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Florida State University Libraries Electronic Theses, Treatises and Dissertations The Graduate School 2006 Geochemical and Geochronological Investigations in the Southern Appalachians, Southern Rocky Mountains and Deccan Traps. Reshmi Das Follow this and additional works at the FSU Digital Library. For more information, please contact [email protected] THE FLORIDA STATE UNIVERSITY COLLEGE OF ARTS AND SCIENCES GEOCHEMICAL AND GEOCHRONOLOGICAL INVESTIGATIONS IN THE SOUTHERN APPALACHIANS, SOUTHERN ROCKY MOUNTAINS AND DECCAN TRAPS. By RESHMI DAS A Dissertation submitted to the Department of Geological Sciences in partial fulfillment of the requirements for the degree of Doctor of Philosophy Degree Awarded: Fall Semester, 2006 The members of the Committee approve the Dissertation of Reshmi Das defended on 1st November, 2006. ____________________________________ A. Leroy Odom Professor Directing Dissertation ____________________________________ Jeffrey Chanton Outside Committee Member ____________________________________ Stephen A. Kish Committee Member ____________________________________ Vincent J. M. Salters Committee Member ____________________________________ James F. Tull Committee Member Approved: _____________________________________________ Professor A. Leroy Odom, Chair, Geological Sciences The Office of Graduate Studies has verified and approved the above named committee members. ii To my first geology teacher, Dr. Mohan Chand Baral, who made me fall in love with Geology. iii ACKNOWLEDGEMENTS Many people have contributed to the making of this dissertation. First and foremost I would like to acknowledge Prof. Leroy Odom, my advisor who not only supported me academically and financially for five years but also lent personal support whenever required. This dissertation would have been impossible without him. Prof. James Tull supervised my project on the southern Appalachian and words are inadequate to express his contribution in development of this dissertation. Prof. Stephen Kish always had an answer for my most difficult question and Prof. Vincent Salters helped me with the modeling part of the chemical data. Prof. Munir Humayun, Prof. Tapas Bhattacharyya and Dr. Michael Bizimis, though not a part of my committee, supervised portions of my research. I am indebted to them for their time, technical support and patience. The field work required for this dissertation was partly funded by EDMAP component of the National Geologic Mapping Act 2002. The Isotope Geochemistry Division at the National High Magnetic Field Laboratory has been a wonderful place to work, and I would like to thank Ted Zateslo and Afi SachiKocher for their technical support. I would also like to thank all of the geochemistry graduate students for the long chats, the little favors, and camaraderie along the way. My parents made the greatest sacrifice of their life by letting their only child go ahead with her life and providing the best educational opportunity. Last but not the least is my best friend and husband Subhajit who inspired me to continue my career in academia. His extremely active academic and intellectual presence kept me going and will continue to do so. iv TABLE OF CONTENTS List of Tables ................................................................................................ List of Figures ................................................................................................ Abstract ...................................................................................................... vii viii x 1. Geochemical and Geochronological Constraints on the Origin and Evolution of the Eastern Blue Ridge, Southern Appalachians. .......................................... 1 1.1 Introduction........................................................................................... 1 1.2 Regional Overview ............................................................................... 3 1.3 Study Area and General Description of Lithotectonic Units................ 4 1.4 Purpose of the Study ............................................................................. 7 1.5 Analytical Techniques .......................................................................... 8 1.6 Results ................................................................................................ 10 1.7 Nd Model Age and Detrital Zircon Ages of the Metasediments of the Eastern Blue Ridge .......................................................................................... 14 1.8 Mulberry Rock Gneiss .......................................................................... 15 1.9 Discussion ............................................................................................. 16 2. Kilometers Scale Strontium Isotopic Homogenization During Metamorphism: A Case Study in the Tres Piedras Granite, New Mexico.................................... 62 2.1 Introduction........................................................................................... 2.2 General Geology ................................................................................... 2.3 Tres Piedras Granite.............................................................................. 2.4 Previous Work ...................................................................................... 2.5 Results ................................................................................................ 2.6 Discussion ............................................................................................. 2.7 Conclusion ............................................................................................ 62 63 64 65 69 70 74 3. Trace Element and Lead Isotopic Studies of the Kutch Volcanics of Northwest India ................................................................................................ 87 3.1 Introduction........................................................................................... 3.2 Geological Setting................................................................................. 3.3 Deccan Stratigraphy.............................................................................. 87 88 89 3.4 Previous Work-Sr and Nd Isotopic Data .............................................. 3.5 Analytical Technique ............................................................................ 3.6 Results ................................................................................................ 3.7 Comparison with the DVP .................................................................... 3.8 Discussion ............................................................................................. 3.9 Conclusion ............................................................................................ APPENDICES 90 91 91 93 94 95 ................................................................................................ 107 A Analytical Techniques .......................................................................... B Sample Locations.................................................................................. C Photomicrographs ................................................................................. 107 118 120 REFERENCES ................................................................................................ 122 BIOGRAPHICAL SKETCH .............................................................................. 149 vi LIST OF TABLES Table 1.1: Major Oxide concentration in weight percent. ................................... 26 Table 1.2: Trace Element and REE concentration (in ppm) of the Pumpkinvine Amphibolites (PCF) and the Galts Ferry Gneiss (GFG) ............................................................ 27 Table 1.3: Trace Element and REE concentration (in ppm) of the Hillabee Greenstones (HG) and Hillabee Dacite (dacite) ....................................................................................... 28 Table 1.4: Sr and Nd Isotopic Data...................................................................... 29 Table 1.5: Rb-Sr isotopic data for whole rock samples....................................... 30 Table 1.6: Nd Model Age of metasediments from eastern Blue Ridge ............... 30 Table 1.7: U-Pb analysis of Galts Ferry Gneiss zircons by LA-MS-ICPMS ...... 31 Table 1.8: U-Pb analysis of detrital zircons from meta-sandstone by LA-MS-ICPMS 33 Table 1.9: U-Pb analysis of Mulberry Rock Gneiss zircons by LA-MS-ICPMS 35 Table 1.10: REE concentrations of eastern Blue Ridge metasediments.............. 36 Table 2.1: Regional Geologic history of the north-central New Mexico ............ 67 Table 2.2: Stratigraphic nomenclature and lithologic description of supracrustal Proterozoic rocks of New Mexico........................................................................................... 68 Table 2.3: Chemical analysis, norm and modes of the Tres Piedras Granite ...... 75 Table 2.4: Rb-Sr whole rock analysis of Tres Piedras Granite............................ 76 Table 2.5: Rb-Sr analysis of the mineral phases of the Tres Piedras Granite...... 77 Table 2.6: Tres Piedras Granite zircon analysis by LA-MC-ICPMS .................. 78 Table 3.1: Stratigraphic nomenclature and thickness of the southwestern Deccan Formations ................................................................................................ 97 Table 3.2: Trace element and Pb-isotope results ................................................. 98 Table 3.3: Elemental ratios of Re'union type source at 65 Ma ............................ 94 vii LIST OF FIGURES Figure 1.1: Generalized geological map of southern Appalachians ................... 37 Figure 1.1A: Generalized geologic map of eastern Alabama and western Georgia Blue Ridge terranes ................................................................................................ 38 Figure 1.2: Bimodality of Hillabee Greenstone sequence and Pumpkinvine Creek Formation ................................................................................................ 39 Figure 1.3: Total alkali vs. silica diagram for the Hillabee Greenstone sequence and Pumpkinvine Creek Formation ........................................................................... 40 Figure 1.4A: Hillabee dacite and GFG classification using Shand's index ........ 41 Figure 1.4B: HG and PCF amphibolite plotted on AFM diagram ...................... 41 Figure 1.5: Ti vs Zr plot of greenstones- dacites and PCF amphibolite - GFG .. 42 Figure 1.6: Co-variation diagrams of relatively immobile elements for HG and PCF amphibolites ................................................................................................ 43 Figure 1.7: Tectonic discrimination diagrams for the PCF amphibolites and HG 44 Figure 1.8: La/10-Y/15-Nb/8 diagram for the PCF amphibolite and HG .......... 45 Figure 1.9: Tectonic discrimination diagram for GFG and Hillabee dacite ....... 46 Figure 1.10: Spider diagram plot of PCF amphibolite and HG .......................... 47 Figure 1:11 CI chondrite normalized plot of PCF amphibolite and HG ............ 48 Figure 1.12: CI chondrite normalized plot of GFG and Hillabee dacite ............ 49 Figure 1.13: Initial Sr vs. epsilon Nd isotopic plot ............................................. 50 Figure 1.14: Rb-Sr isochron of GFG samples .................................................... 51 Figure 1.15: Two chemical groups of GFG defined by REE concentrations. .... 52 Figure 1.16: U-Pb ages of GFG zircons ............................................................. 53 Figure 1.16A: CL images of GFG zircons .......................................................... 54 Figure 1.17: Nd isotopic evolution diagram of the eBR metasediments ............ 55 viii Figure 1.18: NASC normalized REE plot of the eBR metasediments ............... 56 Figure 1.19: U-Pb ages of the detrital zircons of eBR ........................................ 57 Figure 1.20: Rb-Sr whole rock age of Mulberry Rock Gneiss ........................... 58 Figure 1.21: U-Pb ages of Mulberry Rock Gneiss zircons ................................. 59 Figure 1.22: Epsilon Nd of Ordovician arc felsic magmas of the Appalachians. 60 Figure 1.23: Model for passive mechanism of lithosphere extension ................ 61 Figure 1.24: Model for accretionary orogen in southern Appalachians during Ordovician ................................................................................................ 61 Figure 2.1: Map of Proterozoic basement uplifts and associated major faults in north-central New Mexico ................................................................................................ 80 Figure 2.1a: Map of Tres Piedras Granite sampling locations ........................... 80 Figure 2.2: Whole rock Rb-Sr isochron of Tres Piedras Granite ........................ 81 Figure 2.3: Rb-Sr mineral isochron of Tres Piedras Granite .............................. 83 Figure 2.4: U-Pb ages of Tres Piedras Granite zircons by LA-MC-ICPMS ...... 85 Figure 2.5: LA-ICPMS measurement of concentration ratios across a feldspar grain from Tres Piedras Granite ................................................................................................ 86 Figure 3.1: Structural-tectonic features of southern Asia and Indian Ocean basin 99 Figure 3.2: Initial Nd vs. initial Sr isotope plot of the Kutch samples ............... 100 Figure 3.3: Trace element-REE plot of the Kutch basalts ................................. 101 Figure 3.4: La/Nb vs Ce/Pb and La/Sm vs Sm/Yb in the Kutch basalts ............ 102 Figure 3.5: Pb and Sr isotopic variation in Kutch Basalts .................................. 103 Figure 3.6: Melting model of garnet peridotite to produce the alkali basalts ..... 104 Figure 3.7: Mixing model for the Kutch tholeiites to generate the trace element-REE pattern ................................................................................................ 105 Figure 3.8: Mixing model for the Kutch tholeiites to generate the Sr and Nd isotope ratioserate the trace element-REE pattern ............................................................................ 106 ix ABSTRACT In the southernmost Appalachians, bimodal volcanics of Pumpkinvine Creek Formation (PCF) and its proposed equivalent, the Hillabee Greenstone (HG) have indistinguishable ages (~460 Ma) and trace element-REE pattern similar to an arc/back-arc type setting. 143Nd values of felsic members of the HG and PCF indicate involvement of Grenville crust during petrogenesis. U-Pb dates (900-1500 Ma) of detrital zircons in PCF meta-sandstone cluster around 1100 Ma. Nd-model ages of the Ashland-Wedowee Supergroup metasediments range between 943-1439 Ma and cluster around 1000 Ma. Rb-Sr whole rock and U-Pb zircon dates of the Mulberry Rock Gneiss also demonstrate an Ordovician age (~460 Ma). It is concluded that the PCF-HG arc formed on the Laurentian continental margin on Ashland-Wedowee sediments during Ordovician and remained outboard of the continent until final closure during Alleghenian orogeny. Geochronological investigations of the Tres Piedras Granite of northcentral New Mexico have revealed a sharp discordancy between Rb-Sr whole-rock and U-Pb zircon ages. Analyses of fifty individual zircons (most concordant) by LA-MS-ICPMS yield a ~1730 Ma magmatic crystallization age. Rb-Sr whole-rock isochron ages from separate localities are 1490+/-20 Ma and 1497+/-42 Ma. Sphene/whole-rock/biotite isochron ages and initial 87Sr/86Sr ratios from separate localities are indistinguishable from those of whole-rock isochrons. In both cases feldspar plots above the isochrons and appeared to be an open system as evidenced by 4% difference in 87Sr/86Sr in the core and rim of feldspar. Taken altogether, these data are interpreted to reflect a large (kilometer) scale redistribution and rehomogenization of strontium isotopes during an independently, well-documented metamorphic event in the region. The geochemical character of the Kutch volcanics, northwest of Deccan Traps, India, have been investigated in order the magma's origin. Sr-Nd-Pb isotopic ratios and trace element patterns identifies three end members: Reunion plume-type alkali basalts, Mahabaleshwar-type alkali basalts and crustally contaminated tholeiites. The first type of alkali basalts that can be generated by very low degree of partial melting (1.6-1.8%) of Reunion plume like source at garnet stability field; the tholeiites can be explained by crustal contamination of Indian-MORB like magma. High 207Pb/204Pb (15.61-15.83) ratio of the tholeiites agrees well with the Pb-isotopes of local Archean crust. x CHAPTER ONE GEOCHEMICAL AND GEOCHRONOLOGICAL CONSTRAINTS ON THE ORIGIN AND EVOLUTION OF THE EASTERN BLUE RIDGE, SOUTHERN APPALACHIANS. 1.1 Introduction Forty years ago Tuzo Wilson (1966) proposed the theory of the Wilson Cycle by studying the eastern margin of North America. At its simplest, the Wilson Cycle describes the opening and closing of ocean basins - plates rift into pieces and diverge away from each other and new ocean basins form in between, then there is a reversal in motion, convergence of the two rifted plates followed by a plate collision, and mountain building. The rock records of the eastern margin of North America preserves evidence of at least two complete Wilson cycles– starting with breakup of the supercontinent Rodinia and ending with the assembly of the supercontinent Pangaea and finally rifting of Pangea and opening of the modern Atlantic Ocean (Thomas, 2005). The breakup and rifting of Rodinia is well documented in the stratigraphic rock records of the eastern margin of North America that include synrift sedimentary and igneous rocks, post-rift unconformity and early post-rift sedimentary strata. Isolation of rifted Laurentia was complete by early Cambrian (ca. 530 Ma) (Thomas, 2005). The early extensions span between 760–650 Ma (e.g., Aleinikoff et al., 1995; Hogan and Gilbert, 1998; Thomas et al., 2000; Cawood and Nemchin, 2001; Owens and Tucker, 2003) followed by pervasive rifting (620–545 Ma) along the eastern margin of Laurentia and terminating with the evidence of few late stage rifting of microcontinents between 540–530 Ma. The Appalachian mountain chain was created by broadly three Paleozoic orogenies after the rifting of supercontinent Rodinia: the Ordovician-Silurian Taconic Orogeny, the Devonian Acadian orogeny and the Pennsylvanian-Permian Alleghanian orogeny. Evidence of deformation due to Taconic orogeny was first recognized in the Hudson valley of New York and the affected areas stretch from western Newfoundland south to about New York State. Deformation caused by Taconic Orogeny are also identified much farther south in Georgia and Alabama. Taconic orogeny was dominated by A-type subduction against the Laurentian margin resulting in the closure of the Proto Atlantic (also called the Iapetus Ocean). Island arcs forming due to the subduction were obducted onto the Laurentian margin. Few examples include Cowrock, Cartoogechaye and Tugaloo terranes in the southern Appalachians and the Chopawamsic, Potomac and Bultimore terranes in the central and northern Appalachians (Hatcher et al., 2006). The arc building and amalgamation produced numerous plutons (Steltenpohl et al., 2005; Bream, 2003), penetrative deformation, and as high as granulite facies metamorphism (Hatcher and Butler, 1979; Eckert et al., 1989; Moecher et al., 2004). Taconic orogeny generated a huge amount of new crust that was amalgamated along the southeastern margin of Laurentia. The Acadian Orogeny occurred during the early Devonian Period. It was a continuing collision to the island arc of the Taconic Orogeny (Hatcher, 1989). The Acadian Orogeny was primarily a continent-continent collision between Laurentia (North America) and Baltica (Europe). The 1 Northern Iapetus Ocean was completely closed and the Acadian orogeny re-deformed some of the rocks previously affected by the Taconic Orogeny, and produced a mountain belt in , Nova Scotia, Newfoundland, Maine, New York, England, Ireland, Scandinavia and Greenland.that are as high as the Himalayas The Acadian Orogeny of North America is equivalent to the Caledonian Orogeny in Europe. Following the collision of Europe and North America, the newly formed Appalachian Mountains started to shed tremendous amounts of sediment westward into the foreland. The sediments were primarily conglomerates, quartz arenites, breccias, arkose sandstones etc (Harrison, 2002, Hatcher, 2005). Near the mountains, the sedimentary rocks were primarily terrestrial like river conglomerates and red shale, sands, etc. Westward, these rocks undergo facies changes into black shales and beach sands (Baranoski and Riley, 1987). In Pennsylvania and New York, the Catskill Delta represents the terrestrial-marine component (Van Tassell, 1987). In other parts of North America, sediment erosion from the Appalachian Mountains wasn’t a dominating event and the shallow marine seaways existed supporting large amounts of carbonate sedimentation and oolites and reefs formation (Drivet and Mountjoy, 1997). The Alleghenian Orogeny of the southern Appalachians was the third and last of the great orogenies that affected eastern North America during the Paleozoic. The final collision of North America with Africa, transported a huge composite crystalline thrust sheet – the Blue Ridge Piedmont megathrust sheet that included pre-Alleghanian metamorphic rocks to at least 350 Km inside the North American craton (Hatcher et al., 2006). The Alleghenian orogeny was completed by 265 Ma with the formation of Pangea and completion of the first (Paleozoic) Wilson cycle. The Alleghenian suture is marked by the boundary between the easternmost subsurface Suwannee terrane and Carolina terrane. From the fossil assemblage of the cover sequence (Gondwana basement covered by fossiliferous Ordovician to Devonian sandstones and shales) it seems likely that the Suwanee originated as part of African Gondwanaland (Mueller et al., 1994; Pojeta et al., 1976) and remained on the Laurentian side during breakup of Pangea. The three episodes of orogeny are identified in the Appalachians but they are not substantiated along the entire length of the mountain chain. (Moecher et al., 2004: Hatcher, 1978; Hatcher et al., 1989). The southern Appalachian mountain chain records protracted continental orogeny related to multiple extensional and contractional events. These events have been attributed to a range of convergent tectonic mechanisms, including subduction, arc accretion, and continental collision. Both direct and oblique convergence mechanisms have been proposed, and subduction is interpreted to have been directed both toward and away from Laurentia, the core of North America that existed at the time (Hatcher et al., 1987). The eastern Blue Ridge province was affected by all of the major events that shaped the southern Appalachian orogen and is a key link in elucidating the assembly of the Appalachian portion of the North American continent. However, important aspects of the evolution of the eastern Blue Ridge remain obscure. This work presents new data on zircon contained within the metasediments in the eastern Blue Ridge of Alabama and Georgia. The geochemical and geochronological data have bearing upon the formation, evolution and accretion of the most inboard ‘suspect’ terranes in the southern Appalachians and will compare this sequence chemically with the Hillabee Greenstone of the Talladega belt-western Blue Ridge. These data will have implication for the genesis of the 2 eastern Blue Ridge in the southern Appalachians and provide the implicit constraints on the tectonic history of this region. 1.2 Regional Overview The Blue Ridge province of the southern and central Appalachians includes low to high grade metamorphic rocks that have been thrust northwestward over the unmetamorphosed sedimentary rocks of the Valley and Ridge province. It is divided into eastern and western portions by the east-dipping Hayesville and related faults in the northeast and by the Allatoona fault and Hollis Line Fault in the southeast (Figs. 1.1, 1.1A). The post-metamorphic Allatoona fault and Hollis Line fault places the structurally complex eastern Blue Ridge assemblage of chemically immature clastic metasedimentary rocks, mafic to ultramafic bodies, variably deformed felsic intrusive rocks and Grenville basement (eg.Tullahh Falls Formation) over the more mature metasedimentary rocks and basement of the western Blue Ridge (Rankin, 1975; Hatcher, 1978). Regionally the eastern Blue Ridge is a composite terrane that may include parts of Laurentian outer margin cover sequence as well as accreted components of accretionary prism, ophiolitic and island arc affinity. The eastern Blue Ridge (EBR) is generally similar lithologically to the adjacent Piedmont. The Piedmont is also an exotic composite terrane that extends some 700 km southwestward from the frame of the Sauratown Mountains window in North Carolina to the Coastal Plain overlap in Alabama (Hatcher, 2002). The Brevard fault zone, that separates the eastern Blue Ridge from the Piedmont, is a major structure with a protracted, polyphase history of displacement (Hatcher, 1978). The western Blue Ridge is demonstrably part of Laurentia, but the origin of the EBR is controversial. The presence of alpine peridotite, mafic-ultramafic complexes has led some to suggest that the EBR may be an exotic, far-traveled terrane that was accreted during Paleozoic orogeny, with the mafic rocks being remnants of the closed ocean basin (for example, Horton, Drake, and Rankin, 1989; Willard and Adams, 1994; Zen, 1981; Williams and Hatcher, 1983; Shaw and Wasserburg, 1984). The Pine Mountain Window (Fig. 1.1) is exposed in the Piedmont on the Alabama promontory. It exposes 1.1 Ga Laurentian Grenvillian basement of granulite- and upper-amphibolite-facies granitic gneisses and a kyanite-sillimanite-grade platformal cover sequence (Hatcher, 2006). The easternmost accreted terrane in the central and southern Appalachians is the Carolina superterrane situated to the east of Piedmont (Fig. 1.1). During Rodinia rifting numerous microcontinents were formed in the prevailing ocean between Laurentia and Gondwana, proximal to Gondwana, loosely termed as Peri-Gondwanan terranes. The Carolina superterrane was formed by amalgamation of a series of these microcontinents of peri-Gondwana derivation and accreted to the Laurentia during the mid-Paleozoic (Rast and Skehan, 1983). The Carolina superterrane is primarily composed of Neoproterozoic mafic to felsic volcanics of volcanic arc origin and some volcaniclastic sedimentary rocks. Several plutons ranging in age between 550 to 600 Ma intrude the arc complex (Hibbard et al., 2003). This entire assemblage was metamorphosed to upper amphibolite to greenschist facies during Cambrian. This is evidenced by ~ 530 Ma plutons that intrude the metamorphic rocks (Dennis and Wright, 1997). Cambrian and Ordovician (identified from fossil assemblages) clastic sedimentary rock overlies the arc assemblage (Samson, et al., 1982; Koeppen et al., 1995). The western Carolina 3 superterrane is metamorphosed to a higher grade with a 350-360 Ma metamorphic overprint and contains numerous Devonian to Carboniferous younger granitoids and gabbros (McSween et al., 1991; Hibbard et al., 2003) probably intruded during the docking of the Carolina terrane with the Laurentia during mid to late Paleozoic. 1.3 Study Area and General Description of Lithotectonic Units The study area includes the Hillabee greenstone belt, Ashland-Wedowee belt, Mulberry rock gneiss and the Pumpkinvine Creek Formation (PCF) of the southern Appalachians (Fig. 1.1A). Fresh unaltered samples were collected using several criteria including quality of mapping and documented tectonic or stratigraphic significance of individual units. Hillabee Greenstone: The Hillabee Greenstone (HG) belt structurally occurs on top of the lower greenschist facies Talladega Group in the Talladega belt. The Talladega Belt represents the most outboard Laurentian margin cover sequence and the Talladega Group is the frontal metamorphic thrust sheet of the southernmost Appalachians of Alabama and western Georgia. Like the PCF it is also a sequence of bimodal volcanics (the maximum thickness is ~2.6 Km at places ) with the mafic (~75%) and felsic members (~25%) being tholeiitic metabasalt and calcalkaline metadacite/rhyolite respectively. The metavolcanic complex is in a thrust contact with the underlying fossiliferous shallow marine Devonian to earliest Mississipian (?) metasedimentary rocks of the Talladega belt. The mafic member is primarily low-potassium tholeiitic metabasalt and basaltic meta-andesite referred to as greenstone herein and the felsic member is metadacite referred as Hillabee dacite herein (Tull et al., 2006). Unlike PCF there is not much of metasedimentary units present in the HG belt. Rare, thin layers of micaceous quartzite and sericite phyllite occur locally within the metavolcanic complex and are the only non-volcanic member of the HG belt. Textural evidences like rarely preserved ophitic, intergranular, and porphyritic textures is indicative of the extrusive nature of the mafic rocks in form of basalt flows and/or basaltic ash (Tull and Stow, 1979, 1980; Stow, 1982). The felsic volcanics that are interlayered with the mafic rocks occur in form of thick (upto 165 m) mineralogically and chemically homogeneous, laterally continuous, tabular sheet-like bodies. Compositionally they are porphyritic meta dacite/rhyolite. The metadacite shows bands of polycrystalline quartz alternating with finegrained mica and opaque minerals. The rock contains porphyroclasts of hornblende, actinolite, albitic plagioclase, and quartz (Tull et al., 1998) interpreted to be phenocrysts. Strain is evidenced by undulose extinction of quartz grains and weakly kinked lamelleae of plagioclase. Hillabee dacite extended over an area greater than 100's of sq. km. and are interpreted as largevolume pyroclastic ash flow crystal tuffs (Tull et al., 1998). Samples of the greenstone and Hillabee dacite were collected for chemical analysis. Structurally overlying the Talladega belt is the EBR allochthon. 