.18 68.85 69.45 71.32 Ti02 0.93 0.97 0.82 0.54 0.62 0.54 Al203 15.43 13.35 15.01 13.58 14.75 13.18 Fe203 10.28 8.25 6.60 4.87 3.05 2.47 MnO 0.16 0.14 0.09 0.04 0.05 0.06 MgO 7.92 4.40 1.52 0.66 1.27 1.44 CaO 10.72 7-36 3.56 2.49 3.74 2.84 Na20 2.59 3.59 3.73 4.03 5.30 6.18 K20 0.27 0.60 2.44 2.46 0.53 0.25 P2O5 0.10 0.13 0.25 0.11 0.15 .0.09 LOI 2.33 2.05 1.60 1.53 1.00 1.05 Rest 0.16 0.15 0.18 0.20 0.12 0.08 Total 99.53 99.08 98.98 99.36 100.03 99.50 Ba 70 85 535 810 150 55 Rb 6 12 129 70 23 8 Sr 167 149 145 109 168 121 Pb 12 36 6 10 8 12 Th 2 6.90 10.80 6.10 11.70 15.60 U <1 1.00 1.40 1.30 2.60 1.90 Zr 48 208 272 448 312 142 Nb <2 6 10 8 12 8 Y 19 48 39 38 69 26 La 8 19.50 29.00 24.50 16.40 18.00 Ce 10 45.00 61.00 53.00 44.00 47.00 Nd - 27.00 30.00 27.50 29.00 27.00 Sm • 5.00 5.60 6.10 7.70 6.00 Eu • 1.62 1.83 2.19 2.20 1.19 Gd - 6.00 5.90 5.90 8.90 5.50 Tb - 1.03 0.96 0.92 1.49 0.84 Ho - 1.40 1.25 1.20 2.20 L 1.00 Yb - 3.75 3.80 3.80 7.40 2.50 Lu - 0.74 0.63 0.65 1.16 0.38 Sc . 34.00 16.40 17.00 17.00 16.20 V 240 148 88 26 48 50 Cr 288 124 16 6 11 28 Mn 1300 1100 720 340 415 445 Ni 80 36 6 <2 6 18 Cu 70 22 10 2 <2 12 Zn 106 122 68 38 42 40 Sn <5 5 <5 <5 5 <5 Ga 15 17 20 18 17 12 As <1 2 3 2 4 13 Sb • 0.15 0.45 0.30 0.60 0.20 Cs _ 1.50 3.60 2.00 1.20 0.10 9.40 4.90 Hf m 6.40 6.40 8.10 125.76 120.45 109.41 Total REE - -111.04 139.97 1.10 0.81 0.62 Eu/Eu* — 0.90 0.97 4.36 1.50 4.87 a 5.16 LaN/YbN 3.51 44.69 49.05 Law 21.80 53.13 79.02 66.76 0.14 0.07 Rb/Sr 0.04 0.08 0.89 0.64 0.15 0.15 Rb/Ba 0.09 0.14 0.24 0.09 0.61 0.1C 0.04 K20/Na20 0.10 0.17 0.65 ASl 0.64 0.67 0.98 0.98 0.92 0.841 Sample No. R15936 FJ1595E R1594C R1595-1• • HI 5942 R15939 R15941 1 Rock name Granite Granite Granite! Granitei liranite Granite Granite Si02 74.84 75.24 75.4C 75.55» 75.5S> 75.78 76.04 Ti02 0-19 0.18 0.19 0.2C' 0.1 I£ 0.16I 0.19 Al203 12.04 11.98 12.25 12.13 12.3E» i2.iei 12.29 Fe203 2.28 2.61 2.10 2.24 1.37 1.84 1.99 MnO 0.02 0.03 0.02 0.03 0.01 O.C 0.03 MgO 0.14 0.3 0.15 0.23 0.25 0.13 0.17 CaO 0.72 , 0.46 0.87 0.77 0.81 0.26 0.73 NazO 3.47 3.64 3.62 3.64 3.88 3.87 3.75 K20 3.99 3.65 3.93 3.98 3.69 4.35 3.96 P205 0.02 0.03 0.03 0.04 0.03 0.03 0.03 LOI 1.24 1.41 2.19 0.87 1.07 0.89 0.81 Rest 0.16 