University of Florida Thesis Or Dissertation

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PALEOMAGNETISM AND GEOCHRONOLOGY OF THE PRECAMBRIAN MAFIC DYKE SWARMS OF INDIAN SUB-CONTINENT AND SEDIMENTARY FORMATIONS OF MONGOLIA AND KAZAKHSTAN IN CENTRAL ASIA – UNDERSTANDING THE PRECAMBRIAN PALEOGEOGRAPHY AND TECTONIC EVOLUTION OF INDIA AND CENTRAL ASIAN MICROCONTINENTS

VIMAL ROY PRADHAN

A DISSERTATION PRESENTED TO THE GRADUATE SCHOOL OF THE UNIVERSITY OF FLORIDA IN PARTIAL FULFILLMENT OF THE REQUIREMENTS FOR THE DEGREE OF DOCTOR OF PHILOSOPHY

UNIVERSITY OF FLORIDA

2011

To my parents, wife and my son for their unconditioned love, motivation, patience, sacrifice, encouragement and support without which this dissertation would not have been possible.

ACKNOWLEDGMENTS

I would like to express my sincere appreciation to Dr. Joseph G. Meert, the chairperson of my graduate supervisory committee, for his scientific guidance, encouragement during my graduate study, and financial support. I would also like to extend my gratitude to the members of my committee, Dr. Neil Opdyke, Dr. David

Foster, Dr. Ray Russo and Dr. James Fry, for consistently guiding me through this research and offering fantastic advice and critical review. I especially want to thank Joe for encouraging creative scientific thinking and teaching me so many interesting things about Paleomagnetism, field geology and scientific writing.

Words are not enough to thank my parents, Shanti Sharma and Harish Chandra

Pradhan, and my wife Ritu Sharma, for all of their help along the way and being supportive during my difficult times. Special appreciation goes to my elder brother

Krishan Roy Pradhan and sister-in-law Meera Sharma, my father-in-law Lalit Mohan

Sharma and mother-in-law Pushpa Sharma for their love and encouragement at every step that has influenced my success. I want to thank all of my friends in geology- we have really been through everything together, those from the very beginning of my undergraduate geology in India and the ones here at UF with the highest ambitions.

There are some very crucial people that contributed significantly to the science involved in this research- in lab training, analyses, field work, scientific discussion and review. They are George Kamenov, Kainian Huang, Shawn Malone, Warren Grice,

Daniel Gorman, Trond Torsvik, Laura Gregory, Erfan Ali Mondal, Linda Sohl and Luke

Gommermann. I especially thank Dr. Manoj Pandit and Dr. Paul Mueller for their consistent guidance and help in the successful completion of this program. I want to thank the Geological Sciences faculty for guiding me for the past five years and

teaching me vast amounts of geology. Finally, I thank my best friends throughout the past few years who have always lended a seemingly interested ear to my enthusiasm for science, and with whom I have had some of the greatest experiences- Alex Hastings and Jennifer Gifford. Funding for this research was provided by the National Science

Foundation grants EAR04-0901; EAR05-08597 and EAR09-10888 to Dr. Joseph Meert.

TABLE OF CONTENTS

page

ACKNOWLEDGMENTS ...... 4

LIST OF TABLES ...... 9

LIST OF FIGURES ...... 10

LIST OF ABBREVIATIONS ...... 14

ABSTRACT ...... 16

CHAPTER

1 INTRODUCTION ...... 18

2 A CAUTIONARY NOTE ON THE AGE OF THE PALEOMAGNETIC POLE OBTAINED FROM THE HAROHALLI DYKE SWARMS, DHARWAR CRATON, SOUTHERN INDIA ...... 21

Introductory Remarks...... 21 Paleomagnetism of the Harohalli Alkaline Dykes ...... 23 Geochronology of the Harohalli Dykes...... 25 Discussion ...... 27 Summary ...... 29

3 INDIA‘S CHANGING PLACE IN GLOBAL PROTEROZOIC RECONSTRUCTIONS: A REVIEW OF GEOCHRONOLOGIC CONSTRAINTS AND PALEOMAGNETIC POLES FROM THE DHARWAR, BUNDELKHAND AND MARWAR CRATONS ...... 39

Introductory Remarks...... 39 Geologic Setting ...... 40 The Aravalli and Bundelkhand Cratons ...... 41 The Dharwar Craton ...... 43 Methods ...... 45 Paleomagnetic Methods ...... 45 Geochronologic Methods...... 46 Paleomagnetic and Geochronologic Results ...... 47 Gwalior Traps ...... 47 Paleomagnetic Results ...... 48 Ananatapur Dykes Region ...... 49 Geochronologic Results ...... 50 Paleomagnetic Results ...... 51 Malani Igneous Suite Basement and Early Plutonism ...... 52 Discussion ...... 55

Paleoproterozoic Results (2.5Ga–1.6Ga) ...... 55 Mesoproterozoic Results (1.6Ga-1.1Ga) ...... 58 Neoproterozoic Results (1000-570 Ma) ...... 62 Summary ...... 64

4 PALEOMAGNETIC AND GEOCHRONOLOGICAL STUDIES OF THE MAFIC DYKE SWARMS OF BUNDELKHAND CRATON, CENTRAL INDIA: IMPLICATIONS FOR THE TECTONIC EVOLUTION AND PALEOGEOGRAPHIC RECONSTRUCTIONS ...... 85

Introductory Remarks...... 85 Geological Setting and Previous Work ...... 87 Sampling and Methodology ...... 90 Paleomagnetic Methods ...... 90 Geochronological Methods ...... 91 Results ...... 93 Geochronological Results ...... 93 Paleomagnetic Results ...... 94 Rock Magnetic Results ...... 95 Discussion ...... 96 ~ 1.1 Ga Paleomagnetism ...... 96 Age implications for the Bhander-Rewa sequence of the Upper Vindhyan ...... 97 India in Rodinia supercontinent at 1100 Ma ...... 99 ~ 2.0 Ga Paleomagnetism ...... 106 Tectonic evolution of the Aravalli and Dharwar protocontinents ...... 106 Implications for Columbia ...... 107 Summary ...... 112

5 GEOCHRONOLOGIC AND PALEOMAGNETIC CONSTRAINTS ON THE LATE NEOPROTEROZOIC – EARLY PALEOZOIC VOLCANO-SEDIMENTARY UNITS OF THE CENTRAL ASIAN OROGENIC BELT MICROCONTINENTS. .... 132

Introductory Remarks...... 132 Geological Setting ...... 137 Lesser Karatau Microcontinent Block ...... 137 Baydaric Microcontinent Block ...... 142 Greater Karatau Microcontinental Block ...... 146 Talas Karatau Microcontinental Block ...... 149 Previous Studies ...... 150 Paleomagnetic Studies ...... 150 Geochronological Studies...... 153 Methods and Sampling ...... 156 U-Pb Zircon Geochronology ...... 156 Paleomagnetic Studies ...... 157 Results ...... 159 Geochronological Results ...... 159

Paleomagnetic Results ...... 160 Baydaric block ...... 160 Lesser Karatau block ...... 166 Discussion ...... 168 ~750 Ma Reconstruction ...... 169 ~530 Ma Reconstruction: ...... 172 Summary ...... 175

CONCLUSION ...... 211

LIST OF REFERENCES ...... 216

BIOGRAPHICAL SKETCH ...... 259

LIST OF TABLES

Table page

2-1 Paleomagnetic results from alkaline dykes near Harohalli ...... 37

2-2 Geochronologic data for sample (A) I5143 ...... 38

3-1 Paleomagnetic results from the Gwalior volcanics ...... 81

3-2 Anantapur geochronologic results ...... 82

3-3 Anantapur paleomagnetic results ...... 83

3-4 Barmer geochronologic results ...... 84

4-1 Bundelkhand paleomagnetic results ...... 128

4-2a Geochronologic results from the Bundelkhand older suite of dykes (I9GS-13) 129

4-2b Geochronologic results from the Mahoba dykes (GDM and GDM1) ...... 130

4-3a Paleomagnetic poles at ca. ~1.1 Ga ...... 131

4-3b Paleomagnetic poles at ca. ~2.0 Ga ...... 131

5-1 Radiometric age data from the Late Neoproterozoic felsic/volcanic units in various CAOB microcontinents ...... 203

5-2a Kazakhstan Paleomagnetic Results (High temp. components HTC1 and HTC2) ...... 204

5-2b Baydaric Paleomagnetic results (High temp. components, HTC1 & HTC2) ..... 206

5-3a Geochronologic results from the Talas – Karatau ...... 208

5-3b Geochronologic results from the Greater – Karatau ...... 208

5-4 Paleomagnetic poles for the major cratonic blocks (Neoproterozoic – Cambrian) ...... 209

LIST OF FIGURES

Figure page

2-1 Map of the Harohalli region in Karnataka showing the location of dikes used in this study (after Dawson and Hargraves, 1994)...... 31

2-2 Stereoplots of directions obtained from the Harohalli dikes...... 32

2-3 U/Pb isotopic analysis of the dyke sample I5143...... 33

2-4 U/Pb concordia diagram for concordant zircons from sample I5145...... 34

2-5 Recalculated rubidium/strontium isochron analysis of data reported by Devaraju et at. (1995)...... 35

2-6 Reconstruction at ~1.15 Ga based on our new age and data reported in Pesonen et al. (2003)...... 36

3-1 Generalized tectonic map of Indian subcontinent: Precambrian cratons, mobile belts and lineaments...... 66

3-2 Sketch map of the major units in the Aravalli-Bundelkhand craton, NW India. ... 67

3-3 Sketch map of the Bundelkhand craton showing mafic dyke swarms including the Great Dyke of Mahoba...... 68

3-4 Paleomagnetic and rock magnetic results for the Gwalior traps...... 69

3-5 Stereoplot of site mean directions from our study,the mean of Athavale (1963) Klootvijk (1974) and McElhinny et al. (1978)...... 70

3-6 Sketch map of the Eastern Dharwar craton...... 71

3-7 U/Pb isotopic analysis for the Anantapur dyke sample I595...... 72

3-8 Paleomagnetic and rock magnetic results for the Anatapur dyke samples...... 73

3-9 Stereoplot of site mean directions from our study (Anantapur site I595), the mean of Kumar and Bhalla (1983, 1984) and Poornachandra Rao (2005)...... 74

3-10 Sketch map showing Precambrian stratigraphic units of the Aravalli Mountain Region and further west in NW India ...... 75

3-11 U/Pb isotopic analysis for the Barmer sample Bar05...... 76

3-12 Paleomagnetically-based reconstruction at 1.83 Ga after (Pesonen et al., 2003)...... 77

3-13 Reconstruction at ~1.2 Ga based on our new age from Harohalli Paleomagnetic pole and data reported in Pesonen et al. (2003)...... 78

3-14 Reconstruction at ~1.0 Ga using our new paleomagnetic data from the ~1.0 Ga dated Anantapur dykes of the Dharwar craton, India...... 79

3-15 Paleogeographic reconstruction at 580 Ma from the Harohalli dykes overprint...... 80

4-1 Generalized geologic and tectonic map of the Indian sub-continent with the Precambrian mafic dyke swarms ...... 115

4-2 Sketch map of the major units in the Bundelkhand craton, NW India (modified after Malviya et al., 2006)...... 116

4-3 Field photographs of the Bundelkhand dykes, Central India...... 117

4-4 Backscattered and cathodoluminiscence images of the zircons/uranium bearing minerals for the Bundelkhand dykes...... 118

4-5 U/Pb isotopic analysis of the dyke samples I9GS_13 and great dyke of Mahoba (GDM and GDM1)...... 119

4-6 Orthogonal vector plots from the NW-SE and ENE-WSW trending dykes of the Bundelkhand craton...... 120

4-7 Baked contact test on NW-SE trending dyke I925...... 121

4-8 Rock magnetic results for the Bundelkhand dykes...... 122

4-9 Stereoplots for the three sets of dykes from the Bundelkhand craton...... 123

4-10 Schmidt projection of virtual geomagnetic poles (VGP) from the Mahoba dykes, Majhgawan kimberlite and the Bhander-Rewa...... 124

4-11 Paleomagnetically based reconstruction at ~1.1 Ga...... 125

4-12 Tectonic and paleogeographic interpretation of the paleomagnetic data at ~2.0 Ga for Indian sub-continent...... 126

4-13 Paleomagnetically-based reconstruction at ~2.0 Ga (after Pesonen et al., 2003)...... 127

5-1 Location map of the Central Asian Orogenic Belt within Eurasia...... 178

5-2 Sketches showing the contrasting models for the evolution of the Central Asian Orogenic belt...... 179

5-2 Sketches showing the contrasting models for the evolution of the Central Asian Orogenic belt (Contd.)...... 180

5-3 Stratigraphic column for the Lesser Karatau sequence...... 181

5-4 Simplified geologic map of Karatau ridge...... 182

5-5 Simplified tectonostratigraphic map of Mongolia showing the main tectonic units to the north and south of Main Mongolian Lineament...... 183

5-6 Stratigraphic column for the Baydaric block, central Mongolia...... 184

5-7 Detailed stratigraphic column of the Riphean to Early Paleozoic sedimentary units of the Greater Karatau block, south Kazakhstan...... 185

5-8 Detailed stratigraphic column of the Talas Karatau block in the Kyrgyzstan .... 186

5-9 Geologic/Tectonic sketch map of the Lesser Karatau region in southern Kazakhstan...... 187

5-10 Geologic/Tectonic sketch map of the Baydaric block, central Mongolia...... 188

5-11 U/Pb isotopic analysis of the volcanic tuff sample KT6-1B ...... 189

5-12 Rock magnetic results for Tsagaan Oloom, Bayan Gol and Tamdy suite of rocks...... 191

5-13 Orthogonal vector plots from the carbonate – clastic rocks of the Bayan Gol and Tsagaan Oloom Formations in the Baydaric microcontinent, ...... 192

5-14 Overall site mean directions for the high temperature component (HTC1) and (HTC2) from the Bayan Gol and Tsagaan Oloom formations...... 193

5-15 Stereoplots of the HTC1 component for the Tsagaan Oloom and Bayan Gol formations and the Tamdy suite...... 194

5-16 Equal area projection of the paleomagnetic poles for average site location of the combined Tsagaan Oloom and Bayan Gol formations (HTC1)...... 195

5-17 Inclination-only fold test and the strike test for the HTC2 component of the Tsagaan Oloom and Bayan Gol formations...... 196

5-18 Equal area projection of the paleomagnetic poles for average site location of the component (HTC2) of the Tsagaan Oloom and Bayan Gol formations. . .. 197

5-19 Orthogonal vector plots from the carbonates and dolostones of the Tamdy Series in the Lesser Karatau block ...... 198

5-20 Equal area projection diagrams for the Tamdy suite of rocks (Lesser Karatau, Kazakhstan)...... 199

5-21 Equal area projection of the paleomagnetic poles calculated for average site location of the Early Cambrian – Ordovician Tamdy suite, Lesser Karatau. . .. 200

5-22 Paleogeographic reconstruction for the Neoproterozoic (~750 Ma) interval (simplified from Li et al., 2008) ...... 201

5-23 Paleogeographic reconstruction for Tommotian time (~530 Ma; simplified from Meert and Lieberman, 2008) ...... 202

LIST OF ABBREVIATIONS

2σ two sigma units (expression of error)

α95 circle of 95% confidence about the mean

AF Alternating field

APWP Apparent Polar Wander Path

40Ar/39Ar ratio of argon isotopes 40 and 39

BCP Baicaoping Formation

BDC Bundelkhand craton

BGM Bundelkhand Granite Massif

BKGC Bundelkhand Granitoid Complex

°C degrees Celcius ca. circa

CAOB Central Asian Orogenic Belt

ChRc Characteristic remanent magnetization

CL cathodoluminescence

CS Colider Suite

DC Dharwar craton

Ga Giga annum (Latin: billion years)

GDM Great Dyke of Mahoba

IGRF90 1990 International Geomagnetic Reference Field

IMSLEK Collection of cratons: India, northeastern Madagascar, Sri Lanka, East Antarctica and the Kalahari craton

IRM Isothermal Remanence Magnetization k kappa precision parameter

LA-MC-ICPMS Laser Ablation Multi Collector Inductively Coupled Plasma Mass Spectrometer

Ma Mega annum (Latin: million years)

MIS Malani Igneous Suite

μm micro-meter mT millitesla

MSWD Mean squared weighted deviates

NRM Natural remanent magnetization

Pb-Pb lead-lead isotope geochronology

Rb-Sr rubidium-strontium isotope geochronology

SIMS Secondary Ion mass Spectrometry

TcC Curie temperature on cooling

TcH Curie temperature on heating

THO Trans-Hudson Orogenic belt

TRM Thermal Remanent Magnetization

U-Pb uranium-lead isotope geochronology

VGP Virtual Geomagnetic Pole

Abstract of Dissertation Presented to the Graduate School of the University of Florida in Partial Fulfillment of the Requirements for the Degree of Doctor of Philosophy

Vimal Roy Pradhan

August 2011

Chair: Joseph G. Meert Major: Geology

The first paper presented in this dissertation attempts to examine the igneous history of the younger suite of Harohalli alkaline dykes intruding the Dharwar craton in southern India providing robust constraints on the age of the Harohalli alkaline dyke swarm paleomagnetic pole at ~1192 Ma. A tentative reconstruction at ~1.2 Ga using the Harohalli pole places India at intermediate to high latitudes with Laurentia and

Australia in a configuration quite different from archetypal Rodinia.

The second paper documents new paleomagnetic and geochronological results from the Anatapur alkaline mafic dykes of the Dharwar, Paleoproterozoic Gwalior traps of the Bundelkhand and Malani Igneous Suite of the Marwar cratons of the Indian subcontinent. The paleomagnetic results from our previous studies on the Harohalli alkaline dykes, Upper Vindhyan sequence and Majhgawan kimberlite are also reviewed in an attempt to constrain the paleogeography of the Indian subcontinent from 1.8 Ga to

580 Ma.

Third paper in this dissertation documents new paleomagnetic and geochronological data from the Precambrian mafic dykes intruding granitoids and supracrustals of the Archean Bundelkhand craton (BC) in northern Peninsular India and is significant in constraining the position of India at 2.0 and 1.1 Ga. The paleomagnetic data from the ~ 2.0 Ga Bundelkhand dykes and the paleomagnetic data from the

Bastar/Cuddapah suggest that the North and South Indian blocks of the Peninsular

India were in close proximity by at least 2.5 Ga. The paleomagnetic and geochronological results from Mahoba dyke swarm are significant in constraining the age of the Upper Vindhyan strata to >1000 Ma.

The fourth manuscript documents new paleomagnetic and geochronologic data from the Baydaric block (Mongolia) and Lesser Karatau block (Kazakhstan) in an attempt to enhance our understanding of the complex origin of the microcontinental blocks that form the Central Asian Orogenic Belt (CAOB). The combined analysis of the paleomagnetic and geochronological data from the carbonate rocks of the Tamdy

Series of the Lesser Karatau block and the Bayan Gol and Tsagaan oloom Formations of the Baydaric microcontinents, island arc complexes and other cratonic blocks is consistent with a tropical archipelago of microcontinents in a peri-Siberian configuration.

CHAPTER 1 INTRODUCTION

Proterozoic paleogeographic reconstructions and tectonic studies are an interesting topic of debate. The cratonic blocks of East Gondwana that include India,

Antarctica, Australia, Madagascar, Sri Lanka and the Seychelles (Meert, 2003) represent an important element for the Proterozoic reconstructions and tectonic studies.

There is considerable controversy surrounding the disposition of the East Gondwana elements during the Paleoproterozoic and Mesoproterozoic. These controversies include, but are not limited to; (a) how were the East Gondwana blocks configured in the

Columbia and Rodinia supercontinents? (b) Was East Gondwana a long-lived cohesive unit during the Proterozoic? (c) How were the elements of East and West Gondwana united in the Neoproterozoic-Cambrian? There is a prolific body of scientific literature regarding these questions (Rogers et al, 1995; Meert and Van der Voo 1996; Weil et al,

1998; Powell and Pisarevsky, 2002; Zhao et al., 2004, 2006; Meert, 2003; Meert and

Torsvik, 2003; Pesonen et al., 2003; Veevers, 2004; Collins and Pisarevsky, 2005;

Squire et al., 2006; Li et al., 2008; Meert and Lieberman, 2008); however, there are areas where data are sparse and the gaps prevent definitive answers to the above mentioned questions. Research on the Indian subcontinent provides an important window into this problem as India is both accessible and contains targets of the appropriate age. This dissertation reports on several studies of the alkaline and mafic dyke swarms from the Dharwar craton along with mafic dykes and volcanic rocks from the Bundelkhand craton of Peninsular India. Mafic dyke intrusion in the Indian peninsular shield is widespread and well exposed in the Aravalli- Bundelkhand cratons in the north and northwest, Singhbhum in the east, and the Dharwar and Bastar cratons

in the south and southeast (Figure 1). These dyke swarms range in age from

Paleoproterozoic to Mesoproterozoic and thus provide target rocks for the key paleomagnetic and geochronologic studies during the Precambrian.

Chapters 2, 3 and 4 in this dissertation document the new paleomagnetic and geochronologic data generated in this research in combination with the previously reported data by our group involved in this project and focuses on the two major

Precambrian cratonic blocks of the Indian sub-continent namely the Aravalli-

Bundelkhand craton in the north (North Indian Block) and the Dharwar craton in the south (South Indian Block). The data generated are utilized to decipher the tectonic evolution of the cratonic blocks of the Indian sub-continent during Precambrian. In addition, the data are further incorporated with, and compared to, data that exists from other continents to generate the global paleogeographic maps documenting the position of India in the proposed Proterozoic super-continental assemblies such as Columbia and Rodinia and its subsequent coalescence in the southern Gondwanan landmass.

The second project (Chapter 5) in this dissertation targets the Late Neoproterozoic to Early Paleozoic volcano-sedimentary sequences of the Central Asian Orogenic belt

(CAOB) microcontinents to understand their complex origin and tectonic evolution. The

Central Asian Orogenic Belt or Altaids also known as the Ural-Mongol Fold belt (UMFB) is the largest orogenic system in the Eurasian landmass and was the locus of the growth of Eurasia through the Neoproterozoic and Paleozoic. The kinematic history of the CAOB is highly controversial due to myriad conflicting models on its tectonic evolution (see Zonenshain et al., 1990; Mossakovsky et al., 1993; Kheraskova et al.,

2003; Didenko et al., 1994; Filipova et al., 2001; Şengör and Natal‘in, 1996; Yukubchuk

et al., 2001; 2002; Stampfli and Borel, 2002; Puchkov, 2000; Windley et al., 2007; Xiao et al., 2010; Levashova et al., 2011a, b). We present new U/Pb zircon ages from the

Late Neoproterozoic felsic volcanic sequences lying unconformably below a late

Neoproterozoic diamictite deposit at Greater Karatau (Kazakhstan) and Talas Karatau

(Kyrgyzstan). New paleomagnetic data are also reported from the Early Paleozoic carbonate beds of the Tamdy Series in the Lesser Karatau microcontinent in southern

Kazakhstan along with data from the carbonate sequences in the Tsagaan-Oloom and

Bayan-Gol Formations of the Baydaric microcontinental block in central Mongolia. The new radiometric ages and the paleomagnetic directions are compared with existing data from other CAOB microcontinental blocks during Late Neoproterozoic-Cambrian to improve our understanding of the origin and tectonic evolution of these microcontinents.

The first three manuscripts (Chapters 2, 3 and 4) included in this dissertation are either published (Pradhan et al., 2008; Pradhan et al., 2010) or submitted articles in peer reviewed journals (Pradhan et al., under review; Precambrian Research) whereas the last manuscript (Chapter 5) will be submitted to Tectonophysics within the next few weeks.

CHAPTER 2 A CAUTIONARY NOTE ON THE AGE OF THE PALEOMAGNETIC POLE OBTAINED FROM THE HAROHALLI DYKE SWARMS, DHARWAR CRATON, SOUTHERN INDIA

Introductory Remarks

The Precambrian basement of the south Indian peninsular shield is made up of granite greenstone belts which are crisscrossed by several generations of dolerite, and alkaline dykes (Halls, 1982). These mafic intrusive dykes, particularly from the Harohalli region of the Dharwar craton exposed south of Bangalore city, are the source of some high-quality paleomagnetic data from India for the Proterozoic (Fig. 1). Accordingly, a rather robust paleomagnetic dataset had been generated by various groups of workers for these Meso- to Neoproterozoic mafic intrusives (Naqvi et al., 1974; Hargraves and

Bhalla, 1983; Radhakrishna and Mathew, 1993; Dawson and Hargraves, 1994;

Radhakrishna and Mathew, 1996; Halls et al., 2007). The main problem is that there is a paucity of robust radiometric ages for the younger of these paleomagnetically significant dykes. The local geological setting of the Harohalli region consists of a basement gneissic complex (~2900 Ma BGC-Peninsular gneisses) and the elongated body of the

Closepet Granite (~2500 Ma), both trellised by various dyke intrusions. Separate geochronological studies had been carried out and two distinct sets of dykes have been reported in the area by various workers. The older set of dykes consists of a series of

~E-W trending doleritic dykes that were dated by Rb-Sr whole rock methods at

2370±230 Ma by Ikramuddin and Stueber(1976) and have been more recently dated by

U/Pb methods on baddeleyite grains at 2365.5±1.1 Ma by French et al. (2004).

Additional work on correlative dykes by Halls et al. (2007) on baddeleyites yielded a robust U/Pb age of 2370±1 Ma. Halls et al. (2007) refers to these older dykes as the

Bangalore dyke swarm and we follow that terminology in our paper. The younger dykes are referred to as the Harohalli alkaline dykes.

The Harohalli alkaline dykes, trending in ~N-S directions (Fig. 1) were dated by

Ikramuddin and Stueber (1976) at 814 ± 34 Ma (Rb-Sr whole rock) and 810 ± 25 Ma (K-

Ar whole rock). Later Anil-Kumar et al. (1989) reported a whole rock-feldspar Rb-Sr age of 823 ± 15 Ma. These results were considered to represent strong evidence for an 850-

800 Ma age for the alkaline dykes of Harohalli. The ages are cited by numerous authors investigating the assembly of Gondwana, the breakup of Rodinia as well as geodynamic models positing rapid true polar wander (Meert and Powell, 2001; Meert, 2003; Meert and Torsvik, 2003; Li et al., 2004; Maloof et al., 2006; Piper, 2007). Igneous intrusive events in the Dharwar craton are not limited to the Harohalli dykes. A series of dykes and igneous intrusions can be found in many areas throughout the Dharwar craton, but most have poor age control. We call attention to some recent studies with robust geochronology from the Dharwar craton. Kumar et al. (2007) reported precise U/Pb perovskite and Rb-Sr phlogopite ages of ~1.1 Ga for the kimberlites from eastern

Dharwar craton, which they linked with a short-lived mantle plume activity that took place in this region. Dykes from the Bastar and Dharwar cratons were recently shown to be part of a1883-1891 Ma large igneous province by French et al. (2007). Patranabis-

Deb et al. (2007) examined igneous rocks capping the Chhattisgarh basin and argued that this particular Purana basin was some 500 Ma older than assumed. The U/Pb ages on the detrital zircons obtained from the Chattisgarh basin are in the range from 990-

1020 Ma. The present study of the Harohalli alkaline dykes attempt to examine the igneous history of the younger suite of dykes using LA-ICP-MS geochronologic methods

and provide robust constraints on the age of the Harohalli alkaline dyke swarm paleomagnetic pole.

Paleomagnetism of the Harohalli Alkaline Dykes

The dykes sampled for geochronologic study were also sampled for paleomagnetic study. Samples were drilled in the field using a water cooled gasoline- powered hand drill and oriented using both sun and magnetic compass. Due to deep tropical weathering in the region, samples were removed from ‗boulders‘ of dyke that looked to be in place. In the case of dykes I5145 and I5150 there was no ambiguity about the in-situ nature of the sampled material. I5143 was sampled carefully, but it was not always clear whether or not the material sampled was in place or boulders. Samples were then cut and measured using a Molspin spinner magnetometer at the University of

Florida. Pilot samples were subjected to detailed stepwise alternating field (AF) or thermal demagnetization and the remaining samples were treated based on the best response.

Dyke I5143 yielded no consistent paleomagnetic directions although individual samples were well behaved. The results confirmed our suspicion that the boulders we sampled were not all in-situ. Dawson and Hargraves (1994) also sampled this dyke but did not obtain any useful paleomagnetic results. Dyke I5145 gave consistent directions that were to the north and intermediate up with a mean declination=4.6° and inclination=-36.2° (k=21, α95=12.6°; n=8; Fig. 2a; Table 1). Dyke I5150 shows a low coercivity direction to the north and shallow up (Table 1; Dec=6.4°, I=-8.1°, k=253,

α95=3.8°, n=7). This direction is distinct from the present earth‘s field direction, but comparable to the B-direction obtained by Halls et al. (2007). Halls et al. (2007) interpreted the ubiquitous B-component observed in the Bangalore swarm as an

Ediacaran-Cambrian remagnetization (Fig. 2e). The remaining I5150 dyke samples were treated by AF demagnetization. Dyke I5150 yielded a well-defined high coercivity grouping with a mean declination=52.3° and inclination= -85.7° (k=134, α95=4.8°, n=8) with a resultant virtual geomagnetic pole at 7.3° N, 70.6° E (Fig. 2b and 2c). The direction and pole are similar to the directions reported from the adjacent dyke (site 25;

Fig. 1) by Radhakrishna and Mathew (1996).

Two separate high-latitude paleomagnetic poles have been published for the

Harohalli/Bangalore dykes. The paleomagnetic pole for Bangalore suite of dykes was reported by Dawson and Hargraves (1994) at 9.5° S, 242.4° E (α95=9.0°). Multiple studies on the Harohalli alkaline dykes (Dawson and Hargraves, 1994; Radhakrishna and Mathew, 1996) yield a grand mean result with a declination of 2.3° and an inclination of +83.6° (k=34, α95=8.4°; n=10 including one new site from our visit in 2005;

Table 1; Fig. 2d). The resultant mean paleomagnetic pole is located at 24.9° S, 258° E

(A95=15°). We note that while the overall mean result from these dykes has a good grouping and small α95, four of the ten dykes had α95 values exceeding 15° and 4 dykes had only 3 samples. Nevertheless, this pole is typically cited as a key ~823 Ma pole for India and has been used to make a variety of tectonic/geodynamic conclusions

(Li et al., 2004; Maloof et al., 2006).

The paleomagnetic poles obtained from the ~2370 Ma Bangalore dyke swarm overlap with the younger alkaline suite and this led Dawson and Hargraves (1994) to speculate about possible remagnetization of the Bangalore suite of dykes. A more detailed examination of coeval dykes throughout the region by Halls et al. (2007) concluded that the older Bangalore suite of dykes carry a near primary magnetization

that is fortuitously coincident with the directions from the younger Harohalli alkaline dykes.

Geochronology of the Harohalli Dykes

For the current study, four mafic dyke samples were analyzed for their U/Pb isotopic signatures (Fig. 1). Zircon grains were concentrated through standard gravity and magnetic techniques from pulverized samples. The doleritic samples were first crushed, disk milled and sieved to a fraction of grain size < 125μm to liberate as many zircon grains as possible. These were then rinsed, followed by water table treatment with slow sample feed rates. This was followed by heavy liquid mineral separation with multiple agitation periods to reduce the number of entrapped grains in the lower density fraction. The final step involved repeated passes on a Frantz Isodynamic separator up to a current of 1.0 A (2-4° tilt). Approximately 30-50 fresh looking, euhedral to almost anhedral zircon grains were handpicked from the two samples labeled I5143 and I5145 under an optical microscope to ensure the selection of only the clearest grains and fractions of grains, mounted first in resin and then polished to expose median sections.

The plugs were sonicated and cleaned in nitric acid to remove any common Pb surface contamination. U-Pb concentrations were collected and analyzed using the University of

Florida Nu LA-MC-ICPMS system. Data calibration and drift corrections were based on the Forest Centre-1 (Duluth Gabbro) zircon standard dated at 1099.1±0.5 Ma (Paces and Miller, 1993). FC-1 which is used as the zircon standard for calculation of all the ages are typically large, clear, colorless, inclusion free and sector zoned. A laser beam diameter of 40μm and laser energy density of ~10 J/cm2 was applied to analyze the samples I5143-a, b and c. A few of these analyses were discarded on account of high levels of observed discordance interpreted to reflect incorporation of common lead.

U/Pb isotopic analyses were conducted at the Department of Geological Sciences

(University of Florida) on a Nu Plasma multicollector plasma source mass spectrometer equipped with three ion counters and 12 Faraday detectors. The MC–ICPMS is equipped with a specially designed collector block for simultaneous acquisition of 204Pb

(204Hg), 206Pb and 207Pb signals on the ion-counting detectors and 235Uand 238U on the Faraday detectors (see Simonetti et al., 2005). Mounted zircon grains were laser ablated using a New Wave 213 nm ultraviolet beam. During U/Pb analyses, the sample was decrepitated in a He stream and then mixed with Ar-gas for induction into the mass spectrometer. Background measurements were performed before each analysis for blank correction and contributions from 204Hg. Each sample was ablated for ~30 s in an effort to minimize pit depth and fractionation. Data reduction and correction were conducted using a combination of in-house software and Isoplot (Ludwig, 1999).

Additional details can be found in Mueller et al. (2008).

Geochronologic results: U/Pb ages from the zircon were determined for two out of five Harohalli alkaline dyke samples. The first dyke (I5145) yielded zircon blades and fragments. A second N-S trending dyke (I5143) near Harohalli yielded one well-faceted zircon and several fragments/tips of zircon (Fig. 3a). Six laser spots from three different zircons (sample I5143) yielded a concordant U/Pb age of 1192±10 Ma (2σ; MSWD=2.4;

Fig. 3a). Six spots on 3 fragments and tips (sample I5143) yielded discordant ages ranging from 580-1139 Ma. A discordia fit using these grains (forced through the three concordant grains noted above) yielded a lower intercept age of 506 +18/-19 Ma and an upper intercept at 1212 +20/-21 Ma (MSWD=1.15; Fig. 3b). The lower age may represent Pb-loss event associated with the final assembly of Gondwana (Collins et al.,

2007; Meert, 2003) and the upper intercept is within error of the 1192±10 Ma crystallization age of the dyke. Analytical data are given in Appendix A. Dyke I5145 yielded 15 grains of zircon which gave a concordant age of 2941±14.2 Ma (2σ;

MSWD=2.2; Fig. 4). The implication of this age is discussed below. Analytical data are given in Appendix B.

Discussion

The paleomagnetic pole from the alkaline suite of the Harohalli dykes has been updated and recalculated at 24.9° S, 258° E (A95=15°). The paleomagnetic directions from 10 dykes show a dual-polarity magnetization. The mean of 8 normal polarity dykes yields a declination=10.4° and an inclination of +81.1° (Table 1). Two reverse polarity dykes yielded a mean declination=61.5° and an inclination of -84.3°. A reversal test

(McFadden and McElhinny, 1990) was negative (critical angle λc=7.0°; observed angle

λo=15.8°) signifying that the angle between the normal and reverse directions is much larger than required for a common mean direction. Because of the limited number of reversely magnetized sites, the test may simply be an indication that those two dykes did not adequately average secular variation.

The previous age estimates for the alkaline dykes at Harohalli are called into question by our current study. Although only 2 N-S trending dykes contained dateable material, neither of these dykes gave ages in the range from 800-900 Ma. The most concordant grains yielded an age of 1192±10 Ma which is far older than all previously published ages on the Harohalli and supposedly correlative dykes (Ikramuddin and

Stueber, 1976); Anil-Kumar et al., 1989 and Devaraju et al., 1995). The study by

Ikramuddin and Stueber (1976) used only whole rock samples to derive their K-Ar and

Rb-Sr ages. While whole-rock Rb-Sr and K-Ar ages may yield reasonable age

estimates, the geochemical/geological requirements for whole-rock dating-especially for older mafic dykes, are strict and rarely met (Bernard-Griffiths, 1989). The most detailed geochemical and geochronological study of related dykes was conducted by Devaraju et al., 1995). The study by Devaraju et al. (1995) on alkaline dykes from southern India considered equivalent to the N-S trending alkaline dykes at Harohalli yielded two distinct ages. According to that study, a regression on 6 whole rock samples yielded an isochron with an age of 1046±21 Ma (2σ) and an initial 87Sr/86Sr ratio of 0.70476 ±

0.0001. Devaraju et al. (1995) rejected the data from one of the dykes and arrived at an isochron age of 832±40 Ma for the alkaline dykes.

We have re-analyzed the data reported in the paper by Devaraju et al. (1995) using the ISOPLOT (Ludwig, 2003) programme and were unable to reproduce the ages based on the data reported in the paper. A robust regression (Ludwig, 2003) through all

6 whole rock data points yields an age of 883 +130/-220 Ma (MSWD=76;

(87Sr/86Sr)i=0.7065; Fig. 5a). According to their analysis a better fit was obtained by removing one sample from the group. We find a slightly poorer ‗errorchron‘ is generated from a robust regression of these data at 806 +160/-320 Ma ((87Sr/86Sr)i = 0.7079) Fig.

5b). Although both overlap with previously cited ages for the Harohalli dyke swarm neither is a particularly good fit.

Data from dyke I5145 (Fig. 1) yielded 15 concordant zircons with a concordant age of 2941 ± 14.2 Ma (MSWD=2.2; Fig. 4). There are no known dykes of this age in peninsular India, but the peninsular gneisses are of this approximate age (Mojzsis et al.,

2003). Our interpretation is that this dyke contains a suite of inherited zircons from the basement region. Paleomagnetic data from this dyke show only an northerly and

shallow-up direction similar to the B-directions reported by Halls et al. (2007) and also the low-coercivity and high temperature directions in dyke I5150 (Fig. 2e and Table 1).

We interpret this as an overprint of ~Ediacaran age and it may correspond to the slightly discordant ~590 Ma ages observed in sample I5143 (Fig. 3b; Appendix A).