4 Pumpkinvine Creek Formation: The Pumpkinvine Creek Formation is a linear belt of metavolcanic rock sequence extending from near the Alabama-Georgia border into northern Georgia for a length of about 100 Km. It is included as part of >220 Km long Dahlonega Gold Belt (DGB) terrane that is strike continuous and in the same structural position with the AshlandWedowee belt (Fig. 1.1, 1.1A). The importance of the PCF lies with the fact that it is structurally adjacent to the Laurentian portion of the Appalachian making it the most inboard ‘suspect’ terrane (because of uncertain affinity, lack of basement it is considered to be “suspect” with respect to Laurentia) in this portion of the southern Appalachians. Lithologically the PCF is a sequence of bimodal volcanics in the outcrop scale and mapscale containing fine grained amphibolite (referred as PCF amphibolites herein), felsic gneiss (termed as Galts Ferry Gneiss or GFG member) and minor amounts of pelitic schists (Canton Schist) and subordinate metamorphosed sandstone. The mafic and the felsic units are interlayered from centimeter scale to tens of meters. The Canton schist possibly represent background sedimentation (Holm, 2006) interpreted from its sporadic occurrence and in sharp contact with the metavolcanic rocks of the PCF. McConnell (1980) identified discontinuous but regionally mappable banded iron formation interlayered with the PCF amphibolites. Due to lack of any ordered stratigraphic occurrence of the three units of the PCF (mafic, felsic and the metasediments) they are all considered different facies of the PCF. The amphibolite is fine grained, the major mineral phases being amphibole (75-80%) and plagioclase (~20%) and the accessory mineral phases include epidote, garnet and actinolite and/or chlorite (retrograde assemblage of amphibole). Typical volcanic textures include relic amygdule filled with epidote and plagioclase phenocrysts are common. Pillow structures have also been reported by several workers (McConnell, 1980; Abrams and McConnell, 1984). The felsic metavolcanic lithology (GFG) occurs in two modes. In places it forms thin (0.1-0.5 m) layers possibly representing rhyolitic ash eruptions. The second type of occurrence is as thick, lenticular bodies up to 2 km thick extending along strike. Major mineral composition of the GFG is plagioclase and quartz and enclaves of intermediate to mafic composition rich in amphibole is common. Accessory minerals include muscovite and to a lesser degree garnet. The larger bodies of felsic gneiss are likely large felsic eruptions. The third and volumetrically least significant lithology of the PCF is the aluminous Canton Schist, composed mainly of quartz sericite/muscovite ± garnet and biotite. This likely represents background sedimentation during periods of volcanic quiescence (Holm, 2006). Presence of extremely low sediments in archetypal back-arc basins such as the Lau basin-Havre trough (portion of the Tonga-to-New Zealand back-arc system) have been reported by Gamble and Wright, 1995. In the northern portion of the study area the Chattahoochee fault (a late structure in the kinematic sequence) demarcates the upper boundary of the PCF and emplaces higher grade, migmatitic, predominantly metasedimentary units of the Sandy Springs Group (Higgins and McConnell, 1978) structurally above the PCF. The units structurally above the Pumpkinvine Creek Formation, described as the New Georgia Group (Abrams and McConnell, 1981), contain a mixture of pelitic and chemical sediments (iron-bearing quartzite), minor amphibolite and 5 locally altered ultramafic pods. These units also hosts numerous intrusive bodies varying in composition from gabbroic to intermediate and felsic gneisses (eg. ~430 Ma Austell gneiss dated by Higgins et al. 1997) and many more unnamed intrusive bodies. These units likely correlate to the southwest into Alabama with part of the Ashland-Wedowee belt (Fig. 1.1). The intrusive units are similar in age and composition to the granitoids (including the 460 Ma Kowaliga and Zana granites) of Alabama as described by Russell et al. (1987) and Drummond et al. (1997). The age of the New Georgia Group and correlative units of the Ashland-Wedowee belt to the southwest is likely late-Proterozoic to earliest Paleozoic, and may represent slope-rise sediments along the eastern margin of Laurentia (Tull ,1978). The intrusives might have formed because of subduction-related slab melting beneath Laurentia (ex. Elkahatchee Quartz Diorite at 490 Ma) or later magmatic events during Taconic and Acadian orogenies (Drummond et al.,1997). Samples were collected from both the felsic (GFG) and mafic units of the PCF and Canton Schist. Ashland Wedowee Belt: East of the Talladega belt and west of the Brevard fault zone in Alabama and west Georgia is a broad belt of medium to upper amphibolite facies schists, gneisses and amphibolites of the Ashland – Wedowee belt. These rocks are interpreted to be late Precambrian metamorphosed sedimentary rocks and minor basalts (Thomas et al., 1980). The Ashland does not continue beyond the state line but the Wedowee and Emuckfaw extend northeast into Georgia (Fig. 1.1A) Exposed over much of the northern Alabama and western Georgia Piedmont (Fig 1.1, 1.1A) is a distinctive graphite bearing metasedimentary sequence of slate, phyllite, quartzite and schist, called the Wedowee Group, interpreted as the regressive phase of a more complex sedimentary cycle involving many other rock units (Neathery, 1973). Protoliths for the metamorphic lithologies possibly are arenite, siltstone and claystone interpreted to be a deep water environment deposit. The Ashland Super Group (Tull, 1978) underlies the Wedowee Group and is characterized by metabasalts called the Poe Bridge Mountain amphibolites mostly concentrated in the lower part of the sequence and metasediments with an aggregate thickness of few thousand meters. The lower Ashland can be separated from the other units with relative ease because it is the only stratigraphic unit containing abundant metabasalts (amphibolites). Ashland Super Group is a thick sequence of metasedimentary rocks. Garnetiferous biotite schist and siliceous graphitic muscovite schist are common. Less abundant rock types include kyanite-quartz schist, sillimanite-quartz schist and quartzite (Vernon, 1973). Although arial separation and apparent lithologic differences prohibit any direct correlation, McConnell and Abrams (1984) speculated that the rocks of the new Georgia are at least in part equivalent to the Ashland Supergroup. This is based primarily on the fact that both the New Georgia Group and the Ashland Supergroup contain metavolcanic rocks and similar types of sulfide hosted ore deposits. The authors also suggested that the rocks defined as Wedowee formation in Alabama (Tull, 1978) are equivalent to the rocks of the Sandy Springs Group. This correlation is based on lithologic similarities and the association of both Sandy Springs Group 6 and Wedowee Group with volcanic bearing rock groups (i.e., New Georgia Group and Ashland Supergroup, respectively) Fresh samples of quartzite, muscovite schist, garnetiferrous biotite schist from the Ashland and Wedowee Group were analyzed for Nd isotopic composition to determine the model age of the provenance. Mulberry Rock Gneiss: Holm (2001, 2001a) studied the Mulberry Rock Gneiss (MRG), that occurs as a structural window (eyelid window). The MRG occurs within a recess where the eastern Blue Ridge rocks and ‘terrane bounding Allatoona fault are tightly folded into the structural recess’ (Holm, 2001, 2001a). The unique position of the MRG (Fig. 1.1A) caused previous workers to conclude that the yet undated rocks were Grenville basement to the Talladega belt and the Talladega Group lies unconformably on the basement. Higgins et al., 1988 described them as equivalent to the nearby Grenville aged Corbin metagranite. Holm 2001, 2001a described MRG as medium grained, slightly metaluminous, two mica granite where muscovite percent dominates over biotite. The other major minerals are quartz, plagioclase and K-feldspar. Zircon, epidote and allanite occus as accessory phases. Tectono-stratigraphic models of the MRG structural recess are tested in the light of the Rb-Sr whole rock age, U/Pb zircon dates and Nd model ages of the MRG. The theories that are tested are either 1) whether MRG represents Grenville basement to the Talladega Group or 2) it is one of the numerous Paleozoic intrusion within the eastern Blue Ridge terrane. Fresh samples of MRG was collected to look at the U-Pb age of the zircon, Rb-Sr whole rock age and Nd model age to check the emplacement age as well as its connection with the Grenville basement. 1.4 Purpose of the Study This study will focus on the geochemical and geochronological constraints on the origin and evolution of different tectonic divisions of the eastern Blue Ridge, southern Appalachians. A major problem in Appalachian geology is the lack of Taconian deformation in the southern Appalachians. The eastern margin of Laurentia was curved into bays and promontories during opening and closing of the Iapetus ocean. Talladega belt (now within Alabama recess) originated as a continental promontory. During continental collision, the protruding promontories should first record the deformation. The absence of any Taconic deformation events in the Talladega belt rocks (outermost preserved portions of the Laurentian margin in the southern Appalachian) implies that a major promontory of the Laurentian margin might have escaped the effects of Taconic orogeny (Tull, 1998). As in the central Appalachians, it is difficult to appeal to subduction beneath a proven Ordovician arc as a deformation mechanism for the southern Appalachians. This study will use geochemical and geochronological data supported by detailed field maps to investigate whether or not the PCF formation can qualify for the Ordovician Arc. The PCF and Hillabee, exist only as fragments, unlike the Taconic arcs of the New England Appalachians (e.g. Bronson Hill arc) and the Ordovician northern Appalachian arcs (e.g. Chopawamsic Terrane) that commonly include typical arc related features such as voluminous arc-related sediments, accretionary wedge, and evidence of accretion in the foreland. If the PCF and Hillabee are the remnants of the Ordovician volcanic arc, the tectonic setting, timing and 7 style of the Taconic orogeny in southern Appalachians and formation of the arc-related terranes took place in an environment very different from the Ordovician arcs of the northern and central Appalachians. Possibly the arc in the southern Appalachian formed during Ordovician but remained outboard of the Laurentian margin until terminal closure of the Iapetan Ocean. The goal of this research is to propose a tectonic model for the southern Appalachians during Taconic orogeny. The study will be based on new detrital/magmatic zircon ages, major and trace element analysis and Sr and Nd isotopic compositions of major tectonic units from southern Appalachians tectonic divisions west of the Brevard fault zone. 1.5 Analytical Techniques 1.5.1 ICP MS Trace Element Analysis Trace elements for the PCF amphibolites, HG, GFG and Hillabee metadacite were determined by solution ICP-MS analysis using a ThermoFinnigan Element. Small chips of each sample were handpicked and crushed in agate mortar and pestle. Thirty mg of each sample was weighed into screw top Teflon beakers and dissolved with 2 ml of 3:1 distilled HF-HNO3. After drying on a hot plate at 120oC, the samples were allowed to reflux overnight in concentrated HNO3. Samples were then dried again and brought to a final volume of 60 ml with 2% HNO3. The sample solutions were further diluted in the ICP vials for target concentration of 100 ppm TDS. One ppb of Indium was added as internal standards and drift corrections for each analyzed mass were applied by interpolating with the internal standard. The solution ICPMS analyses of the amphibolites and greenstones were calibrated against a single solution of well-characterized Hawaiian basalt USGS standard BCR-1 prepared identically to the samples. For the felsic samples G2 was used as the USGS standard. BHVO-1 was also prepared like a sample and calibrated against the BCR-1 and G2 to check the precession of measurement. 1.5.2 Sr-Nd Isotopic Analysis Fresh whole rock samples chosen for whole rock analysis of Sr and Nd were powdered in an agate mortar and dissolved in a 3:1 mixture of 2X distilled HF: HNO3. Separation of Sr and Nd followed ion exchange procedures employing 4.5 ml of AG50W-X8 9-cm bed-length ion exchange resin. Nd was separated as a bulk REE fraction and eluted in 6 N HCl. Nd and the REE fraction was further separated on a 1.2 ml, 6 cm bed-length column of Ln resin SPS. Measurements were made on a Finnigan-MAT 262 mass spectrometer. Strontium measurements are reported relative to the measured value of the E&A Sr standard of 87Sr/86Sr = 0.708000 ± 11 (2sd, n = 15) and Nd relative to the measured LaJolla standard 143Nd/144Nd = 0.511848 ± 11 (2sd, n = 12). n= number of analysis, each analysis being the average of 100 ratios. 1.5.3 Rb/Sr Ratio Measurement in ICP MS A precise method for Rb and Sr ratio measurements for samples (GFG) with low Rb/Sr ratios (<10) has been developed. Fresh whole rock samples were powdered in a agate mortar and 8 dissolved in a 3:1 mixture of 2X distilled HF: HNO3 and then diluted to obtain Rb and Sr concentrations in the 10 ng/ml range. These solutions were then directly measured without elemental separation of Rb and Sr from the rock matrix or from one another. Measurements of the peaks: 85Rb+, 86Sr+, and 88Sr+, were performed by ICP-MS, using a ThermoFinnigan Element in low-resolution mode. Ion count rates were in the 10-100 MHz range. Isobaric interference of 86 Kr+ from the Ar plasma gas was monitored on acid blank solutions, and was found to be negligible (<100 Hz). The precision on standard solutions was ±0.2%. The method was applied to two USGS rock standards, G-2 and AGV-1, and yielded Rb/Sr ratios that were accurate to within 2-4%. These data are within the reported uncertainties of the Rb/Sr ratios of the two standards (Gladney et al., 1988). 1.5.4 Rb/Sr Ratio Measurement by Isotope Dilution The MRG samples have a higher range of Rb/Sr ratios and were measured by standard ID technique. Approximately 100 mg of powder was spiked with isotopically enriched tracers of 84 Sr and 87Rb. The samples were then dissolved in screw cap teflon beakers with distilled 3:1 HF and Nitric acid. The dissolved samples were converted into a chloride form and Rb and Sr were separated using a cation exchange column containing AG 50W-X8 resin with a volume elution of 2.5N HCl. The total analytical blanks for Sr and Rb were less than 100 picogram and 1 nannogram respectively. Isotopic analysis were made on a Finnigan-MAT 262 mass spectrometer. Rb and Sr were loaded on W- single filament with TAPH (10g H2O + 0.5g TaCl5 + 3g 0.1N H3PO4 + 0.5g conc. HF) solution. The measured Sr isotopic compositions were corrected for the isotopic composition of the spike and normalized to a 87Sr/86Sr value of 0.1194. 1.5.5 LA-ICP-MS Single Zircon Analysis Samples were collected from GFG, MRG and a thin meta-sandstone unit interlayered with the Canton Schist. At each sample locality, 1–2 kg of fresh whole rock samples were collected. Samples were prepared for analysis using standard crushing and separation techniques, including heavy liquids and magnetic separation. Apparently inclusion-free zircons were then hand-picked under a binocular microscope. At least 100 zircons from each sample were mounted in epoxy and polished. U–Pb analyses were performed at University of Arizona, Tucson with a Micromass Isoprobe multicollector Inductively Coupled Plasma Mass Spectrometer (ICP-MS) equipped with nine Faraday collectors, an axial Daly collector, and four ion-counting channels. Zircons were ablated with a New Wave DUV 193 nm Excimer laser ablation system. All analyses were conducted in static mode with a laser beam diameter ranging from 30 to 50 μm. Contribution of Hg to the 204 Pb mass position was removed by subtracting measured background values. Isotopic fractionation was monitored by analyzing a Srilankan zircon standard, which has a concordant TIMS age of 564 ± 4 Ma (Dickinson and Gehrels, 2003). This standard was analyzed once for every five unknowns in detrital grains and once for every three unknowns in magmatic zircons. Uranium and Thorium concentrations were monitored by analyzing a standard (NIST 610 Glass) with approximately 500 ppm Th and U. The calibration correction used for the analyses was 2– 9 3% for 206Pb / 238U and approximately 2% for 206Pb / 207Pb (2-sigma errors). The lead isotopic ratios were corrected for common Pb, using the measured 204Pb, assuming an initial Pb composition according to Stacey and Kramers (1975) and uncertainties of 1.0, 0.3 and 2.0 for 206 Pb / 204Pb, 207Pb / 204Pb, and 208Pb / 204Pb, respectively. For samples that are younger than 1 Ga ages are considered reliable if five or more analyses performed in different grains yield overlapping 206Pb / 238U ages. This strategy is used because of the low precision of 206Pb / 207Pb ages for young grains, making concordance/discordance a poor criteria for determining reliability. Clustering is also a better criteria for reliability than concordance given that Pb loss and inheritance in young systems can create concordant ages that are significantly younger or older than the true ages. Such analyses could be concordant but would not define a cluster, and would accordingly be rejected as unreliable. Three samples were analyzed for U-Pb geochronology in this study, GFG, MRG and detrital zircon from a meta-sandstone. Ages of 50-60 zircon grains were measured from each igneous sample and ~100 grains were measured from the sedimentary sample. Results and errors are reported in Tables 1.7, 1.8 and 1. 9. Each line in the Table represents a spot analysis. 1.6 Results 1.6.1 Major and trace-element geochemistry The major and trace element concentrations of the PCF, greenstone, GFG and Hillabee dacite are shown in Tables 1.1, 1.2 and 1.3. The bimodal nature in terms of SiO2 of both mafic and felsic components of the PCF and the HG sequence is shown in a histogram (Fig.1.2). It is important to note that there is no intermediate component associated with these units. Using the total alkalis vs. SiO2 diagram (Fig.1.3) by Lebas et al. (1986), the PCF amphibolites and greenstone samples plot within the basalt field and the GFG and Hillabee dacite plots in the rhyolite and dacite field respectively. Tull et al., (1978) reported some basaltic andesite from the Hillabee that are plotted in the diagram too. The GFG composition is extremely sodic (Table 1.1) and is strongly peraluminous (Fig. 1.4A). The Hillabee dacites also plot in the peraluminous field but they are more near the boundary between metaluminous and peraluminous field. Traditionally it is thought that strongly peraluminous magmas are derived from mature pelitic sediments and produce S-type igneous rocks. Miller (1985) describes more likely sources of these rock types to be continentally derived immature sediments and intermediate to felsic igneous rocks as well as certain products of partial melting in metaluminous mafic sources. Miller (1985) describes criteria for identifying S-type granites (Chappell and White, 1974) and pelitic parentage of igneous rocks. The GFG and Hillabee metadacite do not fit either criteria and when compared to a typical pelitic composition (Gromet et al., 1984), show depletion in rare earth element (REE) chemistry. In the AFM diagram (Fig. 1.4B) the amphibolites of both the Hillabee and PCF defines a tholeiitic trend where as the felsic components follow the calc alkaline trend. In the Ti vs. Zr plot (Fig. 1.5A & 1.5B) the felsic components and the amphibolite shows two parallel trends depicting some amount of differentiation within each group but indicates that the felsic 10 component is possibly not derived by simple fractional crystallization of the associated mafics. Further discrimination of the PCF and greenstone using trace element concentrations will help elucidate its origin. Commonly used geochemical parameters in distinguishing arc and marginal basin rocks include light rare earth element (LREE) fractionation leading to enrichment and Nb-Ta anomalies relative to MORB values. High field strength elements (HFSE) provide values typically unmodified by moderate grade metamorphism. These elements are used as robust parameters in discrimination of tectonic environment (Cabanis and Lecolle, 1989; Meschede, 1986; Pearce and Norry, 1979; Pearce et al., 1984; Shervais, 1982; Wood 1980). 1.6.1a PCF Amphibolites and Greenstones. Because the rocks of the PCF have undergone amphibolite grade metamorphism and the greenstone belongs to the lower green schist facies, it is essential to check the mobility of elements during regional metamorphism and subsequent alteration. Using covariation diagrams (Kim et al., 2003) plotting significant elements against HFSE such as Zr, a linear trend suggests relative immobility of those elements. Figure 1.6 illustrates linear trends between Nb, Y, V, and the HFSE Zr in PCF and Hillabee mafic facies, thus signifying immobility of the elements used for tectonic discrimination and their utility in identifying tectonic environment of emplacement. Several discrimination diagrams have been developed for mafic rocks based on HFSE including Wood (1980), Pearce et al. (1984), Meschede (1986), Cabanis and Lecolle (1989), and Shervais (1982). Discrimination diagrams (Figs. 1.7A,B) using HFSE including Meschede (1986) and Pearce et al. (1984) help constrain the PCF amphibolites and greenstone as island-arc related yielding compositions in the range of N-MORB to volcanic arc basalt. Cabanis and Lecolle (1989) use the relationship between La and Nb to help discriminate between volcanic arc signatures and N-MORB signatures and include a field for BABB (Fig. 1.8) where PCF amphibolite and greenstone samples fall. High La relative to MORB is a signature of SSZ compositions (Hawkins, 1995, 2003). Decoupling of large ion lithophile elements (LILE) and HFSE in the supra subduction zone environment, termed “flux melting” (Gill, 1981), yields enrichment of LILE such as Ba, Rb, and Sr in BABB and is likely due to dehydration fluids associated with the subducting slab while HFSE or HREE remain unaffected because of their immobility in aqueous fluids (Hawkins, 1995, 2003). Flux melting paired with decompression melting melts more of the hydrous mantle wedge than either process alone can generate (Xia et al., 2003). When normalized to Primitive Mantle, PCF amphibolites display a modest LILE enrichment and HFSE near Primitive Mantle values (Fig. 1.10). Another typical BABB characteristic is the negative Ta-Nb anomaly (Kim et al., 2003). Rare earth element (REE) patterns of the PCF are extremely flat exhibiting minimal fractionation, which is typical of tholeiitic basalt (Fig. 1.11). The REE’s are considered relatively immobile and would not yield a supra subduction zone signature. The greenstone samples shows enrichment in LILE relative to the Primitive Mantle and a Nb depletion but the primary difference with the PCF samples being Pb enrichment which is one of the important features of the arc rocks (McCulloch and Gamble, 1991) (Fig 1.10). 3 out of 7 samples show moderate Zr and Ti anomaly, another characteristics of arc basalts. The REE 11 pattern are flat, similar to the PCF amphibolites though they show higher range of variability possibly due to olivine fractionation at the source (Fig. 1.11). Due to geochemical similarities and similar tectonic settings the PCF and Hillabee volcanic sequence might represent different parts of the same arc-back arc system. 1.6.1b Galts Ferry Gneiss and Hillabee dacite. Maniar and Piccoli (1989) recommend the use of major element data in tectonic discrimination of granitoids, however, elemental mobility in multiply deformed and metamorphosed terranes in the southern Appalachians limits this approach. In an analogous approach to that of basaltic rocks, Pearce et al. (1984) use multi-element plots emphasizing HFSE as parameters and normalize samples to hypothetical ocean ridge granite (ORG) in order to discriminate tectonic fields. In Nb versus Y and Rb versus Y + Nb (Fig. 1.9) and Yb versus Ta and Yb + Ta versus Rb (not shown), the GFG and the Hillabee metadacite plots in the volcanic arc granite field. All GFG samples exhibit moderate light rare earth element (LREE) enrichment and a modest Eu anomaly (Fig. 1.12) except for sample SC318G that is a very thin (<0.1 m) felsic layer and was probably an ash deposit. A couple of samples show slight positive Eu anomaly that might be due to the presence of plagioclase phenocrysts. The Hillabee dacites are LREE enriched with (Ce/Yb)n ranging from 11.2 to 14.3. The resulting REE pattern thus display a distinct negative LREE slope typical of calc-alkaline dacites in orogenic belts (e.g. Gill,1976). Absence of Eu anomaly indicate that plagioclase fractionation is not the cause of geochemical variations. 1.6.2 Sr and Nd Model Values The Nd and Sr isotope systems are most useful in providing insight as to the nature of, and mixing between, sources during magmagenesis. The Sr and Nd values of the PCF amphibolites, GFG, greenstone and Hillabee dacite are shown in Table 1.4. The measured Sr and Nd isotopes are recalculated at 460 Ma (Russell, 1978; Thomas et al, 2001). The PCF amphibolite samples exhibit epsilon Nd (T460 Ma) values of +3.3 to +7.7 and the greenstone exhibit εNd (T460 Ma) values ranging between +2.01 to +6.84, typical of a depleted mantle source composition (Table 1.4). The GFG epsilon Nd (T460 Ma) values of –3.21 to 4.65 and the even more negative Hillabee dacites (ε Nd at T460 Ma between –3.9 to –5.5) indicate that this magma was not derived directly from fractional crystallization of the magma that formed the amphibolite protolith. It is likely evidence of sampling of a more isotopically evolved component (Fig. 1.13). This may include supra subduction zone fluids, subducted sediments, assimilation of the older basement rocks or melting of the basement. 1.6.3 Rb-Sr Whole Rock Age Rb-Sr whole rock analyses of the GFG (Table 1.5) are employed to help constrain timing of crystallization and possibly provide insight into petrogenesis of the felsic facies of the PCF. Fig. 1.14 is a plot of the 87Sr/86Sr vs. 87Rb/86Sr data. For reference a hypothetical 391 Ma isochron is drawn. Out of the eight samples analyzed, they appear to cluster into two groups with different initial 87Sr/86Sr ratios. The two groups can be separated based on their REE pattern too (Fig. 12 1.15). Four samples with low 87Sr/86Sr initial (~0.705), similar to the associated amphibolites have a relatively flat LREE and HREE pattern (La/Yb = 1.2 to 4.2) where as the samples with higher initial 87Sr/86Sr (~0.708) displays a steeper slope for the LREE (La/Yb = 4.2 to 23.7). It is likely that the magmagenesis is complex and may tap into multiple sources. The GFG samples possibly define two isochrons of similar ages (424±25 Ma and 450±66 Ma) with significant errors but differing initial 87Sr/86Sr. The age within the error margin might represent either the crystallization age or the regional metamorphic age of 366 ± 25 Ma (Russell, 1978) of the EBR metasediments. A lower 87Sr/86Sr initial value of 0.7059 is similar to the associated (PCF) metabasalts and is likely closely related to the primary melt by fractional crystallization of the metabasalts and subsequent eruption (Fig. 1.14). A higher 87Sr/86Sr initial value of 0.7082 proves the existence of another, more radiogenic/isotopically-evolved, source bearing differing initial ratios that may include a mixed (mantle source plus recycled crust and/or sediment) source, possibly by assimilation and fractional crystallization (AFC). There are also the possibilities of Sr mobility due to subsequent thermo-tectonic events, or initial 87Sr/86Sr variability in the magma source. The Hillabee metadacites yielded a Rb-Sr whole rock age of 395 ± 20 Ma with a initial 87Sr/86Sr ratio of 0.7082 (Durham, 1993). This age is very close to K-Ar whole rock age of 399 ± 17 Ma of the slates of the Talladega Slate Belt reported by Kish, 1990 and within the range of uncertainty for the whole rock Rb-Sr age of 366 ± 25 Ma for the Ashland-Wedowee metasediments and the K-Ar age of 382 ± 14 Ma for hornblende in the Hillabee metadacite reported by Russell, 1978. Thus Durham interpreted the Rb-Sr whole rock age of the metadacites as the age of lower green schist facies regional metamorphism. The initial 87Sr/86Sr ratio of 0.7082 is similar to the higher initial values of the GFG. The high initial Sr isotopic ratio observed for the GFG and Hillabee metadacite is within the range of ratios observed by Pushkar et al., 1973 in the Lesser Antilles. The higher ratios (0.709) observed in the Lesser Antilles was explained by assimilation of crustal material, either continental detritus or marine sediments. 1.6.4 Zircon Data Incorporation of older crustal material by the PCF volcanics can be tested by looking at the zircon inheritance (if any) of the GFG (Table 1.7). Zircons from the GFG samples were analyzed optically under a petrographic microscope and a scanning electron microscope (SEM) in backscattered electron (BSE) mode. Approximately 50 zircons were analyzed and almost all zircons yielded Ordovician ages and displayed igneous morphologies (e.g. euheadral crystals) with no detectable optical zoning. When plotted in a histogram the modal class ranged from 440 to 480 Ma (206Pb/238U age) that comprised of ~ 35 grains (Fig 1.16). A weighted 206Pb/238U average of the 35 grains yielded an age of 460.2 ± 9.7 Ma. This agrees with the 463 ± 3 Ma age of the GFG (near Allatoona Dam) reported by Thomas et al., (2001). A few grains (n~4) under the Cathode Luminescence (CL) image show the presence of cores in the zircons (Fig. 1.16A). The core age is Cambrian. No older zircon (Grenville age) was found. Zircons from Hillabee metadacites were not analyzed in this study but Russel, 1978; Russel et al., 1984, McClellan and Miller (2000) and Tull et al., (2006) reported an Ordovician age (~460-470 Ma) for the complex. Some of the ion probe data set on the Hillabee metadacite (McClellan et al., 2005) shows older Cambrian core at ~500 Ma., i.e. 30-40 m.y. older than the interpreted Hillabee crystallization age similar to that observed for the GFG. 13 1.7 Nd Model Age and Detrital Zircon Ages of the Metasediments of the eastern Blue Ridge. A direct approach of testing the sediment source is looking at their Nd-model age. Nd and Sm may undergo appreciable fractionation during extraction of crustal material from the mantle but are typically not fractionated during process such as erosion and metamorphism. Sm-Nd isotopic data for a rock can be used to calculate a model age representing either a crustal extraction or average crustal residence age. The significance of the model age depends on the proportion of the juvenile verses ancient sources contributing to the rock. A TDM age is the time when a sample and the mantle source had identical isotopic composition and a TCHUR represents the extraction of the source rock from the Chondritic Uniform Reservoir (CHUR). In practice, Nd isotopic data and Nd model ages are used as chemical finger prints to distinguish between crustal provinces and terranes within an orogen (e.g. Patchett & Ruiz, 1989; Bennett & DePaolo, 1987). Table 1.6 summarizes isotopic data from the metasediments of the Ashland-Wedowee belt and the Canton Schist that is interlayered with the PCF. The Ashland-Wedowee sediments generally exhibit a range in TCHUR ages from 964 m.y to 1119 Ma with one sample at 1439 Ma. Canton Schists range in age between 943 to 1137 Ma. In order to make the Nd model age internally more consistent a depleted mantle (DM) extraction model was tested. Nd isotopic values and Sm/Nd ratios of the DM for the range of ages for the depletion event of the two-stage evolution model were taken from Salters and Stracke 2004. Models were tested for a minimum depletion age of 1.8 Ga, which is similar to the minimum age expected from the Pb-isotope systematics in MORB (Hart, 1984; Tatsumoto, 1978). A maximum depletion age of 3.4 Ga inferred from the Re-Os isotope systematics. A 2.2 Ga was based on the assumption that continental crust extraction occurred at 2.2 Ga (Chase and Patchett, 1988; Condie, 2000; Galer et al., 1989) and if the DM is assumed to be complementary to the continental crust a similar average age is expected. The 3.4 Ga depleted mantle formation fits the model best (Fig. 1.17) and 12 out of 15 metasedimentary samples analyzed range in age between 1514 – 1580 Ma, two samples are slightly younger at 1450 Ma and 1482 Ma and one sample is 1892 Ma (Table 1.6). The difference in the model age might be due to difference in source character. In order to test the hypothesis the REE’s of the metasediments are plotted in a North American Shale normalized diagram (Fig. 1.18, Table 1.10). The normalized REE patterns of the metasedimentary samples are flat, do not show any visible difference and are 0.5 to 2.5 times more enriched than the North American Shale with the exception of the two young samples that have the highest concentrations of all the REEs. The sample with the oldest age has a flat REE pattern similar to the North American Shale. A suite of detrital zircons (n~100) separated from a meta-sandstone interlayered with the Canton Schist were analyzed in the same method as the GFG samples (Table 1.8). A concordia diagram and histogram derived from LA-MC-ICPMS ages of the detrital zircon grains shown in Fig 1.19. The dominant feature of the detrital zircon age distribution is a pronounced cluster of ages between 918 and 1190 which contains nearly eighty percent of the ages. Zircons of this age range are clearly the dominant signature of the Grenville terrane, as they are the most abundant age group. Age peaks at 511 Ma and 589 Ma (n=3) are somewhat curious as they are not ages representative of Laurentian zircon populations. In Newfoundland, Cawood and Nemchin (2001) 14 report synrift (572-628 Ma) detrital zircon with a local igneous source. Older peaks of 1351 Ma and 1462 Ma and also the TDM age of 1439 of one Ashland sample correspond to the ages of the eastern Granite-Rhyolite province (Becker at al., 2005). 1.8 Mulberry Rock Gneiss 1.8.1 Rb-Sr Results The Rb-Sr whole rock isochron for MRG is shown in Figure 1.20. The analytical error is smaller than the sizes of individual data points. The calculated isochron age is 467 Ma with an uncertainty of 16 Ma at 95% confidence level. The range of 87Rb/86Sr, from 5.1 to 51.9 is very large in this body (Table 1. 5). This suggests that any metamorphic events that led to the current fabric of the gneiss did not chemically or isotopically homogenize the body, at least not sufficiently to produce an isochron reflecting the time of metamorphism. More likely the linear isochron reflects the time of crystallization of the MRG from a granitic magma. The initial 87Sr/86Sr is 0.7076±0.0036. This value suggests that the source of the MRG magmas involve some fraction of crustal rocks. 1.8.2 Nd Model Ages. Table 1.4 summarizes the Nd-isotope data that exhibit a range in εNd of the MRG between –1.6 to –2.8. Thus it can be interpreted that the MRG had recycled older (Laurentian?) crust. 1.8.3 U-Pb Zircon Ages. Zircon grains (n~70) were separated from the MRG samples and analyzed by LA-MC-ICPMS. The 206Pb/238U ages show quiet a range with the highest number of grains plotting in the age range between 420 and 450 Ma (histogram plot of Figure 1.21, Table 1.9). Inherited components (?) and/or partially re-set older zircon xenocrysts have a range of ages from 450 to 600 Ma, with one concordant grain at 1068 ±20 Ma (Pb-Pb age-See Table). Pb loss seems to be a problem too for these samples with 30% of the grains ranging in age between 300 and 400 Ma. On a concordia diagram the error ellipses cluster around ~450 Ma with the upper intercept of the discordia line at 1049 Ma and the lower intercept at 345 Ma. The upper intercept represents Grenville inheritance while the lower intercept might be lead loss due to a thermal event. Thus in light of the Rb-Sr whole rock data, Nd isotope data and U-Pb zircon analysis it can be concluded that MRG is not part of the Grenville basement, however it might have originated by anatectic melting of older Grenville-aged rocks. 