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Florida State University Libraries Electronic Theses, Treatises and Dissertations The Graduate School 2009 Superposed Fault Systems of the Southernmost Appalachian Talladega Belt: Implications for Paleozoic Orogenesis in the Southern Appalachians Clinton Ivan Barineau Follow this and additional works at the FSU Digital Library. For more information, please contact [email protected] FLORIDA STATE UNIVERSITY COLLEGE OF ARTS AND SCIENCES SUPERPOSED FAULT SYSTEMS OF THE SOUTHERNMOST APPALACHIAN TALLADEGA BELT: IMPLICATIONS FOR PALEOZOIC OROGENESIS IN THE SOUTHERN APPALACHIANS By CLINTON IVAN BARINEAU A Dissertation submitted to the Department of Geological Sciences in partial fulfillment of the requirements for the degree of Doctor of Philosophy Degree Awarded: Summer Semester, 2009 The members of the committee approve the dissertation of Clinton I. Barineau defended on June 25th, 2009. __________________________________ James F. Tull Professor Directing Dissertation __________________________________ George W. Bates Outside Committee Member __________________________________ A. Leroy Odom Committee Member __________________________________ Stephen A. Kish Committee Member Approved: _____________________________________ A. Leroy Odom, Chair, Geological Sciences ____________________________________ Joseph Travis, Dean, Arts and Sciences The Graduate School has verified and approved the above-named committee members. ii To Diedre, Lauren, and Zach, for their love and never ending encouragement. iii ACKNOWLEDGEMENTS I owe an incalculable debt of gratitude to a huge group of people, without whom I could have never completed this document and my graduate work. First, I am indebted to my advisor, Jim Tull, who spent countless hours patiently guiding me through my academic efforts over the past 15 years. I count him as much a friend as I do an advisor. Jim has been an excellent source of career advice and I am indebted to him for his efforts in helping me procure a faculty position. His enthusiasm towards field geology, in general, and the southern Appalachians, specifically, has inspired me over the years and I am lucky to have learned my trade from someone with his knowledge and expertise. If my career as a geologist is half as successful as his, I shall count myself fortunate. I have interacted with a number of faculty members at FSU during the course of my undergraduate and graduate studies, all who have provided me with valuable assistance and insights into my academic pursuits. Thanks to Roy Odom for his knowledge of southern Appalachian geology and willingness to share his thoughts and opinions with me. I am also grateful for his time and assistance at the National High Magnetic Field Laboratory during preparation of mineral samples for isotopic analysis. Neil Lundberg has always been willing to share his expertise on the Taiwan collisional orogen, especially the sedimentological record, and his feedback on these aspects of my dissertation has proved invaluable. Steve Kish has proven a wealth of information on the metamorphic history of the region and I am indebted to him for his thoughts and opinions on metamorphism and metamorphic ages in the southern Appalachian orogenic belt. Jennifer Georgen supplied a wealth of knowledge on tectonic processes and geophysics, and her feedback on subduction initiation processes was greatly appreciated in preparing this manuscript. Bill Parker was always willing to share his knowledge on sedimentation processes and I have benefited much over the years from his tutelage. Tapas Bhattacharyya’s knowledge of Himalayan tectonics was extremely valuable and his willingness to act as a sounding board for my ideas is greatly appreciated. I am especially indebted to George Bates, Department of Biological Sciences, who agreed on short notice to act as my outside committee member and who iv has provided me with countless hours of challenge on the racquetball court (usually trouncing me soundly) during my intermittent breaks from writing. Lastly, I would like to thank the collective faculty in the Department of Geological Sciences for providing me with a superb education. Over the past 15 years you have inspired me to endeavor to become both a quality researcher and teacher. I have immensely enjoyed my interactions with a number of FSU staff members, graduate students, and alumni over the years, and I have greatly benefited from their knowledge and friendship. Tami Karl and Mary Gilmore have always bent over backwards to help me with the bureaucratic maze that makes up a university. My discussions with Li Li, Stephen Palmes, Chul Lim, Lance Johnson, Jiun-Yee Yen, Martin Balinsky, Deb Hilton, Haitham Baggazi, Pallov Pal, Stephen Whiting, and many other students in the structure lab have taught me much during my graduate career. I am grateful to Reshmi Das for her help with mineral separation techniques and isotopic geochemistry and Ted Zateslo for his tireless efforts to solve computer and other technical problems. I am especially grateful to three FSU Geology alumni: Chris HolmDenoma, who shared an office with me and spurred much of my thinking on the Taconic orogeny; Mark Groszos, my colleague at Valdosta State University who shared his knowledge of the southern Appalachians during our daily commute between Tallahassee and Valdosta; and David Allison at the University of South Alabama, who has been an endless source of information regarding the geology of the AshlandWedowee belt, database management, and the construction of digital maps. I would also like to extend my thanks to the Department of Geological Sciences, which provided partial funding of my graduate studies through teaching and research assistantships for a number of years. Additionally, a significant percentage of my field work was funded through grants from the U.S. Geological Survey Educational National Mapping Cooperative, the National Science Foundation, Sigma Xi, the Robert C. Lang III Memorial Scholarship, and FSU Department of Geological Sciences Banks Award. Additionally, I have been fortunate to have the opportunity to teach at a number of institutions during my graduate studies, including Florida State University, Tallahassee Community College, Valdosta State University, Thomas University, and Columbus State University, all of which have provided me with invaluable experience as an educator. v Finally, I am most grateful to my family, all whom have been unceasingly supportive (both emotionally and financially) through my entire academic career. To my parents I owe a debt of gratitude for instilling in me a curiosity of the world around me and for teaching me to persevere in pursuit of my dreams. However, I owe the biggest thanks of all to my wife, Diedre, who has sacrificed much for me to complete my graduate work, given me tireless encouragement and support through the years, and enriched my life beyond all expectations. Without you and our children, this would hold little meaning for me. vi TABLE OF CONTENTS List of Figures.......................................................................................................... x Abstract ................................................................................................................... xii 1. INTRODUCTION................................................................................................ 1 1.1 Purpose and scope................................................................................. 1 1.2 Methodology ........................................................................................... 3 2. GEOLOGY OF THE SOUTHERNMOST APPALACHIANS ............................... 5 2.1 General Overview................................................................................... 5 Foreland Fold and Thrust Belt ........................................................... 6 Talladega Belt (western Blue Ridge) ................................................. 9 Alternative Interpretations of the Talladega belt................................. 14 Ashland-Wedowee belt (western Blue Ridge).................................... 18 Ashland-Wedowee belt plutonic rocks ............................................... 20 2.2 Tectonic models for the southern Appalachians ..................................... 27 Iapetan Rifting and peri-Laurentian Microcontinents.......................... 27 Taconic orogenesis............................................................................ 31 Acadian orogenesis ........................................................................... 31 Alleghanian orogenesis...................................................................... 32 3. HOLLINS LINE FAULT SYSTEM AND HILLABEE THRUST............................. 34 3.1 Hollins Line: General Overview .............................................................. 34 3.2 Roof, Floor, and Splay Thrusts ............................................................... 35 3.3 Hillabee Thrust Overview........................................................................ 39 3.4 Nature of the Hillabee Thrust.................................................................. 40 3.5 Alternative Interpretations of the Hollins Line Fault System ................... 43 vii 3.6 Alternative Interpretations of the Hillabee Thrust.................................... 48 3.7 Alternative Interpretations, Shear Strain, and Metamorphism ................ 53 4. TACONIC OROGENY OF THE SOUTHERN APPALACHIANS ........................ 56 4.1 Arc-continent collisions ........................................................................... 56 4.2 Northern Appalachian Geology............................................................... 58 4.3 Southern Appalachian Geology .............................................................. 60 Talladega belt (Western Blue Ridge) ................................................. 60 Ashland-Wedowee belt (Eastern Blue Ridge).................................... 61 Arc-Continent Collisional Orogenesis – Talladega/Ashland-Wedowee Belts................................................................................................... 64 Arc-Continent Collisional Orogenesis – GA/NC/SC/TN Blue Ridge... 66 4.4 Accretionary Orogens ............................................................................. 70 4.5 Lithotectonic Elements of the Southern Appalachians............................ 73 Dahlonega Gold Belt.......................................................................... 74 Cowrock Terrane ............................................................................... 74 Cartoogechaye Terrane ..................................................................... 75 Western Tugaloo Terrane (eastern Blue Ridge) ................................ 76 Eastern Tugaloo Terrane (western Inner Piedmont).......................... 76 Cat Square Terrane (eastern Inner Piedmont)................................... 77 4.6 The Southern Appalachian Taconic Orogeny ......................................... 77 Latest Cambrian(?) to earliest Ordovician (>500 to ~485) ................. 78 Earliest Ordovician to middle Late Ordovician (<485 to ~455 Ma)..... 78 Late Ordovician and Younger Events (<455 Ma) ............................... 80 5. SUBDUCTION INITIATON ALONG THE EARLY PALEOZOIC LAURENTIAN MARGIN.................................................................................................................. 83 5.1 Models of Subduction Initiation............................................................... 83 5.2 Development of the Alabama Promontory Subduction Zone .................. 87 viii 5.3 Evolution of the Alabama Promontory Subduction Zone ........................ 93 6. CONCLUSIONS ................................................................................................. 99 APPENDIX .............................................................................................................. 103 REFERENCES........................................................................................................ 109 BIOGRAPHICAL SKETCH...................................................................................... 136 ix LIST OF FIGURES Figure 1.1: Southern Appalachian terranes, faults, and Gulf-Atlantic Coastal Plain unconformity............................................................................................................ 2 Figure 1.2: Geologic map of the study area (Alabama portion of the Talladega belt) showing 7.5’ quadrangle names, location of the post-metamorphic Hollins Line fault footwall duplex, Talladega Group, Hillabee Greenstone and pre-metamorphic Hillabee thrust fault. .............................................................................................................. 4 Figure 2.1: Generalized geologic map of the southern Appalachians. .................. 5 Figure 2.2: Summary of Paleozoic geologic events and stratigraphy of the Laurentian Alabama promontory and adjacent continental margin. .......................................... 8 Figure 2.3: Map showing King and Keller (1992) collection localities of metachert samples for chert crystallinity study......................................................................... 15 Figure 2.4: Map showing location of Elkahatchee sample used for SHRIMP RG U-Pb dating. ..................................................................................................................... 22 Figure 3.1: Generalized geologic map of the Hollins Line fault system and Hillabee thrust showing footwall-hanging wall stratigraphy and stratigraphic cutoffs, roof-floorsplay thrusts and fault bound horses, and regional geology.................................... 36 Figure 3.2: Lower hemisphere stereographic projection of poles to asymmetric extensional shear bands and net slip bearing determined from S-C fabric in a single exposure adjacent to the Hollins Line roof thrust, Lay Dam, 7.5' Quadrangle, AL... 39 Figure 3.3: Stereographic analysis of 60 poles to S0-S1 of the Lay Dam formation, Ross Mountain, 7.5' quadrangle, Alabama.............................................................. 46 Figure 3.4: - Stereographic analysis of 26 poles to S1 foliation planes of the Hillabee Greenstone, Ross Mountain, 7.5' quadrangle, Alabama. ........................................ 47 Figure 3.5: Stereographic analysis of 80 poles to S0-S1 planes of the Lay Dam formation, Cheaha Mountain, 7.5' quadrangle, Alabama. ....................................... 49 Figure 3.6: Stereographic analysis of 34 poles to S1 foliation planes of the Hillabee Greenstone, Cheaha Mountain, 7.5' quadrangle, Alabama..................................... 49 Figure 5.1: Evolution of the Ouachita embayment along southern Laurentia following Early Cambrian seafloor spreading at the Ouachita rift. .......................................... 87 Figure 5.2: Middle Cambrian juxtaposition of contrasting lithosphere across the Oklahoma-Alabama transform. ............................................................................... 89 x Figure 5.3: Evolution of the Puysegur subduction zone. ........................................ 94 Figure 5.4: Lithospheric collapse at the continent-ocean boundary and potential paths of trench propagation. ............................................................................................. 