Summary

The Harohalli alkaline dykes pole has been treated as a robust paleomagnetic pole for the Indian continent in spite of the poor age controls on the dyke. Previous estimates on the dykes resulted in broadly consistent Rb-Sr and K-Ar whole rock ages between 800-850 Ma. We report here the first attempt to date the dykes using the U-Pb method and have obtained a concordant age of 1192±10 Ma (2σ; MSWD=2.4) for one of the largest alkaline dykes in the area. A second dyke yielded 15 concordant zircons with an age of 2941±14.2 Ma (2σ; MSWD=2.2). This age is considered to reflect inheritance from the basement rocks in the region.

We treat our new age with some caution as it requires an almost 400 Ma revision in the age of the Harohalli alkaline dykes. We also note that the dyke which yielded the

1192 Ma age did not yield paleomagnetic directions and the dyke that yielded paleomagnetic directions did not yield any useful minerals for geochronology. Since the

Harohalli dykes intrude the peninsular gneiss (2.9 Ga) and the Closepet granite (2.5 Ga) these would be the most likely sources of inherited zircons. Inheritance of 1192 Ma zircons is considered an ad-hoc explanation for the age obtained here.

An interesting implication of our new age is that it would place India adjacent to both Australia and Laurentia in a ca. 1.2 Ga reconstruction (Fig. 6; Pesonen et al.,

2003). Lastly, we note that several extreme geodynamic models (inertial interchange and true polar wander events) have been proposed based in part on presumed ~823

Ma age for the Harohalli alkaline dykes (Li et al., 2004; Maloof et al., 2006). If our new age data are indeed accurate, then those geodynamic models become less well- constrained. We are continuing our work on the Harohalli alkaline dykes, but we urge caution in the uncritical use of the extant paleomagnetic and geochronologic results from these dykes.

Acknowledgements: This work was supported by a grant from the US National

Science Foundation to JGM (EAR04- 09101). We wish to thank George Kamenov and

Warren Grice for help with U-Pb data reduction and to Laura Gregory & Richard Ernst for a constructive review of the manuscript.

Figure 2-1. Map of the Harohalli region in Karnataka showing the location of dikes used in this study (after Dawson and Hargraves, 1994). Site 25 (white box) is a sample from Radhakrishna and Joseph (1996). Geochronologic samples that yielded zircon are I5143 and I5145.

Figure 2-2. Stereoplots of directions obtained from the Harohalli dikes (a) Stereoplot of directions obtained from the dike at Site I5145. (b) AF demagnetization of sample I5150-7b from site I5150 showing removal of a northern and shallowly directed overprint and a great-circle trend towards an easterly and steeply up direction. (c) Individual samples and mean direction for site I5150. (d) Compiled results from 3 studies on the alkaline dikes at Harohalli (overall mean direction is shaded). (e) Stereoplot of B-component directions reported here and in Halls et al. (2007) converted to a common site location at 12.6 N, 77.4 E. In all stereoplots open/closed circles represent up (-)/down(+) inclinations.

Figure 2-3. U/Pb isotopic analysis of the dyke sample I5143. (a) U/Pb concordia diagram for concordant zircons from sample I5143. inset: ‗tuffzirc‘ analysis +22.2 (Ludwig, 2003) for the concordant grains yielding an age of 1184.7 /-19.7. (b) Discordia plot using fragmentary zircons from I5143 and the 6 concordant +20 analyses in (a). Upper intercept age is 1212 /-21 and the lower intercept +18 age is 506 /-19. Photo of zircon grain from plug 1 grains 1-3 (Appendix A).

Figure 2-4. U/Pb concordia diagram for concordant zircons from sample I5145. inset: ‗tuffzirc‘ analysis (Ludwig, 2003) for the concordant grains yielding an age of +29.5 2924.3 /-22.3.

Figure 2-5. Recalculated rubidium/strontium isochron analysis of data reported by Devaraju et at. (1995). (a) Recalculated rubidium-strontium isochron analysis of data reported by Devaraju et al., 1995 for alkaline dikes in Karnataka using all data reported in that paper and (b) recalculated rubidium-strontium isochron analysis of data in the Devaraju et al. (1995) paper removing sample M-80. Both results differ from ages reported in the original paper.

Figure 2-6. Reconstruction at ~1.15 Ga based on our new age and data reported in Pesonen et al. (2003).

Table 2-1. Paleomagnetic results from alkaline dykes near Harohalli Site N Dec Inc k 95 Plat Plong Reference 18 5 018.4 83.9 065 9.5 23.8 S 261.6 E Dawson & Hargraves, 1994 21 6 032.0 82.0 028 12.8 25.6 S 266.8 E Dawson & Hargraves, 1994 17 5 030.9 75.2 077 09.0 35.8 S 274.7 E Radhakrishna & Mathew, 1996 19 3 337.9 73.9 011 38.0 39.9 S 243.3 E Radhakrishna & Mathew, 1996 20* 4 200.5 68.5 043 19.0 23.4 N 244.0 E Radhakrishna & Mathew, 1996 21 3 355.1 79.1 018 29.0 33.4 S 255.5 E Radhakrishna & Mathew, 1996 22 3 000.5 65.7 239 08.0 54.5 S 258.2 E Radhakrishna & Mathew, 1996 24 4 043.7 71.5 012 28.0 35.1 S 285.5 E Radhakrishna & Mathew, 1996 25 3 067.4 -82.8 028 13.0 06.9 N 04.4 E Radhakrishna & Mathew, 1996 I5150-low 7 006.4 0-8.1 253 3.8 72.1 S 056.2 E This study-B Component I5150-high 8 052.3 -85.7 134 4.8 07.3 N 070.6 E This study-A Component I5145-high 8 004.6 -36.2 021 12.6 57.0 S 069.5 E This study-B Component Mean- B** 12 002.0 0-1.6 027 8.4 76.5 S 068.8 E This study and Halls et al., 2007 Mean 8 N 010.4 81.1 033 9.8 29.7 S 261.0 E This study-A Mean 2 R 061.5 -84.3 ------07.1 N 067.4 E This study-A Overall Mean 10 002.3 83.6 034 8.4 24.9 S 258.0 E This study-A *Recalculated pole originally reported incorrectly at 7.8 S, 251.3 E. **Calculated for a site at 12.6N, 77.4 E. Dec=declination, Inc=inclination, k=kappa precision parameter; 95=cone of 95% confidence about the mean Direction, Plat=Paleopole latitude, Plong=Paleopole longitude

Table 2-2. Geochronologic data for sample (A) I5143 1σ 1σ 1σ Age Age Age % Grain 207Pb/206Pb error 206Pb/238U error *207Pb/235U error (Ma) 1σ (Ma) 1σ (Ma) 1σ Disc Plug-1 206Pb/238U 207Pb/235U 207Pb/206Pb I5143_1 0.06345 0.00008 0.15998 0.00089 1.6473 0.098 1194 17.9 1206 12.3 1226 12 1 I5143_2 0.06273 0.00022 0.15849 0.0012 1.4725 0.089 1185 18.6 1191 12.9 1204 14 1 I5143_3 0.06201 0.0002 0.15489 0.00093 1.48487 0.090 1160 17.6 1167 12.4 1181 14 1 I5143_4 0.04810 0.00016 0.07357 0.00057 0.49256 0.008 576 9.5 593 8.3 659 15 3 I5143_5 0.04927 0.00018 0.07477 0.00044 0.52290 0.007 584 9.3 611 8.2 710 15 4 I5143_6 0.04828 0.00012 0.07429 0.00029 0.50903 0.013 581 8.9 599 7.8 667 14 3 I5143_7 0.05003 0.00022 0.07528 0.00064 0.52139 0.010 587 9.9 620 8.9 742 16 5 I5143_9 0.06004 0.00063 0.12643 0.0013 0.71417 0.260 957 17 1007 14 1117 24 5 I5143_10 0.05977 0.00073 0.12033 0.0023 0.96380 0.064 913 21 972 18 1108 27 6 Plug-2 I5143_04 0.06352 0.00172 0.15217 0.00103 1.30551 0.151 1207 14 1214 21 1227 54 1 I5143_05 0.06266 0.00182 0.14637 0.00193 0.81900 0.138 1165 18 1177 23 1200 58 1 Plug-3 I5143_24 0.06484 0.00018 0.16658 0.00320 1.60018 0.073 1178 24 1180 18 1185 23 0

(B) I5145 1σ 1σ 1σ Age Age Age % Grain 207Pb/206Pb error 206Pb/238U error *207Pb/235U error (Ma) 1σ (Ma) 1σ (Ma) 1σ Disc 206 238 207 235 207 206 Plug-1 Pb/ U Pb/ U Pb/ Pb I5145_01 0.20139 0.00094 0.4241 0.01065 11.57250 0.73620 2859 126 2941 53.5 2998 23 3 I5145_03 0.19433 0.00330 0.42238 0.00966 10.67222 0.36281 2867 124 2910 54 2940 35 1 I5145_05 0.19540 0.00150 0.42128 0.00929 11.26985 0.29659 2856 123 2911 52 2949 25 2 I5145_06 0.19727 0.00110 0.41022 0.01147 10.87436 0.52806 2770 127 2884 55 2965 24 4 I5145_08 0.19254 0.00113 0.43914 0.00241 11.24917 0.28656 2996 115 2954 47 2925 24 -1 I5145_09 0.20006 0.00131 0.41286 0.00903 11.10058 0.56857 2781 122 2902 53 2987 24 4 I5145_10 0.19294 0.00147 0.4288 0.02725 11.29180 1.81756 2918 185 2924 75 2929 25 0 I5145_16 0.19276 0.00098 0.41524 0.00485 11.41154 0.39880 2819 114 2883 49 2927 23 2 I5145_17 0.18764 0.00198 0.44235 0.00401 11.47673 0.35060 3034 116 2944 48 2884 25 -3 I5145_18 0.19368 0.00105 0.44900 0.00312 11.85102 0.33358 3066 117 2987 47 2935 24 -3 I5145_22 0.19287 0.00121 0.44343 0.00448 12.53204 1.11569 3027 117 2968 48 2928 23 -2 I5145_23 0.19191 0.00098 0.42950 0.00866 12.65479 3.22206 2926 122 2923 51 2920 24 0 I5145_25 0.19459 0.00221 0.41241 0.00443 11.09926 0.31495 2793 113 2881 50 2942 25 3 I5145-26 0.19624 0.00087 0.42662 0.00344 11.56610 0.39893 2892 114 2930 48 2956 26 1 I5145-27 0.19455 0.00183 0.44014 0.01144 11.65435 0.66839 2997 130 2964 53 2942 24 -1

CHAPTER-3 INDIA‘S CHANGING PLACE IN GLOBAL PROTEROZOIC RECONSTRUCTIONS: A REVIEW OF GEOCHRONOLOGIC CONSTRAINTS AND PALEOMAGNETIC POLES FROM THE DHARWAR, BUNDELKHAND AND MARWAR CRATONS

Introductory Remarks

The East Gondwana continent is comprised of India, Australia, East Antarctica,

Madagascar and Sri Lanka. The tectonic history leading to the amalgamation of East

Gondwana can be evaluated through the use of well-dated paleomagnetic poles.

Madagascar, Sri Lanka and the exposed regions of East Antarctica are dissected by numerous orogenic belts related to Gondwana assembly and therefore target rocks in these regions are poor candidates for paleomagnetic study. In contrast, both India and

Australia contain numerous unaltered sequences of sedimentary and igneous rocks that may be used to establish their drift histories prior to the final fusion of Gondwana.

In the past few years, we have focused on improving the Meso-Neoproterozoic drift history of India through the acquisition of robust paleomagnetic and geochronologic data from key sedimentary and igneous sequences. In this paper, we summarize those findings (Pradhan et al., 2008; Gregory et al., 2006, 2009; Malone et al., 2008) along with new preliminary geochronologic and paleomagnetic data from the Anantapur dikes

(Dharwar craton), Gwalior volcanics (Bundelkhand craton) and the Malani Igneous Suite

(Marwar craton). These new data, in combination with recently published paleomagnetic and geochronologic results lead to the proposal of a new set of paleogeographic maps for India. The paper also evaluates several controversial proposals regarding the geodynamic evolution of East Gondwana.

Geologic Setting

Four distinct cratonic blocks are recognized in the Indian shield; (1) the Aravalli and Bundelkhand craton in northwestern and central regions; (2) the Bastar craton in south-central region; (3) the Singhbhum craton in eastern region; and (4) the Dharwar craton covering the southern half of peninsular India (Naqvi et al., 1974). These cratonic regions are bordered by orogenic belts (Fig. 1). The region to the west of the Aravalli–

Delhi Proterozoic belts in NW India exposes well developed Phanerozoic sequences, however, it is one of the least understood in terms of ‗Precambrian basement‘. This terrane is also known as the Marwar craton.

In addition to the cratonic rocks and adjacent orogenic belts, many regions contain purported Meso- to Neoproterozoic sedimentary basins. These so-called ―Purana‖ basins (Fig. 1) include the Vindhyan, Chhattisgarh, Prahnita-Godavari, Cuddapah,

Bhima, Indravati, Badami and Kaladgi basins. Age constraints are generally lacking in most of these basins; however, Chakraborty (2006) tentatively called all of these basins

‗classic examples of Proterozoic intracontinental basins‘ within India and remarked on similarity in their lithologies, depositional environments and stratigraphic architecture.

Recent work in both the Vindhyan and Chhattisgarh basins (the results come from these basins only) supports Meso-Paleoproterozoic ages for these basins (Malone et al.,

2008; Patranabis-Deb et al., 2008; Basu et al., 2008; Das et al., 2009). Each of the above mentioned cratons has also been intruded by a number of igneous bodies including large swarms of mafic dykes, kimberlites, lamproites and granitoids. The ages of many of these bodies are still poorly constrained, however some recent geochronologic work (French et al., 2008; Halls et al., 2007; Gregory et al., 2006;

Pradhan et al., 2008) have begun to sort out these distinct intrusive events in the evolution of the Indian subcontinent.

The Aravalli and Bundelkhand Cratons

The Aravalli and Bundelkhand cratons (Fig. 2) are bounded to the northeast by the

Mesoproterozoic Vindhyan Basin and Indo- Gangetic alluvium and to the south by the northern edge of the Deccan Traps. The Great Boundary Fault (GBF) separates the

Aravalli cratonic block to the west and the Bundelkhand–Gwalior block to the east of the

GBF. The NE-trending Aravalli–Delhi Fold Belt (NDFB) defines the western boundary of the Aravalli/Bundelkhand craton(s). The Bundelkhand and Aravalli cratons are also separated from the Bastar and Singhbhum cratons in the south by the Narmada–Son lineament (Goodwin, 1991; Naqvi and Rogers, 1987).

Most of the Aravalli craton is underlain by the 3.3 Ga Banded Gneiss Complex

(hereafter BGC; Wiedenbeck et al., 1996; Roy and Kröner, 1996). The BGC is composed of migmatites, gneisses, schists, amphibolites, pelites and metasedimentary rocks. The BGC in the Aravalli region is bounded by the Aravalli–Delhi Fold Belts. Age constraints on both the Aravalli and Delhi Fold Belts are only poorly constrained to 2.5–

1.9 Ga and 1.8–0.85 Ga, respectively (Gupta et al., 1980, 1997). The ages of metamorphism in the Aravalli craton have tighter constraints. Roy et al. (2005) argue that the main pulse of metamorphism took place between 1725 and 1621Ma at the outset of the Delhi Orogenic Cycle. Buick et al. (2006) also obtained metamorphic ages of ~1720Ma for part of the Delhi Belt that was formerly thought to be reworked high- grade metamorphic rocks of Archaen age. Buick et al. (2006) also report a younger metamorphic episode that took place between ca.950 and 940Ma and suggested that these ages may be part of a larger metamorphic and igneous event (Pandit et al., 2003;

Tobisch et al., 1994). Support for this metamorphic event is also observed in detrital zircon spectra from the Sonia and Girbakhar sandstones in the Marwar Supergroup of

Rajasthan (Malone et al., 2008). The best constrained magmatic events took place between 1711 and 1660Ma (Kaur and Mehta, 2007; Kaur et al., 2009) and appear to be coincident with the main phase of metamorphism documented by both Buick et al.

(2006) and Roy et al. (2005). Wiedenbeck et al. (1996) analyzed ion microprobe

207Pb/206Pb zircon isotopic data on Berach Granite and suggested a ~2.5 Ga stabilization age for the southern segment of the Aravalli craton, based on the uniformity of the Late Archaean and Early Proterozoic crystallization ages.

Other tectono-metamorphic events in the Aravalli region include a metallogenic event at around 990Ma followed by a tectonothermal event between 990 and 836Ma

(Deb et al., 2001; Roy, 2001). The last major episode of Neoproterozoic igneous activity took place during the emplacement/eruption of the Malani igneous rocks between ca.770 and 750Ma (Torsvik et al., 2001a; Gregory et al., 2009). The Malani province is overlain by the sedimentary sequences of the late Neoproterozoic (to Early Cambrian)

Marwar Supergroup (Pandit et al., 2001).

The Bundelkhand craton lies to the east of the Aravalli–Delhi Fold Belt (Fig. 3).

The most conspicuous feature of the region is the Bundelkhand Igneous Complex that intrudes enclaves of schists, gneisses, banded iron formations, mafic volcanic rocks and quartzites (Goodwin, 1991). Ages of the enclaves are not known, but there are a few ages on the granites that intrude them. The Bundelkhand granite is dated to

2492±10Ma (Mondal et al., 2002), and is therefore contemporaneous with the intrusion of the Berach Granite in the Aravalli craton dated ~2500Ma (Sivaraman and Odam,

1982; Tucker, personal communication). Numerous mafic dykes of unknown age intrude the Bundelkhand Igneous Complex. Rao (2004) suggests that most of the mafic dikes were emplaced in two phases, one at 2.15 Ga and the second at 2.0 Ga based on the

40Ar/39Ar age determination of the dolerite dykes.

Overlying the Bundelkhand granite are metasedimentary rocks of the 1854±7Ma

Hindoli Group (Deb et al., 2002) that some consider equivalent to the Gwalior Group.

These metasedimentary rocks are slightly older than the depositional ages of sedimentary sequences in the Vindhyan Basin (ca.1700–1050 Ma; Ray et al., 2002,

2003; Rasmussen et al., 2002; Sarangi et al., 2004; Malone et al., 2008). The

1073±13.7Ma Majhgawan kimberlite (Gregory et al., 2006) intrudes the lower Vindhyan as well as the Kaimur sandstone in the lower part of the Upper Vindhyan sequence.

The Dharwar Craton

The Dharwar craton (DC) in southern India (Fig. 1) is the largest and one of the most extensively studied of the Precambrian cratons of Indian Peninsular shield.

Together with the Southern Granulite Terrane (SGT), the Dharwar craton forms the

Dravidian Shield covering an area of more than 238,000km2 (Goodwin, 1991). To the east, the DC is bounded by the Eastern Ghats Mobile Belt (EGMB) and to the south by the charnockites and khondalites of the SGT. The NE margin is bordered by the

Godavari Rift, and the Narmada–Son lineament marks the northern boundary (Rogers,

1985). The DC is truncated on the west due to the earlier separation of Madagascar from India during the breakup of Gondwanaland (Agarwal et al., 1992) and a small piece of the Western Dharwar craton may now occupy north-central Madagascar

(Tucker et al., 1999). The northwest segment of the Dharwar craton is blanketed by the extensive basaltic flows of the Deccan Traps (Cretaceous–Eocene).

The Dharwar craton is divided into western and eastern domains on the basis of lithological variations, differences in volcano-sedimentary facies, magmatism and metamorphic characteristics (Ramakrishnan and Vaidyanadhan, 2008; Ramakrishnan,

1994; Peucat et al., 1993; Radhakrishna and Vaidyanadhan, 1997). The N–S trending linear outcrop of 2.55–2.51 Ga Closepet granite (Friend and Nutman, 1991) marks the division between the Eastern and Western Dharwar cratons (Ramakrishnan and

Vaidyanadhan, 2008; Naqvi and Rogers, 1987). The Western Dharwar craton (WDC) is dominated by tonalite–trondhjemite Peninsular Gneisses (3.4–2.7 Ga; Jayananda et al.,

2000, 2006) and supracrustals/greenstone schist belts dating to (3.3–2.6 Ga; Anil

Kumaret al., 1996; Trendall et al., 1997a,b;Nutmanet al., 1996). The Eastern Dharwar craton (EDC), on the other hand, is predominantly composed of 3.0–2.55 Ga granites and gneisses, with subordinate linear greenstone belts of limited dimensions

(Balakrishnan et al., 1990; Vasudev et al., 2000; Jayananda et al., 2000; Chadwick et al., 2000; Chardon et al., 2002; Meert et al., 2010).

Meso-Neoproterozoic rocks of igneous and sedimentary origin are adjacent to the

EDC in the large crescent shaped Cuddapah basin (Fig. 1) overlying the Peninsular

Gneiss (Pichamuthu, 1967; Nagaraja Rao et al., 1987). The Dharwar craton experienced widespread mafic magmatism during the Proterozoic along with the intrusion of Meso-Neoproterozoic diamondiferous kimberlites and lamproites (Rao and

Pupper, 1996; Murthy and Dayal, 2001; Chalapathi Rao et al., 2004). In addition to the kimberlites and lamproites, there are other mafic intrusives including metadolerites/metanorites, tholeiitic and alkali-olivine basaltic dykes forming dense E–W and NNW to NW trending swarms crosscutting the greenstone and gneissic basement

of the Dharwar craton (Murthy et al., 1987; Kumar and Bhalla, 1983; Radhakrishna and

Joseph, 1996; Chalapathi Rao et al., 2005). The E–W trending mafic dykes are well developed in the eastern portion of the craton around the Cuddapah basin. The ENE trending dykes form the densest cluster along the southern margin of the basin and these can be traced further west, toward the Closepet granite batholith as widespread albeit impersistent bodies (Murthy, 1995). Radiometric data on these doleritic dykes suggest at least three major episodes of dyke emplacement in the region that occurred at 1.9–1.7, 1.4–1.3 and 1.2–1.0 Ga (Murthy et al., 1987; Padmakumari and Dayal, 1987;

Mallikarjuna Rao et al., 1995; Chatterji and Bhattacharji, 2001). A high precision baddeleyite age of 1885.4±3.1Ma for the Pullivendla mafic sill from the Cuddapah basin has been interpreted to coincide with the widespread ~1.9 Ga basaltic magmatism occurred in the widely separated Large Igneous Provinces (French et al., 2008). The whole rock Rb–Sr isochron age of 2370±230Ma (Ikramuddin and Stuber, 1976) and a

Sm–Nd isochron age of 2454±100Ma (Zachariah et al., 1995) for other easterly trending dykes are further refined by a highly precise U–Pb baddeleyite age of 2367±1Ma (Halls et al., 2007). These age data suggest that most of the easterly dykes in the vicinity of

Cuddapah are Paleoproterozoic in age and intruded during two major magmatic episodes centered at ~2.1 and ~2.4 Ga.

Methods

Paleomagnetic Methods

All paleomagnetic samples were collected using a water-cooled portable drill. The samples were oriented using both sun and magnetic compass and readings were corrected for local magnetic declination and deviations. Subsequently they were cut into cylindrical specimens, and measurements were made on either a Molspin® spinner

magnetometer or a 2G 77R Cryogenic magnetometer at the University of Florida.

Samples were stepwise demagnetized by using either thermal or alternating field (AF) methods. Samples with a very high initial NRM were treated in liquid nitrogen baths prior to thermal or AF treatment to remove more viscous multidomain magnetism.

Linear segments of the demagnetization trajectories were analyzed via principal component analysis (Kirschvink, 1980) using the IAPD software (Torsvik et al., 2000).

Representative sample fragments from each site were ground into a fine powder and analyzed on a KLY-3S susceptibility bridge with a CS-3 heating unit in order to characterize the magnetic carriers in the samples.

Geochronologic Methods

We analyzed two alkaline to doleritic dyke samples from Anantapur district,

Andhra Pradesh and two samples of granitic material from the Barmer district in

Rajasthan for their U–Pb isotopic signatures. Using standard gravity and magnetic separation techniques, zircon grains were concentrated from pulverized samples in various laboratories at the University of Florida. The samples were first crushed, then disk milled and sieved to <80 µm (dolerites) or >80µm (granites) grain size. The fractions were then rinsed, followed by water table treatment with slow sample feed rates. This was followed by heavy liquid mineral separation with multiple agitation periods to reduce the number of entrapped grains in the lower density fraction. Finally, the sample is repeatedly passed on a Frantz Isodynamic magnetic separator up to a current of 1.0A (2–4◦ tilt). Approximately 20–30 fresh looking (clear), euhedral to nearly anhedral zircon grains were handpicked from the two samples of the same site labeled

I595 (dolerite) and 10 subhedral to euhedral zircon grains from the Barmer sample under an optical microscope to ensure the selection of only the clearest grains and

fractions of grains. Further hand-picking of the grains reduced the number to only four to five good grains from the Anantapur dolerites and eight grains from the Barmer granitoids. The zircons were then mounted in resin and then polished to expose median sections. Further sonication and cleaning of the plugs in nitric acid (HNO3) helped to remove any common-Pb surface contamination.

U–Pb isotopic analyses were conducted at the Department of Geological Sciences

(204Hg), 206Pb and 207Pb signals on the ion-counting detectors and 235Uand 238U on the Faraday detectors (see Simonetti et al., 2005). Mounted zircon grains were laser ablated using a New Wave 213 nm ultraviolet beam. During U–Pb analyses, the sample was decrepitated in a He stream and then mixed with Ar-gas for induction into the mass spectrometer. Background measurements were performed before each analysis for blank correction and contributions from 204Hg. Each sample was ablated for ~30 s in an effort to minimize pit depth and fractionation. Data calibration and drift corrections were conducted using the FC-1 Duluth Gabbro zircon standard. Data reduction and correction were conducted using a combination of in-house software and Isoplot

(Ludwig, 1999). Additional details can be found in Mueller et al. (2008).

Paleomagnetic and Geochronologic Results

Gwalior Traps

The Gwalior traps of the Bundelkhand craton are exposed near the city of Gwalior

(Fig. 3) and are assigned to the Morar Subgroup (Gwalior Group). The traps are found at two levels within the Morar Subgroup and are separated by sedimentary rocks. The

lower part of the traps consists of two or more eruptions and the upper zone of the traps consists of a 160-m-thick sequence of multiple flows. The rocks have experienced little deformation and have a low 3–5º northerly dip. The Gwalior traps are correlated with the

Bijawar traps (Athavale et al., 1963; Chakrabarti et al., 2004). Crawford and Compston

(1969) reported a Rb–Sr age of 1798±120Ma (using the revised 87Rb decay constant of

1.42×10−11 year−1) for the upper traps. Mafic sills in the Gwalior basin have Rb and K–

Ar ages of 1775–1790Ma (Ramakrishnan and Vaidyanadhan, 2008). While none of these ages are robust estimates, we also note that some preliminary paleomagnetic data from 1883Ma dykes (dated by French et al., 2008) indicate slightly higher paleolatitudes and suggest that the northern cratons of India may have undergone slow rotational and latitudinal motion between 1883 and 1780 Ma.

Paleomagnetic Results

We resampled the Gwalior traps in the Bundelkhand craton. Previous studies on the Gwalior traps were based on a limited number of samples from the upper traps in the Gwalior Fort region (Athavale et al., 1963; Klootwijk, 1974; McElhinny et al., 1978).

Both Klootwijk (1974) and McElhinny et al. (1978) described a shallow low-coercivity component and a slightly steeper high-coercivity component directed toward the east and down. These authors used blanket demagnetization rather than vector analysis.

We sampled four new sites of the trap volcanics (see Table 1) and applied stepwise demagnetization techniques. One of these sites (41) was a recently opened quarry that contained 3 distinct flow horizons in the lower traps (A–C). Fig. 4 shows typical demagnetization results from the Gwalior volcanic rocks. The mean results from all studies are shown in Fig. 5. Our results largely confirm the previous studies although our directions are somewhat shallower than the high-coercivity components reported by

Klootwijk (1974) and McElhinny et al. (1978). The Gwalior traps are all of a single polarity and our combined tilt corrected mean result has a declination = 73.9º and an inclination of +4.4º (k = 22, α95=11.2º). The paleomagnetic pole was calculated using a site location of 26ºN, 78ºE and is located at 15.4ºN, 173.2ºE (dp = 5.6º, dm= 11.2º; see

Table 1).

Rock magnetic studies were also performed to constrain the magnetic mineralogy of the dyke samples. Most of the sites show unblocking temperatures ranging between

550 and 590 ºC (Fig. 4a–d) indicative of titano-magnetite. A sharp drop in susceptibility occurred in most of the samples during heating at temperatures between 550 and

580ºC; suggests the presence of magnetite (Fig. 4e and f). Unfortunately, exposures of the Gwalior volcanics are limited and there are no field tests that can be conducted to ascertain whether or not the remanence is primary, so the results should be viewed with caution. In addition, the age cited above by Crawford and Compston (1969) has a large error (120 Ma). Our best estimate for the age of the Gwalior pole is thus ~1.8 Ga.

Ananatapur Dykes Region

The Anantapur district lies to the west of the Cuddapah basin (Fig. 6a and b).

Granitoids, gniesses and schists of Archaean to Paleoproterozoic age (2.9–2.1 Ga) constitute the basement lithogies in the area. The basement is intruded by a number of dyke swarms of different ages (1.9–1.7 Ga, 1.5–1.3 Ga and 1.2–1.0 Ga) and directions

(NE–SW, ENE–WSW or NW–SE). Some of these dykes are over 150m wide and tens of kilometers in length. The larger dykes can be seen as persistent linear ridges, surrounding the Paleo-Mesoproterozoic intracratonic Cuddapah basin (Karunakaran,

1971; Halls, 1982; Drury, 1984; Murthy et al., 1987; Poornachandra Rao, 2005). Most of these dykes abruptly terminate at the boundary of the Cuddapah basin and are thus

thought to be older than the basin (Kumar and Bhalla, 1983; Balakrishna et al., 1979;

Bhattacharji and Singh, 1981). A close scrutiny of the field, petrographic, geochemical and radiometric data on these dykes suggest three major phases of dyke emplacement

(all poorly dated) between Paleo-Neoproterozoic (1900–1700, 1500–1300 and 1200–

1000 Ma) around the Cuddapah basin (Chalapathi Rao et al., 2005). The majority of these dyke swarms are dolerite and gabbro along with a few intrusions of peridotite, syenite and granophyres. The dykes are fine to coarse-grained with subophitic to ophitic or granular textures and tholeiitic to alkaline in compositions (Balakrishna et al., 1979;

Kumar and Bhalla, 1983; Halls, 1982; Murthy, 1987; Murthy et al., 1987; Chalapathi Rao et al., 2005). The Anantapur gneissic basement is also intruded by a diamondiferous kimberlitic and lamproitic cluster at ~1100Ma (Chalapathi Rao et al., 2004). The older

(~1900 Ma) of the igneous intrusions are thought to coincide with the thermal events responsible for the initiation and development of intracratonic Cuddapah basin of the

Peninsular India. We sampled the E–W to ENE trending granular to coarse grained, alkaline to tholeiite dykes from the southwestern region of Anantapur for zircon dating

(site I595) and paleomagnetic studies (site I589, I594 and I5103; Fig. 6). Kumar and

Bhalla (1983) also conducted a paleomagnetic study on ENE–WSW and NE–SW trending doleritic dykes in the region.

Geochronologic Results

U–Pb ages from the zircon were determined for the E–W trending alkaline dolerite dyke sample I595 (Fig. 6). Only one of the two samples yielded two well-faceted zircon and several fragments/tips of zircon (Fig. 7). Six laser spots from the two euhedral zircons (sample I595) yielded a concordant U–Pb age of 1027.2±13Ma (2σ; MSWD=

5.0; Fig. 7a). One of the six spots was slightly more discordant (3%) than the other five

(all less than 1% discordant) and an age calculated from these five spots yielded a more precise age of 1025.6±3.8Ma (2σ; MSWD= 1.7; Fig. 7b; Table 2).

Paleomagnetic Results

The NRM of most of the Anantapur dykes formed well grouped clusters and yielded fairly consistent paleomagnetic directions that further improved with cleaning.

Our sites yielded intermediate to shallow inclination that are similar to those obtained by previous workers from the area (Kumar and Bhalla, 1983; Rao, 2005; Table 3). Dyke

I595 that yielded a concordant U–Pb age of 1025.6±3.8Ma also yielded a mean declination = 76.9º and inclination =−50.6º (k = 19.23, α95=12º; N=9; Fig. 8).

Previous paleomagnetic, geochemical and petrologic characteristics studies indicated at least three phases of dyke emplacement in the region (Murthy, 1987;

Kumar and Bhalla, 1983; Chalapathi Rao et al., 1996; Poornachandra Rao, 2005). The combined mean direction of site I595 and dyke (ii) of Kumar and Bhalla (1983) is intermediate and up to the east with a declination = 77º and inclination =−66º (k = 35,

α95=15º; N= 4). Paleomagnetic results on the E–W to N–E trending dykes reported by

Poornachandra Rao (2005) from the vicinity of Anantapur district also yielded consistent intermediate up directions to the east (Fig. 9, Table 3). The overall mean from eight of these sites including ours (I595) gave a VGP at 10ºN and 211ºE with a mean declination = 65º and inclination =−57º (k = 31, α95=10º; Fig. 8, Table 3). The older set of E–W trending dykes yielded an overall mean VGP at 28ºN and 176ºE with a mean declination = 61º and inclination =−0.5º (k = 33, α95=14º; Fig. 8, Table 3).

Thermal demagnetization on our samples indicated unblocking temperatures between 580 and 590 ºC, consistent with low-Ti magnetite as the main carrier of magnetization (Fig. 9a–d). This was confirmed in the Curie temperature runs (Fig. 9e

and f) that show a sharp loss in magnetic susceptibility at temperature ranges of 570–

590 ºC that is characteristic of magnetite. In the absence of any robust field tests to confirm a primary magnetization, we note that the results presented in this paper should be considered preliminary.

Malani Igneous Suite Basement and Early Plutonism

The Malani Igneous Suite (MIS) is exposed over a large area in NW India (Fig.

10); however, most of the region is covered by alluvial sands and the outcrop density is very low. The MIS felsic volcanics and granites are seen as isolated tors, ridges and hills. The mutual relationships between different phases of MIS magmatism are unclear due to the cover and not much is known about the ‗basement‘ for the MIS.

Geochronologic results on the rhyolitic phase of Malani activity is taken from a 771±5Ma

U–Pb age (Gregory et al., 2009). The duration of magmatism is unclear though Gregory et al. (2009) suggested that it lasted from ~771 to 751Ma on the basis of correlation

(and U–Pb ages) with igneous activity in the Seychelles (Torsvik et al., 2001b). Van

Lente et al. (2009) suggested a connection between the coeval Sindreth and Punagarh felsic volcanism of southwestern Aravalli craton and MIS on the basis of the U–Pb isotopic ages of 767–761Ma reported from the Sindreth rhyolites. The closing age of

Malani magmatism is less clear. Granitic bodies such as the Jalore and Siwana are considered to be a part of the MIS on the basis of reported Rb–Sr ages (727–698 Ma;

Rathore et al., 1996, 1999), geological and geochemical considerations (Eby and

Kochhar, 1990; Pandit and Amar Deep, 1997). Laul and Balakrishnan (2007) reported a

Sm–Nd isochron age of 813±13Ma from the Siwana granite, which is approximately 90 million years older than the previously reported Rb–Sr whole rock ages (see Dhar et al.,

1996; Rathore et al., 1996). Recently, Just et al. (submitted for publication) have argued

for a time interval of ~50 million years between Erinpura Granite (pre-Malani) and the

MIS, on the basis of monazite ages of Erinpura granite.

The basement rocks for MIS in the southwestern part are represented by ground level outcrops ~50km west of Barmer town (Fig. 10). These rocks can be described in terms of three temporal assemblages, trondhjemitic gneiss, migmatized diorite and granodiorite (Pandit et al., 1999). The oldest unit is a well-foliated trondhjemitic gneiss

(quartz, biotite, oligoclase–andesine and subordinate K-feldspar). The trondhjemitic gneiss is characterized by high silica (>73%), moderate alumina (13.6–14.4%), low to moderate MgO (0.5–0.74%) and CaO (1.54–2.23%) and a high Na2O/K2O ratio (1.76–

2.23). The migmatized diorites are exposed only within depressions and streambeds and include a fine-grained amphibole rich mafic component and a plagioclase-rich relatively coarse grained neosome.

The youngest assemblage of the supposed pre-Malani basement, the Harsani granodiorite, is exposed as minor NE-trending knolls to the east of village Harsani. The

Harsani granodiorite is a medium grained and crudely foliated rock with predominant plagioclase (oligoclase), quartz, perthite and hornblende. The granodiorites are geochemically distinct from the trondhjemitic gneisses in terms of relatively lower silica

(61.24–67.4%), moderate to high alumina (15.42–15.84%), higher MgO (1.26 to 1.56%) and CaO (2.75–3.06%) and lower Na2O/K2O ratios (1.16–1.81). However, both are clearly distinct from the undeformed, highly siliceous, alkali-rich, CaO- and MgO-poor

Malani rhyolites (including felsic tuffs) and granites.

The trondhjemitic gneiss is cut by Malani rhyolite dykes. Although no direct relationship between trondhjemitic gneisses and granodiorites could be observed due to

extensive sand cover, some indirect field relations have been utilized by Pandit et al.

(1999) to assign an older Archean age to the trondhjemitic gneiss and a pre-Malani age to the granodiorite. We analyzed zircons from the Harsani granodiorite for U–Pb geochronology to further constrain the lower age limit of MIS activity.