1.9 Discussion 1.9.1 Petrogenesis of Modern and Ancient Arc/Back-Arc Systems – Origin of the Bimodal Volcanics 15 It has been expressed that arc and back-arc basins appear to be simple tectonic environments, yet they often display tectonic and geochemical heterogeneity on the scale of kilometers and sometimes even within a single volcanic center (Gamble and Wright, 1995). Behind the arc, a back-arc basin is typically developed. This is the site of MORB-like volcanism that creates thin ocean type crust in an extentional tectonic environment behind the volcanic arc. It has been suggested that felsic volcanism accompanies early rifting in the back-arc setting (Marsaglia, 1995) yielding bimodal volcanism with an arc-like signature during these early stages, whereas the more mature back-arc basin volcanism displays mid-ocean ridge basalt (MORB)-like characteristics (Xia et al. 2003). In Xia et al.’s (2003) model of an Ordovician back-arc basin in the northern Qilian Mountains of China, it is suggested that early back arc rifting does not influence normal arc magmagenesis in the mantle. During initial rifting, some of the arc melts are diverted to the young rift axis yielding the possibility of incorporation of supra-subduction zone (SSZ) components and melting of the ambient mantle. A substantial backarc extension is necessary before induced mantle upwelling is sufficient to generate backarc-spreading-ridge lavas (White and McKenzie, 1989). From this perspective, it is impossible to generate backarcspreading-ridge lavas during the earliest island-arc rifting stages of back-arc-basin formation. Continued back-arc spreading and the consequential relative movement away from the locus of arc magmagenesis should yield normal mid-ocean ridge basalt (N-MORB) -like composition of the newly formed crust. Descriptive terminology of back-arc basin basalts (BABB) is variable in that it includes an array of arc-like to MORB-like compositions. Taylor (1995) suggests that modern back-arc volcanism exhibits diversity in the generation of magmas at back-arc spreading ridges and feels that it is not feasible to describe compositional differences in basalts relating to stages of maturity of a back-arc basin. There are probably three first-order components that influence the rock chemistry of evolving back–arc basins including melting of the incipient island arc during initial rifting, supra subduction zone fluids associated with the dehydration melting of the subducting oceanic slab and decompression melting associated with the mantle wedge. The first two components should be most important during the initial and early back-arc spreading and should decrease over back-arc maturation as the spreading axis progresses away from the arc and SSZ fluid source. Many back-arc basins, ancient and modern, show variation along the axis of back-arc spreading from the incipient rifting, yielding island arc-like basalts, to a mature back-arc where basalts are MORB-like (Xia et al., 2003; Gamble and Wright, 1995; Hawkins, 1995 and 1976) giving rise to the term back-arc basin basalt (BABB). It has also been shown that MORB-like basalts can occur concomitantly and proximal to arc-like eruptions in the initial spreading of an active back-arc (Fryer, 1992). Modern or young analogues of back-arc bimodal magmatism are limited. Cole et al. (1995) describe the back-arc volcanism in the Taupo Volcanic Zone, located at the North Island of New Zealand as the southwest extension of the oceanic back-arc Havre Trough where magmagenesis of a large volume of rhyolite here is likely composed of crustal materials incorporated with a mantle source. The Okinawa Trough exhibits a bimodal petrogenetic nature. Shinjo and Kato (2000) put forth a model explaining the origins of basaltic, rhyolitic and hybrid magmas generated at the Okinawa Trough. The authors have identified two types of rhyolite (and associated basalts) and hybrid andesite related to different petrogenetic processes. While the Okinawa Trough is a continental back-arc basin, the modes of magmagenesis are limited to fractional crystallization of the associated basalts, assimilation and fractional crystallization (AFC) of basaltic magma, and hybridization of evolved basalts and rhyolite derived from 16 fractional crystallization. All these processes involve a lower gabbroic crust. Silurian bimodal volcanism is described in the northern Appalachian Coastal volcanic belt of New England and New Brunswick. While it is documented as a Silurian continental extensional environment (Van Wagoner et al., 2002), identification as a back-arc basin remains enigmatic because of continental within plate geochemical signatures and island arc-like Nb anomalies. It is common for bimodal volcanism and the majority of highly silicic volcanism to occur in the early stages of rifting in continental regimes by underplating and anatexis of pre-existing continental crust and also in intra-oceanic regimes. Ringwood (1977) described the petrogenesis of a tholeiitic suite of rocks concomitantly with a calc-alkaline suite that is derived from partial melting of the tholeiitic magmas in source regions comprised of either subducted oceanic crust or mantle wedge materials and yield an initially high silica content. Subduction is a complex process that produces some characteristics igneous association with distinctive pattern of the major and trace element chemistry. Andesites and basaltic andesites dominate the island arc rock spectrum. Basalts, dacites and rhyolites are subordinate. In case of the PCF and the Hillabee greenstone sequence mostly basalts, dacites and rhyolites are present. Tull et al., (1978) reported some basaltic andesite in the Hillabee greenstone sequence. This apparent discrepancy in the lithologic association could be explained by the deep erosion that the PVC and the Hillabee sequence suffered and are preserved today as thin slivers of the remnant arc. If the exposed plutons of the eastern Blue Ridge represents the core of the arc (Drummond et al., 1997) then the entire arc carapace is eroded which might have been the andesitic and basaltic andesitic members of the Ordovician arc. The petrogenetic relationship between the felsic and mafic facies of the PCF and Hillabee is important in understanding the arc/back-arc evolution. It is difficult to determine whether these facies follow the same trend between tholeiitic and calc-alkaline because there are no intermediate values between the suites. The mafic facies displays a tholeiitic trend whereas the felsic GFG and the Hillabee dacite displays a calc-alkaline trend as shown in the AFM diagram (Fig. 1.4B). Shinjo and Kato (2000) use mass balance models to identify the contribution of fractional crystallization and assimilation/fractional crystallization (AFC) in the Okinawa Trough bimodal back-arc volcanic suite. It is important to identify a parameter useful in discriminating the process of generation of the felsic component. Rollinson and Windley (1980) use the relationship of Zr and Ti as an indicator of fractional crystallization as a source mechanism of felsic melt during magmagenesis. If fractional crystallization of felsic material from a mafic source is the main contributor to the genesis of the felsic magmas, the felsic material would either fall along the same trend-line as the mafic material or, after Ti-bearing minerals such as magnetite or titanite crystallize out, the felsic material would fall off the trend but with increasing Zr content. In figure 1.5A and 1.5B, it is clear that the felsic does not fall along the differentiation by fractional crystallization trend of the spatially related mafic material. It is unlikely that fractional crystallization is responsible for the generation of the felsic magma. It is important to examine possible source mechanisms for the GFG and the dacites. As it is unlikely that fractional crystallization is the major factor in genesis of the GFG, it is important to explore the possibility of supra subduction zone (SSZ) components such as dehydration and partial melting of a subducting slab, mantle wedge components and later decompression melting at a spreading ridge or partial melting of arc crust or a preexisting older continental crust. Further 17 discrimination of the PCF and greenstone by major and trace element will help elucidate its origin. The major chemical difference between the two felsic members, the GFG and the Hillabee dacites is in their silica content and alkali content. GFG has low content of potassium (K2O<2%) and high Na2O making it a sodium rhyolite/ trondhjemite. The flat REE pattern of the GFG ((La/Sm)n ranges from 1.54 to 4.47) also indicates fractionation from a mafic source. The Hillabee dacites have flat HREE and enriched LREE ((La/Sm)n ranges from 4.60 to 5.73) typical of calc alkaline magmas. Thus the dacites might be generated by a different process than the GFG which involved either partial melting of lower continental crust and/or interaction with the Laurentian continental crust. Because back-arc environments are complex and the compositions of magmas produced in these environments are variable with back-arc maturity, it is difficult to distinguish the contributing source to the GFG and Hillabee dacite magmas here. Because back-arc basins in their initial stages often rift an existing arc, it is conceivable that the GFG and Hillabee dacite magma includes supra subduction zone fluids, mantle wedge materials, and assimilation of the rifted arc. The possible variability in the source material makes it difficult to complete a mass balance of the felsic volcanics based on typical epsilon Nd values of these sources. Initial 87Sr/86Sr values of the PCF amphibolite vary between 0.7040 and 0.7069, though mostly cluster around 0.7045. The greenstone has a spread between 0.7039 to 0.7063 with a cluster around 0.7050. Variability could easily be caused by Sr or Rb mobility during metamorphism. Initial 87Sr/86Sr of the GFG samples are scattered between 0.7046 and 0.7084 while a low population, <0.706, and a high population, >0.708, exist (Figure 1.13). The population distribution within the GFG can be derived from Sr mobility during metamorphism, but could also further indicate that different sources contributing to the magma. GFG samples bearing similar initial 87Sr/86Sr values (~0.705) to that of PCF amphibolites might represent a liquid line of decent from PCF magma, whereas initial 87Sr/86Sr values >0.708 are likely derived from a more isotopically evolved source. The metadacites however has a smaller spread for initial 87Sr/86Sr ranging between 0.7069 to 0.7082 and distinctly crystallized from a more evolved source than the associated greenstones. Many of the Ordovician arc-related rocks of the Appalachians show evidence of interaction with Laurentian (Grenville-aged) continental crust. The abundance of inherited zircons and very negative epsilon Nd values attest to the close interaction of these arcs with the preexisting Grenville/Laurentian crust (Coler et al. 2000). It is important to attempt to frame the possible relationships to other concomitant arc terranes in order to determine a possible reference frame for arcs eventually accreted onto Laurentia. Most Appalachian arc-related data is available from the northern portion of the orogen, and in a compilation, by Coler et al. (2000) and references therein, of available Nd isotopic data, a Nd isotopic evolution diagram (Fig. 1.22) shows that most northern Appalachian (and two southern Appalachian) arcs and arc-related felsic magmas sample at least Mesoproterozoic, Grenvillerelated crust like that of the Hillabee metadacites. This suggests that these arcs sampled a Laurentian crustal substrate. Only one arc, Exploits, does not have an epsilon Nd value that involves a large portion of Grenville crust. The GFG also shows similar Nd-isotopic characters 18 like that of Exploits arc. While this is not conclusive in excluding the PCF from being genetically related to Laurentian crust, it strongly suggests little to no input of Grenville or older materials incorporated into the magma. While trying to characterize the trace element content of the arc lavas it might be expected that the arc lavas represent a mix of three main components, depleted mantle wedge, subducted sediments and (altered) mafic oceanic crust in case of intra oceanic arc and continental crust in case of continental margin arcs. It has been recognized for sometime (McCulloch and Gamble, 1991) that a characteristics feature of island arc basalt is the depletion of high field strength (HFS) elements such as Nb, Ta and to a lesser extent Zr, Ti, Yb, Sc and Ni relative to large ion lithophile elements (LIL) such as Rb, Ba, Pb, Sr, U and Th. In the PCF amphibolites and greenstones there is modest depletion of Nb, and Zr, Ti in some samples. There is enrichment in Ba, Pb in the greenstone which are considered as signatures from dehydrating slab. Pb enrichment can partly be attributed to the sulfide mineralization in the greenstone. Sr enrichment in both the amphibolites and greenstones might be due to epidotization. The REE pattern of both the amphibolites and the greenstones are flat like that of MORB. Considering the flat REE pattern and modest enrichment of LIL elements with slight Nb a volcanic back-arc source is preferred. There is good geochemical evidence for transport of LIL enriched volcanic arc melt into BAB (Saunders and Tarney, 1984; Nakamura et al., 1989). 1.9.2 Tectonic Relationships On the basis of the new data, and taking into consideration the timing and nature of the interaction of oceanic and continental tectonic elements, a revised Cambrian –Ordovician tectonic history of the Laurentian margin is proposed in the southern Appalachians of Georgia and Alabama. There are still insufficient data to constrain many of the fundamental features of Laurentian margin tectonic development in the southern Appalachians and there are many aspects that remain equivocal. The sequence of events in the beginning of the first Wilson cycle can be summarized as follows: Grenville orogeny at ca 1 Ga resulting in amalgamation of the Rodinian supercontinent and extensive intrusion of syn tectonic granites (Karlstrom et al., 2001). Neoproterozoic rifting begun 240 Ma later in the eastern margin of Laurentia with the opening of Iapetus and culminating at ca. 560 Ma (Cawood et al., 2001). Breakup of Rodinia was multistage (e.g., Colpron et al., 2002 ; Harlan et al., 2003) spanning over ~230 m.y. and involving different parts of Laurentia. Magmatic events that record northward-migrating rift-pulses in the Appalachian Blue Ridge are the 760–745 Ma Crossnore event in North Carolina (Su et al., 1994 ; Aleinikoff et al., 1995 ). Other syn rift intrusions during Iapetus opening (e.g., Badger and Sinha, 1988 ; Aleinikoff et al., 1995) are Robertson River (735– 725 Ma), Battle Mountain events of Virginia Blue Ridge (705– 680 Ma) (Tollo and Hutson, 1996 ; Tollo and Aleinikoff, 1996 ; Tollo et al., 2004 ), Catoctin eruptive event (572–564 Ma) etc. The significant thickness of the metasedimentary package of the Ashland-Wedowee belt and equivalents of the southern Appalachians has caused previous workers to suggest a marginal basin/slope-rise facies origin for the rocks associated with the Proterozoic rifting of Rodinia, adjacent to the southeastern margin of Laurentia (Drummond et al., 1994 and 1997). Cawood and Nemchin (2001) identified 4 distinct zircon populations in the upper Neoproterozoic to 19 Ordovician Laurentian margin sedimentary sequence in the Newfoundland Appalachians that record a cycle of ocean opening and closing. The four identified groups are as follows (1) Archean zircons, age range 2850 -2600 Ma. Potential source- Laurentian hinterland. Age corresponds to zircon crystallization during major magmatic and tectonothermal events in the Superior craton. (2) Paleoproterozoic zircons, age range 1950 - 1750 Ma. Age corresponds to craton margin orogenic belts (e.g., Ungava, New Quebec, and Torngat). (3) Mesoproterozoic to early Neoproterozoic zircons, age range 1450 - 950 Ma. Age corresponds to Grenville orogen lithologies. (4) Neoproterozoic zircons, age range, 760 - 570 Ma. Age corresponds to numerous syn-rift igneous intrusions along Laurentian margin, now preserved within the Appalachian orogen. The stratigraphic thickness of the Ashland-Wedowee Supergroup in the southern Appalachians of Alabama and Georgia reaches >10 Km at places. The lithologic association of the AshlandWedowee Supergroup, predominantly deep water turbidite metasediments and subordinate intercalated tholeiitic rift/ocean floor basalts points towards a significant sediment source. The possibilities include 1. the sediments represents slope-rise sequence of a rifted continental margin that prograded out on the oceanic crust (now preserved as the tholeiites) 2. intra oceanic complex. The later option is improbable due to the relative proportion of sediments to basalt. The Ashland-Wedowee Supergroup is in fault contact and tectonically lies immediately above the Laurentian outer shelf rocks. Presence of Middle Proterozoic zircon xenocrysts in numerous plutons (product of anatectic melts of EBR units) that intrude the Ashland-Wedowee sediments indicates that Alabama EBR formed near or at the Laurentian rifted margin. It has been suggested that the Ashland-Wedowee Supergroup sediments qualifies as the best and possibly only candidate for the late Precambrian clastic slope/rise outer margin prism along the Alabama promontory during Rodinia rifting (Tull, 1978; Thomas et al.,1980; Stow et al., 1984; Drummond et al.,1988; 1994; 1997). In case of the Ashland-Wedowee group sediments because there are no distinguishing characteristics like fossil assemledge or basement clasts to ascertain the age and composition of the basement beneath the Ashland Supergroup, Nd model ages (based on CHUR extraction) were computed to determine the source characteristics of the Ashland-Wedowee Belt sediments. Majority of the samples exhibited ages between 950-1150 Ma, with one outlier at 1440 Ma. The ages are typically of Laurentian Grenville basement (the older age is similar to the GraniteRhyolite province) and suggests that Ashland-Wedowee belt sediments were supplied from Laurentian margin. Increasing Nd isotopic data now exist in the literature for crystalline rocks, specially the granitoids of the Blue Ridge. The Nd isotopic data for the metasediments overlap the field defined for the Grenville basement data (recalculated at 700 Ma, approximate time of deposition of the Ashland metasediments) defined by Carrigan et al., (2003) and Hatcher et al., (2004) (Fig. 1.17). The oldest TDM age reported by Bream et al., (2004) is 1.9 Ga from the Dahlonega Gold belt. Eastern Blue Ridge metagraywacke and quartzites from Tallulah Falls Formation ranged in age between 1.24 to 1.61 Ga (Bream et al., 2004). The TDM ages of the metasediments obtained from Ashland Supergroup and Canton Schist agrees well with the previous data. 20 The detrital zircon data is obtained from one sample within the Pumpkinvine Creek formation that shows a pronounced age cluster between 918 to 1190 Ma which is similar to the Grenville magmatic event. Few pre-Grenville components like Middle Proterozoic grains (1351-1462 Ma) are also present. In the absence of detrital zircon data on the Ashland Wedowee samples it is difficult to interpret the old age of 1892 Ma but middle Proterozoic and late Archean grains have been identified from eastern Blue Ridge sediments and the Dahlonega Gold belt sediments by Bream et al., 2004. Majority of the detrital zircons analyzed from Dahlonega Gold belt metasandstones reflect a North American source except for one sample that contains Gondwana (?) (1.8-2.0 Ga) zircons, in addition to Grenville and older North American detrital zircons (Bream, 2003). The Gondwana ages have not been reproduced in the other analyzed samples (Bream, 2003; Hatcher et al., 2006). The zircons in the one sample with Gondwana affinity could have been derived from the Laurentian Penokean orogen that occurred during the same time. During Rhodinia supercontinent rifting, numerous crustal fragments and an array of complex depositional basins between Laurentia and the South American craton was created. Within the southern Appalachian crystalline core there is exists strong evidence for both Laurentian and exotic sediment sources. The predominating detrital zircon population is aged between 1.0–1.25 Ga from the metasediments of southern Appalachian that were inboard of the Carolina terrane. This evidence suggests that these basins were placed near a Grenvillian source. This interpretation is consistent with the nonconformities that exist between many basement and cover contacts in the Blue Ridge. Orthoamphibolites commonly interlayered with graphite-quartz schist and quartzite and garnetquartzite/schist are common in the upper part of the Ashland Group. They range in thickness from centimeters to several 100 m and make up about 7% of these equivalent groups (Tull et al., 2006). A volcanic or plutonic protolith for the amphibolite is suggested from field relationships (interlayered, locally gradational, concordant contacts with surrounding metasediments) and textural evidence (orthopyroxene relics, surrounded by coronas of hornblende). Geologic setting combined with geochemical, textural, and mineralogic criteria led geologists to interpret these rocks to be tholeiitic ocean-floor basalts intercalated with deep water metasediments. (Tull, 1978; Thomas et al., 1980; Stow et al., 1984; Drummond et al., 1988). Small lenses and pods of ultramafic rocks are locally present in the amphibolite (Reynolds, 1973; Mies and Dean, 1994). Lithologic contacts between this thick pile of amphibolites with intercalated metasediments indicates that they are not faulted, but are interlayered with metasedimentary and other possible metavolcanic lithologies. There is complete lack of any primary facing data in these units, but it is highly unlikely that this >10 km thick sequence is regionally overturned. It is suggested that the amphibolites originated as ocean floor tholeiites and are related to the initial rifting of the supercontinent Rhodinia. Celar Sengor and Burke, 1978 described the process of evolution of continental rifts can occur in two ways: active rifting or passive rifting. 1. Active rifting – it is initiated by upwelling of magma in the asthenosphere thus often called mantle-generated rifting. In case of active rifting the zone of uplift and the zone of thinning extend for many kilometers far beyond the rift margin. 2. Passive rifting – it is caused by extensional forces in the lithosphere thus referred to as lithosphere-generated rifting. The zone of uplift in passive rifting and the zone of thinning are on the surface of the earth and confined to the rift margin. Fig. 1.23 summarizes the common 21 concepts conceived for passive extension models. Plate tectonics is the primary driving force for lithospheric extension in the passive model. The buoyant sub-lithospheric mantle passively upwells beneath the thinned area (McKenzie, 1978). Subsidence followed by sedimentation occurs continuously, both during syn-rift (initial subsidence) and post-rift (long-term subsidence) stages. Thus mantle melting in this case is a late consequence of the dynamic stretching of lithosphere (White and McKenzie, 1989). The amagmatic rifted margins (ARM) typically have a weak crust dominated by faults and abundant serpentinite. Different rift margins are characterized by difference in magma budgets (extremely low magmatism for the ARM). Thus the lithosphere beneath also vary from ‘depleted mantle beneath Volcanic Rifted Margins and undepleted mantle beneath ARMs’ (Stern, 2004). Though Cenozoic examples of passive rifted margins are rare but there are several well documented passive margins initiated rifting during the Mesozoic time. Examples include the west Iberia-Newfoundland conjugate margins and Alpine Tethyan margins. In the former case the final rifting phase and sea floor spreading was during Early Cretaceous (~133Ma) (Whitmarsh et al., 2001) and the later experienced the final phase of rifting during Triassic/Early Jurassic presaging the opening of the Liguria-Piemonte ocean where sea floor spreading dates back to 160- 165Ma ago (Whitmarsh et al., 2001). Both the above mentioned cases are characterized by magma-poor margins with margin-parallel deep-water zones that indicates successive stages of margin evolution. The continental crust along the margin is thinned and dissected by numerous low angle detachment faults. Sub-continental mantle exhumation follows faulting and typically the mafic melts volumetrically increases oceanward and merges with the oceanic crust (Dean et al., 2000; Boillot et al., 1989).The current lithologic association and stratigraphic thickness of the Ashland-Wedowee belt syn-post rift sediments and the presence of a very low percent of amphibolites might be related to the initiation of passive margin rifting of the Rhodinia supercontinent during the Neo Proterozoic time at least in the southern part of the Appalachians. The occurrence of granitic intrusion is widespread within the Ashland Wedowee belt that were emplaced between ~490-430 Ma (Russell, 1978; Russell et al., 1987; Higgins et al., 1997). Drummond (1997) suggests that the oldest intrusive Elkahatchee Quartz Diorite (EQD) formed by subduction and melting of a westward dipping oceanic slab beneath Laurentia during the earliest Ordovician. The younger granitoids (like Rockford type) are thought to accompany discrete dynamothermal metamorphic events (Russell, 1978). Zircons from the Ordovician-early Silurian Austell (Higgins et al. 1997) and Mulberry Rock gneisses include a fraction of inherited Grenville aged (~1.1 Ga) zircons and zircon cores suggesting that these formed while incorporating or intruding into Laurentian-affinity, Grenvillian crust. Cambrian to Late Ordovician plutonic bodies within the EBR likely formed the core of a continental margin arc (Drummond et al., 1997). Tull et al., (2006) suggested that as the Proterozoic- EBR strata served as the arc basement of subduction-generated magmas associated with a continental arc and the present erosional level would place this sequence several kilometers beneath the former arc volcanic carapace. If a ~460 Ma Hillabee and PCF were cogenetic with the plutonic complexes, then it could have constituted part of that overlying volcanic carapace in an arc to back-arc position. If the main part of the volcanic arc overlay the current EBR it would have been stripped by erosion and be absent at the current erosion level. A cogenetic association of the Hillabee and PCF with the EBR plutonic arc core suggests that B-type subduction may have spanned a period 22 of ~35 m.y., based upon the ages of the plutons. A similar span may also be reflected in the GFG and Hillabee dacite zircon data with Cambrian core at ~500 Ma., i.e. 30-40 Ma older than the interpreted GFG and Hillabee dacite crystallization age. Thus, in modeling subduction events related to a cogenetic EBR and PCF-Hillabee, the time span of subduction is assumed to be ~30-40 Ma. It has been suggested by previous workers (e.g. McConnell, 1980; McConnell and Abrams, 1984; Higgins et al., 1988; McClellan et al., 2005; Holm and Das, 2005) that the Hillabee Greenstone of Alabama and the PCF are equivalent units. The Hillabee and the PCF have very similar geochemical characteristics and occupy a similar structural position adjacent to the Talladega Belt-western Blue Ridge. Geochemical, isotopic and geochronological evidence suggests that the PCF and Hillabee are both Ordovician volcanic arc/back arc that marked the beginning of Taconic orogeny in the southern Appalachians. Both the GFG and the Hillabee dacite lack any inherited Grenville zircon but the ε Nd values for the dacites are more negative that the GFG indicating some interaction with older crust. The negative ε Hf values (Tull, personal communication) of the Hillabee dacite zircons attests the interaction of the Hillabee with an older crust. If the PCF and Hillabee are parts of the same arc-back arc system then the subduction zone might have stepped across a transform fault. In that case PCF were formed due to subduction against an oceanic crust where as the Hillabee formed on the continental crust. The background sedimentation that accompanied the arc building was receiving detritus from Laurentia as evidenced by the Nd model age of the Canton schist and detrital zircon of Grenville affinity in the meta-sandstone of PCF. However, there is no evidence of deformation along the southeastern Laurentian margin until the Devonian-Mississippian. Thomas (2004) suggests that because there is no deformation in the palinspastic location of the Talladega belt during the Taconic or Acadian orogenies and that the accreted terranes remained at or southeast of the rifted margin of Laurentia until accretion during the Alleghenian orogenic event. The Hillabee is in thrust contact with the Talladega Group and has the same metamorphic grade and concordant metamorphic fabric. Thus the time of thrusting is thus bracketed by the age of the uppermost member of the Talladega Group and metamorphic age. The table (Table I) below summarizes the available age data and their possible interpretations regarding the metamorphic age of the EBR and thrusting age of HG on Talladega belt. The (327-333) Ma 40Ar/39Ar biotite, hornblende ages of the amphibolite facies EBR rocks are interpreted to be metamorphic age of Talladega belt caused by emplacement of EBR over Talladega belt (Kish, 1990, Steltenpohl et al., 2005). This interpretation seems unlikely because if EBR is emplaced directly on cold Talladega belt sediments the hot basal part of the EBR would show evidence of extensive retrograde metamorphism which is not observed in the field (Tull et al., 2006). 23 Table I : Available age data and their possible interpretations regarding the metamorphic age of the EBR and thrusting age of HG on Talladega belt. Material dated Age Dating References Possible age interpretation technique Upper Erin 245-375 Fossil Tull et al., Deposition age of Erin shale Shale Ma assemblage 1989; and the youngest Gastaldo, stratighaphic age of the 1995 Talladega belt. Talladega belt ~370-400 K/Ar, Rb/Sr Wampler et The oldest age (older than slate ages, some Ma and 40Ar/39Ar al., 1970; the age of the Erin shale) from Erin shale. whole- rock Tull, 1982; possibly is from detrital Kslate ages Kish, 1990; bearing minerals in the Durham, 1993 metapelites (Kish, 1990). 40 39 Talladega Belt 321-334Ma Ar/ Ar McClellan et Cooling age and not a peak metasedimentary analysis of al., 2005 metamorphic age. and metaigneous white micas. rock Hornblende and 333.8 ±1.7 40Ar/39Ar ages Kish, 1990, Metamorphic age of Biotite from Ma Steltenpohl et Talladega belt caused by EBR al., 2005 emplacement of EBR over 327.4 ±1.6 amphibolite Talladega belt. Ma facies rocks. Syn to post ~ 366-370 U-Pb zircon ages Russell, 1978; Peak metamorphic age of and Rb/Sr Russell et al., EBR. kinematic Ma whole rock ages. 1987; igneous Steltenpohl et intrusions in al., 2005 EBR. Similar metamorphic ages obtained from EBR and Talladega Belt suggest that metamorphism and accretion of the Hillabee Greenstone, and the PCF, were synchronous with and completed during the Alleghenian orogeny. Palinspastic restoration of the paleotectonic position of the Alabama/Georgia eastern Blue Ridge places it eastward to the present location of the Pine Mountain internal basement massif. This has been done by strike perpendicular restoration of Talladega belt carbonate shelf strata and lower Paleozoic foreland (Thomas, 2004). A Laurentian affinity of the Ashland-Wedowee belt sediments makes it a suitable candidate for the clastic continental slope/rise outer margin prism that was forming over the Neoproterozoic – Cambrian margin of southeastern Laurentia. From the available stratigraphic and structural data and in the light of the new geochemical and geochronological data the following constrains are placed while trying to predict a model for the origin and evolution of the eastern Blue Ridge, southern Appalachian: A) the Ashland – Wedowee sediments is interpreted as slope rise deposit adjacent to Laurentia, B) the HG and PCF formation are part of the same Ordovician suprasubduction arc-back arc system probably involving Mesoproterozoic continental crust C) HG currently lies directly above Laurentian outer shelf strata whereas PCF formation is a part of the eastern Blue Ridge sequence D) 24 Mulberry rock gneiss is one of the numerous Ordovician-Silurian pluton and is not a Grenville basement in the Talladega belt. The geochemical signature that the HG and PCF formed as part of a continental arc system, their present structural position, and age similarity of the eastern Blue Ridge plutonism argue strongly that the HG-PVC is native to Laurentia and was not part of an exotic island arc. If the PCF is equivalent to the Greenstone, then accretion must take place after the Devonian/early Mississippian and prior to metamorphism constrained by 40Ar/39Ar ages of 334-320 Ma (Barineau and Tull, 2001; McClellan et al. 2005). The PCF and Hillabee, exist only as fragments, specific to the back-arc region, unlike the Taconic arcs of the New England Appalachians that commonly include typical arc related features such as voluminous arc-related sediments, accretionary wedge, and evidence of accretion in the foreland. Figure 1.24 summarizes the tectonic model proposed for the southern Appalachians during the Ordovician Taconic Orogeny in this study. If the PCF and Hillabee are the remnants of the Ordovician volcanic arc, the tectonic setting, timing and style of the Taconic orogeny in southern Appalachians and formation of the arc-related terranes took place in an environment very different from the Ordovician arcs of the northern and central Appalachians.The arc in the southern Appalachian formed during Ordovician but remained outboard of the Laurentian margin until terminal closure of the Iapetan Ocean due to continent-continent collision between Laurentia and Africa. 