96 Figure 5.5: Lithospheric collapse along the Alabama-Oklahoma transform, propagation through transitional crust at the continent-ocean boundary and development of the Alabama promontory subduction zone and continental margin arc......................... 98 xi ABSTRACT Two major thrust systems are located along the eastern Blue Ridge – western Blue Ridge boundary in the Alabama and Georgia Appalachians. The pre-metamorphic Hillabee thrust is near the trailing edge of the Talladega belt thrust sheet and is associated with emplacement of a backarc volcanic suite atop rocks of the Paleozoic Laurentian shelf. The post-metamorphic Hollins Line fault system emplaces metasedimentary rocks of the Neoproterozoic Laurentian rifted margin atop the Talladega belt amalgamated shelf-volcanic terrane. These three, fault bounded lithotectonic belts are important for unraveling the geologic history of this part of the Appalachian orogen and have widespread implications for the tectonic history of the southern Appalachians in general. Following Neoproterozoic rifting of the Rodinian supercontinent and deposition of a thick metasedimentary package of rocks along an amagmatic rifted margin, now comprising the Ashland-Wedowee belt in the eastern Blue Ridge of Alabama and Georgia, shelf strata of the Kahatchee Mountain and Sylacauga Marble Groups (Talladega belt) indicate establishment of a passive margin along the Cambrian trailing edge of this portion of Laurentia. Intrusion of continental slope-rise deposits within the Ashland-Wedowee belt by latest Cambrian(?) to Middle Ordovician arc related plutons was immediately preceded by subduction initiation outboard of the trailing margin continental hinge zone as a result of passage of the Ouachita mid-ocean ridge along the southern margin of the Alabama promontory and foundering of Iapetan oceanic lithosphere in the weak transitional crust at the continent-ocean boundary. Rollback of the newly-formed, westward-dipping subduction zone resulted in extension within the overriding Laurentian plate and backarc spreading, leading to formation of the bimodal metavolcanic Hillabee Greenstone (Alabama-Georgia), Pumpkinvine Creek formation (Georgia), and other backarc affinity volcanic suites along structural strike of the Appalachians to the northeast. Continued westward subduction of Iapetan seafloor beneath the attenuated Laurentian margin resulted in extensive suprasubduction plutonism intruding the marginal rift basin and distal trailing margin sediments of the eastern Blue Ridge of Alabama and Georgia and farther to the northeast. Uplift of xii Cambrian and Ordovician shelf strata in the evolving backarc resulted in deposition of the southern Appalachian Blount-Taconic clastic wedge, which sourced intrabasinal uplifts along the westward facing (craton-facing) side of the Hillabee backarc, and perhaps localized thrust belts at the continental edge of the backarc associated with lithospheric compression during contractional phases of backarc evolution. Following Ordovician accretionary orogenesis (Taconic orogeny) in Alabama, Georgia, and farther northeast, deposition of the Talladega Group in the Silurian(?) to earliest Mississippian(?) indicates stabilization of the most outboard portion of the Laurentian shelf margin. Latest Devonian-Mississippian collision along the Alabama promontory, however, resulted in emplacement of the Hillabee Greenstone terrane atop the adjacent Devonian-earliest Mississippian(?) shelf strata of the Talladega Group along a thin-skinned thrust fault, the Hillabee thrust, possibly as a result of collision between the Ouachita arc and the Laurentian margin. Following movement along the pre-metamorphic Hillabee thrust, progressive dynamothermal metamorphism affected both the amalgamated Talladega belt (Hillabee and Talladega-Sylacauga MarbleKahatchee Mountain Groups) and Ashland-Wedowee belt. Alleghanian terminal orogenesis (continental collision), culminating in the late Paleozoic Pangean supercontinent, resulted in telescoping of the Laurentian margin and emplacement of rifted margin metasediments and associated arc plutons atop the Talladega belt along the post-metamorphic Hollins Line footwall duplex thrust system. xiii CHAPTER 1 INTRODUCTION 1.1 Purpose and scope Compression at active plate margins often results in the translation of significant bodies of rock along low angle thrust faults and subsequent lithospheric shortening and thickening. In the Appalachian orogenic belt of eastern North America, thrust faults commonly represent fundamental tectonic boundaries across which terranes with separate geologic histories have been juxtaposed. Examination of these terrane bounding faults, as well as the rocks of their foot and hanging walls, can often shed light on the constraints and kinematics under which these faults developed, as well as the geologic history of the juxtaposed terranes. This study examines two such terrane bounding faults in the southernmost Appalachians of Alabama (Fig. 1.1 and Plates 1 through 4), the pre-metamorphic Hillabee thrust and post-metamorphic Hollins Line fault system. Both faults developed during Devonian and later growth of the Appalachian mountain chain, and are significant in that they accommodated significant displacement and resulted in the juxtaposition of different tectonic elements of the Neoproterozoic – Paleozoic Laurentian margin (geologic time scale used here by Gradstein et al., 2004). In the case of the Hillabee thrust, a backarc bimodal volcanic suite was emplaced along this fault onto Neoproterozoic to uppermost Devonian-lowermost Mississippian(?) shelf sediments which formed along the trailing Laurentian margin. The combined package of rock was then metamorphosed to lower greenschist facies and is collectively termed the Talladega belt. This terrane was then duplexed by the post-metamorphic Hollins Line fault system, which emplaced probable Neoproterozoic rift-related metasedimentary rocks and Paleozoic intrusions of the hanging wall eastern Blue Ridge onto rocks of the Talladega belt and its duplexed stratigraphy of the Hollins Line footwall. Study of the combined lithotectonic elements associated with each fault provides considerable insight into evolution of Laurentia following breakup of the Neoproterozoic 1 Figure 1.1 – Southern Appalachian terranes, faults, and Gulf-Atlantic Coastal Plain unconformity. Modified from Hatcher et al., 2007. 2 supercontinent Rodinia and ending in amalgamation of the late Paleozoic supercontinent Pangea – a complete Wilson cycle (Wilson, 1966). 1.2 Methodology Geologic mapping at a 1:24,000 scale is vital to understanding the petrologic, stratigraphic, and structural relationships of rocks within and between lithotectonic terranes. Geologic data from the eastern and western Blue Ridge of Alabama has been compiled into a single database and resultant geologic map. The data utilized in this work spans twenty-two 7.5 minute USGS topographic quadrangles in eastern Alabama (Fig. 1.2) and was gathered by a number of researchers, including this author, James F. Tull, Thornton L. Neathery, Jonathan W. Mies, David T. Allison, Lance W. Johnson, Chul Lim, Connie Bocz-Garrett, and William W. Carter. This combined dataset, gleaned from published and non-published geologic maps and manuscripts, was compiled into a Microsoft Access™ database designed and hosted by Dr. David T. Allison, Department of Geography and Geology, University of South Alabama. This database is a cooperative effort between workers at Florida State University, Columbus State University, and the University of South Alabama. The database was then utilized for both structural analysis and the construction of geologic maps. Efforts to credit the appropriate workers for their contributions have been made in both construction of the database and in preparation of this manuscript. All geologic maps utilized in this manuscript were created using Autodesk Map 3D™, version 2004 or later. 3 Figure 1.2 - Geologic map of the study area (Alabama portion of the Talladega belt) showing 7.5’ quadrangle names, location of the post-metamorphic Hollins Line fault footwall duplex, Talladega Group, Hillabee Greenstone and pre-metamorphic Hillabee thrust fault. 4 CHAPTER 2 GEOLOGY OF THE SOUTHERNMOST APPALACHIANS 2.1 General Overview Figure 2.1 - Generalized geologic map of the southern Appalachians. Stippled pattern = Mesozoic and Cenozoic rocks of the Gulf-Atlantic Coastal Plain. Modified after Tull et al., 2007. The geologic history of the southernmost Appalachians includes elements of the Proterozoic supercontinent, Rodinia, rifting and drifting of the ancient North American continent, Laurentia, and Paleozoic orogenesis resulting in the development of the Appalachian Mountains and assembly of the late Paleozoic supercontinent, Pangea. In Alabama, these elements are preserved in three distinct lithotectonic terranes: the foreland fold and thrust belt of the Appalachian foreland, the Talladega belt of the western Blue Ridge, and the Ashland-Wedowee belt of the eastern Blue Ridge (Fig. 2.1). While rocks of the foreland fold and thrust belt were not involved in this research, 5 a general stratigraphy is presented here because of their correlatives with stratigraphy of the Talladega belt and implications for margin-wide tectonic events. Foreland Fold and Thrust Belt The foreland fold and thrust belt of Alabama includes a package of rocks ranging from Cambrian to Pennsylvanian in age (Thomas and Drahovzal, 1973; Kidd, 1975), known from the Appalachian and Black Warrior basins (Fig. 2.1). These rocks can be linked to North America both by fossil assemblages and correlative stratigraphy, formed on Laurentian continental lithosphere, and record Cambrian drifting of Laurentia to subsequent Paleozoic orogenesis (Thomas, 1991; Thomas, 2004). The TalladegaCartersville fault truncates the structural top of the foreland sequences in Alabama, separating them from hanging wall rocks of the Talladega belt to the southeast. The oldest strata from the foreland sequences of Alabama (Fig. 2.2) consist of the Lower Cambrian Weisner Formation, a siliciclastic unit deposited after cessation of significant rift related activity along the Alabama promontory. The Weisner Formation represents the uppermost unit of the drift facies Chilhowee Group in Alabama and Georgia, which developed along the trailing Laurentian margin during the late Neoproterozoic-Early Cambrian (Simpson and Eriksson, 1989). The Weisner is stratigraphically overlain by the Lower Cambrian Shady Dolomite, dominated by dolostones, and in turn by the Rome Formation, dominated by siliciclastics, both of which are interpreted as drift facies strata (Thomas and Drahovzal, 1973; Thomas, 1991). Atop the Rome lies the Middle to upper Cambrian Conasauga Formation, dominated by siliciclastics in western portions of the Appalachian basin and carbonates in eastern sections (Thomas, 1991). Within the Birmingham graben, coeval Conasauga strata are highly variable and are interpreted as having been deposited in basins experiencing synsedimentary basement faulting (Thomas, 1991; Thomas et al., 2000). The Conasauga, therefore, records intracratonic destabilization of the Laurentian passive margin and has been interpreted as marking the end of an ‘immature’ stage of margin drift (Glumac and Walker, 2000). This tectonic instability has been attributed to interaction between the northern terminus of the Ouachita mid-ocean ridge and the Alabama continental promontory across the Alabama-Oklahoma transform (Thomas, 1991; Thomas et al., 2000), however, an alternate explanation of cratonic instability 6 associated with dynamic topography due to initiation of a subduction zone outboard of the Middle Cambrian Alabama promontory will be explored later in this manuscript (Chapter 5). Following deposition of the Conasauga and equivalent strata, a widespread carbonate bank, represented by the Furongian to Lower Ordovician Knox Group, formed across the Laurentian margin. Some rocks of the Knox Group were subsequently eroded by the post-Knox unconformity, generally interpreted as a response to the Middle Ordovician Taconic orogeny (Bayona and Thomas, 2003; 2006). Middle and Upper Ordovician rocks of the Alabama foreland vary from clastic dominated in the east to carbonate dominated in the west. The clastics are generally interpreted as synorogenic deposits associated with the Blountian ‘phase’ of the Taconic orogeny (Athens Shale, Rockmart Slate, and Chota Formation), which prograde over carbonate facies rocks (Lenoir and Little Oak Limestones) to the west. These strata are succeeded by a series of clastics (Greensport Formation, Colvin Mountain Sandstone, and Sequatchie Formation), except in the most distal portions of the foreland, where the entire sequence (Chickamauga Limestone and Sequatchie Formation) is represented by carbonates. Coeval with Middle Ordovician synorogenic deposits, extensive ash deposits (k-bentonites) were deposited throughout the Appalachian basin. Palinspastically restored isopach maps of two of these, the Milbrig and Deicke, indicate a source in the vicinity of the Alabama promontory (Kolata et al., 1998) and are geochemically indicative of Plinian eruptions associated with a continental margin arc (Tucker, 1992; Tucker and McKerrow, 1995; Samson et al., 1989). Portions of the Middle and Upper Ordovician sequences are cut by a postOrdovician unconformity, especially toward the more internal portions of the orogenic belt (Thomas and Drahovzal, 1973; Bayona and Thomas, 2003; Thomas and Bayona, 2005). Synorogenic clastic wedge deposits of the Athens Shale and Rockmart Slate dominantly source intrabasinal sediments, but not sediments derived from advancing thrust sheets containing arc terranes (Bayona and Thomas, 2006), as proposed in the arc obduction model commonly used to explain the Taconic orogeny in the southern Appalachians. Silurian rocks of the foreland fold and thrust belt are represented by the Red Mountain Formation, a clastic-dominated sequence in the southeast grading into 7 8 Figure 2.2 - Summary of Paleozoic geologic events and stratigraphy of the Laurentian Alabama promontory and adjacent continental margin. carbonates in the northwest, which is unconformably absent in the more southeastern portions of the foreland. A Lower to Middle Devonian unit, the Frog Mountain Sandstone, consists mainly of clastics with lesser carbonate and chert components and is overlain by the Chattanooga and Maury shales (Upper Devonian and Lower Mississippian respectively). This strata is in turn overlain by shelf deposits of the Lower Mississippian Fort Payne Chert and Tuscumbia Limestone, which immediately precede synorogenic deposits associated with the Ouachita and Alleghenian orogens. Clastic wedge deposits from the developing Alleghanian orogeny to the east of the foreland are represented by the Upper Mississippian Floyd Shale-Pennington Formation and Pennsylvanian Gizzard Group-Sewanee Sandstone. Essentially coeval with Alleghanian synorogenic deposits, a separate clastic wedge from the developing Ouachita orogeny to the southwest is represented by the Mississippian Floyd ShaleParkwood Formation and Pennsylvanian Pottsville Formation (Thomas and Drahovzal, 1973; Bayona and Thomas, 2003; Thomas and Bayona, 2005). In summary, the foreland fold and thrust belt of Alabama records stabilization of the rifted Laurentian margin in the Early Cambrian and continual passive margin shelf sedimentation punctuated by episodic rift-related fault activity and/or dynamic topography from far field subduction stresses through the Early Ordovician. Middle and Late Ordovician orogenesis (Taconic) resulted in deposition of a prograding clastic wedge across stable margin strata of the foreland, while Silurian rocks suggest a resumption of passive margin sedimentation prior to erosion and truncation by an Early Devonian unconformity. Lower Devonian to Lower Mississippian rocks record restabilization of the Laurentian margin and resumption of shelf sedimentation, while the youngest units within the Alabama foreland document the onset of Carboniferous orogenesis, with thrust loading from the southwest (Ouachita) and east (Alleghanian) resulting in two