0.15 0.16 0.17 0.14 0.17 0.16 Total 99.11 99.68 100.91 99.85 99.38 99.66 100.15 Ba 680 660 735 815 580 890 645 Rb 95 59 97 100 54 79 128 Sr ^ 39 48 54 45 48 38 39 Pb 6 6 6 4 6 2 8 Th 19.20 ; 20.00 18.60 19.00 20.00 19.00 21.00 U 2.10 ' 4.00 2.50 4.00 3.00 4.00 3.00 Zr 230 226 234 240 226 226 234 Nb 12 12 10 14 12 12 12 Y 57 1 56 59 53 48 52 59 La 46.00 42 31.50 38 38 34 42.00 Ce 96.00 ' 95 76.00 90 90 75 93.00 Nd 44.50 - 41.00 - - - 44.00 1 " Sm 8.00 1 8.10 - 9.30 Eu 1.25 - 1.20 - - - 1.15 Gd 9.00 ; 9.00 - - - 9.40 Tb 1.38 1.43 - - , 1.48 Ho 1.65 - 1.80 - - - 1.85 Yb 5.40 - 6.30 - - - 6.00 Lu 0.87 0.98 - - - 0.98 Sc 11.80 1 11.50 - .- - 11.80 V 2 i <2 <2 4 6 <2 4 Cr 4 ' <2 3.3 <2 <2 <2 2.9 Mn 130 i 305 165 200 145 40 220 Ni <2 <2 <2 <2 <2 <2 <2 Cu <2 i <2 <2 2 2 <2 <2 Zn 18 | 22 18 20 16 6 26 Sn <5 <5 <5 <5 <5 <5 <5 Ga 17 ! 16 17 16 16 16 17 As 2 1 2 2 3 2 2 Sb 0.20 i 0.35 - - - 0.40 Cs 0.80 I 1.10 - - - 2.00 Hf 7.20 i 7.60 - - - 7.80 Total REE 214.05 176.81 - - - 209.16 Eu/Eu* 0.45 0.43 - - - 0.37 LaN/YbN 5.76 3.38 - - - 4.73 LaN 125.34 1114.44 85.83 103.54 103.54 92.64 114.44 Rb/Sr 2.44 ! 1.23 1.80 2.22" 1.13 2.08 3.28 Rb/Ba 0.14 ' 0.09 0.13 0.12 •0.09 0.09 0.20 1.07 K20/Na20 1.15 1.00 1.09 1.09 0.95 1.12 ASI 1.06 ; 1.11 1.04 1.04 1.04 1.05 1.04 w 296 ^^^^~ Appendix 4.3 Whole rock geochemical data from the Taknar Rhyolite (oxides, wt% and traces, ppm). Sample No. R15952 R15950 R15951 R15948 R15949 Rock name Rhyolite Rhyolite Rhyolite Rhyolite Rhyolite SiOz 75.75 75.96 76.07 76.39 77.90 Ti02 0.13 0.16 0.14 0.14 0.14 Al203 12.10 12.64 12.37 12.03 10.91 Fe203 2.10 1.57 1.58 1.71 4.07 MnO 0.00 0.01 0.01 0.05 0.06 MgO 0.22 0.45 0.25 0.45 0.64 CaO 0.13 0.23 0.18 0.35 0.13 Na20 2.84 3.81 4.88 3.60 0.37 K20 5.35 3.78 3.31 3.83 3.32 P205 0.03 0.03 0.03 0.04 0.06 LOI 0.91 1.33 0.63 1.03 2.27 Rest 0.14 0.23 0.12 0.17 0.15 Total 99.70 100.20 99.57 99.79 100.02 Ba 795 1020 545 945 705 Rb 119 91 76 130 120 Sr 38 93 45 63 12 Pb 6 12 6.00 12 6 Th 20 23 19.80 17 19 U 3 3 2.6 4 4 Zr 124 144 138 132 130 Nb 8 10 8 8 10 Y 37 34 49 41 44 La 12 40 28.50 18 36 Ce 25 80 63.00 45 75 Nd - - 28.00 - - Sm - - 6.10 - - Eu - - 0.54 - - Gd - - 6.90 - - Tb - - 1.16 - - Ho - - 1.45 - - Yb - - 4.80 - - Lu - - 0.73 - - Sc - - 4.00 - - V 4 6 4 6 2 Cr <2 <2 3 <2 <2 Mn 20 100 60 375 495 Ni <2 <2 <2 <2 <2 Cu <2 2 <2 <2 34 Zn 6 404 8 68 64 Sn <5 10 <5 <5 5 Ga 14 16 15 16 14 As 1 5 2 <1 2 Sb - - 0.50 - - Cs - - 0.50 - - Hf - - 4.90 - - Total REE - - 84.48 - - 0.25 - Eu/Eu* - - M Laj«(/YbN - - 4.01 - " 49.05 98.09 LaN 32.70 108.99 77.66 Rb/Sr 3.13 0.98 1.69 2.06 10.00 Rb/Ba 0.15 0.09 0.14 0.14 0.17 1.06 8.97 K20/Na20 1.88 0.99 0.68 ASI 1.13 1.17 1.04 1.12 2.46| Appendix 4.4 Whole rock geochemical data from the Kuh Mish intrusions,(oxides, wt% and traces, ppm). Sample Nq. R15929 . R15932 R15930 R15956 R15934 Rock name Gabbro Qtz monzodiorite Qtz monzodiorite Qtz monzodiorite Qtz monzodiorite Si02 45.75 51.85 53.18 54.37 60.71 Ti02 0.13 1.08 0.56 0.31 0.41 Al203 17.23 14.99 15.63 15.96 14.57 Fe203 4.19 11.77 9.59 8.92 7.12 MnO 0.07 0.17 0.17 0.14 0.13 MgO 11.81 4.30 5.99 . 6.16 3.77 CaO 16.60 8.01 6.84 8.43 6.95 Na20 0.68 3.15 3.93 2.80 2.99 K20 0.07 0.62 0.62 0.57 0.80 P205 0.01 0.18 0.09 0.03 0.09 LOI 2.83 3.19 3.07 2.18 1.54 Rest 0.20 0.14 0.12 0.12 0.09 Total 99.57 99.45 99.79 99.99 99.17 Ba 3 125 85 165 135 Rb 1 10 8 9 14 Sr 108 191 228 166 133 Pb <2 <2 <2 <2 <2 Th <0.1 <1 1 <1 1 U <0.1 <1 <1 <1 <1 Zr 2 62 40 20 40 Nb <2 2 <2 <2 <2 Y 3 22 14 6 15 La i 0.10 6 8 4 4 Ce 0.50 15 10 <5 5 Nd i 0.96 - • - - Sm 0.50 - - - - Eu 0.14 - . - - - Gd 0.75 - - - - Tb i 0.18 - - - - Ho 0.30 - - - - Yb 0.43 - - - - Lu 0.07 - - - - Sc * 41.00 - - - - V 104 316 272 280 182 Cr 805 6 62 22 14 Mn 665 1480 1490 1270 1170 Ni 242 14 26 26 14 Cu 122 100 6 72 16 Zn 22 90 56 64 64 Sn <5 <5 <5 <5 <5 Ga 9 16 13 13 13 As 1 <1 <1 <1 1 Sb ! 1.60 - - - - Cs S <0.1 - - - - Hf ! 0.10 - - - - Total REE 3.93 - - - - Eu/Eu* 0.70 ------LaN/YbN 0.16 - 21.80 10.90 10.90 LaN j 0.27 16.35 Rb/Sr 0.00 0.05 0.04 0.05 0.11 Rb/Ba 0.33 0.08 0.09 0.05 0.10 0.20 0.27 K20/Na20 0.10 0.20 0.16 0.78 0.79 .. .-. _, , 0.73 0.80 298 Appendix 4.4 (Continued): Sample No. R15926 R15936 R15927 R15937 R15933 R15928 R15935 R15931 Rock name Grd. Grd. Grd. Grd. Grd. Grd. Grd. Grd. Si02 63.93 70.66 71.58 72.59 73.30 73.34 73.58 75.96 Ti02 0.51 0.30 0.36 0.24 0.24; 0.22 0.25 0.15 A1203 15.00 13.92 13.59 13.03 13.20 13.12 13.58 12.52 Fe203 5.66 3.37 3.66 