25 Table 1.1. Major Oxide concentration in weight percent. Sample Unit SiO2 Al2O3 Fe2O3T MgO M68 M201B M243 M268 M271 M272 SM1100 M67 D1 D2 BG995 BG999 Y 23 C Y 24 Y 131 NT 911 BH 166 SC 318G TA 7 US 41A Y 126 Y 172 TA 16 TA 17 SC 254 SC 318A AC 78 BH 57 G G G G G G G Metadacite Metadacite Metadacite Metadacite Metadacite GFG GFG GFG GFG GFG GFG GFG GFG PCF PCF PCF PCF PCF PCF PCF PCF 47.46 47.07 48.04 49.24 51.04 51.21 48.05 67.71 66.45 66.32 65.70 65.90 75.28 73.64 76.75 74.13 78.11 76.29 73.88 73.79 50.29 49.12 47.79 50.47 43.85 50.26 48.51 48.34 16.44 15.57 15.40 14.17 13.88 14.76 13.90 15.44 16.30 16.34 16.90 16.74 13.67 14.31 14.56 14.83 13.37 12.61 14.19 12.83 15.12 15.63 17.19 14.93 17.67 14.37 14.47 16.05 CaO 10.70 8.53 13.58 12.54 10.07 10.74 12.93 7.55 11.26 11.73 8.51 12.42 13.15 6.44 10.93 9.21 8.28 12.31 10.75 11.28 14.21 4.03 1.89 4.26 3.95 1.81 3.63 4.17 1.97 3.56 3.47 1.87 3.62 3.89 1.62 3.89 1.98 0.40 1.62 3.95 0.71 1.39 2.13 0.29 0.13 2.66 0.47 1.34 2.05 1.28 0.86 2.90 1.22 0.74 2.39 0.70 2.49 3.85 1.43 1.43 11.37 7.66 11.46 13.97 4.28 10.00 9.60 8.56 13.63 11.60 6.96 11.20 14.01 4.98 16.11 13.18 7.76 9.10 12.47 7.69 12.78 11.24 9.26 10.70 Na2O K2O TiO2 P2O5 MnO % LOI 1.77 2.02 2.86 2.36 2.31 3.11 1.16 3.29 3.97 4.12 3.98 3.76 6.63 4.18 3.80 3.93 3.64 6.00 4.59 5.51 2.57 3.82 1.93 3.06 1.28 3.79 2.32 2.86 0.195 0.166 0.185 0.164 0.204 0.154 0.196 0.061 0.070 0.060 0.060 0.070 0.020 0.071 0.031 0.030 0.010 0.030 0.051 0.071 0.162 0.184 0.145 0.163 0.223 0.173 0.172 0.154 Wt % determined by ACME laboratory by ICP- emission spectrometry. Wt % corrected for LOI (Loss of ignition) Total iron reported as Fe2O3T G - Hillabee Greenstone Metadacite – Hillabee dacite GFG - Galts Ferry Gneiss PCF - Pumpkinvine Creek Amphibolites 26 0.05 0.22 0.05 0.04 0.15 0.07 0.04 2.77 2.88 2.93 3.37 3.12 0.26 1.42 1.95 2.30 0.46 0.04 1.17 0.35 0.14 0.15 0.07 0.06 0.07 0.04 0.06 0.05 0.98 1.22 1.35 1.18 1.54 0.68 0.37 0.36 0.45 0.54 0.39 0.41 0.14 0.27 0.21 0.22 0.19 0.23 0.20 0.27 1.07 2.39 0.94 1.34 1.47 1.18 1.18 1.06 0.082 0.114 0.123 0.092 0.153 0.041 0.021 0.091 0.080 0.110 0.100 0.080 0.010 0.061 0.010 0.051 0.031 0.020 0.051 0.031 0.101 0.347 0.062 0.142 0.203 0.081 0.121 0.113 2.1 1.8 1.8 1.6 1 1.6 1 1.9 1.0 1.3 0.9 1.2 1 1.3 1.8 1.3 2.4 1.6 1.2 1.8 1 2 2.4 1.6 1.5 1.6 1.2 2.9 Table 1.2. Trace Element and REE concentration (in ppm) of the Pumpkinvine Creek Amphibolites (PCF) and the Galts Ferry Gneiss (GFG). Y 126 Y 172 TA 16 TA 17 Unit Li Rb Sr Y Zr Nb Cs Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Yb Lu Hf Ta Pb Th U Sc Ti V Cr Co Ni Cu Zn Ga SC SC AC 78 BH 57 254 318A PCF PCF PCF PCF PCF PCF PCF PCF _ _ _ _ _ _ _ _ 2.60 1.80 2.40 0.50 0.80 2.10 2.70 0.90 165 327 200 198 646 106 223 229 31.6 50.8 19.8 30.5 43.9 27.4 29.9 25.2 66.4 188.4 51.2 81.6 100.3 67.8 69.0 57.2 3.3 9.3 2.5 4.4 4.1 2.1 2.8 2.6 0.10 0.08 0.20 0.05 0.20 0.13 0.20 0.18 14.9 25.7 59.5 13.7 8.6 12.3 13.9 40.2 4.9 12.6 3.6 5.6 7.3 3.9 3.8 3.1 12.5 29.9 8.6 12.8 14.6 9.1 9.8 8.3 1.89 4.07 1.28 1.99 2.40 1.47 1.53 1.26 10.20 21.70 6.70 10.90 13.40 8.30 8.60 7.40 3.50 6.10 2.40 3.50 4.20 3.00 2.80 2.60 1.05 2.31 0.83 1.24 1.62 0.85 1.09 0.85 4.22 7.35 2.99 4.38 5.46 3.82 4.16 3.44 0.78 1.26 0.50 0.77 1.09 0.70 0.72 0.66 4.99 8.22 3.19 5.04 6.73 4.37 4.87 4.19 1.05 1.64 0.64 1.07 1.43 0.94 1.02 0.86 3.02 4.84 1.94 2.95 4.36 2.71 3.02 2.39 3.08 4.74 1.90 2.97 4.56 2.63 2.73 2.54 0.46 0.76 0.28 0.43 0.71 0.44 0.45 0.39 2.00 5.60 1.70 2.30 3.00 2.00 2.30 1.90 0.30 0.60 0.30 0.30 0.50 0.20 0.20 0.20 0.20 0.80 0.10 0.30 0.50 0.20 0.40 0.20 0.50 1.00 0.30 0.40 0.60 0.30 0.10 0.30 0.20 0.40 0.10 0.10 0.60 0.10 0.10 0.10 41 35 36 41 42 44 47 49 6419 14315 5648 8042 8825 7067 7099 6359 258.0 371.0 232.0 315.0 479.0 314.0 326.0 295.0 _ _ _ _ _ _ _ _ 43.6 32.0 39.0 48.0 49.7 55.6 57.5 74.1 74 22 94 35 51 43 79 98 11.6 109.1 31.9 59.9 28.6 73.6 105.9 150.3 9 29 18 20 9 14 10 27 17.3 25.8 16.3 19.3 29.1 16.5 20.0 17.5 27 Y23C Y24 Y 131 NT BH SC TA7 US41 911 166 318G GFG GFG GFG GFG GFG GFG GFG GFG 5.36 4.72 1.41 4.64 7.96 1.00 2.52 5.33 13.90 39.68 60.73 75.75 12.96 0.40 33.25 10.52 58 89 64 129 99 25 194 48 25.0 14.4 8.4 9.8 11.2 9.1 8.5 34.9 94.6 44.6 75.4 45.6 26.7 33.8 57.0 44.9 7.6 6.2 5.9 5.3 1.7 3.5 3.2 5.0 0.38 0.27 0.40 0.72 0.13 0.04 0.26 1.05 53.0 336.1 461.0 627.6 115.8 15.5 399.5 64.9 32.0 8.4 4.7 9.7 5.6 2.8 20.0 10.7 64.6 21.6 35.9 14.1 14.7 6.5 38.6 24.6 7.10 2.44 1.06 1.88 1.44 0.94 3.07 3.04 24.72 9.53 3.69 6.72 5.70 4.31 9.42 12.92 4.62 2.27 0.82 1.35 1.28 1.15 1.48 3.42 0.51 0.58 0.31 0.55 0.28 0.37 0.58 0.71 6.18 2.35 2.47 1.60 1.46 1.03 3.39 3.23 0.72 0.40 0.18 0.23 0.23 0.21 0.24 0.69 4.07 2.78 1.36 1.44 1.60 1.53 1.21 5.27 0.84 0.59 0.31 0.31 0.35 0.32 0.25 1.15 2.59 1.80 1.00 1.00 1.14 0.99 0.80 3.65 2.48 1.74 1.02 1.03 1.11 1.01 0.84 3.47 0.36 0.27 0.17 0.17 0.18 0.17 0.14 0.53 2.66 1.43 2.21 1.37 0.83 0.83 1.57 1.37 1.09 0.70 0.65 0.56 0.32 0.37 0.88 1.03 4.28 20.59 19.14 20.13 8.04 0.92 6.79 2.08 15.32 2.85 8.44 4.17 2.47 0.89 9.42 1.69 6.47 1.95 2.45 3.44 0.90 0.21 4.97 0.58 4.76 4.71 4.39 3.43 6.91 15.36 4.96 10.43 862 1992 1302 1441 530 1348 1219 1618 17.2 23.4 12.1 29.2 21.1 3.6 41.9 19.9 0.71 1.80 1.67 2.51 2.99 1.05 3.61 2.98 10.8 20.9 12.2 11.2 16.8 12.1 53.4 51.0 0.1 2.7 0.7 2.4 2.5 3.5 2.2 0.7 23.3 10.4 5.2 1.4 156.9 2.5 0.9 1.4 11.6 50.3 24.9 21.2 17.4 40.8 14.5 45.0 1.7 10.7 14.1 19.6 3.6 0.5 12.1 2.0 Table 1.3. Trace Element and REE concentration (in ppm) of the Hillabee Greenstone (G) & Hillabee dacite (dacite). Unit Li Rb Sr Y Zr Nb Cs Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Yb Lu Hf Ta Pb Th U Sc Ti V Cr Co Ni Cu Zn Ga M68 M201b M243 G G G 6.89 12.44 9.17 0.89 3.52 0.40 106 180 153 19 23 31 64 53 62 3.28 3.33 4.45 0.12 0.05 0.03 33 26 235 3.05 4.61 8.52 7.91 10.32 14.81 1.23 1.72 3.15 6.46 8.51 14.51 2.06 2.63 4.14 0.75 0.98 1.42 1.69 2.16 3.09 0.44 0.53 0.80 3.35 4.04 5.88 0.71 0.83 1.19 2.05 2.44 3.41 1.77 2.08 2.85 0.26 0.29 0.37 0.65 0.24 0.73 0.27 0.23 0.33 1.04 0.65 1.20 0.25 0.24 0.58 0.07 0.09 0.17 39.7 35.3 46.2 5734 6864 7831 240 235 311 565 781 506 47 59 57 116 299 100 60 66 126 73 79 82 14 16 16 M268 G 4.71 0.52 134 24 73 3.52 0.06 154 4.16 10.44 1.57 8.03 2.56 0.97 2.09 0.53 4.06 0.86 2.50 2.06 0.29 0.85 0.29 0.83 0.41 0.12 41.1 6763 273 513 54 151. 74 77 14 M271 G 8.84 6.02 140 33 52 6.82 0.32 24 8.97 21.26 3.02 14.05 3.97 1.20 3.55 0.79 5.59 1.16 3.36 2.82 0.41 1.58 0.47 4.14 1.64 0.38 41.6 8985 319 103 55 55 81 105 17 M272 Sm1100 G G 6.08 4.53 0.75 0.27 106 64 15 8 58 56 0.84 0.35 0.03 0.05 23 36 1.21 0.82 3.64 2 0.61 0.31 3.48 1.71 1.29 0.66 0.54 0.35 1.01 0.55 0.31 0.17 2.54 1.42 0.55 0.31 1.68 0.92 1.50 0.82 0.23 0.13 0.44 0.26 0.11 0.09 1.02 0.38 0.13 0.09 0.04 0.03 38.6 39.2 3950 2117 240 162 481 243 41 67 119 111 58 145 63 52 14 10 28 M67 D1 D2 BG995 BG999 dacite dacite dacite dacite dacite 12.65 123 226 19 71 10.55 2.10 337 18.81 42.30 5.30 19.56 4.41 1.02 4.69 0.64 3.62 0.69 2 1.91 0.28 2.47 1.14 13.74 12.22 4.63 10.8 2085 63.2 49.4 20.2 16.7 1.9 41.6 18 15.26 158 181 7 455 5.72 3.49 530 19.43 39.76 4.18 13.43 2.12 0.53 1.02 0.17 1.37 0.28 0.84 0.73 0.14 12.81 0.60 6.87 9.04 1.94 5.0 806 26.0 40.4 4.5 16.7 1.5 16.2 3.6 6.78 127 181 14 537 8.51 2.94 478 25.98 53.64 5.76 20.04 3.53 0.59 1.48 0.28 2.51 0.50 1.50 1.25 0.23 16.45 0.73 10.04 11.78 2.67 7.2 1138 31.7 129.8 5.9 28.4 56.5 19.4 3.4 11.55 125 195 11 504 6.95 2.34 429 22.06 47.04 4.84 16.64 2.96 0.54 1.27 0.24 2.10 0.42 1.26 1.06 0.19 15.32 0.56 10.20 12.12 2.40 7.2 1002 32.5 122.1 5.7 25.8 45.4 20.7 3.3 8.95 163 184 11 650 6.45 3.13 441 24.19 48.56 5.13 17.21 3.01 0.53 1.28 0.24 2.16 0.44 1.32 1.12 0.20 20.15 0.64 7.53 13.34 3.28 6.9 958 29.3 104.7 5.7 20.9 149.4 20.9 3.4 Table 1.4. Sr and Nd Isotopic Data. Sample Units M68 G M201b G Sm Nd Rb Sr (ppm) (ppm) (ppm) (ppm) 2.06 6.46 0.89 106.12 2.63 8.51 3.52 180.43 147 Sm/144Nd 143 Nd/144 Nd 87 Rb/86Sr 87 Sr/86Sr (87Sr/86Sr)I εNd present εNd 460 0.1970 0.512939 (10) 0.0245 0.705891 (09) 0.705732 5.87 5.86 0.1911 0.512932 (08) 0.0571 0.705639 (09) 0.705270 5.74 6.07 M243 G 4.14 14.51 0.40 153.05 0.1760 0.512717 (15) 077 0.705932 (08) 0.705882 1.54 2.76 M268 G 2.56 8.03 0.52 133.95 0.1971 0.512923 (12) 0.0113 0.705118 (12) 0.705045 5.56 5.54 M271 G 3.97 14.05 6.02 139.61 0.1745 0.512674 (10) 0.1263 0.706134 (12) 0.705317 0.70 2.01 M272 G 1.29 3.48 0.75 105.74 0.2291 0.513086 (09) 0.0207 0.706485 (11) 0.706351 8.74 6.84 Sm1100 G 0.66 1.71 0.27 64.11 0.2394 0.51302 (09) 0.0125 0.703986 (09) 0.703905 7.45 4.95 M 67 HD 4.41 19.56 123.1 226.1 0.1392 0.512192 (10) 1.5937 0.718392 (08) 0.708080 -8.70 -5.55 D1 HD 2.12 13.43 158.4 180.7 0.0972 0.512103 (10) 2.5671 0.723548 (11) 0.706939 -10.44 -4.59 D2 HD 3.53 20.04 127.2 180.8 0.1088 0.512136 (12) 2.0612 0.721219 (18) 0.707883 -9.79 -4.63 BG995 HD 2.96 16.64 125.5 195.2 0.1100 0.512146 (14) 1.8829 0.720451 (07) 0.708269 -9.60 -4.51 BG999 HD 3.01 17.21 163.1 184.3 0.1080 0.512169 (13) 2.5910 0.724355 (09) 0.707591 -9.15 -3.94 Y 126 PCF 3.50 10.20 2.60 164.6 0.2118 0.51296 (11) 0.0463 0.70441 (14) 0.704106 6.28 5.33 6.64 Y 172 PCF 6.10 21.70 1.80 326.7 0.1735 0.512908 (10) 0.0161 0.70474 (12) 0.704631 5.27 TA 16 PCF 2.40 6.70 2.40 200.1 0.2211 0.512937 (16) 0.0351 0.70459 (10) 0.704360 5.83 4.40 TA 17 PCF 3.50 10.90 0.50 198.3 0.1982 0.512938 (24) 074 0.70413 (10) 0.704080 5.85 5.77 6.38 SC 254 PCF 4.20 13.40 0.80 646.6 0.1935 0.512955 (14) 036 0.70693 (09) 0.706906 6.18 SC 318A PCF 3 8.30 2.10 106 0.2231 0.512887 (12) 0.0580 0.70579 (10) 0.705419 4.86 3.31 AC 78 PCF 2.80 8.60 2.70 223.3 0.2009 0.513045 (10) 0.0354 0.70666 (12) 0.706430 7.94 7.70 BH 57 PCF 2.60 7.40 0.90 229.1 0.2169 0.513029 (17) 0.0115 0.70482 (13) 0.704746 7.63 6.45 Y 23 C GFG 5 26.40 13.90 57.68 0.1169 0.512233 (09) 0.7055 0.712751 (08) 0.708186 -7.90 -3.21 Y 24 GFG 3 12.10 39.68 88.70 0.1530 0.512510 (09) 1.3098 0.714149 (08) 0.705674 -2.50 0.07 Y 131 GFG 1.10 5.10 60.73 63.46 0.1331 0.512379 (10) 2.8025 0.722402 (07) 0.704270 -5.05 -1.32 NT 911 GFG 2.40 12.20 75.75 128.82 0.1214 0.512382 (11) 1.7220 0.71674 (09) 0.705598 -4.99 -0.57 BH 166 GFG 1.40 6 12.96 99.32 0.1440 0.512514 (10) 0.3821 0.710751 (07) 0.708279 -2.42 0.68 SC 318G GFG 1.60 4.80 0.40 24.77 0.2057 0.512750 (12) 0.0472 0.706339 (10) 0.706034 2.18 1.66 TA 7 GFG 1.30 10.10 33.25 193.69 0.0794 0.512257 (16) 0.5027 0.710996 (10) 0.707743 -7.43 -0.54 US 41A GFG 3.80 14.20 10.52 0.1652 0.512781 (09) 0.6465 0.709467 (09) 0.705284 2.79 47.67 4.65 εNd 438 MR 1 MRG 5.5 27.1 0.1253 0.512292(10) -6.75 -2.83 MR 3A MRG 5.7 20.8 0.1691 0.512449(10) -3.69 -2.18 MR 2a MRG 3.1 16.2 0.1181 0.512321(09) -6.18 -1.87 MR 2b MRG 3.7 18.7 0.1221 0.512285(12) -6.89 -2.79 MR 3b MRG 6.5 25.5 0.1573 0.512432(16) -4.02 -1.86 MR 4a MRG 6.5 23.0 0.1744 0.512462(12) -3.43 -2.21 MR 4b MRG 5.8 23.1 0.1550 0.512435(09) -3.96 -1.67 MR 4c MRG 6.2 22.6 0.1693 0.512461(13) -3.45 -1.95 Type of units: G - Hillabee Greenstone HD – Hillabee dacite PCF - Pumpkinvine Creek Amphibolites GFG - Galts Ferry Gneiss MRG – Mulberry Rock Gneiss 2σ errors on 143Nd/144 Nd and 87Sr/86Sr ratios reported as last two significant digits. 143Nd/144 Nd measured ratio normalized to 146Nd/144 Nd = 0.7219. 143Nd/144 Nd present day = 0.512638 and 147Sm/144 Nd present day = 0.1967 Initial ε values were calculated from the present day 143Nd/144NdCHUR = 0.512638 and 147Sm/144NdCHUR = 0.1967. Initial 87Sr/86Sr values were calculated for present day bulk earth value of 87Sr/86Sr = 0.7047, and 87Rb/86Sr = 0.0847. 29 Table 1.5. Rb-Sr isotopic data for whole rock samples. Sample MR 1 MR 3A MR 2a MR 2b MR 3b MR 4a MR 4b MR 4c Y 23 C Y 24 Y 131 NT 911 BH 166 SC 318G TA 7 US 41A Unit Rb (ppm) Sr (ppm) MRG 201 106 MRG 315 24.2 MRG 219 53.2 MRG 262 19.5 MRG 312 28.6 MRG 309 18.5 MRG 277 25.5 MRG 315 17 Rb/Sr 1.90 13.03 4.11 13.41 10.91 16.68 10.87 18.55 87 GFG GFG GFG GFG GFG GFG GFG GFG 0.227 0.437 0.864 0.447 0.127 0.019 0.163 0.200 0.71275 0.71415 0.72240 0.71674 0.71075 0.70634 0.71100 0.70947 14.1 47.5 59.8 86.7 13.1 0.5 32.8 10.6 62.2 108 69.2 194 103 26.8 201 53 Sr/86Sr 0.74242 0.78318 0.96096 0.95364 0.91850 1.01209 0.91588 1.04346 87 Rb/86Sr 5.13 11.76 38.49 36.40 32.12 44.58 30.33 51.95 0.7136 1.4265 2.7073 0.3500 0.0483 0.4674 0.5991 4.3788 Type of units: GFG - Galts Ferry Gneiss MRG – Mulberry Rock Gneiss 87 Sr/86Sr measured ratio normalized to 86Sr/88Sr = 0.1194. Precision of replicate whole rock determinations is 0.01% for 87Sr/86Sr. 87 Rb/86Sr ratio of MRG determined by isotope dilution and that of GFG measured by ICP-MS. Rb and Sr concentrations determined by ICP-MS. Table 1.6. Nd Model Age of metasediments from eastern Blue Ridge. Sample Lithology Sm (ppm) 3.73 3.62 8.06 2.48 7.10 5.65 5.96 4.34 7.20 8.92 3.89 15.52 13.65 17.01 7.49 7.56 Nd (ppm) 20.03 19.41 41.48 12.51 43.80 29.39 32.19 21.65 36.72 48.09 18.67 87.85 78.27 101.64 40.03 40.84 147 Sm/144Nd 143 Nd/144 Nd TCHUR TDM3.4 (Ma) (Ma) 1048 1545 1055 1552 1016 1545 1439 1892 1119 1541 984 1514 1043 1538 1023 1573 1022 1557 1032 1530 993 1580 964 1450 943 1482 1137 1568 1028 1532 1095 1580 RA401 Meta Sediments Ashland Wedowee Belt 0.11484 0.512073 (10) RA 401(rerun) Meta Sediments Ashland Wedowee Belt 0.11510 0.512071 (12) RA1005 Meta Sediments Ashland Wedowee Belt 0.11987 0.512124 (14) RA 1021 Meta Sediments Ashland Wedowee Belt 0.12246 0.511934 (10) RA 1029 Meta Sediments Ashland Wedowee Belt 0.10005 0.511926 (09) RA 422A Meta Sediments Ashland Wedowee Belt 0.11864 0.512132 (09) RA536 Meta Sediments Ashland Wedowee Belt 0.11418 0.512071 (10) RA 550 Meta Sediments Ashland Wedowee Belt 0.12368 0.512146 (14) RA 587 Meta Sediments Ashland Wedowee Belt 0.12098 0.512128 (14) RA701b Meta Sediments Ashland Wedowee Belt 0.11443 0.512079 (16) RA1034 Meta Sediments Ashland Wedowee Belt 0.12853 0.512192 (10) Sm639 Meta Sediments Ashland Wedowee Belt 0.10907 0.512082 (10) SC25 Canton Schist PCF 0.10768 0.512085 (10) Sc33 Canton Schist PCF 0.10330 0.511939 (10) NG057 Canton Schist PCF 0.11546 0.512088 (11) DR136 Canton Schist PCF 0.11422 0.512043 (12) PCF - Pumpkinvine Creek Amphibolites. 2σ errors on 143Nd/144 Nd ratios reported as last two significant digits. 143Nd/144 Nd measured ratio normalized to 146 Nd/144 Nd = 0.7219. TDM ages calculated as in Salters and Stracke (2004). TCHUR were calculated assuming the present day 143Nd/144NdCHUR = 0.512638 and 147Sm/144NdCHUR = 0.1967. 30 Table 1.7. U-Pb analysis of Galts Ferry Gneiss zircons by LA-MS-ICPMS CORRECTED CONCENTRATIONS AND RATIOS CALCULATED AGES & 1-S SD RANDOM ERRORS 207 206 206 U (ppm) Pb Pb* ± (%) Pb* ± (%) error 206Pb* ± (Ma) 207Pb* ± (Ma) 206Pb* ± (Ma) 204 Pb 219 375 64 210 226 347 170 627 475 317 447 326 285 794 954 148 313 237 173 593 341 183 404 681 337 282 270 190 610 206 285 246 419 41 53 116 173 231 35 76 84 60 127 8718 9917 5578 3833 9485 13517 13614 7768 4088 7301 22930 3943 5351 3773 25160 5310 10278 14475 14257 2779 14965 19486 9046 6720 7041 9301 11697 14622 8184 11679 2451 14898 2549 2246 6566 5384 7466 12157 1501 9176 5774 3336 6848 235 238 U 0.5732 0.5507 0.5768 0.6532 0.5747 0.5577 0.6312 0.5002 0.4935 0.5901 0.6592 0.5317 0.5632 0.5355 0.5888 0.5606 0.5405 0.5343 0.5878 0.4710 0.6273 0.5790 0.4888 0.4406 0.5473 0.5256 0.5628 0.6058 0.5307 0.5253 0.5220 0.5714 0.3914 0.7638 0.6233 0.5919 0.5970 0.5781 0.7258 0.5814 0.6223 0.7106 0.5567 U 4.64 3.21 13.74 6.70 6.77 3.93 5 5.42 6.26 2.54 2.84 6.86 3.77 1.83 3.02 7.35 5.62 2.60 6.44 9.20 4.79 4.64 2.16 2.24 5.84 5.34 6.71 4.16 3.74 4.25 11.23 3 10.69 19.03 18.43 14.38 5.73 5.70 21.97 14.04 16.85 11.25 10.71 0.0737 0.0705 0.0698 0.0820 0.0759 0.0728 0.0777 0.0652 0.0638 0.0742 0.0825 0.0723 0.0733 0.0659 0.0734 0.0765 0.0695 0.0694 0.0771 0.0612 0.0819 0.0743 0.0637 0.0581 0.0721 0.0667 0.0720 0.0775 0.0672 0.0720 0.0742 0.0727 0.0485 0.0817 0.0734 0.0744 0.0785 0.0750 0.0767 0.0746 0.0741 0.0741 0.0732 corr. 0.70 1.66 3.80 4.27 0.72 1.87 1.37 5.21 3.67 1.47 2.07 2.92 1.52 1.39 2.59 1.29 5.06 1.17 2.01 4.82 2.25 0.75 0.87 1.79 4.52 2.21 6.02 0.87 3.37 0.71 1.08 1.83 7.83 5.78 1.57 1.21 4.47 0.82 4.54 4.24 1.86 5.14 1.42 31 0.15 0.52 0.28 0.85 0.11 0.48 0.27 0.96 0.59 0.58 0.73 0.43 0.40 0.76 0.86 0.18 0.90 0.45 0.31 0.52 0.47 0.16 0.40 0.80 0.77 0.41 0.90 0.21 0.90 0.17 0.10 0.61 0.73 0.30 0.09 0.08 0.78 0.14 0.21 0.30 0.11 0.46 0.13 238 235 U 458.5 439.3 435.5 508.2 471.9 453.3 482.7 407.6 399.2 462.6 511.4 450.2 457.0 411.5 457.1 475.6 433.3 433.1 478.8 383.2 505.8 463.2 398.3 364.6 449.2 416.7 448.4 481.5 419.4 448.5 461.9 452.6 305.4 506.5 456.9 462.8 487.5 466.4 476.8 464.3 461.0 460.9 455.4 207 U 3.1 7.1 16.0 9.7 3.3 8.2 6.4 20.6 14.2 5.2 10.2 12.7 6.7 5.5 11.4 5.9 21.2 4.9 9.3 17.9 6.0 6.3 3.4 6.3 19.6 8.9 26.1 4.0 13.7 3.1 4.8 8.0 23.4 28.2 6.9 5.4 21.0 3.7 20.9 19.0 8.3 22.9 6.3 460.1 445.5 462.5 510.4 461.1 450.0 496.9 411.8 407.3 470.9 514.1 432.9 453.6 435.5 470.1 451.9 438.8 434.7 469.5 391.9 494.4 463.9 404.1 370.7 443.2 428.9 453.4 481.0 432.3 428.7 426.5 458.9 335.4 576.2 491.9 472.1 475.3 463.2 554.1 465.4 491.3 545.1 449.4 Pb* 17.2 11.6 51.1 17.1 25.1 14.3 19.7 18.3 21.0 9.6 11.5 24.2 13.8 6.5 11.4 26.8 20.0 9.2 24.2 29.9 18.8 17.3 7.2 6.9 21.0 18.7 24.6 15.9 13.2 14.9 39.1 11.1 30.5 83.9 72.0 54.3 21.8 21.2 94.1 52.5 65.7 47.5 38.9 468 477 599 521 408 433 563 436 454 517 526 342 441 564 534 333 467 443 424 443 433 472 438 409 412 495 478 478 502 324 240 491 549 862 658 517 417 448 886 471 635 915 419 102 61 287 90 151 77 105 33 113 46 43 141 77 26 34 164 54 52 137 174 94 101 44 30 83 107 66 90 36 95 258 53 159 379 397 316 80 126 449 297 363 206 238 Table 1.7 continued CORRECTED CONCENTRATIONS AND RATIOS CALCULATED AGES & 1-S SD RANDOM ERRORS 207 206 206 206 U (ppm) Pb Pb* ± (%) Pb* ± (%) error Pb* ± (Ma) 207Pb* ± (Ma) 206Pb* ± (Ma) 204 Pb 164 77 53 39 76 46 174 72 527 5656 3936 3215 1646 6954 3292 5871 6875 20054 235 238 U 0.5476 0.5948 0.5757 0.7062 0.6275 0.7175 0.5549 0.6262 0.5720 U 9.79 7.63 21.86 22.29 16.37 10.24 9.70 18.31 1.90 0.0749 0.0735 0.0751 0.0767 0.0748 0.0741 0.0676 0.0729 0.0737 corr. 2.33 3.96 3.33 2.43 4.85 1.58 1.37 4.71 0.70 0.24 0.52 0.15 0.11 0.30 0.15 0.14 0.26 0.37 238 235 U 466.0 457.4 467.1 476.6 465.2 460.9 422.2 453.9 458.9 207 U 10.5 17.5 15.0 11.1 21.8 7.0 5.6 20.6 3.1 443.4 474.0 461.7 542.5 494.5 549.2 448.2 493.8 459.3 Pb* 35.2 28.9 81.3 93.9 64.2 43.5 35.2 71.7 7.0 328 555 435 830 633 935 584 683 462 216 142 486 467 338 208 209 381 39 U concentration has an uncertainty of ~15%. Isotope ratios are corrected for Pb/U fractionation by comparison with standard zircon with an age of 564+/-4 Ma. Initial Pb composition interpreted from Stacey and Kramers (1975), with uncertainties of 1.0 for 206Pb/ 204Pb and 0.3 for 207 Pb/204Pb. 32 Table 1.8. U-Pb analysis of detrital zircons from meta-sandstone by LA-MS-ICPMS Isotopic ratios U 206 (ppm) 204 215 138 174 173 822 77 57 37 58 283 145 169 75 66 134 68 201 84 507 144 407 334 240 162 582 126 39 194 299 321 336 363 278 298 438 734 32 178 422 70 193 1072 351 Pb Pb 16371 14899 6931 17202 55166 8342 5443 4455 6783 5968 9405 18110 11913 7035 10834 7642 16751 4102 40410 18133 41030 33630 22274 19728 54839 9808 2046 6197 35883 28831 46426 13379 24956 25680 59290 52263 4568 28587 25585 3864 31361 73111 79746 207 Pb* ± (%) 235 206 Pb* ± (%) error 238 U 2.028 1.875 1.653 2.121 1.753 1.786 2.123 1.792 1.749 0.666 1.681 3.637 2.205 1.849 2.273 2.539 2.067 1.901 1.860 1.749 2.801 1.790 1.960 1.727 2.078 0.808 1.559 1.652 1.817 2.408 2.168 1.743 1.778 1.729 1.555 1.753 1.994 3.132 1.640 2.559 2.112 1.938 1.687 Apparent ages (Ma) U 2.03 4.86 4.79 3.72 0.99 12.49 7.04 12.77 6.16 8.06 4.30 2.29 4.31 7.92 2.59 5.25 3.10 9.70 1.66 4.79 1.57 1.73 1.77 4.10 1.16 9.96 18.57 5.46 2.85 2.42 2.47 2.86 3.10 2.54 5.87 1.22 16.41 3.19 5.70 9.15 4.11 1.57 5.33 0.1934 0.1752 0.1713 0.1970 0.1706 0.1683 0.1922 0.1780 0.1729 0.0819 0.1691 0.2773 0.2011 0.1707 0.2122 0.2216 0.1934 0.1761 0.1796 0.1717 0.2336 0.1772 0.1879 0.1741 0.1906 0.0955 0.1493 0.1712 0.1807 0.2120 0.1970 0.1721 0.1755 0.1717 0.1555 0.1686 0.1757 0.2471 0.1597 0.2134 0.1912 0.1797 0.1649 corr. 0.90 0.70 0.92 1.20 0.70 3.70 1.10 2.30 1.10 4.40 1.20 1.10 1.20 1.70 1.10 0.91 1.30 6.80 0.70 1.60 0.70 1.40 0.90 1.70 0.70 0.90 10.10 2.91 0.90 1 1.10 0.90 0.70 0.80 5.70 0.70 2.80 2.20 5.10 5.10 1.30 1.40 4.40 0.44 0.14 0.19 0.32 0.71 0.30 0.16 0.18 0.18 0.55 0.28 0.48 0.28 0.21 0.42 0.17 0.42 0.70 0.42 0.33 0.45 0.81 0.51 0.41 0.61 0.09 0.54 0.53 0.32 0.41 0.45 0.32 0.23 0.32 0.97 0.57 0.17 0.69 0.90 0.56 0.32 0.89 0.83 33 206 Pb* ± (Ma) 238 Pb* ± (Ma) 235 U 1140.2 1040.8 1019.4 1159.4 1015.6 1002.8 1133.6 1056.1 1028.4 507.5 1007.4 1578.2 1181.6 1016.2 1241.0 1290.6 1139.8 1045.8 1064.9 1021.6 1353.8 1051.9 1110.0 1035.1 1125.0 588.0 897.3 1019.2 1071.0 1239.8 1159.5 1024.0 1042.6 1021.8 931.9 1004.7 1043.7 1423.7 955.6 1247.1 1128.0 1065.7 984.2 207 1125.2 1072.4 990.8 1155.8 1028.4 1040.4 1156.4 1042.6 1027.1 518.7 1001.7 1557.8 1182.6 1063.4 1203.9 1283.5 1138.2 1081.5 1067.2 1026.9 1355.9 1042.0 1102.1 1018.7 1141.7 601.6 954.2 990.4 1051.6 1245.1 1171.1 1024.6 1037.7 1019.5 952.7 1028.6 1113.5 1440.6 985.7 1289.0 1152.9 1094.4 1003.8 Pb* ± (Ma) 207 U 9.4 6.7 8.7 12.7 6.6 34.4 11.5 22.4 10.5 21.5 11.2 15.4 13.0 16.0 12.4 10.6 13.6 65.7 6.9 15.1 8.5 13.6 9.2 16.3 7.2 5.1 84.6 27.4 8.9 11.3 11.7 8.5 6.7 7.6 49.5 6.5 27.0 28.1 45.3 57.8 13.5 13.8 40.2 206 Pb* 13.8 32.2 30.3 25.7 6.4 81.5 48.6 83.4 39.8 32.7 27.4 18.2 30.1 52.2 18.3 38.3 21.2 64.6 11.0 31.0 11.7 11.2 11.9 26.4 7.9 45.2 115.4 34.5 18.7 17.4 17.2 18.4 20.1 16.3 36.3 7.9 111.4 24.5 35.9 66.9 28.4 10.5 34.0 1096 1137 928 1149 1056 1120 1199 1014 1024 568 989 1530 1184 1162 1138 1272 1135 1154 1072 1038 1359 1021 1086 984 1173 653 1088 927 1012 1254 1193 1026 1027 1014 1001 1080 1252 1466 1053 1360 1200 1152 1047 36 96 97 70 14 239 137 256 123 147 84 38 82 153 47 101 56 137 30 91 27 20 30 76 18 213 314 95 55 43 44 55 61 49 29 20 318 44 51 147 77 14 61 Table 1.8 continued Isotopic ratios U 206 (ppm) 204 213 214 35 316 385 277 69 160 202 68 112 213 422 261 521 126 542 52 137 421 246 344 496 170 173 141 238 319 282 501 169 395 90 755 881 110 500 112 75 100 Pb Pb 10787 18812 5637 28898 26025 11477 6065 16561 10405 3995 10297 21588 13076 7741 5602 10457 49461 5254 3599 14335 12829 29772 74843 18221 21924 16160 29975 43306 44281 30093 16376 14769 8141 25311 53077 4675 63345 17631 7521 8509 207 Pb* ± (%) 235 206 Pb* ± (%) 238 U 26 2.239 2.203 2.972 1.703 1.719 1.805 1.608 1.539 1.680 1.692 1.666 2.266 1.994 1.571 1.511 1.764 2.942 2.413 2.242 27 2.302 1.735 1.663 2.661 2.142 2.533 1.726 2.142 2.037 1.957 1.579 1.580 2.873 1.679 1.547 1.874 2.146 2.040 1.427 Apparent ages (Ma) U 3.55 2.30 9.51 3.45 2.23 3.86 7.46 6.56 7.46 7.74 8.76 2.70 2.23 4.62 3.97 6.52 1.15 7.68 6.73 2.13 8.70 1.36 1.66 6.17 2.41 5.11 2.09 2.37 2.22 1.40 2.71 1.95 5.72 1.29 1.53 6.17 1.72 4.65 6.36 9.34 0.1894 0.2069 0.1935 0.2322 0.1691 0.1692 0.1675 0.1637 0.1543 0.1590 0.1669 0.1708 0.1903 0.1861 0.1526 0.1528 0.1702 0.2496 0.2154 0.2006 0.1814 0.2020 0.1722 0.1705 0.2235 0.2017 0.2205 0.1671 0.2015 0.1851 0.1801 0.1534 0.1583 0.2270 0.1590 0.1475 0.1786 0.1985 0.1884 0.1400 error corr. 2 1.10 3.40 3 1.80 2.30 1.30 1.80 6.20 1.51 1.90 0.90 1.80 2.51 3.70 1.90 0.70 0.71 2.34 1.70 8.30 0.70 0.70 4.10 1.70 1.90 1.20 1.40 0.70 1.20 1.90 0.70 2.50 0.80 1.30 2.70 1 2.50 1.20 5.20 0.56 0.48 0.36 0.87 0.81 0.60 0.17 0.27 0.83 0.19 0.22 0.33 0.81 0.54 0.93 0.29 0.61 0.09 0.35 0.80 0.95 0.52 0.42 0.66 0.70 0.37 0.57 0.59 0.32 0.86 0.70 0.36 0.44 0.62 0.85 0.44 0.58 0.54 0.19 0.56 206 Pb* ± (Ma) 238 Pb* ± (Ma) 235 U 1118.6 1212.3 1140.6 1346.4 1007.4 1007.9 998.5 977.8 925.3 951.4 995.1 1016.8 1123.3 1100.6 915.6 916.8 1013.6 1436.5 1257.8 1179.0 1074.8 1186.3 1024.5 1015.4 1300.8 1184.8 1284.8 996.1 1183.6 1095.3 1068.0 920.5 947.3 1318.8 951.2 887.3 1059.4 1167.7 1112.8 845.1 207 1117.7 1193.5 1182.0 1400.6 1010.0 1015.8 1047.3 973.5 946.4 1001.1 1005.6 995.7 1202.0 1113.5 958.9 935.1 1032.6 1392.8 1246.7 1194.3 1118.0 1213.0 1021.9 994.7 1317.8 1162.6 1281.6 1018.4 1162.6 1128.0 1100.9 962.3 962.6 1375.1 1000.8 949.5 1072.2 1163.9 1129.3 900.3 Pb* ± (Ma) 207 U 20.6 12.2 35.6 36.5 16.8 21.5 12.0 16.3 53.4 13.3 17.5 8.5 18.6 25.4 31.6 16.2 6.6 9.1 26.7 18.3 82.2 7.6 6.6 38.5 20.0 20.6 14.0 12.9 7.6 12.1 18.7 6.0 22.0 9.5 11.5 22.4 9.8 26.7 12.3 41.2 206 Pb* 24.1 16.2 66.5 26.2 14.3 24.8 48.8 41.1 46.0 49.3 55.9 17.1 15.7 31.3 24.6 39.9 7.4 58.3 48.3 14.9 59.1 9.6 10.7 39.2 17.8 35.4 15.2 15.3 15.4 9.6 18.2 12.1 35.6 9.7 9.7 38.1 11.4 32.2 43.4 55.8 1116 1160 1259 1484 1015 1033 1150 964 996 1111 1028 949 1346 1139 1059 978 1073 1326 1228 1222 1203 1261 1016 949 1346 1121 1276 1067 1124 1192 1166 1059 998 1464 1111 1097 1098 1157 1161 1038 59 40 174 32 27 63 146 129 84 152 173 52 26 77 29 127 18 148 124 25 52 23 30 95 33 95 33 39 42 14 38 37 105 19 16 111 28 78 124 157 U concentration has an uncertainty of ~15%. Isotope ratios are corrected for Pb/U fractionation by comparison with standard zircon with an age of 564+/-4 Ma. Initial Pb composition interpreted from Stacey and Kramers (1975), with uncertainties of 1.0 for 206Pb/ 204Pb and 0.3 for 207 Pb/204Pb. 34 Table 1.9. U-Pb analysis of Mulberry Rock Gneiss zircons by LA-MS-ICPMS CORRECTED CONCENTRATIONS AND RATIOS U (ppm) 206 Pb 204 Pb 245 283 444 149 661 515 309 558 218 1625 920 375 296 794 794 823 1776 1024 1409 375 793 316 287 842 379 459 962 264 1005 147 648 273 248 285 390 672 316 980 272 805 262 1204 224 464 1079 574 836 1486 4205 2680 4416 1194 3255 1692 1394 6381 3075 3898 8917 2248 1637 6017 2804 23984 1844 2165 3143 1183 5679 1694 1970 6576 2027 1735 1609 2484 1176 1056 3007 4770 6596 1087 2381 2701 2116 2674 3246 1813 3850 654 2528 1028 207 Pb* ± (%) 235 Pb* U 6.56 10.54 6.60 8.30 4.06 13.32 7.08 8.06 8.43 1.81 4.26 4.58 6.32 5.75 7.08 4.37 2.17 1.43 7.57 5.29 4.72 11 3.89 5.52 8.38 3.81 4.69 7.21 3.26 10.96 6.03 10.85 8.38 8.27 4.95 4.60 8.25 4.40 7.35 7.47 5.33 6.45 8.46 15.99 5.89 6.24 0.0757 0.0748 0.0707 0.0720 0.0668 0.0491 0.0678 0.0658 0.0689 0.0709 0.0666 0.0710 0.0698 0.0674 0.0675 0.0684 0.0578 0.1707 0.0675 0.0689 0.0618 0.0503 0.0713 0.0672 0.0673 0.0683 0.0630 0.0729 0.0703 0.0727 0.0625 0.0762 0.0707 0.0686 0.0684 0.0655 0.0717 0.0664 0.0619 0.0680 0.0770 0.0685 0.0764 0.0613 0.0718 0.0498 CALCULATED AGES & 1-S SD RANDOM ERRORS ± (%) error 238 U 0.6710 0.5953 0.5569 0.5949 0.5370 0.4206 0.5412 0.5439 0.5681 0.5627 0.5398 0.5712 0.5501 0.5507 0.5560 0.5322 0.4522 1.7656 0.5576 0.5283 0.4868 0.4192 0.5384 0.5533 0.5616 0.5170 0.5101 0.6279 0.5761 0.6074 0.5045 0.6552 0.6108 0.5001 0.5360 0.5379 0.5829 0.5389 0.4984 0.5737 0.6107 0.5612 0.5598 0.5680 0.5916 0.3923 206 corr. 4.16 1.69 1 2.74 1.86 8.38 1.39 3.31 2.11 1.34 3.29 1 1.14 1.24 3.25 1.77 1.67 1 2.01 2.46 2.68 6.12 1.01 1.22 1.94 1.86 2.96 2.79 2.06 1.82 2.94 1.49 2.83 2.14 1.28 1.93 2.53 2.90 2.27 2.91 1.82 2.15 3.92 2.19 2.26 2.56 0.64 0.16 0.15 0.33 0.46 0.63 0.20 0.41 0.25 0.74 0.77 0.22 0.18 0.22 0.46 0.41 0.77 0.70 0.26 0.47 0.57 0.56 0.26 0.22 0.23 0.49 0.63 0.39 0.63 0.17 0.49 0.14 0.34 0.26 0.26 0.42 0.31 0.66 0.31 0.39 0.34 0.33 0.46 0.14 0.38 0.41 35 206 Pb* ± (Ma) 238 Pb* ± (Ma) 235 U 470.8 465.3 440.6 448.4 417.3 309.3 423.3 410.9 429.9 441.9 415.7 442.5 435.3 420.7 421.5 426.6 362.3 1016.2 421.1 430.0 386.9 316.8 444.3 419.3 419.9 426.3 394.0 453.6 438.4 452.8 391.2 473.4 440.6 427.7 427.0 409.1 446.6 414.6 387.7 424.1 478.6 427.3 474.7 383.7 447.3 313.3 207 521.4 474.3 449.5 474.0 436.5 356.5 439.3 441.0 456.8 453.3 438.3 458.8 445.1 445.5 449.0 433.3 378.9 1032.9 450.0 430.7 402.7 355.5 437.4 447.2 452.6 423.2 418.5 494.8 462.0 481.9 414.8 511.7 484.1 411.8 435.8 437.0 466.3 437.7 410.6 460.4 484.0 452.3 451.4 456.8 471.9 336.1 Pb* ± (Ma) 207 U 18.9 7.6 4.3 11.9 7.5 25.3 5.7 13.2 8.8 5.7 13.2 4.3 4.8 5.0 13.3 7.3 5.9 9.4 8.2 10.3 10.1 18.9 4.3 4.9 7.9 7.7 11.3 12.2 8.7 8.0 11.1 6.8 12.1 8.9 5.3 7.7 10.9 11.7 8.6 12.0 8.4 8.9 18.0 8.2 9.8 7.8 206 Pb* 26.7 39.9 24.0 31.4 14.4 40.1 25.3 28.9 31.0 6.6 15.2 16.9 22.8 20.7 25.7 15.4 6.9 9.2 27.5 18.6 15.7 33.0 13.8 20.0 30.6 13.2 16.1 28.2 12.1 42.1 20.5 43.6 32.3 28.0 17.5 16.3 30.9 15.6 24.8 27.7 20.5 23.6 30.8 58.9 22.2 17.9 749 518 496 600 539 676 524 601 595 512 559 541 496 576 592 469 482 1068 601 435 494 617 401 593 622 406 556 690 581 623 548 687 696 323 483 587 565 561 542 646 510 582 335 843 593 496 107 229 144 170 79 222 152 159 177 27 59 98 137 122 137 89 31 20 158 104 86 198 84 117 176 74 79 142 55 234 115 230 168 182 106 91 171 72 153 148 110 132 170 331 118 126 Table 1.9 continued U (ppm) CORRECTED CONCENTRATIONS AND RATIOS 207 206 206 Pb Pb* ± (%) Pb* ± (%) 204 235 Pb 1151 311 269 165 269 566 298 169 705 299 176 664 158 236 1904 1362 550 296 364 187 121 238 U 1131 168 430 963 2564 875 610 576 620 1551 5945 6502 5733 2538 3008 1343 1414 6614 755 10115 2230 U 0.4870 0.6287 0.5408 0.5382 0.4156 0.5200 0.5449 0.9735 0.6253 0.5337 0.5503 0.4763 0.6770 0.6055 0.5652 0.6211 1.2266 0.5906 0.7355 0.6460 0.5724 3.41 9.17 9.92 13.39 15.17 6.95 12.19 10.42 7.20 7.29 9.50 9.13 11.09 10.17 3.97 2.96 4.26 5.15 7.47 7.66 11.61 0.0589 0.0532 0.0545 0.0653 0.0498 0.0633 0.0606 0.0954 0.0724 0.0641 0.0711 0.0645 0.0878 0.0759 0.0698 0.0750 0.1089 0.0750 0.0724 0.0784 0.0770 error CALCULATED AGES & 1-S SD RANDOM ERRORS 207 206 Pb* ± (Ma) Pb* ± (Ma) Pb* ± (Ma) 206 238 corr. 2.47 1.89 3.25 5.34 6.84 4.59 4.75 5.48 4.07 2.81 2.06 7.63 1.12 2.64 2.44 1.21 3.01 0.95 3.65 4.19 0.93 0.73 0.21 0.33 0.40 0.45 0.66 0.39 0.53 0.57 0.39 0.22 0.84 0.10 0.26 0.61 0.41 0.71 0.19 0.49 0.55 0.08 235 U 207 U 368.9 334.6 342.3 408.1 313.3 396.2 379.3 587.4 450.6 400.8 443.3 403.0 542.7 471.7 435.3 466.6 666.4 466.6 451.1 486.8 478.2 8.9 6.2 10.9 21.1 20.9 17.6 17.5 30.8 17.7 10.9 8.8 29.8 5.8 12.0 10.3 5.4 19.1 4.3 15.9 19.6 4.3 402.9 495.3 439.0 437.2 352.9 425.2 441.7 690.3 493.2 434.3 445.2 395.5 525.0 480.7 454.9 490.6 812.8 471.3 559.8 506.0 459.6 Pb* 11.3 36.0 35.4 47.6 45.3 24.1 43.7 52.2 28.1 25.8 34.2 29.9 45.5 39.0 14.6 11.5 23.9 19.4 32.2 30.5 42.9 603 582 984 594 622 585 781 720 696 616 455 352 449 524 555 604 51 97 191 267 293 113 237 192 127 145 206 113 246 216 68 59 494 620 594 368 112 134 139 261 U concentration has an uncertainty of ~15%. Isotope ratios are corrected for Pb/U fractionation by comparison with standard zircon with an age of 564+/-4 Ma. Initial Pb composition interpreted from Stacey and Kramers (1975), with uncertainties of 1.0 for 206Pb/ 204Pb and 0.3 for 207 Pb/204Pb. Table 1.10 REE concentrations of the metasediments of the eastern Blue Ridge. RA401 RA 401 RA701b RA1034 Sc33 NG057 DR136 RA 422A RA536 RA 550 RA 587 RA1005 RA 1029 RA 1021 Sm639 SC25 La 36 31 73 30 203 38 56 27 33 20 35 65 73 18 120 144 Ce 55 36 180 84 327 93 127 35 71 45 74 88 95 58 208 132 Pr 5.4 5.1 13.5 5.0 30.3 10.8 10.5 7.6 8.5 5.6 9.5 11.0 12.0 3.4 25 23 Nd 20 19 48 19 102 40 41 29 32 22 37 41 44 13 88 78 Sm 3.7 3.6 8.9 3.9 17.0 7.5 7.6 5.6 6.0 4.3 7.2 8.1 7.9 2.5 16 14 Eu 0.69 0.76 0.88 0.64 1.65 1.28 1.11 1.18 1.02 1.01 1.94 1.66 1.39 0.51 1.81 1.65 Gd 1.20 2.08 3.14 1.46 5.49 2.64 2.65 3.03 3.25 2.59 4.09 4.93 4.59 1.72 4.43 3.16 Tb 0.26 0.35 0.60 0.34 1.25 0.50 0.51 0.55 0.57 0.55 0.75 1.00 0.85 0.33 1.06 1.00 Dy 1.90 1.85 6.25 5.38 9.89 4.14 3.73 3.11 3.29 3.99 4.39 6.47 5.60 2.47 10.87 8.57 Ho 0.32 0.30 1.29 1.16 1.92 0.82 0.74 0.49 0.58 0.83 0.72 1.21 1.11 0.47 2.10 1.55 Er 0.81 0.77 4.10 3.55 6.16 2.49 2.27 1.16 1.75 2.54 1.71 3.63 3.28 1.34 6.18 4.43 Yb 0.69 0.72 3.20 2.49 6.37 2.25 2.04 0.94 2.34 2.80 1.23 4.64 3.31 1.37 4.26 3.57 Lu 0.12 0.13 0.58 0.46 1.23 0.40 0.36 0.15 0.43 0.47 0.18 0.82 0.54 0.23 0.68 0.61 36 Study area (Fig 1.1 A) Figure.1.1. Generalized geological map of the Southern Appalachian. CBR – Central Blue Ridge; DGB – Dahlonega Gold Belt. Modified after Holm & Das, 2006. 37 Figure 1.1A. Generalized geologic map of eastern Alabama and western Georia Blue Ridge terranes showing Fig1.1A the tectonic relationship between the PCG and HG. Modified after Holm and Das, 2006 (in review). WBRwestern Blue Ridge; PCF-Pumpkinvine Creek Formation; MRG-Mulberry Rock Gneiss; AG-Austell Gneiss; HG-Hillabee Greenstone; EQD-Elkahatchee Quartz Diorite. 38 Figure 1.2. Geochemical representation of bimodality found in the HG and Hillabee dacite and PCF amphibolites and felsic gneiss (GFG). (A) SiO2 weight % histogram of the felsic and mafic components of the HG belt. (B) SiO2 weight % histogram of the felsic and mafic components of the PCF. Data from this study, Durham, 1993 and McConnell and Abrams, 1984. 39 Figure 1.3. Total Alkalis versus SiO2 (Lebas et al. 1986) place the mafic component of the PCF (purple filled diamond) and HG (blue filled square) in the basalt field while the GFG (purple open diamond) falls in the rhyolite field and the Hillabee dacite (blue open square) falls in the dacite field. Data from Durham, 1993 and McConnell and Abrams, 1984 are plotted too. Literature data symbols: GFG (pink open plus sign), Hillabee dacite (blue open multiply sign), PVC amphibolite (pink filled plus sign), greenstones (blue open multiply sign). 40 Figure 1.4A. Classification of the Hillabee dacite and GFG using Shand’s index of Maniar and Piccoli (1989). Both the felsic components display a peraluminous affinity. Symbols and data source same as Figure 1.3. Figure 1.4B. AFM diagram after Irvine and Barager (1971) showing tholeiitic trend of HG and PCF amphibolites and calc-alkaline trend of Hillabee dacite and GFG. Symbols same as Figure 1.3. FeO* is the total FeO. 41 A Ti (ppm) 100000 10000 Expected compositio n of felsic component if derived from fractional crystallizat ion of mafic 1000 100 10 100 1000 Zr (ppm) 100000 Ti (ppm) B 10000 Expected composition of felsic component if derived from fractional crystallizatio n of mafic material. 1000 100 10 100 1000 Zr (ppm) Figure.1.5. Ti vs. Zr plot to illustrate trends of A) HG and Hillabee dacite B)PCF amphibolites and GFG and expected trend of dacite and GFG if they were derived by fractional crystallization of their mafic counterparts. Symbols same as Figure. 1.3. 42 Ti (ppm) 100000 10000 1000 10 60 110 160 210 Zr (ppm) 10 Nb (ppm) 9 8 7 6 5 4 3 2 1 0 10 60 110 160 210 160 210 Zr (ppm) 60 Y (ppm) 50 40 30 20 10 0 10 60 110 Zr (ppm) Figure.1.6. Co-variation diagrams of relatively immobile elements for the HG (blue filled squares) and PCF amphibolites (pink filled diamonds). Ti, Nb and Y vs. the HFSE Zr. 43 A B Figure 1.7. Tectonic discrimination diagrams for the PCF amphibolites and HG. (A) Zr/4-Nb*2-Y diagram (Meschede, 1986). (B) Y vs. Nb diagram (Pearce et al., 1984). OIB- ocean island basalt, WP Alk- withinplate alkaline; WP Th- within-plate tholeiite; E-MORB- enriched midocean ridge basalt; N-MORB- normal mid-ocean ridge basalt; VABvolcanic arc basalt. Symbols same as Figure 1.6. 44 Figure1.8. La/10-Y/15 Nb/8 diagram (Cabanis and Lecolle, 1989). IAT- island-arc tholeiite, OFB- ocean-floor basalt, MORB- mid-ocean ridge basalt, BABB- back-arc basin basalt, VAT- volcanic-arc tholeiite, NMORB- normal mid-ocean ridge basalt, EMORB- enriched mid-ocean ridge basalt, Cont. Continental. Symbols same as Figure 1.6. 