prograding clastic wedges to the northeast and west, respectively (Thomas and Bayona, 2005). Talladega Belt (western Blue Ridge) The allochthonous Talladega belt of Alabama and western Georgia is a fault bounded block of lower greenschist facies rocks that formed along the Neoproteroic- 9 lower Paleozoic margin of Laurentia. At its structural base, the Talladega belt hanging wall is separated from anchimetamorphic rocks of the Appalachian foreland fold and thrust belt along the Talladega-Cartersville fault (Figure 2.1). At its structural top, the Talladega belt is juxtaposed against the overlying middle to upper amphibolite facies Ashland-Wedowee belt (AWB) of the eastern Blue Ridge across the Hollins line fault system. Rocks of the Talladega belt (Fig. 2.2) include Lower Cambrian to uppermost Devonian-lowermost Mississippian(?) Laurentian shelf strata and a pre-metamorphic structurally emplaced Ordovician bimodal volcanic suite. The structural and stratigraphic base of the Talladega belt consists of more than two kilometers of siliciclastics, the Kahatchee Mountain Group, correlative with the Laurentian drift facies, Lower Cambrian Chilhowee Group of the Alabama foreland. These rocks are conformably overlain by approximately three kilometers of carbonates, the Sylacauga Marble Group, which correlate with the Laurentian, Cambrian-Ordovician carbonate bank Shady Dolomite through Knox Group (Tull et al., 1988; Johnson and Tull, 2002). An unusual lithology within the Sylacauga Marble Group consists of arkosic sandstone within the carbonate-rich Shelvin Rock Church Formation (Johnson and Tull, 2002). These siliciclastics may be associated with local dissection of underlying feldspathic units and recycling of their clastic components, and if so, would seem to coincide temporally with rocks of the Conasauga, which record intracratonic destabilization of the Laurentian carbonate platform. If so, it would suggest that much of the Alabama promontory, from the Birmingham graben to the palinspastic location of the Talladega belt at or just beyond the Laurentian shelf edge, was affected by mild tectonism during the Middle or earliest late (Furongian) Cambrian. A low angle unconformity, the sub-Lay Dam Unconformity (Shaw, 1970; Cook, 1982; Tull et al., 1988; Tull, 2002), separates the underlying carbonate and clastic sequences from the superposed Talladega Group. The dominantly metaclastic Talladega Group has been interpreted as a middle Paleozoic successor basin which developed unconformably atop the Laurentian margin, Cambrian-Lower Ordovician carbonate shelf (Tull and Telle, 1989; Tull and Groszos, 1990). Rocks of the Kahatchee Mountain, Sylacauga Marble, and Talladega Groups can be stratigraphically and paleontologically linked with Laurentia and represent the southernmost and most 10 outboard preserved fragment of the Laurentian shelf (Butts, 1926; Tull et al., 1988; Gastaldo, 1995; Tull, 2002). The lowermost unit of the Talladega Group, the Lay Dam Formation, consists of >2 km of metaturbidite, arkosic conglomerate, and diamictite (Shaw, 1970; Carrington, 1973; Tull and Telle, 1989; Lim, 1998; Tull, 2002) deposited atop the sub-Lay Dam unconformity. Diamictites contain clasts derived from both the underlying Sylacauga Marble Group and Kahatchee Mountain Group in addition to granite and granitic gneiss derived from Grenville basement (Telle et al., 1979) and probably represent sediment deposited in the proximal portions of a submarine fan derived from uplifts to the southsoutheast (Shaw, 1970; Carrington, 1973; Tull and Telle, 1989; Lim, 1998; Tull, 2002). The age of the Lay Dam is constrained by both the age of the underlying and overlying strata, as well as conodonts from it’s uppermost section, and could be as old as Silurian. However, the nature of sediments within the Lay Dam suggests rapid deposition and it is likely that the entire sequence is early Devonian in age (Tull et al., 1988; Tull and Telle, 1989; Tull, 2002). Conformably overlying the Lay Dam Formation is the Lower Devonian Butting Ram/Cheaha Quartzite, consisting dominantly of quartzose metasandstone and metaconglomerate greater than 450 m in its thickest sections (Carter, 1985; Barineau, 1999a; 1999b; Tull, 2002). The Butting Ram/Cheaha equivalents represent stabilization of the Laurentian margin (Tull, 2002) and contain a diverse group of marine invertebrates of Early Devonian age (Butts, 1926; Carrington, 1973). The Butting Ram/Cheaha, together with the underlying Lay Dam Formation, is most likely equivalent to the Devonian Frog Mountain Formation of the Alabama foreland (Tull, 2002). In the southwest portions of the Talladega belt, the Butting Ram Quartzite grades into the overlying Jemison Chert, whereas it’s equivalent, the Cheaha Quartzite, grades into the overlying Jemison equivalent Erin Slate in the northeastern Talladega belt (Tull, 2002). Bearce (1973) mapped stratigraphy above the Cheaha Quartzite to the northeast of the Erin type section in Cleburne and Calhoun counties, naming it the Chulafinee Schist. However, because the three units (Jemison Chert, Erin Slate, Chulafinee Schist) can be mapped across the extent of the belt and are facies equivalents (Tull, 2002), the term “Chulafinee Schist” should be abandoned, with the 11 earlier terminology of Butts (1926) taking precedence. The age of the Erin-Jemison is constrained by fossils in the lowermost strata, a diverse assemblage of Oriskany Group marine invertebrates of probable middle Emsian or older age (Tull, 2002). Additionally, the presence of a rare plant fossil, Periastron, from the uppermost Erin (Gastaldo et al., 1993), indicates a Late Devonian (Famennian) to earliest Mississippi (Tournaisian) age for the youngest part of the stratigraphy. The Erin-Jemison interval is probably equivalent to the Chattanooga and Maury shales of the Alabama foreland (Tull, 2002) Structurally above the Talladega Group lies the Hillabee Greenstone, a bimodal metavolcanic sequence. Compositionally, the Hillabee consists dominantly of mafic phyllites-greenstones and subordinate quartz metadacites. Mafic phyllites and greenstones are interpreted as low potassium tholeiites and basaltic andesites based on major and trace element studies, and represent extensive basaltic lavas and basaltic ash deposits (Tull et al., 1978; Tull and Stow, 1980; Stow, 1982). Calc-alkaline metadacites and metarhyodacites, locally making up as much as 25% of Hillabee stratigraphy, can be traced for >20 km and range from 50 to >150 m in thickness. Individual metadacite layers are interpreted as laterally extensive felsic ash falls and welded ash flow crystal tuffs, in some cases covering >300 km2 and representing >50 km3 of eruptive material (Tull and Stow, 1980; Tull et al., 1998; Tull et al., 2007). In the central portion of the study area, at Pyriton, Alabama, massive, strata-bound sulfide zones have been mapped near the tectonic base of the Hillabee (Stow and Tull, 1979; Tull and Stow, 1982). Major, trace, and rare earth element analyses indicate that metavolcanic rocks of the bimodal Hillabee Greenstone formed in a suprasubduction setting, most likely in the backarc region of a continental margin arc (Tull et al., 1978; Tull and Stow, 1979, 1980; Durham et al., 1990; Tull et al., 1998; Barineau et al., 2006; Holm-Denoma, 2006; Tull et al., 2007) A number of geologists (Neathery, 1973; Tull et al., 1978; Tull, 1979; Tull and Stow, 1980) have argued that the Hillabee post-dated underlying units based on structural and metamorphic observations, interpreting it as a mid to late Paleozoic volcanic suite in stratigraphic contact with the underlying Talladega Group. In contrast, analyses of zircon from a Hillabee metadacite unit (Russell, 1978; Russell et al., 1984) provided 207Pb/206Pb ages ranging from 444 Ma to 462 Ma, indicating that the Hillabee 12 must be in fault contact with the structurally underlying Talladega Group. Subsequent ion microprobe U/Pb analyses of HG metadacite zircons confirmed this Middle to early Late Ordovician age of crystallization for the Hillabee and, therefore, its tectonic position above rocks of the Paleozoic Laurentian trailing margin (McClellan and Miller, 2000; McClellan et al., 2005, Tull et al., 2007; McClellan et al., 2007). Metadacite units of the Hillabee were erupted during the same time frame as Ordovician (Whiterockian) volcanism and K-bentonites of the Appalachian foreland basins (Kolata et al., 1996). Following emplacement of the Hillabee volcanic suite, the combined Laurentian shelf strata and Hillabee were subjected to Neo-Acadian or earliest Alleghanian lower greenschist facies metamorphism and deformation, which produced a pervasive, axial planar penetrative cleavage and resulted in isoclinal mesoscopic folding of Talladega belt lithologies. Post-metamorphic deformation within the study area includes thrusting along the Hollins Line fault system and emplacement of the structurally overlying Ashland-Wedowee (Jefferson terrane of Horton et al., 1989) belt, followed by large scale cross-folding of both thrust sheets (e.g. Millerville antiform) (Tull, 1977; 1978; Moore et al., 1983; Tull, 1984, 1995). In summary, rocks of the Talladega belt record the earliest Cambrian Laurentian drift phase (Kahatchee Mountain) following breakup of Rodinia and establishment of a carbonate platform (Sylacauga Marble), punctuated by Middle to earliest Late (Furongian) Cambrian tectonism (Shelvin Rock Church). Following uplift and erosion of these Laurentian shelf strata in Early Devonian time, stabilization of the margin is marked by the Devonian to Early Mississippian(?) strata of the Talladega Group, with resumption of tectonism in latest Devonian or Early Mississippian time concomitant with emplacement of an Ordovician backarc volcanic suite most likely outboard of the Laurentian continental hinge zone. The autochthonous history of the Talladega belt terminated with deformation and lower greenschist facies metamorphism of the amalgamated shelf rocks and volcanic suite. The geologic history of the Talladega belt is similar in many respects to that of the Alabama foreland, with three major differences. First, the Talladega belt lacks unambiguous Silurian strata and the post-Knox Taconic unconformity. Second, rocks of the Talladega belt Laurentian shelf (Kahatchee Mountain through Erin/Jemison strata) 13 are in pre-metamorphic faulted contact with a volcanic terrane (Hillabee Greenstone) which must have lain outboard of the Laurentian continental shelf. Finally, the Talladega belt experienced more significant deformation and is at a higher metamorphic grade than the Alabama foreland. All three of these differences may be attributed to the Talladega belt’s more outboard position on the Laurentian margin in the Paleozoic and, therefore, its resultant more internal position in the developing Appalachian orogenic belt. Alternative Interpretations of the Talladega belt Higgins and Crawford (2008) present an alternative interpretation of the Talladega belt in Alabama and Georgia, in which the Talladega belt is separated into an eastern and western section, roughly divided at the location of Interstate 65 near Jemison, Alabama. In their alternative interpretation, rocks of the “western Talladega slate belt” (Jemison Chert, Butting Ram Quartzite, and Lay Dam/Mohorn Creek metamudstone) represent relatively unmetamorphosed stratigraphic sequences of the Valley and Ridge, while rocks of the “eastern Talladega slate belt” consists of no less than six fault bounded structural panels separating rocks of the Lay Dam Formation from “Permian-Ordovician silicious mylonites” and the Hillabee Greenstone in a 1-3 km wide “Hillabee Fault Zone”. In this interpretation, the Kahatchee Mountain Group and Sylacauga Marble are mapped as part of the Valley and Ridge, while rocks of the Talladega Group (Lay Dam Fm., Cheaha Quartzite/Butting Ram Sandstone, and Erin Slate/Jemison Chert) are interpreted as “phyllonite”, “S-C mylonite”, and “silicious mylonite” derived from the Hillabee Greenstone during shearing along the “Hillabee fault zone”. In their interpretation, Higgins and Crawford (2008) argue that separation of the “eastern” and “western” Talladega belt is accommodated by displacement “along one of a known set of nearly north-trending faults” 100-150 meters east of Interstate 65 near Jemison, Alabama. This “alternative” interpretation of the geology of the Talladega belt has a number of significant problems, which I have outlined below. 1) Higgins and Crawford (2008) argue that stratigraphy in the “eastern” and “western” Talladega belt, especially that of the Jemison Chert, cannot be correlated across their proposed north-trending “fault” east of Interstate 65. On 14 their geologic map, stratigraphy of the Jemison Chert ends on the west side of Interstate 65, and rocks mapped by other workers as Jemison Chert on the east side of Interstate 65 are labeled “siliceous mylonite” of the “Hillabee fault zone”. However, in a detailed study of chert crystallinity in this particular area, King and Keller (1992) collected more than 300 samples of the Jemison Chert from 33 separate outcrops (Fig. 2.3) on both sides of Interstate 65 and the proposed Jemison, AL Approximate location of I-65 Figure 2.3 – Map showing King and Keller (1992) collection localities of metachert samples for chert crystallinity study. Approximate location of Interstate 65 and community of Jemison, Alabama, shown for reference. Map adapted from King and Keller, 1992. north trending “fault”. King and Keller (1992) note that all samples from the Jemison Chert (both east and west of Interstate 65) had a “distinct metamorphic crystal texture” and that none of the samples showed “anhedral, cryptocrystalline texture indicative of non-metamorphic or sedimentary chert.” They also note that, where bulk strain has resulted in coarser granoblastic textures, the Jemison Chert could be texturally described as a quartzite. However, they use the term “metachert” for the entire unit, regardless of the degree of quartz recrystallization. 15 Using scanning electron microscopy to analyze quartz crystallinity, King and Keller (1992) concluded that bulk strain and metamorphic grade increased from west to east in the study area (Fig. 2.3), consistent with interpretations of increasing metamorphic grade from the southwestern to the northeastern end of the Talladega belt, where the Barrovian chlorite zone gives way to rocks within the biotite zone near the Alabama-Georgia state line (Tull and Holm, 2005). 2) Higgins and Crawford (2008) argue that the Kahatchee Mountain Group and Sylacauga Marble Group, traditionally included as part of the lower greenschist facies Talladega belt (Tull et al., 1988; Johnson and Tull, 2002), are actually part of the “Valley and Ridge rocks that share anchizone metamorphism, slaty cleavage, and other characteristics with accepted Valley and Ridge rocks adjacent to them on the north”. However, they fail to note that both the Kahatchee Mountain Group and Sylacauga Marble Group rocks included in their Valley and Ridge stratigraphy east of Interstate 65 contain a well documented lower greenschist facies mineral assemblage (chlorite-actinolite-epidote), well above that of the anchizone Valley and Ridge rocks. This makes lumping the Kahatchee Mountain and Sylacauaga Marble Groups with the lower grade Valley and Ridge rocks an untenable hypothesis. 