New geochronologic results: A total of 10 zircon grains/pieces were recovered from the Harsani granodiorite (Fig. 11). Only one of the grains was euhedral and well- faceted and we were able to obtain two laser spots on this zircon. These two spots yielded a concordant age of 827.0±8.8Ma (2σ; 0.96 probability of concordance). Five other spots on three grains (tips) yielded a concordant age of 786.4±5.6Ma (2σ) with a relatively large MSWD= 9.6. In addition, four other grains yielded highly discordant ages with a lower intercept of 613±16Ma and an upper intercept age of 2060 + 37/−38Ma

(MSWD= 1.6; Table 4).

Our interpretation of these data is that the 827.0±9 Ma age closely approximates the time of intrusion of the granodiorite and the 786.4±5.6Ma may relate to a disturbance of the U–Pb system at the onset of Malani volcanism. This is consistent with the reported ages of the coeval intermediate granitoid rocks (tonalites) forming the basement of Punagarh volcanics (Van Lente et al., 2009) that are correlated with the

Erinpura granites (Choudhary et al., 1984; Deb et al., 2001; Laul and Balakrishnan,

2007) and preceded the out pouring of Malani rhyolites at around 771Ma (Gregory et al., 2009). The discordant zircons reflect inheritance of an older basement component of

Paleoproterozoic age. We place less significance on the lower intercept age but do note that others have suggested an Ediacaran-age thermal disturbance in the region

(Rathore et al., 1999). Thus, it is thought that the main phase of rhyolitic magmatism in

the Malani province occurred over a relatively short interval (~20–30 Ma)and therefore the paleomagnetic pole derived from the Malani sequence should show very little latitudinal scatter (assuming normal plate motion speeds). Indeed, Gregory et al. (2009) show that data from mafic dykes are identical to that observed in the rhyolites.

Additional paleomagnetic studies on the older phases of magmatism may help constrain motions between ~800 and 771 Ma.

Discussion

Paleoproterozoic Results (2.5Ga–1.6Ga)

A combined result for the poorly dated Paleoproterozoic Gwalior traps yields a paleomagnetic pole for India at 15.4ºN, 173.2ºE (dp = 5.6º, dm= 11.2º) using the reported age of ~1.8 Ga (Crawford and Compston, 1969; Ramakrishnan and

Vaidyanadhan, 2008). Our reconstruction using the new Gwalior paleomagnetic pole places Indian subcontinent at equatorial position with a palaeolatitude of 2.2±5.5ºN. The other Paleoproterozoic poles used for our ~1.8 Ga reconstruction come from Laurentia,

Baltica, Amazonia, Australia and Kalahari (Fig. 12). Our reconstruction follows the paleogeographic configuration of Laurentia, Baltica and Amazonia at 1.83 Ga in the proposed model of ―Hudsonland‖ (Pesonen et al., 2003; Williams et al., 1991). The intermediate latitudinal position of Laurentia in our reconstruction is defined by a mean

VGP from the Sparrow dykes, Churchill Province and the Coronation Geosyncline

(Pesonen et al., 2003). Meert (2002) places Laurentia at polar latitudes in his ―closest approach‖ paleomagnetic model for the Paleo-Mesoproterozoic supercontinent

―Columbia‖ at ~1.7–1.5 Ga based on the paleomagnetic data from the Trans Hudson

Orogenic belt (THO) of the Canadian Shield (Symons and MacKay, 1999). The absence of any rigorous field tests from THO makes the data less reliable. Our intermediate

latitude position for Laurentia at ~1.8 Ga is in accord with the paleomagnetically consistent paleogeography for Columbia proposed by Bispo-Santos et al. (2008) using the ~1.8 Ga Dubawnt Group paleomagnetic pole of the Churchill Province. Although we do note that other interpretations are also possible for the position of Laurentia within

Columbia configuration for this time period, in the absence of more robust data from the

THO, we prefer the intermediate latitude position for Laurentia. The position of Baltica is constrained to be close to Laurentia by the well-defined paleomagnetic pole from the

Haukivesi lamprophyres of Finland (Neuvonen et al., 1981). The Amazonian craton was located in the southern hemisphere within the Columbia supercontinent adjacent to

Baltica and its position is defined by the mean VGP given in Pesonen et al. (2003). The position is also supported by a later configuration of Amazonia using the results obtained from the 1790Ma Colider Suite (CS Pole) by Bispo- Santos et al. (2008). A

Paleoproterozoic (~1.8–1.85 Ga) connection between India, North China and Laurentia was proposed on the basis of the presence of contiguous unmetamorphosed and undeformed radiating mafic dyke swarms within Southern peninsular India, North China and Canadian Shield of North America (Hou et al., 2008; French et al., 2008). In their classic Columbia model, Zhao et al. (2004, 2005, 2006) also placed India and North

China adjacent to each other, based on geologic similarities between the two cratons.

The presence of similar aged orogenic belts in India and North China further reinforces the argument of their contiguity although, detailed geologic arguments favoring their proximity are lacking during Mesoproterozoic (Rogers, 1996; Rogers and Santosh,

2002; Zhao et al., 2002, 2004, 2006; Santosh et al., 2003, 2007, 2009a,b).

We do not include North China in our ~1.8 Ga reconstruction due to the paucity of paleomagnetic data from North China at ~1.8–1.9 Ga. Lastly, we do note that our

Gwalior pole for India lacks robust age determination and while making any paleogeographic reconstruction using this pole, our interpretation relies upon the U–Pb age of 1798±120 (Crawford and Compston, 1969). It is interesting to note that the VGPs from the older (~1.9 Ga) mafic dykes of the southern Bastar craton and nearby

Cuddapah basin from the adjacent Dharwar craton, India (French et al., 2008; Clark,

1983; and our own work) and our ~1.8 Ga Gwalior traps do show a similar near- equatorial position for Indian subcontinent although the declinations differ by ~60º.

The paleoconfiguration of Australia is not well defined during Paleoproterozoic.

The paleopole position at 1.82 Ga for the Plum Tree Creek volcanics in the Pine Creek

Orogen of northwestern Australia used in our reconstruction, falls at 29ºS, 195ºE and yields a low to intermediate paleolatitudinal position for northwestern Australia block

(Idnurm and Giddings, 1988; Kruse et al., 1994; Idnurm, 2004). The low latitudinal position for Australia during Paleoproterozoic is also supported by the paleomagnetic data obtained by Schmidt and Williams (2008) from the Elgee siltstone and Pentecost sandstone of Kimberley Group. Pesonen et al. (2003) speculated that Australia was a part of ―Hudsonland‖ (Columbia) at 1.77 Ga. Paleomagnetic data from the Kalahari craton is sparse for the Paleoproterozoic like many other continental blocks. At ~1.8 Ga, the Kalahari craton was undergoing prolonged crustal accretion to form the proto-

Kalahari craton by 1750Ma according to the model by Jacobs et al. (2008).We use the only available 1.8–1.85 Ga Sebanga Poort dyke pole of Jones et al. (1976) in our

reconstruction that position proto-Kalahari at intermediate to high latitudes. However, we are aware that this pole is not paleomagnetically well constrained.

Mesoproterozoic Results (1.6Ga-1.1Ga)

The Mesoproterozoic period witnessed the disintegration of the proposed

Paleoproterozoic supercontinent Columbia and the subsequent amalgamation of the

Early Neoproterozoic landmass Rodinia (Moores, 1991; Dalziel, 1991; Hoffman, 1991;

Rino et al., 2008). The position of the Indian subcontinent in Rodinia configuration is debated. In archetypal Rodinia, India was attached to Australia and East Antarctica, similar to its Gondwanan configuration (Dalziel, 1991; Hoffman, 1991; Moores, 1991; Li et al., 1996; Torsvik et al., 1996; Weil et al., 1998). Later workers challenged this configuration on paleomagnetic and geological grounds and suggest that India was never a part of Rodinia supercontinent (Fitzsimons, 2000; Meert, 2003; Powell and

Pisarevsky, 2002; Torsvik et al., 2001a,b). In this paper, we review our updated ~1.2 Ga paleomagnetic pole from the Harohalli alkaline dyke swarm at 24.9ºS, 258ºE (α95=15º;

Pradhan et al., 2008) in an attempt to better constrain the position of India during this period. Pradhan et al. (2008) reported concordant U–Pb age of 1192±10Ma that is far older than all previously published ages on the Harohalli and supposedly correlative dykes (Ikramuddin and Stuber, 1976; Anil Kumar et al., 1989; Devaraju et al., 1995).

Discordant ages ranging from 580 to 1139Ma was also calculated the zircon fragments of sample I5143 (Pradhan et al., 2008). Slightly discordant zircons form a tight grouping near 580–600Ma region of concordia (Pradhan et al., 2008).

This new U–Pb radiometric age necessitates an almost 400Ma revision in age of the Harohalli alkaline dykes. Fig. 13 shows 1.2 Ga reconstruction using our new

Harohalli pole for India and other key paleomagnetic poles from Laurentia, Amazonia,

North China and Australia. Laurentia is positioned at intermediate latitudes following the mean paleomagnetic pole calculated at ~1156Ma (Pesonen et al., 2003). Amazonia is juxtaposed with the Llano segment of Laurentia‘s Grenville orogen using the recently dated ~1200Ma Nova Floresta mafic sills pole reported by Tohver et al., 2002 (also see

Cordani et al., 2009; Elming et al., 2009). Tohver et al. (2004a,b, 2005a,b, 2006) suggested a common Rodinian paleogeography for Amazonia and southern Laurentia at least since 1.2 Ga with Amazonian craton undergoing lengthy sinistral strike slip motion along the Grenville margin (i.e. present-day eastern North America). The North

China (NC) craton occupies equatorial latitudes in our reconstruction using the mean

BCP (Baicaoping Formation) pole at 1.2 Ga (Zhang et al., 2006). Within their APWP matching of NC with Laurentia during 1250–750 Ma, the 1200Ma BCP pole falls between 1270 and 1100Ma Laurentian poles (Zhang et al., 2006). However, these authors do suggest that the distantly located paleopoles from North China and Laurentia for 1200–1400Ma periods imply non-coherence between them at 1.2 Ga. The combined paleomagnetic data from the mafic and metamorphic intrusive rocks of the Albany-

Fraser orogen (Mt. Barren Group, Bremer Bay and Whalebone Plateau and Fraser dyke) place Australia at higher latitudes in ca 1.2 Ga configuration (Pisarevsky et al.,

2003). An interesting observation from the paleomagnetic data of the ~1.2 Ga Harohalli dykes of the Dharwar craton is that it places India at polar latitudes adjacent to Australia although not in a traditional ‗East Gondwana‘ configuration.

~1.1-1.0 Ga reconstructions: The end of Mesoproterozoic and beginning of the

Neoproterozoic era is recognized as the significant time period for the initiation of the breakup of the proposed Rodinia supercontinent followed by the assembly of

Gondwana. We review our previously reported paleomagnetic and geochronologic results from the Upper Vindhyan sediments (Malone et al., 2008) and the Majhgawan kimberlite (Gregory et al., 2006) of Central India along with the newly reported geochronologic and paleomagnetic data from the Anantapur mafic dyke swarms, and the possible ~580Ma paleomagnetic overprint of the Harohalli dykes of the Dharwar craton of the peninsular India to evaluate the paleogeographic configuration of Indian subcontinent during Neoproterozoic (~1.0 Ga to ~580 Ma).

The Vindhyan Basin in central Peninsular India is a large intracratonic sedimentary basin that provides target area to conduct the necessary paleomagnetic and geochronologic studies for Meso-Neoproetrozoic. In our efforts to constrain the depositional history of the basin and provide key paleomagnetic poles from India for the

Neoproterozoic, we considered analyzed contemporaneous geochronologic and paleomagnetic results from the Upper Vindhyan sediments and the Majhgawan kimberlite that intrudes the lower part of the Upper Vindhyan sedimentary rocks

(Gregory et al., 2006; Malone et al., 2008).

Malone et al. (2008) studied detrital zircon populations from the Bhander and

Rewa Groups in the Vindhyan rocks of Rajasthan sector along with samples from the

Lower Marwar Supergroup in Rajasthan to constrain the age of the Upper Vindhyan.

Malone et al. (2008) noted that the youngest population of zircons from the Upper

Bhander is older than 1000 Ma. This observation, coupled with the similarity in paleomagnetic directions from the Upper Vindhyan and the 1073.5±13.7Ma age of the

Majhgawan kimberlite (Gregory et al., 2006), it is concluded that the Upper Vindhyan sedimentation was completed by ~1000 Ma, a result consistent with recent

geochronologic data from the Chhattisgarh basin to the south (Patranabis-Deb et al.,

2008; Basu et al., 2008). The paleomagnetic data from the Bhander and Rewa Groups of the Upper Vindhyan sequence yield pole positions that overlap with the Majhgawan kimberlitic pole (Gregory et al., 2006).

Our new geochronologic results of the tholeiitic to alkaline Anantapur dykes

(1025.6±3.8 Ma; Fig. 7b) fall into a similar age bracket as those of magmatic events around Cuddapah basin. The overall mean of paleomagnetic directions from eight sites including ours (I595) from the Anantapur region gave a VGP at 10◦N and 211ºE

(α95=10º; Fig. 9, Table 2). Within the archetypal Rodinia configuration at around 1.1–

1.0 Ga, Laurentia formed the core of the supercontinent with East Gondwana components lying at its present-day western margin while Baltica and Amazonia occupy its present-day eastern margin (Moores, 1991; Dalziel, 1991; Hoffman, 1991). However, the lack of reliable paleomagnetic data from the East Gondwana blocks prevents its precise placement within Rodinia. The paleogeographic reconstruction using our newly dated ~1.0 Ga Anantapur dyke pole places India at a paleolatitude of 37.6◦ N (Fig. 14).

The other key paleomagnetic poles used for this configuration come from Laurentia,

Siberia, Baltica, Australia, Congo-Sao Francisco and Kalahari. The occurrence of

Grenvillian age collisional features in NW Baltica, NW Amazonia and NW Congo indicate the final docking of these blocks to Laurentia at 1.05 Ga (Pesonen et al., 2003).

We used the Bangemall sills paleopole for Australia that is close to the AUSMEX fit of Wingate et al. (2002) at 1.05Ma (Fitzsimons, 2002). Although India can be placed in close proximity to Australia due to longitudinal uncertainties, it is clear that the configuration does not favor an East Gondwanan configuration for these two blocks.

The position and orientation of Siberia is considered controversial within Rodinia.

Siberia is placed either adjacent to the northern margin of Laurentia (Hoffman, 1991;

Condie and Rosen, 1994; Rainbird et al., 1998) or within Australia–Siberia–Laurentia fit along the western margin of Laurentia (Sears and Price, 2000).

Neoproterozoic Results (1000-570 Ma)

The end of the Neoproterozoic era witnessed the breakup of the Rodinia supercontinent and amalgamation of Gondwana around 530Ma (Meert, 2001, 2003;

Powell and Pisarevsky, 2002; Meert and Lieberman, 2004; Torsvik et al., 1996). There are several competing models for the assembly of Gondwana that are reviewed in

Meert and Lieberman (2008). East Gondwana is widely assumed to have been a single coherent entity since the end of the Mesoproterozoic (Powell et al., 1993; Dalziel, 1997;

Weil et al., 1998; Yoshida, 1995; Yoshida et al., 2003). The apparent continuity of a

Grenville-age (1100–1000 Ma) metamorphic belt along the Greater India–East

Antarctica–Australia margin has generally been taken as evidence for no substantial movements of these cratonic blocks from their East Gondwana fit since 1.0 Ga. Meert and Van der Voo (1997) and Meert (2003) challenged the assumption of a coherent

East Gondwana and suggested a multiphase amalgamation of the various cratonic blocks through a series of orogenic events spanning the interval from ~750 to ~530 Ma.

Fitzsimons (2000) demonstrated that the proposed hypothesis of a single continuous late Mesoproterozoic orogenic belt along the African, Indian and Australian margins of

Antarctica in Gondwana was incorrect and showed that there are three distinct late

Mesoproterozoic–early Neoproterozoic orogenic belts.

Paleomagnetic data from the coeval Malani Igneous Suite (MIS) of India and

Mundine well dyke swarms of Australia–Mawson continental blocks also indicates that

these were not united until after 750Ma (Torsvik et al., 2001a,b; Malone et al., 2008;

Gregory et al., 2009; Wingate and Giddings, 2000). Our paleomagnetic data from the

Malani dykes (Gregory et al., 2009) place India and the Seychelles at higher latitudes than coeval poles from Australia (Wingate and Giddings, 2000). These three robust paleomagnetic results (Mundine dykes, Malani Igneous Suite and Mahe dykes) argue strongly against a coherent East Gondwana at 750Ma and explain the younger Pan-

African belts between these cratons as the result of a final Ediacaran-age collision.

We also report a paleogeographic reconstruction at ~580 Ma. Paleogeographic reconstructions for this interval of time are highly controversial (Meert et al., 1994;

Torsvik et al., 1996; Kirschvink et al., 1997) in part due to the apparent rapid motion of several cratonic blocks. Whether or not this rapid motion is due to plate motion, true polar wander or inertial interchange true polar wander is beyond the scope of this review; however, discussions regarding each of these mechanisms can be found in

Kirschvink et al. (1997), Torsvik et al. (1998), Evans (1998), Meert (1999) and Meert and Tamrat (2004).

We use the average of our new paleomagnetic pole 76.5ºS, 68.8ºE (α95=15º) obtained from the overprint of the Harohalli dykes (Pradhan et al., 2008; Halls et al.,

2007) and the mean Banganapalli paleomagnetic pole 73.48ºN, 233.63ºE (α95 = 5.2º) from the Kurnool Group (Goutham et al., 2006) to position India. The paleoconfiguration using this new Indian pole and other key paleomagnetic poles from the constituent cratonic nuclei of East Gondwana at 580Ma (Fig. 15) places India at equatorial latitudes. We follow the model of Meert (2003) and place part of East Antarctica (Rayner

Complex and Prince Charles Mtns.) adjacent to the Eastern Ghats margin of India. The

other paleomagnetic poles used in our reconstruction at ~580Ma come from Laurentia and the West Gondwanan blocks including Amazonia, West Africa, Congo-Sao

Fansisco, Rio de-Plata and Kalahari (as in Tohver et al., 2006; Gray et al., 2006;

Cordani et al., 2009).

On the basis of the paleomagnetic data from Australia–Mawson continental blocks, Powell and Pisarevsky (2002) suggest that these were not united until after 750

Ma. India and Western Australia could be in contact as early as 610 Ma, supported by the existence of Darling mobile belt, a sinistral shear that is consistent with an oblique collision of Greater India along the Western Australia margin (Powell et al., 1993). If our new pole is correct (and indeed ~580 Ma), then it is possible that India and Australia were either in contact or very close to one another by this time. Constraints on the positions of the western Gondwana blocks in our reconstruction are not based on paleomagnetic data and so the apparent discordance between the eastern Gondwana blocks and those from western Gondwana is an artifact of our placing the western blocks adjacent to present day eastern Laurentia during the opening of the Iapetus

Ocean (Fig. 15).

Summary

The new results reported in this paper from the Gwalior traps, the Anantapur dykes and the Harsani granodiorite, coupled with the previous studies done by our group emphasize that revisions in the Precambrian paleogeographic history of the

Indian subcontinent are still in the early stages. We have presented some possible paleoreconstructions of India at ~1.8 Ga, 1.2 Ga, 1.0 Ga and 580 Ma. With the exception of the 580Ma reconstruction, support for a traditional East Gondwana configuration with Himalayan margin of India adjacent to western Australia is precluded.

Although most of our newer data need additional support, they are consistent with the model of Gregory et al. (2009) wherein the assembly of ―East Gondwana‖ is coincident with the assembly of greater Gondwana during the Ediacaran–Early Cambrian time.

Acknowledgements: This work was supported by a grant from the US National

Science Foundation to J.G. Meert (EAR04-09101). We wish to thank George Kamenov and Warren Grice for their help with the U–Pb data reduction.

Figure 3-1. Generalized tectonic map of Indian subcontinent: Precambrian Cratons, Mobile belts and Lineaments. AFB = Aravalli Fold Belt, DFB = Delhi Fold Belt, EGMB = Eastern Ghat Mobile Belt, SMB = Satpura Mobile Belt, NSL = Narmada Son Lineament, CIS = Central Indian Suture and BPMP = Bhavani Palghat Mobile Belt (Modified from Vijaya Rao and P.R. Reddy, 2001).

Figure 3-2. Sketch map of the major units in the Aravalli-Bundelkhand craton, NW India (after Naqvi and Rogers, 1987 and Ramakrishnan and Vaidyanadhan, 2008).

Figure 3-3. Sketch map of the Bundelkhand craton showing mafic dyke swarms including the Great Dyke of Mahoba (grey-dashed line; after Malviya et al., 2006).

Figure 3-4. Paleomagnetic and rock magnetic results for the Gwalior traps. (a) Thermal demagnetization behavior of Gwalior volcanic at site 38. (b) Alternating field demagnetization behavior of Gwalior volcanic at site 39. This sample shows a great circle trajectory towards the East and down. (c) Alternating field demagnetization behavior of volcanic rocks at site 40. (d) Thermal demagnetization behavior of Gwalior volcanic at site 41. Curie temperature runs for selected samples in this study (e) Gwalior dyke sample I541 showing a heating Curie temp of 591.7 ºC and a cooling Curie temperature of 589.1 ºC; (f) Gwalior dyke sample I544 with a heating Curie temperature of 588.5 ◦C and a cooling Curie temperature of 585.1 ºC.

Figure 3-5. Stereoplot of site mean directions from our study (white shading) the mean of Athavale (1963, yellow shading) Klootvijk (1974, red shading) and McElhinny et al. (1978, blue shading). The overall mean based on 9 sites yields an in-situ declination of 74.3º and an inclination of 5.8º. Correction for regional tilt (strike: 90, dip: 4) yields a tilt-corrected mean of 73.9º/4.4º.

Figure 3-6. Sketch map of the Eastern Dharwar craton. (a) Archaen assemblages associated with cratonization are shown here. Abbreviations for schist belts are as follows: Sa = Sandur, KKJH = Kolar-Kadiri=Jonnagiri-Hutti, RPSH+ Ramagiri- (Penakacheria-Sirigeri)-Hungund, and VPG = Veligallu-Raichur- Gadwal superbelt. The dotted and dashed line indicates possible location under the basin. Dashed lines represent Paleozoic to more recent sedimentary cover (Modified from Naqvi and Rogers, 1987). Dyke intrusions into the EDC are H&B=Harohalli and Bangalore swarm; A=Anantapur swarm; M=Mahbubnagar swarm; H=Hyderabad swarm. (b) Enlarged sketch Map of the Anantapur region in Andhra Pradesh showing the location of dikes used in this study including the Great dyke of Bukkapatnam (after Murthy, 1995). Geochronologic samples that yielded zircon are I594 and I595 (blue boxes). These sites are also studied for paleomagnetic analyses.

Figure 3-7. U/Pb isotopic analysis for the Anantapur dyke sample I595. (a) U/Pb concordia diagram for the six spots from the concordant zircons from sample I595 yielding an age of 1027.2  13 Ma (2; MSWD=5.0). Photo of zircon grains from plug I595 grains 1-6 (Appendix A). (b) U/Pb concordia diagram for the five spots from the concordant zircons from sample I595 yielding an age of 1025.6  3.8 Ma (2σ; MSWD=1.7). Photo of zircon grains from plug I595 grains 1-5 (Appendix B).

Figure 3-8. Paleomagnetic and rock magnetic results for the Anatapur dyke samples. (a) Thermal demagnetization behavior of Anantapur dyke sample I590-9b. (b) Alternating field demagnetization behavior of Anantapur dyke sample I590-8. (c) Thermal demagnetization behavior of Anantapur dyke sample I595-5A. (d) Alternating field demagnetization behavior of the dyke sample I595-5B. Curie temperature runs for selected samples in this study (e) Anantapur dyke sample I595 showing a heating Curie temp of 561.7 ºC and a cooling Curie temperature of 565.1 ºC; (f) Anantapur dyke sample I5102 with a heating Curie temperature of 569.7 ºC and a cooling Curie temperature of 562.4 ºC. In all the stereoplots open/closed circles represent up(-)/down(+) inclinations.

Figure 3-9. Stereoplot of site mean directions from our study (Anantapur site I595), the mean of Kumar and Bhalla (1983, 1984) and Poornachandra Rao (2005). The best overall mean based on 7 sites yields an in-situ declination of 66º and an inclination of -56º (α = 12; k = 27; shaded yellow).

Figure 3-10. Sketch map showing Precambrian stratigraphic units of the Aravalli Mountain Region and further west in NW India with sampling area of Harsani granodiorite boxed (adapted from GSI publications and other published work).

Figure 3-11. U/Pb isotopic analysis for the Barmer sample Bar05. (a) U/Pb concordia diagram for the five spots from three concordant zircon fragments and tips from sample Bar-05 yielding an age of 786.4  5.6 Ma (2) with a relatively large MSWD=9.6. (b) U/Pb concordia diagram for the two spots from two concordant zircons from sample Bar05 yielding an age of 827.0  8.8 Ma (2; 0.96 probability of concordance). (c) Discordia plot using other four fragmentary zircons from Bar05. The lower intercept age is 613  16 Ma and the upper intercept age is 2060 +37/-38 Ma (MSWD=1.6).

Figure 3-12. Paleomagnetically-based reconstruction at 1.83 Ga after (Pesonen et al., 2003). The Bundelkhand craton is positioned according to paleomagnetic data from the Gwalior volcanics reported in this paper. (Palaeolongitudes are unconstrained). The reconstruction places India at equatorial positions with mid latitudinally located Laurentia within the Columbia Supercontinent.

Figure 3-13. Reconstruction at ~1.2 Ga based on our new age from Harohalli Paleomagnetic pole and data reported in Pesonen et al. (2003) that places India at polar latitudes contiguous to Australia as compared to mid latitudinal positions of Laurentia, Siberia and Amazonia with North China at the equatorial positions (Paleolongitudes are unconstrained).

Figure 3-14. Reconstruction at ~1.0 Ga using our new paleomagnetic data from the ~1.0 Ga dated Anantapur dykes of the Dharwar craton, India. The paleomagnetic poles for other continental blocks are taken from Pesonen et al. (2003). Our reconstruction places India (Anantapur dyke pole) and Australia (Bangemall sill pole) in contiguous positions.

Figure 3-15. Paleogeographic reconstruction based on the combined paleomagnetic poles from the younger discordant 580 Ma Harohalli overprint (Pradhan et al., 2008) and the coeval Banganapalli Quartzite of the Kurnool Group (Goutham et al. 2006). It places India at shallow equatorial latitude in close proximity to a part of Antarctica but separated from Australia via Mawson Sea. The other half of the East Antarctica microcontinent is attached to Australia based on its Elatina dyke paleomagnetic pole. The Cratonic blocks within West Gondwana with Laurentia rifting away from Amazonia and Rio Plata (RP). Kalahari is separated from Congo-SF and Amazonia via Adamastor and Khomas oceans. (modified from Gray et al., 2006).

Table 3-1. Paleomagnetic results from the Gwalior volcanics

Site/Study N Dec Inc k A95 Athavale et al., 1963 07 70.0° 3.0° 18.0° Klootwijk, 1974 23 78.0° 34.0° 369.0 05.0° McElhinny et al., 1978 12 78.5° 25.8° 048.0 06.3° Site 38- This study 13 81.3° 11.7° 089.4 04.4° Site 39-This study 08 73.5° 10.8° 044.0 08.5° Site 40-this study 12 71.8° 0.0° 195.0 03.1° Site 41A-this study 10 71.2° -10.3 052.0 06.7° Site 41B-this study 05 72.5° -19.3 588.0 03.2° Site 41C-this study 06 72.9° -3.8 079.0 07.6° Lower Trap Mean 21 72.2° -11.1 107.2 12.0° Upper Trap Mean 75 75.4° 14.2° 034.9 11.6° Overall Mean 96 74.3° 5.8° 022.0 11.2° Tilt-Corrected Mean 96 73.9° 4.4° 022.0 11.2°

N = No. of samples used; k=kappa precision parameter α95=cone of 95% coincidence about the mean direction, Site 41A, B and C are from the lower traps all other are from the upper traps.

Table 3-2. Anantapur geochronologic results 207Pb/ 1σ 206Pb/ 1σ *207Pb/ 1σ 206Pb/ 207Pb/ 207Pb/ % Grain 1σ 1σ 1σ 206Pb error 238U error 235U error 238U 235U 206Pb Disc RHO I595_1 0.067712 0.0001 0.14783 0.0009 1.40979 0.0820 1002 19 1013 13 1039 7.9 1 0.98 I595_2 0.067349 0.0002 0.14872 0.0008 1.27409 0.1000 1008 19 1013 13 1028 8.4 1 0.98 I595_3 0.066836 0.0001 0.14760 0.0014 1.70992 0.1900 1001 20 1004 14 1012 7.8 0 0.98 I595_4 0.067759 0.0001 0.15586 0.0009 1.45505 0.0230 1052 20 1048 13 1040 7.7 0 0.98 I595_5 0.067216 0.0001 0.15135 0.0008 1.35673 0.0240 1024 19 1023 13 1024 7.4 0 0.98 I595_15 0.068380 0.0002 0.14698 0.0009 1.75198 0.1400 997 19 1015 13 1058 10.0 2 0.97

Table 3-3. Anantapur paleomagnetic results Site N Dec Inc k a95 Plat Plong Reference Younger Dykes (E-W) DT2 5 76 .0° -34 .0 14 .8 16.3° 08°N 189°E G.V.S. Poornachandra Rao, 2005 D8 6 69.0° -48.0 72.4 06.7° 10°N 201°E G.V.S. Poornachandra Rao, 2005 D29 5 50.0° -66.0 19.9 15.7° 13°N 226°E G.V.S. Poornachandra Rao, 2005 D30A 5 51.0° -46.0 30.1 11.4° 25°N 208°E G.V.S. Poornachandra Rao, 2005 DT1 5 54.0° -69.0 27.0 12.1° 08°N 228°E G.V.S. Poornachandra Rao, 2005 D32 5 258.0° 57.0 10.0 20.0° 00°N 027°E G.V.S. Poornachandra Rao, 2005 Kumar and Bhalla, 1983 + This Dyke (ii) + I5145 4 sites 76.7° -66.0 35.0 15.0° 02°S 217°E study Dyke (i) 5 57.0° -69.0 52.0 07.0° 08°N 225°E Kumar and Bhalla, 1983

Older Dykes E-W Dyke (iii) 12 64.0° -7.0 142.0 08.0° 24°N 178°E Kumar and Bhalla, 1983 Dyke (iv) 6 53.0° -8.0 142.0 06.0° 34°N 183°E Kumar and Bhalla, 1983 I594 14 46.0° 5.0 27.9 07.6° 45°N 177°E This Study I5103 7 79.0° 1.3 12.1 18.1° 11°N 169°E This Study I589 8 241.0° -6.4 36.0 09.4° 29°N 171°E This Study

Overall Mean- Younger 8 dykes 65.2° -57.0 31.0 10.0° 10°N 211.4°E This Study Overall Mean-Older 5 dykes 61.0° -00.5 33.0 14.0° 28°N 176.0°E This Study

Dec=declination; Inc=Inclination; k=kappa precision para α95=cone of 95% coincidence about the mean Direction, Plat=Pole latitude; Plong=Pole Longitude

Table 3-4. Barmer geochronologic results % 207Pb/206P 1σ 206Pb/238 1σ *207Pb/235 1σ 206Pb/238 1 207Pb/235 1 207Pb/206P 1 Dis Analysis b error U error U error U (Ma) σ U (Ma) σ b (Ma) σ c Concordant (830 Ma) Bar05_14 0.104559 0.0004 0.796001 0.007 0.053964 0.0003 830 8 827 10 819 13 0 Bar05_17 0.103866 0.0005 0.758269 0.021 0.054239 0.0001 824 9 826 9 830 8 0

Concordant (786 Ma) Bar05_2 0.098418 0.0007 0.723274 0.012 0.053568 0.0001 784 9 789 10 804 8 1 Bar05_3 0.098691 0.0006 0.763438 0.011 0.053557 0.0001 786 9 790 9 803 9 1 Bar05_4 0.097705 0.0007 0.718155 0.009 0.053207 0.0001 778 9 781 9 789 8 0 Bar05_11 0.098071 0.0005 0.844853 0.034 0.053664 0.0002 781 8 788 9 807 11 1 Bar05_19 0.096751 0.0005 0.713415 0.011 0.053189 0.0001 771 8 776 9 789 8 1

Discordant grains Bar05_1 0.083196 0.0004 0.638326 0.011 0.055020 0.0001 669 7 714 8 859 8 6 Bar05_8 0.206235 0.0049 2.68271 0.095 0.093854 0.0010 1545 39 1696 47 1887 20 9 Bar05_9 0.203406 0.0033 2.66302 0.076 0.094257 0.0010 1526 28 1688 37 1895 20 10 Bar05_10 0.241246 0.0052 3.36606 0.120 0.101459 0.0014 1774 42 1893 52 2027 25 6

CHAPTER-4 PALEOMAGNETIC AND GEOCHRONOLOGICAL STUDIES OF THE MAFIC DYKE SWARMS OF BUNDELKHAND CRATON, CENTRAL INDIA: IMPLICATIONS FOR THE TECTONIC EVOLUTION AND PALEOGEOGRAPHIC RECONSTRUCTIONS

Introductory Remarks

The Indian subcontinent holds a key position when attempting to unravel the intricacies of Precambrian paleogeography, such as the Paleoproterozoic Columbia supercontinent (Rogers, 1996; Rogers and Santosh, 2002; Meert, 2002; Santosh et al.,

2003; Zhao et al., 2004a; b), the Meso-Neoproterozoic Rodinia supercontinent

(McMenamin and McMenamin, 1990; Meert and Torsvik, 2003; Meert and Powell, 2001;

Li et al., 2008), and the Ediacaran-Cambrian Gondwana supercontinent (Meert, 2003;

Meert and Van der Voo, 1996; Burke and Dewey, 1972; Pisarevsky et al.,2008). India offers target rocks that are both accessible and of appropriate age for conducting paleomagnetic and geochronologic investigations (Meert, 2003; Powell and Pisarevsky,

2002; Meert and Powell, 2001; Pesonen et al., 2003). These target rocks include

Precambrian mafic dykes and dyke swarms that intrude the Archean-Paleoproterozoic cratonic nuclei, as well as volcanic and sedimentary successions exposed in the

Dharwar and Aravalli protocontinents of the Indian peninsular shield (figure 1).

The correlation of mafic dyke swarms in terms of their distribution, isotopic age, geochemistry and paleomagnetism is critical in order to fully evaluate Precambrian plate reconstructions and possible configuration of supercontinents (Halls, 1987; Van der Voo and Meert, 1991; Halls, 1995; Mertanen et al., 1996; Park et al., 1995; Bleeker and

Ernst, 2006; Ernst and Srivastava, 2008; Piispa et al., 2011). The Indian peninsular shield is traversed by numerous Precambrian mafic dyke swarms (Drury, 1984; Murthy,

1987, 1995; Ramachandra et al., 1995; French et al, 2008, 2010; Meert et al., 2010;

Pati et al., 1999; 2008; Pradhan et al., 2008; Ernst et al., 2008; Srivastava and Gautam,

2008; Srivastava et al., 2008). These dykes intrude the granite-greenstone terranes of the major Indian cratonic nuclei, namely Dharwar in the south, Bastar in the east central, Singhbhum in the east and the Aravalli and Bundelkhand in the north-west and the north (Halls, 1982; figure 1). The dykes are the focus of considerable scientific attention aimed at defining their geochemical and geophysical characteristics, geochronology, paleomagnetism and tectonic controls on their emplacement (Devaraju,

1995, Radhakrishna and Piper, 1999, Srivastava et al., 2008; Halls et al, 2007; French et al., 2008, 2010; Pradhan et al., 2008, 2010; Pati et al., 1999; 2008; Piispa et al.,

2011; Ratre et al., 2010; Ernst et al., 2008; Ernst and Bleeker, 2010; Meert et al., in press). Paleomagnetic and geochronologic studies on these Precambrian mafic dyke swarms as well as the volcanic and sedimentary successions of the Gwalior, Cuddapah and Vindhyan basins, provide insight into India‘s changing paleogeographic position during the critical Paleo- to Mesoproterozoic interval and may, in some circumstances, allow for further age constraints to be placed on the poorly dated sedimentary basins of

India (Pradhan et al., 2010, 2008; Gregory et al., 2006; Malone et al., 2008; Meert et al., in press; Meert et al., 2010; French et al., 2008; French, 2007; Halls et al., 2007; Ratre et al., 2010; French and Heaman, 2010; Piispa et al., 2011).

This study focuses on the Precambrian mafic dyke swarms intruding the

Bundelkhand craton in north-central India (Pati et al., 1999). The most prominent of the

Paleo-Mesoproterozoic (2.5-1.0 Ga) magmatic events are represented by three mafic dyke swarms that intrude the Archean age granitic-gneissic basement of the

Bundelkhand craton. The dykes follow two major trends, suggesting at least two major

pulses of magmatic emplacement within the craton. Based on field cross-cutting relationships, the ENE-WSW trending dykes are the youngest in the craton, and are represented by the ―Great Dyke of Mahoba‖ (figures 3c-d). Previous paleomagnetic studies on these dykes were hindered by poor age constraints (Poitou et al., 2008).

Existing radiometric ages allow an age bracket between 2150 Ma to 1500 Ma for the older suite of mafic dykes (Rao, 2004; Rao et al., 2005; Basu, 1986; Sarkar et al., 1997;

Sharma and Rahman, 1996).

We report new paleomagnetic and geochronologic data on the Bundelkhand mafic dykes from the Mahoba, Banda, Khajurao and nearby areas of the Bundelkhand province in the central India (figure 2). This study offers the first robust age constraints on the paleomagnetic poles calculated from these mafic dykes. These data will ultimately aid in improving our models for the Precambrian tectonic evolution of the

Indian cratons, clarify the role of India in the Columbia and Rodinia supercontinents, and generate data for developing an apparent polar wander path (APWP) for India in this poorly resolved period of Earth history.