2.99 2.32 2.36 2.16 1.33 MnO 0.11 0.07 0.04 0.04 0.04 0.04 0.04 0.02 MgO 2.40 1.02 1.14 0.66 0.99 0.94 1.08 0.37 CaO 5.08 3.78 5.10 2.85 3.01i 2.77 1.26 1.58 Na20 3.31 4.02 3.08 3.87 3.73 3.70 5.63 4.05 K20 1,66 0.48 0.14 0.95 1.62 1.54 0.43 2.48 P2o5 0.10 0.07 0.07 0.06 0.05 0.04 0.04 0.03 LOI 1.50 1.94 1.23 1.58 1.29 1.76 1.56 0.79 Rest 0.11 0.06 0.07 0.07 0.07 0.07 0.03 0.07 Total 99.37 99.69 100.06 98.93 99.86 99.90 99.64 99.35 Ba 215 105 35 210 220 210 35 270 Rb 29 5 1 13 32 27 5 38 Sr 242 160 249 174 97 95 89 60 Pb 2 2 <2 <2 <2 <2 <2 4 Th 2.60 2 1.10 2 2 2 1 2 U 0.90 <1 0.50 <1 <1 <1 <1 <1 Zr 86 56 110 52 60 62 56 94 Nb <2 <2 <2 <2 <2 <2 <2 <2 Y 19 18 24 19 13 12 18 15 La 6.60 10 5.80 6 8 8 6 12 Ce 15.10 5 14.00 10 10 10 10 10 Nd 9.00 - 8.60 - - - - - Sm 2.30 - 2.10 - - - - - Eu 0.70 - 0.67 - - - - - Gd 2.90 - 2.40 - - - - - Tb 0.50 - 0.44 - - - - - Ho 0.65 - . 0.70 - - • - - - Yb 2.05 - 2.55 - - - - - Lu 0.35 • - 0.44 - - - - - Sc 18.90 - 11.70 - - - - - V 132 50 54 20 40 38 16 12 Cr <1 4 2.8 2 12 12 <2 4 Mn 940 575 320 360 360 375 355 195 Ni 4 <2 <2 <2 4 2 <2 <2 Cu 30 2 <2 6 4 16 <2 4 Zn 42 34 12 24 22 22 16 14 Sn <5 <5 <5 <5 <5 <5 <5 <5 Ga 14 12 12 12 11 11 10 10 As <1 <1 2 <1 <1 3 <1 2 Sb 1.15 - 0.65 - - - - - Cs 0.5 - 0.40 - - - - Hf 2.9 - 3.20 - - - - . Total REE 40.15 37.7 - - - - , Eu/Eu* 0.83 - 0.91 - - - • *• - - - - - • LaN/YbN 2.18 1.54 21.80 21.80 16.35 32.70 LaN J, 17.98 27.25 15.80 16.35 Rb/Sr 0.12 0.03 0.00 0.07 0.33 0.28 0.06 0.63 Rb/Ba 0.13 0.05 0.03 0.06 0.15 0.13 0.14 0.14 0.42 0.08 0.61 KzO/NazO 0.50 0.12 0.05 0.25 0.43 A AJ nn9 0.94 1.04 0.99 1.03 1.13 1.25 L^y Cat. No.| Field No. ANU No. Description Locality Age Longtitude Latitude R15900 1-GAR-2 PCW112 A.F.G. Kashmar 43.5 Ma 58° 20' 55" E 35° 19'48" N R15901 1-GAR-3 PCW113 Granodiorite Kashmar 42.8 Ma 58° 21'00" E 35° 20' 27" N R15902 1-KES-7 PCW114 Granodiorite Kashmar 42.8 Ma 58° 22' 38" E 35° 20" 42" N R15903 1-KES-6 PCW115 Granite Kashmar 42.8 Ma 58° 23' 49" E 35° 20' 27" N R15904 1-KES-5 PCW116 Granodiorite Kashmar 42.8 Ma 58° 24 11" E 35° 20' 00" N R15905 1-kES-2 PCW117 Granite Kashmar 42.8 Ma 58° 24' 37" E 35° 19'00" N R15906 1-KES-4 PCW118 Granite Kashmar 42.8 Ma 58° 24' 40" E 35° 19'45" N