3) Higgins and Crawford (2008) argue that the Devonian Butting Ram sandstone is the only stratigraphy of the “western Talladega belt” that extends east of Interstate 65 “and it extends for only 100-150 meters or so east of the Interstate into the eastern Talladega slate belt where it grades into or is cut off by silicious mylonites in the Hillabee fault zone”. However, they fail to note that the type section of the Butting Ram Sandstone at Butting Ram shoals (Butts, 1926) along the Coosa River (prior to construction of Mitchell Dam in 1923) lies more than 4 miles east of Interstate 65 in rocks they interpret as part of the “eastern Talladega slate belt”. This makes it extremely difficult to argue that the Butting Ram Sandstone is truncated by a north-trending “fault” on the east side of Interstate 65, when the type section of this unit lies in rocks they interpret as “silicious mylonites” of the “Hillabee fault zone”. Additionally, it makes untenable the notion that rocks of the “Hillabee fault zone” are derived from shearing of the 16 structurally overlying Hillabee Greenstone, when the Devonian Butting Ram Sandstone, which they recognize as a stratigraphic unit, lies within the proposed “Hillabee fault zone” east of Interstate 65. 4) Higgins and Crawford (2008) argue that the Erin Slate and Jemison Chert cannot be correlated, in part, because “the Jemison is Lower and Middle Devonian” and “the Erin is lower Mississippian”. They fail to note, however, that the well documented fossil assemblage of the Jemison Chert is constrained to the lower 200 meters of the stratigraphic package, while the uppermost strata of the Jemison does not, to date, have a well delineated age (Tull, 2002). In contrast, identification of Periastron from the Erin Slate to the northeast of Jemison (Gastaldo et al., 1993) comes from the uppermost section of the Erin. The most comprehensive work on Periastron from North American rocks was done by Beck (1978) in southwestern Kentucky, where he collected phosphatic nodules from the upper New Albany Shale. Periastron has generally been assumed to be of Kinderhookian or Tournasian age (lowermost Mississippian) by recent workers in North America (Beck, 1978; Gastaldo, 1995) because of the strata in which it occurs in Europe. However, in southwestern Kentucky, where Beck (1978) collected his samples, the Falling Run bed (from which Beck assumed his phosphate nodules came) has been removed by erosion (Ettensohn et al. 1989), leaving open the question from which strata Beck’s samples originated. Interestingly, to the south of Beck’s sample localities, in Central Tennessee, where the Falling Run Bed reappears, this phosphatic nodule bearing lag zone contains conodonts of both latest Devonian and Early Mississippian age (Ettensohn et al., 1989). Thus, it is not inconceivable that Periastron (found in only two other localities outside of North America) could range from latest Devonian to Early Mississippian in age. Regardless of the age of Periastron, however, because rocks of the Jemison Chert can be mapped to the northeast into rocks of the Erin Slate, they are interpreted as facies equivalents, thus making it reasonable to assume that this stratigraphic interval ranges in age from Early to Middle Devonian at its base to latest Devonian or Early Mississippian near its stratigraphic top. If the nearly 1,000 meters of Erin 17 Slate now exposed in the Talladega belt (Tull, 2002) spanned an interval of this nature (~50-60 m.y.), deposition rates would fall in the 1.5 to 2cm per k.y. range, well within sedimentation rates for black shales forming in starved basins. Thus, the age constraints placed upon the Erin-Jemison interval, do not preclude it representing facies equivalents. Ashland-Wedowee belt (western Blue Ridge) The highly allochthonous, middle to upper amphibolite facies Ashland-Wedowee belt (Figure 2.1) of the Alabama/west Georgia eastern Blue Ridge consists predominantly of pelitic rocks and minor orthoamphibolite (Neathery, 1975; Tull, 1978; Muangnoicharoen, 1975; Tull, 1987; Allison, 1992). In Alabama, the belt can be divided into a structurally lower Coosa block to the northwest and structurally higher Tallapoosa block to the southeast. These blocks are separated from one another by a high angle normal fault, the Goodwater-Enitachopco. Steltenphol et al. (2005) argues for fault movement along the Goodwater-Enitachopco between ~344 Ma and ~327 Ma based on 40 Ar/39Ar analysis of white mica from a mylonitinized granite within the fault zone. However, the linear trace of the fault indicates that it was unaffected by late stage crossfolding which deformed pre-existing thrust faults and Pennsylvanian (and probably younger, now eroded strata to the northwest. This strongly suggests that the Goodwater-Enitachopco most likely formed during orogenic collapse following the Alleghanian orogeny (Tull, 1984; Allison, 1992) and that white mica 40Ar/39Ar ages obtained by Steltenpohl et al. (2005) represent cooling ages associated with uplift of the Ashland-Wedowee belt (Tull et al., 2007). The Ashland-Wedowee belt is bounded to the southeast by the Brevard Zone (Figure 2.1). The Ashland Supergroup comprises the structural and, most likely, stratigraphic base of the Ashland-Wedowee belt. Whereas primary facing data are lacking to date in the Ashland-Wedowee, the likelihood of the >10km thick stratigraphic package being overturned in its entirety is unlikely. The lowermost stratigraphy is represented by the Higgins Ferry and equivalent Poe Bridge Mountain Groups (Fig. 2.2), which are dominated by metapelites, but also contain metagreywacke, quartzite and minor (<7%) orthoamphibolite (Tull, 1978; Allison, 1992; Tull et al., 2007). Orthoamphibolites within 18 the Higgins Ferry/Poe Bridge Mountain Groups are interpreted as tholeiitic ocean-floor basalts and associated sills intercalated with deep water sediments (Tull, 1978; Thomas et al., 1980; Whittington, 1982; Stow et al., 1984; Drummond et al., 1988). Structurally above these rocks lie the Hatchett Creek and Mad Indian Groups, which, like the Higgins Ferry/Poe Bridge Mountain, are dominated by metapelites, but which also include calc-silicate gneiss, rhythmically bedded metagreywackes-metapelites, and lensoidal migmatitic pods of granitic rocks. Unlike the underlying strata, the Hatchett Creek/Mad Indian rocks do not include amphibolites. The presence of metamorphosed ocean floor basalts, voluminous pelites, and rhythmites strongly suggest development of the Ashland Supergroup within rift basins along a continental margin experiencing tectonism (Drummond, 1986; Allison, 1992; Tull et al., 2007). Structurally and, most likely, stratigraphically overlying the Ashland Supergroup is the Wedowee Group. The Wedowee is dominated by graphitic schists, with lesser amounts of metagreywacke, and is interpreted as having been originally deposited in an anoxic basin dominated by black shales (Drummond, 1986). Thus, the AshlandWedowee stratigraphy would seem to represent ocean floor basalts interlayered with rifted sediments deposited in continental rift basins (Higgins Ferry/Poe Bridge Mountain), succeeded by submarine fan deposits and aluminous pelites forming along a continental slope (Hatchett Creek/Mad Indian), and finally marine shales deposited in an anoxic basin (Drummond, 1986; Allison, 1992; Tull et al., 2007). The age and tectonic affinity of rocks lying structurally/stratigraphically above the Ashland-Wedowee stratigraphy (Emuckfaw-Canton-Univeter formations) is currently unknown but could be as young as Ordovician based on correlations with metavolcanic rocks of the Pumpkinvine Creek-Barlow Gneiss-Sally Free mafic complex, all of which yielded uppermost Lower Ordovician to Middle Ordovician U-Pb zircon ages (Thomas, 2001; Settles, 2002; Bream, 2003; Holm-Denoma, 2006). The lack of voluminous mafic volcanic and rift-related felsic volcanic rocks suggests formation of the Ashland Supergroup and Wedowee Group within an amagmatic (non-volcanic) rifted margin setting (Louden and Chian, 1999; Bryan et al., 2002). Louden and Chian (1999) divide amagmatic margins into two end members based on studies in the North Atlantic basin: a sediment-starved and sediment-filled 19 margin. In the first, a wide zone of faulted and extremely thinned upper continental crust is dominated by tilted blocks of continental basement with limited sediment cover. In the second, based primarily on studies of the Labrador margin, a limited number of tilted fault blocks exist outboard of the continental hinge and crustal extension is accommodated by a wide zone of thinned lower continental crust. Sediment thickness in this model can be significant (~9 km on the Labrador margin) and the ocean-continent transition zone is wider than that for the sediment-starved model. The great thickness (>10 km) of AWB metasediments suggests a similarity to the Labrador-type margin, thick sedimentary rift model. The stratigraphic character of the Ashland-Wedowee belt, correlation with parts of the eastern Blue Ridge in Georgia and the Carolinas, and its tectonic emplacement atop rocks of the Laurentian shelf (lower greenschist facies Talladega belt) strongly argue that the Ashland-Wedowee belt represents late Prototerozoic-early Paleozoic(?) sedimentary fill and related volcanic rocks along the slope-rise of the rifted Iapetan margin of Laurentia (Barineau and Tull, 2006; Barineau and Tull, 2007; Tull et al., 2007). The lack of voluminous mafic metavolcanic and rift-related felsic volcanic rocks suggest formation within an amagmatic rifted margin setting (Louden and Chian, 1999; Bryan et al., 2002), while the great thickness (>10 km) of metasediments suggests a similarity to the Labrador-type margin, thick sedimentary rift model proposed by Louden and Chian (1999). Ashland-Wedowee belt plutonic rocks The Ashland-Wedowee belt of Alabama-Georgia has been intruded by a number of pre-, syn-, and post-metamorphic I and S-type granitoid plutons, including the Elkahatchee Quartz Diorite, Zana, Kowaliga, Austell, and Sand Hill Gneisses, Rockford and Blakes Ferry granitoids (Fig. 2). Peraluminous rocks of the Elkahatchee, Zana, Kowaliga, and Villa Rica represent trondjhemite-tonalite-granodiorite suites which formed from intrusion of suprasubduction magmas into the rifted margin sediment prism outboard of the Laurentian continental hinge zone in the vicinity of the Alabama promontory (Drummond et al., 1994; Holm-Denoma, 2006). Previous attempts to isotopicallly date these granitoids has suggested Cambrian to Ordovician intrusive ages, 20 although the methods utilized in determining these ages yield large errors and often involve multi-grain analyses that may not be representative of a single age population. However, the numerous Cambrian to Ordovician dates determined for these granitoid bodies suggest they were emplaced between latest Cambrian and Middle Ordovician (Russell et al., 1987; Drummond et al., 1994; Holm-Denoma, 2006), a supposition supported by high resolution ion microprobe analysis of zircons from the Villa Rica Gneiss in the Georgia segment of the Ashland-Wedowee belt, which yielded an age of 458±3 Ma (Thomas, 2001). Emplacement of suprasubduction magmas during this time interval is further supported by Middle Ordovician bimodal metavolcanic terranes of the Hillabee Greenstone and Pumpkinvine Creek Formation, which occupy a structural position between the Ashland-Wedowee belt and Talladega belt-western Blue Ridge, and likely formed in a backarc position to the continental margin arc represented by the Alabama-Georgia granitoids (Holm-Denoma, 2006; Tull et al., 2007). The Elkahatchee Quartz Diorite is typically assumed to have formed during magma genesis in a suprasubduction setting, with the tonalitic, granodioritic, and minor mafic components interpreted as sourcing mid-ocean ridge basalt and minor deep sea sediments (Drummond et al., 1994). Accurate age constraints on the timing of Elkahatchee intrusion would constrain suprasubduction activity in the AshlandWedowee slope-rise rifted margin sedimentary package, however, existing chronologic data from the Elkahatchee (Russell et al., 1987) does not provide this information with the level of confidence necessary to draw definitive conclusions. A Rb-Sr whole rock age of 490 ± 28 Ma and U-Pb age for 5 multi-grain zircon separates show a range of 207 Pb/206Pb ages from 415 to 489 Ma and 206Pb/238U ages from 263 to 388 Ma (Russell et al., 1987). U-Pb data are scattered and discordant, but were interpreted as having an upper intercept of 496 ± 14 Ma by Russell et al. (1987). In an attempt to confirm the isotopic age data on the Elkahatchee from Russell et al. (1987), 17 zircons from a single locality (Fig. 2.4) were analyzed using sensitive high-resolution ion microprobe–reverse geometry (SHRIMP RG) by Paul Mueller, University of Florida (Table 2.1). A single morphological group of 9 prismatic grains ranged in age from 348 to 388 Ma (206Pb/238U) and yielded a standard error of the mean of 371 Ma ± 7 Ma (2σ) (Mueller, personal communication). Like other eastern Blue 21 Ridge plutons of similar age to the north, zircons included numerous inherited cores, plainly visible via cathode luminescence. Three grains with obvious inheritance were analyzed, yielding minimum ages of 480, 694, and 979 Ma, suggesting this sample was emplaced in rocks containing Mesoproterozoic, Neoproterozoic, and lower Paleozoic components. Alleghenian overgrowth was present as a 300 Ma concordant analysis on a single grain, which also exhibited high common Pb and relatively low Th/U, indicative of hydrothermal zircon growth. This may also suggest a period of partial Pb-loss in this population of zircons (Mueller, personal communication). These analyses are Figure 2.4 – Map showing location of Elkahatchee sample used for SHRIMP RG U-Pb dating. Location is proximal to sampling site reported by Russell et al. (1987). compatible with ages reported by Steltenpohl et al. (2005) on a trondhjemitic dike cutting shear zones within the Elkahatchee. Their data also suggested xenocrystic 22 23 Table 2.1 – Zircon U/Pb isotopic data (SHRIMP RG) for the Elkahatchee Quartz Diorite. (Mueller, personal communication). components, with 11 of 16 grains ranging from ~400 to ~1300 Ma, while the 5 youngest ages yielded an age of 369 ± 1.5 Ma (Steltenpohl et al., 2005). The obvious disagreement between the ages obtained via SHRIMP and those obtained by Russel et al. (1987) using U-Pb techniques suggests that Pb-loss, xenocrystic components, and/or higher uranium content in zircon overgrowth compounded the original multi-grain analysis and, additionally, that the zircons analyzed and reported by Russell et al. (1987) may have included samples from rocks that were not of the same genetic origin. This suggests that the Elkahatchee may not be a simple single-phase intrusive body and that determination of an intrusive age will require both further mapping of complex igneous relationships and more comprehensive isotopic work. Because of the ambiguity in the age of the Elkahatchee, it is important to identify other magmatic-volcanic rocks in the southernmost Appalachians that may be subduction related. Tull et al. (2007) and McClellan et al. (2007) reported ages for metadacites of the Hillabee Greenstone of ~460-470 Ma, interpreting the bimodal volcanic suite as part of a back-arc basalt-dacite association. Similarly, the Galts Ferry gneiss of the Pumpkinvine Creek Formation, also a bimodal sequence, is reported to be ~460 Ma (Thomas et al., 2001b; Das, 2006; Holm-Denoma, 2006). Structurally below the Galts Ferry gneiss, the Mulberry Rock Gneiss is exposed in an eyelid window below the Pumpkinvine Creek Formation and eastern Blue Ridge thrust sheet (Holm-Denoma, 2006). Based on its presumed structural/stratigraphic position, the Mulberry Rock has been proposed to be significantly older than the Galts Ferry, representing the Grenvillian basement to the Talladega belt (Higgins et al., 1996). However, Mueller (personal communication) reports a SHRIMP RG weighted mean 206Pb/238U age of 450 ± 4 Ma (2σ) for 6 concordant grains, consistent with laser ablation multiple-collection inductively coupled plasma mass spectrometry (LA-MC-ICPMS) zircon ages reported by Holm-Denoma (2006) and Das (2006). Taken together, these data suggest that active subduction with concomitant development of a back-arc environment and intrusion of magmas, both I and S-type, through lithosphere containing Grenville crustal components began along the southern Appalachian/Alabama promontory continental margin by at least 460-470 Ma (Hillabee Greenstone, Pumpkinvine Creek Formation, 24 Zana-Kowaliga Gneiss) and continued through ~450 Ma. Rb-Sr whole rock and U-Pb multi-grain zircon analyses (Russell et al., 1987), the presence of inherited latest Cambrian-Early Ordovician zircons in both Hillabee Greenstone metadacites (McClellan et al., 2007) and the Elkahatchee Quartz Diorite Gneiss (data reported here), as well as latest Cambrian-Early Ordovician zircons in Pleistocene coastal plain sediments of coastal Florida (Mueller et al., 2008), suggests that plutons of ~500 Ma may be present in the southernmost Appalachians in general and the Ashland-Wedowee belt specifically. A second generation of granitoids, the Zana Granite and Kowaliga Gneiss, include pre-metamorphic compositional granites, quartz monzonites, and granodiorites (Bieler and Deininger, 1987). Russell et al. (1987) assign ages of ~460 Ma to both plutons based on analyses of multi-grain zircon samples. Dating of zircon from the Kowaliga resulted in an upper intercept U/Pb isochron age of 460±19 Ma, with one concordant point yielding ages of 463±3 Ma, 462±4 Ma, and 459±18 Ma (206Pb/238U, 207 Pb/235U, and 207Pb/206Pb respectively). Analyses of zircon samples from the Zana yielded slightly discordant ages, ranging from 451±10 Ma to 464±10 Ma (207Pb/206Pb ages), however samples from both plutons plot on or very close to the same chord, and are thus interpreted as having been emplaced during the same phase of magmatic activity (Russell et al., 1987). As with the Elkahatchee, the Zana-Kowaliga granitoids are I-type and are interpreted as suprasubduction plutons forming the core of a continental margin volcanic arc. Both generally yield negative εNd, although relatively high initial 87Sr/86Sr values suggest the Zana may have been contaminated or partially derived from partial melting of sedimentary sequences within the host rock (Russell et al., 1987; Holm-Denoma, 2006). A third generation of plutons within the Ashland-Wedowee belt is collectively referred to as the ‘Rockford type’ granitoids and includes the foliated/lineated Rockford Granite, Bluff Springs Granite, and Almond pluton. Russell et al. (1987) noted that 11 multi-grain zircon analyses from these plutons did not form linear arrays on concordia diagrams, which they associated with mixing of older inherited zircons. Unlike the Elkahatchee and Zana-Kowaliga plutons, Rockford bodies are peraluminous S-type (Chappell and White, 1974) granitoids, and are interpreted as having been derived from 25 anatectic melting of the Ashland Supergroup country rocks during peak metamorphism (Drummond, 1986; Drummond et al., 1988; Allison, 1992). Russell et al. (1987) assigned a ‘tentative’ age of 366±21 Ma to the synmetamorphic Rockford granitoids (Bluff Springs granite) based on a whole rock Rb/Sr “errorchron”, but noted little confidence in the significance of this age. Sixteen U/Pb single grain analyses on trondhjemitic dikes cutting ductile shear zones in the pre-metamorphic Elkahatchee batholith (Steltenpohl et al., 2005) yielded an upper intercept isochron age of 369.4+1.6/-1.4, interpreted as the time of crystallization of the dike, consistent with the ~366 age reported by Russell et al. (1987). A pluton of questionable Rockford affinity, the trondhjemitic Blakes Ferry, lacks the foliated nature of the Rockford, is subcircular in map outline, lacking the deformation phases present in the pre-metamorphic Elkahatchee and Zana-Kowaliga, and has imparted a contact metamorphic aureole on the surrounding host rocks. A K-Ar muscovite cooling age of 324±11 Ma (Cambell, 1973; Drummond et al., 1994) suggests emplacement between the thermal peak coincident with Rockford granitic intrusions (~370) and uplift (~324 Ma). Therefore, the Blakes Ferry could either represent a postthermal peak intrusion related to the earlier Rockford granitoids, or alternatively, an eclogitic MORB derived magma from a remnant slab associated with Acadian orogenesis (Drummond et al., 1997) Because the Ashland-Wedowee belt is interpreted to represent the rifted margin sedimentary prism along the Laurentian continent; Upper Cambrian(?)-Ordovician suprasubduction plutonism within the eastern Blue Ridge indicates the existence of an early Paleozoic, Andean-type margin along this segment of Laurentia (Drummond et al., 1994; Hibbard, 2000; Holm-Denoma, 2006; Tull et al., 2007). The northeastern extent of this continental margin arc is currently unknown, but could have extended much farther along the Ordovican margin if the Emuckfaw-Canton-Univeter Formations are, in fact, correlative with metavolcanic rocks of the Pumpkinvine Creek-Barlow Gneiss-Sally Free mafic complex and their associated metasedimentary strata (Dahlonega Gold belt), which may have formed in a backarc setting on the Laurentian margin (HolmDenoma, 2006). 26 2.2 Tectonic models for the southern Appalachians The development of the Appalachian orogenic belt is generally attributed to three major tectonic events following Neoproterozoic breakout of Laurentian associated with breakup of Rodinia: the Middle to Late Ordovician Taconic orogeny (Drake et al., 1989), the Late Devonian to earliest Carboniferous Acadian orogeny (Osberg et al., 1989), and the Carboniferous to Permian Alleghanian orogeny (Hatcher et al., 1989a). While examining the Appalachians in this framework oversimplifies evolution of the orogenic belt to some degree, it is nonetheless a useful model for development of the Appalachians. The assembly of the late Paleozoic supercontinent Pangea and development of the Appalachians involve a number of lithotectonic elements including: Mesoproteroic basement of the Grenville orogeny and Rodinian supercontinent (Rankin et al., 1989), rift and drift sequences which developed on the trailing Laurentian margin during the Neoproterozoic and early Paleozoic (Rankin et al., 1989; Tull and Groszos, 1990; Tull, 2002; Thomas, 2006), cover sequences deposited on the margins of the developing orogenic belt (Drake et al., 1989; Osberg et al., 1989; Hatcher et al., 1989b), and volcanic-sedimentary packages which developed outboard of the craton on Laurentian affinity continental blocks and/or oceanic lithosphere, or plates other than Laurentia (Hatcher, 2005; Tull et al., 2007). Iapetan Rifting and peri-Laurentian Microcontinents Stratigraphic and isotopic data strongly suggest separation of eastern Laurentia (directions referenced to current coordinates) from other parts of the Rodinian supercontinent in two stages (Aleinikoff et al., 1995; Li and Tull, 1998; Cawood et al., 2001, Hibbard et al., 2007). Bimodal volcanic sequences in the southern-central Appalachian Mt. Rogers Formation point to an earlier, 760-700 Ma failed rifting event, whereas 572 – 564 Ma tholeiitic flood basalts of the Catoctin Formation suggest that successful rifting leading to Iapetan seafloor spreading occurred in the latest Neoproterozic (Ediacaran, GTS2004 timescale used throughout this manuscript) (Aleinikoff et al., 1995; Gradstein et al., 2004). In the northern Appalachians, rift-related gabbroic dikes of the Blair River Inlier range in age from 581-576 Ma, supporting 27 interpretations of initiation of Laurentian breakout from Rodinia during this time (Miller and Barr, 2004). Following latest Neoproterozoic separation of Laurentia from the former Rodinian supercontinent along the ‘Blue Ridge’ rift (Thomas, 1991), southern Appalachian strata of the Lower Cambrian Chilhowee Group and its equivalents mark the rift-drift transition and establishment of a passive margin (Thomas, 1977, Simpson and Eriksson, 1989; Thomas, 1991; Walker and Driese, 1991). In the Newfoundland Appalachians, the Bradore Formation marks the rift-drift transition, also suggesting Early Cambrian development of the passive margin in this portion of Laurentia, although with less clarity than that observed in the southern Appalachians (Williams and Hiscott, 1987; Cawood et al., 2001). Early Cambrian separation of the Argentine Precordillera from the Ouachita embayment (Thomas, 1991; Thomas and Astini, 1996) and latest Neoproterozic-Early Cambrian separation of the Newfoundland Dashwoods block (Waldron and van Staal, 2001; van Staal et al., 2007) from the Laurentian margin mark rifting events responsible for the development of peri-Laurentian microcontinental blocks. Separation of the Argentine Precordillera from Laurentia in the Early Cambrian was followed by its translation across an open Iapetus ocean and subsequent accretion to the South American craton (Thomas, 1991; Thomas and Astini, 1996; Thomas et al., 2001a). In contrast, latest Neoproterozoic-Early Cambrian separation of the Dashwoods block resulted in creation of only a narrow seaway (Humber seaway) between it and the Laurentian margin, followed by re-accretion of the Dashwoods block to Laurentia during the Ordovician (Waldron and van Staal, 2001; van Staal et al., 2007). Within the Ouachita embayment and conjugate Argentine Precordillera, synrift rocks are not coeval with the latest Neoproterozoic Ocoee rift basin(s) of the Tennessee embayment, but instead were deposited during the Early Cambrian – a time in which the Blue Ridge rifted margin was already in a drift stage (Thomas and Astini, 1996; Thomas et al., 2000). Passive-margin sedimentation along the southern margin of the Alabama promontory (Alabama-Oklahoma transform fault) was not established until Middle to late Cambrian (Furongian) (Gradstein et al., 2004) when post-Iapetan drift 28 facies rocks of the Conasauga Group on the Alabama promontory were well established (Thomas, 1991; Thomas et al., 2000). To the north, some researchers have proposed that rifting of the Dashwoods microcontinental block was part of a margin-wide event that may have separated a large number of individual microcontinents, or a single ‘ribbon’ continent, along the entire proto-Appalachian margin of Laurentia (Cawood et al., 2001; Hibbard et al., 2007). Cawood et al. (2001) argue that an initial Iapetan rifting event occurring at ~570 Ma was followed by a younger rift phase coeval with the 555 – 550 Ma Skinner Cove Formation and Lady Slipper Pluton, and separation of peri-Laurentian blocks from Laurentia. In this model, the lower Cambrian Chilhowee Group, with an extensive along-strike length, does not record the earlier Iapetan passive margin, but instead marks the transition from rifting to drifting following separation of these peri-Laurentian blocks (Cawood, et al., 2001; Hibbard et al., 2007). Cawood et al. (2001) and other researchers (Hibbard et al., 2007; van Staal et al., 2007) argue that the Laurentian conjugate margin, Amazonia, was separated by a wide tract of Iapetan seafloor by 550 Ma, following final breakup of Rodinia at ~570 Ma, clearly predating the Lower Cambrian Chilhowee Group drift strata and requiring a younger rifting event between 540 and 535 Ma. I would suggest, however, that the Amazonia paleomagnetic record and stratigraphic record of the southern Appalachians do not support the interpretation of Neoproterozoic-Early Cambrian rifting of peri-Laurentian microcontinents along the entire margin. The degree of separation between Laurentia and the conjugate Amazonian craton in the latest Neoproterozoic is hampered by limited high precision paleomagnetic data. In a comprehensive survey of available Gondwanan and Laurentian paleomagnetic data, Tohver et al. (2006) indicate that Laurentia and the supposed conjugate Amazonian craton must have separated in an interval between 600 and 500 Ma, associated with opening of Iapetus, but note that the exact timing of this separation if poorly constrained based on existing data sets. As such, the width of Iapetus between Laurentia and Amazonia at the proposed time of Dashwoods separation, ~550530 Ma, is poorly constrained. Thus, it is possible that the tract of Iapetan oceanic crust separating Laurentia and Amazonia remained narrow at 550 Ma and the mid-Iapetan Blue Ridge rift was relatively close to the Laurentian margin at the close of the 29 Neoproterozoic. In fact, the Iapetan rift-drift history along the trailing Laurentian margin may be quite similar to the rift-drift history of younger rifted margins, including that of the Red Sea region and Atlantic margin of North America. In the Gulf of Aden, synrift sedimentary deposits indicate rifting of the Arabian plate from the African craton at ~35 Ma, followed by seafloor spreading at ~16.5 Ma. In the Red Sea, 26 – 18 Ma rift-related volcanic activity was followed by diachronous seafloor spreading at ~4 Ma in the southern Red Sea and modern embryonic seafloor spreading in the northern Red Sea (Wolfenden et al., 2005). In eastern North America, rifting along the Atlantic margin, like the Red Sea, was diachronous. Rifting in the Middle Triassic to Early Jurassic was followed by a rift-drift transition which began first at the southern end of the North American Atlantic margin (~200 Ma), coeval with development of the Central Atlantic Magmatic Province (Schlische et al., 2003), and progressed to the north over a period of 15 m.y. or more (Withjack et al., 1998). The collective Red Sea – Atlantic Ocean record suggests that igneous activity associated with continental breakup could reasonably precede seafloor spreading by 12 to 22 m.y. at any single location within an opening ocean, and that establishment of a laterally extensive basin floored by oceanic crust, especially in diachronous rifting, could take 15 m.y. or longer after establishment of the earliest “passive margin” sediments in the older portions of the ocean basin. Assuming that rocks of the Catoctin, Blair River Inlier, Skinner Cove volcanic sequence and Lady Slipper Pluton represent widespread ~580 to 550 Ma synrift igneous activity (Aleinikoff et al., 1995; Cawood et al., 2001; Miller and Barr, 2004), an Early Cambrian (542-513 Ma) rift-drift transition recorded in rocks of the lower Chilhowee Group in Virginia (Simpson and Erikksson, 1989) represents a time gap of this order. This would support the idea that the Chilhowee represents strata deposited along the drifting margin of Laurentia facing an open Iapetus and not a ‘ribbon-continent’ or array of margin parallel, isolated blocks of peri-Laurentian terranes proposed by some workers (Cawood et al., 2001; Hibbard et al., 2007; van Staal et al., 2007). Additionally, subsidence curves from the Alabama foreland do not support a proposed margin-wide Early Cambrian rifting event. Instead they show steadily decreasing Early Cambriian subsidence rates following successful Neoproterozoic 30 rifting of Rodinia (Thomas et al., 2000), typical of that observed on other rifted margins (Lin et al., 2003). The collective data suggests that separation of the northern Appalachian’s Dashwoods block was a relatively isolated event and not associated with rifting along the entire trailing margin of Laurentia, but instead could have been a result of diachronous seafloor spreading and/or off-axis igneous activity similar to that observed along the Mesozoic rifted Atlantic Newfoundland-Iberian conjugate margins (Jagoutz et al., 2007). Taconic orogenesis The classic Taconic orogeny in the northern Appalachians is interpreted as the result of the subduction of Laurentia beneath a volcanic arc of exotic origin (Drake et al., 1989) or beneath an arc built on peri-Laurentian rifted continental crust (Waldron and van Staal, 2001). This model of A-type subduction of the Laurentian margin has been