Geological Setting and Previous Work

The Archaean-early Proterozoic Bundelkhand craton (BC), commonly known as

Bundelkhand Granite Massif (BGM) is a semi-circular to triangular province that forms the northernmost part of the Indian peninsula (figure 2). It covers an area of ~29,000 km2 and lies between latitudes 24º30‘ and 26º00‘N and longitudes 77º30‘ E and 81º00‘

E (Sharma, 1998). The BC is delimited to the west by the Great Boundary Fault (GBF), to the northeast by the Indo-Gangetic alluvial plains and to the south and southeast by the Narmada-Son lineament (figure 2). The southwestern fringe is marked by relatively small outcrops of Deccan Basalt; additionally, Paleoproterozoic rocks of the Bijawar and

Gwalior Groups are exposed in the SW and NE parts of the BC. The sickle shaped

Vindhyan basin overlies the BC in the south and southeastern sections (Goodwin, 1991;

Naqvi and Rogers, 1987; Pati et al., 2008; Meert et al., 2010). Sharma and Rahman,

(2000) divide the Bundelkhand craton into three litho-tectonic units: (a) the ~3.5 Ga old highly deformed granite-greenstone basement; (b) ~2.5 Ga old multiphase granitoid plutons and associated quartz reefs; and (c) the mafic dykes and dyke swarms. The

BGM represents a significant phase of felsic magmatism associated with a complex of pre-granite sedimentary rocks (Basu, 1986). The basement is represented by a highly deformed granite-greenstone terrane that consists of various enclaves of Archean gneisses, amphibolites, ultramafics, BIF‘s, Tonalite-Trondhjhemite-Gneiss (TTG), marble, calc-silicate rocks, fuchsite quartzites and other metasediments (Roy et al.,

1988; Basu, 1986; Sharma, 1998; Sinha-Roy et al., 1998; Mondal and Zainuddin, 1996;

Mondal et al., 2002). Three different phases of granitoid emplacement are identified in the BGM. In order of decreasing 207Pb/206Pb age, these granites include 2521 ± 7 Ma hornblende granite, 2515 ± 5 Ma biotite granite, and 2492 ± 10 leucogranites and constitute 80-90% of its exposed area (Gopalan et al., 1990; Wiedenbeck and

Goswami, 1994; Wiedenbeck et al., 1996; Mondal et al., 1997, 1998, 2002; Malviya et al., 2004, 2006). Granitoid emplacement in the BC was followed by a number of minor intrusions including widespread pegmatitic veins, porphyry dykes and dyke swarms, and felsic units (rhyolites, dykes of rhyolitic breccias with angular enclaves of porphyries).

Numerous quartz veins of varied size with mainly NNE-SSW and NE-SW trends are observed in parts of the BC representing episodic tectonically controlled hydrothermal activity (Pati et al., 1997; Pati et al., 2007).

The youngest phase of magmatism in the BGM is represented by mafic dykes and dyke swarms that traverse all the above lithologies. More than 700 mafic dykes are known to intrude the granitoid rocks of the BC (Pati et al., 2008). The majority of these dykes trend in a NW-SE direction, with subordinate ENE-WSW and NE-SW trending dykes including the ENE-WSW Great dyke of Mahoba (Basu, 1986; Mondal and

Zainuddin, 1996; Mondal et al., 2002; figure 2). These mafic dykes are subalkaline to tholeiitic in composition and display characteristics indicating continental affinity (Pati et al., 2008). The dykes are commonly exposed as a series of discontinuous and bouldery outcrops extending in length from few tens of meters to more than 17 km. The ‗great dyke‘ of Mahoba has maximum strike length of ≥ 50 km (figure 2 & figures 3c, d). The dykes are generally non-foliated, relatively unaltered and exhibit sharp chilled contacts with the host granitoids (figure 3b). Basu (1986) distinguished at least three generations of dykes based on their cross-cutting relationships. The oldest, coarse grained NW-SE trending suite is cut by ENE-WSW and NE-SW trending medium- grained dykes that include small bodies and lenticles of an aphanitic dolerite. Numerous intermediate to felsic dykes are exposed in the west-central part of the massif between

Jhansi and Jamalpur, including diorite porphyry, syenite porphyry and fine-grained syenite porphyry (Basu, 1986).

The ages for the Archaean rocks in the BGM are poorly constrained due to limited isotopic studies. The oldest rocks in the massif are associated with the TTG magmatism intruding the basement rocks and are assigned an age of 3503 ± 99 Ma

(Rb-Sr isochron; Sarkar et al., 1996). Zircons from the basement gneiss yield an ion microprobe 207Pb/206Pb age of 3270 ± 3 Ma (Mondal et al. 2002). The Bastar, Dharwar

and Aravalli cratons (Sarkar et al., 1993; Wiedenbecke et al., 1996) in peninsular India also contain coeval Archean basement rocks suggesting that these protocontinents played an important role in the earliest stages of nucleation of the Indian shield. The overall stabilization age for the massif has been interpreted to be ~2.5 Ga based on the

207Pb/206Pb ages of the granitoids (Meert et al., 2010; Crawford, 1970; Mondal et al.,

1997, 1998, 2002; Roy and Kröner, 1996). The large scale granitic magmatism in the

BC overlaps with similar events of granite magmatism and mineralization in adjacent

Bastar (2490 ± 10 Ma; Stein et al., 2004) and Dharwar cratons (2510 Ma; Jayananda et al., 2000) indicating widespread granitic plutonism throughout much of the Indian shield at the Archean – Proterozoic boundary. Age control on the Bundelkhand mafic dykes is problematic and is loosely constrained to between ca. 2.1 and 1.5 Ga (Crawford, 1970;

Crawford and Compston, 1970; Sarkar et al., 1997). More recent in-situ 40Ar/39Ar laser ablation data yielded 2150 Ma and 2000 Ma ages for mafic dykes (Rao, 2004; Rao et al., 2005) indicating two pulses of dyke emplacement, however, no details on the locations of the samples were provided in those publications. The majority of the laser spots for the in situ 40Ar-39Ar analysis were in the 2000 Ma range, suggesting that some of the older spot ages may suffer from excess argon within the samples.

Sampling and Methodology

Paleomagnetic Methods

We sampled 27 sites (dykes) and collected a total of ~380 core samples from the mafic dykes intruding the Bundelkhand craton for paleomagnetic analysis (figure 2).

Out of these, only 18 sites (dykes) yielded consistent results and are discussed in this paper (table 1). All samples were drilled in the field using a water cooled portable drill.

The samples were oriented using both sun and magnetic compass and readings were

corrected for local magnetic declination and errors. Samples were returned to the

University of Florida where they were cut into standard sized cylindrical specimens.

These specimens were stepwise demagnetized by using both thermal and alternating field (AF) methods in order to identify the best treatment for isolating vector components within the samples. Alternating field demagnetization was conducted using a home-built

AF demagnetizer and with fields up to 100 mT. Thermal demagnetization was conducted up to temperatures of 600 C with an ASC-Scientific TD-48 thermal demagnetizer. Based upon the magnetic strength of the samples and instrument sensitivity the measurements were made on either a Molspin® spinner magnetometer or a 2G 77R Cryogenic magnetometer at the University of Florida. In samples that showed a very high initial NRM, samples were treated in liquid Nitrogen baths prior to thermal or AF treatment to remove any viscous multi-domain magnetism. Linear segments of the demagnetization trajectories were analyzed via principal component analysis using the IAPD software (Torsvik et al., 2000). Curie temperature experiments were run on representative sample fragments from each site on a KLY-3S susceptibility

Bridge with a CS-3 heating unit. Isothermal Remanence Acquisition (IRM) studies were performed on samples using an ASC-IM30 impulse magnetizer to further characterize the magnetic behavior.

Geochronological Methods

Two samples from the NE-SW trending ―Great Dyke of Mahoba‖ and five samples from the NW-SE trending older suite of mafic dykes from the Bundelkhand craton were processed for geochronology (figure.2). To ascertain the presence of uranium bearing phases, selective samples were thin sectioned and then imaged via scanning electron

microscopy (SEM) for zircon/baddeleyite using the back scattered diffraction (BSD) method. Out of the dozen samples imaged, only two from the NE-SW trending younger suite (GDM and GDM-1) and four (including analyzed sample I9GS-13) from the NW-

SE trending older suite displayed bright spots representative of zircon or baddeleyite, ranging in size between 2-100 µm (figure 4a and 4b). We pulverized sample I9GS_13 from the older suite and Mahoba dyke samples (GDM and GDM-1) for conventional methods of mineral extraction. Using standard gravity and magnetic separation techniques, the zircon grains were concentrated from pulverized samples at University of Florida. The samples were crushed, then disk milled and sieved to < 80 μm grain size fraction. The fractions were then rinsed using Calgon (an anionic surfactant), followed by water table treatment with slow sample feed rates. This was followed by heavy liquid mineral separation with multiple agitation periods to reduce the number of entrapped grains in the lower density fraction. Finally, the samples were repeatedly passed through a Frantz Isodynamic magnetic separator up to a current of 1.0 Amps (2-

4° tilt). Approximately 15-20 clear, euhedral to nearly anhedral zircon grains and zircon fragments were handpicked from the two samples of the NE-SW trending younger

Mahoba suite (GDM and GDM-1) and 25-35 subhedral to euhedral zircon grains from the NW-SE trending older suite (I9GS_13) under an optical microscope to ensure the selection of only the clearest grains and fragments of grains. Further hand-picking of the grains reduced the number to only 7-8 good grains from the Mahoba dolerites and

10-15 grains from the I9GS-13 dyke sample. The grains were then mounted in resin and then polished to expose median sections. Further sonication and cleaning of the

plugs in 5% nitric acid (HNO3) helped to remove any common-Pb surface contamination.

U/Pb isotopic analyses were conducted at the Department of Geological Sciences

(University of Florida) on a Nu Plasma multi-collector plasma source mass spectrometer equipped with three ion counters and 12 Faraday detectors. The LA–ICPMC-MS is equipped with a specially designed collector block for simultaneous acquisition of 204Pb

(204Hg), 206Pb and 207Pb signals on the ion-counting detectors and 235Uand 238U on the

Faraday detectors (Mueller et al., 2008). Mounted zircon grains were laser ablated using a New - Wave 213nm ultraviolet laser beam. During U–Pb analyses, the sample was decrepitated in a He stream and then mixed with Ar-gas for induction into the mass spectrometer. Background measurements were performed before each analysis for blank correction and contributions from 204Hg. Each sample was ablated for ~30 seconds in an effort to minimize pit depth and fractionation. Data calibration and drift corrections were conducted using the FC-1 Duluth Gabbro zircon standard. Data reduction and correction were conducted using a combination of in-house software and

Isoplot (Ludwig, 1999). Additional details can be found in Mueller et al. (2008).

Results

Geochronological Results

The U-Pb ages from the zircon/zircon fragments were determined for the older

crystallization age of the dyke. The regression calculated from nine analyses on four other fragmentary zircons/baddeleyites from sample I9GS_13 yielded an upper intercept age of 2729 ± 83 Ma (2σ; MSWD=4.6; figure 5b; table 2a). This ~2.7 Ga age zircons may reflect the inheritance from the basement rocks in the region (Mondal et al.,

2002). An additional 3 spots from a single zircon yielded a well defined upper intercept age of ~3.3 Ga that is also common in the Bundelkhand craton.

The two Mahoba dyke samples yielded only two well faceted grains and several fragments/tips of zircons (figure 4a). The regression derived from initial 207Pb/206Pb ratios from five laser analysis on three different zircon grains and zircon fragments/tips yielded a minimum concordant age of 1096 ± 19 Ma (2σ; MSWD=0.3; figure 5c; table

2b). The weighted mean 207Pb*/206Pb* age for these five analysis yielded an age of

1113 ± 7 (2σ; MSWD=0.75; probability of concordance=0.56; figure 5d; table 2b).

Paleomagnetic Results

The U-Pb ages from the zircon/zircon fragments were determined for the older

NW-SE trending dyke sample I9GS-13 and the Great dyke of Mahoba samples GDM and GDM1. The dyke sample I9GS_13 yielded a number of well faceted zircons and zircon fragments suitable for U-Pb isotopic analysis. Nine laser analysis on five different euhedral zircons yielded a concordant age of 1980 ± 9 Ma (2σ; MSWD=0.4; probability of concordance=0.55; figure 5a; table 2a) and is interpreted to be the crystallization age of the dyke. The regression calculated from nine analyses on four other fragmentary zircons/baddeleyites from sample I9GS_13 yielded an upper intercept age of 2729 ± 83 Ma (2σ; MSWD=4.6; figure 5b; table 2a). This ~2.7 Ga age zircons may reflect the inheritance from the basement rocks in the region (Mondal et al.,

2002). An additional 3 spots from a single zircon yielded a well defined upper intercept age of ~3.3 Ga that is also common in the Bundelkhand craton.

2b). The weighted mean 207Pb*/206Pb* age for these five analysis yielded an age of

1113 ± 7 (2σ; MSWD=0.75; probability of concordance=0.56; figure 5d; table 2b).

Rock Magnetic Results

Representative results of thermomagnetic analysis (susceptibility vs. temperature) of ENE-WSW and NW-SE trending dyke samples are shown in figure 8 (a-d). The heating and cooling curves of magnetic susceptibility for the ENE-WSW trending

Mahoba dykes as shown in figures 8a, c display two magnetic phases. The heating curves show a prominent peak centered close to 250 ºC - 300 ºC and susceptibility drop above 300° C, indicating the likely existence of pyrrhotite. A rapid decrease in the magnetic susceptibility around 550 ºC - 575 ºC indicates the existence of magnetite

(figures 8a & 8c). The cooling curve shows higher susceptibility which is probably caused by the ex-solution and conversion of Ti-magnetite to pure magnetite.

The rock magnetic studies on the dominantly occurring NW-SE trending older suite of dykes show nearly reversible Curie temperature runs characteristic of magnetite

(figures 8b & 8d). The heating Curie temperature TcH of dyke sample I927-12 is 572.8

ºC, and the cooling Curie temperature TcC is 567.5 ºC (figure. 8d). Isothermal remanence acquisition (IRM) curves along with back field coercivity of remanence from both these ENE-WSW and NW-SE trending mafic dykes also indicate magnetite as the

principal magnetic carrier in the samples. A rapid rise in intensity near saturation at

~0.25-0.3 Tesla was observed for majority of the dyke samples and their magnetization remains constant at higher fields, up to the highest applied field of 1.2 Tesla. The values for the back field coercivity of remanence ranges between 0.07 and 0.1 mT

(figures 8e-f).

Discussion

~ 1.1 Ga Paleomagnetism

The VGP calculated for the ENE-WSW trending Mahoba dykes of the

Bundelkhand craton is significant in terms of its age and the direction. The U-Pb zircon age of 1113 ± 7 Ma for Mahoba dykes falls in the same time bracket as the Majhgawan kimberlite that intrudes both the Lower Vindhyan sequence (~1.6 Ga) and the Baghain sandstone of the Kaimur Group (Upper Vindhyan). Gregory et al. (2006) dated

Majhgawan kimberlite at 1073 ± 13.7 Ma via 40Ar-39Ar analysis of phlogopite phenocrysts and obtained a virtual geomagnetic pole at 36.8 ⁰S, 32.5 ⁰E (dp = 9º; dm =

16.6º; also see Miller and Hargraves, 1994) that is statistically indistinguishable from the virtual geomagnetic pole calculated for the Mahoba dyke swarm in this study (figure 10).

Malone et al. (2008) conducted a paleomagnetic study on the Bhander-Rewa Groups

(Upper Vindhyan) and obtained a mean paleomagnetic pole at 44⁰S, 34.0⁰E (A95 =

4.3). We note that the VGP‘s of the Mahoba dykes generated in this study and the penecontemporaneous Majhgawan kimberlite are nearly identical to the paleomagnetic poles obtained from the Bhander-Rewa Groups in the Upper Vindhyan sequence (figure

10).

Azmi et al. (2008) argue that the paleomagnetic data from the Majhgawan kimberlite and the Bhander-Rewa sequence are indeed age-equivalent and late

Neoproterozoic in age, but that the Majhgawan phlogopites crystallized at depth some

400 million years earlier. We acknowledge that the presence of similar magnetizations in three different units in close proximity may also raise concern about the possibility of remagnetization. However, we note that there is no reported ~1.1 Ga remagnetization event within or in the vicinity of the Bundelkhand craton.

Additional support for the primary nature of magnetization is found in both Upper

Vindhyan sequence and the NW-SE trending older suite of Bundelkhand dykes. Upper

Vindhyan sedimentary units show at least eleven geomagnetic reversals supporting a primary magnetization in these rocks. Similarly, the presence of partial baked contact test yielded by the granitic host rock samples traversed by the mafic dykelets at site

I925 (figure.7e & 7f) and the dual polarity magnetization shown by one of the dykes (site

I443 C) support a primary magnetization in the NW-SE trending older suite of

Bundelkhand dykes.

Age implications for the Bhander-Rewa sequence of the Upper Vindhyan

The age of deposition in Vindhyan basin located to the south of the Bundelkhand

Craton (figure 1) in the northern Indian peninsular shield, has been debated for over 100 years (Oldham, 1893; Auden, 1933; Crawford and Compston, 1970; Venkatachala et al., 1996; Malone et al., 2008; Azmi et al., 2008). The onset of sedimentation in the lower Vindhyan Supergroup is generally well constrained at around 1.6-1.8 Ga by radiometric data (Rasmussen et al., 2002a, b; Ray et al., 2002, 2003 Sarangi et al.,

2004; Kumar, 2001); however the age of the Upper Vindhyan is still enigmatic and highly contentious due to a lack of targets suitable for geochronology, controversial

fossil finds and poorly constrained global stable isotopic correlations (Vinogradov et al.,

1964; Crawford and Compston, 1970; Paul et al., 1975; Srivastava and Rajagopalan,

1988; Chakrabarti, A., 1990; Smith, 1992; Kumar and Srivastava, 1997; Kumar et al.,

2002; De, 2003, 2006; Ray et al., 2002, 2003; Rai et al., 1997; Gregory et al., 2006;

Malone et al., 2008, Azmi et al., 2008).

The Upper Vindhyan sedimentary rocks were typically correlated with the Marwar

Supergroup (Rajasthan); sequences in the Salt Range of Pakistan and with the Krol-Tal

Group of the Lesser Himalayas (McElhinny et al., 1978; Klootwijk et al., 1986;

Mazumdar and Banerjee, 2001). These correlations, however, are based on somewhat similar lithologies and the proximity of the undeformed Marwar and Upper Vindhyan strata. The lithologic comparisons between these units are rather problematic. For example, the evaporite deposits are prevalent within the Marwar and Salt Ranges, but absent in the Upper Vindhyan sequence. Malone et al. (2008) examined detrital zircon suites from both the Upper Vindhyan sedimentary rocks in Rajasthan and the nearby

Marwar Supergroup. The 207Pb/206Pb age distribution observed in the detrital zircon analysis of the Upper Bhander sandstone yielded several noteworthy peaks between

1850 – 1050 Ma (Malone et al., 2008). In contrast, detrital zircons analyzed from the

Sonia and Girbarkhar sandstone of the Marwar Supergroup yielded age population peaks in the 840-920 Ma range, a population completely absent in the Upper Bhander sandstone of the Vindhyan Supergroup. Hence, the results from the detrital zircon geochronology suggest that previous correlations between the two depositional sequences are incorrect. Malone et al. (2008) hypothesized that the closure age of the

Upper Vindhyan sedimentation was no younger than 1.0 Ga based on their

paleomagnetic study on the Bhander-Rewa Groups of the Upper Vindhyan and the detrital zircon work on Bhander-Rewa units and the Marwar Supergroup. Additional support for a late Mesoproterozoic closure age came from a paleomagnetic/geochronologic study of the Majhgawan kimberlite (Gregory et al.,

2006).

The paleomagnetic and geochronological data from the Mahoba dykes generated in this study is germane to the discussion of the sedimentation ages in the Upper

Vindhyan basin. The VGP obtained for the ~ 1113 Ma Mahoba dykes (37.8ºS, 49.5ºE) is nearly identical to the mean paleomagnetic pole from the Bhander – Rewa Groups and the Majhgawan kimberlite (figure 10; see discussion above). Our interpretation of the Mahoba paleomagnetic and geochronologic results therefore lends additional support to the proposal of Malone et al. (2008) and Gregory et al. (2006) that the closure of sedimentation in the Upper Vindhyan sequence is older than ~ 1.0 Ga.

India in Rodinia supercontinent at 1100 Ma

The late Mesoproterozoic (ca. 1100 Ma) has been postulated as the time interval for the formation of the supercontinent Rodinia (McMenamin and McMenamin, 1990;

Dalziel, 1991; Moores, 1991; Hoffman, 1991). The existence of Rodinia is supported by the presence of a number of 1300-900 Ma old orogenic/mobile belts (Dewey and Burke,

1973) and associated geologic links between the various cratonic nuclei (Young, 1995;

Dalziel, 1997; Rainbird et al., 1998; Karlstrom et al., 1999; Sears and Price, 2000;

Dalziel, 2000). However, the geometry/paleogeography and duration of the Rodinia supercontinent remains extremely fluid and controversial due to a paucity of well dated, high quality paleomagnetic poles from various continental blocks forming the

supercontinent (Weil et al., 1998; Meert and Powell, 2001; Meert, 2001; Meert and

Torsvik, 2003; Li et al., 2008).

One of the outcomes of this study is our ability to constrain the paleoposition of

Indian sub-continent in Rodinia configuration at ~ 1.1 Ga. The recent paleomagnetic and geochronologic studies from the Majhgawan kimberlite (Gregory et al., 2006) and

Bhander-Rewa Groups of Upper Vindhyan (Malone et al., 2008) provided high quality poles to constrain the paleogeographic position of India at 1.1 Ga. There are a number of paleomagnetic poles available from other elements of the Rodinia supercontinent that are more or less coeval with those from India. Considering the East Gondwana elements, the key pole for this interval in Australia is the ~1070 Ma doleritic rocks of

Bangemall sills (Wingate et al., 2002). The Bangemall pole achieves a score of Q = 7 in the reliability scheme of Van der Voo (1990).

The paleogeographic position for Laurentia is based on the combined mean paleomagnetic data from the Portage Lake volcanics and Keweenawan dykes (Pesonen et al., 2003; Swanson-Hysell et al., 2009). Based on the extensive field and laboratory tests these poles are inferred to be primary and have precise zircon/baddeleyite U-Pb dates of 1095 Ma (Portage Lake Volcanics) and ~1109 Ma (Keweenawan volcanics), respectively (Halls and Pesonen, 1982; Davis and Sutcliffe, 1985; Goodge et al., 2008).

More recently, Swanson-Hysell et al. (2009) provided high resolution paleomagnetic data from a series of well-dated basalt flows at Mamainse Point, Ontario, in the

Keweenawan Rift and suggested that the previously documented reversal asymmetry for these volcanic rocks is an artefact of the fast motion of the Laurentian plate towards the equator at this interval (also see Meert, 2009). The combined mean of the Portage

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Lake volcanics and the Keweenawan dolerites pole places Laurentia at intermediate paleolatitudes in our ~1.1 Ga reconstruction.

There are two paleomagnetic poles available from the Baltica between ~ 1100 Ma and 1123 Ma. Most of the Rodinia reconstructions at ~ 1100 Ma utilize the paleomagnetic data from the Bamble intrusions in southern Norway to constrain the position of Baltica (Brown and McEnroe, 2004; Stearn and Piper, 1984). A Sm/Nd metamorphic age of 1095 Ma has been reported for the Bamble intrusive rocks intruding the high grade granulite facies. Recently, Salminen et al. (2009) reported a virtual geomagnetic pole for the Baltica from their paleomagnetic and rock magnetic studies on the well dated 1122 ± 7 Ma (U-Pb age; Lauerma, 1995) Salla diabase dykes in northeastern Finland. The palaeolatitudinal position of Baltica constrained by the

Bamble intrusive poles and the Salla dykes VGP indicate a latitudinal difference of almost 90º. This would require an unusually rapid southward drift of the Baltica plate for the short time interval of ~ 28 Ma. However, the Salla diabase dyke pole is highly tentative and requires additional verification due to the inadequate averaging of secular variation (Salminen et al., 2009). Given the tentative nature of the Salla diabase pole, we select the previously reported mean paleomagnetic pole from the Bamble intrusions to place Baltica in our reconstruction (Brown and McEnroe, 2004; Stearn and Piper,

1984).

A key, well-dated paleomagnetic pole, from the Kalahari craton at ~ 1.1 Ga is based on the Umkondo dolerites and the Kalkpunt poles (Hanson et al., 2004a; Gose et al., 2004; 2006; Jacobs et al., 2008). The precisely dated Umkondo dolerite paleomagnetic data from the Kalahari craton are often correlated with Keweenawan

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igneous suite of Laurentia (Hanson et al., 2004a; Gose et al., 2006; Jacobs et al.,

2008). These two coeval paleomagnetic poles are predominantly of one polarity and thus constrain the relative paleolatitudes of the two cratons within the Rodinia configuration. The paleomagnetic data in our reconstruction at ~1.1 Ga are derived from the mean pole based on the results from Timbavati gabbros, the Umkondo-Borg large igneous province and the Grunehogna province of Antarctica after restoring East

Antarctica to its position next to South Africa at ~1.1 Ga (Jones and McElhinny, 1966;

Renne et al., 1990; Hargraves et al., 1994; Jones and Powell, 2001; Powell et al., 2001;

Pesonen et al., 2003; Hanson et al., 2004a; Gose et al., 2006; Jacobs et al., 2008).

The reconstruction discussed herein also utilizes the mean paleomagnetic pole from the granophyres and rhyolite nunataks of Coats Land, Antarctica to evaluate the paleogeographic position of Coats Land in Rodinia and their connection to the Kalahari craton at ~1.1 Ga (Gose et al., 1997). The granophyres and rhyolite nunataks provide indistinguishable mean paleomagnetic directions with precise U-Pb zircon ages of 1112

± 4 Ma (rhyolites) and 1106 ± 3 Ma (granophyres) respectively. These volcanic rocks are interpreted to represent part of the Umkondo-Borg-Keweenawan magmatic province

(Jacobs et al., 2008).

The position of North China Block in late Mesoproterozoic times (~1100 Ma) is constrained by paleomagnetic data from the Tieling formation of the Jixian Province

(Zhang et al., 2006). The paleomagnetic pole appears to be primary having passed both fold and reversal tests; however, the age of the magnetization remains tentative.

Nevertheless, we use these results to position the North China block in our reconstruction.

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The position of the Siberian block is contentious during Rodinia assembly (Sears and Price, 2000; Pisarevsky and Natapov, 2003; Meert and Torsvik, 2003; Li et al.,

2008; Pisarevsky et al., 2008). In a majority of Rodinia configurations, Siberia is placed with its (present-day) southern margin in the vicinity of northern Laurentia (Gallet, 2000;

Pisarevsky and Natapov, 2003; Li et al., 2008). Others argue for either non – inclusion of the Siberian craton within the Rodinia assembly (Meert and Torsvik, 2003; Pisarevsky et al., 2008) or place Siberia adjacent to the present-day Cordilleran margin of Laurentia

(Sears and Price, 2003). We use paleomagnetic data from the Late Mesoproterozoic

Malgina and Linok formations of southeastern Uchur-Maya region and northeastern

Turukhansk region of the Siberia craton (Gallet, 2000) in our reconstruction.

Chemostratigraphic, biostratigraphic and isotopic data constrain the age of these two formations between 1100-1050 Ma and the presence of positive fold and reversal tests, together with a conglomerate test suggest a primary magnetization in these rocks.

Figure 11 shows the paleogeographic reconstruction of Rodinia at ~ 1.1 Ma utilizing the paleomagnetic poles from the above mentioned cratonic blocks. This time period of 1100 Ma is known to record the first phase of Rodinia assembly (Pesonen et al., 2003). Our reconstruction is based on the ―closest approach‖ technique as described in Meert and Stuckey (2002); Buchan et al. (2000, 2001) and Pesonen et al.

(2003), using individual poles to reconstruct the continents to their correct palaeolatitudinal position since paleomagnetism cannot provide longitudinal control.

In our reconstruction, the combined mean of Portage Lake volcanics and

Keweenawan poles (normal + reversed) (Pesonen et al., 2003; Goodge et al., 2008;

Swanson-Hysell et al., 2009) places Laurentia at an intermediate latitudinal position with

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the North China block at its present day north and Australia located further south.

Zhang et al. (2006) suggested a common APW path for the North China Block (NCB) and Laurentia from ca. 1200 Ma to ca. 700 Ma in support of an NCB-Laurentia connection during this interval. Siberia is placed at a distance from Laurentia to accommodate the North China block, with its southern margin facing the northern margin of Laurentia (Rainbird et al., 1998; Pisarevsky et al., 2008). While the position of

Laurentia is well constrained at this interval, the placement and orientation of Siberia and NCB remains poorly understood and requires additional support. Pesonen et al.

(2003) suggested a southerly drift for Baltica away from Laurentia between 1.25 Ga and

1.1 Ga based on the paleomagnetic data from the Sveconorwegian Province (Poorter,

1975; Patchett and Buylund, 1977; Pesonen and Neuvonen, 1981). In our reconstruction, Baltica occupies an independent southerly latitudinal position derived from the combined paleomagnetic pole from the 1095 Ma Bamble intrusives. A tighter fit can be obtained between the Baltica and Laurentia by inverting the polarity of the

Bamble dykes pole; however, this position would require southern Baltica to face the northern margin of Laurentia for which there is little support from the geology. Thus, we follow the southern palaeolatitudinal option of Pesonen et al. (2003) to constrain Baltica in our reconstruction (also see Salminen et al., 2009).

Based on the available paleomagnetic data from Kalahari and Laurentia, Kalahari is located too far south of Laurentia to support any linkage between the two blocks

(Hanson et al., 2004a; Gose et al., 2006). The ~ 1070 Ma Bangemall Sill pole from

Western Australia (Wingate et al., 2002) constrains the Australia-Mawson landmass at low to intermediate palaeolatitudes along the southern margin of Laurentia. In their

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reconstruction, Powell and Pisarevsky (2002) suggested that the Kalahari craton is separated from Laurentia by the Australia – Mawson cratons. They noted that ~ 1080

Ma deformation in the Darling belt along the western margin of Australian margin might link with the coeval deformation in the Natal and Maud belts along the southern margin of Kalahari. Such an interpretation is possible within the framework of Figure 11. This separation of the Kalahari and Laurentia at 1.1 Ga – 1.0 Ga is such that it allows the placement of parts of East Gondwana to the southwest or south of Laurentia (Meert and

Torsvik, 2003; Pesonen et al., 2003). Among the elements of the East Gondwana, the link between India and East Antarctica remains paleomagnetically untested for this interval due to the absence of any paleomagnetic data from East Antarctica at ~ 1100

Ma. However, the proposed geological coherence of the Mawson block with Australia in post-1200 Ma Rodinia reconstructions (Cawood and Korsch, 2008; Payne et al., 2009;

Wang et al., 2008; Kelly et al., 2002; Torsvik et al., 2001) and the disposition of India at low latitudes (this study) do favor a loose association of these cratonic elements, with

India positioned at the southern margin of the Mawson block. Meert et al. (1995) and

Meert (2003) argue that East Gondwanaland was not a coherent entity until late

Neoproterozoic or Early Cambrian (see also Fitzsimons, 2000a; b; Boger et al., 2001,

2000; Powell and Pisarevsky et al., 2002). We have reliable paleomagnetic data from

India and Australia, but to make any robust conclusion our interpretation relies upon additional paleomagnetic data from the East Antarctica craton at this interval.

The time period of ~ 1.1 Ga represents the initial phase of the Rodinia amalgamation. Our paleomagnetic reconstruction at ~ 1.1 Ga is based on the existing database with smaller number of continental blocks well constrained in space and time.

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The palaeoconfiguration of these blocks suggest a loose association within Rodinia with no paleomagnetic poles from majority of the West Gondwana elements.

~ 2.0 Ga Paleomagnetism

Tectonic evolution of the Aravalli and Dharwar protocontinents

A series of collisional events took place during the Paleoproterozoic interval (~2.1

– 1.8 Ga) on a near global scale (Zhao et al., 2002; French et al., 2008). The 2.0 - 1.8

Ga orogenic belts are thought to be the result of accretion of Archaean cratonic nuclei to form numerous supercratons or possibly, a supercontinent (Hoffman, 1988; Gower et al., 1990; 1992; Krapez, 1999; Rogers and Santosh, 2002, 2003; Condie, 2002; Zhao et al., 2002, 2004). The Central Indian Tectonic zone (CITZ) is a linear Paleoproterozoic orogenic belt that represents the locus of collision between the northern Aravalli protocontinent (Aravalli-Delhi belt, Bundelkhand massif and Vindhyan basin) and the southern Dharwar protocontinent (Bastar, Singhbhum, Dharwar cratons and Southern

Indian Granulite terrane; figure. 12a). The connection between the Bastar, Singhbhum and Dharwar cratons of the South Indian block is well established at least by 1.9 Ga

(Ramchandra et al., 1995; French et al., 2008). The temporal control on the collision between the Aravalli and the Dharwar protocontinents is still contentious (Meert et al.,

2010). Stein et al. (2004) proposed an oblique collision between the northern and southern blocks around 2.5 Ga with the initiation of the formation of CITZ at this time.

Other authors argue for a younger age of collision at ~1.9-1.7 Ga between the two blocks (Acharyya et al., 2003).

Our new paleomagnetic data from the ~2.0 Ga NW-SE trending mafic dykes intruding the Bundelkhand craton combined with paleomagnetic data from the ~1.9 Ga mafic dykes of the Bastar craton (Meert et al., 2010; French et al., 2008), Cuddapah

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trap volcanics (Clark, 1982) and a dyke adjacent to the Cuddapah basin in the Dharwar craton (Kumar and Bhalla, 1983) can provide some insight into the timing of the collision between the Aravalli and Dharwar protocontinents (figure12b). The combined result from the ~2.0 Ga NW-SE trending older suite of dykes from the Bundelkhand craton yield a paleomagnetic pole for India at 58.5ºN and 312.5ºE (dp = 4.9º; dm = 9.7º) and places the Bundelkhand craton in an equatorial position (figure 12b). The combined 1.9

Ga paleomagnetic pole for the Dharwar and Bastar cratons falls at 31° N, 330°E (dp =

6.3°; dm = 12.2°) and also places these cratonic nuclei at low latitudes (Meert et al.,

2011). A nearly equatorial position for the Bundelkhand craton is also indicated by the paleomagnetic data from ~ 1.8 Ga Gwalior volcanics of the Bundelkhand craton that generates a paleomagnetic pole for the cratonic nuclei at 15.4⁰ S and 353.2⁰E (dp = 5.6; dm = 11.2; Pradhan et al., 2010). Utilizing the three paleomagnetic poles from the

Bundelkhand dykes, Bastar-Cuddapah dykes and the Gwalior volcanics, an apparent polar wander (APW) path was constructed for the Indian peninsular shield during the

2.0- 1.8 Ga interval(figure 12b). The APWP indicates that India rotated through 80 degrees with only slight changes in paleolatitude during this interval.

Implications for Columbia

The Paleo-Mesoproterozoic history of the Earth is punctuated by the presence of globally distributed 2.1-1.8 Ga collisional orogens related to a proposed pre-Rodinian supercontinent named Columbia (Zhao et al., 2002; Zhao et al., 2004a, b; Condie,

2000, 2002; Rogers and Santosh, 2002; Kusky et al., 2007a; French et al., 2010). In particular, the Columbia model by Zhao et al. (2004) is based primarily on the geological

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connections between globally distributed cratonic nuclei and the presence of the Paleo-

Mesoproterozoic orogenic belts (2.1-1.8 Ga; figure 13a).

Reliable paleomagnetic data provide a quantitative reference frame for documenting the history and dynamics of the plate tectonic regime and constraining the position of various cratonic blocks within the proposed continental assemblies. Any attempt at global reconstructions in the Precambrian is an arduous task due to the scarcity of well-dated paleomagnetic data, polarity ambiguity and the absence of longitudinal control in positioning of the continents (Van der Voo and Meert, 1991;

Meert, 2002; Meert et al., in 2011). In addition, the majority of present day continents are an amalgam of older Archean cratonic nuclei that were not fully welded until

Mesoproterozoic or later time. The paleomagnetic database for the Paleoproterozoic has improved over the past decade as reliable and well-dated poles are becoming more commonplace. Below, we discuss the paleogeographic reconstruction of the

Paleoproterozoic Columbia supercontinent utilizing the most updated paleomagnetic results from various cratonic blocks at ~2.0 Ga (figure 13b; table 3a).

The ~ 2.0 Ga paleomagnetic pole calculated from the NW-SE trending dyke suite of the Bundelkhand craton constrains the position of the North Indian Block (Aravalli protocontinent) in our reconstruction. As noted above, the North Indian and the South

Indian blocks of the peninsular India were in close proximity during 2.1 – 1.9 Ga (also see Meert et al., in press; Stein et al., 2004). Hence, we suggest that the ~ 2.0 Ga pole from the Bundelkhand dykes is representative of a major portion of Peninsular India for this time period.

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The palaeolatitudinal position of the Superior craton of the Laurentian block is constrained at ~2.0 Ga via the paleomagnetic data from the diabase dykes of the Minto block of the northeastern Superior Province (Buchan et al., 2000). The Minto dykes paleomagnetic pole has a positive baked contact test and is precisely dated at 1998 ± 2

Ma by U-Pb isotopic methods (Buchan et al., 2000). Most of the other cratonic nuclei of

Laurentia such as Slave-Rae-Hearne and Wyoming blocks collided with the Superior craton during Trans-Hudson orogeny (2.1-1.8 Ga), although the terminal collision was protracted until ~1.78 Ma (Corrigan et al., 2005). Recent paleomagnetic studies on the

Lac de Gras diabase dykes across the Slave Province of the Canadian Shield suggest that the Slave and Superior cratons were located at similar latitudes at ~2.02 Ga but with different relative orientations than today (Buchan et al., 2009). Due to the absence of coeval paleomagnetic data from other cratonic blocks of Laurentia at ~2.0 Ga, we use the Minto dyke pole as representative for Laurentia assuming that the other Laurentian nuclei were in close proximity during this period.