45 Figure 1.9. Granitoid tectonic discrimination diagrams of GFG (pink open diamonds) and Hillabee dacite (blue open squares). (A) Y vs. Nb diagram (Pearce et al., 1984). (B) Y + Nb vs. Rb diagram (Pearce et al., 1984). VAGvolcanic arc granite, syn-COLG- syn-collisional granite, WPG- within plate granite, ORG- ocean-ridge granite. 46 Rock/Primitive Mantle McDonough and Sun 100.00 PCF 10.00 1.00 0.10 0.01 Rb Ba Th U Nb La Ce Pb Pr Sr Nd Zr Sm Eu Ti Dy Yb Y Lu 100.00 Hillabee Greenstone 10.00 1.00 0.10 0.01 Rb Ba Th U Nb La Ce Pb Pr Sr Nd Zr Sm Eu Ti Dy Yb Y Lu Figure. 1.10. Primitive Mantle- normalized (McDonough and Sun, 1995) spider diagram plot of PCF amphibolites and Hillabee Greenstone. 47 Rock/Chondrites McDonough and Sun 1995 100 10 PCF Amphibolites 1 La Ce Pr Nd Sm Eu Tb Dy Ho Er Yb Lu 100 Hillabee Greenstone 10 1 La Ce Pr Nd Sm Eu Tb Dy Ho Er Yb Lu Figure. 1.11. CI chondrite-normalized (McDonough and Sun, 1995) REE diagram plotting of PCF amphibolites and Hillabee Greenstone. 48 Rock/Chondrites McDonough and Sun 1995 1000 GFG 100 10 1 La Ce Pr Nd Sm Eu Tb Dy Ho Er Yb Lu Yb Lu 1000 Hillabee dacite 100 10 1 La Ce Pr Nd Sm Eu Tb Dy Ho Er Figure. 1.12. CI chondrite-normalized (McDonough and Sun, 1995) REE diagram of GFG and Hillabee dacites. Field for metadacite from Durham, 1993 is shown for comparison. 49 9.00 7.00 HG 5.00 3.00 εNd Hillabee Metadacite 1.00 GFG -1.00 -3.00 PCF Amphibolite -5.00 -7.00 0.703 0.704 0.705 0.706 0.707 0.708 0.709 (87Sr/86Sr)i Figure.1.13. Initial Sr vs. εNd isotopic plot of Pumpkinvine Creek Formation and Hillabee Greenstone Sequence Rocks. 50 0.724 Y 131 391 Ma isochron 0.720 86 Sr/ Sr NT 911 0.716 87 Y 24 Y 23 C 0.712 TA 7 BH 166 US 41A Age = 391±110 Ma 0.708 Initial 87 86 Sr/ Sr =0.7076±0.0020 SC 318G MSWD = 60 0.704 0.0 0.4 0.8 1.2 1.6 87 2.0 2.4 2.8 86 Rb/ Sr 0.720 NT 911 Figure. 1.14. Rb-Sr isochron of the GFG Y 23 C 0.712 BH 166 87 86 Sr/ Sr 0.716 TA 7 0.708 Age = 450±66 Ma Initial 87 86 Sr/ Sr =0.7082±0.0012 MSWD = 2.2 0.704 0.0 0.2 0.4 0.6 0.8 87 1.0 1.2 1.4 86 Rb/ Sr 0.724 Y 131 0.716 86 Sr/ Sr 0.720 87 Y 24 0.712 US 41A Age = 424±25 Ma 0.708 Initial 87 SC 318G 86 Sr/ Sr =0.7059±0.0010 MSWD = 0.8 0.704 0.0 0.4 0.8 1.2 1.6 87 86 Rb/ Sr 2.0 2.4 51 2.8 GFG, normalized to CI chondrite (McDonough & Sun 95) 250 High initial 87Sr/86Sr sample/standard 200 150 100 50 Low initial 87Sr/86Sr 0 La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Yb Lu Figure. 1.15. Chondrite normalized REE plot of GFG samples. The samples define two groups. Light gray with high initial 87Sr/86Sr (~0.708) and dark gray with low initial 87Sr/86Sr (~0.705). 52 data-point error ellipses are 2σ 0.11 650 550 450 0.07 206 Pb/238U 0.09 350 0.05 Concordia Diagram Galts Ferry Gneiss Samples 250 0.03 0.2 0.4 0.6 207 Pb/ 0.8 1.0 235 U 25 Number 20 15 10 5 0 260 300 340 380 420 460 500 540 206Pb/238U age in Ma 510 box heights are 1σ 206Pb/238U age 490 470 450 430 Mean = 460.4±9.7 95% conf. MSWD = 1.7 410 53 Figure. 1.16. U-Pb ages of zircons from GFG plotted in Concordia diagram, histogram and weighted average plot. Bin width for the histogram is 20 Ma. 206Pb/238U ages of zircons falling in the modal class and second highest frequency class after the modal class of the histogram are selected for the weighted average plot. 462+/-5 Ma 457+/-7 Ma 200 micron 463+/-9 Ma 505+/-6 Ma 200 micron Figure. 16.1A. Galts Ferry Gneiss zircon showing magmatic zoning (top) and older core (bottom). 54 0.5135 0.5130 0 143Nd/144Nd 0.5125 500 1000 1500 2000 CHUR 0.5120 143 Nd/144Nd Depleted Mantle extracted at 3.4 Ga. 0.5115 0.5110 0.5105 Grenville basement at 700 Ma. 0.5100 Age in Ma Figure. 1.17. Nd isotopic evolution of the eastern Blue Ridge. CHUR (Hamilton et al., 1983) evolution curve in bold red and Depleted Mantle (Salters and Stracke, 2004) evolution curve in bold blue. The two green evolution curves are for the two outliers with the least TDM model age and the red evolution curve is for the outlier with the oldest TDM model age. Nd isotopic value for the Grenville crust at 700 Ma is shown by the gray shaded box (Carrigan et al., 2003; Hatcher et al., 2004). 55 NASC normalized REE plot of the Metasediments 5.0 4.5 4.0 3.5 3.0 2.5 2.0 1.5 1.0 0.5 0.0 La Pr Nd Sm Eu Tb Dy Ho Er Yb Lu Figure.1.18. North American Shale normalized (Gromet, 1984) REE plot of the metasediments from Ashland Wedowee belt and Canton Schist. Two samples with youngest TDM is shown in green and the one oldest sample is shown in red. 56 0.34 data-point error ellipses are 2σ 0.30 0.26 1400 206Pb/238U 0.22 0.18 1000 0.14 600 0.10 Concordia Diagram Meta-sandstone from PCF 0.06 200 0.02 0 1 2 3 4 207Pb/235U 18 16 14 Number 12 10 8 6 4 2 0 300 500 700 900 1100 1300 1500 1700 207Pb-206Pb age in Ma Figure. 1.19. Concordia Plot and Histogram (bin size = 50 Ma) of the zircons from the meta-sandstone within the Pumpkinvine Creek Formation. 57 1.15 0.95 87 Sr/ 86Sr 1.05 0.85 Age = 467±16 Ma 0.75 Initial 87 86 Sr/ Sr =0.7076±0.0036 MSWD = 7 0.65 0 20 40 87 Rb/86Sr Figure.1.20. Mulberry Rock Gneiss whole rock Rb-Sr 58 60 0.20 1100 data-point error ellipses are 2σ 0.16 206Pb/238U 900 0.12 700 500 0.08 300 0.04 Concordia Diagram Mulberry Rock Gneiss 100 0.00 0.0 0.4 0.8 1.2 1.6 2.0 207Pb/235U 16 14 35 data in this cluster 12 Number 10 8 Lead Loss 6 Inhereted (?) or mixing 4 2 0 200 250 300 350 400 450 500 550 600 Ma Figure 1.21. U-Pb ages of zircons from Mulberry Rock Gneiss plotted in Concordia diagram, and histogram. 59 Hillabee dacite Figure. 1.22. Epsilon Nd of Ordovician arc terrane felsic magmas of the Appalachians. Crust fields and included arc terrane values are modified from Coler et al. (2001) and references therein. The new data from the GFG and Hillabee dacite are compared with the other more northern Appalachian arc terranes and associated units 60 Figure. 1.23 Passive mechanisms of lithosphere extension. From a synthesis and simplification of numerous studies that followed Celar Sengor and Burke, 1978. Modified after Geoffroy, 2005. Figure. 1.24. Accretionary Orogen in southern Appalachians (ca. Ord) 61 CHAPTER TWO KILOMETERS SCALE STRONTIUM ISOTOPIC HOMOGENIZATION DURING METAMORPHISM: A CASE STUDY IN THE TRES PEIDRAS GRANITE, NEW MEXICO 2.1 Introduction More than a half century ago Compston and Jeffery (1959) made a major advance in the budding science of geochronology when they showed that the common discordance between whole-rock and mineral Rb-Sr dates of "granite-like" bodies were artifacts of the use of an initial strontium isotopic composition in the calculation of dates. They revealed that, on the scale of a whole-rock sample, a redistribution of Sr between minerals can occur during metamorphism, so that the whole rock and minerals become homogenized to a common 87Sr/86Sr ratio. Furthermore it was indicated that minerals and whole-rock yielded concordant dates for the time of metamorphism, when this rehomogenized ratio was used as the common strontium in age calculations. Shortly after that Nicolaysen (1961) of the Bernard Price Institute introduced a graphical solution to such systems. Making use of the BPI plot (now referred to as isochron diagram), Nicolaysen demonstrated that different granite-gneiss whole-rocks of the Baltimore Dome had remained closed systems during metamorphism while there had been redistribution of Sr isotopes between the minerals of individual rocks. Multiple whole rock samples yielded isochron ages that provided the time of initial crystallization of the granite, while individual whole rocks and their respective minerals yielded isochrons that date the time of metamorphic rehomogenization of strontium. The scale of this strontium isotopic exchange was that of a whole-rock (decimeters). There are numerous cases in the literature where these early advances in understanding have found application. There are also numerous reports in which whole rocks appear to have been open Rb-Sr systems, and analyses fail to produce the isochrons that should result as a necessary consequence of close systems. Then there are reports in which well defined whole-rock isochrons yield dates less than those obtained from U-Pb dates of zircons. Proposed reasons for such discrepancy between well defined Rb-Sr whole rock isochron ages and zircon ages from the same rocks have been (1) the presence of older zircon xenocrysts (Odom and Fullager, 1984), (2) a systematic loss of radiogenic 87Sr or gain of 87Rb (Bickford and Mose 1975), (3) development of an apparent isochron by mixing (Field and Raheim, 1980), and (4) an undefined "resetting" of an isochron age by metamorphism (Page, 1978). It is inferred that such a resetting would involve an isotopic rehomogenization of strontium on a scale equal to or greater than the distance between samples. At present such large (kilometer) scale rehomogenization yet has been demonstrated. Now geochronological investigations of the Tres Piedras Granite of Northcental New Mexico seem to have revealed a remarkable example. Tres Piedras Granite located near the town of Tres Piedras of the Rio Arriba County is one of the several granites intruded during the growth of Proterozoic continental lithosphere in the southwestern United States between 1.74 and 1.65 Ga (Maxon, 1976, Karlstrom and Bowring, 1988; Bauer and Williams, 1994). Two well documented thermal event is recorded in the area, at ca. 1.6 Ga (Grambling and Dallmeyer, 1993) after the initial continental building period and at 1.48 Ga (Lanzirotti and Hanson, 1997) that affected the plutons in the area. 62 2.2 General Geology Extensive Proterozoic continental lithosphere was formed from Wyoming southwestward till New Mexico during Middle Proterozoic (~1800 – 1600 Ma) that record two major orogenic cycles; the Yavapai Orogeny (~1700 Ma) and the Mazatzal Orogeny (~1650 Ma) (Hoffman, 1988; Karlstrom and Bowring, 1988). Subsequently the rocks underwent regional and contact metamorphism and intrusion of “anorogenic” granitoids at 1400 Ma (Karlstrom et al., 1997; Pedrick et al., 1998; Read et al., 1999; Williams et al., 1999). Controversies exist regarding metamorphic and deformation history of north central New Mexico. While earlier studies (Grambling,1986; Bowring and Karlstrom, 1990; and Nyman et al., 1994) propose one metamorphic event either at 1650 Ma or 1400 Ma, more recent works have identified two deformation events at both 1650 Ma or 1400 Ma (Pedrick et al., 1998; Read et al., 1999; Williams et al., 1999) In many Precambrian – cored uplifts of northern New Mexico and southern Colorado, the ages of major fold and thrust structures, regional foliation and associated metamorphism are loosely constrained between the ~1.70-1.65 Ga age of deformed plutons and the ~ 1.4Ga age of cross-cutting plutons, dykes and cooling of metamorphic minerals (Reed et al., 1987; Gibson and Simpson, 1988; Williams, 1991). Assuming both the events were responsible for the deformation and metamorphism of the rocks, it is still debated which event is responsible for which deformational fabrics. Proterozoic supracrustal rocks are exposed in isolated fault bounded uplifts across New Mexico. Lithologies are dominated by metavolcanic rocks, metaquartzite and pelitic schist. Primary ages range from 1750 to 1650 Ma, generally becoming younger to the south (Bowring and Condie, 1982). The supracrustal rocks were intruded by several granitic suites as well as several younger batholiths prior to metamorphism. The oldest rocks of the Taos, Tusas and Santa Fe Ranges are primarily gneisses and amphibolites. Here the Moppin, Pecos, and Gold Hill complexes have been dated between 1760 to 1720 Ma (Bowring and Condie, 1982, Robertson and Condie, 1989, Bowring, 1986, Silver and Dickinson 1987). Stratigraphically above these are schists, metaconglomerates and amphibolites of the Vadito Group, followed by the Hondo Group. The Hondo group has two formations, lower kilometer thick Ortega Formation and the upper Rinconada Formation mainly composed of schists and phyllites. The Hondo Group is generally interpreted to be a cratonic margin sedimentary sequence that can be correlated from range to range in northern New Mexico (Bauer and Williams, 1989). This stratigraphy is cut by ~1.4 Ga and ~1.7 Ga. granites. (Table 2.1,2.2) (Table 2.1,2.2) 63 2.3 Tres Piedras Granite 2.3.1 Occurrence The Tres Piedras Granite is exposed around the town of Tres Piedras of eastern Rio Arriba County (Fig. 2.1, 2.1a). It is a pink, flesh-colored moderately to well foliated medium grained granite. Crude foliation is defined by mica rich layers. The rock -foliation strikes northwest – southeast. The rocks in this province consist of a thick section of interbedded metasedimentary and metavolcanic units. The major metamorphic rock types exposed at the surface are granite – gneiss, quartzite, hornblende – chlorite schist and phyllite. Tres Piedras Granite is a group of granitic to granodioritic bodies that intrude the metasedimentary units. The granites and the metasediments have been metamorphosed to upper amphibolite facies and deformed during later metamorphic events. The exact shape and size of the granitic – gneiss body is unknown due to subsequent deformation with Tertiary and Quaternary sediments and volcanics covering all but isolated outcrops. From the surface expression of the outcrops it can be seen that the granite unit is elongated in shape and has its longer axis subparallel to the regional foliation tend of the crystalline province. Large outcrops of the Tres Piedras Granite near the town of Tres Piedras and along the Tusas River Canyon to the west of Tres Piedras (Fig. 2.1, 2.1a). The Tres Piedras Granite near the town of Tres Piedras consists in places of granitic inliers whose base is hidden beneath thick basalt flows related to the Rio Grand Rifting. 2.3.2 Petrography & Chemistry The rock is granite, ranging from very close to alkali feldspar granite to true granite. The composition (eye estimation by volume) can be expressed as about 40% potash feldspar, 15% plagioclase, 20% quartz; the rest 25% is made up of biotite, hornblende, muscovite, epidote, sphene, chlorite, and rarely calcite and apatite. Some portions show weak to moderate recrystallization texture attributed to low-grade metamorphism. The overall texture is that of typically igneous allotriomorphic granular with most grains of felsic minerals being anhedral. Only the biotites show a preferred orientation that define the crude foliation. The alkali feldspar rich portions are dominated by large grains of flame and string perthite. In a few perthite grains the exsolved lamellae of albite are oriented in two mutually perpendicular directions. Some perthite and also some homogeneous grains of potash feldspar show thin sodic rims that in many cases merge with the albite lamellae in perthitic grains. Graphic intergrowth of quartz are present in some smaller grains of potash feldspar. Some portions of the rocks show strong alteration of potash feldspar to sericite in crystallographically controlled directions. Some grains show two such orientations. Microcline is abundant in portions, showing cross-hatched twinning. Some microcline grains also show perthitic intergrowths. Potash feldspar grains are the largest among all the minerals. Plagioclase grains are more regular in shape compared to alkali feldspar and quartz grains, and even after strong alteration to epidote and muscovite show lath shaped outline. Lamellar twinning is characteristically present. 64 Biotite (very dark brown, and greenish brown varieties) appears to be of two different generations. Dark brown varieties are of weakly tabular in shape and are cross cut by clean, large, perfectly tabular muscovite grains implying the metamorphic nature of muscovite. These biotites could be pre-metamorphic and igneous in origin. The greenish variety of biotite form composite grains with chlorite, and green hornblende with bluish tinge has grown over the biotite. This assemblage of chlorite – biotite – hornblende is interpreted to be metamorphic. However, most of the grains of hornblende are irregular in shape, and some even do not show any good cleavage. Only a few hornblende grains are subhedral with very good cleavage. These hornblende grains show brownish pleochroism and could be of igneous origin. Some very large muscovite grains show sieve structure defined by large sub-circular inclusions of quartz. These seem to be older than the prismatic smaller, and inclusion free grains of muscovite. Thus the larger grains could be of igneous origin. Quartz grains are all anhedral. Larger grains show a tendency to form aggregates of smaller grains – but no preferred orientations are noticeable. Lozenge-shaped sphene grains, and a few prismatic epidote grains show metamict texture defined by dark rims around the grains. These are common around inclusions inside hornblende. History: Alkali feldspar dominated granite with small amount of biotite, hornblende and muscovite was metamorphosed to lower amphibolite facies. A chemical and normative analysis of the Tres Piedras Granite by Barker (1958) is included in Table 2.3. A comparison of the modal analysis on the same rock by Barker (1958), Bingler (1965) and this study is included in the table too. The Tres Piedras Granite has very low content of calcium and is weakly peralkaline (presence of normative corrundum). 2.3.3 Contact Relation Several intrusive contacts between the Tres Piedras Granite and the surrounding rocks can be seen in the field. Xenoliths of the country rocks are common within the Tres Piedras Granite in some localities. Along the Tusas River Canyon, a large elongated block of muscovitic quartzite approximately 100 meters in thickness is enclosed by granitic-gneiss. Northwest of Tusas River Canyon a body of Tres Piedras Granite is found in which approximately one-third of the total rock is composed of schist xenoliths (Barker, 1958). In a location southeast of the well defined contact, a wide diffuse contact zone between granite-gneiss and muscovite quartzite is found to consist of irregular sills which form a hybrid zone between the two units. To quote Barker (1958), “These sills have absorbed varying amounts of muscovite quartzite and granite. The intrusion was followed by hydrothermal reactions between wall rock and granite which resulted in hybridization of xenolithic material and of granite along the contacts.” 2.4 Previous Work The Proterozoic orogenic belt in the southwestern United States is a grand area to study the growth, stabilization and reactivation of the continental lithosphere. Isotopic and geochronologic data points towards largely juvenile crust of the orogenic belt (Bennett and DePaolo, 1987) assembled between 1.75 and 1.6 Ga. Following termination of orogenic activity ca. 1.6 Ga, the belt was amagmatic until the intrusion of 1.48-1.32 Ga “anorogenic” granitoids in the southwestern United States (Anderson, 1983) and 1.1-1.2 Ga mafic intrusions in the southwestern United States (Keller et al., 65 1989). The Tres Piedras Granite along with numerous other felsic intrusions like the Embudo Granite and the Maquinita granodiorite were intruded during the phase of juvenile crust building between 1.75 and 1.6 Ga. The granite exposed along the Rio Tusas and at Tusas Mountain was called the Tusas granite by Just (1937), who also included under this name the Maquinita granodiorite and granite at Tres Piedras. The granite exposed at Tres Piedras and along Tusas Canyon and the lower Rio Tusas differs from that of Tusas Mountain. The granite at Tusas Mountain appears to be an atypical, fine grained porphyritic granite. Barker (1958) redefined the granite at Tres Piedras and along the Rio Tusas as the Tres Piedras Granite after the excellent exposure in and around that town. Hedge et al. (1977) thought Tres Piedras Granite was probably emplaced during the same intrusive event as the Maquinita Granodiorite as both foliated rock types bear the structural imprint of a single period of regional metamorphism that affected their wall rocks. The age of the granodiorite emplacement was dated by U-Pb zircon ages at 1755 m.y (Silver and Dickinson, 1987). The Tres Piedras Granite is comparable with the Embudo Granite located in north central New Mexico in the Picuris Range of the southern Sangre de Cristo Mountains in more than one aspect. Both of them have intruded the Ortega Quartzite and the Vadito Group. Petrographically and texturally they are similar. Rb- Sr whole-rock isotopic analyses indicate that the Embudo Granite crystallized 1673 +/- 41 Ma. ago (Fullager & Shiver, 1973). This intrusive event is temporally related to the igneous and metamorphic events and formation of regional structures involving the tectonic emplacement of 1.65-1.7 Ga granitoids within the supracrustal rocks (Fullager & Shiver, 1973; Pedrick et al., 1998). Maxon (1976) reported analyses of data from six Tres Piedras Granite zircon concentrations that delineate a chord with an upper intercept on concordia at 1654 Ma. For the U-Pb analysis performed for his study, zircons were obtained from four samples of Tres Piedras Granite type locality (that includes the exposures around the town of Tres Piedras) and one sample from Tusas River Canyon. In addition four zircon separates were obtained from Ortega Quartzite. The zircons all show the usual discordance in which the age pattern is Pb207/Pb206 > Pb207/U235 > Pb206/U238. The Ortega Quartzite samples on the concordia diagram intersects the chord (upper intercept) at 1830 Ma (Maxon, 1976). Throughout northern New Mexico metamorphic conditions are characterized by the co-exsistence of all three aluminosilicate polymorphs. In the Rinconada Formation, garnet-biotite thermometry indicates temperatures of 530 +/- 30 °C at 4 kbar (Holdaway and Goodge, 1990). Currently, the absolute timing of peak metamorphism in northern New Mexico is debated (Bauer and Williams, 1994; Grambling, 1989). The Rb-Sr whole rock and 40Ar-39Ar muscovite and hornblende mineral ages regionally yield ages of ca. 1.4 Ga and younger, implying that much of the metamorphism may be of this age (Mawer et al. 1989; Thompson et al. 1991; Grambling and Dallmeyer, 1993). On the other hand, U-Pb zircon ages of ca. 1.65-1.7 Ga from plutons which cross-cut strongly deformed supracrustal rocks are consistent with earlier fabric development. Karlstrom et al. (1994) suggest that locally a later (1.4 Ga) 500-600 °C, 4-5 kbar metamorphism is superimposed on an earlier (1.6 Ga) 700-800 °C, 7-9 kbar metamorphic assemblage. 66 Barker (1958) reported a pegmatitic hydrothermal event that produced numerous pegmatite injections, including the Harding Pegmatite in north central New Mexico. Gresens (1975) related the emplacement of the pegmatites to the regional metamorphism and associated them genetically with metasomatically altered metarhyolite. Most pegmatite bodies are quartzofeldspathic, with the exception of Li-rich Harding pegmatite. Montgomery (1953) assumed that the source for all pegmatite units is the Embudo Granite, which intrudes bedded Precambrian formations. Long (1972) dated the pegmatites of the La, Madera quadrangle to be 1425 +/- 25 Ma. His age calculation was based on Rb-Sr isotopic data from whole rocks, metamorphic muscovite and pegmatite “book muscovite”. Long (1972) concluded from low error on the age that the igneous intrusion and final stages of metamorphism appear to have ceased within a time interval too short to be resolved by the Rb-Sr method. The staurolite and garnet separates from a sample of the Rinconada Formation from the Picuris Mountains analyzed by Lanzirotti and Hanson (1997) yielded U-Pb ages of about 1461 ± 13 Ma. These data are consistent with metamorphism at 1450 Ma in northern New Mexico which results in porphyroblast growth. Table 2.1. The regional geologic history of the north-central New Mexico is summarized in the following table. AGE in Ma. GEOLOGIC EVENT ~1425 Pegmatitic event following the regional metamorphism. (Long, 1972; Gresens, 1975) ~1460 Regional metamorphism following the felsic intrusions. Peak metamorphic temperature 400-500 degrees C at 4-5 Kb pressure.(Lanzirotti and Hanson, 1997) 1480-1400 Anorogenic granitoid intrusion (Anderson, 1983, Bowring and Karlstrom, 1990) 1600 Regional metamorphism following the granite intrusion. Peak metamorphic temperature 700-800 degrees C at 7-9 Kb pressure.(Grambling and Dallmeyer, 1993) Granites and granodiorites like Maquinita granodiorite, Embudo “type” granite intruding the mafic complex and/or the metasedimentary sequence (Fullager and Shiver, 1973; Gresens, 1975). Tres Piedras Granite intrusion (Maxon, 1976) Deposition of the Hondo Group. It is thought to be a cratonic margin sedimentary sequence. (Bingler, 1974) 1650-1750 <1700 1700 Deposition of the Vadito Group.(Bingler, 1974) 1720-1765 Moppin, Pecos, Gold Hill and similar mafic complex geochemically similar to modern tholeiitic and/or calc alkaline rock found in arc or back arc assemblage (Gabelman, 1988) 67 Table 2.2 Stratigraphic nomenclature and lithologic description of supracrustal Proterozoic rocks of New Mexico. Modified after Bauer and Williams, 1989. Nomenclature Description of Lithology Age (Ma) HONDO GROUP Piedra Lumbre Formation Pilar Formation Phyllites Rinconada Formation Schists and quartzite. Ortega Formation Thick orthoquartzite. VADITO GROUP The unit is divided into a lower member dominated by conglomerate with minor interlayered micaceous quartzite and an upper member dominated by micaceous and feldspathic quartzite. Burned Mt. Formation Massive quartz-feldspar rock characterized by distinctive quartz and feldspar eyes in a fine grained laminated matrix. Marquenas Formation Consists of approximately equal amount of metaconglomerate and quartzite. Glenwoody Formation Approximately 1000 ft of feldspathic schist containing quartz and feldspar megacrysts. MAFIC METAVOLCANIC SEQUENCE 1700 Ma. Pecos Complex 1720 Ma. U-Pb zircon dating of metavolcanic rocks (Bowring and Condie, 1982, Robertson and Condie, 1989). Big Rock Formation Gold Hill Complex Moppin Complex Consists of 1. metavolcanic rocks with minor interbedded volcaniclastic rocks and iron formation; 2. metasedimentary rocks including feldspathic sandstone, shale, volcaniclastic graywacke, carbonate and distal iron formation; 3. Intrusive felsic and mafic rocks of a subvolcanic complex; 4. Younger granitic rocks. A mafic/felsic layered gneiss succession consisting of complexly interlayered feldspathic gneiss, biotitehornblende gneiss, hornblende gneiss and amphibolite. The sequence is interpreted as metamorphosed volcaniclastic rocks interlayered with flows, tuffs and sedimentary rocks. Mafic metavolcanic sequence with chlorite schist and/or amphibolite with lesser components of feldspathic schist and gneiss, muscovite (+/-biotite) schist, metaconglomerate and BIF. Locally exposed primary volcanic and sedimentary structures include pillows, graded beds and relict phenocrysts. The Moppin complex is associated with (intruded by?) several granodiorite bodies collectively called the Maquinita Granodiorite. 68 Metavolcanic rocks from several localities within the Vadito Group have yielded U-Pb zircon isotopic ages of approximately 1700 Ma. (Silver, 1984 as referenced in Bauer and Williams, 1989) 1765 Ma. U-Pb zircon data from a felsic metatuff near Gold Hill (Bowring, 1986) Older than 1755 Ma. U-Pb age of zircons from Maquinita Granodiorite is (Silver and Dickinson 1987) 2.5 Results 2.5.1 Rb-Sr Whole Rock Analysis In the Rb-Sr whole rock analysis of the of the Tres Piedras Granite (Table 2.4) , the following samples were analyzed: eleven samples from the Tres Piedras Granite type locality (near the town of Tres Piedras) and five samples from the Tres Piedras Granite exposed along the Tusas River Canyon. The ages determined by Isoplot 3 (Ludwig, 2001) and plotted in the cubic least square isochron diagram were as follows: Tres Piedras Granite Type Locality – 1490 Ma ± 20 Ma with an initial Sr87/Sr86 ratio of 0.7182 ± 0.0006; Tres Piedras Granite at Tusas River – 1497 Ma ± 42 Ma with an initial Sr87/Sr86 ratio of 0.7147 ± .0012 (Figure 2.2a, b, c). The data indicates that the RbSr system of the Tres Piedras Granite is not much disturbed and the data delineate good isochrons (MSWD <3) for the both the bodies studied. 2.5.2 Rb-Sr Mineral isochron Analyses were made on total rock samples and mineral separates including sphene, feldspar and biotite from Tres Piedras Granite type locality sample and Tusas River Canyon sample. Using the Rb-Sr isochron diagram, the measured ratios 87Sr/86Sr and 87Rb/86Sr are plotted for the two locations. Analytical error is less than the size of the symbol. If all of the minerals have reequilibrated at the time of metamorphism and were subsequently closed they would all lie on a common straight line. The data for the Tres Piedras Granite type locality sample is shown in Figure 2.3, Table 2.5 . The biotite-whole rock-sphene isochron corresponds to an age of 1472±40 Ma with an initial 87Sr/86Sr of 0.724±0.037 with feldspar lying above the isochron (Figure 2.3a). The deviation of the feldspar from the biotite-whole rock-sphene from the Tres Piedras type locality sample may represent incomplete re-equilibration or possibly due to later contamination by radiogenic strontium and/or rubidium loss. The Rb-Sr mineral isochron obtained from the Tusas River Canyon locality yielded an isochron of 1509±40 Ma with initial 87Sr/86Sr =0.731±0.031 (Figure 2.3b). In this case also the feldspar did not fall on the biotite-whole rock-sphene isochron and was lying above the isochron. 2.5.3 U – Pb zircon ages by LA-MS-ICPMS. Zircon were separated from the Tres Piedras Granite type locality and Tusas River Canyon samples and analyzed with LA-MS-ICPMS facility at the University of Arizona (Table 2.6). When plotted on the concordia diagram 20 zircon analyses from the Tres Piedras type locality samples yielded a chord with the upper intercept at 1732.7±4.2 Ma (Figure 2.4a). 29 zircon measurement from the Tusas River Canyon locality delineates a chord in the concordia diagram with the upper intercept at 1729±20 Ma and lower intercept at 90±20 Ma (Figure 2.4b). The zircon populations of the Tusas River Canyon are affected by lead loss more than the Tres Piedras type locality samples. Thus the age of crystallization of the Tres Piedras Granite is ~1730 Ma as evidenced by the LA-MS-ICPMS analysis of the zircons separated from the granite samples from two different localities. 69 2.6 Discussion The U-Pb zircon ages of the Tres Piedras Granite and Tusas River Canyon are significantly different from the Rb-Sr whole rock ages in both the locations-Tres Piedras Granite type locality and Tusas River Canyon. The possible explanations for this difference can be: A) The Rb-Sr whole rock age (~1470 Ma) is the crystallization age of the Tres Piedras Granite and the isotope dilution U-Pb age (~1654 Ma) of zircons (Maxon, 1976) are due to inheritance of xenocrysts. If the Rb-Sr whole-rock age is the true crystallization age of the Tres Piedras Granite, then the 1654 Ma U-Pb age of the zircons might be biased due to incorporation of older zircon xenocrysts. This could occur if 1500 Ma granite assimilated wall rock during emplacement. In that case the 1654 Ma age may represent a “mixing” age between younger 1500 Ma zircon and unreset or partly reset older zircons. The rocks which are in intrusive contacts with the Tres Piedras Granite are predominately quartz-muscovite schists which are believed to be metasomatized rhyolites (Gresens, 1971), and a pluton of calc-alkaline granodiorite (Maquinita Granodiorite). Barker, 1974 determined that the ages of these bodies fall into the 1750-1800 Ma age range as determined by UPb zircon ages. Tres Piedras Granite that intruded the 1890 Ma Ortega Formation (Maxon, 1976) could also assimilate small amount of the quartzite. Therefore these rocks provide a source for the zircons that could be later incorporated into the possibly 1500 Ma Tres Piedras Granite. This problem can be easily resolved by looking at the single zircon ages. Among the 20 zircons analyzed by laser ablation ICPMS in the Tres Piedras Granite sample more than 18 grains are between 97 to 100 % concordant (difference between 206Pb/238U age and 206 Pb/207Pb age). The presence of a high percent of concordant zircon between 1720 and 1730 Ma excludes the possibility of incorporation of xenocrystic zircons. In that case some type of disturbance must have occurred in the Rb-Sr isotopic systems of the Tres Piedras Granite to cause the isochron ages to differ from the zircon ages. This might give rise to two possibilities for this disturbance in the Rb-Sr whole rock system of the Tres Piedras Granite. B) Migration of the isotopes (open system) and lowering of the age by this migration. In an open system, with the lowering of age, there is a need to maintain isochron linearity and the process must involve the migration of either both Rb and Sr or loss or gain of equal amount of Sr or Rb respectively from each sample. The isochron obtained for the whole rock samples show a good linearity of data points. This low spread in the data (MSWD < 3.0) suggests that the change in the 87 Sr/86Sr or the 87Rb/86Sr ratios of each sample due to isotope exchange must be closely proportional to the particular Rb/Sr ratio of the sample. This would require either the addition of a proportional amount of Rb and the removal of a proportional amount of Sr from all samples or both. It is doubtful whether trace element metasomatism would proceed in such an orderly manner and the close alignment of the data points makes implausible demands upon a transport mechanism. C) The internal redistribution or isotopic homogenization at approximately 1500 Ma. 70 Tres Piedras Granite crystallized at ~ 1700 Ma but was later affected by regional metamorphism at ~1500 Ma that redistributed and homogenized the Sr isotope. The high initial 87Sr/86Sr ratio (0.7182 and 0.7147) and good isochron linearity indicate that this homogenization at approximately 1500 Ma was on a scale larger than the size of the whole-rock samples and was at least the size of the domain over which the samples were collected (approx. 