extended into the central and southern Appalachians by some workers (Hatcher, 1987; Drake et al., 1989; Hatcher, 2005), while other workers contend that Taconic orogenesis was the result of Andean-type subduction along the Laurentian margin (Hibbard, 2000; Tull et al., 2007). Regardless of the model invoked, clastic wedge deposition, metamorphism, and deformation of Proterozoic and early Paleozoic Laurentian cover sequences suggest tectonism along the Laurentian margin during the Middle to Late Ordovician (Drake et al., 1989; Moecher et al., 2004; Hatcher, 2005; Bayona and Thomas, 2006; Corrie and Kohn, 2007). Resolution of conflicting tectonic models of the Taconic orogeny in the southern Appalachians will be addressed in more detail later in this manuscript (Chapter 4). Of importance, however, is the lack of significant deformation that could be attributed to the Taconic orogeny in the Talladega belt, suggesting that Ordovician deformation in the southernmost segment of the exposed Appalachians was either confined to portions of Laurentia far outboard of the shelf edge, or mild in its intensity (Tull, 1998; Tull, 2002; Tull et al., 2007). Acadian orogenesis Devonian and early Carboniferous tectonism is typically attributed to Acadian and/or Neo-Acadian orogenesis. Tectonic models for these events include subduction 31 of the Taconic modified Laurentian margin beneath one or more exotic terranes (Osberg et al., 1989; Bradley et al., 2000; Schoonmaker et al., 2005), flat-slab subduction beneath the Laurentian craton (Murphy and Keppie, 2005), terrane accretion along an Andean-type margin (Hatcher, 1987), with the possibility of significant strike-slip tectonics (Ettensohn, 2004). In the southern Appalachians, the most intense phase of Paleozoic deformation and metamorphism has historically been attributed to Taconic orogenesis (Hatcher, 1987; Moecher et al., 2004; Hatcher et al., 2005; Southworth et al., 2005), however, white mica ages from the Talladega belt (Kish, 1990; McClellan et al., 2005; McClellan et al., 2007) and farther to the northeast in the western Blue Ridge of Georgia and North Carolina (Hames et al., 2007) suggest an Acadian to earliest Alleghanian age of metamorphism to at least greenschist facies for portions of the Blue Ridge in the southern Appalachians. Additionally, deformation of Devonian-earliest Mississippian(?) strata in the Talladega belt (Gastaldo, 1995; Tull et al., 1988) indicate an orogenic event of Neo-Acadian or earliest Alleghanian age in this segment of the Appalachians. Alleghanian orogenesis Carboniferous and Permian tectonicsm is generally attributed to the effects of Alleghanian and Ouachitan orogenesis, which resulted in formation of the supercontinent Pangea and terminal closure of the Iapetan and Rheic ocean basins between Laurussia (Laurentia amalgamated with Baltica and one or more microcontinents/arc terranes) and Gondwana (Hatcher et al., 1989a; Stampfli and Borel, 2002; Thomas, 2006). The Alleghanian orogeny telescoped the pre-existing orogenic and cover sequences in the Blue Ridge, resulting in a series of thrust faults which accommodated transport of the Appalachian crystalline core toward the craton and folding/faulting in the Appalachian foreland basins to the west (Hatcher et al., 1989; Thomas and Bayona, 2005). This terminal orogeny in the Appalachian sequence also produced a series of clastic wedges along the developing continental suture (Hatcher et al., 1989; Ettensohn, 2004, Thomas et al., 2004), metamorphism (Hatcher et al., 1989; Wooton et al., 2005), and syncollisional granites (Samson et al., 1995). In the Talladega belt, both the upper and lower bounding thrusts (Talladega-Cartersville fault 32 and Hollins Line fault respectively) as well as large scale cross-folds, formed during this orogenic phase (Tull, 1978). 33 CHAPTER 3 HOLLINS LINE FAULT SYSTEM AND HILLABEE THRUST 3.1 Hollins Line: General Overview The Hollins Line fault represents the fundamental boundary between rocks of the Talladega belt (footwall) and Ashland-Wedowee belt (hanging wall) of AlabamaGeorgia. As such, the Hollins Line is the eastern Blue Ridge – western Blue Ridge boundary in this segment of the Appalachian orogenic belt. The fault-bounded western Blue Ridge terrane is interpreted as the most internal segment of the Appalachians that has definitive links to ancient North America and collectively includes rift, drift, and cover sequences of the trailing Laurentian margin following Neoproterozoic breakup of Rodinia. Rocks of the central and eastern Blue Ridge have been historically viewed as suspect or exotic terranes, were generally subjected to higher grades of metamorphism and deformation, and are more internal to the Appalachian orogenic belt. In recent years, detrital zircon data from central and eastern Blue Ridge rocks (Bream et al., 2004) in Georgia, North Carolina, and Tennessee have suggested a Laurentian affinity for many of these fault bounded terranes, indicating that the central/eastern Blue Ridge should probably be viewed as a composite terrane with varying tectonic origins and histories. The eastern Blue Ridge – central/western Blue Ridge boundary consists of different generations of faults with unclear relationships (Fig. 1.1). To the northeast, in North Carolina and Georgia, the Hayesville fault separates rocks of the western Blue Ridge and rocks of the central Blue Ridge (Hatcher, 1978; Settles, 2002; Hatcher et al., 2005; Hatcher et al., 2007). Relationships between the Hayesville fault to the northeast and Allatoona fault to the southwest at the structural base of the Dahlonega Gold belt and southern terminus of the Hayesville-Soque River thrust sheet in northern Georgia are complex (Settles, 2002). Unlike the complexly deformed Hayesville fault, the relatively straight trace of the Allatoona fault suggests late stage, post-metamorphic displacement along its length (Tull and Holm, 2005; Holm-Denoma, 2006; Tull et al., 2007), although some workers have suggested pre- to syn-metamorphic movement 34 along the Allatoona fault followed by late Paleozoic reactivation (Settles, 2002; Eckert and Hatcher, 2003). To the southwest, the Allatoona fault truncates the postmetamorphic Hollins Line fault system in Alabama, west of the AL-GA state line, evidenced by the post-metamorphic cross-folds affecting the Hollins Line and notably absent in the Allatoona fault of Georgia. Marking a significant topographic, structural, stratigraphic, and metamorphic discontinuity in eastern Alabama, the Hollins Line fault system has been described and studied by a number of workers (Prouty, 1923; Adams, 1926; Neathery and Reynolds, 1973; Tull, 1975, 1977, 1978, 1982; Mies, 1991, 1992; Tull et al., 2007). To the northwest of the Hollins Line fault system, the lower greenschist facies Talladega belt footwall represents the leading metamorphic allochthon in this portion of the Appalachian orogen. Southeast of the Hollins Line, the middle to upper amphibolite facies Ashland-Wedowee belt makes up rocks of the hanging wall. Stratigraphy of the uppermost Talladega Group and Hillabee Greenstone in the Talladega belt footwall, as well as the penetrative metamorphic fabric, is truncated by the Hollins Line, a feature similarly observed in the stratigraphic and structural cutoffs of the hanging wall AshlandWedowee belt (Fig. 3.1). The Hollins Line consists of a roof thrust, floor thrust, and connecting splay thrusts (Figure 3.1). Because splay thrusts from the main roof thrust imbricate rocks of the footwall Talladega belt (Moore and Tull, 1983; Tull, 1994; Tull et al., 2007), the Hollins Line can be classified as a footwall duplex thrust system after the terminology of Boyer and Elliot (1982). 3.2 Roof, Floor, and Splay Thrusts The ~200 km trace of the Hollins Line roof thrust marks a significant topographic, structural, stratigraphic, and metamorphic break between the lower greenschist facies, amalgamated Hillabee Greenstone – Laurentian shelf drift/cover strata of the Talladega belt and the middle to upper amphibolite facies Laurentian attenuated margin rift-drift(?) strata of the Ashland-Wedowee belt (Tull, 1978). Stratigraphy of both the Talladega and Ashland-Wedowee belts, as well as stratigraphically controlled topography in the 35 Figure 3.1 - Generalized geologic map of the Hollins Line fault system and Hillabee thrust showing footwallhanging wall stratigraphy and stratigraphic cutoffs, roof-floor-splay thrusts and fault bound horses, and regional geology. Modified from Tull et al., 2007. 36 vicinity of the Hollins Line, is commonly truncated across the roof thrust (Figure 3.1). Whereas rocks of the hanging wall Ashland-Wedowee belt exhibit little mechanical control on the stratigraphic level of the Hollins Line roof thrust, the mechanically weak Hillabee Greenstone exhibited significant mechanical control over the stratigraphic level of the roof thrust in the Talladega belt footwall. As such, >82% of the trace of the Hollins Line roof thrust abuts rocks of the Hillabee Greenstone in the footwall, with rocks of the Erin-Jemison making up the remaining ~18% of the structural top of the Talladega belt. In general, the Hollins Line roof thrust dips, on average, 15-20° and juxtaposes mica schists of the Ashland-Wedowee belt in the hanging wall against chlorite phyllites (greenstones) of the Hillabee Greenstone (Tull, 1995). Because of the differences in relative resistivity between the easily weathered mafic rocks (greenstones and mafic phyllites) of the Talladega belt and more resistant mica schists and quartzites of the Ashland-Wedowee belt, the fault is commonly marked by a distinct topographic escarpment. However, where amphibolites of the Poe Bridge Mountain/Higgins Ferry Group are juxtaposed against rocks of the Hillabee Greenstone in the footwall, the Hollins Line roof thrust is not generally marked by a change in topography and is typically difficult to locate barring excellent exposure, as the highly weathered greenstone/mafic phyllite is usually indistinguishable from highly weathered amphibolite. Because intense weathering in the study area (>59 inches of precipitation annually) typically limits exposure of Hillabee Greenstone and Poe Bridge Mountain/Higgins Ferry amphibolites to select roadside exposures and some stream cuts, the Hollins Line is not generally visible as a distinct feature in outcrop, but can be commonly delineated to within an area of a few meters to few hundred meters dependant on surrounding outcrop. Mica schists of the hanging wall Ashland-Wedowee belt sometimes exhibit “button” textures (S-C mylonites) associated with rare outcrops in which the Hollins Line can be observed, but as with the footwall Talladega belt rocks, weathering commonly obscures fault-related shear fabrics. The Hollins Line floor and splay thrusts imbricate rock units in the Talladega belt footwall and delineate a number of parautochthonous fault bounded ‘horses’ (Boyer and Elliot, 1982). Horses of the fault duplex are typically stacked in an en-echelon right sense (Fig 3.1B) and vary in length from 2 to 100 km (Tull et al., 2007). Unlike the roof 37 thrust, the Hollins Line floor and splay thrusts are not associated with a metamorphic discontinuity, although Talladega belt stratigraphy and structures are commonly truncated by these faults. Footwall splays commonly sole in and are parallel to stratigraphy of the Erin-Jemison and Hillabee strata, although rocks of the CheahaButting Ram locally mark the structural base of duplex horses. Thus, while the Hillabee Greenstone exerted the dominant mechanical control on the position of the Hollins Line roof thrust, splay thrusts most commonly cut structurally through the Hillabee Greenstone stratigraphy and sole in the metapelites of the Erin-Jemison (>56% of the trace of the floor thrust). Less commonly, splay thrusts cut completely through the Hillabee and Erin-Jemison strata and sole in quartzites of the underlying CheahaButting Ram (~18% of the trace of the floor thrust) which are generally more mechanically competent than the micaceous lithologies of the Erin-Jemison. Although evidence for a thrusting along the Hollins line system is unequivocal, lack of cohesive exposures in proximity to the Hollins Line constituent faults limits outcrop availability for kinematic analysis and, additionally, kinematic indicators collected during this study do not provide a consistent direction of displacement. However, fault decapitation of dextral, map scale folds (half wavelength >1.5 km) in the footwall duplex and floor thrust footwall (parautochthon) during emplacement of the eastern Blue Ridge allochthon along the Hollins Line, coupled with en-echelon right stacking of horses (Fig. 3.1B) within the Hollins Line footwall duplex suggest dextral transpressional movement along the fault system (Tull, 1994; Tull, 1995a, 1995b; Tull et al., 2007). Detailed study of a single outcrop within ~40 meters of the Hollins Line roof thrust (Lay Dam, 7.5’ quadrangle – Fig. 1.2), however, did provide a single unequivocal S-C fabric within rocks of the footwall Talladega Group (Erin-Jemison) with a calculated top to the west along a bearing of 25° N90E, consistent with interpretations of the Hollins Line as accommodating dextral transpressional motion. Additionally, asymmetrical extensional shear bands in this same outcrop, which typically develop at angles of 15° to 25° to the plane of bulk flow within a shear zone (Hanmer and Passchier, 1991) suggest top to the west-northwest displacement along the Hollins Line (Fig. 3.2). Regardless, both the paucity and variable nature of kinematic indicators along the Hollins Line roof, splay, and floor thrusts necessitates further research for a 38 more definitive interpretation of the net slip vector associated with this Alleghanian thrust system. Figure 3.2 – Lower hemisphere stereographic projection of poles to asymmetric extensional shear bands and net slip bearing determined from S-C fabric in a single exposure adjacent to the Hollins Line roof thrust, Lay Dam, 7.5' Quadrangle, AL. 3.3 Hillabee Thrust Overview Geologic mapping within the Talladega belt by a number of workers (Neathery, 1973; Tull et al., 1978; Tull, 1979; Tull and Stow, 1980) indicate remarkable concordancy of the Hillabee Greenstone with stratigraphy of the Talladega Group within both horses of the Hollins Line footwall duplex and the parautochthon (Fig. 3.1). As such, metavolcanics of the Hillabee Greenstone were interpreted as Devonian or younger based on this apparent concordant contact and lack of a metamorphic or fabric break between the Devonian-Lowermost Mississippian(?) Erin Slate-Jemison Chert and 39 the Hillabee Greenstone (Tull et al., 2007). Isotopic dating of metadacites within the Hillabee Greenstone metavolcanic complex (Russell, 1978; Russell et al., 1984; McClellan and Miller, 2000; McClellan et al., 2005, Tull et al., 2007; McClellan et al., 2007), however, has shown conclusively that the Ordovician Hillabee must be in fault contact with the underlying younger Talladega Group metasediments. Because Laurentian shelf rocks of the Talladega belt share the same metamorphic and deformational characteristics as the structurally overlying, fault-emplaced Hillabee Greenstone, this fault must be either pre- or syn-metamorphic in nature (Barineau and Tull, 2001; Barineau et al., 2006; Tull et al., 2007; McClellan et al., 2007). 