A low to intermediate latitudinal position for the Fennoscandian shield during

Paleoproterozoic is based on paleomagnetic data between 2.2 – 1.97 Ga (Pisarevsky and Sokolov, 1999; Neuvonen et al., 1997; Mertanen and Pesonen, 1994). The paleoposition of the Fennoscandia is constrained in our reconstruction utilizing the only available paleomagnetic pole from the 1974 ± 27 Ma (Sm-Nd age) Konchozero peridotite sill (component I) of the Fennoscandia (Pisarevsky and Sokolov, 1999).

According to Bogdanova et al. (1994) the Ukranian shield was probably attached to the Fennoscandia at ca. 2.0 Ga. However, the paleomagnetic data for the Ukranian shield between ca. 2.0 and 1.72 Ga suggest that the final accretion of Fennoscandia to

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the Ukranian shield took place sometimes after ca. 1.8 Ga (Elming et al., 2001). The position of the Ukranian shield at ~2.0 Ga in our reconstruction is derived from the paleomagnetic data from the gabbro diabase dyke of the southern Sarmatia province of

Baltica (Elming et al., 2001).

The Kalahari craton of South Africa consists of the Kaapvaal and the Zimbabwe blocks welded together along the Limpopo belt sometime between 1.9 Ga and 2.06 Ga

(De Wit et al., 1992; Lubnina et al., 2010b). Paleomagnetic data for the Zimbabwe craton at ~2.0 Ga are lacking and therefore its connection with the Bushveld complex of the Kaapvaal craton is also ambiguous (Lubnina et al., 2010b; Soderlund et al., 2010).

We use updated paleomagnetic data from the ~2.02 Ga Vredefort dykes/impact structures of the Kaapvaal craton to constrain its position at ~2.0 Ga in our reconstruction (Salminen et al., 2009; Carporzen et al., 2006; Pesonen et al., 2002).

The metamorphosed Uauã mafic dykes of São Fransisco provide the only paleomagnetic pole for Congo-São Fransisco cratons at ~2.0 Ga (De‘Agrella Filho and

Pacca, 1998). Paleomagnetic, petrographic and geochronologic evidence support a secondary origin of magnetization in these dykes, probably acquired approximately at 2

Ga-1.9 Ga, during the uplift and regional cooling of the Transamazonian cycle

(De‘Agrella Filho and Pacca, 1998).

Nomade et al. (2003, 2001) reported a paleomagnetic pole from the Oyapok-

Campoi river zone (OYA in their manuscript) of the French Guiana shield of the

Amazonian craton with a magnetization age of 2036 Ma. Other paleomagnetic poles that exist for the Amazonian craton at ~2.0 Ga come from the Encrucijada pluton of the

Venezuela block (Onstott and Hargraves, 1981). The statistical precision on the Rb-Sr

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ages of the Encrucijada granites are poor and the ages are considered rather inconsistent. The two poles are significantly different and, on the basis of the difference, it was argued that there was northward movement along the Pisco Jurua fault zone during the 2036 - 2000 Ma periods (Onstott and Hargraves, 1981; Nomade et al., 2001). Although not united but lying in close association, we consider the OYA poles to constrain the position of the Amazonia craton in our ~2.0 Ga reconstruction.

Utilizing all the available paleomagnetic data from the cratonic blocks mentioned above, we attempt to test the proposed Columbia configuration of Zhao et al. (2002,

2004a,b). Figure 13(a) represents the typical Columbia configuration of Zhao et al.

(2004) based on the correlation of the coeval orogenic belts (2.1 - 1.8 Ga) within its constituent cratonic blocks. The Zhao et al. (2004) configuration is not supported by the existing paleomagnetic database at ca. 2.0 Ga (figure13b; table 3a). The majority of these continental nuclei occupy low to intermediate latitudinal positions as they did in the Kenorland assembly at ca. 2.45 Ga (Salminen et al., 2009; Pesonen et al., 2003).

The only resemblance to the Columbia model by Zhao et al. (2004) is exhibited by the juxtaposition of the parts of Amazonia and West-Africa (figure 13b; table 3a).

In the traditional Columbia configuration, India is placed adjacent to East

Antarctica and North China along the present day south-western margin of Laurentia

(Zhao et al., 2004). The Kalahari craton of South Africa, on the other hand is positioned along the north-western margin of Australia (figure 12a). In our reconstruction, the new paleomagnetic data from ~2.0 Ga Bundelkhand dykes (this study) constrain the position of India at nearly equatorial latitudes with Laurentia occupying low to intermediate latitudes. The Kaapvaal block of the Kalahari craton can be placed at mid-latitudes

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either north or south depending upon polarity choice. Baltica did not exist as a coherent block at ~2.0 Ga and the available paleomagnetic data from the Fennoscandian and the

Ukrainian shield position these two blocks at mid-latitudes with no obvious connection to

Laurentia (Salminen et al., 2009).

The African and South-American continental fragments represent a loose connection in the configuration at ~2.0 Ga and may show the only resemblance to the archetypal Columbia model of Zhao et al. (2004). The low southerly latitudinal positions indicated by the paleomagnetic data for Amazonia, West Africa and Congo-Sao

Fransisco blocks may indicate that these were juxtaposed at ~2.0 Ga (Nomade et al.,

2001, 2003). The presence of coeval Trans-Amazonian and Eburnean orogenic events in Amazonia (Hartmann, 2002) and West Africa (Ledru et al., 1994) respectively, may lend additional support to their connection at 2.1-1.9 Ga.

Although our reconstruction at ~2.0 Ga is based on limited number of paleomagnetic poles from a small number of cratonic blocks, most of the poles are reasonably constrained in both time and space. Thus, we cannot reject the existence of a ―Columbia-type‖ supercontinent during the Paleoproterozoic; though the independent drift of most of the continental blocks at ~ 2.0 Ga enables us to reject some of the relationships between cratonic blocks suggested in the Columbia model by Zhao et al.

(2004).

Summary

(1). The paleomagnetic and geochronological results on the mafic dyke swarms intruding the Archean - Paleoproterozoic Bundelkhand craton presented in this paper are significant not only in providing reliable age constraints for these magmatic events, but also for improving our understanding of the tectonic evolution of the Central Indian

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shield. The three distinctive paleomagnetic directions obtained in this study for the mafic dyke swarms suggest three different pulses of emplacement within the BC. The precise U/Pb isotopic zircon age of 1980 Ma for the NW-SE trending dyke (sample

I9GS-13) and 1113 Ma for the ESE-WSW trending great dyke of Mahoba support our paleomagnetic analysis and the geologic correlations. (2) This study provides a refined virtual geomagnetic pole for the Mahoba suite of dykes at 37.8ºS, 49.5ºE

(dp/dm=10.8/18.3). This geomagnetic pole correlates well with the VGP generated for the Majhgawan kimberlite and the poles from the Bhander-Rewa Groups of the Upper

Vindhyan sequence (Gregory et al., 2006; Malone et al., 2008). Our interpretation of the Mahoba paleomagnetic and geochronologic results lend further support to the argument that the closure age for the Upper Vindhyan rocks is older than 1000 Ma. (3)

The paleomagnetic results from the ~2000 Ma NW-SE trending dyke suite of the

Bundelkhand craton along with other published data from the 1800-2000 Ma interval indicates a close proximity between the major elements of Peninsular India (Aravalli-

Bundelkhand and Dharwar-Bastar and Singhbhum cratons). This would suggest that the younger tectono-magmatic events recorded in the CITZ represent crustal-scale reactivation along an existing zone of weakness. (4) We also provide paleogeographic reconstructions at 2.0 Ga and 1.1 Ga. Our ~2.0 Ga pole from the NW-SE trending

Bundelkhand dykes favors a low latitudinal position for India with mid latitudinal disposition of Laurentia and the Kaapvaal block. The proposed connection of India to

East Antarctica and North China (Zhao et al., 2004) remains untested due to the absence of paleomagnetic data from North China and East Antarctica cratons at ~2.0

Ga. Laurentia and India can be placed in proximity to each other but in a vastly different

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configuration than that suggested by Zhao et al. (2004) based on the geometric relationships of the radiating dyke swarms. The global paleogeography at ~2.0 Ga favors scattered independent continental blocks as suggested by Pesonen et al. (2003; also see Salminen et al., 2009) and does not support the Columbia model of Zhao et al.

(2004). (5) With reference to Rodinia, one of the interesting implications of our results from ~1113 Ma Mahoba dyke pole in conjunction with the coeval Majhgawan kimberlite and Bhander-Rewa poles is that it infers a low latitudinal position for India with Australia

– Mawson blocks located at mid latitudes. The link between India and East Antarctica remains paleomagnetically untested for this interval due to the absence of any paleomagnetic data from the East Antarctica block. However, the proposed geological coherence of the Mawson block with Australia in post-1200 Ma configurations (Cawood and Korsch, 2008; Payne et al., 2009; Wang et al., 2008; Kelly et al., 2002; Torsvik et al., 2001) and the disposition of India at low latitudes (this study) might indicate a loose

East Gondwana fit, with India positioned further south of East Antarctica.

Acknowledgements: This work was supported by a grant from the US National

Science Foundation to J.G. Meert (EAR09-10888). We wish to thank Shawn Malone for the careful proofreading of the manuscript. We would also like to thank Dr. Rajesh

Srivastava and an anonymous reviewer for the thorough and thoughtful review of the manuscript.

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Figure 4-1. Generalized geologic and tectonic map of the Indian sub-continent with the Precambrian mafic dyke swarms cross-cutting various Archean cratonic blocks. Ch: Chattisgarth Basin; CIS: Central Indian Shear Zone; GR: Godavari Rift; M: Madras Block; Mk: Malanjkhand; MR: Mahanadi Rift; N: Nilgiri Block; NS: Narmada-Son Fault Zone; PC: Palghat-Cauvery Shear Zone; R: Rengali Province and Kerajang Shear Zone; S: Singhbhum Shear Zone; V: Vindhyan Basin (modified after French et al., 2008)

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Figure 4-2. Sketch map of the major units in the Bundelkhand craton, NW India (modified after Malviya et al., 2006). The asterisks on the map show the sites sampled for both paleomagnetic and geochronologic analysis (NW-SE trending dyke I9GS-13 and ENE-WSW trending Great Dyke of Mahoba). Inset figure shows the tectonic map of Indian sub-continent.

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Figure 4-3. Field photographs of the Bundelkhand dykes. (a) Photo of a NW-SE trending dyke (I925) exposed in an outcrop ~50 km away from Mauranipur district, M.P. (b) granitic host rock at the same site I925 trellised by numerous mafic dykelets. (c & d) Photos of ENE-WSW trending dyke I923 exposed in a quarry in Mahoba district, Madhya Pradesh, Central India.

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Figure 4-4. Backscattered and cathodoluminiscence images of the zircons/uranium bearing minerals for the Bundelkhand dykes. (a) Backscattered electron (BSE) and cathodoluminescence (CL) images of selected zircons/uranium bearing minerals from ENE-WSW trending Mahoba dykes. Scale bars in µm. (b) Backscattered electron (BSE) and cathodoluminescence (CL) images of selected zircons/uranium bearing minerals from NW-SE trending dykes. Bar length corresponds to 10-100 µm.

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Figure 4-5. U/Pb isotopic analysis of the dyke samples I9GS_13 and great dyke of Mahoba (GDM and GDM1). (a) Tera- Wasserburg concordia diagram for the nine spots from five concordant zircon/baddeleyite grains and tips from sample I9GS_13 yielding an age of 1980.1  9.4 Ma (2) with a MSWD=0.4. (b) Weatherhill plot obtained from 9 analyses on a population of four fragmentary zircons/baddeleyites from NW-SE trending dyke sample I9GS_13 yielding an upper intercept age of 2729  83 Ma (2σ; MSWD=4.6) (c) Tera-Wasserburg concordia intercept age of 1096  19 Ma (2σ; MSWD=0.3) derived from the regression through the uncorrected data from the initial 207Pb/206Pb for the five analysis on three zircon grains and tips from sample GDM & GDM1 (d) The weighted mean 207Pb*/206Pb* age for the five analysis from sample GDM & GDM1 yielded an age of 1113 ± 7 (2σ; MSWD=0.75; probability of concordance=0.56).

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Figure 4-6. Orthogonal vector plots from the NW-SE and ENE-WSW trending dykes of the Bundelkhand craton showing typical characteristic remanent magnetization directions. (a) Thermal demagnetization behavior of ENE-WSW trending dyke sample I923-1A. (b) Alternating field demagnetization behavior of ENE-WSW trending dyke specimen I923-4B. (c) Thermal demagnetization behavior of NW-SE trending dyke sample I925-22A. (d) Alternating field demagnetization behavior of NW-SE trending dyke sample I925-7B. Solid squares represent projections on the horizontal plane indicated by ‗H‘; open squares represent projection onto a vertical plane indicated by ‗V‘.

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Figure 4-7. Baked contact test on NW-SE trending dyke I925. (a) Thermal demagnetization behavior of the dyke specimen I925-7A (b) Thermal demagnetization behavior of the granitic host rock specimen I925-39A about 1.5 m away from the main dyke showing similar directions as the main dyke.

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Figure 4-8. Rock magnetic results for the Bundelkhand dykes. (a -b) Curie temperature runs for ENE-WSW trending dyke samples I923-5 & 6 in this study. Sample I923-5a shows TcH of 575.1 ºC and TcC of 578.5 ºC and I923-6 shows TcH = 562.1ºC & TcC = 550.5ºC. (c - d) Curie temperature runs for the NW-SE trending older dyke suite samples I925-26 (TcH = 548.2ºC & TcC = 525.5ºC) and I927-12 (TcH = 572.8ºC & TcC = 567.5ºC). (e) Isothermal Remanence Magnetization (IRM) acquisition and back field demagnetization behavior of the two ENE-WSW dyke specimens I923-4B and 8B. (f) Isothermal Remanence Magnetization (IRM) acquisition and back field demagnetization behavior of two NW-SE dyke samples I925-6B and 11B. All samples saturate at about 0.25-0.3 T. Coercivity of remanence values ranged from 0.07 to 0.1 Tesla.

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Figure 4-9. Stereoplots for the three sets of dykes from the Bundelkhand craton. (a) Stereoplot of site mean directions from the 12 NW-SE trending dykes from the older suite (Dec = 155.3º and Inc = -7.8º). (b) Stereoplot of site mean directions from ENE-WSW Mahoba suite of dykes (Dec = 24.7º and Inc = - 37.9º). (c) Stereoplot of site mean directions from NE-SW trending third set of dykes (Dec =189.3º and Inc = 64.5º) with a palaeolatitude of 46.4ºN.

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Figure 4-10. Schmidt projection of virtual geomagnetic poles (VGP) from the Mahoba dykes (this study) and the Majhgawan kimberlite (Gregory et al., 2006) correlated to the mean paleomagnetic pole of the Bhander-Rewa Groups of the Upper Vindhyan Sequence (Malone et al., 2008). The locations of these sites are shown in the inset map of India (red square = Mahoba dykes; magenta square = Majhgawan kimberlite; and green square = Bhander- Rewa).

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Figure 4-11. Paleomagnetically based reconstruction at ~1.1 Ga based on our combined virtually geomagnetic pole (VGP) from ~1113 Ma Mahoba dykes of the Bundelkhand craton (this study); Majhgawan kimberlite pole (~1073 Ma; Gregory et al., 2006) and Bhander-Rewa pole (Malone et al., 2008) and data reported in Pesonen et al. (2003) that places India at low latitudinal position to ~ mid latitudinal positions for Australia (Bangemall sill pole) (Paleolongitudes are unconstrained; see table 3a).

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Figure 4-12. Tectonic and paleogeographic interpretation of the paleomagnetic data at ~2.0 Ga for Indian sub-continent. (a) Generalized tectonic map of India showing the North Indian Block (Aravalli protocontinent) and South Indian Block (Dharwar protocontinent) docked together along Central Indian Tectonic Zone (CITZ). (b) Positions of northern Aravalli-Bundelkhand protocontinent (~2.0 Ga) and southern Dharwar protocontinent (~1.9 Ga) of the Indian peninsular shield based on the paleomagnetic poles from the ~2.0 Ga NW-SE trending Bundelkhand dyke suite, ~1.9 Ma Keskal dykes/Cuddapah dykes of the Bastar/Dharwar cratons and ~1.8 Ga Gwalior volcanics. The figure also shows the APW path for India utilizing the paleopoles from ~ 2.0 Ga – 1.8 Ga.

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Figure 4-13. Paleomagnetically-based reconstruction at ~2.0 Ga (after Pesonen et al., 2003). The Bundelkhand craton is positioned according to paleomagnetic data from the ~1987 Ma NW-SE trending dykes reported in this paper. The reconstruction places India at the equatorial positions with mid latitudinally located Laurentia (Sl: Slave and Wy: Wyoming cratons). The other continental blocks used in the reconstruction are Amazonia, Congo-SF, Kp: Kaapvaal, Fennoscandia, Ukraine, WA: West Africa (Palaeolongitudes are unconstrained; see table 3b)

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Table 4-1. Bundelkhand paleomagnetic results Site/Study Latitude/Longitude N/n Dec Inc Κ α95 VGP VGP dp/dm latitude longitude Older Suite (NW-SE) Site I442 25.430ºN, 78.321ºE 5 141.0º - 9.0º 011.3 07.2° 47.2ºN 325.7ºE 3.7/7.3 Site I443* Dyke A 25.249ºN, 78.398ºE 2 139.0º - 12.6º 070.0 11.1º 46.5ºN 329.8ºE 5.8/11.3 Dyke B 25.249ºN, 78.398ºE 4 141.4º 5.8º 070.0 11.1º 43.2ºS 137.1ºE 5.6/11.1 Dyke C** 25.249ºN, 78.398ºE 2 339.0º - 21.0º 070.0 11.1º 48.5ºS 110.5ºE 6.1/11.7 Site I444 25.243ºN, 78.397ºE 7 174.0º - 14.8º 016.0 15.5º 71.4ºN 277.3ºE 8.1/15.9 Site I454 25.085ºN, 80.026ºE 7 156.5° - 17.3º 050.0 8.6° 62.4ºN 318.2ºE 4.6/8.9 Site I455 24.567ºN, 79.547ºE 18 149.0° - 21.0º 047.3 05.1° 57.6ºN 330.6ºE 2.8/5.5 Site I925 24.568ºN, 79.547ºE 34 151.3° -19.5º 034.0 04.3° 59.1ºN 326.5ºE 2.3/4.5 Site I927 25.010ºN, 80.279ºE 27 167.5° -15.9º 042.0 04.3° 69.3ºN 297.6ºE 2.3/4.4 Site I929 25.113ºN, 80.275ºE 9 168.2° -10.3º 107.0 05.0° 67.1ºN 291.8ºE 2.6/5.1 Site I930 25.112ºN, 80.277ºE 16 164.0º -13.2º 071.3 04.4° 66.1ºN 302.7ºE 2.3/4.5 Site I931 25.179ºN, 79.557ºE 10 152.9° -8.8º 046.2 07.2° 56.8ºN 315.3ºE 3.7/7.3

Mahoba Dykes (ENE-WSW) Site I451 25.173ºN, 79.512ºE 23 014.6º -31.7º 025.0 06.1º 45.3ºS 059.5ºE 3.8/6.9 Site I452 25.173ºN, 79.513ºE 13 027.6º -48.7º 232.0 02.7º 29.1ºS 052.1ºE 2.3/3.6 Site I923 25.114ºN, 79.240ºE 20 014.2º -37.0º 125.0 03.3º 42.2ºS 061.2ºE 2.3/3.9 Site I924 (Great circles) 25.185ºN, 79.564ºE 10 040.7º -29.2º - 02.8º 33.1ºS 031.0ºE 1.7/3.1

NE-SW (third dyke suite) Site I448 25.083ºN, 79.451ºE 20 200.0º 60.0º 042.4 14.3º 21.5ºS 063.3ºE 16.3/21.6 Site I928 25.038ºN, 80.298ºE 20 175.0º 68.0º 048.6 04.7º 13.8ºS 083.5ºE 6.6/7.9

Overall Mean - Older (NW-SE) 25.00ºN,80ºE 141/12 155.3° -7.8º 021.4 09.6° 58.5ºN 312.5ºE 4.9/9.7

Overall Mean – Mahoba (ENE- 25.17ºN,79.5ºE 76/4 024.7º -37.9º 035.9 15.5º 38.7ºS 049.5ºE 9.5/16.3 WSW)

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Table 4-2a. Geochronologic results from the Bundelkhand older suite of dykes (I9GS-13) Ratios Ages 207Pb/ 206Pb/ *207Pb/ 206Pb/ 207Pb/ 207Pb/ % Grain ± 2σ ± 2σ ±2σ ±2σ ±2σ ±2σ 206Pb 238U 235U 238U(Ma) 235U(Ma) 206Pb(Ma) Disc RHO I9GS_13_2 0.1222 0.0006 0.34907 0.0089 5.88 0.15 1930 43 1959 22 1989 8 3 0.98 I9GS_13_6 0.1230 0.0008 0.35774 0.0120 6.07 0.21 1971 57 1985 30 2000 12 1 0.98 I9GS_13_7 0.1218 0.0013 0.36297 0.0111 6.10 0.20 1996 53 1990 28 1983 19 -1 0.95 I9GS_13_8 0.1210 0.0007 0.36787 0.0104 6.14 0.18 2019 49 1995 25 1970 10 -3 0.98 I9GS_13_9 0.1212 0.0006 0.37220 0.0110 6.22 0.19 2040 52 2007 26 1975 9 -3 0.98 I9GS_13_10 0.1220 0.0008 0.36876 0.0094 6.20 0.16 2023 44 2004 23 1985 11 -2 0.97 I9GS_13_11 0.1215 0.0006 0.35601 0.0082 5.96 0.14 1963 39 1970 20 1978 9 1 0.98 I9GS_13_14 0.1218 0.0006 0.35741 0.0099 6.00 0.17 1970 47 1976 24 1983 9 1 0.98 I9GS_13_15 0.1214 0.0006 0.35357 0.0078 5.92 0.13 1952 37 1964 20 1977 9 1 0.98 I9GS_13_16 0.1211 0.0007 0.35077 0.0081 5.86 0.14 1938 39 1955 21 1972 10 2 0.97 I9GS_13_17 0.1213 0.0006 0.35355 0.0089 5.91 0.15 1951 42 1963 22 1975 9 1 0.98 I9GS_13_21 0.1883 0.0008 0.45398 0.0221 11.79 0.58 2413 98 2588 45 2728 7 12 0.99 I9GS_13_23 0.1928 0.0007 0.47373 0.0190 12.59 0.51 2500 83 2649 38 2766 6 10 0.99 I9GS_13_24 0.1939 0.0008 0.46927 0.0187 12.55 0.50 2480 82 2646 37 2776 6 11 0.99 I9GS_13_26 0.2572 0.0020 0.47423 0.0696 16.82 2.47 2502 301 2924 136 3230 12 23 1.00 I9GS_13_27 0.2611 0.0007 0.63422 0.0162 22.83 0.59 3166 64 3220 25 3253 4 3 0.99 I9GS_13_28 0.2615 0.0008 0.56674 0.0566 20.44 2.04 2894 231 3112 95 3256 5 11 1.00 I9GS_13_30 0.1941 0.0007 0.50044 0.0099 13.39 0.27 2616 43 2708 19 2777 6 6 0.98 I9GS_13_31 0.1900 0.0007 0.48971 0.0130 12.83 0.34 2569 56 2667 25 2743 6 6 0.99 I9GS_13_33 0.1864 0.0008 0.48857 0.0105 12.56 0.28 2564 45 2647 21 2711 7 5 0.98 I9GS_13_34 0.1896 0.0012 0.45310 0.0204 11.85 0.54 2409 90 2592 42 2739 10 12 0.99 I9GS_13_35 0.1837 0.0006 0.46585 0.0121 11.80 0.31 2465 53 2588 24 2686 6 8 0.99 I9GS_13_37 0.1852 0.0009 0.47595 0.0122 12.16 0.32 2510 53 2616 24 2700 8 7 0.98 I9GS_13_38 0.1888 0.0016 0.49585 0.0112 12.91 0.31 2596 48 2673 23 2732 14 5 0.98

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Table 4-2b. Geochronologic results from the Mahoba dykes (GDM and GDM1) Ratios Ages 207Pb/ 206Pb/ *207Pb/ 206Pb/ 207Pb/ 207Pb/ % Grain ± 2σ ± 2σ ±2σ ±2σ ±2σ ±2σ 206Pb 238U 235U 238U(Ma) 235U(Ma) 206Pb(Ma) Disc RHO GDM_5a 0.07697 0.0006 0.17605 0.0051 1.868 0.06 1046 28 1070 20 1120 17 7 0.96 GDM_6a 0.07697 0.0007 0.17509 0.0040 1.858 0.05 1041 22 1066 16 1120 18 7 0.93 GDM_7a 0.07671 0.0006 0.17551 0.0041 1.856 0.05 1043 23 1066 16 1114 16 6 0.95 GDM_8a 0.07654 0.0006 0.17623 0.0041 1.860 0.05 1047 23 1067 16 1109 17 6 0.94 GDM_18a 0.07631 0.0006 0.18242 0.0041 1.919 0.05 1081 22 1088 16 1103 16 2 0.94

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Table 4-3a. Paleomagnetic poles at ca. ~1.1 Ga Continent Pole name Age Plat Plong A95 Reference

India Mahoba dykes 1113 ±7 Ma 37.8⁰S 49.5⁰E 15.5⁰ This study Laurentia Keweenawan dykes mean 1099 Ma 36.0⁰N 188.0⁰E 3.0⁰ Swanson-Hysell et al., 2009; + Portage volcanics Pesonen et al., 2003

Baltica Mean Baltica 1093 Ma 01.0⁰N 208.0⁰E 16.0⁰ Pesonen et al., 2003 Australia Bangemall sill pole 1070 Ma 83.7⁰S 129.0⁰E 8.3⁰ Wingate and Giddings, 2002 Kalahari Mean Kalahari 1096 Ma 65.0⁰S 225.0⁰E 5.0⁰ Gose et al., 2006; Pesonen et al., 2003 Siberia Malgina +Linok Formation 1078 Ma 25.0⁰S 231.0⁰E 3.0⁰ Gallet et al., 2000

Table 4-3b. Paleomagnetic poles at ca. ~2.0 Ga Continent Pole name Age Plat Plong A95 Reference India NW-SE Bundelkhand 1987 ± 8.6 Ma 58.5ºN 312.5ºE 9.6⁰ This study dykes Laurentia Minto Dykes 1998 ± 2 Ma 38ºN 174.0ºE 10.0º Buchan et al., 1998 Fennoscandia Konchozero sill, Karelia 1974 ± 27 Ma 14.2ºS 282.0ºE 11.0º Pisarevsky and Sokolov, 1999

Ukraine Gabbro Monzonite + 2000 Ma 53ºN 142.0ºE - Pesonen et al., 2003; Elming et Gabbro diabase al., 2001a Amazonia Oyapok tonalites + 2019 Ma 42ºN 181.0ºE 4.0º Pesonen et al., 2003 Encrucijada pluton Congo-SF Uauá dykes 1983 ± 31 Ma 24ºN 331.0ºE 4.0º D‘Agrella-Filho and Pacca, 1998 Kalahari Vredefort dykes 2023 Ma 27ºN 31.0ºE 7.0º Salminen et al., 2009; Pesonen et al., 2002 West-Africa Liberia M 2050 ± 6 Ma 18ºS 89.0ºE 13.0º Onstott and Dorbor, 1987

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CHAPTER-5 GEOCHRONOLOGIC AND PALEOMAGNETIC CONSTRAINTS ON THE LATE NEOPROTEROZOIC – EARLY PALEOZOIC VOLCANO-SEDIMENTARY UNITS OF THE CENTRAL ASIAN OROGENIC BELT MICROCONTINENTS.

Introductory Remarks

The Eurasian supercontinent is comprised of Europe, Siberia, North China, South

China, Tarim and a multitude of smaller microcontinental blocks dissected by complex orogenic belts (Figure 1a). The Ural-Mongol Fold belt (UMFB) also known as Central

Asian Orogenic Belt (CAOB) is one of the largest orogenic systems within the Eurasian landmass and was the locus of the growth of Eurasia throughout the Neoproterozoic and Paleozoic. The central part of the CAOB has a complex tectonic history and a clear understanding of its early evolution is crucial for the understanding of the amalgamation of the Eurasian supercontinent throughout the Paleozoic.

The Central Asian Orogenic Belt (CAOB) stretches for nearly 10,000 km from the

Urals through Kazakhstan and Tian Shan to Altai and Mongolia to the Okhotsk Sea in the east of Russia (Sengör et al., 1993; Windley et al., 2007; Figure1a). The CAOB is fringed by the Siberian and East European cratons in the north and by the Tarim and

Sino-Korean cratons in the south (Windley et al., 2002, 2007; Figure 1a). The CAOB resulted from a long and complex history of accretionary orogenesis from the late

Proterozoic – Mesozoic (ca. 1000 – 250 Ma; Segnör et al., 1993; Segnör and Natal‘in

1996; Windley et al., 2007) and is composed of a variety of tectonic units, including numerous microcontinental blocks with Precambrian basement, accretionary complexes, ancient island arcs, ophiolites, passive margin sequences and rare ancient gneissic massifs (Figure 1b). It is further characterized by an extended history of subduction (Neoproterozoic to Late Permian-early Triassic) and ‗A-type‘ (Mesozoic to

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Cenozoic) magmatism (Mossakovsky et al., 1993; Sengör et al., 1993; Sengör and

Natal‘in 1996a, b; Kröner et al., 2007; Windley et al., 2007). Late Neoproterozoic to

Early Paleozoic terrigenous clastic or carbonate rocks are now found juxtaposed within these volcanic complexes, accretionary wedges and flysch sequences (Mossakovsky et al., 1993; Jahn et al., 2000; Jahn et al., 2004). The CAOB reaches its maximum width and structural complexity in Kazakhstan (Puchkov, 1997, 2000; Levashova et al., 2009).

The kinematic history of the CAOB is controversial due to myriad conflicting models of tectonic evolution (see Zonenshain et al., 1990; Mossakovsky et al., 1993; Kheraskova et al., 2003; Didenko et al., 1994; Filipova et al., 2001; Şengör and Natal‘in, 1996;

Yukubchuk et al., 2001; 2002; Stampfli and Borel, 2002; Puchkov, 2000; Windley et al.,

2007; Xiao et al., 2010; Levashova et al., 2011a, b).

Figures 2a-e represent sketches of several contrasting models that attempt to describe the evolutionary history of the CAOB microcontinents (also see Xiao et al.,

2010). There are many different schools of thought regarding the origin of the CAOB microcontinents although they can be broadly lumped in two categories. (1) The first group of models regard the CAOB as an ancient analogue of the modern western

Pacific (Zonenshain et al. 1990; Mossakovsky et al. 1993; Didenko et al. 1994;

Kheraskova et al. 2003; Khain et al. 2003; Kuzmichev et al. 2005, Windley et al., 2007;

Xiao et al., 2008; Yang et al., 2011). These authors argue for the existence of an archipelago of scattered microcontinents, island arcs and oceanic basins within the

Paleoasian Ocean in the late Neoproterozoic-early Paleozoic. The Vendian-Early

Cambrian closure of the Paleoasian Ocean via the docking of various blocks to Siberia,

Baltica or to the growing Kazakhstan continent resulted in the formation of major

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portions of the CAOB. (2) Another group of workers advocate the existence of a long- lived volcanic arc system (Şengör and Natal‘in, 1996, 2004; Puchkov, 2000; Yakubchuk,

2008; Yakubchuk et al., 2002; Stampfli and Borel, 2002). The key aspects of these hypotheses and their variants are as follows: (a) Şengör and Natal‘in (1996; 2004) argue for an extensive ―Kipchak-Tuva-Mongol Arc‖ connecting Siberia and Baltica during the interval from 542 Ma – 250Ma. Accretionary wedge complexes were formed as a result of the westward subduction of the oceanic crust under the Kipchak arc.

Successive roll-back of the arc in the Cambrian to mid-Silurian gave rise to the Khanty–

Mansi backarc Ocean. The ocean was closed by Late Carboniferous due to the differential rotation of Siberia and Baltica and the fragments of the ancient arc gradually coalesced into a continent-sized landmass named ―Kazakhstania‖. (b) The CAOB blocks appear to have originated solely from the margins of Siberia and Baltica.

Yakubchuk et al. (2001, 2005) and Yakubchuk (2002, 2004) modified the Kipchak arc model by assuming the occurrence of two parallel island arcs. They emphasized that the strike-slip motion and imbrication of the island arc fragments are responsible for the formation of the Kazakhstania continental block. (d) Puchkov (2000) and Stampfli and

Borel (2002) suggest a rather complicated configuration of the island arc in the early

Paleozoic, but they do not ascribe any importance to strike-slip motions. (e) Jahn et al.

(2000) and Jahn (2004) and Kovalenko et al. (2004) suggested that the initial formation of the CAOB was associated with the breakup of the Rodinia supercontinent by action of the South Pacific superplume.

Despite the existence of myriad models for the Phanerozoic evolution of CAOB, the Neoproterozoic – Cambrian sequences on many of these microcontinents with

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crystalline Precambrian basement bear striking stratigraphic, faunal and lithologic similarities that may provide additional insights into the tectonic history of the CAOB

(Ankinovich, 1962; Zubtsov, 1971; Mossakovsky et al., 1993; Kröner et al., 2007;

Levashova et al., 2011a; Meert et al., 2011a). A majority of the CAOB microcontinents are located in the western part of Kazakhstan and to the south of the Siberian craton

(Figure 1b; e.g. Xiao and Kusky, 2009; Rojas-Agramonte et al., 2010; Meert et al.,

2011a). There are several lines of evidence that support a common origin/association for some of the Precambrian blocks:

(1) Microcontinental domains, including the Kokchetav, Ulutau, North Tian Shan,

Talas-Karatau, Tuva-Mongol and the Baydaric consist primarily of crystalline basement of Archean to Paleoproterozoic age overlain by thick piles of younger Meso-

Neoproterozoic sedimentary and meta-sedimentary sequences (Kozakov et al., 1993;

2007; Degtyarev and Ryazantsev, 2007; Kröner et al., 2007; 2010; Demoux et al., 2009;

Salnikova et al., 2001). (2) Late Neoproterozoic (815-700 Ma) coeval bi-modal volcanic sequences of rhyolitic, trachy-rhyolitic and mafic composition are predominant in several

CAOB microcontinents (Chumakov, 2009a; Levashova et al., 2010, 2011a, b; Meert et al., 2011a). There are contrasting views on the origin of these volcanic complexes.

One group of workers interpret the volcanic units as subduction-related (Mossakovsky et al., 1993, Kheraskova et al., 1995; Levashova et al., 2011a, b; Salnikova et al., 2001) whereas others suggest that all the volcanic complexes are rift related (Ilyin, 1990,

Burasnikov and Ruzhentsev, 1993; Khain et al., 2003). (3) The older Neoproterozoic complexes are overlain by latest Neoproterozoic to early Paleozoic sedimentary sequences and by carbonate-clastic sequences that show significant similarities in their

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stratigraphic, faunal and lithological characteristics (Ankinovich, 1962; Zubtsov, 1971;

Khain et al., 2003; Degtyarev and Ryazantsev, 2007; Chumakov, 2009a; Popov et al.,

2009; MacDonald, 2009; Meert et al., 2011a). (4) The notable occurrence of the late

Neoproterozoic glacial diamictites found at one or more stratigraphic levels on the

CAOB microcontinents (Chumakov, 1978, 2009a; Korolev and Maksumova, 1984;

Lindsay et al., 1996; Meert et al., 2011a; MacDonald, 2009) and phosphorite layers that occur at the Ediacaran-Cambrian boundary on some of these blocks are considered as the marker horizons for correlating these domains (Eganov and Sovetov, 1979; Korolev and Maksumova, 1984; Meert and Lieberman, 2008; Popov et al., 2009). (5) Although limited, paleomagnetic data from the Late Neoproterozoic volcanic rocks and the earliest Paleozoic sedimentary cover sequences on these blocks can be used to place them at similar paleolatitudes (Evans et al., 1996; Kravchinsky et al., 2001; 2010;

Levashova et al., 2010; 2011a, b). Based on the above similarities, Levashova et al.

(2011a) suggested these CAOB microcontinents were part of two larger continent sized

―Kazakhstania‖ and ―Mongol‖ domains (also see Degtyarev and Ryazantsev, 2007 and references therein).

In this paper, we present new U-Pb zircon ages from Late Neoproterozoic felsic volcanic sequences lying unconformably below a Neoproterozoic diamictite deposit at

Greater Karatau (Kazakhstan) and Talas Karatau (Kyrgyzstan). We also report paleomagnetic data from the Early Paleozoic carbonate beds of the Tamdy Series in the

Lesser Karatau microcontinent in southern Kazakhstan and the carbonate sequences within the Tsagaan-Oloom and Bayan-Gol Formations of the Baydaric microcontinental block in central Mongolia. The new radiometric ages and the paleomagnetic directions

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reported in this paper are compared with existing data from other CAOB microcontinental blocks during Late Neoproterozoic-Cambrian to improve our understanding of the origin and tectonic evolution of these microcontinents.