50 sq. mt. of collecting area at each exposure). The best fit line yields an age that seem to correspond well with Lanzirotti and Hanson’s (1997) staurolite and monazite age of amphibolite grade metamorphism between at 1480 Ma. from the Picuris range. If each collecting location represents a single domain in which the whole-rock isochron has undergone resetting due to metamorphism, the data then appear to indicate that limits can be set on the extent of isotopic homogenization for the Tres Piedras Granite. D) Statement of preferred interpretation. If the Tres Piedras Granite has undergone isotopic homogenization on the outcrop scale, it should show similar behavior on the mineral scale too. Both the Tres Piedras Granite sample and the Tusas River Canyon sample define a biotite-whole rock-sphene isochron of 1472±40 Ma and 1509±40 Ma respectively. The initial 87Sr/86Sr of the two locations are similar within error (0.724±0.027 and 0.731±0.031). Thus isotopic homogenization of between 1470 and 1500 Ma is evidenced in the RbSr whole rock as well as the biotite-whole rock- sphene isochron. Feldspar analyzed in both the samples fall off the isochron and lies above it in both cases. The deviation of the feldspar grains from the biotite-whole rock-sphene isochron is possibly due to incomplete equilibration of the feldspar or later contamination. Thus if the feldspar remained an open system in the Tres Piedras type locality with respect to rubidium and/or strontium it might have developed a concentration gradient with respect to rubidium and/or strontium. In order to check any inhomogenity in the feldspar grains (separated from Tres Piedras type locality sample), the Rb and Sr concentrations were measured across single feldspar grains by laser ablation in Finnigen element ICP-MS. Several grains were measured and the Rb/K, Sr/K and Rb/Sr ratio of one traverse of a single grain are shown in Figure 2.5. None of the grains showed any concentration gradient from core to rim. The local variability of the ratios can be attributed to a small degree of sericitization of feldspars. In the absence of any concentration gradient of Rb and Sr it was necessary to evaluate whether or not the feldspars were homogenized with respect to Sr isotope. Approximately 100 mg of feldspar grains with well developed crystal faces were hand picked under a binocular microscope. They were partially dissolved in cold distilled 3:1 HF-HNO3 acid for 15 minutes which dissolved 20 % (by weight) of the grains. This dissolved the rims of the feldspar grains and the remaining partially dissolved grains were similarly treated with distilled 3:1 HFHNO3. The core and rim solutions of the feldspars were passed through the cation column to separate Sr that was measured for the 87Sr/86Sr ratio in the TIMS. The 87Sr/86Sr ratio obtained for the rim is 0.88429 ± 0.00011 and that for the cores is 0.85071 ± 0.00011 indicating that with respect to the Sr isotopic ratio the feldspar grains are inhomogeneous, possibly causing the feldspar to deviate from the 1472 Ma mineral isochron of the Tres Piedras type locality sample. Assuming that the feldspars were homogenized at the 1500 Ma metamorphic 71 event, the Δ87Sr/86Sr of the rim and the core observed at present can be caused by a difference of 0.56 in the Rb/Sr ratio of the rim and the core at 1500 Ma. Discordance in the Rb-Sr whole rock age and U-Pb zircon ages of granitic plutons are common in the literature, but the cause of this difference is rarely complete isotopic homogenization of Sr isotope during a later metamorphic or metasomatic event. In most cases the discordant Rb-Sr whole rock age is explained by resetting, which rarely corresponds to the regional metamorphic age. In other cases the U-Pb isotope dilution zircon ages are older than the true crystallization age due to incorporation of older xenocrystic zircons. Examples include the Crossnore Complex of the southern Appalachians (e.g. Davis et al., 1962; Rankin et al., 1969; Odom and Fullager, 1984), World Beater Complex of California (Lanphere et al., 1963), Mount Isa of Australia (Page, 1987)etc. Rb-Sr data on the Neoproterozoic alkalic to peralkaline rift related granites of the Crossnore plutonic suite yielded and age between 681-706 Ma (Odom and Fullager, 1984) whereas the zircon ages ranged between 690 and 820 Ma (e.g. Davis et al., 1962; Rankin et al., 1969; Odom and Fullager, 1984). It was established that there exists at least two zircon populations of different ages (Odom and Fullager, 1984) and more recent studies (Su et al., 1994) determined the age of the Crossnore plutonic complex to be 741 Ma, with inherited pre-Grenville source components of 1424 Ma that corresponds well with the whole rock Nd model ages of 1300-1500 Ma. In the case of the Tres Piedras Granite the laser ablation study of single zircon clarified the absence of any inherited zircon xenocrysts that might have caused the difference in age between the Rb-Sr whole rock and U-Pb (isotope dilution) age of the zircons. Page (1978) conducted a detailed study of the Rb-Sr whole-rock and mineral and U-Pb zircons of the Proterozoic silicic volcanic rocks and granitic intrusions from near Mt. Isa, northwest Queensland, Australia. U-Pb zircon ages within the basement igneous succession show that the oldest recognized crustal development was the out pouring of acid volcanics at 1865 Ma which were intruded by coeval granites and granodiorites (Kalkadoon Granite) whose U-Pb age is ~1862 Ma. All of these rocks are altered in various degrees by low grade metamorphic events that can be related to the emplacement of a syntectonic granite batholith (Wonga Granite) at 1670 Ma ago. The rocks that significantly predate this earliest recognized metamorphism have had their primary RbSr whole rock systematics profoundly disturbed as evidenced by 10 to 15 % lowering of most RbSr isochron ages and many of the lowered ages (some of which are in conflict with unequivocal geological relationships) are within the 1600-1700 Ma interval. Page (1978) argued that the processes that modified the Rb-Sr whole rock ages are complex and may have been related to the first major greenschist metamorphism (1670-1625 Ma) and a younger greenschist facies event (1620-1490 Ma) superimposed further isotopic re-distribution for Rb-Sr isotopes. The Kalkadoon Granite gave different initial 87Sr/86Sr ratios for different sets of mineral isochrons thus producing parallel isochrons for samples from different locations of the same pluton. Thus, Page attributed this to source characteristics rather than a secondary incomplete homogenization feature. He explained the parallel isochrons with different initial ratios produced by different pulses of magmatism in the same granite body. The initial 87Sr/86Sr ratios obtained from the whole-rock analysis of Tres Piedras Granite and Tusas River Canyon range from 0.7147 to 0.7182 (though they are indistinguishable within error margin). 72 The two 87Sr/86Sr ratios observed in the Tres Piedras Granite can be generated in 250 Ma with a 87 Rb/86Sr ratio of 3 and 4 respectively with a starting 87Sr/86Sr ratio of 0.704 of the primary magma. If we assume a fixed elemental ratio of 87Rb/86Sr =2.8 (average of all the analyzed whole rocks) the initial 87Sr/86Sr ratio required are 0.705 and 0.708 respectively. Partial isotopic homogenization of the minerals during a metamorphic event is common also. In the World Beater Complex of Paramint Range in California the Precambrian and Paleozoic rocks have been metamorphosed in late Mesozoic time. K-Ar ages on biotite gives ages ranging from 103-130 Ma (Lanphere et al., 1963). Rb-Sr isotopic studies of all the constituent minerals (apatite, plagioclase, K-feldspar, muscovite and biotite) and their associated total rock yielded biotite-total rock isochrons indicating ages ranging from 64 to 156 Ma (Lanphere et al., 1963). In each case apatite and muscovite was lying below and above the isochron and the observation was explained by incomplete isotopic homogenization of these two mineral phases. Discrepancy in age Rb-Sr whole rock and U-Pb zircon age is also recorded in the granites and rhyolites of St Francosis Mountain of southeastern Missouri. The U-Pb zircon ages are in the range of ~1530 Ma where as the Rb-Sr whole rock ages are between 1273 Ma to 1408 Ma (Bickford and Odom, 1969; Mose and Bickford, 1972). Initial 87Sr/86Sr of all the rocks are characterized by large errors caused in part by scatter of analytical points (effect more pronounced particularly for the volcanic rocks) but mostly by the high 87Rb/86Sr ratios that are so characteristics of these rocks. These uncertainties in the initial ratios do not appreciably affect ages computed from the isochron, but they make the initial 87Sr/86Sr ratios virtually useless for petrogenetic interpretations. Rb-Sr ages of the volcanic rocks show a significant positive correlation with Sr concentrations (Mose and Bickford, 1972) suggesting that Sr loss is responsible for lowering the Rb-Sr age determination. Thus it was concluded by Bickford and Mose (1975) that the igneous rocks of the St Francosis Mountain were formed at ~ 1500 Ma and some subsequent event disturbed the Rb-Sr system. As the rocks were not affected by regional metamorphism the authors postulated that Sr loss could have occurred during widespread hydrothermal alteration of the rocks, an event suggested by the generally turbid feldspar, alteration of mafic minerals and occurrence of epidote. A hydrothermal event is also suggested by iron mineralization in the area as typically hematite and magnetite replace the volcanic rocks. Though there is evidence of hydrothermal event (pegmatite injections) in the north central New Mexico at ~1425 Ma but the Tres Piedras Granite do not show any sign of hydrothermal alteration either in hand specimen or under the microscope. Diversity of age clusters obtained from regional metamorphic areas are very often interpreted as the result of isotopic homogenization that causes the "resetting" and "rejuvenation" of age values. Largescale isotopic homogenization is rarely attained; and when it does, it requires extensive reaction and chemical exchange with a pervasive fluid representing one end member of fluid-rock interaction. During metamorphism fluids result due to dehydration reactions and are quickly removed from the system pervasively in case of permeable rocks and through cracks and fractures in impermeable or less permeable rocks. It is more common that fluid – rock interaction will be in local scale and patchy (Cartwright and Valley, 1992; Rumble et al., 1983; Rye et al., 1976; Sheppard and Schwarcz, 1970; Valley and O'Neil, 1984) rather than to be in regional scale. 73 The Tres Piedras Granite has been affected by regional metamorphism and there is evidence of alkali metasomatism during the intrusion of pegmatite (Bingler, 1974). Regional metamorphism is the most widespread of any type of metamorphism occurring over broad expanses of deeper levels of the crust that involves reconstitutive changes in fabric and composition of the rock body. Reconstitutive metamorphic changes occur in the solid state, in the presence of usually minor amounts of aqueous or carbonic fluids. If the homogenization process of the Tres Piedras Granite represents such a solid state reconstitution it will be interesting to study the O-isotope partition as that might suggest the extent to which Si, Al, O participated in the homogenization and whether the process should be regarded as an exchange process or a recrystallization process. 2.7 Conclusion A fundamental problem in geochronology in the interpretation of the discordant mineral and total rock ages which are commonly obtained in areas which have had a polymetamorphic histrory. However the age discordance itself provides a valuable tool in the study of the petrogenesis and evolution of a polymetamorphic terrane. There can be three possibilities that can explain the discordance of the Rb-Sr whole-rock age and the U-Pb zircon age: A) the U-Pb zircon ages is biased due to incorporation of older zircons B) the Tres Piedras Granite was an open system and Rb-Sr whole rock age is lowered by the migration of the element(s) C) internal redistribution or isotopic homogenization of the Sr isotopes at 1500 Ma. U-Pb analysis of single zircons by Laser ablation MC-ICPMS and presence of a large number of concordant zircons between 1720-1730 Ma excluded the possibility of presence of xenocrystic zircons. If the U-Pb zircon ages of ~1720 Ma indicate the time of crystallization, then some type of disturbance must have occurred in the Rb-Sr isotopic systems of the Tres Piedras Granite to cause the type locality and Tusas River Canyon isochron ages to differ from the zircon age. Open system behavior is not favored as that would require equal loss or gain of Sr and Rb respectively from all the samples. The good linearity of the isochron with minimum spread of the data along with the high initial 87Sr/86Sr ratio favors the possibility for this disturbance is internal redistribution or isotopic homogenization at 1500 Ma. The high initial ratio and good isochron fit of the whole rock samples of Tres Piedras Granite from two locations indicate that this homogenization at approximately 1500 Ma. was on a scale larger than the size of the whole rock samples and was atleast the size of the domain over which the samples were collected (approx. 100 sq. mt. of collecting area at each exposure). The mineral isochron in both the locations corresponds well with the whole-rock Rb-Sr age. The strontium isotopic composition of sphene and biotite was almost completely homogenized. The degree of strontium isotopic homogenization is less complete for feldspar. The 1500 Ma. isotopic homogenization age seems to correspond well with Lanzirotti and Hanson (1997) staurolite and monazite age of amphibolite grade metamorphism at 1480 Ma. from Picuris range. If each location represents a single domain in which the whole rock isochron has undergone resetting due to metamorphism, the data then appears to indicate that limits can be set on the extent of isotopic homogenization for the Tres Piedras Granite. 74 Table 2.3. Chemical analysis, norm and modes of the Tres Piedras Granite. Chemical Analysis (Barker, 1958) Norm (Barker, 1958) Tres Piedras Granite (n=16) Tres Piedras Granite SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO Na2O K2O P2O5 H2O CO2 TOTAL 77.25 0.14 11.56 0.55 0.92 0.03 0.11 0.27 2.65 5.59 0.06 0.53 0 99.66 Quartz Orthoclase Albite Anorthite Magnetite Ilmenite Enstatite Ferrosilite Corundum TOTAL 39.5 33.5 22.5 1.1 0.7 0.3 0.3 1.1 0.6 98.6 Modes Quartz Microcline Albite Muscovite Biotite Total 75 Barker, 1958 Bingler, 1965 This Study (n=6) This Study (n=5) Tres Piedras Granite Tres Piedras Granite Tres Piedras Granite 43 38 13 3 3 100 47 16 23 10 4 100 44 31 24 1 100 Tusas River Canyon 45 43 12 1 101 Table 2.4 Rb-Sr isotopic data obtained by mass spectrometric analysis of whole rock samples of the Tres Piedras Granite. SAMPLE Rb, ppm Sr, ppm TPGT 1 TPGT 2 TPGT 3 TPGT 4 TPGT 5 TPGT 6 TPGT 7 TPGT 8 TPGT 9 LT-02-05 LT-02-4B TPGT 11 TPGT 12 TPGT 13 TPGT 14 TPGT 15 139.6 55.7 68.3 88.6 145.5 64.1 63.9 180.9 71.1 140.5 20.1 124.6 109.2 74.5 128.9 155.4 177.2 135.7 183.1 152.1 165.3 158.6 158.4 85.8 150.0 60.17 224.3 156.5 158.3 189.6 130.3 162.9 Rb87/Sr86 2.250 ± 0.0007 1.169 ± 0.0010 1.062 ± 0.0007 1.661 ± 0.0005 2.515 ± 0.0030 1.150 ± 0.0009 1.149 ± 0.0014 6.090 ± 0.0012 1.350 ± 0.0014 6.538 ± 0.0005 0.301 ± 0.0013 2.272 ± 0.0025 1.958 ± 0.0012 1.119 ± 0.0007 2.827 ± 0.0005 2.725 ± 0.0005 Sr87/Sr86 0.7669 0.7432 0.7410 0.7526 0.7706 0.7406 0.7427 0.8455 0.7473 0.8593 0.7228 0.7623 0.7563 0.7381 0.7740 0.7722 Sr/86Sr measured ratio normalized to 86Sr/88Sr = 0.1194. Precision of replicate whole rock determinations is 0.001% for 87Sr/86Sr. 87 Rb – Sr Whole Rock Ages Location Type Locality Tusas River Canyon Age 1490 +/- 20 m.y 1497 +/- 42 m.y (Sr87/Sr86)0 0.7182 +/- 0.0006 0.7147 +/- 0.0013 76 ± 0.0008 ± 0.0010 ± 0.0007 ± 0.0008 ± 0.0006 ± 0.0006 ± 0.0009 ± 0.0003 ± 0.0004 ± 0.0004 ± 0.0009 ± 0.0008 ± 0.0006 ± 0.0004 ± 0.0003 ± 0.0005 LOCATION Type Locality Type Locality Type Locality Type Locality Type Locality Type Locality Type Locality Type Locality Type Locality Type Locality Type Locality Tusas River Canyon Tusas River Canyon Tusas River Canyon Tusas River Canyon Tusas River Canyon Table 2.5 Rb-Sr isotopic data obtained by mass spectrometric analysis of mineral phases of the Tres Piedras Granite. Rb87/Sr86 Sr87/Sr86 Whole Rock Sphene Feldspar Biotite 6.538 ± 0.0005 0.0496 ± 0.0015 1.1298 ± 0.0010 43.437 ± 0.0019 0.85934 ± 0.0004 0.72753 ± 0.00011 0.88281 ± 0.00021 1.64287 ± 0.00036 Type Locality LT-02-05 Whole Rock Sphene Feldspar Biotite 1.616 ± 0.0003 0.0869 ± 0.0019 0.6662 ± 0.0020 45.642 ± 0.0012 0.78519 ± 0.00018 0.73495 ± 0.00011 0.83400 ± 0.00021 1.74287 ± 0.00035 Tusas River Canyon LT-02-2 SAMPLE *87Sr/86Sr normalized assuming 86Sr/88Sr = 0.1194 Rb – Sr Mineral Isochron Ages Location Type Locality Tusas River Canyon Age 1472±40 Ma 1509±40 Ma (Sr87/Sr86)0 0.724±0.037 0.731±0.031 77 LOCATION Table 2.6 Tres Piedras Granite Type Locality Zircon analysis by LA-MC-ICPMS CONCENTRATIONS AND RATIOS U (ppm) 206Pb U/Th 207Pb* ±(%) 206Pb* ±(%) Error 204 235 238 Pb U U corr 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 6534 970 469 1866 1173 2255 1709 824 3050 4845 393 1270 563 1166 1595 1131 1185 665 1357 221 36729 16625 6226 18184 13776 9443 8253 10174 48107 47685 14182 27442 13121 58470 54324 23449 12221 23459 25612 5957 6.2 4.8 4.2 6.0 6.1 4.3 4.8 4.7 4.4 5.0 2.7 5.3 5.1 4.3 6.1 4.0 5.9 4.6 5.3 1.6 4.4560 4.4470 4.4389 4.4732 4.4396 4.4486 4.4822 4.4605 4.5093 4.5139 4.5318 4.4704 4.4517 4.6671 4.4458 4.6495 4.4481 4.4458 4.4625 4.4105 1.90 2.24 2.15 1.54 1.78 1.81 2.37 2.54 2.25 1.75 3.57 1.98 1.72 1.57 1.45 2.25 1.42 3.66 2.43 1.43 0.3045 0.3041 0.3051 0.3034 0.3058 0.3061 0.3054 0.3053 0.3080 0.3080 0.3101 0.3075 0.3049 0.3185 0.3012 0.3172 0.3049 0.3017 0.3034 0.3027 1.00 2.00 1.79 1.17 1.46 1.49 2.14 2.33 2.01 1.43 3.42 1.71 1.39 1.21 1.05 2.02 1.00 3.52 2.21 1.00 0.53 0.89 0.83 0.76 0.82 0.82 0.90 0.92 0.90 0.82 0.96 0.86 0.81 0.77 0.72 0.90 0.71 0.96 0.91 0.70 206 Pb 238 U 1713.7 1711.9 1716.9 1708.4 1720.4 1721.5 1718.4 1717.5 1731.1 1731.3 1741.5 1728.7 1715.8 1782.6 1697.7 1776.4 1716.0 1699.7 1708.3 1705.0 APPARENT AGES (Ma) ±(Ma) 207Pb ±(Ma) 206Pb ±(Ma) 235 207 U Pb 15.0 30.1 27.0 17.6 22.0 22.5 32.3 35.1 30.5 21.7 52.2 25.9 20.9 18.8 15.7 31.4 15.1 52.6 33.2 15.0 1722.8 1721.1 1719.6 1726.0 1719.8 1721.5 1727.7 1723.7 1732.7 1733.5 1736.8 1725.5 1722.0 1761.4 1720.9 1758.2 1721.3 1720.9 1724.0 1714.3 15.7 18.5 17.8 12.8 14.8 15.0 19.7 21.1 18.7 14.5 29.7 16.4 14.3 13.1 12.0 18.8 11.8 30.4 20.1 11.9 1734 1732 1723 1747 1719 1721 1739 1731 1735 1736 1731 1722 1730 1736 1749 1737 1728 1747 1743 1726 30 18 22 18 19 19 19 19 18 18 19 18 19 18 18 18 18 18 18 19 U concentration has an uncertainty of ~15%. Isotope ratios are corrected for Pb/U fractionation by comparison with standard zircon with an age of 564+/-4 Ma. Initial Pb composition interpreted from Stacey and Kramers (1975), with uncertainties of 1.0 for 206Pb/ 204Pb and 0.3 for 207Pb/204Pb. 78 Table 2.6 continued.Tusas River Canyon Zircon analysis by LA-MC-ICPMS U (ppm) 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28 29 200 2489 843 969 2110 1154 1286 294 1272 442 1836 209 201 361 1159 2107 377 180 755 1496 468 571 425 1251 1321 2743 647 933 446 CONCENTRATIONS AND RATIOS Pb U/Th 207Pb* ±(%) 206Pb* ±(%) 204 235 238 Pb U U Error corr 5893 2.4 4.3267 2.54 0.3004 1.00 1049 16.6 1.1977 2.96 0.0937 2.68 2663 3.6 4.4154 8.63 0.3005 8.35 1917 3.1 2.4435 4.15 0.1687 3.88 1290 2.3 1.3191 2.16 0.0940 1.89 2078 3.0 2.3748 2.62 0.1628 2.41 1411 4.8 1.6717 3.94 0.1238 3.74 22711 3.7 4.4075 1.64 0.3005 1.30 884 3.2 1.7329 1.93 0.1191 1.41 6619 3.8 3.9825 3.14 0.2690 2.92 968 2.4 1.2998 2.84 0.0949 2.54 17484 3.8 4.2284 3.24 0.2956 2.28 22725 3.0 4.3606 2.59 0.2985 2.07 8388 3.1 4.4141 1.77 0.3015 1.00 3788 4.2 4.2985 3.87 0.2984 3.73 1864 4.0 1.3994 3.56 0.0951 2.74 33317 3.5 4.3879 1.59 0.2966 1.04 22135 3.4 4.2764 2.35 0.2997 1.26 4301 2.6 3.2171 2.17 0.2193 1.91 914 3.6 1.6348 3.28 0.1151 2.73 25241 4.0 4.4092 2.15 0.2985 1.90 8897 4.2 4.4012 8.29 0.3017 8.22 6057 3.9 4.4613 3.32 0.3062 3.16 1254 2.7 2.2571 2.36 0.1524 1.98 2441 3.0 2.3126 2.52 0.1626 2.27 902 4.0 1.4066 2.51 0.1067 1.43 4189 3.3 4.3407 3.61 0.3004 3.47 1754 3.3 2.2604 8.82 0.1576 8.65 9223 3.7 3.8949 2.04 0.2685 1.76 0.39 0.91 0.97 0.93 0.88 0.92 0.95 0.79 0.73 0.93 0.90 0.70 0.80 0.56 0.96 0.77 0.65 0.54 0.88 0.83 0.88 0.99 0.95 0.84 0.90 0.57 0.96 0.98 0.86 206 206 Pb 238 U 1693.7 577.5 1694.2 1004.9 579.2 972.6 752.7 1694.2 725.8 1536.2 584.7 1669.7 1684.2 1698.8 1683.8 586.0 1674.8 1690.0 1278.3 702.5 1684.3 1699.9 1722.2 914.7 971.7 653.9 1693.5 943.4 1533.4 APPARENT AGES (Ma) ±(Ma) 207Pb ±(Ma) 206Pb ±(Ma) 235 207 U Pb 14.9 14.8 124.4 36.1 10.5 21.8 26.6 19.4 9.7 39.9 14.2 33.5 30.7 14.9 55.3 15.4 15.3 18.7 22.2 18.2 28.2 122.8 47.8 16.9 20.5 8.9 51.7 75.9 24.0 1698.5 799.6 1715.2 1255.5 854.2 1235.1 997.8 1713.8 1020.8 1630.6 845.6 1679.6 1704.9 1715.0 1693.1 888.7 1710.1 1688.8 1461.3 983.7 1714.1 1712.6 1723.8 1199.0 1216.2 891.7 1701.1 1200.0 1612.6 21.0 16.4 71.6 29.9 12.5 18.7 25.0 13.6 12.4 25.5 16.3 26.6 21.4 14.7 31.9 21.1 13.1 19.4 16.8 20.7 17.8 68.7 27.5 16.6 17.8 14.9 29.8 62.1 16.4 1704 1482 1741 1715 1657 1728 1584 1738 1722 1755 1611 1692 1730 1735 1705 1743 1754 1687 1738 1679 1751 1728 1726 1755 1681 1539 1711 1697 1718 43 24 40 27 19 19 23 18 24 21 24 43 29 27 19 42 22 37 19 33 18 19 18 23 20 39 19 31 19 U concentration has an uncertainty of ~15%. Isotope ratios are corrected for Pb/U fractionation by comparison with standard zircon with an age of 564+/-4 Ma. Initial Pb composition interpreted from Stacey and Kramers (1975), with uncertainties of 1.0 for 206Pb/ 204Pb and 0.3 for 207Pb/204Pb. 79 Figure 2.1. Map of Proterozoic basement uplifts and associated major faults in north-central New Mexico. After Bauer and Williams, 1989. Figure 2.1a. Map of Tres Piedras Granite sampling locations (Modified after Maxon, 1976) 80 Whole Rock Isochron of Tres Piedras Granite from the Type Locality 0.86 0.84 TPGT 8 0.82 LT-02-05 87Sr/86Sr 0.80 0.78 TPGT 5 TPGT 1 0.76 Age = 1490±20 Ma Initial 87Sr/86Sr =0.7182±0.0006 MSWD = 2.2 0.74 LT-02-4B 0.72 0.70 0 2 87Rb/86Sr 4 6 0.76 TPGT 4 TPGT 9 TPGT 7 TPGT 3 TPGT 2 TPGT 6 87Sr/86Sr 0.74 0.72 Age = 1490±20 Ma Initial 87Sr/86Sr =0.7182±0.0006 MSWD=2.2 0.70 0.0 0.4 0.8 1.2 1.6 2.0 2.4 2.8 87Rb/86Sr Figure 2.2. Whole rock Rb-Sr isochron of Tres Piedras Granite type locality samples. Isochron plotted using Isoplot 3.0 (Ludwig, 2001) 81 Whole Rock Isochron of Tres Piedras Granite from Tusas River Canyon 0.82 Age = 1497±42 Ma Initial 87Sr/86Sr =0.7147±0.0012 MSWD = 0.49 87Sr/86Sr 0.80 0.78 TPGT 14 0.76 TPGT 15 TPGT 11 TPGT 12 0.74 TPGT 13 0.72 0.70 0.0 0.4 0.8 1.2 1.6 87Rb/86Sr 2.0 2.4 2.8 Figure 2.2c. Whole rock Rb-Sr isochron of Tusas River Canyon samples. Isochron plotted using Isoplot 3.0 (Ludwig, 2001) 82 1.9 1.7 Biotite 87 Sr/86Sr 1.5 1.3 1.1 Feldspar 0.9 Tres Piedras Type Locality Age = 1472±40 Ma Initial 87Sr/86Sr =0.724±0.037 MSWD = 9 WR 0.7 Sphene 0.5 0 10 20 30 40 50 87 Rb/86Sr 1.9 1.7 Biotite 1.3 87 Sr/86Sr 1.5 1.1 Feldspar 0.9 Tusas River Canyon Age = 1509±40 Ma Initial 87Sr/86Sr =0.731±0.031 MSWD = 20 WR 0.7 Sphene 0.5 0 10 20 30 87 40 50 60 86 Rb/ Sr Figure 2.3. Sphene-whole rock-biotite mineral isochron of Tres Piedras Granite type locality and Tusas River Canyon. Isochron plotted using Isoplot 3.0 (Ludwig, 2001) 83 1.9 Biotite 1.7 Biotite 1.3 87 86 Sr/ Sr 1.5 1.1 0.9 WR 0.7 Age = 1503±65 Ma 87 86 Initial Sr/ Sr =0.732±0.030 MSWD=15 WR Sphene 0.5 0 10 20 30 87 40 50 60 86 Rb/ Sr Figure 2.3 continued. The mineral separates (sphene, biotite) and wholerock from both Tusas River Canyon and Tres Piedras Granite type locality sample plotted together. Isochron plotted using Isoplot 3.0 (Ludwig, 2001) 84 data-point error ellipses are 2σ 1800 0.32 1600 206 Pb/ 238 U 0.28 1400 0.24 Tres Piedras Type Locality Intercepts at 1732.7±6.7 Ma MSWD = 0.92 1200 0.20 1000 0.16 1.5 2.5 3.5 207 Pb/ 4.5 235 U data-point error ellipses are 2σ. 0.4 1800 U 0.3 206 Pb/ 238 1400 0.2 1000 Tusas River Canyon Intercepts at 90±20 & 1729±20 Ma MSWD = 3.5 600 0.1 200 0.0 0 1 2 3 207 Pb/ 4 5 6 235 U Figure 2.4. Concordia plot of U-Pb analysis of single zircons analyzed by LAMC-ICPMS. Concordia plotted using Isoplot 3.0 (Ludwig, 2001) 85 0.1 Rb/K Sr/K ratio of cps 0.01 0.001 0.0001 0.00001 0 200 400 600 800 1000 Distance in microns Rb/Sr (cps ratio) 1000 100 10 0 200 400 600 800 1000 Distance in microns Figure 2.5. LA- ICPMS measurement of concentration ratios across a feldspar grain from Tres Piedras Type Locality. 86 CHAPTER THREE Trace Element & Pb Isotope Studies Of The Kutch Volcanics Of NW-India 3.1 Introduction The Deccan Traps were formed at the end of the Mesozoic era by outpouring of enormous lava to form a large continental flood basalt province over vast areas (~500,000 km2) of Western, Central, and Southern India. Compositionally the Deccan lavas are primarily tholeiitic and alkali basalts with subordinate carbonatites that erupted from multiple centers of the Western Ghats, India, the most prominent being a shield volcano like structure. The enormous size and a 65 Ma (K/T boundary) eruption age of the Deccan trap magmas have made it a current and most interesting topic of research. The Deccan Volcanic Province (DVP) is thought to be linked to the Re’union plume, which is responsible for the volcanic activity on Re’union Island in the Indian Ocean (Duncan, 1981; Morgan, 1981). All current models of plate reconstruction indicate that the Indian subcontinent drifted northward subsequent to the break up of Gondwanaland, its western margin passing over the newly initiated Re’union hotspot at around 60–66 Ma (Karmalkar et al. 2000). Thus Deccan volcanic province records the first magmatism from the plume head of the Re’union hotspot. Although plume impact is a widely accepted hypothesis for the Deccan Volcanic Province (DVP), details such as the exact location of the plume, the duration of plume-related volcanism, interaction of the plume with the ambient lithosphere-asthenosphere and the nature of the plumeaffected mantle are still debated. Age data on the alkaline rocks from the DVP reveal that the alkaline magmatic activity overlapped with the main period of tholeiitic volcanism at ~66 Ma. (Pande et al., 1988). However Basu et al., (1993) noticed that the alkaline magmatism north of the DVP is slightly older (~3 Ma.) compared to the areas further south. The alkaline rocks in Kutch region of Gujarat State (NW India), which lies to the NNW of the main DVP are nowhere seen to be in direct contact with the Deccan tholeiites (Fig.3.1). This is in contrast to the minor alkaline activity on the west-coast of India farther south, which is intrusive into the Deccan Lavas (Karmalkar et al. 2000). The occurrence of alkaline rocks in Kutch and the geographic disposition of the Kutch lavas in relation to the main tholeiitic province, is therefore significant (Karmalkar et al., 2000, 1998). If the alkaline rocks of Kutch represents the initiation of DVP then it would also represent the first melts generated from the Re’union plume head. In this study, I report trace element, REE and Pb isotope data on 11 samples of alkaline and tholeiitic basalts from Kutch region. Available Nd and Sr isotopic data on the same set of samples (Bizimis, unpublished data) identifies three potential mantle components; Re´union plume, continental lithosphere and asthenosphere (Indian MORB-like). This study intends to illuminate whether or not the trace element and the Pb isotope ratios, in conjunction with the SrNd isotope ratios can identify the end member components of Kutch lava and will compare and contrast the chemical and isotopic signatures of the Kutch volcanic rocks with the main DVP. 87 Samples of the alkali basalts from Dharam hill, close to Nakhatrana in the Kutch area, give an age of 68.5+2 Ma which place their eruption prior to the peak activity of Deccan volcanism at 65 m.y. (Venkatesan et al., 1986). Tholeiitic basalts of the Deccan Traps that outcrop in the Kutch area have yielded an age of 66.8+0.3 Ma (Pande et al.,1988). Thus the ages are within error of each other. This work is a part of a larger project that proposes to precisely date the age (40Ar– 39 Ar) of Kutch volcanism in order to determine the temporal relationship with the main DVP. If the Kutch lava marks the beginning of the Deccan eruption, it will be important to study the chemical and isotopic character of these lavas in order to determine the early chemical characters of the Re´union plume and it’s interaction with the subcontinental lithospheric mantle and the asthenosphere (Indian MORB like). 3.2 Geological Setting Recent stratigraphic reconstructions and isotopic age determinations indicate the following sequence of Gondwana breakup: Gondwanaland breakup into eastern and western component and departure of the India–Madagascar from Africa occurred at around 160 Ma (Plummer & Belle, 1995). Following the initial rifting a number of mantle plumes or hotspots played significant role in further continental breakup (Storey et al.,1995). India–Madagascar rifted from Antarctica at around 130 Ma (Chand et al., 2001) followed by breakup between India–Seychelles and Madagascar c. 88 Ma (Storey et al., 1995). The breakup between –Seychelles and Madagascar is linked to the development of the Marion hotspot. Beginning of the rift between India–Seychelles and Madagascar is dated at 96 Ma and its termination at about 84 Ma by Plummer (1995). During this Late Cretaceous break-up event, the western margin of India presumably rifted off the eastern margin of Madagascar; however, equivalent Cretaceous volcanism is not well documented from western India. Possible exceptions include mafic dykes in mainland south-west India (Radhakrishna et al., 1994, 1999), and the acid volcanic rocks of the St. Mary islands (Valsangkar et al., 1981) that form a chain of small islands trending NW– SE, off the western coast of central India. Naganna (1966) argued that the acid volcanic rocks represented an early phase of Deccan flood basalt province (expected age of crystallization being slightly greater than 65 Ma), but whole rock K–Ar ages range between 80.3 ± 1.7 Ma to 97.6 ± 2.3 Ma (Valsangkar et al., 1981) suggest that this magmatism might be better correlated with an older event related to India–Madagascar break-up (Fig.3.1). Next development in the sequence was rifting of Seychelles from India. Seychelles Plateau consists of a sliver of continental crust that is traced back to some 700 Ma. (Plummer & Belle,1995). Presumably Re’union hotspot was the cause of separation at about 65 Ma at the Cretaceous/Tertiary (K/T) boundary. Gnos et al. (1997) corroborate the timing of separation and reported a counterclockwise movement of India at about the same time. This was the approximate time of Deccan volcanic event, development of the Carlsberg Spreading Ridge and initial opening of the north-western Indian Ocean basin (Plummer & Belle, 1995). It is postulated that the Deccan volcanic sequence in India was a consequence of the passage of the northerly drifting Indian subcontinent over the Re’union starting plume in the Late Cretaceous (Morgan, 1981). This plume also produced the three-rift Cambay triple junction, the three arms being the West Coast graben belt (also known as Kachehh Rift), the Narmada-Tapi rift zone and the Cambay rift (Fig.3.1). 88 The Kutch region is situated to the northwest of the main Deccan Volcanic province and proximal to the Cambay Triple junction (Fig.3.1). It lies to the south west of the zone of low seismic velocities which is interpreted as fossil plume head of the Deccan Plume (Kennett and Widiyantoro, 1999). The tectonic evolution of the Kutch region has been attributed to the rifting process and the region accordingly marks the site of paleo-rift graben whose evolutionary history dates back to the Mesozoic times (Biswas, 1987). Consistent with the northward passage of Greater India over the Reunion hotspot in the latest Cretaceous, the geochemical stratigraphy of the south- western Deccan exhibits a southwardyounging progression of volcanism (Beane et al. 1986, Devey et al., 1986). The stratigraphic relationship of flows in the little-studied northwestern part of the province to those in the southwest is unknown, but at least some of the former may represent stratigraphically lower (earlier) levels of the overall volcanic succession (Mahoney et al., 1985). An important difference between the few northwestern areas that have been studied and the southwestern Deccan formations is that alkalic and transitional flows are common in the northwestern sections, whereas the southwestern lavas are almost exclusively tholeiitic basalts (Mahoney et al., 1985). In order to ascertain the exact relation of the Kutch volcanics with the main DVP precise dates are required. Without obvious stratigraphic relationship of the Kutch volcanics with the main DVP it is important to chronologically place the Kutch volcanics in the broader picture. If they are slightly older than the Deccan volcanics they might represent the earlier phase of volcanism from the Re’union hotspot; if the age of the volcanics date back to the Jurassic the Kutch volcanism might be related with the Marion hotspot. This work will analyze the trace element and Pb isotope data in conjunction with the available Sr-Nd data on the 11 alkaline and tholeiitic samples from Kutch to chemically compare them with the main DVP and identify their affiliation. 