3.4 Nature of the Hillabee Thrust The Hillabee Greenstone can be mapped along strike for more than 230 km, from outliers in the Gulf Coastal Plain southwest of Jemison, Alabama, northeast to the Alabama-Georgia state line near Hightower, Alabama (northeastern limit of the study area) and ultimately south of Tallapoosa, Georgia (Heuler and Tull, 1988; Heuler, 1993). The tectonic base of the Hillabee is in contact with only the uppermost unit of the Talladega Group stratigraphy over this entire length – the upper Devonian to lowermost Mississippian(?) Erin Slate-Jemison Chert (Tull et al., 2007). This relationship holds true in both horses of the Hollins Line footwall duplex and in parautochthonous rocks of the Hollins Line floor thrust, Talladega belt footwall. Additionally, internal stratigraphy of the Hillabee (extensive metadacites (Fig. 3.1C) and massive sulfide zones) is not only internally concordant, but is also parallel to internal stratigraphy of the Erin-Jemison in both the parautochthon and the footwall duplex along the entire length of the Talladega belt. Finally, the fault separating the Hillabee from the underlying Erin-Jemison remains essentially parallel to stratigraphy of the Cheaha Quartzite-Butting Ram Sandstone beneath the Erin-Jemison in both horses of the footwall duplex and the parautochthon (Tull et al., 2007). This remarkable concordancy suggests that the fault along which the Hillabee was emplaced paralleled stratigraphy of both the hanging and footwall, resulting in a flat-on-flat geometry of considerable length and areal extent (Barineau and Tull, 2001; Tull et al., 2007). 40 Radiometric ages of Hillabee metadacite crystallization and Talladega belt metamorphism, the geometric relationships between the Hillabee and underlying Talladega Group, and geochemical interpretations of the tectonic origin of the Hillabee suggest the following characteristics for the Hillabee thrust separating the Hillabee Greenstone and underlying Talladega Group shelf metasediments: 1) Because the Talladega belt can be palinspastically restored to at least the present location of the Pine Mountain belt, rocks of the Talladega Group must have formed at or beyond the Paleozoic Laurentian shelf edge (Ferrill and Thomas, 1988; Thomas, 2004). Because there are no volcanic rocks of coeval age in either the Talladega belt or foreland fold and thrust belt to the northwest, and because rocks of the Hillabee metavolcanic suite are fault emplaced, they must be exotic to the Laurentian shelf at the palinspastic location of the Talladega Group. Additionally rocks of the Silurian(?)-Lower Devonian Lay Dam Formation within the Talladega Group, which formed unconformably atop strata of the underlying Sylacauga Marble and Kahatchee Mountain Groups, cannibalized both Cambrian-Ordovician carbonate and clastic rocks of the underlying strata, as well as Grenville basement, but show no evidence of entraining debris from an Ordovician arc terrane. Thus, the Hillabee Greenstone (Hillabee terrane) cannot be native to the Talladega belt Laurentian shelf and must have formed outboard of the Laurentian shelf edge and continental hinge zone (Barineau and Tull, 2006; Tull et al., 2007). 2) Because the youngest metasediments of the Talladega Group (Erin-Jemison) are at least uppermost Devonian and possibly as young as lowermost Mississippian, emplacement of the Hillabee and active movement along the Hillabee thrust can be no older than 375 to 359 Ma (Butts, 1926; Gastaldo, 1995; Tull, 2002; Tull et al., 2007). 3) Because of the extensive flat-on-flat geometry of the Hillabee thrust and the remarkable concordance between strata in the Talladega Group and Hillabee Greenstone, the Hillabee thrust must be a thin-skinned thrust that transported rocks of the Hillabee backarc from their palinspastic location outboard of the Laurentian shelf onto cover sequences of the Laurentian shelf (Talladega Group) 41 without significantly deforming the hanging or footwall stratigraphy (Tull et al., 2007). Three scenarios have the potential to explain this lack of fault related (Hillabee thrust) deformation. First, this might suggest that transport of the Hillabee terrane did not involve translation over significant thrust ramps as the fault cut up section from the level of the Hillabee backarc to the level of the Laurentian shelf margin (Talladega Group). Alternatively, if the Hillabee thrust did involve translation of the Hillabee terrane over significant thrust ramps, then evidence for deformation of hanging wall rocks (Hillabee Greenstone) has been removed by post-metamorphic Hollins Line thrusting. Finally, it is possible that Hillabee thrust related deformation is evident, but has not yet been recognized in the Talladega belt (see section 3.5 for a discussion of alternative views of the Hillabee thrust). The cryptic nature of the Hillabee thrust is the subject of ongoing research by this author and others (Tull et al., 2007; McClellan et al., 2007). 4) Because the Talladega belt was metamorphosed between 370-366 Ma (approximate metamorphic age of the adjacent eastern Blue Ridge allochthon) and 334-320 Ma (40Ar/39Ar age for Talladega belt white mica – see Tull et al., 2007 for discussion), the Hillabee thrust must have ceased to be active prior to the youngest possible metamorphic age. This, coupled with point #2 above, indicates that the Hillabee was emplaced between 375 and 320 Ma in the most liberal interpretation (Barineau and Tull, 2001; Tull et al., 2007; McClellan et al., 2007). 5) Because the Hillabee and metasediments of the Talladega belt share a common metamorphic and deformational history, with no evidence for earlier thermal or deformational events in either body of rock, and because isotopic ages of Hillabee metadacites indicate crystallization of the Hillabee during the lower Middle Ordovician (~468 to 472 Ma), rocks of the Hillabee must have remained, essentially, tectonically undisturbed following eruption of the maficfelsic volcanic suite in a continental margin backarc and their emplacement atop rocks of the Laurentian shelf (from ~470 Ma to a period between 375 and 320 Ma, i.e., 95 to 150 million years). 42 3.5 Alternative Interpretations of the Hollins Line Fault System Mies (1992) interprets rocks of the entire Hillabee Greenstone as occurring within a large scale shear zone (“Hillabee mixed zone”) between rocks of the AshlandWedowee belt and Talladega belt (associated with Hollins Line thrusting) based on work mainly in the Hightower and Ross Mountain 7.5’ quadrangles at the northeastern end of the study area (Fig. 1.2). Mies argued that massive greenstones of the Hillabee occurred as discontinuous pods, and that associated mafic phyllites (phyllonites according to Mies) were tectonically derived from intense shearing of the Hillabee under greenschist facies during Alleghanian movement along the Hollins Line and emplacement of the eastern Blue Ridge allochthon. Additionally, Mies argued that silicious muscovite phyllites within the “Hillabee Mixed Zone” resulted from shearing of footwall quartzites and micaceous quartzites within the upper Talladega Group. Based on these lithologic-stratigraphic interpretations, coupled with identification of “mylonitic foliation” in the study area, Mies argued that the Hollins Line fault in southern Cleburne County, Alabama “is a broad ductile shear zone, a few hundred to nearly 1,000 m (500 to 3,000 ft) thick, that binds the Talladega slate belt and the eastern Blue Ridge belt”. Furthermore, Mies argued that the Hollins Line “shear zone” consists of the Hillabee “mixed zone”, (an undifferentiated, tectonically mixed Hillabee Greenstone and Hillabeederived phyllonite), in addition to various shear zone tectonites derived from both the Ashland-Wedowee hanging wall rocks and Talladega belt footwall rocks (Mies, 1992). Higgins and Crawford (2008) compound Mies’ interpretation by claiming that the well mapped stratigraphy of the Butting Ram-Cheaha Quartzite and Erin-Jemison are all part of a “Hillabee fault zone”, which they correlate with Mies’ “Hillabee Mixed Zone”, but expand to include all stratigraphy between the stratigraphic top of the Lay Dam Formation and base of the Ashland-Wedowee belt. The current study contradicts certain aspects of Mies’ interpretation, as well as Higgins and Crawford’s interpretations for the following reasons: 1) Mafic phyllites of the Hillabee Greenstone are not tectonically derived “phyllonites” from discontinuous pods of massive greenstone and, as such, are 43 not indicative of Hollins Line-related shearing, except in the immediate vicinity of the Hollins Line roof, splay, or floor thrusts. Metabasalts commonly consist of both foliated and non-foliated (greenschist vs. greenstone) components (e.g. Catoctin in Loudoun County, Virginia or greenstones of the Eagles Bay assemblage, south central British Columbia) (Southworth et al., 1993; Höy, 1998). Geochemical analyses of massive and foliated greenstones of the Hillabee show little compositional variation between the various metamafic lithologies and suggest that textural differences in Hillabee foliation (non to weakly foliated greenstone vs. foliated mafic phyllite) are most likely the result of selective development of penetrative cleavage or initial differences between basalt and basaltic ash layers sourcing the same magmas (Tull et al., 1978; Tull and Stow, 1980). Additionally, in the central portion of the study area, in the vicinity of Millerville, extensive layers of Hillabee metadacite can be mapped for more than 20 km along strike. Intense shearing of the entire Hillabee would have significantly disrupted primary layering such as this (in addition to strata-bound sulfide zones) and made it impossible to trace such layers for any significant distance. The fact that primary layering within the Hillabee (metadacites and strata-bound sulfide deposits) can be mapped for significant distances at many localities within the Talladega belt makes it unlikely that the entire Hillabee is an intensely sheared, tectonically-derived lithologic unit across its entire strike length as suggested by Mies. 2) Silicious micaceous ‘phyllonites’ supposedly derived from the Talladega belt footwall (as mapped by Mies) are, in actuality, primary compositional layers of Talladega Group stratigraphy. Both the Cheaha Quartzite-Butting Ram Sandstone and overlying Erin Slate-Jemison Chert include micaceous quartzites and quartzose sericite phyllites. Over the length of the Hillabee thrust, as discussed previously, rocks of the Hillabee Greenstone overly the Erin-Jemison. Additionally, interlayered mafic phyllites and dacite units, as well as rare and thin (typically less than 1 meter ) sericite phyllites within the Hillabee have experienced pervasive penetrative foliation and associated deformation during Neoacadian or earliest Alleghanian metamorphism. In the absence of easily 44 identified primary structures, these compositional variations might be confused with tectonically derived phyllonites. While Mies does recognize that the Talladega belt footwall may in fact contain primary bedding in the study area, because no primary structures “such as cross beds, graded bedding, ripple marks, etc.” were observed, no efforts were made by Mies “to distinguish original bedding and gneissocity”. This failure to identify primary compositional layering in both the Hillabee and underlying Talladega Group is partly a function of the nature of Talladega Group stratigaphy within the area in question. The most easily recognized stratigraphic unit within the Talladega belt, the Cheaha Quartzite-Butting Ram sandstone, is typically the prominent ridge-forming rock unit in this region. From the Cheaha Mountain 7.5’ quadrangle to the southwest of Mies’ study area, this stratigraphic unit thins from more than 400 meters thick to less than 10 meters in the Ross Mountain and Hightower quadrangles. Additionally, while the Cheaha-Butting Ram typically consists of interlayered metasandstones and metaconglomerates to the southwest, this same stratigraphy typically lacks coarse conglomeratic layers in the extreme northeast of the study area and is most commonly a fine-grained quartzite or micaceous quartzite likely deposited in more distal parts of the sedimentary basin than equivalent strata to the southwest. Finally, easily recognized strata of the ErinJemison (Cook, 1982; Tull, 1982) common in rocks to the southwest (laterally continuous foliated ‘papery’ quartzites bounded by variably quartzose sericite phyllites, for example) are generally absent in the eastern Ross Mountain and Hightower 7.5’ quadrangles, contributing further to Mies’ interpretation of pervasive ‘phyllonites’ and lack of primary compositional layering. 3) While shear fabrics and related deformational features associated with Hollins Line roof, splay, and floor thrusts are certainly present in the area, a comparative structural analysis of rocks both within and outside of the region adjacent to the Hollins Line shows that many of the deformational features attributed to ductile shearing by Mies are, in fact, associated with syn- and post-metamorphic deformation predating emplacement of the eastern Blue Ridge allochthon along the Hollins Line. If shear fabrics and deformational features of the “Hillabee 45 Mixed Zone”, as interpreted by Mies, were associated with a wide zone (up to 1 km) of Alleghanian Hollins Line shearing, then those deformational fabrics should be distinct from deformational features deeper in the Talladega belt footwall, where clear evidence of compositional layering far from potential Hollins line shearing exists. Stereographic analysis of data from the Ross Mountain, 7.5’ Figure 3.3 - Stereographic analysis of 60 poles to S 0 -S 1 of the Lay Dam formation, Ross Mountain, 7.5' quadrangle, Alabama. quadrangle, including poles to bedding (S0) and metamorphic foliation (S1) in the Lay Dam formation, where primary compositional layering can be easily recognized and pervasive foliation is axial planar to isoclinal folding of primary layering, clearly show polyphase deformation (Fig. 3.3) associated with Neoacadian or earliest Alleghanian metamorphism-deformation and subsequent 46 late Paleozoic cross-folding (Tull, 1984). Because the Lay Dam lies well below the structural level of the Hollins Line ‘shear zone’ proposed by Mies, it should Figure 3.4 - Stereographic analysis of 26 poles to S 1 foliation planes of the Hillabee Greenstone, Ross Mountain, 7.5' quadrangle, Alabama. have a very different structural character than rocks within the “Hillabee Mixed Zone”, which according to Mies, only exhibit structural characteristics of Hollins Line associated Alleghanian shearing and post-metamorphic cross-folding. However, stereographic analysis of Ross Mountain quadrangle poles to metamorphic foliation in the Hillabee (Fig. 3.4) show essentially the same phases of deformation, with rocks of the Hillabee exhibiting predominant northeast striking-southeast dipping beds and foliation resulting from syn-metamorphic isoclinal folding and subsequent cross folding. This same relationship can be observed in stereographic analysis (Figs. 3.5 and 3.6) of bedding and foliation planes within equivalent rocks to the southwest (Cheaha Mountain, 7.5’ quadrangle). 47 4) If the stratigraphy of the Butting Ram-Cheaha Quartzite and Erin-Jemison constituted “siliceous mylonites” of a “Hillabee fault zone”, as interpreted by Higgins and Crawford (2008), this it should be virtually impossible to map stratigraphy along strike of the Talladega belt for any length. Yet, the Butting Ram Sandstone-Cheaha Quartzite, which holds up a prominent topographic ridge over the length of the Talladega belt, can be mapped for virtually the entire length of the belt, albeit with variations in stratigraphic thickness. Additionally, the contacts between the Lay Dam Formation, Butting Ram-Cheaha, and ErinJemison, which are well exposed at a number of locations (e.g. Cheaha Mountain and Bulls Gap), show clear gradational contacts between one stratigraphic package and another indicating that the boundaries between these strata represent concordant sedimentary contacts – not fault bounded thrust sheets of a “Hillabee fault zone”, as proposed by Higgins and Crawford (2008). 