Geological Setting

Lesser Karatau Microcontinent Block

The Lesser (Maly) Karatau microcontinent is a small domain within the larger

Central Asian collage of microcontinents (Figure 1b). It is located 350 km to the southwest of Lake Balkhash in southern Kazakhstan, within the Karatau-Talas folded zone of the Ulutau-Sinian structural belt that extends southeastward from the Ulutau

Mountains of central Kazakhstan into northern China (Sergeev and Schopf, 2010). The stratigraphic order of various units of the Lesser Karatau block from oldest to youngest is: the Kokdzhot, Bolshekeroi, Zhanatass, Koksu, Malokaroi and Tamdy Series (Figure

3; Eganov et al., 1986; Sergeev, 1989; Sergeev and Schopf, 2010). The

Paleoproterozoic schist, slates and phyllites of the Kokdzhot Series and thick terrigenous flysch of the Bolshekeroi Series forms the base of the Lesser Karatau

Range (Figure 3; Sovetov, 1990). The tectonic contacts between the Kokdzhot and

Bolshekeroi Series are poorly exposed and complexly faulted making it difficult to unravel the relative chronology between these two units. Inherited zircons within the

~760 Ma Kurgan tuffs yielded ages equal to ~2.0 Ga or older for the crystalline basement (Levashova et al., 2011a; Meert et al., 2011a). The ~1.5 km thick Koksu

Series disconformably overlies the Bolshekeroi strata throughout the Lesser Karatau range except in the Zhanatass block, and is comprised of alternating conglomerates gravellites and polymict sandstones.

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The basement rocks are overlain by Neoproterozoic-Early Paleozoic volcaniclastic and carbonate-clastic rock cover exposed in a series of repetitive NW trending thrust sheets (Sovetov, 1990). Of these, the Malokaroi Series conformably overlies the Koksu

Group, and is divisible into three Formations with conformable contacts between them:

(a) the lowermost Aktugai Formation (~30-200 m thick) is composed of arkosic sandstones and gravellites; (b) intermediate Chichkan Formation (up to 120 m thick) comprised of fine grained clastic rocks, stromatolitic limestones, cherts and black mudstones; and (c) the uppermost Kurgan Formation (≥500 m thick) predominantly made up of oligomictic and polymictic sandstones interbedded with tuffs and tuffaceous units (Eganov et al., 1986; Sergeev and Schopf, 2010). The volcano-sedimentary

Kurgan rocks are unconformably overlain by carbonate rocks of the Ediacaran-Lower

Ordovician aged Tamdy Series. The Tamdy Series consists of limestones and dolomites, chert, shales, phosphorites and lensoid iron-manganese horizons (Seydalin and Sul‘din, 1977). The Tamdy Series is subdivided into three suites: (a) the lowermost terrigenous and dolomitic Kyrshabakty Suite; (b) intermediate phosphatic and carbonaceous Chuluktau Suite and; (c) the uppermost limestone and dolostone- dominated Shabakty Suite (Figure 3). Direct radiometric age constraints are lacking for most of the Tamdy Series with the exception of the Chuluktau and Shabakty Suites where several distinct fossil assemblages provide Cambrian-Ordovician ages (Eganov et al., 1986).

The lowermost Kyrshabakty Suite rests unconformably over the underlying Kurgan

Formation and has a variable thickness ranging from a few meters to over 150 meters.

The basal unit of the Kyrshabakty Suite is a conglomeratic sequence made up of poorly

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rounded tuffaceous clasts from the underlying Proterozoic rocks in a mudstone- sandstone sized matrix. Eganov and Sovetov (1979) described this conglomerate unit as a ‗tilloid‘ and Meert et al. (2011a) confirmed a glacial origin for this unit. The tillite is capped by a widespread 10-12m thick pink dolomite that is further overlain by a terrigenous and carbonate mixed sequence. The upper part of the Kyrshabakty Suite appears to be conformable with an 8-60m thick unit, variously referred as ―Lower

Dolomite‖ or the ―Berkuty Member‖ of the Chuluktau Suite; however, in other cases the

Kyrshabakty is absent or the conformable nature of the contacts is difficult to ascertain

(Eganov et al., 1986; Popov et al., 2009; Sergeev and Schopf, 2010; Meert et al.,

2011a). Chumakov (2009) considered the basal tilloid member overlying the Kurgan

Formation of the Lesser Karatau Range analogous to the ―Baykonurian‖ glaciations in

Kazakhstan and Kyrgyzstan and assigned a tentative Late Vendian-Nemakit-Daldynian age to this glaciation (~550-540 Ma; Eganov and Sovetov, 1979; Chumakov, 2010). In contrast, Meert et al., (2011a) argue for a Late Cryogenian (~635 Ma) age for this glacial unit corresponding to the Nantuo (South China), Ghaub (Namibia) and Marinoan

(Australia) glaciations based on chemostratigraphic correlation.

The middle part of the Tamdy Series, the Chuluktau Suite is typically a few tens of meters thick phosphate-siliceous layer, that contains small-shelly fossils of the

Protohertzina anabarica zone in the so called ‗Berkuty Member‖ (Eganov et al., 1986;

Popov et al., 2009), Pseudorthoteca costata in the phosphorite zone (Tommotian stage;

534 - 530 Ma), Rhombocorniculum cancellatum and Bercutia cristata in the uppermost carbonate strata of the Chuluktau (Atdabanian stage; 530 - 524 Ma). The uppermost

Botomian-Amgan Stage (524 – 502 Ma), thick carbonate sequence unconformably

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overlying the Chuluktau Suite is referred as the Shabakty Suite that contains massive dolomites with characteristic stromatolitic and argillaceous structures and trilobite fossils including Hebediscus orientalis, Ushbaspis limbata, Redlichia-chinensis and Kootenia gimmerlfarbi (Mambetov and Repina, 1979; Eganov and Sovetov, 1979; Popov et al.,

2009).

The Karatau Range demonstrates a complex tectonic history of folding and thrusting, particularly between Neoproterozoic and Cenozoic times. The family of fault systems in the Karatau range includes a series of parallel faults including the Main

Karatau fault and Bolshoi Karoi fault. The Main Karatau fault is a continuation of the northwestern Talas-Fergana fault (southern Kyrgyzstan) originated in Mid- to Late

Ordovician and is consistent with major compressional deformation resulted from a collision to the NE of Karatau (Figure 4; Alexeiev, 1998).

There are two different phases of folding observed in the Karatau Range during

Paleozoic time (Allen et al. 2001). The first phase of folding is related to one of two orogenies taking place in adjacent areas of Central Asia in the late Paleozoic. The collision between the northern, passive margin of the Tarim microcontinent with the southern margin of Asia began in the latest Devonian in SW China and produced an orogen preserved in the southern parts of modern Tian Shan block (Allen et al., 1993).

In the south, subduction under the Kazakhstan continent that began in the mid-

Carboniferous led to the convergence and collision of Kazakhstan with the Turan

(Karakum) and Tarim continents (Windley et al., 2007). The regional termination of the carbonate sedimentation across Kazakhstania is Late Carboniferous in age and so is the most of the collision related deformation in the western Tian Shan block (Burtman,

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1975). An alternative cause of the first phase of the Late Paleozoic deformation is the tightening of the Kazakhstan Orocline as a result of the opposing motion of the Siberia and Tarim continents indicated by both structural (Zonenshain et al., 1990; Şengör and

Natal‘in, 1996) and paleomagnetic data (Levashova et al., 2003). The second phase of folding to affect the Upper Paleozoic strata in the Karatau Range may correlate with the

Late Permian dextral strike-slip deformation identified to the southeast on the Talas-

Fergana fault and in much of the Kazakhstan, Baltica, Siberia and northwest China caused by differential movements and rotations of the main continental blocks during the collisional and post-collisional times (Burtman, 1980; Buslov et al., 2003).

The Early Carboniferous and Permian was also the time of major metallogenesis in Kazakhstan and Central Asia. Most Cu-Mo porphyry, Ag-Sb and Sn-W mineralization, skarn and epithermal deposits of Kazakhstan were formed during this time (Heinhorst et al., 2000; Pavlova and Borisenko, 2009). During the Late Permian

(~281 Ma; K-Ar age; Andreeva et al., 1984) there was extensive Ag-Sb, Sn-Ag and Sn- sulfide mineralization in the Talas – Karatau (Babakhan and Kumyshtag ore clusters) and Lesser Karatau (Kurgan cluster) blocks (Pavlova and Borisenko, 2009) of central

Kyrgyzstan and southern Kazakhstan, respectively. This Late Permian metallogenic event correlates with the time of alkaline mafic magmatism, volcano-plutonic series of trachybasalt, trachyte, shoshonite and rhyolites. Magmatic arc rocks in many cases unconformably overlie the previously amalgamated tectonic units, and Late

Carboniferous to Permian granites of various types is responsible for generation of the bulk of the continental crust in the area (Mossakovsky et al., 1993; Heinhorst et al.,

2000).

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Baydaric Microcontinent Block

The Baydaric (Dzabkhan) microcontinent located in Central Mongolia is bounded by the Late Neoproterozoic-Cambrian subduction-related complexes of the Lake

(Ozerny) tectonic zone (527-522 Ma; Kovalenko et al., 1996) to the west and south and by Bayankhongor Neoproterozoic ophiolites (Figure 5; 665 ± 15 Ma; Kovach et al.,

2005) to the north and east, respectively. It is separated from the Tuva-Mongolian microcontinent in the north by the Balnay fault (Figure 5; Levashova et al., 2010 and the references therein).

The Archean – Paleoproterozoic basement of most of the Baydaric block is represented by two highly metamorphosed complexes namely, Baidaragin and

Bumbuger (Figure 6; Kozakov et al., 2007). The structural relationship between the two complexes is uncertain; however, on the basis of Nd model ages, Kozakov et al. (1997) speculated that the Bumbuger metasediments were most likely derived from the

Neoarchean Baidaragin complex. The basement rocks, the Bumbuger and the

Baidaragin complexes are intruded by Paleoproterozoic granitic rock bodies at ~1825 ±

5 Ma (U/Pb zircon age; Kotov et al., 1995) and experienced high grade metamorphism at 1800-1850 Ma (Khain et al., 2003). The basement in the eastern part of the microcontinent is overlain by the Neoproterozoic carbonates and terrigenous sediments of the Ulzitgol Formation. The Dzabkhan Volcanics (~2000 m in thickness) overly the crystalline basement and contain andesitic-basalts, andesites and felsic volcanic rocks with numerous pyroclastic tuffs and ignimbrites (Khomentovsky and Gibsher, 1996;

Levashova et al., 2010). The reported U-Pb zircon ages for the Dzabkhan volcanics range between ~805-770 Ma (Levashova et al., 2010).

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The late Neoproterozoic – early Cambrian sedimentary sequences overly the

Dzabkhan Formation with both erosional and angular unconformity in the Dzabkhan river valley section (Figure 6; Khomentovsky and Gibsher, 1996; MacDonald et al.,

2009; Levashova et al., 2010). The sedimentary units (in ascending order) are represented by the Vendian to Nemakit - Daldynian sequences of the Tsagaan Oloom

Formation (620 – 543 Ma) and the clastic-carbonate rocks of the Nemakit – Daldynian to Atdabanian (543 – 520 Ma) Bayan Gol and Salaany Gol Formations (Figure 6;

Khomentovsky and Gibsher, 1996). Macdonald et al. (2009) describes the lowermost part of the ~1500m thick Tsagaan Oloom Formation by 6-500m thick Maikhan Ul member that is composed of tillites and clastic sedimentary rocks (Figure 6a). The

Maikhan Ul member is overlain by ~570 m thick Tayshir Member composed of limestone, limestone marl, rhythmites and massively bedded grainstone units with microbial laminites and giant ooids. In the Khongoryn and Tsagaan blocks, a ~20 m thick Khongoryn diamictite layer unconformably overlies the Tayshir Member. The

Khongoryn diamictite layer is capped by the buff to pink colored "Ol" carbonate that is composed of largely recrystallized micropeloidal dolostone, rhythmites and grainstone layers (Figure 6a). The middle member of the Tsagaan Oloom Formation is Ulaan

Bulagyn member ranging in thickness from 100m to over >500m of massive limestones and dolostones. The uppermost Zunne Arts Member (~250m thick) is composed of columnar stromatolites (Boxonia grumulosa), phosphatized shales, phosphorites, dolomites and limestone rhythmites with characteristic ichnofossils (Phycodes pedum,

Rusophycus cf. avalonensis; Figure 6a; Goldring and Jensen, 1996).

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The Early Cambrian Bayan-Gol Formation (~1000m thick; Figure 6b) conformably overlies the Tsagaan-Oloom Formation and is represented by terrigenous sediments mainly sandstones and siltstones interbedded with clay-rich limestones (Khomentovsky and Gibsher, 1996). The Bayan Gol Formation is characterized by rich diversity of ichnogenera (Phycodes pedum, Helminthoida cf. miocenica, Palaeophycus tubularis,

Rusophycus cf. avalonensis, Monomorphichnus, Treptichnus Pbifurcus, Cochlichnus ,

Hormosiroidea, Planolites, Didymaulichnus miettensis; Goldring and Jensen, 1996;

Lindsay et al., 1996), small shelly fossils (Anabarites tripartitus , Protohertzina unguliformis, Angustiochrea magna, Palaeosulcachites biformis, Purella cristata, Purella panda, Tommotia applanata, Salanacus voronovi, Salanacus cristatus, Dokidocyathus regularis, Lapworthella tortuosa, Khairkhania rotata, Torellella curvae, Ovalitheca mongolica, Sachites proboscideus; Khomentovsky and Gibsher, 1996) and calcimicrobial patch reefs (Kruse et al., 1996). The characteristic small shelly fossils of the Bayan Gol Formation are thought to correlate well with the Nemakit – Daldynian

(Anabarites and Purella antiqua Zones) to Early Tommotian (Nochoroicyathus sunniginicus and Dokidocyathus regularis Zones) biozones in Siberia (Braiser et al.,

1996; Khomentovsky and Gibsher, 1996). The terrigenous rocks of the Bayan Gol

Formation are overlain by an additional 400m thick Atdabanian Salaany Gol Formation

(525 – 520 Ma) composed of massive carbonate sequences with archeocyathid reefs

(320 m; Voronin et al., 1982).

The oldest deformation in southwestern Mongolia is Early Cambrian in age (~573-

540 Ma; 40Ar/39Ar mica ages) and is related to the northward thrusting of an eclogite mélange of the Tsakhir Uul and Khantaishir ophiolite over the Lake Zone basement

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(Figure 5; Levashova et al., 2010; Štipská et al., 2010). The accretionary complex made up of Tsakhir Uul and Khantaishir formations obducted on the Lake Zone may represent the suture zone of a small oceanic domain closed during Early Cambrian times (Kröner et al., 2010). Thrusting and faulting also affected the late

Neoproterozoic–early Cambrian sedimentary cover and formed a series of nappes

(Khomentovsky and Gibsher, 1996).

A second phase of deformation in Mongolia is represented by a Late Devonian –

Early Carboniferous (~365-330 Ma; 40Ar/39Ar muscovite ages) ―terrane‖ accretion represented by the E-W directed emplacement of the Chandman arc magmas in the

Gobi-Altai Zone (Figure 5; Lehmann et al., 2010). The tectonic emplacement of ophiolites (nappe stacking) in the Trans-Altai Zone over Devonian sediments occurred during late Devonian to Early Carboniferous E-W shortening (Lehmann et al., 2010).

The Middle-Late Devonian (395-360 Ma) interval is also known for the collision of Altai-

Mongolia (AM) and Gorny-Altai (GA) terranes along the Charysh-Terekta-Ulagan-Sayan suture-shear zone (CTUSs), followed by the emplacement of voluminous granitoids

(Glorie et al., 2011 and references therein). According to Glorie et al. (2011), this collision led to large-scale Late Devonian – Early Carboniferous strike-slip displacements along the CTUSs followed by the terminal closure of the Paleo Asian

Ocean (PAO) at Late Carboniferous – Permain boundary after the collision of Baltica,

Siberia and Paleo-Kazakhstan (~295 Ma).

Epithermal ore mineralization events related to Sn-W and Co-As deposits are of common occurrence in the Yustid cluster (Gorny-Altai) of SW Mongolia in middle

Paleozoic (Figure 5; ~355 Ma; 40Ar/39Ar muscovite ages; Pavlova and Borisenko,

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2009). These metallogenic events are genetically linked to the Early Carboniferous alkaline mafic and (355 ± 8.6 Ma) granitoid intrusions of the Yustid complex

(Gonverdovskiy and Rudnev, 2000). Early Carboniferous time also coincides with the terminal collision between the Tuva-Mongolian block and the Siberian block and related deformation (Berzin et al., 1994).

The third phase of tectonic deformation in the Baydaric microcontinent is related to the Permian to Late Triassic N-S shortening responsible for steep and tight folding of all the Lake Zone units; including the Carboniferous and Permian sedimentary cover sequences and lower grade metamorphism of the Khantaishir ophiolites in the Gobi-

Altai Zone (Lehmann et al., 2010). These Permian tectonothermal events (290 – 270

Ma) concentrated in two major E-W trending belts (Gobi - Altai and Gobi – Tian Shan bimodal complexes) produced crustal melting and migmatites via major crustal scale thermal perturbations within the region (Kovalenko et al., 1995). The Late Permian and

Triassic ore mineralization events (~266 Ma – 235 Ma) produced siderite – arsenite, native bismuth, pyrrhotite, chalcopyrite, cobaltite and scheelite deposits in Mongolia

(Borisenko et al., 1984; Pavlova and Borisenko, 2009). Most of these epithermal mineralizing events are linked to the felsic magmatic plutons, hydrothermal veins and sedimentary rock formations. The last phase of tectonism in the Baydaric microcontinent is linked to the beginning of the closure of the Mongol-Okhotsk Ocean in

Late Triassic – Early Jurassic period (220 – 190 Ma; Zonenshain et al., 1990).

Greater Karatau Microcontinental Block

The Greater Karatau microcontinent commonly known as Bolshoi Karatau lies in the south-west part of the Karatau ridge in southern Kazakhstan (Figure 1b).

Lithostratigraphically, Greater Karatau consists of shallow marine carbonates and rift-

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type volcanic rocks of Riphean age, Vendian shales and Early Paleozoic deep marine carbonates, cherts and siliciclastic turbidites (Abdulin et al., 1986; Popov et al., 2009;

Chumakov, 2009). The stratigraphic order of various units of the Greater Karatau block from oldest to youngest is: The Barkyly and Kaynar suites, Ulutau Group, Kurumsak

Formation and Kokbulak Formation (Ankinovich, 1962; Seydalin and Sul‘din, 1977;

Chumakov, 2009; Figure 7). The lowermost Bakyrly and Kaynar suites are made up of volcano-sedimentary and intrusive igneous rocks (Ankinovich, 1962; Chumakov, 2009).

The U-Pb zircon ages known for the volcanic rocks correspond to 690 ± 15 Ma and 705

± 10 Ma (Sudorgin, 1990; Korolev and Maksumova, 1984) whereas the youngest of the intrusive rocks are dated at 720 ± 20 Ma (U-Pb zircon age; Kiselev, 2001).

Unconformably overlying the Barkyrly suite is the Vendian (~700 Ma) Ulutau Group.

The Ulutau Group consists of the Rang, Kokshok and Karagur, Aksyumbe and

Baykonur Formations containing conglomerates, tillites, sandstones, siltstones, carbonaceous - clay and chlorite - clay shales, finely laminated limestones and dolomites (Chumakov, 2009). The uppermost formation of the Ulutau Group is the Late

Vendian to Early Cambrian Baykonur Formation consisting of glacial conglomerates, overlain by shales, sandstones and dolomites. The Kurumsak and Kokbulak

Formations, consisting of vanadium bearing shale units overlie the Baykonur Formation with transgressive unconformity (Seydalin and Sul‘din, 1976; Chumakov, 2009). A

Nemakit-Daldynian to Botomian age (544 – 509 Ma) has been assigned to the

Kurumsak and Kokbulak Formations based on the occurrence of microfossils and trilobites within the rocks.

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The lithologic units of the Greater Karatau and Lesser Karatau blocks are correlated based on the geochronologic, chemostratigraphic, biostratigraphic, and geologic similarities (Eganov and Soveotv, 1979; Kholodov and Paul, 1999; Chumakov,

2009). The bimodal magmatic rocks of the basal Barkyly and Kaynar Suites in Greater

Karatau (705 – 720 Ma; Korolev and Maksumova, 1984) are roughly coeval to the

Neoproterozoic Kurgan Formation (~766 Ma; Levashova et al., 2011a) of the Lesser

Karatau microcontinent. The Neoproterozoic to Early Paleozoic thick carbonate-clastic sequences overlying these volcanic units show striking similarities and strengthen the geological correlations between the two blocks (Ankinovich, 1961; Zubtsov, 1971).

Among these, the marker horizons are the Late Neoproterozoic glacial diamictites

(Chumakov, 1978; Korolev and Maksumova, 1984; Chumakov, 2009; Meert et al.,

2011a) and the phosphorite layers (Korolev and Maksumova, 1984; Meert and

Lieberman, 2008). The late Vendian to Early Cambrian glacial Baykonur Formation of

Greater Karatau has been correlated with the Aktas tillite at Lesser Karatau (Eganov and Sovetov, 1979; Chumakov, 2009). The tentative age of this glacial horizon was recently updated to Marinoan (650 Ma) by Meert et al. (2011a).

The vanadium-bearing siliceous shale sequence of the Kurumsak Formation of

Greater Karatau is correlated with the phosphorite bearing Chuluktau Formation of

Lesser Karatau (Kholodov and Paul, 1999). Similarly, the trilobite zones within the

Upper Cambrian carbonate rocks of the Kokbulak and Kamal Formations of the Greater

Karatau described by Ergaliev (1983) closely resemble the fauna of eastern Tian Shan and Upper Cambrian Shabakty Formation of the Tamdy Series of Lesser Karatau

(Popov et al., 2009).

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Talas Karatau Microcontinental Block

The Talas Karatau zone is located to the south of Greater Karatau microcontinent and within the Tian Shan mountain range, Kyrgyzstan (Figure 1b). Major exposures of

Precambrian and early Paleozoic rocks occur within thrust sheets (Alekseev, 1997;

Khudolei and Semiletkin, 1992). The three thrust-bounded regions are known as the

Uzunakhmat (SW) and Karagoin (NE) separated by the central Kumyshtak volcano- sedimentary sequence (Figure 8; Korolev and Maksumova, 1980; Kiselev and Korolev,

1981; Abad et al., 2003). The Uzunakhmat thrust sheet contains a 3- km thick sequence of (oldest to youngest) the Bakair, Karabura and Uzunakhmat formations that are thought to be of Lower to Middle Riphean age (1100 – 1600 Ma; Figure 8; Kiselev and Korolev, 1981). The lithology of the Uzunakhmat sheet consists of ﬂysch sediments that underwent dynamic greenschist metamorphism. The lower part of the

Uzunakhmat rock sequence has a carbonate-terrigenous composition and the upper part is wholly terrigenous.

The middle to upper Riphean Karagoin thrust sheet contains the Karagoin Group that is composed (oldest to youngest) of the Chondzol, Tagyrtau, Chydygolot, Birbulak,

Urmaral, Chokutash and Kyzylbel Formations (Figure 8; Korolev and Maksumova,

1980). The lower stratigraphic levels of Karagoin group include sandstones and limestones of the Tagyrtau and Chondzol formations. The rocks of Tagyrtau are conformably overlain by the ~ 2 km thick sequences of carbonates and rhythmic units of terrigenous flysch sediments of Chydygolot, Birbulak, Urmaral, Chokutash and Kyzylbel

Formations. The rhythmic units are composed of gravellites, calcareous sandstone and siltstone, calcareous shale, slates and limestone.

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The middle Kumyshtag thrust sheet consists of coarse proximal flysch sediments of Neoproterozoic time accumulated in the continental rift basin in the region. The flysch deposits are interrupted by the accumulation of fine-grained carbonate- terrigenous deposits and trachyte, trachyandesite, and dacite tuffs. The rhyolite and trachyrhyolite tuffs of this unit were dated to 680 ± 20 Ma by U-Pb method (Kiselev,

2001). All the aforementioned formations in the thrust sheet discussed above are unconformably overlain by Ordovician limestones within the Talas Ala Tau block (Figure

8; Abad et al., 2003).

Previous Studies

Paleomagnetic Studies

Paleomagnetic studies on the evolution of the CAOB have concentrated on the relatively undeformed and low metamorphic grade Neoproterozoic to Paleozoic sections

(Evans et al., 1996; Alexyutin et al., 2005; Kravchinsky et al., 2001; 2010; Van der Voo et al., 2006; Wang et al., 2007; Abrajevitch, 2007; 2008; Bazhenov et al., 2008;

Levashova et al., 2003; 2007; 2010; 2011a, b). The available paleomagnetic data from

Mongolian Neoproterozoic to early Paleozoic units come from the studies that were conducted on the Tommotian-Atdabanian Bayan Gol and Salaany Gol Formations

(Evans et al., 1996 and Kravchinsky et al., 2001), the Vendian – Nemakit-Daldynian

Tsagaan-Oloom Formation (Kravchinsky et al., 2001), the Ediacaran - Early Cambrian

East Sayan rocks from Tuva-Mongolia terrane (Kravchinsky et al., 2010) and the

Neoproterozoic Dzabkhan volcanic sequence, from the Baydaric microcontinent

(Levashova et al., 2010). Levashova et al. (2011a) also report paleomagnetic data from the ~766 Ma Kurgan Formation of the Lesser Karatau microcontinent in southern

Kazakhstan.

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The limestone units of the Early Cambrian Bayan-Gol (and Salaany Gol)

Formation of the Baydaric block were studied paleomagnetically by Evans et al. (1996).

They reported a positive fold test for the carbonates of the Salaany Gol formation but the exact timing of the magnetization was ambiguous. A depositional paleolatitude of

44° was computed for the Salaany Gol formation and was compared to the APWP of the Siberian craton. The age of the magnetization in that study was interpreted to date to either (a) just prior to the accretion of Dzabkhan basin to the Siberian craton or (b)

(more likely) post accretion during Silurian-Devonian time. In a second study of the

Bayan Gol and Tsagaan Oloom Formations, Kravchinsky et al. (2001) isolated an intermediate temperature magnetization B-component and a dual-polarity high temperature magnetization C-component, both directions distinct from the directions described by Evans et al. (1996). The B-component of Kravchinsky et al. (2001) yielded a positive fold test and a paleolatitude of 68° N or S corresponding to a tilt-corrected inclination of -79°. This B-component was interpreted as an Early Carboniferous or Late

Permian remagnetization direction based on a comparison to the reference Siberian

APWP. The dual-polarity, low-inclination high temperature C-component indicated a nearly equatorial position for the Baydaric microcontinent. It is important to note that the paleolatitude (~44°) generated for the high temperature magnetization component of the

Salaany Gol Formation determined by Evans et al. (1996) lies quite close to the great circle joining the C and B paleopoles of Kravchinsky et al. (2001). Kravchinsky et al.

(2001) interpreted this relationship as the product of incomplete separation between a primary high temperature component and a remagnetized intermediate temperature component. More recently, Kravchinsky et al. (2010) reported paleomagnetic results

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from the Eastern Sayan rocks of the Tuva-Mongolia terrane. In addition to a predominant Paleozoic overprint from the Dunzhugur Formation, they obtained a magnetization direction (interpreted as primary) from the sedimentary sequences of the

Bokson Formation that was supported by the fold and reversal tests. If both of these low inclination results are primary, then the directions support tropical paleolatitudes for both the Baydaric and Tuva-Mongolia blocks during Ediacaran – Early Cambrian time.

Late Neoproterozoic felsic and bimodal volcanic units were also studied for from the Baydaric and Lesser Karatau microcontinents. Levashova et al. (2010) constrained the position of the Baydaric domain to the north of Siberia, during 770-805 Ma at a paleolatitude of 47 ± 14 ºN/S, based on the paleomagnetic results from the Dzabkhan volcanics in Central Mongolia. They further suggested an association of the Baydaric microcontinent with either India; South China or Tarim plates during Late

Neoproterozoic time (also see Levashova et al., 2011b).

Among the Kazakhstanian microcontinents, the only reliable Neoproterozoic-early

Paleozoic paleomagnetic study comes from the Lesser Karatau block in central

Kazakhstan. Levashova et al. (2011a) studied the ~770 Ma volcano-sedimentary units of the Kurgan formation of the Lesser Karatau domain for paleomagnetism and geochronology. A paleolatitude of 34.2 ± 5.3º N or S was reported for Lesser Karatau.

Levashova et al. (2011b) compared the paleolatitudes of the Lesser Karatau and

Baydaric blocks and reviewed the tectonostratigraphic correlations between CAOB microcontinents and other major cratons. Using paleomagnetic and geologic filters these authors postulated Tarim, India or South China as the possible candidates for the origin of both the Kazakhstan and Mongol domains.

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Geochronological Studies

Kröner et al. (2007) demonstrated that many of the CAOB microcontinents contained a crystalline basement of Archean to Paleoproterozoic age. Those microcontinents included the Baydaric, Tian Shan, Ulutau, Kokchetav, Karatau-Talas and Chu Illi microcontinents (Figure 1b). Kozakov et al. (1993) reported an upper intercept U/Pb zircon age of 2646 ± 45 Ma for the tonalitic-gneissic basement of the

Baydaric Group of the Baydaric microcontinent. The Baydaric Group of rocks is intruded by 1825 ± 5 Ma Paleoproterozoic granitic intrusions (Kotov et al., 1995). A Rb-

Sr whole rock isochron age of 3153 ± 57 Ma and a U/Pb zircon age of ~2000 Ma are reported for the Early Precambrian Gargan block of the Tuva-Mongolia microcontinent

(Badarch et al., 2002). The gneissic basement of the central part of the Tuva-

Mongolian block has a protolith 207Pb/206Pb zircon age of 1868 ± 3 Ma (Badarch et al.,

2002). The Chu-Illi microcontinent in southern Kazakhstan (Figure 1b) has an Archean basement of granitoid gneisses, granites and foliated amphibolites. The U-Pb SHRIMP ages on magmatic zircons of a granite-gneiss sample from Uzunbulak area of the Chu

Illi block yielded an upper concordia intercept age of 2791 ± 24 Ma, interpreted to be the emplacement age of the protolith (Figure 1b; Kröner et al., 2007).

The Kokchetav microcontinent existed as an isolated continental block within the

Paleoasian Ocean in Late Precambrian and Cambrian time (Figure 1b). Its continental crust is made up of a granite gneiss basement overlain by carbonate-terrigenous sediments. The oldest basement complexes date back to the Paleoproterozoic (2.6 –

2.0 Ga), based on the U-Pb geochronology on the metamorphic zircons extracted from the quartz-garnet micaceous schists of the Berlykskaya Suite of the Kokchetav block

(Shatagin et al., 2001). Zircons from the biotite-amphibole bearing schist from the

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Kuilyu suite of the Middle Tian Shan domain yielded a 207Pb/206Pb isochron age of

2612 ± 18 Ma (Kiselev et al., 1993). A U/Pb zircon age of 2570 ± 50 Ma on these micaceous schists is consistent with the above mentioned 207Pb/206Pb ages (Kroner et al., 2007 and references therein).

Table 1 lists the existing radiometric ages of Neoproterozoic felsic (dominantly rhyolitic and trachyrhyolitic) and bimodal volcanic units from many of these CAOB microcontinental blocks (Figure 1b). These include the late Neoproterozoic volcanic rocks of the Koksu series from the Ulutau microcontinent, the Dzabkhan volcanic series from the Baydaric block, the Kaynar Formation from the Greater Karatau, the Kurgan

Formation from the Lessser Karatau, the Kopa Formation from the Chu-Illi microcontinent, the Bolshoy Naryn Formation from the Central Tian Shan, and the

Altynsyngan Formation from the Aktau-Mointy and Jungaar domains (Figure 1b; table

1).

Levashova et al. (2011a) conducted a geochronologic and paleomagnetic investigation on the Kurgan Formation of the Lesser Karatau microcontinental block in central Kazakhstan. The U/Pb zircon analysis of the rhyolite and tuff sequences at

Lesser Karatau yielded concordant ages of 831 ± 15 Ma and 766 ± 7 Ma. The age of the tuff sequence is interpreted to be 766 ± 7 Ma whereas the rhyolite sample yielding

831 Ma is interpreted as a reworked tuff with detrital zircons (Levashova et al., 2011a;

Meert et al., in 2011a).

The isotopic ages known from the volcanic rocks at the base of the Baykonur

Formation of the Ulutau microcontinental block in central Kazakhstan are 690 ± 15 Ma and 705 ± 10 Ma (U/Pb zircon ages; Sudorgin, 1992; Korolev and Maksumova, 1984).

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Another U/Pb zircon age of 720 ± 20 Ma is known for the youngest of the intrusive rocks within the complex (Kiselev, 2001). The Baykonur Formation is the upper subdivision of the Ulutau Group in central Kazakhstan and of Dzhetym Group in Kyrgyzstan and overlies the older volcano-sedimentary complex and intrusive rocks. Chumakov, (2009) correlated the Baykonur Formation of the Ulutau mountains in central Kazakhstan,

Greater Karatau in southern Kazakhstan and Dzhetym Group in central Kyrgyzstan based on the chemostratigraphic, geologic and paleontologic similarities.

The basement rocks on some of the CAOB microcontinents are younger. The available U/Pb zircon age for the basement volcanic rocks of the Aktau-Mointy microcontinent is 800 ± 11 Ma (Kozakov et al., 1993). Similarly, the U-Pb zircon ages known for the volcanic rocks of the Greater Karatau basement correspond to 690 ± 15

Ma and 705 ± 10 Ma (Sudorgin, 1990; Korolev and Maksumova, 1984). The youngest of the intrusive rocks in Greater Karatau block are dated at 720 ± 20 Ma (U/Pb zircon age; Kiselev, 2001). However, the young zircons found in the basement rocks do not preclude older rocks being present in these domains.

The 206Pb/238U and 207Pb/235U analysis of the rhyolite samples taken from the upper and lower sections of the Dzabkhan volcanic near Tsagaan-Gol river constrains the age of the Dzabkhan volcanics between ~805-770 Ma (Levashova et al., 2010).

Additional support for the Late Neoproterozoic age for the volcanic sequences of the

Baydaric microcontinent is provided by U/Pb SHRIMP zircon ages from the ash flow tuffs within the Dzabkhan Formation of 777 ± 6 Ma (Zhao et al., 2006). The Greater

Naryn Group of the Central Tian Shan (CTS) microcontinent, southern Kazakhstan is made up of Neoproterozoic felsic and subordinate mafic volcanics that have been dated

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at 830 ± 20 Ma (Kiselev and Maksumova, 2001), 764 ± 4 Ma (Kröner et al., 2009) or

~830 Ma (Jourquin, 2010) and are interpreted as a former marginal Andean type magmatic belt of this continent originated from Eastern Gondwana (Bakirov and

Maksumova, 2001; Biske and Seltmann, 2010; Jourquin, 2010). The similarities of the

Neoproterozoic acid volcanic and tillite formations and Lower Paleozoic cover within the

CTS implies that outcrops of Precambrian basement favors a common association with the oher microcontinental blocks (Zubtsov et al., 1974; Windley, et al., 2007; Alekseev et al., 2009).

Methods and Sampling

U-Pb Zircon Geochronology

Several samples from the Greater Karatau (Kazakhstan) and Talas Karatau

(Kyrgyzstan) microcontinents were collected for U/Pb geochronological analysis. Out of these, sample GK6-3B is a rhyolitic tuff (42.42°N, 70.22°E) and was collected from the base of the Neoproterozoic section at Greater Karatau. Sample KT6-1B was taken from a volcanic sequence (42.18°N, 71.46°E) of Neoproterozoic Kumyshtag unit of the Talas

Karatau microcontinental block (Figure 1b). Standard gravity and magnetic separation techniques were used to separate zircon grains from pulverized samples at University of

Florida. The samples were first crushed, then disk milled and sieved to < 80 μm grain size. The fractions were then rinsed using Calgon (an anionic surfactant) followed by water table treatment with slow sample feed rates. This was followed by heavy liquid

(TBE and MEI) mineral separation with multiple agitation periods to reduce the number of entrapped grains in the lower density fraction. Finally, the sample was repeatedly passed through a Frantz isodynamic magnetic separator using a current of 1.0 amp (2-

4° tilt). Approximately 20–25 fresh looking (clear), euhedral to nearly anhedral zircon

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grains were handpicked from the least magnetic fraction of the two samples GK6-3B and KT6-1B under an optical microscope to ensure the selection of only the clearest grains and fractions of grains. Further hand-picking of the grains reduced the number to

10-15 grains. The zircons were mounted in resin and then polished to expose median sections. Further sonication and cleaning of the plugs in 5% nitric acid (HNO3) helped to remove any common-Pb surface contamination.

U–Pb isotopic analyses were conducted at the Department of Geological Sciences

(University of Florida) on a Nu Plasma multicollector plasma source mass spectrometer equipped with three ion counters and 12 Faraday detectors. The LA-MCICP-MS is equipped with a specially designed collector block for simultaneous acquisition of 204Pb

(204Hg), 206Pb and 207Pb signals on the ion-counting detectors and 235Uand 238U on the Faraday detectors (see Simonetti et al., 2005). Mounted zircon grains were laser ablated using a New - Wave 213nm ultraviolet beam. During U–Pb analyses, the sample was decrepitated in a He stream and then mixed with Ar-gas for induction into the mass spectrometer. Background measurements were performed before each analysis for blank correction and contributions from 204Hg. Each sample was ablated for ~30 s in an effort to minimize pit depth and fractionation. Data calibration and drift corrections were conducted using the FC-1 Duluth Gabbro zircon standard. Data reduction and correction were conducted using a combination of in-house software and

Isoplot (Ludwig, 1999). Additional details can be found in Mueller et al. (2008).