R15907 1-KES-3 PCW119 Tonalite Kashmar 42.8 Ma 58° 24' 53" E 35° 19'30" N R15908 2-KA-8 PCW120 Granodiorite Kashmar 42.8 Ma 58° 26'15" E 35° 20' 45" N R15909 1-KA-2 PCW121 Granite Kashmar 42.8 Ma 58° 27' 30" E 35° 21'36" N R15910 1-KA-1 PCW122 Granite Kashmar 42.4 Ma 58° 27' 42" E 35° 21'00" N R15911 3-KA-1 PCW123 Tonalite Kashmar 42.8 Ma 58° 27' 42" E 35° 18'28" N R15912 2-KA-7 PCW124 Tonalite Kashmar 42.8 Ma 58° 27' 53" E 35° 20'10" N R15913 2-KA-4 PCW125 Granite Kashmar 42.8 Ma 58° 28' 00" E 35° 19'45" N R15914 2-KA-2 PCW126 A.F.G. Kashmar 42.8 Ma 58° 28' 05" E 35° 19'21 "N R15915 2-KA-1 PCW127 Granodiorite Kashmar 42.8 Ma 58° 29' 28" E 35° 18'36" N R15916 2-KA-6 PCW128 A.F.G. Kashmar 42.8 Ma 58° 32' 30" E 35° 18' 21" N R15917 1-QP-6 PCW129 Granodiorite Kashmar 42.8 Ma 58° 37' 08" E 35° 18'51" N R15918 1-QP-4 PCW130 Granite Kashmar 42.5 Ma 58° 38' 07" E 35° 20' 00" N R15919 1-BAH-1 PCW131 A.F.G. Kashmar 42.8 Ma 58° 40' 28" E 35° 17' 36" N R15920 1-NY-3 PCW132 A.F.G. Kashmar 42.8 Ma 58° 42' 30" E 35° 18'30" N R15921 1-FO-5 PCW133 Granite Kashmar 42.8 Ma 58° 44' 07" E 35° 18' 54" N R15922 1-FO-2 PCW134 Granite Kashmar 42.8 Ma 58° 46'15" E 35° 18" 00" N R15923 1-BQ-4 PCW135 Granite Kashmar 42.8 Ma 58° 49'12" E 35° 18'32" N R15924 1-AZ-3 PCW136 Tonalite Kashmar 42.8 Ma 58° 51'30" E 35° 19'20" N R15925 1-AZ-1 PCW137 Granodiorite Kashmar 42.8 Ma 58° 52' 20" E 35° 17'33" N R15926 1-NAM-1 PCW138 Granodiorite Kuh Mish Tertiary 57° 19'50" E 36° 07' 20" N R15927 1-DAR-1 PCW139 Granodiorite Kuh Mish Tertiary 57° 26' 20" E 35° 59' 30" N R15928 1-KUM-1 PCW140 Granodiorite Kuh Mish Tertiary 57° 39' 00" E 35° 53' 00" N R15929 2-KUM-1 PCW141 Gabbro Kuh Mish Tertiary 57°37'10"E 35° 53' 00" N R15930 2-KUM-3 PCW142 Qtz Monzodiorite Kuh Mish Tertiary 57° 37' 50" E 35° 51* 50" N R15931 3-KUM-1 PCW143 Granodiorite Kuh Mish Tertiary 57° 40' 00" E 35° 54' 50" N 35° 54* 30" N R15932 3-KUM-2 PCW144 Qtz Monzodiorite Kuh Mish Tertiary 57° 37' 50" E 35° 53' 30" N R15933 3-KUM-4 PCW145 Granodiorite Kuh Mish Tertiary 57° 41'00" E 35° 53' 10" N R15934 3-KUM-5 PCW146 Qtz Monzodiorite Kuh Mish Tertiary 57° 44 10" E 57° 42'10" E 35° 54' 00" N R15935 3-KUM-7 PCW147 Granodiorite Kuh Mish Tertiary 57°37'10"E 