3.3 Deccan Stratigraphy Recent stratigraphic classification of the western Deccan (referred as the main DVP or the southwestern Deccan in this study) between the Igatpuri and Amboli area (hatched area in Fig. 3.1) has divided the lava sequence to eleven formations with a maximum combined thickness exceeding 3 km, based on both field markers and general geochemical similarity of the flows (Cox and Hawkesworth, 1985; Devey and Lightfoot, 1986). The basalts lie upon and are largely surrounded by Precambrian continental crust, much of it Archean. 40Ar–39Ar dating of samples covering most of the exposed stratigraphy in the western Deccan indicates an age of ~66 Ma with no measurable difference between the upper and lower part of the succession, implying <2 Ma period of eruption for the major phase of Deccan volcanism (Duncan and Pyle, 1988). Table 3.1 summarizes the stratigraphic nomenclature and thickness of the southwestern Deccan basalt formations. 89 3.4 Previous Work – Sr and Nd Isotopic Data Sr and Nd isotopes have been measured on 7 samples of alkali basalts with olivine phenocrysts and 4 tholeiites from the Kutch region (Bizimis, personal comm.). Figure 3.2 compares the initial Sr and Nd isotopes of the Kutch lavas with the main DVP basalts. Sr and Nd isotope data for the main formations of the southwestern Deccan from Cox & Hawkesworth, 1985, Lightfoot & Hawkesworth, 1988, Peng et al.,1994, Peng et al., 1998 are plotted in Figure 3.2. Ambenali basalts with the highest initial 143Nd/144Nd (0.512640 to 0.512914) and the lowest (87Sr/86Sr)I (0.7038 to 0.7060) has been identified as the least contaminated with continental material. Nd and Sr isotopic data of the 10 formations (all except Desur of Table 3.1) define two general trends; the Ambenali-Poladpur-Bushe trend that also includes lavas from Khandala, Bhimasankar, Thakurvadi, Neral and Igatpuri –Jawhar and Ambenali-Panhala-Mahabaleshwar trend (not all the formations are shown in Fig.3.2). Lavas of the Bushe formation are characterized by very high (87Sr/86Sr)I (0.7134 to 0.7202) and very low 143Nd/144Nd initial (0.511889 to 0.511684). They appear to be contaminated by an Archean, broadly granitic crustal end-member (Peng et al., 1994). The Poladpur, Khandala, Bhimasankar, Thakurvadi, Neral and Igatpuri –Jawhar lavas define a Sr-Nd array lying between those of Ambenali and Bushe formations and appear to have been contaminated to intermediate degrees. The Mahabaleshwar formation, directly below Ambenali exhibits an isotopic array that overlaps that of the Ambenali but extends to low 143Nd/144Nd (0.512816 to 0.512258) and possesses much lower (87Sr/86Sr)I (0.7040-0.7055) than the Bushe. The Ambenali-Mahabaleshwar trend has been proposed to reflect the mixing of isotopically Ambenali like magmas with one of the following components: a) ancient granulite crust (Mahoney, 1988; Lightfoot et al., 1990) or b) old, stabilized continental lithospheric mantle (Lightfoot & Hawkesworth, 1988; Lightfoot et al., 1990). The question of which of these components was involved is still the subject of investigation. In summary, most of the formations possess Nd-Sr isotopic compositions reflecting significant continental lithospheric influences in their origin with Bushe reflecting the highest degree of contamination. The main exception is the 500 m thick Ambenali Formation in the upper part of the sequence, which is relatively uncontaminated. However, the Ambenali basalts are not equivalent to modern oceanic island products of the Re’union hotspot. Instead, the trace element characteristics of the least-contaminated Ambenali lavas resemble those of transitional MORB, and their isotopic signatures have been interpreted as a mixture of modern Re’union Island-type, broadly Indian MORB-like, and minor continental lithospheric components (Mahoney, 1988). The isotope results obtained from the least differentiated, silica-undersaturated samples from the complexes of Barmer and Bhuj in the northeastern Deccan (Simonetti et al., 1998) are plotted in Figure 3.2 for comparison with the main DVP lavas. The samples from Bhuj fall between the fields defined for present-day Indian MORB and that representing the composition of Re´union ‘plume’ mantle at ~65 Ma. The melilitites from Barmer contain higher radiogenic Sr and lower Nd isotope values than the samples from Bhuj. In addition, the melilitites from Barmer contain similar initial 87Sr/86Sr ratios but slightly lower initial 143Nd/144Nd ratios than the composition of the Re´union plume. Picritic and basaltic lavas from three drillholes in the northwestern Deccan Traps, near Kutch area (Peng & Mahoney, 1995) are plotted to compare the Nd-Sr isotopes with the main Deccan lavas. The age-corrected 143Nd/144Nd values of the drillhole lavas vary from 90 0.512855 to 0.512419, and 87Sr/86Sr, varies from 0.70414 to 0.70784; these ranges, although substantial, are smaller than observed in the southwestern Deccan. In the age corrected Sr-Nd space the Kutch lavas show three groups (Fig. 3.2). The tholeiites follow the Ambenali-Poladpur-Bushe trend with low initial 143Nd/144Nd (0.512377 to 0.512621) and high (87Sr/86Sr)I (0.70566 to 0.70894). Five of the analyzed alkali basalts plot in very close cluster (143Nd/144Nd) i varying between 0.512839 to 0.512885 and (87Sr/86Sr) i ranging from 0.70363 to 0.70379 between the fields defined by CIR and Re’union plume at 65 M.a. Two of the alkali basalt samples plot generally on the Ambenali-Mahabaleshwar trend but they have lower (87Sr/86Sr)I than the Mahabaleshwar samples and might represent a different source than the rest of the alkali basalts. 3.5 Analytical Technique Eleven samples from the Kutch region have been analyzed for their trace elemental composition and lead isotope, using ICPMS and TIMS (MAT262) respectively and the results are shown in Table 3.2. Trace elements for the alkali basalts/tholeiites were determined by solution ICP-MS analysis using a ThermoFinnigan Element ICPMS. Small chips of each sample were handpicked and crushed with agate mortar and pestle. 30 mg of each sample was weighed into screw top Teflon beakers and dissolved with 2 ml of 3:1 distilled HF-HNO3. After drying on a hot plate at 120oC, the samples were allowed to reflux overnight in concentrated HNO3. Samples were then dried again and brought to a final volume of 60 ml with 2% HNO3. The sample solutions were further diluted in the ICP vials for target concentration of 100 ppm TDS. 1 ppb of Indium was added as internal standards and drift corrections for each analyzed mass were applied by interpolating with the internal standard. The solution ICPMS analyses were calibrated against a single solution of well-characterized Hawaiian basalt standard BHVO-1 prepared identically to the samples. BCR-1 was also prepared like sample and calibrated against the BHVO-1 to check the precession of measurement. Pb isotopic composition was determined in the Finnigan MAT-262. 120 mg of sample powder was leached for ~20 min in cold 6 N HCl in an ultrasonic bath, rinsed several times with Quartz Distilled (QD) water and dissolved in 5 ml of 3-1 distilled HF-HNO3. After drying on a hot plate at 120oC, the samples were allowed to reflux overnight in concentrated HNO3 for 3 times, dried and fluxed again in 5 ml of distilled 6 N HCl twice and finally dried in concentrated HBr twice. Pb was separated using a two-column procedure (Abouchami et al., 1999) and run on a single commercial Re filament using MAT 262. Total blanks were better than 0.3 ng for Pb; and repeated (n=18) analysis of isotopic standard NBS 981 yielded 206Pb/204Pb = 16.90 ± 0.02, 207 Pb/204Pb = 15.45 ± 0.02, 208Pb/204Pb = 36.60 ± 0.04. 3.6 Results The trace element, REE and Pb isotope data are shown in Table 3.2. Concentrations of selected trace elements of the Kutch samples arranged in order of ascending compatibility and normalized to primitive-mantle and CI chondrite concentrations are shown in Fig 3.3a and 3.3b respectively. 91 The samples can broadly be divided into two groups, the alkali basalts with olivine phenocrysts and the tholeiites. 3.6.1 Alkali Basalts Alkali basalts are characterized by an overall smooth trace element pattern with decrease in absolute abundance from Nb to Yb (Fig. 3.3a) and a small range in HREE and Y contents (Fig. 3.3b). Compared to MORB concentrations these samples show enrichment in Nb and LILE. All are enriched in LREE. The ratios of Tb/Yb vary between 0.59 and 0.87 compared to the value of 0.18 in MORB and suggests fractionation of HREE. The ratios of Ti/Zr for the Kutch alkali basalts vary between 61.5 and 104.6 whereas the Zr/Y varies from 8.4 to 13.2 and Ti/Y ranges from 582 to 1015. For most primitive “MORB”, the ratios Ti/Zr, Zr/Y, and Ti/Y are about 100.5, 1.95 and 196 respectively (Salters and Stracke, 2004). The Indian Ocean MORB however is reported to have a mean Ti/Zr ratio of 84 and therefore is an exception to the values given for NMORB. The (La/Sm)n ratio for the Kutch samples is in the range of 2.7 – 4.7, which is higher than the range of 0.86– 1.5 given for MORB. The Lu/Hf values for the Kutch samples range from 0.0299 to 0.0523, which are much lower than for MORB-source mantle (~0.3). All the Kutch alkali basalts show positive Ba and Sr anomalies and a prominent negative Yb anomaly. Overall, the LREE are variously enriched over HREE ((La/Yb)n=21.7–57.3). No Europium anomaly is observed in the samples, and the scatter in trace element abundances maybe regarded as primary, or as reflecting crystal fractionation involving predominantly olivine and clinopyroxene (Simonetti et al., 1998). Due to the lack of Eu anomaly plagioclase is not considered to be involved in fractional crystallization. Fisk et al. (1988) have reported similar trends for the rocks of Reunion Island, which they have related to melt, differentiation. The incompatible element patterns of the alkaline rocks from Kutch region are broadly consistent with a dominant OIB type source component. All of the alkali basalts have high abundances of Nb, with low Zr/Nb (3.0–5.1) typical of OIB and contrary to the higher values reported for MORB (6–10). The Ce/Pb (23.5-26.3) ratios for the Kutch alkaline rocks are consistent with the relatively constant values for these ratios reported in OIB (25±5) (Sun and McDonough, 1989). The isotopic data for some samples from Kutch area characterized by lower initial 87Sr/86 Sr (0.70357–0.70396) and higher initial 143 Nd/144 Nd (0.512839–0.512885) ratios. These ratios closely correspond to those for Reunion. Most of the Kutch alkali basalts show a narrow range of (La/Yb)n (21.7–28.0)(except the two most enriched samples) and (Tb/Yb)n ratios (2.6–3.8) but have low and uniform HREE ((Yb)n contents 0.52 to 0.82). This is consistent with partial melting of variable degrees with garnet being present in residual phase. A garnet signature is also evident from the conjunction of fractionated HREE and low Y, Sc and HREE contents in magmas with Ni >150 ppm. Lu/Hf data may also be used to place some constrains on the source mineralogies. This is because the partitioning systematics between melt and residual mantle for Lu and Hf are strongly affected by the presence of garnet. The Lu/Hf values for the Kutch samples range from 0.030 to 0.052, these lower values imply melting in the garnet stability zone. 92 3.6.2 Tholeiites Considering the continental setting in which the Kutch samples have been emplaced, crustal contamination is a distinct possibility. In Fig. 3.3a, the average composition of the continental crust has been plotted along with the tholeiite samples. They are characterized by a LREE pattern similar to that of the alkali basalts, but flatter HREE (Fig. 3.3b). They show a pronounced positive Pb anomaly and only slightly positive La anomaly like that of the average continental crust. Negative Nb anomaly is similar to the continental crust but they show only modest negative Ti anomaly unlike that of the continental crust. (Th/Nb)n ranges between 5 and 7 as against 6.18 in continental crust (Rudnick and Gao, 2003). The alkali basalt samples on the contrary are characterized by lack of Nb and Pb anomalies, lower HREE, and (La/Sm)n higher than the bulk continental crust. The La/Nb ratios of the Kutch tholeiites (1.32 to 1.57) are comparable with those of the continental crust (~1.5). The indices of crustal contamination such as Ce/Pb or Rb/Sr ratio for the Kutch tholeiits vary from 6.86 to 12.09 and 0.08 to 0.15, respectively, as against values of 4.13 and 0.12 for the continental crust (Rudnick and Gao, 2003). These similarities with the bulk continental crusts are considered to have been produced by significant crustal contamination. 3.7 Comparison with the DVP Figure 3.4 illustrates Ce/Pb vs. La/Nb and La/Sm vs. Sm/Yb in the Kutch rocks. For comparison the CIR, Re’union, southwestern Deccan lavas (or lavas from the main DVP) and Northwestern lava fields are shown. Ce - Pb and La - Nb have similar partition coefficient during mantle melting and any difference in the ratio will illustrate the source character rather than degree of melting and/or fractional crystallization. On the contrary the La/Sm vs. Sm/Yb defines the entire slope of the REE and as the elements have different partition coefficient a difference in the ratio will be due to different degrees of melting and/or fractional crystallization. In the Ce/Pb vs. La/Nb plot the tholeiites shows a crustal signature with high La/Nb and low Ce/Pb. The alkali basalts plot in the Re’union field. For comparison the drill hole lavas from northwestern Deccan and lavas from the main DVP are plotted also. They show a wide variation, ranging from Re’union type source to extreme crustal contamination. In the La/Sm vs. Sm/Yb plot the alkali basalts plot above the Re’union field with the two most enriched alkali basalts having the highest La/Sm and Sm/Yb ratio. The DVP and drillhole samples plot at a much lower Sm/Yb ratio (<3.5) though they display a wide range of La/Sm ratio (1.7 to 8.5). The tholeiites plot on the DVP field. A plot of present day 206Pb/204Pb vs. 207Pb/204Pb (Fig. 3.5a) shows the data listed in Table 3.2, in addition to the fields for Archean basement, Indian MORB and the Re´union plume component. As in the case for their initial Nd and Sr isotope data (Fig. 3.2), the 5 alkali basalt samples from Kutch contain Pb isotope ratios that plot on and between the Re´union and Indian MORB fields (Fig. 3.5a). One of the two most enriched alkali basalt that follow the Ambenali-Mahableshwar trend in the initial Nd-Sr isotope space falls in between the Re´union and Indian MORB fields; the other one falls slightly below the array. The Pb isotope ratios for tholeiite samples from Kutch are clearly more radiogenic compared with the remaining samples. The tholeiite samples are enriched in 207Pb except for BH-2 which is enriched both in 207Pb and 206 Pb. When compared 93 with the southwestern Deccan lavas the alkali basalts overlap with the Ambenali in the 206 Pb/204Pb vs. 207Pb/204Pb space. A (87Sr/86Sr)I vs. 207Pb/204Pb plot of the southwestern Deccan shows two trends diverging from the CIR-Re’union overlapping field (Fig. 3.5b). 1. Mahabaleshwar trend with increasing (87Sr/86Sr)I with progressively decreasing 207Pb/204Pb 2. crustally contaminated trend with increasing (87Sr/86Sr)I and 207Pb/204Pb. The alkali basalts plot at the intersection of the two trends on the Re’union field where as the tholeiites fall on the crustal contamination trend. 3.8 Discussion In the La/Sm vs. Sm/Yb plot the alkali basalt samples plot at a higher ratio than the Re’union plume indicative of a very low degree of partial melting of Re’union type source. The primitive mantle normalized ratios of (Sm/Yb)n in the Kutch rocks are 4–6. This type of values are observed when melting initiates in presence of residual garnet (Ellam, 1992). Here we present a batch-melting model and incremental batch-melting model of garnet peridotite in the garnet stability field and spinel stability field of Re’union type source. Since the Re’union plume is not directly accessible, its composition has to be inferred via constraints derived from rocks, whose origin and composition are related to the plume related magmatism, like the Deccan volcanics. The isotopic composition of the 5 alkali basalt lavas that are thought to be produced by partial melting of the Re’union plume is used to estimate parent-daughter ratios (e.g., Sm/Nd, Rb/Sr, Lu/Hf) and the respective element concentrations in the Re’union plume (e.g., Rb, Sr, Sm, Nd, Lu, Hf, etc). The age corrected isotopic composition of alkali basalts of the Kutch directly reflects the isotopic composition of its source, the Re’union plume at 65 Ma, and is the most reliable constraint on the trace element ratio formed by the parent and daughter element of the isotopic system considered. For example, by measuring 143Nd/144Nd in alkali basalts, the ratio of the parent element Sm and the daughter element Nd (Sm/Nd) can be calculated with some information on the isotopic evolution of the Re’union plume. These calculated parent-daughter ratios form the framework for the estimates of other trace element ratios. The present day isotopic ratios are used to calculate the elemental ratios of the Re’union plume assuming the average age of the Depleted Mantle to be 2.2 Ga (Chase and Patchett, 1988; Condie, 2000) is given in Table 3.3. Table 3.3. Estimating the source composition of the Kutch alkali lavas. Present day Reservoir ratio. At 65 Ma source ratio (87Sr/86Sr) (87Rb/86Sr) (143Nd/144Nd) (147Sm/144Nd) 0.7045 0.082182 0.512638 0.703597 0.0526 0.512867 (176Hf/177Hf) (176Lu/177Hf) 0.1967 0.282772 0.0332 0.2213 0.2831094 0.0432 Present day (87Rb/86Sr) calculated assuming (87Sr/86Sr)BABI of 0.69989 at 4.57 Ga. (87Sr/86Sr)65, (143Nd/144Nd)65, (176Hf/177Hf)65 (Bizimis, personal communication) ratios are age corrected average of 5 alkali basalt samples measured for Kutch volcanics. Resulting elemental ratios are (Rb/Sr)65 = 0.01817 (Sm/Nd)65 = 0.36608 and (Lu/Hf)65 = 0.3039. 94 Figure 1 from Salters and Stracke (2004) is used for estimates of the individual elements and the inter-relationship between the estimates of the different elements. Lu concentration has been choosen as the ‘‘anchoring point’’ of the trace element pattern and fixed at Bulk Silicate Earth (0.0675 ppm) of McDonough and Sun, (1995). With the Lu concentration as a starting concentration, the Lu/Hf ratio derived from the Hf-isotope systematics is used to estimate the Hf concentration. The Hf/Sm ratio of the Re’union is assumed to be chondritic and thus the Sm concentration of the Re’union plume is calculated from the Hf concentration. Nd concentration is calculated from the Sm/Nd ration of Re’union at 65 Ma. Again Nd/Sr ratio of the Re’union is assumed to be chondritic that gives the Sr concentration and hence Rb concentration. Thus after Lu, Hf, Sm, Nd, Sr and Rb are anchored the elements in between are assigned values to generate a smooth curve in a Primitive Mantle normalized spider diagram. Thus the La, Sm and Yb concentrations are determined to be 0.35, 0.317 and 0.42 ppm respectively. In a Primitive Mantle normalized plot of (La/Sm)n vs. (Sm/Yb)n (Fig. 3.6a and 3.6b) continuous lines are for melting of Fertile Lherzolitic mantle. The Kutch alkali basalts plot can be produced by approximately 1.6% batch melting of a garnet peridotite source with 53% olivine, 8% orthopyroxene, 34% clinopyroxene and 5% garnet. Incremental batch melting of the same source yield similar results of approximately 1.6 to 1.8% melting of the source. The two extremely enriched alkali basalts cannot be produced with similar source and possible was derived from a different source. For the tholeiites, normalize incompatible element patterns of contaminated lavas should reflect the patterns of both their source and contaminant (whether a bulk rock, partial melt, etc.) weighted by the relative contribution of each. Patterns for the Kutch tholeiites show several important features typical of continental crust with negative Nb spikes and strongly positive Pb peaks and either they lack negative Ti spikes or have modest negative Ti anomalies. Published analyses of Precambrian crustal rocks (e.g., Weaver and Tarney, 1980, 1981 ; Thompson et al., 1983; Volpe and MacDougall, 1990) indicate that high-SiO2 types typically have very prominent negative Ti spikes in their incompatible element patterns, as well as marked enrichment in the highly incompatible elements other than Nb. Many basic amphibolites and basic to intermediate granulites are enriched in the highly incompatible elements and show only slight negative Ti spikes or lack them altogether (e.g. Weaver et aI., 1977; Khandelwal and Pandaya, 1988; Gopalan et al., 1990). For example, Weaver et al. (1977) identified four groups of Archean granulites in the Madras area of southern India: basic granulites, intermediate granulites, chanockites, and khondalites. Incompatible element patterns for these rocks show that the basic and intermediate varieties possess patterns qualitatively comparable to those of the Kutch tholeiites. Patterns of the basic granulites have negative Nb spikes, a pronounced positive Pb peak, and lack negative Ti spikes; patterns for the intermediate granulites, have higher Ba contents. We now apply a simple model to assess whether crustal assimilation associated with fractional crystallization can plausibly account for the isotopic and incompatible element characteristics of the common signature lavas (Fig. 3.7a and 3.7b). For the model the magma temperature was assumed to be 12000C and the temperature at the base of the crust is assumed to be 8000C. We selected the uncontaminated end-member to be similar to either CIR or a primitive Re’union type 95 parent magma. To fit the trace element patterns of the common signature lavas, we added variable amounts of a given contaminant composition to that of the uncontaminated parental magma. Several types of bulk rock contaminants were tried. The best value (i.e. by best visual fits) were achieved with bulk rock mixture of 75% CIR basalt with 18% basic granulite (n=7) and 6% charnockite (n=28) analyzed by Weaver et. al., 1978. A fractional crystallization (after the Rayleigh’s fractional crystallization model) of 18% olivine was required to generate the observed pattern. In absence of the major element chemistry it was difficult to pinpoint the nature of the fractionating crystals (olivine, clinopyroxene and/or plagioclase). Olivine fractionation gave the best visual fits. The trace element pattern for the most contaminated tholeiite, BH-2 could be generated by mixing 75% CIR basalt with 20% charnockite and 5 % basic granulite (n=10) and fractionating 30% olivine. In absence of the major element concentrations of the Kutch tholeiites, a simple calculation was done to check the SiO2 content of a contaminated MORB with 20% charnockite. Assuming 46% SiO2 concentration for CIR basalts, 20% bulk charnockite (70% SiO2) was mixed. The resultant rock will have a 50.8% SiO2 which is still within the range of silica concentrations of the Deccan basalts. Different proportions of granulite, charnockite and khondalites were mixed with the Re’union magma to test the plume as a plausible end member but we failed to generate the observed Nb anomaly and slight Ti anomaly. Neither the specific chemical nor isotopic characteristics of the lower crustal rocks located beneath the Deccan are known. However it is interesting that the optimum mixture of the mantle and granulite and charnockite end-members suggested by trace element calculations also yield appropriate Nd-Sr isotopic values. When mixed with granulites and charnockites from Krishnagiri, South India (west of Madras Granulites) the isotopic composition of the tholeiites could be generated by 18-24% mixing of the Krishnagiri granulites and charnockites analyzed by Peucat et.al., 1989 (Fig. 3.7c). Lead isotope and uranium abundance data available for the Krishnagiri charnockites and granulites have much lower 207Pb/204Pb and 206Pb/204Pb ratios than it is observed for the Kutch volcanics. However Chakrabarti and Basu (2006) measured some impact breccias from Lonar Crator that have lower 206Pb/204Pb and 208Pb/204Pb ratios but higher 207Pb/204Pb ratios compared to the host basalts. They argued that a major component of the Lonar impact breccias was derived from melting of Archean basement rocks. The basement beneath the Lonar region is believed to be similar to the Dharwar craton of peninsular India. Based on their similar Pbisotopic compositions with the breccia rocks, Chakrabarti an Basu suggested that Archean Chitradurga Group of rocks of this craton to be present in the basement beneath the Deccan lavas of the Lonar region. Thus the lead isotopic composition of the Kutch tholeiites could have been modified by the contamination of the Archean basement rocks. 3.9 Conclusion The Kutch lava shows three types of end members, OIB - type end member, Indian MORB and Archean crust. The trace element pattern and Sr-Nd-Pb isotope ratios of the alkali basalts are similar to that of OIBs and in absence of precise age data it is difficult to pinpoint it’s source (Re’union or Marion hotspot). Elemental data indicate that these lavas are probably the products of comparatively small degrees of partial melting, less than for the southwestern Deccan lavas. If 96 the Re’union plume is responsible for the Kutch lavas that erupted early in the sequence, gradual increase with time in the degree of partial melting in the Deccan (Beane et al. 1986, Devey et al., 1986) is likely to be at least partly responsible for regional north-to-south decreases in average Nb/Zr, Nb/Y. The two most enriched samples have lower (143Nd/144Nd)I than the rest of the alkali basalts. The isotopic ratios of those two samples can be explained by mixing of charnockites with either Re’union or Indian MORB type end members but that fails to explain the trace element pattern. The two enriched samples might have a different source and/or genesis than the rest of the alkali basalt and needs to be examined further in details. The tholeiites from the Kutch are significantly contaminated by crustal material of granulitic and charnockitic composition. The trace element pattern could be explained by mixing Indian MORB with the Archean crust from South India. Assuming Re’union as the uncontaminated endmember source failed to produce the observed Nb anomaly in the tholeiites. The Sr-Nd isotope ratios were compatible with our model too. Thus if Kutch lavas are the first phase of eruption of the DVP then it’s important to note that both Re’union and Indian MORB played significant role during the early phase of Deccan eruption. Table 3.1. Stratigraphic nomenclature and thickness of the southwestern Deccan formations (modified after Peng et al., 1994). Subgroup Formation Max. Thickness (87Sr/86Sr)I WAI Desur ~100 m 0.7072 – 0.7080 Panhala >175 m 0.7046 – 0.7055 Mahabaleshwar 280 m 0.7040 – 0.7055 Ambenali 500 m 0.7038 – 0.7060 Poladpur 375 m 0.7053 – 0.7083 Bushe 325 m 0.7134 – 0.7202 Khandala 140 m 0.7068 - 07107 Bhimashankar 140 m 0.7067 – 0.7076 Thakurvadi 650 m 0.7066 – 0.7112 Neral 100 m 0.7082 - 0.7104 Igatpuri-Jawar >700 m 0.7089 - 0.7124 LONAVALA KALSUBAI 97 Table 3.2. Trace element and Pb-isotope results. Sample BH 19.3 BH 12.3 BH 13.1 BH 16.2 Alkali Alkali *Alkali *Alkali Basalt Basalt Basalt Basalt 206 19.02 18.83 18.37 Pb/204Pb 18.58 207 15.53 15.56 15.50 Pb/204Pb 15.55 208 40.18 39.07 38.37 Pb/204Pb 39.14 8.86 11.95 5.58 8.35 Li 55.22 103.24 27.70 25.51 Rb 877 1516 606 572 Sr 25.75 27.74 19.38 20.82 Y 290 366 188 178 Zr 95.97 118.50 40.62 34.34 Nb 3.42 2.32 0.59 1.13 Cs 1251 2321 470 418 Ba 65.41 87.11 28.91 27.54 La 122.73 157.85 61.13 58.34 Ce 13.82 17.49 7.52 7.02 Pr 55.29 67.03 31.76 29.88 Nd 10.40 11.61 6.63 6.31 Sm 3.65 4.35 2.29 2.19 Eu 15.45 17.89 8.36 7.78 Gd 1.48 1.62 0.96 0.92 Tb 5.98 6.09 4.30 4.40 Dy 0.97 1.00 0.72 0.76 Ho 2.30 2.37 1.77 1.94 Er 1.79 1.86 1.41 1.55 Yb 0.23 0.23 0.18 0.21 Lu 6.68 7.57 4.27 4.10 Hf 5.58 6.32 2.63 2.15 Ta 4.77 5.98 2.48 2.44 Pb 9.51 11.01 3.68 3.32 Th 2.07 2.41 0.82 0.73 U 21.11 19.72 22.73 26.28 Sc 22875 22478 19676 16223 Ti 243.1 224.9 307.1 298.8 V 522.6 334.0 514.7 668.6 Cr 74.9 67.7 58.6 63.3 Co 516.1 373.9 251.6 323.5 Ni 52.0 59.8 73.8 74.7 Cu 130.9 130.7 104.4 100.8 Zn BH 7.1 Alkali Basalt 18.94 15.55 39.02 7.15 22.27 774 33.21 280 57.04 0.64 514 44.76 95.20 11.58 48.68 10.02 3.36 13.19 1.50 7.01 1.23 3.08 2.51 0.33 6.22 3.41 3.69 5.76 1.42 32.19 19336 298.1 510.5 63.6 270.1 58.8 117.1 * two most enriched alkali basalt samples. 98 N 1.1 Alkali Basalt 18.64 15.53 38.76 6.37 42.09 765 22.37 206 43.53 0.64 478 36.55 72.62 8.61 36.11 7.47 2.57 9.92 1.07 4.99 0.83 2.09 1.63 0.22 4.84 2.86 2.95 5.08 1.12 27.50 16581 306.9 712.1 73.7 380.7 75.8 107.5 BH 18.1 BH 17.1 BH 2 BH 3.1 BH 3.2 Alkali Tholeiite Tholeiite Tholeiite Tholeiite Basalt 18.68 18.71 20.09 18.31 18.34 15.52 15.62 15.83 15.61 15.66 38.75 39.06 41.10 38.75 38.90 6.68 5.93 9.30 6.72 6.20 29.41 10.82 37.32 22.50 20.58 694 141 247 207 187 21.46 26.90 45.27 40.39 36.35 220 77 262 153 133 49.17 4.69 24.62 14.18 12.63 0.63 0.50 0.60 0.61 0.56 530 120 337 326 298 36.42 7.39 34.34 19.59 16.72 74.91 15.85 72.41 39.05 33.94 8.97 2.07 8.64 4.68 4.20 37.40 9.56 35.61 19.27 17.59 7.64 2.71 8.04 4.45 4.08 2.65 0.95 2.45 1.55 1.45 10.67 2.93 9.46 5.58 5.38 1.13 0.57 1.31 0.88 0.82 4.77 4.34 8.12 6.32 5.80 0.80 0.94 1.57 1.36 1.25 1.95 2.88 4.45 4.21 3.89 1.59 2.48 3.60 3.69 3.49 0.20 0.38 0.52 0.57 0.53 4.97 1.99 6.22 3.61 3.20 3.24 0.31 1.30 0.88 0.78 3.19 1.96 5.99 5.43 4.95 5.02 1.69 7.77 3.84 3.26 1.16 0.35 1.70 0.79 0.70 23.97 52.46 34.45 43.50 40.21 16649 5784 17381 9330 8705 296.6 362.8 473.1 389.0 359.6 750.6 282.6 125.1 6.7 5.7 80.5 55.0 41.6 48.2 45.8 554.4 116.3 54.7 28.7 26.1 84.7 151.4 219.4 204.6 190.2 125.3 99.2 127.4 119.2 112.5 Kutch Volcanics - Study Area Main DVP Figure.3.1. Tectonic and structural features of southern Asia and the Indian Ocean basin (based on Mahoney et al., 2002 and Sheth, 2005). Abbreviations for localities are: Br, Barmer, M, Mundwara; D, Dhandhuka; B, Bombay; R, Rajahmundry; G, Goa dykes; KK, Karnataka-Kerala dykes; SMI, St. Mary's Islands volcanics. Mahoney et al., 2002 modeled Re’union hotspot track showing expected ages in Ma. 99 0.5131 0.5130 143Nd/144Nd 0.5129 CIR Re'union Bhuj Ambenali 0.5128 0.5127 Barmer NW-Deccan 0.5126 0.5125 Poladpur 0.5124 Igatpuri-Jawhar 0.5123 0.5122 0.702 Mahabaleshwar Khandala Bushe 0.704 0.706 0.708 0.710 87Sr/86Sr Figure. 3.2. Initial Nd vs. initial Sr isotope plot of the Kutch samples (Bizimis, per. comm.), which are compared with data for Central Indian Ridge (CIR) (bold blue field—data taken from Cohen et al., 1980; Cohen & O’Nions, 1982; Ito et al., 1987; Mahoney et al., 1989; Rehkaemper,1997; Nauret, 2006), Re´union (bold blue field—data from Dupre´ & Alle`gre, 1983; Fisk et al., 1988; Fretzdorff & Haase, 2000), Northwestern Deccan (red field—data from Peng & Mahoney, 1995; ). For the southwestern Deccan Ambenali and Mahabaleshwar data field (green field), Poladpur, Khandala, Igatpuri-Jawhar (black field) taken from Cox & Hawkesworth, 1985; Lightfoot & Hawkesworth, 1988; Peng et al.,1994; Peng et al., 1998. Alkaline picrites and basaltic flows from Bhuj and Barmer is from Simonetti et al., 1998. Kutch alkali basalts are represented by red circles, the two most enriched samples of the alkali basalt are red squares and the tholeiites are orange triangle. 100 3a Sample Normalized to Primitive Mantle (McDonough & Sun, 1995) 1000.000 BH 2 BH 3-1 BH 3-2 100.000 sample/standard BH 17-1 BH 7-1 BH 13-1 BH 16-2 10.000 BH 18-1 N 1-1 BH 12-3 BH 19-3 1.000 N-MORB AVG C ONT. C R US T 0.100 Rb Ba Th 3b U Nb La Ce Pb Pr Sr Nd Zr Sm Eu Ti Dy Yb Y Lu Sample normalized to CI chondrite (McDonough & Sun 95) 200 BH 2 BH 3-1 100 sample/standard BH 3-2 50 BH 17-1 BH 7-1 20 BH 13-1 10 BH 16-2 BH 18-1 5 N 1-1 BH 12-3 2 BH 19-3 La Ce Pr Nd Sm Eu Tb Dy Ho Er Yb Lu Figure 3.3 a) trace element compositions of Kutch basalts normalized to the Primitive Mantle. Red lines are the tholeiites, black lines are alkali basalt and blue lines are the two most enriched alkali basalts. N-MORB (Salters and Stracke, 2004) and Average Continental Crust (Rudnick and Gao, 2003) are also plotted for comparison. b) The REE of the same set of samples are plotted on a CI normalized plot. 101 30.5 Ce/Pb 25.5 20.5 15.5 10.5 5.5 0.5 0.5 1 0.5 2.5 La/Nb 1.5 2 2.5 4.5 6.5 6.5 Sm/Yb 5.5 4.5 3.5 2.5 1.5 0.5 La/Sm 8.5 Main DVP 30. 5 25. 5 20. 5 Figure.3.4. La/Nb vs. Ce/Pb and La/Sm vs. Sm/Yb in the Kutch Basalts. Data source same as Figure.3.2. Symbols for Kutch samples same as Figure. 3.2. 102 Re'union 15. 5 10. 5 CIR 5. 5 0. 5 0. 5 1 1. 5 La / N b 2 2. 5 NW-Deccan 15.9 Bushe 15.8 Archean Basement 207Pb/204Pb 15.7 NW-Deccan Ambenali 15.6 Re'union 15.5 CIR 15.4 15.3 Mahabaleshw ar 15.2 16 16.5 17 17.5 18 18.5 19 19.5 20 20.5 21 206Pb/204Pb 16 15.9 Bhimsankar Igatpuri-Jaw har 15.8 207Pb/204Pb 15.7 Re'union 15.6 15.5 NW-Deccan Ambenali CIR Khandala Pahala 15.4 15.3 15.2 Mahabaleshw ar 15.1 15 0.702 0.704 0.706 0.708 0.710 0.712 87Sr/86Sr Figure. 3.5. Pb and Sr isotopic variation in Kutch Basalts. Symbols same as Figure. 3.2. Fields drawn from same data set as Figure. 3.2. 103 8 Garnet Stability Field (Sm /Y b )n 6 Spinel Stability Field 4 0.1% 0.6% 1.6% 2 0.1% 0.6% 1.6% 0 0 1 2 3 4 5 6 7 8 (La/Sm)n Figure.3.6a. Batch-melting model of garnet peridotite and spinel peridotite. Continuous lines are for percent melting of Fertile Lherzolitic mantle. Mineral Fraction, Partition Coeficient and Melt modes from Salters and Stracke, 2004. Symbols same as Figure 3.2. (Sm /Yb)n 8 6 4 0.5% 1% 2% 5% 2 0 0 1 2 3 4 5 6 (La/Sm)n Figure.3.6b. Incremental Batch-melting model of garnet peridotite. Continuous lines are for percent melting of Fertile Lherzolitic mantle. Mineral Fraction, Partition Coeficient and Melt modes from Salters and Stracke, 2004. Symbols same as Figure 3.2. 104 100.000 BH 3.1 3.7a BH 3.2 BH 17.1 Mixing Line 10.000 1.000 0.100 Rb Ba Th Nb La Ce Pb Sr Zr Ti Y 100.000 BH-2 Mixing Line 3.7b 10.000 1.000 Rb Ba Th Nb La Ce Pb Sr Zr Ti Y Figure. 