Paleomagnetic Studies

We sampled a total of 68 sites primarily from the massive dolomites and limestone beds at four different localities of the younger Chuluktau and Shabakty Suites of the

Tamdy Series in southern Kazakhstan (Figure 9). The sites were sampled from the

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sedimentary sections in a stratigraphic order ranging in age between Ediacaran to lower

Ordovician, where each site represented a coherent (and sometimes thick) sedimentary package (Table 2a). In addition, the field work carried out in 2005, covered drilling within 15 sites from the limestone units of the Tsagaan-Oloom Formation and 10 sites from the terrigenous sediments of the Bayan-Gol Formation of Baydaric microcontinent in Central Mongolia (Figure 10). The sampling strategy involved collecting samples from the sedimentary sections in a stratigraphic order. The Tsagaan Oloom formation has a stratigraphic age of Vendian to Nemakit - Daldynian (620 – 543 Ma). The other

10 sites from Bayan Gol formation are Nemakit-Daldynian to Atdabanian (543 – 520

Ma) in age (Table 2b). All paleomagnetic samples were either drilled in the field using a water-cooled portable gasoline drill or block sampled. The block samples from the

Tsagaan Oloom sequence were collected from a monoclinal sequence (with variable dips) near Tayshir. Samples were oriented in the field using both sun and magnetic compass and readings were corrected for local magnetic declination. Samples were studied in the paleomagnetic laboratories of the University of Florida in Gainesville and the geological Institute of the Russian Academy of Sciences in Moscow. At University of Florida, cylindrical specimens of standard dimensions 2.2 cm x 2.5 cm (~1000 core samples) were stepwise demagnetized either by thermal or alternating field (AF) methods and measured for the natural remanent magnetization (NRM) on a 2G 77R cryogenic magnetometer. Pilot samples were subjected to detailed stepwise thermal or alternating field (AF) demagnetization and the remaining samples were treated using thermal demagnetization because it resulted in clearer demagnetization trajectories. In

Moscow, cubic specimens of 8 cm3 volume were sawed from hand blocks and were

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demagnetized in 15 - 20° C increments up to 685° C in a homemade oven. Specimens were measured with a JR-4 spinner magnetometer with a noise level of 0.1 mA m-1.

No systematic difference was observed between the samples that were treated in

Gainesville or Moscow, and the data have been pooled.

Representative sample fragments from selected sites were ground into a fine powder and analyzed on a KLY-3S susceptibility bridge with a CS-3 heating unit in order to characterize the temperature-susceptibility behavior of the samples.

Demagnetization results were plotted on orthogonal vector diagrams (Zijderveld, 1967).

The best possible linear segments of the demagnetization trajectories were analyzed via principal component analysis using the IAPD paleomagnetic software (Torsvik et al.,

2000). Components isolated from samples were used to calculate site-means (Fisher,

1953).

Results

Geochronological Results

U-Pb ages from zircon/zircon fragments were determined for the tuff sample KT6-

1B from Talas Karatau microcontinent and the rhyolitic sample GK6 – 3B from the

Greater Karatau microcontinent. The cathodoluminescence (CL) images of the zircons and zircon fragments from these samples show these grains to be prismatic, occasionally acicular with less pronounced magmatic zoning (Figure 11a & c inset).

Twelve laser spots on ten different acicular zircons/zircon fragments from the volcanic tuff sample KT6-1B (Talas Karatau) yielded a well defined 206Pb/238U concordant age of

726.6 ± 4.9 Ma (2σ error, MSWD = 3.0; Figure 11a). The 206Pb/238U concordia age for the volcanic tuff (KT6-1B) is consistent with the weighted mean 206Pb/238U age of 725.7

± 6.3 Ma (2σ error, MSWD = 2.7; Figure 11b) and is interpreted to be the age of

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deposition for the volcanic tuff. Twelve laser spots on 9 different acicular zircons/zircon fragments from the rhyloitic tuff sample GK6-3B (Greater Karatau) yielded age data that are slightly discordant. The data points are aligned along a regression line in the Tera-

Wasserburg diagram with an upper intercept age of 822 ± 65 Ma (2σ error, MSWD =

1.8; Figure 11c). The weighted mean 207Pb/206Pb age for the twelve data points of the sample GK6-3B is 799.7 ± 7.4 Ma (2σ error, MSWD = 1.7; Figure 11d) and is interpreted to be the time of rhyolite eruption. Detailed analytical data are given in

Table-3a and Table 3b respectively.

Paleomagnetic Results

Baydaric block

Representative results of the thermomagnetic analysis (susceptibility vs. temperature) of the carbonate samples of the Tsagaan Oloom Formation are shown in

Figures 12(a-b). The Curie temperature curves are nearly reversible except a few that display an increase in susceptibility during the cooling phase (Figure 12b). This is attributed to mineralogical changes at high temperatures producing pure end-member magnetite (Dunlop and Ozdemir, 1997). The thermal demagnetization of majority of these samples also show unblocking temperatures ranging between 550 - 580º C

(Figure 13a–b) characteristic of magnetite as the principal carrier of the remanent magnetization.

For the terrigenous Bayan Gol formation, the Curie temperature curves show

Curie points at 570-580 °C (as inflections) and again at 680-690°C suggesting that the remanence may be carried by both magnetite and hematite.(Figures 12c-d). This is consistent with the directional changes observed in these samples above 580 C

(Figure 13d). We also note that the heating and cooling curves are nearly reversible in

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the hematite range, but show a large change in susceptibility at temperatures below 600

C indicative of mineralogical changes during heating (Figure 12c, d).

Paleomagnetic results: The low and high unblocking temperature (HTC1) components of the Vendian to Nemakit - Daldynian Tsagaan Oloom and Early

Cambrian Bayan Gol formations of the Baydaric block show similar demagnetization behavior and hence, we discuss the results of their demagnetization experiments together. After removal of the low temperature component during stepwise demagnetization at 250 - 300°C, an intermediate to high unblocking temperature component (HTC1) is resolved from the carbonate-clastic samples of the Tsagaan-

Oloom and Bayan Gol formations at temperatures ranging between 300 and 525 -

640°C (Figure 13a-d). This HTC1 component shows a single polarity, steep upward inclination. The overall mean direction for HTC1 is: Dec = 252; Inc = -68 (α95=8.9º; k=19.5; n = 15; Figure 14a; table 2a) in in-situ coordinates and Dec = 255º, inclination =

-82º (α95=4º; k=91; n = 15; Figure 14a; table 2a) upon correcting for 70% of tilt. When combined with the site means of the B-component of Kravchinsky et al. (2001), the overall mean direction has a Dec = 235.2; Inc = -68.4 (α95=6.9º; k=19.2; n = 24; Figure

15a; table 2a) in in-situ coordinates and Dec = 270.2º, inclination = -81.5º (α95=3.7º; k=64; n = 24; Figure 15a; table 2a) upon correcting for 100% of tilt. The precision parameter (k) for HTC1 increases upon untilting from 19.2 to 65 (maximum k = 94.4 at

70%; Figure 15a). The fold test becomes positive at 50% of unfolding and remains positive at 100% unfolding with inclination values ranging between -75.8° and -81.5°.

The age of the magnetization in a thrust sheet may be constrained by a comparison of the observed magnetization direction with a stable reference apparent

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polar wander path (APWP; Van der Voo, 1993). The Siberian platform has remained a single tectonic unit largely undeformed since the Late Precambrian and has a fairly well constrained paleomagnetic record from the Early Cambrian to the Permo-Triassic boundary (Cocks and Torsvik, 2007). Except the gap of 85 Ma between the Early

Silurian and the Permo-Triassic boundary (at 360 Ma and 275 Ma), the remainder of the

Paleozoic-Early Mesozoic Siberian APWP is better constrained and hence, we use this stable external reference frame of the Early Cambrian - Triassic APWP of the adjacent

Siberian platform (Cocks and Torsvik, 2007) as a comparitive tool to assess the age of magnetization from the Bayan Gol and Tsagaan Oloom Formations of the Baydaric block and the Tamdy Suite of rocks in the Lesser Karatau microcontinent (discussed in the next section).

The syn-folding magnetization inclination of component HTC1 from the Tsagaan

Oloom and Bayan Gol Formations yield a tilt-corrected range of inclination between -

75.8 and -78°. The swathe of possible pole positions generated for the HTC1 component of the Baydaric block corresponding to the mean inclination of -77° lies along a small circle centered on the sampling locality (Figure 16). When compared to the reference Paleozoic APWP for the Siberian craton, the small circle lies very close to the Permian (~248 Ma) and the ill-defined Devono-Carboniferous path (~360-275 Ma) of the Siberian platform (Figure 16). As discussed earlier, the second phase of tectonic deformation in Baydaric microcontinent is related to the Late Devonian – Early

Carboniferous terrane accretion represented by the E-W directed emplacement of the

Chandman arc magmas in the Gobi-Altai Zone (Figure 5; Lehmann et al., 2010). The tectonic emplacement of ophiolites (nappe stacking) in the Trans-Altai Zone over

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Devonian sediments occurred during late Devonian to Early Carboniferous E-W shortening (Lehmann et al., 2010). The same time interval also overlaps with the collision of Altai-Mongolia (AM) and Gorny-Altai (GA) terranes along the Charysh-

Terekta-Ulagan-Sayan suture-shear zone (CTUSs) in SW Mongolia, followed by the emplacement of voluminous granitoids (Glorie et al., 2011 and references therein). This collision led to large-scale Late Devonian – Early Carboniferous strike-slip displacements along the CTUSs followed by the terminal closure of the Paleo Asian

Ocean (PAO) at Late Carboniferous – Permain boundary after the collision of Baltica,

Siberia and Paleo-Kazakhstan (~295 Ma). Although the small circle generated for the

HTC1 component of Tsagaan Oloom and Bayan Gol Formations also lies close to the

Late Permian – Triassic paleopoles from Siberian platform there is no known major deformational event in the Baydaric block at this time. Hence, we suggest that the pre/syn-folding magnetization component (HTC1) was acquired during the Late

Devonian – Early Carboniferous deformational (folding and thrusting) events in the region.

Seven other sites in the Tsagaan Gol river section yielded a high unblocking temperature component (HTC2) with shallow, up/down and northerly to north-easterly paleomagnetic directions (Figure 14b; table 2a). The mean inclination obtained for these seven sites is -18° (α95=11°; k = 31) in geographical coordinates and 7° (α95 =

17°; k = 13.4) after tilt-correction. When combined with the individual site means of the

C-component of Tsagaan-Oloom Formation of Kravchinsky et al. (2001), the mean inclination obtained is -14° (α95=21°; k = 5.5) in in-situ coordinates and 7° (α95=21°; k =

5.8) in stratigraphic coordinates (Figure 14b; table 2a).

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In addition to HTC1, three other sites from the Tommotian red calcareous siltstones (B58, B512 & B513) of the Bayan Gol formation yielded a shallow high temperature component, HTC2. At site B58, thermal demagnetization of the reddish – green calcareous siltstone samples, yielded both high unblocking temperature component HTC1 at 550 and 630°C and component HTC2 at temperatures between

630 – 685°C (Figure 13d). Previous studies conducted on the sandstone and clay rich limestone sequences of Early Cambrian Bayan Gol Formation also yielded a high temperature C – component at temperatures 600-690°C during thermal demagnetization and was interpreted to represent primary magnetization for the Bayan

Gol formation (Kravchinsky et al., 2001). We note that the low inclination HTC2 paleomagnetic results from the red silts in this study are similar to the C – component of the site mean directions obtained by Kravchinsky et al. (2001; Figure 14c; table 2a).

The overall mean paleomagnetic direction after inverting the individual site means of the

C – component of Kravchinsky et al. (2001), is defined by a mean inclination of -3.5° in tilt-corrected coordinates (Figure 14c; table 2a).

The HTC2 direction shows a reasonable consistency in terms of inclination, but has somewhat scattered declinations. The scattering of declinations can be explained in several ways. First, the difference in declination between the HTC2 directions of the

Bayan Gol and Tsagaan Oloom formations could be the result of continental motion and rotation during the time period of deposition. Secondly, the difference could be explained as the result of block rotations within the individual thrust sheets or lastly that the identified components are merely a semi-coherent grouping of spurious magnetizations at high temperature. Distinguishing from among these options is difficult

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given the limited number of sites; however, some information might be obtained by looking at the relationship between strike and declination as well as by applying an inclination only fold test.

Figure 17 (a) shows an inclination - only fold test on the combined HTC2 component and the C – component of Kravchinsky et al. (2001) of the Tsagaan Oloom

(n = 10) and Bayan Gol (n=9) Formations (Watson and Enkins, 1993). The test shows an increase in the precision parameter (k) from 10.74 to a maximum of 15.8 at 50% unfolding and yield a mean inclination of -1.6° ± 6.5. When compared to the Siberian platform, the swathe of HTC2 paleopoles calculated for this mean inclination of -1.6° for the Bayan Gol and Tsagaan Oloom Formations, overlap with the Vendian to early

Cambrian paleopoles of Siberia supporting a near primary magnetization (Figure 19).

To test whether the magnetization was affected by rotation of the individual thrust sheets (see figures 14b and 14c), a strike test (Schwartz and Van der Voo, 1983) was performed using the generalized structural trends. The slope of the linear regression line in the strike test is 0.067 and does not support a pre rotation acquisition of magnetization although the dataset are limited (Figure 17b).

A few sites (e.g. B58) yielded both steeply inclined magnetization component

HTC1 (Devonian – Carboniferous age) with no evidence of block rotations, and the shallow magnetization component HTC2 (630 - 680° C; Figure 13d). Thus, we interpret that the HTC2 component is a near primary magnetization component acquired during

Vendian to early Cambrian followed by an Early Cambrian thrusting event leading to the observed block rotations in the region whereas the HTC1 component is acquired later

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during the Late Devonian – Early Carboniferous deformational (folding and thrusting) events in the region.

Lesser Karatau block

Out of 68 sites sampled in Tamdy suite, only 38 sites yielded consistent paleomagnetic directions. The intensity of the natural remanent magnetization (NRM) of these carbonates ranged between 0.1 mA/m and 1.1 mA/m. The majority of the samples yielded a low temperature component (LTC) aligned along present day geomagnetic field that was removed by 300º C. Apart from this low temperature component, an intermediate to high temperature component (HTC1) was also successfully isolated from the samples at unblocking temperatures ranging between

550º C and 580º C (Figure 19 a-d), consistent with low-Ti magnetite as the main carrier of magnetization. The presence of low Ti-magnetite is supported by Curie temperature analyses where nearly reversible heating and cooling curves show a sharp decrease in magnetic susceptibility at temperature ranges of 585–590º C, characteristic of magnetite

(Figure 12e-f). The overall mean of the HTC1 component yielded an in-situ declination of 209.5º and an inclination of -55.5º (α95=3.8º; k=43; Figure 20b; table 2a). Correction for the bedding tilt yielded a tilt-corrected mean of 214º/-22.3º (α95=9.5º; k=8; Figure

20b; table 2a). A significant decrease in the precision parameter (k) is observed from

47 (in-situ) to 8 after tilt correction, indicating a post-folding magnetization component

(Figure 15b).

The swathe of possible paleolatitudes generated for the HTC1 component of the

Lesser Karatau block corresponding to the mean inclination of -55° lies along a small circle centered on the sampling locality (Figure 21). When compared to the reference

Paleozoic APWP for the Siberian craton, the paleolatitudes for the Lesser Karatau block

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overlaps with the Late Devonian - Early Carboniferous (380 - 360 Ma) and Late Permian

(250 Ma) paleopoles of the Siberian platform. As stated above, the Late Devonian and

Early Carboniferous was the time of key metallogenic and tectonic events in

Kazakhstan and Central Asia (Heinhorst et al., 2000; Pavlova and Borisenko, 2009) that coincide with the generation of new magmatic arcs and oroclinal bending (Levashova et al., 2003; Yakubchak, 2004; Degtyarev, 2011). The Late Permian is known for the extensive Ag-Sb, Sn-Ag and Sn-sulfide mineralization in the Talas – Karatau (Babakhan and Kumyshtag ore clusters) and Lesser Karatau (Kurgan cluster) blocks (Pavlova and

Borisenko, 2009) of central Kyrgyzstan and southern Kazakhstan, respectively. This

Late Permian metallogenic event correlates with the time of alkaline mafic magmatism and with volcano-plutonic emplacement of trachybasalt, trachyte, shoshonite and rhyolites in the area (Yakubchak et al., 2004; Pavlova and Borisenko, 2009; Degtyarev,

2011). We feel this is the most likely time of HTC1 acquisition at Lesser Karatau. At two other sites (e.g. K650 and K651), we resolved a second high-temperature component (HTC2) distinct from the HTC1 overprint with a mean declination=8º and inclination= -35.3º (Figure 20c; table 2b). A paleolatitude of ~20º N or S was calculated corresponding to the inclination = -35°. These low-inclination results, independent of any field tests or robust statistics, are difficult to interpret. We do note that similarly low- inclination results were obtained in the Baydaric block (see below and Kravchinsky et al., 2001) and at Tuva-Mongolia (Kravchinsky et al., 2010) and it is possible that HTC2 represents a near primary remanence placing Lesser Karatau at similar latitudes with

Siberia, the Baydaric microcontinent and Tuva-Mongolia in the Cambrian.

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Discussion

The new geochronological data from the rhyolite and volcanic tuff units of the

Greater Karatau and Talas Karatau microcontinents (this study) are ~coeval with Late

Neoproterozoic magmatic events (~730-850 Ma) reported from other CAOB microcontinents. There is considerable debate regarding the tectonic setting of these felsic and bimodal volcanic events. One group of workers interpret these volcanic units as subduction-related (Mossakovsky et al., 1993, Kheraskova et al., 1995; Levashova et al., 2011a, b; Salnikova et al., 2001) whereas others suggest that all the volcanic complexes are rift related (Ilyin, 1990, Burasnikov and Ruzhentsev, 1993; Khain et al.,

2003). In spite of this debate, these units are often correlated and considered to be part of two larger (and perhaps proximal) domains named ―Kazakhstania‖ and ―Mongolia‖ during this interval (also see Levashova et al., 2011a).

Furthermore, the similarities among the lithologic, stratigraphic, geologic, faunal and tectonic characteristics of the Late Neoproterozoic – Early Paleozoic volcano- sedimentary sequences of the CAOB microcontinents offer target domains to test for further correlation. Models for the evolution of the CAOB can incorporate these data and allow for further comparisons to potential neighboring plates like Siberia, Baltica,

North China, South China, Tarim, India and Australia (Korolev and Maksumova, 1984;

Esakova and Zhegallo, 1996; Khomentovsky and Gibsher, 1996; Li et al., 2002, Khain et al., 2003; Xiao et al., 2004; 2010; Chumakov, 2009a, b; Levashova et al., 2010;

2011a; Shu et al., 2010). Based on the available paleomagnetic and geochronological data for these microcontinents and the adjacent cratonic blocks, we present tentative paleogeographic maps for the two intervals: (1) ~750 Ma (Neoproterozoic), and (2)

~530 Ma (Early Cambrian).

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~750 Ma Reconstruction

There are a number of caveats in positioning various cratonic blocks in any global

Proterozoic paleogeographic reconstruction. Among these, the paucity of the paleomagnetic data, the inherent polarity ambiguity of the paleomagnetic data and the lack of longitudinal control on the positioning of cratonic blocks/continents are of major concern (Meert, 2002, Van der Voo and Meert, 1991). In spite of these difficulties, reconstructions are sometimes made based on the similarities in the geologic, lithostratigraphic, paleontologic and geochronological data from the constituent continental blocks (Rogers and Santosh, 2002, 2009; Zhao et al., 2002, 2004; Li et al.,

2008; Meert et al., 2011b and references therein). In attempting to present a reasonable paleogeography at 750 Ma we have tried to support our paleomagnetic placement of the cratonic blocks with additional data.

The time interval of 750-850 coincides with the hypothesized break-up of the

Neoproterozoic supercontinent Rodinia (1100 Ma – 750 Ma; Meert and Torsvik, 2003; Li et al., 2008). The Neoproterozoic to Paleozoic volcano-sedimentary units of the Greater

Karatau, Talas Karatau and Lesser Karatau domains are suggested to be dismembered segments of the same microcontinent (Seydalin and Sul‘din, V.A., 1977; Windley et al.,

2007; Chumakov, 2009; Levashova et al., 2011a). In this context, paleomagnetic data from the Kurgan Formation (766 ± 7 Ma) of the Lesser Karatau block can be reasonably applied to Greater Karatau and Talas Karatau. Figure 22 shows the paleogeographic reconstruction at ~750 Ma, based on the available paleolatitudes of the Lesser Karatau block of Kazakhstan, Baydaric block of Mongolia and other major Rodinian plates

(Table 4; Levashova et al., 2011a, b and references therein).

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The geologic and geochronological correlations between the Kazakhstan and

Mongolian microcontinents and the Tarim craton are best documented at

Neoproterozoic times (800 – 700 Ma; Chumakov, 2009b; Rojas-Agarmonte et al., 2010;

Levashova et al., 2010, 2011a, b; Meert et al., 2011a; Long et al., 2010). The new U/Pb isotopic ages from the Neoproterozoic felsic and bimodal volcanic units of the Greater

Karatau and Talas Karatau are similar to the 820-750 Ma bimodal and A-type igneous rocks in the Quruqtagh domain of Tarim craton (Zhang et al., 2006; Deng et al., 2008;

Zhu et al., 2009; Shu et al., 2010). Due to the aforementioned issues regarding the tectonic setting of these volcanic rocks, the geochronological correlation between the

Neoproterozoic volcanic sequences of the Karatau and Baydaric blocks to those in the

Tarim and South China cratons (Shu et al., 2011 and references therein) becomes less reliable.

Paleomagnetic data from the Tarim and the CAOB domains can be interpreted to show a close relationship between these blocks during Neoproterozoic times. The palaeolatitudinal position of the Tarim craton at 800 Ma time is constrained to mid latitudes (~43° N/S) based on the Aksu dykes pole (Chen et al., 2004). The paleolatitude of 47° N or S for the Dzabkhan Formation for the Baydaric block and ~34°

N or S for the Kurgan Formation of the Lesser Karatau block also indicate mid-latitudinal positions for those regions (Levashova et al., 2010; 2011a).

Levashova et al. (2011b) hypothesized that the Kazakhstan and Mongol domains had originally belonged either to Tarim, India or South China cratons at 750-800 Ma, based on a comparative analysis of the paleomagnetic, geologic, geochronological and paleontologic data from the Baydaric and Lesser Karatau blocks to other adjacent

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Eurasian blocks (also see Levashova et al., 2010). In particular a strong correlation between the Tarim and Mongol domains was supported by the analysis of the zircon age populations in the Mongol domains (Rojas-Agarmonte et al., 2010). When compared to the potential source regions in Siberia, the Tarim craton, the North China craton and regions within northeastern Gondwana, Rojas-Agarmonte et al. (2010) found the best correlation with age populations from the Tarim craton. Siberia and North

China were ruled out as provenance sources for the zircon populations in the Mongol and Kazakh domains due to the lack of major crust forming events in Siberia and North

China after the Early Riphean. (Figure 22; table 4; Li et al., 2008).

The Tarim block is commonly placed adjacent to the Cathaysia block of South

China based on the correlation of the detrital zircon age population peaks observed for the rift related bimodal volcanism on the South China craton at 830-795 Ma and 780-

745 Ma (Li et al., 2002) and the age distribution from the Tarim block (Shu et al., 2010).

The available paleomagnetic poles at ~750 – 820 Ma place South China at mid to high latitudes position (Li et al., 2004; Evans et al., 2004) further supporting a close affinity to the Tarim craton (Figure 22). Li et al. (2008) argued for a close affinity between

Australia and the Tarim blocks in his Rodinia model based on the coeval ca. 820–800

Ma lamprophyre dykes and kimberlite pipes in the Kimberley craton of Western

Australia (Pidgeon et al., 1989), the 755 Ma Mundine Well dyke swarm in the northwestern Pilbara craton (Wingate and Giddings, 2000), A-type granitic magmatism in the Leeuwin Block of southwestern Australia (Collins, 2003) and the 820-750 Ma bimodal volcanic rocks in the Tarim. The highly reliable Mundine dykes pole (~755 Ma)

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from Australia constrains the position of Australia to latitudes less than ~30 degrees

(Wingate and Giddings, 2000).

~530 Ma Reconstruction:

The break-up of Neoproterozoic supercontinent Rodinia (850-750 Ma) laid the foundation for the formation of the supercontinent Gondwana at ~530 Ma (Meert, 2003;

Meert and Torsvik, 2003). The interval between the disintegration of Rodinia and terminal assembly of Gondwana marks a time of major tectonic reorganization, the kinematics of which remain highly speculative due to the scarcity of quality paleomagnetic data. The Ediacaran to Early Paleozoic is also significant for the final docking of various CAOB microcontinents to form the larger Kazakhstania and Mongol domains.

The majority of the CAOB microcontinents are characterized by the occurrence of thick carbonate and terrigenous sedimentary sequences of late Neoproterozoic to early

Paleozoic age that show considerable stratigraphic, paleontologic and geologic similarities with the coeval sections on the margins of Siberia, Tarim and South China cratons (Korolev and Maksumova, 1984; Esakova and Zhegallo, 1996; Khomentovsky and Gibsher, 1996; Khain et al., 2003; Chumakov, 2009; Levashova et al., 2011a, b,

Meert et al., 2011a). Among these, the important ones are the marker horizons like

Neoproterozoic glacial diamictites (Chumakov, 1978; Korolev and Maksumova, 1984;

Meert et al., 2011a) and the Ediacaran – Early Cambrian phosphorite units (Ilyin, 1990;

Meert and Lieberman, 2008) that lie at several stratigraphic levels of these microcontinental and cratonic blocks. These units have been correlated based on the faunal similarities like small shelly fossils in the Nemakit – Daldynian and Tommotian stages (Esakova and Zhegallo, 1996). Repina (1985) included all these coeval

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sedimentary units in the same trilobite province at the Atdabanian and Botomian stages.

In the following paragraph, we discuss the existing paleomagnetic constraints along with our new paleomagnetic results from the Ediacaran to early Cambrian carbonate-clastic sequences of these units to propose a paleogeographic configuration of the CAOB fragments and other major cratonic blocks.

In addition to the overprint HTC1, a high temperature component of magnetization,

HTC2 is isolated at two other sites on the Lesser Karatau microcontinent with an average inclination of -35.3º corresponding to a paleolatitude of ~ 20º. The paleomagnetic component (HTC2) obtained for the red – calcareous silt units of the

Bayan Gol Formation and seven other limestone units of the Tsagaan Oloom Formation are distinct from HTC1. When compared to the Ediacaran to Early Triassic APWP of the Siberian platform, the paleolatitude for these red silts and limestone beds of the

Bayan Gol and Tsagaan Oloom Formations indicates a near equatorial position further supporting a peri-Siberian position. Our new estimates are consistent with the paleomagnetic results of Kravchinsky et al. (2001) on the Vendian to Early Cambrian sections of the Tuva Mongolia block that also indicate an equatorial to low south paleolatitude position for those units (also see Pecheresky and Didenko, 1995).

Another paleomagnetic study on the Early Riphean to Cambrian rocks from the Bokson

Formation of the Tuva – Mongolia block yielded a palaeolatitudinal position of ~12° N or

S for the Eastern Sayan terrane of Tuva Mongolia block (Kravchinsky et al., 2010).

Based on the overlap of the analogous age poles of the two blocks on Siberian APWP, the Tuva-Mongolia domain was placed adjacent to the Siberian platform (Kravchinsky et al., 2010).

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Figure 23 shows the palaeogeographic map at Tommotian (~530 Ma) time constraining the palaeolatitudinal position of the Baydaric, Lesser Karatau and Tuva-

Mongolia domains along with other major cratonic blocks including Tarim, South China,

Siberia, Baltica, Laurentia and other Gondwanan terranes (table 4). Based on the available paleomagnetic data, the Gondwana supercontinent stretched from South Pole

(South America) to low northern tropical latitudes (Australia). Siberia, Baltica and

Laurentia existed as independent blocks at low-mid paleolatitudes, separated by oceans. Tarim occupied mid-paleolatitudes in proximity to South China and Australia.

Levashova et al. (2010) suggested a hypothetical large shallow tropical sea/basin in the

Late Neoproterozoic – Early Cambrian times with an archipelago of island arcs and microcontinental blocks to the north of Siberia. The new paleomagnetic data from the

Tsagaan-Oloom and Bayan Gol Formations (Baydaric block) and the Tamdy Suite

(Lesser Karatau) suggests that this hypothetical basin was located in warm tropical latitudes in close proximity to Siberia in the south whereas Tarim, South China and

Australia are located further away in the north. The available geological, paleontological and geochemical data also supports the existence of the carbonate-clastic sedimentary sequences overlying the Precambrian crystalline basement (microcontinents) in warm, shallow tropical paleolatitudes. Replina (1985) correlated the Siberian platform with the

CAOB microcontinents, Tarim, South China and Australia at Early Paleozoic times (542

– 520 Ma) based on the occurrence of characteristic Nemakit-Daldynian - Tommotian shelly fossils and benthic trilobite fauna.

Our reconstruction at ~530 Ma supports the models that regard the evolution of

CAOB as an ancient analogue to the modern western Pacific margin, advocating the

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existence of an archipelago of scattered microcontinents, various short-lived island arc systems and oceanic basins within a Paleoasian Ocean at Early Paleozoic time in

Central Asia (Figure 2a,b,c,e; Zonenshain et al., 1990; Mossakovsky et al., 1993;

Didenko et al., 1994; Kheraskova et al, 2003; Khain et al., 2003; Windley et al., 2007;

Kroner et al., 2010; Levashova et al., 2011a). We do note that the palaeolatitudinal disposition of the Baydaric and the Tuva-Mongolia blocks proximal to the Tuva-

Mongolia arc of Şengör and Natal‘in (1996) in our ~530 Ma reconstruction does support the models proposed by Şengör and Natal‘in, 1996; 2004; Yakubchak et al., 2004 and others. But, the paleoposition of the Lesser Karatau domain away from the proposed

Kipchack arc (Şengör and Natal‘in, 1996) makes their model less plausible (also see

Windley et al., 2007).

Summary

New paleomagnetic data and U-Pb isotopic ages are reported for the

Neoproterozoic –Early Paleozoic volcano-sedimentary rocks from the Lesser Karatau,

Greater Karatau (central Kazakhstan) and Talas Karatau (Kyrgyzstan). The LA-ICP-MS analyzed U-Pb concordant zircon age of the rhyolite sample at Talas-Karatau is 726 ±

4.9 Ma. The felsic tuff at Greater Karatau yielded a weighted mean 207Pb/206Pb zircon age of 799.7 ± 6.3 Ma. These new ages are coeval with several other

Neoproterozoic felsic and bimodal units from the CAOB fragments and support arguments for a common origin or association of these blocks during the

Neoproterozoic. The paleomagnetic analysis of the late Neoproterozoic volcanic units from the Dzabkhan Formation (Baydaric) and Kurgan Formation (Lesser Karatau) indicate a close affinity of the Baydaric and Lesser Karatau domains to Tarim, South

China or India at ~750 Ma. However, a drift of these microcontinental blocks towards

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the lower latitudes is evident from the paleomagnetic analysis of the Early Paleozoic sedimentary cover of the Lesser Karatau and Baydaric domains that yield a tropical - low latitudinal position for these blocks (Figure 22). The paleomagnetic data from the

Tommotian to Ordovician carbonate-rocks of the Tamdy Series in the Lesser Karatau yielded a paleolatitude of ~20º that may indicate a peri-Siberian position of this block at this interval. Similarly, the new paleomagnetic results from the Tsagaan Oloom and

Bayan Gol Formations are also consistent with low-latitude results from Tuva-Mongolia and the Siberian platform (Kravchinsky et al. 2001; 2010). We can summarize the following:

(1) During Neoproterozoic time (~ 750 Ma), some of the CAOB microcontinents

(Baydaric and Karatau) belonged to one of the north Rodinian plates like India, Tarim or

South China and was positioned at mid latitudes. (2) During Ediacaran – early

Cambrian time, the Baydaric, Tuva-Mongolian and perhaps Karatau microcontinents were located at tropical paleolatitudes in close proximity to Siberia as indicated by the paleomagnetic analyses discussed in this paper. Based on the above discussion, we suggest that at least some of the CAOB components (Tuva-Mongolia, Baydaric, Talas

Kararatau, Lesser Karatau and Greater Karatau) were situated at low-latitudes during the Late Neoproterozoic – Early Cambrian, as an island archipelago in a peri-Siberian configuration. Our interpretation supports elements of both models for the CAOB evolution. The position of the Tuva-Mongolia and Baydaric blocks is consistent with the

Tuva-Mongol arc in the Şengör and Natal‘in, (1996) model. Although the position of LK is based on very limited data, the low-latitude position is incompatible with an origin as part of the ―Kipchak arc‖.

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Acknowledgements: We thank Mikhail Bazhenov for his assistance with the fieldwork/logistics in Kazakhstan and Kyrgystan in 2006. This work was supported by a

US National Science Foundation grant to J.G.Meert (EAR05-08597) and by a grant from the Russian Foundation of Basic Research (07-05-00021) and Program 8 of the Earth

Science Division, Russian Academy of Sciences to N.M. Levashova.

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Figure 5-1. Location map of the Central Asian Orogenic Belt within Eurasia. (a) Location of the Central Asian orogenic belt (CAOB) within Eurasia. (b) Major microcontinents are labeled as follows: AM, Aktau-Mointy; CM, Central Mongol; CN, Chatkal-Naryn; CT, Central Tian Shan; GK, Greater Karatau; CH, Chu Illi; Ju, Junggar; KO, Kokchetav; LK, Lesser Karatau; NT, North Tian Shan; DA, Derbi-Arzubey; M, Muya; TM, Tuva-Mongol; UL, Ulutau; TB, Tarbagatay BA, Baydaric. U-Pb ages for the volcanic sequences at Greater Karatau*, Lesser Karatau**, Talas***, Aktau-Mointy† and the Baydaric microcontinents are taken from this study, Levashova et al. (2010, 2011a), Sudorgin, 1992; Korolev and Maksumova, 1984; Kiselev and Maksumova, 2001, Kozakov et al., 1993; Sovetov, 2008, Zhao et al., 2006. The green stars show our geochronological sampling sites while insets shows the paleomagnetic locations.

178

Figure 5-2. Sketches showing the contrasting models for the evolution of the Central Asian Orogenic belt. (a) Model by Zonenshain et al. (1990) shows a single Neoproterozoic landmass of Kazakh microcontinents which rifted and remained a loose agglomerate of smaller domains on the equator at ~550 Ma between Siberia and South China; (b) Berzin and Dobretsov, (1994) places Siberia in the southern hemisphere at ~750 Ma, with would be microcontinents of Mongol domain to the north of Siberia. At ~550 Ma these microcontinents became welded to Siberia;

179

Figure 5-2. Sketches showing the contrasting models for the evolution of the Central Asian Orogenic belt (Contd.). (c)Mossakovsky et al. (1993) shows that in the Late Neoproterozoic (1 Ga – 650 Ma), all the CAOB microcontinents were part of the East Gondwana supercontinent, while Siberia straddled the equator on the other side of a 7000 km-wide ocean. At 650-550 Ma, the continent was torn and dissected into a large number of blocks that finally got attached to Siberia by 440 Ma; (d) Sengör and Natal'in, (1996) advocate an extensive ―Kipchak-Tuva-Mongol Arc‖ connecting Siberia and Baltica during the interval from 542 Ma – 250Ma. Successive roll-back and closing of the Khanty–Mansi backarc Ocean by Late Carboniferous due to the differential rotation of Siberia and Baltica and the fragments of the ancient arc gradually coalesced into a continent-sized landmass named ―Kazakhstania‖; (e) model by Kheraskova et al. (2003) supports Mossakovsky et al. (1993) and states that the the rifting of Kazakhstan units from Gondwana took place at ~600 Ma.

180

Figure 5-3. Stratigraphic column for the Lesser Karatau sequence showing the relative locations of the Late Neoproterozoic –Early Paleozoic volcanic and sedimentary units. The δ13C stable isotopic curve is also shown (details in Meert et al., 2011a). The glacial sequence is interpreted as Marinoan in age (~635-650 million years old) and the large negative excursion below the phosphate layer is assigned to the Shuram/Wonoka anomaly. GTS=ICS2004 time scale, R2000=Russian 2000 time scale nomenclature Meert et al., 2011). Thickness ranges are given in meters.

181

Figure 5-4. Simplified geologic map of Karatau ridge showing the main tectonic units of Greater (Bolshoi) and Lesser (Malayi) Karatau blocks (modified after Chakabaev, 1979).

182

Figure 5-5. Simplified tectonostratigraphic map of Mongolia showing the main tectonic units to the north and south of Main Mongolian Lineament (modified from Badarch et al., 2002).

183

Figure 5-6. Stratigraphic column for the Baydaric block, central Mongolia showing the relative locations of the Late Neoproterozoic –Early Paleozoic volcanic and sedimentary units. (a) Detailed stratigraphic column of the Vendian Tsagaan Oloom Formation with three distinct members composed of massive carbonates, phosphorites and characteristic stromatolitic and ichnofossils (modified from MacDonalds, 2009). (b) Detailed stratigraphic column of Lower Cambrian Bayan Gol Formation composed of terrigenous sedimentary units with characteristic ichnofossils (modified from Khomentovsky and Gibsher, 1996). Thickness ranges are given in meters.

184

Figure 5-7. Detailed stratigraphic column of the Riphean to Early Paleozoic sedimentary units of the Greater Karatau block, south Kazakhstan (modified from Chumakov, 2009). The various units include shallow marine carbonates and rift-type volcanic rocks of Riphean age, Vendian shales and Early Palaeozoic deep marine carbonates, cherts and siliciclastic turbidites. The Thickness ranges are given in meters.

185

Figure 5-8. Detailed stratigraphic column of the middle to late Riphean terrigenous and Early Cambrian to late Ordovicain limestone units of the Talas Karatau block in the Kyrgyzstan (modified from Abad et al., 2003). The thickness ranges are given in meters.

186

Figure 5-9. Geologic/Tectonic sketch map of the Lesser Karatau region in southern Kazakhstan (modified from Levashova et al., 2011a), showing our paleomagnetic sampling locations within the Shebakty and Chuluktau suites of Tamdy Series (sites A, B and C in red).

187

Figure 5-10. Geologic/Tectonic sketch map of the Baydaric block, central Mongolia (modified from Levashova et al., 2010), showing our paleomagnetic sampling locations along the Bayan Gol and Tsagaan Gol river sections within the Bayan Gol and Tsagaan Oloom Formations (sites A, B and C in black).