35° 54" 20" N R15936 3-KUM-8 PCW148 Granodiorite Kuh Mish Tertiary 35° 52' 50" N R15937 3-KUM-10 PCW149 Granodiorite Kuh Mish Tertiary 57° 42' 10" E 57° 50' 47" E 35° 22' 45" N R15938 BOR PCW150 Granite Bornavard 123.8 Ma 57° 52' 40" E 35° 22' 40" N R15939 2-BOR-4 PCW151 Granite Bornavard 117.8 Ma 57° 50' 50" E 35° 22' 35" N R15940 1-BOR-2 PCW152 Granite Bornavard 117.8 Ma 57° 46' 50" E 35° 22'15" N R15941 3-BOR-4 PCW153 Granite Bornavard 111.8 Ma 57° 55' 40" E 35° 24' 45" N R15942 SAR-11 PCW154 Granite Bornavard 117.8 Ma 57° 55'12" E 35° 23' 30" N R15943 1-BOR-3 PCW155 Granodiorite Bornavard 149.2 Ma 57° 53' 05" E 35° 23' 35" N R15944 1-BOR-5 PCW156 Tonalite Bornavard 149.2 Ma 57° 51'05" E 35° 23' 40" N R15945 3-BOR-1 PCW157 Tonalite Bornavard 149.2 Ma 57° 48' 33" E 35° 22' 55" N R15946 3-BOR-3 PCW158 Granodiorite Bornavard 145.6 Ma 57° 52' 34" E 35° 23'15" N R15947 BOR-1 PCW159 Granodiorite Bornavard 152.8 Ma 57° 52' 30" E 35° 22' 00" N R15948 1-BOR-1 PCW160 Rhyolite Taknar Mesozoic 300 R15949 TAK-b PCW161 Rhyolite Taknar Mesozoic 57° 48' 42" E 35° 22' 00" N R15950 TAK-4 PCW162 Rhyolite Taknar Mesozoic 57°45'10"E 35° 21' 15" N R15951 TAK-6a PCW163 Rhyolite Taknar Mesozoic 57° 45' 25" E 35° 22' 55" N R15952 TAK-7 PCW164 Rhyolite Taknar Mesozoic 57° 47' 20" E 35°21'45"N R15953 1-B-1 PCW165 Granodiorite Bornavard 149.2 Ma 57° 51'05" E 35° 23" 00" N R15954 2-BOR-1 PCW166 Granite Bornavard 117.8 Ma 57° 49' 08" E 35° 23' 40" N R15955 TAK-1 PCW167 Granite Bornavard 117.8 Ma 57° 47' 38" E 35° 22'10" N R15956 3-KUM-6 PCW168 Qtz Monzodiorite Kuh Mish Tertiary 57° 42' 00" E 35° 53' 20" N R15957 1-FO-4 PCW169 Granite Kashmar 42.8 Ma 58°46'15"E 35° 17'55" N R15958 1-BQ-3 PCW170 Granite Kashmar 42.8 Ma 58° 49' 12" E 35° 18'32" N R15959 1-AZ-4 PCW171 Granodiorite Kashmar 42 .8 Ma 58° 51'30" E 35° 19'21" N Abbreviations: Cat. Catalogue Number, Gar = Garmab, KES = Kesrineh, KA = Kashmar, QP = Quch Plang, BAH = Baharieh, NY = Nay, FO = Forsheh, BQ = Baq Qaleh, AZ = Azqand, NAM = Namin, DAR = Darin, KUM = Kuh Mish, BOR = Bornavard, SAR = Sarborj, TAK = Taknar, B = Bijvard, ANU = Australian National University, PCW = Paul Carr Wolongong, A.F.G. = Alkali feldspar granite, QTZ = quartz monzodiorite