3.7. Primitive Mantle normalized incompatible element diagram of the Kutch tholeiites compared with mixing different proportion of MORB with charnockites and granulites. a) the pattern is best explained by assimilating 18% basic granulite and 6% charnockite with 75% Indian MORB and fractionating 18% olivine. b) 20% charnockite and 5% granulite is mixed with Indian MORB and 30 % olivine is fractionally crystallized to explain the pattern for the most contaminated sample. Trace element 105 concentration of the granulites and charnockites taken from Weaver, 1978. 0.5135 CIR 0.513 Re'union 0.5125 17 % 24% Granulite & Charnockite 143Nd/144Nd 21 % 0.512 Granite, Tonalite 0.5115 0.511 0.5105 0.51 0.702 0.704 0.706 0.708 0.71 0.712 0.714 0.716 0.718 0.72 87Sr/86Sr Figure. 3.7c. The range of Sr and Nd isotopic composition (at 65 m.y.) is shown for the charnockite, granulite, granite and tonalite from Southern India (Peucat, 1989). The observed isotopic ratios for the Kutch tholeiites can be generated my mixing 17-24% of granulite and charnockite of varying isotopic ratios. Symbols same as Figure. 3.2. 106 APPENDIX A. Analytical Techniques. A.1. Sr and Nd chemical separation and mass spectrometry For Nd and Sr isotopic analysis ca. 50-60 mg of sample powder was weighed in a Savillex screw top beaker and dissolved in 3:1 HF:HNO3 mixture for about 48 hours at about 1000C. To assure complete dissolution of phases like zircon, apatite, sphene etc the samples were dissolved in HF/ HNO3 mixture (3:1) in high-temperature and -pressure Teflon steel-jacketed (Parr) bombs for 48 hrs at 1800C. The bombs are left to cool down in the next 48 hours before opening. The samples are then dried on the hot plate for 24 hrs at 1200C to remove SiF4 and H2O that is produced by the reaction of SiO2 with HF (4HF + SiO2 = SiF4 + H2O). To remove the last drop of acid the temperature is raised to 1800C to 2000C for several hours. After complete evaporation of SiF4 the sample stopped fuming. When the beaker is completely free from condensing acid droplets, the temperature is lowered to 1250C and 5-6 ml of 6N HCl is added to the dry sample cake, re-dissolved and left to dry. This step is repeated 3 times to ensure the removal of remaining fluorides. In the next step, the sample is re-dissolved in 0.5 ml of 2.5N HCl. The solution with some of the un-dissolved residue is transferred into centrifuge tube and centrifuged for 6 – 9 minutes till a clear solution is achieved. This solution is now ready to load on the cation columns to separate Sr from the REE’s. Half of the sample solution (0.25 ml) in 2.5N HCl is loaded on quartz glass cation columns filled with approximately 3.7 cm3cation resin. The bulk sample is eluted with approximately 18 ml of 2.5 N HCl. Sr is collected in 5 ml of 2.5N of HCl. After the Sr is collected 4 ml of 6N HCl is passed through the columns and the rare earth elements are collected in another 6 ml 6N HCl. Resin material may leak through the columns. To ensure complete decomposition of resin material Sr fractions are dried down with 3 drops of HClO4 twice. Then 3 drops of concentrated nitric acid is added and dried to ensure complete nitrate conversion. The process is repeated thrice. The Sr fraction is loaded on single W filaments as a tiny drop of 1 microliter 0.25N HNO3 mixed with equal amounts of the TAPH (10g H2O + 0.5g TaCl5 + 3g 0.1N H3PO4 + 0.5g conc. HF) solution activator. It has been well tested that the TAPH solution enhances ionization of Sr as well as binds it well on the filament to retain the sample long enough for the measurement. Sr measurements were performed on a Finnigan 262 multicollector mass spectrometer. The measurements were performed in the dynamic mode with one peak switch. The signal intensities were obtained after online fractionation correction based on the power law and the 86Sr/88Sr ratio of 0.11940. 107 The signal intensities for 88Sr were always greater than 1 Volt. The granites and gneisses had some Rb leaking into the Sr fraction even after column separation. 87Rb signal interferes with 87Sr .Care was taken so that 85Rb signal was less than 300 counts per second (cps), to avoid signal interferences. Corrections to the 87Sr signal required for 87 Rb interference can be neglected when 85Rb signal is less than 300 cps. In most cases repeated “flashing” of the filament at high temperatures for a few seconds burns off the excess Rb. The reproducibility of the measurements is controlled by replicate measurements of the E & A Sr isotope standard. The average measured 87Sr/86Sr ratio was 0.708006 (0.000019, two standard deviations, n = 10). At least 10 blocks (100 ratios) were taken for each measurement and it was ensured that the internal precision is always better than the external reproducibility of the standard (i.e. less than 0.000019, 2 sigma error). Nd was collected in a bulk REE fraction from the cation columns in 6N HCl. Nd and the REE fraction was further separated on a 1.2 ml, 6 cm bed-length column of Ln resin SPS. After collection, the Nd fraction is dried and is ready for measurement. The Nd fraction is dissolved in ca. 1 microliter 0.25N HCl, mixed with equal amounts of 0.25N H3PO4 and loaded as a single drop on a zone-refined Re filament. The Nd measurements were performed using the double Re filament method. Potential interferences on Nd are Ba, La, Pr and Ce and their oxides. They arise from LREE and LILE enriched rocks where Nd separation was incomplete. The most common interfering mass on Nd is 130Ba16O at mass 146. To control this, 138Ba (the most abundant Ba isotope) and its oxide at mass 154 was monitored. If the interference is high, the sample filament is carefully “flashed” at higher temperatures until Ba is burned off. The presence of La, Pr, Ce does not produce interferences but decreases the ionization efficiency of Nd. Since the samples studied here had high Nd concentrations this was not a problem. Nd measurements were performed on a Finnigan 262 multicollector mass spectrometer. The measurements were performed in the dynamic mode with one peak jump. The signal intensities were obtained after online fractionation correction based on the power law and the 146Nd/144Nd ratio of 0.721903. External reproducibility was measured on the La Jolla Nd standard. During the period of analysis the average was 143Nd/144Nd = 0.511846+ 0.000016 (2 s.d., n = 20). For the Nd isotope compositions, the initial ε values were calculated using the formula: ⎡ (143Nd / 144Nd ) TS ⎤ − 1⎥ * 1000, were the ratio with the superscript T is the initial T 143 144 ⎣ ( Nd / Nd ) CH ⎦ ratio at the time of eruption, and the subscript CH is the composition of the chondritic earth at time T. T ε Nd =⎢ Sr and Nd blanks were measured in ICP-MS. An empty beaker was treated like a sample and processed along with the samples. In the subsequent step, a droplet of 0.25N HCl was added to the beaker in the loading bench as was done or the samples. This droplet was 108 dried and picked up with 3 ml 1% HNO3 and transferred to a clean vial for the ICP measurement. At this point all the Sr and Nd present in the beaker representing both the procedural and reagent blank is in the vial. In order to determine the machine background the 2% wash solution is used. The blank solution was compared against standard solutions of 1 and 10 ppt Nd and 100 ppt Sr. The machine background intensity was around 3% of the Nd intensity in the blank. Thus the blank intensity was sufficiently higher than the machine background. During the blank measurements it was ensured that the machine background stays sufficiently low (<10 cps on 146Nd was the primary monitor). The two blanks determined this way were 9.5 pg and 11.8 pg. This blank contribution is trivial with respect to the Nd concentrations used here (always several hundred nanograms). For Sr the two blanks measured were 120 and 130 pg. A.2. Pb chemical separation and mass spectrometry For Pb isotopic determination approximately 100 mg sample were dissolved in HF/HNO3 (3:1). Samples were put on a hot plate in a clean air flow box for at least 24h at 1000C to complete the dissolution. They are then opened and let dry at that temperature approximately 1000C to evaporate the HF. Lead chemistry was performed following the techniques described in Abouchami et al., 1999. The samples were then dried three times in 1 ml of concentrated HNO3, two times in 6N HCl and two times in 1 ml of 9 N HBr. After dissolution and drying, samples are picked with 0.5 to 1 ml of Solution A (7 ml H2O + 1 ml 2N HBr + 2 ml 2.5N HNO3) and loaded in 100μl Teflon columns filled with anion resin. The bulk of the sample is eluted with 0.5 ml of Solution A and 0.2 ml of Solution B (7.85 ml H2O + 0.15 ml 2N HBr + 2 ml 2.5N HNO3) and Pb is collected with 0.7 ml Solution B. This process is repeated to ensure a pure Pb fraction. Pb isotope measurements were performed on the Finnigan 262 MS. Samples were picked up with 0.2N HBr, mixed with silica gel and H3PO4 and loaded on commercial grade single Re filaments. Since Pb does not have a pair of non-radiogenic isotopes, no online fractionation correction can be performed. To establish the range of fractionation and correct for it, several measurements were performed on the NBS 981 Pb standard. To be able to accurately compare the fractionation observed on the standard to that of the samples, both standards and samples were run at the same temperatures (1230 to 12500C) and at the same load (50-75 ng of Pb). The technique involved a fully automated heat up and run in order to further control fractionation variations arising from different heating up and running procedures between standards and samples. The results for the NBS 981 Pd standard were (n = 14): 208 Pb/204Pb = 36.523(+0.023 1 standard deviation) 207 Pb/204Pb = 15.435(+ 0.007) 206 Pb/204Pb = 16.894(+ 0.005) These values result in a fractionation correction (Todt, et. al., 1993) of 0.12% per amu determined from analysis of 14 standards. 109 Pb blanks were measured with the ICP-MS. An empty beaker was treated as a sample and processed along with the samples. At the end, a droplet of 0.2N HBr was added to the beaker in the loading bench as was done or the samples. This droplet was dried and picked up with 3 ml 1% HNO3 and transferred to a clean vial for the ICP measurement. The two measured Pb blanks were 0.3 ng and 0.27 ng. A.3. ICP-MS analyses 30 mg of sample powder was weighed in a Savillex beakers and dissolved in 3:1 HF:HNO3 mixture for about 48 hours at about 800C. After 2 days the samples are evaporated to dryness at about 1000C, re-dissolved in 7N HNO3 and left in an ultrasonic bath for 30 minutes. At this stage most samples were completely dissolved. In case of incomplete dissolution samples were heated to 175-2000C for several hours and placed in an ultrasonic bath for ~10 minute. This process was repeated several times. This assured complete dissolution of all samples. Samples were then transferred to clean HDPE bottles (pre-cleaned with 1:1 HNO3 for several days), diluted with 2% HNO3 for a total dissolved solid content of 100 ppm. Inductively coupled plasma mass spectrometry (ICP-MS) (Finnigan MAT ELEMENT high-resolution ICP-MS) at the NHMFL was used to determine the concentration of 32 trace elements. CD1E interface (“guard electrode” or platinum shield between the load coil and torch) was used to attain a sensitivity of about 700,000 cps/ppb 115In for sample gas and auxiliary gas flow rates of about 1 l/min. A Teflon spray chamber and Teflon tubing was used along with a Teflon nebulizer from Elemental Scientific, the later was set at a self aspirating mode with a flow rate of 100 μl/min. The washout characteristics improved significantly by use of Teflon for the inlet system compared to a conventional setup consisting of a combination of Tygon tubing and glass nebulizer and spray chamber. Replacing glass by Teflon allowed the use of a mixture of 2% HNO3-0.03N HF as cleaning solution in between sample runs. The background for “sticky” elements such as Nb, Zr, Hf, Pb, Th, and U dropped significantly by a factor of 20-200 after using HF in the cleaning solution and generally improved the washout characteristics. The blank levels for most elements were <10-100 cps. A 90-second sample uptake time was chosen in combination with a washout time of 120 seconds between samples. Samples were measured in small sequences, consisting of 5 unknowns (4 samples and 1 blank) bracketed by two rock standards, (in this case BIR-1 for mafic rock samples and G2 for felsic rock samples). Thus the analysis time was kept short and the effects of signal drift were considerably reduced. Signal drift was usually <5% between the first and last sample of each sequence. Signal drift was corrected based on an internal standardization with Indium (all samples and standards are spiked with In to a concentration of 1 ppb) and a linear interpolation between the external rock standards for each analyzed isotope. The reference concentrations of BIR-1 are Jochum et al., 2005. To check for accuracy and precision, repeat measurements were made on one unknown, and on the rock standard BHVO-I. Based on repeat measurements of the unknown samples and the rock standard precision for most elements is about 2% (1σ, Table A1). 110 Reproducibility for the rare earth elements (REE) is generally better than 2%. For Nb, Ta, Zr, Y, Hf, Rb, Sr, and Th, reproducibility is generally better than 4-5%, and Sc, Ti, V, Cr, Co, Ni and U reproduced to better than 5-7% (Table A1). The average concentration of BHVO-I agrees well with the values given by Jochum et al. (2005) (Table A1). Using CD1E interface has both advantages and disadvantages. The CD1E interface improves sensitivity about five times and thus sample solutions with low total dissolved solid content can be used which in turn improves the signal drift to a considerable degree (about four times less signal drift with a 250 ppm solution compared to 500-1000 ppm solutions over the courses of a two hour measurement). The disadvantage of using CD1E interface is that it can increase oxide formation up to two times. During the course of analysis the major oxide interferences on some of the rare earth elements (REE) such as Eu, Gd and Tb were135BaO on 151Eu, 141PrO on 157Gd, 143NdO on 159Tb. The problem of oxide formation was dealt in the following ways: a) oxide interferences on samples enriched in heavy rare earth elements (HREE) are usually insignificant up to oxide formation levels of several percent, b) oxide formation is approximately constant between samples with similar matrices. Thus using matrix-matched external standard with similar Ba, Pr and Nd concentrations compared to the analyzed samples (as in case of BIR-1 and amphibolites), oxide formation is to a large degree cancelled out by applying the external rock standard correction. However, especially where samples are more enriched in light REE and very incompatible elements than the external rock standard, the effects of oxide formation can be significant, as, for example, between BIR-1 and a LREE enriched rock such as BHVO-I. Table 1A compares our results for the BHVO-I standard with those measured by Jochum et al. (2005). Though majority of the elements measured in BHVOI agree to within the analytical uncertainty of the values of Jochum et al. (2005), measured Gd concentrations in BHVO-I are significantly higher. A.4. LA-ICP-MS Single Zircon Analysis (method description provided by Dr. Victor Valencia at University of Arizona, Tuscon) Samples were collected from GFG, MRG and a thin meta-sandstone unit interlayered with the PCF amphibolites. At each sample locality, we collected 1–2 kg of fresh whole rock. Samples were prepared for analysis using standard crushing and separation techniques, including heavy liquids and magnetic separation. Apparently inclusion-free zircons were then hand-picked under a binocular microscope. At least 100 zircons from each sample were mounted in epoxy and polished. Single zircon crystals were analyzed in polished sections with a Micromass Isoprobe multicollector ICP-MS equipped with nine Faraday collectors, an axial Daly detector, and four ion-counting channels (Kidder et al., 2003) at the University of Arizona, Tucson. The LA-ICP-MS analyses involve ablation of zircon with a New Wave DUV193 Excimer laser (operating at a wavelength of 193 nm) using a spot diameter of 35 to 50 microns. The ablated material is carried in argon gas into the plasma source of a Micromass Isoprobe, which is equipped with a flight tube of sufficient width that U, Th, and Pb isotopes are measured simultaneously. All measurements are made in static mode, using Faraday detectors for 238U, 232Th, 208Pb–206Pb, and an ion-counting channel 111 for 204Pb. Ion yields are 1 mV per ppm. Each analysis consists of one 20-second integration on peaks with the laser off (for backgrounds), 20 one-second integrations with the laser firing, and a 30-second delay to purge the previous sample and prepare for the next analysis. The ablation pit is 20 microns in depth. Common Pb (initial-age corrected) correction is performed by using the measured 204Pb and assuming an initial Pb composition from Stacey and Kramers (1975) (with uncertainties of 1.0 for 206Pb/204Pb and 0.3 for 207Pb/204Pb). Measurement of 204Pb is unaffected by the presence of 204Hg because backgrounds are measured on peaks (thereby subtracting any background 204Hg and 204Pb), and because very little Hg is present in the argon gas. For each analysis, the errors in determining 206Pb/238U and 206Pb/204Pb result in a measurement error of several percent (at 2-sigma level) in the 206Pb/238U age. The errors in measurement of 206Pb/207Pb are substantially larger for younger grains due to low intensity of the 207Pb signal. The 207Pb/235U and 206Pb/207Pb ages for younger grains accordingly have large uncertainties. Interelement fractionation of Pb/U is generally <20%, whereas isotopic fractionation of Pb is generally <5%. In-run analysis of fragments of a large zircon crystal from a Sri Lanka pegmatite (e.g. Dickinson and Gehrels, 2003) with known age of 564 ± 4 Ma (2-sigma error) is used to correct for this fractionation (generally run every third measurement). The uncertainty resulting from the calibration correction is generally 3% (2-sigma) for both 207Pb/206Pb and 206Pb/238U ages. The crystallization ages reported in this thesis are weighted averages of individual spot analyses. The Paleozoic ages interpreted from the ICP-MS analyses are based on 206 Pb/238U ratios because errors of the 207Pb/235U and 206Pb/207Pb ratios are significantly greater. This is due primarily to the low intensity (commonly <1 mV) of the 207Pb signal from these young, U-poor grains. For grains older than 800 Ma, both 207Pb/206Pb and 206 Pb/238U are reported. A.5. Rb/Sr ratios measured in ICP-MS All measurements for Galts Ferry Gneiss samples with Rb/Sr ratio less than 5 were performed on a Finnigan Element 1. Sample preparations were carried out in a Class 300 clean laboratory at the NHMFL. Gravimetric standards were prepared from commercial Rb and Sr 10.00±0.05 µg/ml stock solutions (High-Purity Standards™). The error introduced by mixing and dilution of the stock solutions was trivial compared to the uncertainty on the reported accuracy of the gravimetric solutions (±0.5%). Rock samples were digested in Savillex™ PFA beakers using PFA-distilled HF-HNO3 acids. The HF was dried down and the residual salts taken up in concentrated HNO3. These solutions were diluted in quartz-distilled ultrapure water to 2% HNO3 solutions, stored in acid-washed 125ml LDPE bottles. Solutions were introduced into the ICP source using an ESI™ low-flow 100 µl/min PFA nebulizer with an ESI™ PFA spray chamber. A CETAC ASX-100 autosampler was used 112 for automated sample handling. Analyte concentrations were set at levels sufficient to provide count rates >107 cps that were measured in Analog mode on the Element1 detection system. An analysis of a blank solution (2% HNO3) was performed to correct the Analog background and to subtract off any memory contributions and 86Kr+ interferences (<104 cps). The magnet was settled at mass 85, and the peaks of interest (m/e= 85, 86, 87, and 88) were acquired using the EScan option by scanning the Element’s accelerating voltage in Low Resolution mode (R=300). An isobaric interference of 87Rb on 87Sr limited the value of that peak. Ratios were calculated for 85 Rb/86Sr, 85Rb/88Sr and 86Sr/88Sr. Measured ion intensity ratios are plotted in Figure A1 against the gravimetric values of the standard solutions. A calibration curve is obtained that converts measured ion intensity ratios to concentration ratios. Figure A1 shows the calibration curves obtained for a representative set of measurements of the gravimetric standards for both 85Rb/86Sr and 85Rb/88Sr. Both ratios provide excellent measures of the Rb/Sr concentration ratio of the solution, so the final reported Rb/Sr is an average of the two measurements. To appreciate the precision obtained, Figure A1 also shows the % difference defined as, ⎡ ( Rb / Sr ) measured ⎤ %difference = ⎢ − 1⎥ * 100, ⎣⎢ ( Rb / Sr ) gravimetric ⎦⎥ as a function of the gravimetric value for Rb/Sr ratios derived from 85Rb/86Sr and 85 Rb/88Sr ratios. Precisions on some average ratios approach ±0.2%, and precision was always better than ±0.4%. Further, precision of the 86Sr/88Sr was about ±0.4%, as well. A.6. Rb/Sr ratio measurement by isotope dilution. The Mulberry Rock Gneiss samples had a huge variation of Rb/Sr ratio and thus standard isotope dilution technique was used to determine the 87Rb/86Sr ratios. Whole rock samples were cut into small pieces with a rock saw to remove any traces of weathered surface. The piece were washed in deionized water in a sonic bath for about 15 minutes, rinsed in methanol and air dried. The pieces were pulverized in a Siebtecknic rotary grinder fitted with a tungsten carbide mortar and approximately 100 mg of crushed samples weighed in teflon screw cap beakers and spiked with a mixed spike of 87Rb - 84Sr and dissolved in distilled 3:1 HF:HNO3 and similar chemical procedure was followed as described in A1. The sample solution was then loaded into 12 ml glass column with Dowex AG50W X-8 (200-400) cation exchange resin and Rb and Sr were eluted in 2.5 N HCl. Sr measurement was performed on a Finnigan MAT262 Thermal Ionization Mass Spectrometer following the same procedure described in A1. Rb was measured on the same instrument between 900 and 1000 degrees centigrade. Three blank analysis of Sr yielded 0.12 nanograms, 0.13 nanograms, 0.196 nanograms of Sr and two Rb bank yielded 0.005 microgram and 0.002 microgram of Rb. These values are considered insignificant for all analysis. In Isoplot 3 (Ludwig, 2001) program used for age calculation the age data experimental errors of 0.01% for (87Sr/86Sr)N ratios and 1% for 87 Rb/86Sr ratios were used. Age calculation were made using a 87Rb decay constant of 1.42*10-11 year-1. 113 A.7. Rb/Sr ratio measurement by LA-ICP-MS Feldspar grains with well preserved crystal faces were hand picked under a binocular microscope. Preferably unaltered orthoclase feldspar grains were picked (identified by their flesh pink color) so that the major oxide specially the SiO2 concentration can be estimated for internal standard. The feldspar grains were washed in water and methanol, mounted in epoxy and left overnight for the epoxy to dry. The epoxy mount was then polished in 240-321-400-600 grit sand papers and finally in 1 micron diamond polish. Polished surfaces were analyzed for major and trace elements by laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) at the NHMFL Plasma Analytical Facility. A New Wave UP213 laser ablation system (213 nm UV) connected to a Finnigan Element™ ICP mass spectrometer was used for the measurements. The sample chamber was flushed with helium gas at 800 ml/min, and additional argon makeup gas (980 ml/min) was teed into the sample line to the ICP torch. The sample surface was scanned in raster mode at a rate of 5 µm/s using a 20 µm diameter beam, with 10 Hz repetition rate and 50% power output. The mass spectrometer recorded intensities from the ablated material at the mass values of interest in a series of 275 mass sweeps. Each line scan was 1 mm in length. All of the measured isotopes (7Li, 23 Na, 25Mg, 27Al, 29Si, 39K, 43Ca, 49Ti, 53Cr, 55Mn, 57Fe, 85Rb, 88Sr, 133Cs, 138Ba and 208Pb) were recorded in the same set of analyses. Blank subtractions (average of 2 measurements with the laser off) were made. Following blank subtraction and correction by the instrumental sensitivity factors determined on MPI-DING (ATHO-G) reference glass, the elemental ratios to SiO2 were converted to concentrations assuming SiO2=70% in orthoclase, then normalized to 100%. Detection limits for each analysis were calculated from the 3σ uncertainty of blanks run during the analytical session. Standards used to convert intensities to concentrations included the MPI-DING (ATHOG) reference glass (Jochum et al. 2006). Table A2 gives elemental concentrations for the standards used in the calibration. Relative sensitivity factors (RSF) ⎡C ⎤ element / C SiO2 ⎢ ⎥ RSF = ⎢ I element / I Si 29 ⎥ ⎣ ⎦ for each of the elements analyzed were calculated from the concentrations given in Table A2 and from measured background-subtracted intensity ratios. Background-subtracted intensity ratios were then converted to elemental ratios (normalized to SiO2) and multiplyed by the RSFs. For the Feldspar major elements were calculated from oxide ratios of Si, Ti, Al, Fe, Mg, Ca, Na, K and Mn normalized to 100%, neglecting minor Cr. 114 Table A1: Repeat analysis of one felsic sample (D2) and USGS standard BHVO1 run as an unknown is compared with the published concentrations from Jochum et al. (2005). D2 Li Rb Sr Y Zr Nb Cs Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Yb Lu Hf Ta Pb Th U Ti V Cr Co Ni Cu Zn Ga D2-rerun conc (ppm) conc (ppm) 6.79 7.02 127 123 181 182 14.03 14.54 537 564 8.59 8.88 2.94 3.03 478 479 25.98 26.18 53.64 53.96 5.76 5.92 20.04 20.63 3.53 3.55 0.59 0.58 1.48 1.41 0.28 0.30 2.51 2.59 0.50 0.52 1.50 1.55 1.25 1.28 0.23 0.23 16.45 17.02 0.73 0.77 10.04 10.34 11.78 12.12 2.67 2.72 1138 1182 31.70 33.71 130 123 5.87 5.70 28.40 26.77 56.48 53.92 19.42 20.31 3.41 3.55 % difference 3.23% -3.15% 0.67% 3.56% 4.73% 3.33% 2.95% 0.26% 0.77% 0.59% 2.73% 2.89% 0.51% -0.68% -4.88% 5.04% 3.10% 2.73% 3.33% 2.40% 1.67% 3.33% 4.90% 2.84% 2.80% 1.71% 3.70% 5.98% -5.75% -3.00% -6.10% -4.75% 4.40% 3.95% 115 BHVO1 BHVO1 measured published conc (ppm) conc (ppm) 9.02 384 26.34 168 18.06 0.09 129 15.41 37.46 5.27 24.22 5.90 1.96 5.94 0.93 5.33 0.98 2.52 2.06 0.28 4.11 1.11 2.09 1.16 0.40 8.26 395 26.4 174 17.7 0.102 132 15.4 37.9 5.3 24.7 6.13 2.1 6.39 0.959 5.38 0.969 2.54 1.99 0.271 4.26 1.17 2.13 1.23 0.409 % difference 8.42% -2.97% -0.23% -3.54% 2.00% -8.63% -2.65% 0.08% -1.19% -0.65% -1.98% -3.88% -7.16% -7.54% -3.12% -1.03% 0.66% -0.90% 3.39% 3.21% -3.63% -5.41% -1.71% -6.30% -1.56% Table A2: Concentrations and relative sensitivity factors (RSFs) for elements analyzed in this study. Major element oxides in wt. %, trace elements in ppm. Concentrations from Jochum et al. (2006). Concentration Sensitivity Factor SiO2 (%) TiO2 Al2O3 FeO MgO CaO K2O Na2O MnO Cr (ppm) Li Rb Sr Cs Ba Pb 75.6 0.255 12.2 3.27 0.103 1.7 2.64 3.75 0.106 6.1 28.6 65.3 94.1 1.08 547 5.67 1 5.31 34.4 2.92 4.09 0.1450 0.0239 131 115 0.0377 0.0101 0.0201 0.0149 0.0295 0.0138 0.0128 Table A3: End member composition of the MORB and crust used for mixing calculation to produce the Deccan tholeiites. Charnockites Basic Granulite Basic Granulite Intermediate Granulite Weaver et al., 1978 Weaver et al., 1978 Weaver et al., 1978 Rb Ba Th Nb La Ce Pb Sr Zr Ti Y 100 879 19 18 36 67 11 166 277 2218.15 23 17 400 3 13 28 51 7 210 99 5995 41 N-MORB Weaver et al., 1978 Hofmann, 1988 4 5 77 3 4 9 19 5 113 56 5575.35 20 320 7 13 39 71 9 180 158 9412.15 30 1.26 13.87 0.19 3.51 3.90 12.00 0.49 113.20 104.24 9681.93 35.82 116 16.00 14.00 y = 7.297245x + 0.011180 R2 = 0.999998 85Rb/86Sr 12.00 10.00 8.00 6.00 4.00 2.00 0.00 0.00 0.50 1.00 1.50 2.00 2.50 Rb/Sr gravim etric Figure A1. Measured ion intensity ratios are plotted against the gravimetric values of the standard solutions (top two diagrams). Percent difference as a function of the gravimetric value for Rb/Sr ratios derived from 85Rb/86Sr and 85 Rb/88Sr ratios (bottom diagram). 2.00 1.80 y = 0.901111x + 0.000924 R2 = 0.999995 1.60 85Rb/86Sr 1.40 1.20 1.00 0.80 0.60 0.40 0.20 0.00 0.00 0.50 1.00 1.50 2.00 2.50 2.0 2.5 Rb/Sr gravim etric 0.50% 0.40% % diff 85Rb/86Sr 0.30% % diff 85Rb/88Sr % difference 0.20% 0.10% 0.00% -0.10% 0.0 0.5 1.0 1.5 -0.20% -0.30% -0.40% -0.50% Rb/Sr grav ime tric 117 SAMPLE LOCATIONS Sample UTM_X UTM_Y RA1021 569864 3608902 RA422A 565758 3642576 RA1034 574925 3602673 RA536 548477 3632004 RA401 558126 3643810 RA701B RA587 RA1005 562161 3645773 565771 3648668 548889 3637505 RA550 557502 3637769 RA1029 RA677 SM 639 570609 3606141 585733 3661572 T18S, R10E, SE1/4, NE1/4, Sec 11 T20S, R6E, NE1/4 NE1/4 Sec 9 T20S, R6E, NE1/4 NE1/4 Sec 30 T22N, R14E, SE1/4 SW1/4 Sec 11 T20S, R6E, NE1/4 NE1/4 Sec 9 T20S, R6E, NW1/4 NE1/4 Sec21 T23N, R14E, SE1/4 SE1/4 Sec 34 T19S, R7E, SE1/4 SE1/4 Sec 24 T20S, R6E, NW1/4 NW1/4 Sec 28 SM 1100 M 68 M 243 M 268 M 271 M 272 M 201B M 67 D1 D2 BG 995 BG 999 Y23C Y24 Y131 NT911 BH166 591872 589674 591868 589664 694526 694602 694507 694405 702698 3676201 3678330 3676204 3678337 3763654 3762308 3763688 3763844 3770971 SC318G TA7 US41A Y126 738617 726207 708751 695269 3758025 3733565 3774172 3762310 Quadrangle Elmore Flag Mt. Wetumpka Mitchell Dam Mitchell Dam NW Flag Mt. Flag Mt. Mitchell Dam NW Mitchell Dam NW Wetumpka Goodwater Micaville Unit Emuckfaw-Heard Higgins Ferry Emuckfaw-Heard Higgins Ferry Higgins Ferry Description Garnet quartzite Garnet graphite quartzite Metagreywacke Graphitic quartzite Graphitic quartzite Higgins Ferry Higgins Ferry Higgins Ferry Metagreywacke Metagreywacke Ga+Mu+Bi+Pl schist Higgins Ferry Garnet quartzite Emuckfaw-Heard Higgins Ferry Metagreywacke Graphitic quartzite Kyanite-garnet schist Porters Gap Hillabee Greenstone Massive greenstone Porters Gap Jamison East Hillabee Greenstone Massive greenstone w/sulfides and diabasic texture Hillabee Greenstone Greenstone/mafic phyllite Porters Gap Hillabee Greenstone Greenstone Porters Gap Hillabee Greenstone Greenstone Jamison East Hillabee Greenstone Foliated greenstone/mafic phyllite Foliated greenstone Lineville West Hillabee Greenstone Porters Gap Hillabee Dacite Metadacite Bulls Gap Bulls Gap Bulls Gap Bulls Gap Yorkville Yorkville Yorkville Hillabee Dacite Hillabee Dacite Hillabee Dacite Hillabee Dacite Galts Ferry Gneiss Galts Ferry Gneiss Galts Ferry Gneiss Galts Ferry Gneiss Galts Ferry Gneiss Metadacite Metadacite Metadacite Metadacite Trondhjemite Trondhjemite Trondhjemite Trondhjemite Trondhjemite Galts Ferry Gneiss Galts Ferry Gneiss Galts Ferry Gneiss Pumpkinvine Amphibolite Pumpkinvine Amphibolite Pumpkinvine Amphibolite Pumpkinvine Amphibolite Trondhjemite Trondhjemite Trondhjemite Amphibolite Burnt Hickory Ridge South Canton Taylorsville Acworth Yorkville Y172 693064 3756302 Yorkville TA16 722287 3734143 Taylorsville TA17 722104 3739266 Taylorsville 118 Amphibolite Amphibolite Amphibolite Sample UTM_X UTM_Y Quadrangle SC254 739757 3758454 South Canton SC318A 738617 3758025 AC78 709193 3774914 BH57 707261 3773328 SC 25 Sc 33 DR 136 NG 057 MRsamples 685058 3737238 690404 3738984 691697 3746739 692162 3744326 TPGT 1 TPGT 2 TPGT 3 TPGT 4 TPGT 6,7,8,9 TPGT 1115 413293 4056664 413583 4056310 412209 4057522 412512 460952 412512 460952 404378 4052232 Unit Description Pumpkinvine Amphibolite Amphibolite South Canton Pumpkinvine Amphibolite Amphibolite Acworth Pumpkinvine Amphibolite Amphibolite Burnt Hickory Pumpkinvine Amphibolite Ridge Amphibolite Canton Schist Garnet muscovite schist Canton Schist Garnet muscovite schist Drakestown Canton Schist Garnet muscovite schist New Georgia Canton Schist Garnet muscovite schist New Georgia Mulberry Rock Granite samples collceted from Gneiss quarry New Georgia Mulberry Rock Granite samples collceted from Gneiss quarry Treas Piedras Tres Piedras Granite Granite Treas Piedras Tres Piedras Granite Granite Treas Piedras Tres Piedras Granite Granite Treas Piedras Tres Piedras Granite Granite Treas Piedras Tres Piedras Granite Granite Las Tablas 119 Tres Piedras Granite Granite 100 μ 100 μ Top: Perthites in Tres Piedras Granite samples. 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Yang, H.J., Frey, F.A., and Clague, D.A., 2003, Constraints on the source components of lavas forming the Hawaiian North Arch and Honolulu volcanics: Journal of Petrology, v. 44, p. 603-627. 148 BIOGRAPHICAL SKETCH Reshmi Das joined the Department of Geological Sciences, Florida State University (GS, FSU) in Fall 2001. Prior to that she completed her Bachelor of Science and Master of Science programs at the University of Calcutta, India. She has presented papers at various conferences organized by the Geological Society of America (GSA) and the American Geophysical Union (AGU). Reshmi received several grants from Educational Mapping Program (EDMAP), GSA, FSU-Congress of Graduate Students (COGS) etc. Her research interests include geochemistry, structural geology, and petrology. Reshmi also served as the instructor for the Dynamic Earth course at GS, FSU for several terms and was nominated for the FSU Outstanding Teaching Assistant Award 2005. As a part of her doctoral dissertation Reshmi did three projects on tectonic evolution of southern Appalachian, isotope homogenization of strontium in northern New Mexico, and geochemical nature of the first pulses of Deccan lava in India. Reshmi now joins the National High Magnetic Field Laboratory (NHMFL), Tallahassee, Florida as a postdoctoral fellow. Reshmi loves to travel, paint, and read. 149