188

Figure 5-11. U/Pb isotopic analysis of the volcanic tuff sample KT6-1B and GK6-3B (a) U/Pb concordia diagram for the twelve laser spots from ten zircons/zircon fragments and tips from the volcanic tuff sample KT6-1B of the Talas Karatau yielding an age of 726.6  4.9 Ma (2) with a MSWD=3.0. (b) The weighted mean 206Pb/238U age of 725.7  6.3 Ma (2; MSWD=2.7) for the Talas Karatau sample KT6-1B. (c) Tera and Wasserburg Concordia diagram for six spots from four zircons/zircon fragments from the rhyolite tuff sample GK6-3B of the Greater Karatau yielding an upper intercept age of 761  17 Ma (2; MSWD=1.2). (d) The weighted mean 207Pb/206Pb age of 788.4  6.8 Ma (2; MSWD=1.4) for sample GK6-3B of Greater Karatau. Inset shows the Cathodoluminescence (CL) images of selected zircons/uranium bearing minerals from the volcanic tuff sample KT6-1B and GK6-3B at Talas Karatau. Scale bars in µm.

189

Figure 5-11. Continued.

190

Figure 5-12. Rock magnetic results for Tsagaan Oloom, Bayan Gol and Tamdy suite of rocks. (a-b) Curie temperature runs for the Vendian age Tsagaan Oloom formation of the Baydaric block, (a) sample T54-17 shows TcH = 565ºC & TcC = 585ºC; (b) sample T514-1 shows TcH = 595ºC & TcC = 600ºC and; (c-d) Curie temperature runs for the representative terrigenous clastic rock samples of the Bayan Gol formation of the Baydaric microcontinent, (c) terrigenous sample B58-5B shows TcH of 630ºC and TcC of 635ºC; (d) red calcareous silt sample B513-2 shows TcH of 665ºC and TcC of 655ºC; (e-f) Curie temperature runs for the representative carbonate rock samples of the Tamdy Suite in the Lesser Karatau microcontinent. (e) Sample K632-19 shows TcH of 555 ºC, and the cooling Curie temperature TcC of 558 ºC; (f) sample K652-5 shows TcH of 575 ºC, and the cooling Curie temperature TcC of 580 ºC.

191

Figure 5-13. Orthogonal vector plots from the carbonate – clastic rocks of the Bayan Gol and Tsagaan Oloom Formations in the Baydaric microcontinent, central Mongolia, showing typical characteristic remanent magnetization directions. (a-b) Thermal demagnetization behavior of the terrigenous rock samples T55 – 2 and T510-35 from the Tsagaan Oloom Formation. (c-d) Thermal demagnetization behavior of the clastic rock specimens B59-32 and B58-9A from the Bayan Gol Formation; Solid squares represent projections on the horizontal plane indicated by ‗H‘; open squares represent projection onto a vertical plane indicated by ‗V‘. Dashed lines denote isolated components labeled as in the text.

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Figure 5-14. Overall site mean directions for the high temperature component (HTC1) and (HTC2) from the Bayan Gol and Tsagaan Oloom formations; (a) in –situ coordinates (left) and after 70 % of tilt-correction (right). (b) Overall site mean directions for the high temperature component (HTC2) from this study and the C-component of Kravchinsky et al., (2001) of the Tsagaan Oloom Formation, in –situ coordinates (left) and after tilt-correction (right). α95 circles of confidence are shown. (c) Overall site mean directions for the high temperature component (HTC2) from the red silts (red beds) and the inverted site means of the individual sites of the C-component of the Kravchinsky et al., (2001) of the Bayan Gol Formation, in –situ coordinates (left) and after tilt-correction (right). α95 circles of confidence are shown. Other notations are same as in figure 10 (see table 2b).

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Figure 5-15. Stereoplots of the HTC1 component for the Tsagaan Oloom and Bayan Gol formations and the Tamdy suite. (a) Stereoplots of the combined HTC1 site mean directions (circles) for the Tsagaan Oloom and Bayan Gol rock formations with overall mean direction shown in yellow shading; in-situ coordinates (left) and tilt-corrected coordinates (right). Incremental fold test showing variations in the precision parameter (k) as a function of percent tilt correction are shown (on the left; based on McElhinny, 1964). The precision parameter shows a significant maximum at 70% unfolding. (b) Stereoplots of the HTC1 site mean directions (circles) for Tamdy suite. Incremental fold test showing variations in the precision parameter (k) as a function of percent tilt correction are shown that indicates a post – folding magnetization (on the left; based on McElhinny, 1964).

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Figure 5-16. Equal area projection of the paleomagnetic poles calculated for average site location of the combined Vendian Tsagaan Oloom and early Cambrian Bayan Gol formations (HTC1). The swathe of possible paleolatitudes generated for the HTC1 component corresponding to the mean inclination of - 77.8° lies along a small circle centered on the sampling locality (blue square). The small circle encloses the Late Permian paleolatitudes for the Siberian platform at their 95% ellipses of confidence (shaded in gray). The ages of the Siberian poles are indicated in Ma.

195

Figure 5-17. Inclination-only fold test and the strike test for the HTC2 component of the Tsagaan Oloom and Bayan Gol formations. (a) Inclination - only fold test plotting kappa versus percent unfolding for the HTC2 component of the nineteen sites of the Tsagaan Oloom (n=10) and Bayan Gol (n=9) Formations. The maximum kappa is observed at 50% unfolding. (b) Strike test for the HTC2 component of the nineteen sites of the Tsagaan Oloom and Bayan Gol Formations plotting Strike (structural trend) versus Declination with a 0.067 slope of the regression line.

196

Figure 5-18. Equal area projection of the paleomagnetic poles calculated for average site location of the combined high temperature component (HTC2) of the Vendian Tsagaan Oloom and early Cambrian Bayan Gol formations. The swathe of possible paleolatitudes generated for the HTC2 component corresponding to the mean inclination of -1.6° lies along a small circle centered on the sampling locality (blue square). The small circle encloses the late Ediacaran – Early Cambrian paleolatitudes for the Siberian platform at their 95% ellipses of confidence (shaded in gray). The ages of the Siberian poles are indicated in Ma.

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Figure 5-19. Orthogonal vector plots from the carbonates and dolostones of the Tamdy Series in the Lesser Karatau block showing typical characteristic remanent magnetization directions. (a) Thermal demagnetization behavior of carbonate sample K640 – 6; (b) Thermal demagnetization behavior of carbonate rock sample K645 – 2; (c) Thermal demagnetization behavior of the limestone core sample K650 – 1; (d) Thermal demagnetization behavior of carbonate sample K650 - 6. Solid squares represent projections on the horizontal plane indicated by ‗H‘; open squares represent projection onto a vertical plane indicated by ‗V‘. Dashed lines denote isolated components labeled as in the text.

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Figure 5-20. Equal area projection diagrams for the Tamdy suite of rocks (Lesser Karatau, Kazakhstan). (a) Site mean directions of the low temperature component (LTC) in –situ coordinates (left) and after tilt-correction (right), (b) Site mean directions of the high temperature component (HTC1), in –situ coordinates (left) and after tilt-correction (right), (c) Site mean directions for the high temperature component (HTC2) from the two sites K650 and K651, in –situ coordinates (left) and after tilt-correction (right). For clarity the confidence circles are not shown. Yellow circle is the overall mean for LTC, HTC1 and HTC2, with the confidence circles. Closed (open) symbols represent downward (upward) inclinations (see Table 2a).

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Figure 5-21. Equal area projection of the paleomagnetic poles calculated for average site location of the Early Cambrian – Ordovician Tamdy suite, Lesser Karatau. The swathe of possible paleolatitudes generated for the HTC1 component corresponding to the mean inclination of -56° lies along a small circle centered on the sampling locality (pink square). The small circle encloses the late Devonian to early Carboniferous and late Permian paleolatitudes for the Siberian platform at their 95% ellipses of confidence (shaded in gray).The Early Cambrian to Triassic paleopoles for the Siberian platform comes from the compilation of Cocks and Torsvik (2007) with their 95% ellipses of confidence. The ages of the Siberian poles are indicated in Ma.

200

Figure 5-22. Paleogeographic reconstruction for the Neoproterozoic (~750 Ma) interval (simplified from Li et al., 2008) showing presumable palaeolatitudinal position of Central Asian Orogenic belt microcontinents based on the two paleomagnetic results from the Kurgan Formation, Lesser Karatau (LK) and the Dzabkhan Formation, Baydaric block (BA). See table 4, for the paleopole details of other continental plates.

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Figure 5-23. Paleogeographic reconstruction for Tommotian time (~530 Ma; simplified from Meert and Lieberman, 2008) showing presumable palaeolatitudinal location of Central Asian Orogenic belt microcontinents with thick carbonate covers and surrounding island arcs. BA = Baydaric block, LK = Lesser Karatau, TM = Tuva Mongolia, Tam = Tarim. See the paleopole information in table 4.

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Table 5-1. Radiometric age data from the Late Neoproterozoic felsic/volcanic units in various CAOB microcontinents S.No. Sampling Area /Microcontinental Formation/ Rock Type/ Method Age Reference Block Series Compositon 1 Ulutau microcontinent Baykonur Liparite porphyry 206Pb/238U 705 ± 10 Ma Korolev and Formation Maksumova, Rhyolite 1984; 206Pb/238U 690 ± 15 Ma Sudorgin, 1992; Intrusive rocks 206Pb/238U 720 ± 20 Ma Kiselev, 2001 2 Central Tian Shan microcontinent Bolshoy-Naryn Rhyolite tuffs 206Pb/238U 830 ± 20 Ma Kiselev and Zone Maksumova, 764 ± 4 Ma 2001; Kroner et al., 2009 3 Chu - Ili microcontinent Kopa Formation Dacite/Rhyolitic 207Pb/206Pb 775 ± 0.8 Ma Kröner et al., rocks 2007 4 Tuva-Mongolia terrane Sarkhoy Series Rhyolite whole rock 718 ± 30 Ma Kuzmichev and (East Sayan) 87Rb/86Sr Baykaite, 1994 5 Baydaric microcontinent Dzabkhan Rhyolite rocks 206Pb/238U 773 ± 3.6 Ma Levashova et al., Formation 803.4 ± 8 Ma 2010 Ash flow tuffs SHRIMP 777 ± 6 Ma Zhao et al., 2006 206Pb/238U 6 Lesser Karatau microcontinent Kurgan Volcanic tuff 206Pb/238U 766 ± 7 Ma Levashova et al., Formation 2011 Volcanic tuff 206Pb/238U 779 ± 17 Ma Sovietov, 2008 7 Greater Karatau microcontinent Kainar Formation Rhyolite rocks 206Pb/238U 814 ± 12 Ma This study 8 Talas Karatau microcontinent Talas zone Volcanic tuff 206Pb/238U 771 ± 17 Ma This study

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Table 5-2a. Kazakhstan paleomagnetic results (High temperature components HTC1 and HTC2) HTC1 In-situ Tilt-corrected Site/Study S/D N Dec Inc Dec Inc Κ α95 Stratigraphic Age Site K62 336/78 6 208 -59 227 013 83 7 Late Riphean Site K63 334/80 14 216 -46 222 028. 42 8 Late Riphean 0.0 Site K67 295/22 9 194 -64 196 -42.0 436 3 Tommotian Site K68 323/17 13 209 -34 212 -18.0 22 9 Tommotian Site K69 322/13 4 208 -51 214 -32.0 17 22 Tommotian Site K611 290/14 7 228 -29 226 -15.0 34 11 Tommotian/Atdabanian Site K612 295/20 3 206 -54 206 -24.0 157 10 Tommotian/Atdabanian Site K613 305/17 8 188 -69 200 -53.0 37 9 Tommotian Site K617 280/25 2 203 -57 198 -32.0 25 25 Nemakit-Daldynian Site K619 105/20 9 208 -58 230 -76.0 168 5 Tommotian/Atdabanian Site K620 063/15 23 208 -62 229 -67.0 133 7 Tommotian/Atdabanian Site K621 051/12 9 200 -68 228 -69.0 293 3 Tommotian/Atdabanian Site K623 040/10 8 200 -59 217 -61.0 78 6 Tommotian/Atdabanian Site K624 062/22 6 215 -59 256 -62.0 71 8 Toyonian/Botomian Site K627 013/18 6 227 -46 223 -38.0 59 12 Tommotian/Atdabanian Site K628 345/22 3 203 -63 224 -46.0 106 12 Tommotian/Atdabanian Site K629 018/29 4 216 -58 246 -41.0 19 22 Tommotian/Atdabanian Site K634 286/42 9 192 -42 196 -0.7 79 6 Late Cambrian Site K636 286/42 7 200 -57 198 -15.0 14 16 Middle Cambrian Site K637 286/42 9 207 -49 203 -7.7 192 4 Late Cambrian Site K638 290/41 5 190 -58 194 -17.0 43 10 Late Cambrian Site K639 286/42 8 210 -51 205 -9.8 39 10 Toyonian/Botomian Site K640 295/54 8 213 -34 212 20.0 50 9 Toyonian/Botomian Site K641 299/45 8 209 -61 204 -20.0 330 4 Toyonian/Botomian Site K642 299/45 9 209 -46 209 -1.2 167 5 Toyonian/Botomian Site K643 300/45 8 208 -58 212 -11.0 150 5 Late Cambrian/Early Ordovician Site K645 302/52 8 210 -49 210 -0.7 54 8 Arenig -Tremadoc Site K646 302/52 13 211 -62 211 -10.0 21 12 Arenig -Tremadoc Site K647 300/45 3 203 -55 206 -10.0 224 8 Arenig -Tremadoc Site K648 292/40 2 236 -45 225 -9.0 159 20 Arenig -Tremadoc Site K649 320/75 9 203 -63 219 12.0 16 13 Arenig -Tremadoc Site K656 322/74 5 213 -74 226 -16.0 59 10 Llandeilo Site K660 320/78 5 251 -72 237 5.0 35 13 Llandeilo Overall - 200 209 -56 - - 43 4 Mean Tilt- - 200 - - 214 -22.0 8 10 corrected Mean

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Table 5-2a. Continued HTC2 In-situ Tilt-corrected Site/Study S/D N Dec Inc Dec Inc Κ α95 Stratigraphic Age Site K650 324/76 12 15 26 11 -35 177 3.0 Arenig -Tremadoc Site K651 324/76 8 13 22 5 -36 111 5.8 Arenig - Tremadoc Overall Mean 324/76 20 14 24 - - 772 9.0 Tilt-corrected 324/76 20 - - 8 -35 511 11.0 Mean N=number of samples used; S/D = Strike/Dip; k=kappa precision parameter; a95= cone of 95% confidence about the mean direction. HTC1, High temperature component 1; HTC2, High temperature component 2.

205

Table 5-2b. Baydaric paleomagnetic results (High temp. components, HTC1 & HTC2) Tsagaan In-situ Tilt-corrected Oloom (HTC1) Site/Study S/D N Dec Inc Dec Inc Κ α95 Stratigraphic Age Site T52 102/12 19 219 -76 274 -83 220 2.3 Varangerian Site T53 277/43 35 346 -63 215 -70 80 3.0 Ediacaran Site T54 132/15 20 220 -65 218 -80 887 1.1 Ediacaran Site T55 126/23 16 230 -65 289 -84 303 2.1 Ediacaran Site T56 110/28 38 216 -66 329 -82 110 2.2 Nemakit-Daldynian Site T58 135/20 6 223 -61 220 -81 138 5.7 Nemakit-Daldynian Site T59 135/20 6 232 -69 298 -86 253 4.8 Nemakit Daldynian Site N4000 223/31 30 304 -50 279 -79 20 6.0 Ediacaran (Tayshir monocline) Overall Mean - 170 248 -71 - - 16 14.0 Tilt-corrected - 170 - - 253 -83 117 5.1 Mean

Bayan Gol (HTC1) B54 225/40 16 294 -52 209 -76 78 4.3 Tommotian B55 210/38 69 294 -53 201 -86 105 1.7 Tommotian B56 160/14 9 227 -64 204 -76 207 3.6 Tommotian B57 157/14 18 235 -61 226 -75 226 2.3 Tommotian B58 100/30 9 234 -61 302 -69 62 6.2 Tommotian B59 100/30 40 252 -69 277 -82 23 5.8 Tommotian B510 132/30 16 228 -62 235 -77 453 1.7 Nemakit-Daldynian Overall Mean - 177 255 -64 - - 25 12.0 Tilt-corrected - 177 - - 264 -81 30 11.0 Mean

Overall mean - 347 252 -68 - - 19 8.9 Bayan Gol + Tsagaan Oloom (HTC1) Tilt-corrected - 347 - - 258 -82 53 5.3 Mean Kravchinsky 9 209 -66 284 -80 117 4.8 Mean

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Table 5-2b. Continued Tsagaan In-situ Tilt-corrected Oloom (HTC2) Site/Study S/D N Dec Inc Dec Inc Κ α95 Stratigraphic Age Site T51 116/66 12 345 -25.0 345.0 27.0 50.0 6 Varangerian Site T510 172/53 17 349 -6.0 354.0 -4.6 31.0 7 Ediacaran Site T511 145/50 4 7 -16.0 14.0 24.0 11.0 20 Ediacaran Site T512 168/44 13 358 -23.0 10.0 -10.0 16.0 9 Ediacaran Site T513 140/50 5 15 -13.0 10.0 27.0 19.0 11 Ediacaran Site T514 150/50 5 343 -15.0 350.0 0.0 29.0 14 Ediacaran Site T515 165/40 9 13 -25.0 17.0 -15.0 18.0 12 Ediacaran Overall Mean 53 359 -18.0 - - 31.0 11 Tilt-corrected 53 - - 2.8 7.0 13.4 17 Mean Kravchinsky 4 59.5 -1.1 58.8 5.3 19.2 21 Mean Red silts (HTC2) B58 100/30 6 192 -1.5 192.0 -28.0 21.0 14 Tommotian B512 166/43 6 297 13.0 299.0 -20.0 22.0 15 Tommotian B513 180/54 3 304 32.0 300.0 -15.0 68.0 15 Tommotian Kravchinsky et al. (2001) C- comp. sites 6-1 241/14 9 115 10.7 116.1 1.0 3.7 29 6-2 102/30 8 114 -24.0 99.4 -27.0 11.5 19 6-3 90/24 8 273 -7.8 276.0 -6.3 5.3 27 6-4 1/25 6 128 9.5 128.0 -10.3 6.0 30 6-5 77/28 12 128 5.1 126.0 -16.3 6.3 19 6-6 117/35 8 132 21.0 142.0 9.3 8.2 23 Overall Mean - 66 291 1.6 - - 4.5 27 Tilt-corrected - 66 - - 292.0 -3.5 5.0 26 Mean N=number of samples used; S/D=strike/Dip; k=kappa precision parameter; a95= cone of 95% confidence about the mean direction. HTC1, High temperature component 1; HTC2, High temperature component 2.

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Table 5-3a. Geochronologic results from the Talas – Karatau 207Pb/ 206Pb/ *207Pb/ 206Pb/ 207Pb/ 207Pb/ % Grain ±2σ ±2σ ±2σ 2σ 2σ 2σ 206Pb 238U 235U 238U(Ma) 235U(Ma) 206Pb(Ma) Disc RHO KT6_1B_8 0.0630645 0.0012 0.11619 0.0020 1.0103 0.03 709 12 709 13 710 42 0 0.66 KT6_1B_11 0.0627034 0.0016 0.11764 0.0024 1.0170 0.03 717 14 712 17 698 56 -3 0.62 KT6_1B_16 0.0647159 0.0009 0.12007 0.0022 1.0714 0.03 731 13 739 13 765 31 4 0.79 KT6_1B_17 0.0644539 0.0017 0.12041 0.0021 1.0701 0.03 733 13 739 16 757 54 3 0.58 KT6_1B_24 0.0631658 0.0017 0.11917 0.0020 1.0379 0.03 726 11 723 16 714 57 -2 0.51 KT6_1B_30 0.0643488 0.0010 0.11865 0.0023 1.0527 0.03 723 14 730 13 753 34 4 0.77 KT6_1B_31 0.0633916 0.0010 0.12035 0.0020 1.0519 0.02 733 11 730 12 721 33 -2 0.72 KT6_1B_32 0.0629193 0.0013 0.11811 0.0014 1.0247 0.02 720 9 716 12 706 44 -2 0.52 KT6_1B_37 0.0638259 0.0008 0.11984 0.0025 1.0547 0.03 730 14 731 12 736 27 1 0.85 KT6_1B_42 0.0642634 0.0009 0.11785 0.0024 1.0443 0.03 718 14 726 13 750 29 4 0.80 KT6_1B_43 0.0639639 0.0015 0.11939 0.0027 1.0530 0.03 727 15 730 17 740 48 2 0.70 KT6_1B_44 0.0640428 0.0018 0.12261 0.0020 1.0827 0.03 746 11 745 17 743 59 0 0.51

Table 5-3b. Geochronologic results from the Greater – Karatau 207Pb/ 206Pb/ *207Pb/ 206Pb/ 207Pb/ 207Pb/ % Grain 206 2σ 238 2σ 235 2σ 238 2σ 235 2σ 206 2σ Pb U U U(Ma) U(Ma) Pb(Ma) Disc RHO GK6_3B_5 0.0653976 0.0003 0.11837 0.0023 1.0673 0.02 721 10 737 8 787 11 8 0.94 GK6_3B_6 0.0656101 0.0003 0.11904 0.0023 1.0768 0.01 725 9 742 8 794 11 9 0.93 GK6_3B_7 0.0656078 0.0004 0.11921 0.0023 1.0784 0.02 726 10 743 8 794 13 9 0.92 GK6_3B_11 0.0651884 0.0004 0.11746 0.0023 1.0557 0.02 716 10 732 8 780 13 8 0.92 GK6_3B_21 0.0657516 0.0005 0.11590 0.0025 1.0507 0.02 711 14 730 12 792 20 10 0.91 GK6_3B_28 0.0650683 0.0004 0.12047 0.0026 1.0808 0.02 733 10 744 8 777 12 6 0.93 GK6_3B_37 0.0656451 0.0003 0.11679 0.0025 1.0571 0.02 712 10 732 8 795 11 10 0.95

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Table 5-4. Paleomagnetic poles for the major cratonic blocks (Neoproterozoic – Cambrian) S. Cratonic domain Age range Paleopole Paleo- Reference No. (unit) (Ma) Latitude (degree)

Nº Eº A95 1. Kazakhstan Kurgan Fm., Lesser 766 ± 7 - - - 34.2 ± 5.3° Levashova et al., 2011 Karatau N/S Tamdy suite (HTC2), 490 - 630 - - - ~20° N/S This study Lesser Karatau

2. Mongolia Dzabkhan Fm., ~770-800 - - - 47° ± 14º, Levashova et al., 2010 Baydaric block N/S Tsagaan Oloom Fm. 544-530 - - - 3.5° N/S This study (HTC2) Bayan Gol Fm. 530-520 - - - 1.8° N/S This study; Kravchinsky HTC2 + C et al., 2001 component Bokson Fm., Tuva 518 - 630 -39 98 12.7 1° ± 8° N/S Kravchinsky et al., 2010 Mongolia 3. Siberia Karagas Group 850–740 12 97 10 22-8N Metelkin et al., 2005

Nersinsky Complex ~740 -37 122 11 2N-19S Metelkin et al., 2005 Mean Ediacaran ~560 -35 77 6 3-16S Shatsillo et al., 2006 Pole Redkolesny Fm. ~550 -61 68 5 30-42S Shatsillo et al., 2006 Mean Nemakit- ~540 -60 115 7 25-41S Shatsillo et al., 2006 Daldynian pole Mean Early ~525 -48 151 8 19-36S Shatsillo et al., 2006 Cambrian pole 5. Tarim Aksu dykes 807 ± 12 19 128 6 ~43N Chen et al., 2004 Baiyixi Fm. ~ 740 17 194 4 6S Huang et al., 2005 Sugetbrak Fm. ~551 19 150 8 Zhang et al., 2006 6. South China Xiaofeng Dikes 807 ± 10 14 91 11 ~69N Li et al., 2004 Liantuo Fm. 748 ± 12 4 161 13 ~37N Evans et al., 2004 Nantuo Fm. ~740 0 151 5 ~43N Rui and Piper, 1997 Meishucun Fm. ~ 525 9 31 10 ~14N Lin et al., 1985 Tianheban Fm. ~511 -7 10 23 ~12S Lin et al., 1985

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Table 5-4. Continued S. Cratonic domain Age Paleopole Paleo- Reference No. (unit) range Latitude (Ma) (degree)

Nº Eº A95 7. Australia Hussar Fm. 800-760 62 86 10 3N-19S Pisarevsky et al., 2007 Walsh Tillite 770-750 -21 102 45N–19N Mundine Dikes 755 ± 3 45 135 4 31-8N Wingate and Giddings, 2000 Elatina Formation 600–620 39 186 9 26-1N Sohl et al., 1999 Brachina Formation ~580 33 148 16 44-20N McWilliams and McElhinny, 1980 Lower Arumbera ~550 46 157 4 30-6N Kirschvink, 1978 Todd River ~530 43 160 7 32-9N Kirschvink, 1978

8. Baltica Hunnedalen Dikes ~848 -41 222 10 59-87S Walderhaug et al., 1999 Egersund Dikes ~608 -31 224 15 50-81S Walderhaug et al., 2007 Tornetrask Formation ~535 –56 296 12 24-52S Torsvik and Rehnstrom, 2001 8. Laurentia Wyoming Dikes 782 ± 8 13 131 4 23N-36S Harlan et al., 1997 Tsezotene Sills and 779 ± 2 2 138 5 16N-34S Park et al., 1989 Dikes Kwagunt Formation 742 ± 6 18 166 7 41N-6S Weil et al., 2004 Franklin Dikes 723 ± 4 5 163 5 27N-19S Heaman et al., 1992; Park, 1994 Long Range Dikes 620–610 19 355 18 34N-6S Murthy et al., 1992; Kamo and Gower, 1994 Callander Complex 575 ± 5 -46 121 6 34-81S Symons and Chiasson, 1991 Catochin A Basalt 564 ± 9 -42 117 9 32-77S Meert et al., 1994 Sept_Iles B Complex 564 ± 4 -44 135 5 27-74S Tanszyk et al., 1987 9. Congo - SF Mbozi complex 755± 25 46 325 9 Meert et al., 1995

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CHAPTER 6 CONCLUSION

The work included in this dissertation encompasses the results of the extensive paleomagnetic and geochronologic investigation carried out on the Precambrian mafic dykes of the Indian sub-continent and the Late Neoproterozoic – Early Paleozoic volcano-sedimentary sequences of the Baydaric and Lesser Karatau microcontinents of the Central Asian Orogenic Belt (CAOB). The first task of this work was to generate new well dated paleomagnetic poles for India during Precambrian in order to constrain its position in the proposed supercontinental assemblies of Mesoproterozoic Columbia and Neoproterozoic Rodinia.

We were able to obtain six new well-dated poles for India: (1) the older NW-SE trending dykes of the Bundelkhand craton in Central India yielded an overall paleomagnetic pole of 58.5ºN and 312.5ºE (dp/dm=6.6º/7.9º); (2) A combined result for the poorly dated Paleoproterozoic Gwalior traps yielded a paleomagnetic pole for India at 15.4°N, 173.2°E (dp = 5.6°, dm= 11.2°) using the reported age of ~1.8 Ga (Chapter 3;

Crawford and Compston, 1969; Ramakrishnan and Vaidyanadhan, 2008); (3) the paleomagnetic pole from the 1192 Ma alkaline suite of the Harohalli dykes was updated and recalculated at 24.9° S, 258° E (A95=15°); (4) a 1113 Ma VGP of 38.7ºS and

49.5ºE (dp/dm = 9.5º/16.3º) from the NE-SW trending Mahoba dykes of the

Bundelkhand craton was obtained; (5) the 1025.6±3.8 Ma tholeiitic and alkaline

Anantapur dykes yielded a VGP of 10°N and 211°E (a95=10°); and (6) an average paleomagnetic pole of 76.5◦S, 68.8°E (a95=15°) obtained from the overprint of the

Harohalli dykes (Pradhan et al., 2008; Halls et al., 2007). The new paleomagnetic data

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generated in this work have significant implications in understanding the Precambrian paleogeography of Indian sub-continent.

The paleomagnetic results from the ~2000 Ma NW-SE trending dyke suite of the

Bundelkhand craton (Chapter 4) along with other published data from the 1800-2000

Ma interval supports a close proximity between the major elements of Peninsular India

(Aravalli-Bundelkhand and Dharwar-Bastar and Singhbhum cratons). If correct, it would indicate that the younger tectono-magmatic events recorded in the CITZ represent crustal-scale reactivation along an existing zone of weakness. Our ~2000 Ma pole from the NW-SE trending Bundelkhand dykes is the first robust paleomagnetic pole reported for India at this time period. In a ~2000 Ma paleogeographic reconstruction, the

Bundelkhand dykes pole indicates a low latitudinal position for India with a mid latitudinal disposition of Laurentia and the Kaapvaal block in a vastly different configuration than that suggested by Zhao et al. (2004) based on the geometric relationships of the radiating dyke swarms. The global paleogeography at ~2000 Ma favors scattered independent continental blocks as suggested by Pesonen et al. (2003; also see Salminen et al., 2009) and does not support the Columbia model of Zhao et al.

(2004).

Our reconstruction at ~1800 Ma utilizing the new Gwalior paleomagnetic pole places Indian subcontinent at equatorial position with a paleolatitude of 2.2 ± 5.5° N. A near equatorial position for India is also indicated by the VGPs from the older (~1900

Ma) mafic dykes of the southern Bastar craton and nearly Cuddapah basin from the adjacent Dharwar craton, India (French et al., 2008; Clark, 1983, Meert et al., 2011), although the declinations differ by ~60°.

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The new U-Pb ages of 1192 Ma for the Harohalli paleomagnetic pole of the

Dharwar craton in southern India update and revise the age of the pole by almost 350

Ma (see discussion in Chapter 2). The Harohalli paleopole is considered as a key pole for India during Early Neoproterozoic (~823 Ma) and has been extensively used in various supercontinental reconstructions for India. If our new age of 1192 Ma for the

Harohalli paleomagnetic pole is indeed accurate, then it makes several extreme geodynamic models (inertial interchange and true polar wander events) proposed based in part on presumed ~823 Ma age for the Harohalli alkaline dykes (Li et al., 2004;

Maloof et al., 2006) less well-constrained. An interesting implication of our new age is that it would place India adjacent to both Australia and Laurentia in a ca. 1.2 Ga reconstruction (Figure 2.6; Pesonen et al., 2003).

The controversy regarding the age of the Upper Vindhyan units is discussed in

Chapter 3 and 4 of this dissertation. The U-Pb age of 1113 Ma in combination with a virtual geomagnetic pole (37.8ºS, 49.5ºE; dp/dm=10.8/18.3) for the NE-SW trending

Mahoba suite of dykes correlates well with the VGP generated for the Majhgawan kimberlite and the paleomagnetic poles from the Bhander-Rewa Groups of the Upper

Vindhyan sequence (Chapter 4; Gregory et al., 2006; Malone et al., 2008). Our interpretation of the Mahoba paleomagnetic and geochronologic results lend further support to the argument that the closure age for the Upper Vindhyan rocks is older than

1000 Ma. With reference to Rodinia, one of the interesting implications of our results from ~1113 Ma Mahoba dyke pole in conjunction with the coeval Majhgawan kimberlite and Bhander- Rewa poles is that it infers a low latitudinal position for India with

Australia-Mawson blocks located at mid latitudes. The proposed geological coherence

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of the Mawson block with Australia in post-1200 Ma configurations (Cawood and

Korsch, 2008; Payne et al., 2009; Wang et al., 2008; Kelly et al., 2002; Torsvik et al.,

2001) and the disposition of India at low latitudes might indicate a loose East Gondwana fit, with India positioned further south of East Antarctica.

The new paleomagnetic and geochronologic results reported from the

Bundelkhand dykes, Gwalior traps, the Harohalli dykes, Anantapur dykes and the

Harsani granodiorite, coupled with the previous studies done by our group emphasize that revisions in the Precambrian paleogeographic history of the Indian subcontinent are still in the early stages and additional future paleomagnetic studies are crucial to better understand the tectonic and paleogeographic evolution of Precambrian India.

The second objective of this work was to provide geochronologic and paleomagnetic support to the various competing models of the CAOB evolution. We believe that geochronological correlation supported by paleomagnetism; stratigraphy and faunal record of various coeval microcontinental blocks of CAOB could play a key role in understanding the origin and evolutionary history of the CAOB during Late

Neoproterozoic-Early Paleozoic times. The LA-ICP-MS analyzed U/Pb concordant zircon age of the rhyolite sample at Talas-Karatau (Kyrgyzstan) is 726.6 ± 4.9 Ma and the felsic tuff at Greater Karatau (Kazakhstan) yielded a weighted mean 207Pb/206Pb zircon age of 788.4 ± 6.8 Ma. These ages from the Talas-Karatau and Greater Karatau microcontinents are consistent with the ages obtained from the upper (~831Ma) and lower (~766 Ma) tuffaceous units in Kurgan formation of the Lesser Karatau (Levashova et al., 2011a; Sovetov et al., 2008). These Neoproterozoic units of Lesser Karatau have already been correlated to the ~750-800 Ma old volcanic rocks in Dzabkhan (Baydaric)

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microcontinent in Mongolia by Levashova et al. (2011a). Furthermore, the paleomagnetic analysis of the late Neoproterozoic volcanic units from the Dzabkhan

Formation (Baydaric) and Kurgan Formation (Lesser Karatau) indicate a close affinity of the Baydaric and Lesser Karatau domains to either Tarim, South China or India at ~750

Ma (also see Levashova et al., 2011a). A drift of these microcontinental blocks towards lower latitudes is evident from the paleomagnetic analysis of the Early Paleozoic sedimentary cover of the Lesser Karatau and Baydaric domains that yield a tropical - low latitudinal position for these blocks. We suggest that at least some of the CAOB components (Tuva-Mongolia, Baydaric, Talas Karatau, Lesser Karatau and Greater

Karatau) were situated at low-latitudes during the late Neoproterozoic – Early Cambrian, as an island archipelago in a peri-Siberian configuration. At the same time, we do recognize that our new paleomagnetic data for the Baydaric and Lesser Karatau blocks during Vendian-Ordovician are limited, and future paleomagnetic studies are crucial for the coeval sedimentary sequences from other CAOB microcontinents, to better understand the tectonic evolution of the Central Asian Orogenic Belt at this interval.

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BIOGRAPHICAL SKETCH

Vimal Roy Pradhan was born in Jaipur, Rajasthan, India on August 31st, 1979 to

Shanti Sharma and Harish Chandra Pradhan. He completed his high school education from the Sainik School Chittorgarh and D.A.V Centenary Public School, Jaipur, India.

He attended the prestigious Maharaja College, Jaipur, India for his undergraduate majoring in the core subjects geology, chemistry and zoology. It was in his undergraduate studies, he got inclined towards the fascinating science of the earth and decided to pursue his MS degree in geology. He completed his MS thesis in sedimentology and paleontology in the year 2004 under the supervision of Dr.

D.K.Pandey. The title of his thesis is ―Microfacies and Depositional Environment of the

Gaj and Dwarka formations (Miocene) exposed near Bhatia, district Jamnagar,

Saurashtra, India‖. He was awarded the best student award for his master‘s thesis by

University of Rajasthan, India. After his MS, he worked as a Junior Research Fellow in

2004-05 on a Project titled ―Evolution of Precambrian mantle in sub-continental NW

India, clues from Geochemistry of mafic dikes‖ awarded to Dr. M.K.Pandit. In December

2005, he went on a month long geological field trip to various parts of the Central and

South India with Dr. Joseph.G. Meert, drilling core samples from the Precambrian mafic dykes and sedimentary Formations for carrying out the paleomagnetic and geochronologic analysis. The short educational lectures and discussions on the dinner table every evening after rigorous field work focussing the basics of Paleomagnetism and paleogeographic reconstructions with Dr. Meert inclined Vimal towards the elusive world of Paleomagnetism and he decided to pursue his Ph.D. program under Dr.

Meert‘s supervision. He then moved to Gainesville, Florida in the spring of year 2006 to begin his Ph.D. program in the Geology Department, University of Florida. In addition to

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the chapters documented in this dissertation, Vimal have contributed to, and is a co- author on, the following publications that are related to his dissertation work. (1) Meert,

J.G., Pandit, M.K., Pradhan, V.R., and Kamenov, G., 2011. Paleomagnetism of the

Paleoproterozoic mafic dykes from the Bastar and Dharwar cratons, India: Preliminary findings, Gondwana Research, 20, 335-343; (2) Meert, J.G., Pandit, M.K., Pradhan,

V.R., Banks, J., Sirianni, R., Stroud, M., Newstead, B., and Gifford, J., 2010.

Precambrian Crustal evolution of Peninsular India: A 3.0 billion year odyssey, Journal of

Asian Earth Sciences, 39, 483-515; (3) Malone, S.J., Meert, J.G., Banerjee, D.M.,

Pandit, M.K., Tamrat, E., Kamenov, G., Pradhan, V.R. and Sohl, L., 2008.

Paleomagnetism and detrital zircon geochronology of the Upper Vindhyan sequence,

Son Valley and Rajasthan, India: A 1000 Ma closure age for the Purana basins,

Precambrian Research, 164, 137-159; (4) Pandey, D.K., Kondo, Y., Jain, R.L., Bahadur,

T. and Pradhan, V.R., 2008. Microfacies and Depositional Environment of the Gaj

Formation (Miocene) exposed near Bhatia, District Jamnagar, Saurashtra, Journal of the Paleontological Society of India, 53, 131-145; (5) Gregory, L.C., Meert, J.G.,

Tamrat, E., Malone, S., Pandit, M.K. and Pradhan, V.R., 2006. A paleomagnetic and geochronologic study of the Majhgawan kimberlite, India: Implications for the age of the

Upper Vindhyan Supergroup, Precambrian Research, 149, 65-75.

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