3.6 Alternative Interpretations of the Hillabee Thrust McClellan et al. (2005), McClellan (2007), and McClellan et al. (2007) argue that the basal contact (“Hillabee Shear Zone”) between the Hillabee Greenstone and ErinJemison formation, as well as metadacites of the Hillabee volcanic complex structurally above the tectonic base between the Hillabee and underlying Erin-Jemison, preserve mylonites and shear-related fabrics/features associated with the pre- to synmetamorphic Hillabee thrust. McClellan (2007) describes the “Hillabee Shear Zone” as a “pre-to synmetamorphic zone of ductile shearing that variably affected both hanging wall and footwall rocks”, while McClellan et al. (2007) argue that “conditions of shearing are constrained to greenschist facies by mineral assemblages and microstructures in the Hillabee metadacite, in which mylonitic textures are particularly well developed”. The implications of McCleallan et al. (2005), McClellan (2007), and McClellan et al. (2007) are two-fold. Since rocks of the Talladega belt, both Talladega Group and Hillabee Greenstone, experienced only a single progressive dynamothermal 48 Figure 3.5 - Stereographic analysis of 80 poles to S 0 -S 1 planes of the Lay Dam formation, Cheaha Mountain, 7.5' quadrangle, Alabama. Figure 3.6 - Stereographic analysis of 34 poles to S 1 foliation planes of the Hillabee Greenstone, Cheaha Mountain, 7.5' quadrangle, Alabama. 49 metamorphic event dated by McClellan et al. (2007) at 334 – 320 Ma, then the “Hillabee Shear Zone”, occurring under greenschist facies conditions, must be syn-metamorphic in nature. Additionally, because McClellan et al. (2005), McClellan (2007), and McClellan et al. (2007) correlate “mylonitic” fabrics within metadacites of the Hillabee Greenstone to shearing along the basal Hillabee thrust, then ductile emplacement of rocks of the Hillabee terrane must have been associated with a “wide zone of deformation” that extends to at least the structural level of the metadacites within the Hillabee Greenstone hanging wall of the Hillabee thrust, a thickness of >1,800 m (base of Hillabee to base of upper dacite) in areas such as the Millerville, 7.5’ quadrangle This research is not supportive of these two interpretations by McClellan et al. (2005), McClellan (2007), and McClellan et al. (2007) for the following reasons: 1) If rocks of the Hillabee Greenstone terrane were emplaced along a synmetamorphic, ductile shear zone, it would be difficult to preserve the flat-on-flat geometry observed along the basal Hillabee contact across the length and width of the Talladega belt (Tull et al., 2007). Talladega belt lithologies contain typical lower greenschist (chlorite-albite-epidote-actinolite zone) facies mineral assemblages (Winkler, 1976; Bucher and Frey, 1994) and formed at temperatures of 300 - 350°C and pressures of 3-4 kb (Tull et al., 1988; Tull et al., 1998). McClellan et al. (2007) argue that “foliation in the sheared Hillabee rocks is slightly oblique to that in quartzose phyllite of the Talladega Group” along exposures in the vicinity of the West Fork of Hatchett Creek, Millerville, 7.5’ quadrangle (McClellan et al., 2005, 2007), whereas Tull et al. (2007) argue that a hanging wall flat – footwall flat geometry is observed across the length of the Hillabee thrust, and that no hanging wall or footwall stratigraphic cutoffs are observed. If the Hillabee thrust (“shear zone” of McClellan et al. (2007)) were in fact, synmetamorphic, it would imply that emplacement of the Hillabee Greenstone terrane was coeval with stacking of one or more thrust sheets (12 to 15 km) above the basal contact with the Erin-Jemison in order to produce the temperatures and pressures necessary to metamorphose the collective Hillabee and Talladega Group. A single thrust sheet of this thickness (>12 km) would imply that emplacement of the Hillabee was associated with thick-skinned 50 thrusting, which generally includes basement (Coward, 1983). However, no deep-seated basement rocks exist in the Talladega belt, suggesting that thickskinned thrusting as a mechanism of Hillabee emplacement is unlikely. Additionally, a thrust sheet of this thickness would have been likely to result in propagation of other latest Devonian-Mississippian age thrust faults well in advance of the position of the Hillabee thrust, thus producing shallower level thrust faults in the Talladega belt or Alabama foreland sequences. There are no candidates for faults of this age in either the Talladega belt or foreland, and thus it is unlikely that faulting along the Hillabee thrust was associated with emplacement of a 12-15 km thrust nappe atop the amalgamated Talladega belt and concomitant metamorphism (Barineau and Tull, 2001; Barineau et al., 2006; Tull et al., 2007). Duplex stacking of multiple thrust sheets to achieve a cumulative thickness of at least 12 km (Boyer and Elliot, 1982) would cause complex deformation of hanging wall strata and dramatic hanging wall and/or foot wall cutoffs such as those observed in synmetmorphic ductile thrust systems in other locations, e.g. Dalradian belt, Ireland (Jolley, 1996). Because of the extensive flat-on-flat geometry observed in the study area (Tull et al., 2007), it is difficult to envision thrust emplacement of either a single or cumulative >12 km thrust sheet(s) coeval with metamorphism. Consequently, the Hillabee thrust is more reasonably associated with thin-skinned emplacement of the Hillabee Greenstone along a shallow, bedding plane-controlled fault as a single thrust sheet, prior to metamorphism of the amalgamated Hillabee-Talladega Group in the latest Devonian-Mississippian (McClellan et al., 2005; Tull et al., 2007; McClellan et al., 2007). What is not evident, however, is the degree to which thrust emplacement of the Hillabee predates the metamorphic age of the Talladega belt, and it is not difficult to conceive of scenarios in which emplacement of the Hillabee Greenstone immediately preceded metamorphism in the Talladega belt. 2) Shear fabrics observed in metadacites of the Hillabee Greenstones are well above the structural base of the Hillabee terrane and do not record deformation phases associated with the basal Hillabee thrust, but instead are associated with 51 dynamothermal metamorphism which affected the entire Talladega belt, following emplacement of the Hillabee terrane. McClellan et al. (2005), McClellan (2007) and McClellan et al. (2007) argue that metadacites within the Hillabee Greenstone are mylonites and record Hillabee thrusting (“shearing”). Mylonitic textures described by McClellan et al. (2007) include “relict quartz grains” with “undulose extinction and subgrain development, surrounded by smaller dynamically recrystallized grains”, in addition to brittle fracturing of feldspars and hornblende and “growth of syndeformational, aligned acicular actinolite, chlorite, and epidote in fractures between segments”. Abundant quartz grains which compose Hillabee metadacites, identified as “relict quartz grains” and dynamically recrystallized quartz by McClellan et al. (2007), are typical of lower greenschist facies metamorphism, as are brittle fracturing of igneous phenocrysts such as feldspar and hornblende. Similar features are observed in metagraywackes of the Lay Dam Formation (Fig. 3.7), in which ductile quartz grains underwent dynamic recrystallization while detrital feldspar grains were fractured under peak greenschist facies metamorphism. “Mylonitic” textures observed in Hillabee metadacites are not indicative of fault zone mylonites (McClellan, 2007; McClellan et al., 2007), which are the result of intense grain size reduction, but instead formed as a result of strain partitioning in the mechanically weaker dacite units adjacent to primary basaltic lithologies of the Hillabee Greenstone during metamorphism (Tull et al., 1998). Protoliths for metadacites of the Hillabee have been interpreted as pyroclastic ash flow deposits (ignimbrites) and as such were originally composed of significant quantities of volcanic glass as well as crystal and lithic components (Tull et al., 1998). The “mylonitic” features associated with metadacite quartz grains observed by McClellan et al. (2007) are the result of dynamic recrystallization of volcanic glass under greenschist facies conditions, and thus represent an increase in the overall grain size of the original ignimbrite matrix – opposite of what would be expected in a fault zone mylonite. Since the Hillabee thrust is most reasonably interpreted as a premetamorphic, thin-skinned thrust, shearing along the basal faulted contact would 52 need to be distinguished from deformation associated with Neo-Acadian or earliest Alleghanian greenschist facies metamorphism observed throughout the Talladega belt. Hillabee metadacites record the same deformation phases as rocks of the underlying Talladega, Sylacauga, and Kahatchee Mountain Groups, including brittle deformation of pre-metamorphic feldspars, mineral stretching lineations, and asymmetric kink folding of predominant foliation in proximity to the post-metamorphic Hollins Line fault system (Tull et al., 1998; Tull et al., 2007). Because the Hillabee cannot be native to the Laurentian shelf, it must have originated a considerable distance from the Talladega Group shelf strata and transport of this thrust sheet may be on the order of 10’s of km. If emplacement of the Hillabee terrane was associated with a greenschist facies ductile shear zone affecting not only the structurally lower greenstones, but additionally, metadacites located 1800 or more meters above the location of the basal Hillabee thrust, then the cumulative shear associated with a displacement of this magnitude should have dismembered any internal stratigraphy within the “Hillabee Shear Zone”. Because both massive stratabound sulfide deposits, as well as extensive metadacite layers, can be mapped continuously across much of the strike of the Hillabee terrane, it would be difficult to characterize the fault at the contact between the Hillabee and Talladega Group as anything other than a discrete shear zone which did not penetrate significantly into either the hanging or footwall. As such, use of the term “Hillabee Shear Zone”, associated with interpretations of the entire Hillabee Greenstone as a fault zone (Mies, 1992; McClellan et al., 2007), should be abandoned in favor of ‘Hillabee thrust’ until unique deformation phases distinct from those associated with regional metamorphism clearly identify a continuous zone of shearing as opposed to the discrete fault zone currently recognized. 3.7 Alternative Interpretations, Shear Strain, and Metamorphism In alternative interpretations of both the Hollins Line fault system and Hillabee thrust, much the shear strain attributed to faulting is not unique to rocks adjacent to the 53 respective faults. The presence of a discreet shear zone implies that rocks within and adjacent to the shear zone are more highly strained than rocks outside of and far from the shear zone (Davis and Reynolds, 1996). In the case of the Hollins Line fault system, rocks adjacent to the roof, floor, and splay thrusts themselves contain shear fabrics (S-C, asymmetric shear bands, etc.) which cut the regional synmetamorphic fabrics and that are unique to the fault zone. However, much of the “phyllonitic” fabric (foliation-regionally developed slaty cleavage) argued for by Mies in the “Hillabee Mixed Zone” are present in rocks far from the fault itself, consistent with an interpretation of these fabrics as phyllitic, synmetamorphic penetrative foliation – not “phyllonitic” fabrics derived from shearing of the Hillabee Greenstone. Metamorphic fabrics within the Hillabee away from the Hollins Line fault system are coaxial and coplanar with metamorphic fabrics in the Erin-Jemison, Cheaha Quartzite, and Lay Dam formations, and formed during progressive strain associated with metamorphism of the entire Talladega belt. If shear strain within the “Hillabee Mixed Zone” formed as a result of Hollins Line thrusting (Mies, 1992), then the entire Talladega belt would have to be considered a fault zone, as the metamorphic fabrics and deformational features present in the Hillabee are ubiquitous to the entire Talladega belt (Figs. 3.3 – 3.6) and its wellmapped, laterally continuous stratigraphy. Although Higgins and Crawford (2008) essentially claim that much of the Talladega belt is, indeed, part of their proposed “Hillabee fault zone”, evidence presented previously (sections 2.1 and 3.5) makes their claim untenable. Similarly, if the arguments presented by McClellan et al. (2007) suggest that metadacites of the Hillabee formed during emplacement of the Hillabee metavolcanic suite along a Hillabee “shear zone”, then deformation features within the entire Talladega belt must have formed during that same event, as these “mylonitic” shear fabrics are coaxial with metamorphic fabrics throughout the Talladega belt and formed under the same lower greenschist facies event. As above, this would imply that the entire Talladega belt is a shear zone associated with emplacement of the Hillabee Greenstone. Arguments that mylonitic fabrics in Hillabee metadacites are fault zone “mylonites” are further compounded by interpretations by McClellan et al. (2007) that metamorphism in the Talladega belt was associated with emplacement of the eastern 54 Blue Ridge (Ashland-Wedowee belt) atop the Talladega belt and accompanying downward heating (Kish, 1990). McClellan et al. (2007) argue that the Talladega belt (combined Kahatchee Mountain Group – Sylacauga Marble Group – Talladega Group – Hillabee Greenstone) experienced only a single prograde lower greenschist facies metamorphic event based on theirs and previous workers isotopic and petrologic work in the Talladega belt (Tull et al. 1978; Tull and Stow 1980; Mies 1992). If the “Hillbee shear zone” of McClellan et al. (2007) occurred under greenschist facies conditions, as they argue, then this must have been the same greenschist facies event caused by emplacement of the eastern Blue Ridge terrane, which implies that the Hillabee Greenstone was emplaced atop the Talladega Group at the same time that Hollins Line thrusting accommodated emplacement of the Ashland-Wedowee belt. This is an especially difficult scenario to envision as the Hollins Line fault system is clearly post metamorphic and no evidence for inverted metamorphic gradients are present in the Talladega belt foot wall or Ashland-Wedowee belt hanging wall along the Hollins Line fault system. Thus, the alternative interpretations by Mies (1992) and McClellan et al. (2007) for the Hollins Line fault system and Hillabee thrust are not supported by this research. Shear fabrics in L-S tectonites of both greenstones and metadacites of the Hillabee can be largely attributed to flattening strain and unidirectional stretching associated with lower greenschist facies metamorphism, except in the immediate vicinity of the Hollins Line fault system. 55 CHAPTER 4 TACONIC OROGENY OF THE SOUTHERN APPALACHIANS 4.1 Arc-continent collisions In the western Pacific, Neogene arc-continent collisions have been used as analogues for ancient collisional orogens. In Timor, oblique subduction of the Australian passive margin beneath the Banda arc between ~3.5 Ma and 2 Ma resulted in the ongoing Banda orogen (Harris et al., 1998; Audley Charles, 2004; Harris, 2006). Farther to the east, in Papua New Guinea, northward subduction of the Australian passive margin beneath an older Melanesian arc terrane resulted in obduction of the Irian ophiolite and collisional orogenesis at ~12 Ma (Cloos et al., 2005). In Taiwan, subduction of the Eurasian passive margin beneath the Luzon arc began at ~8.5 Ma, progressing from north to south (Huang et al., 2006). In each of the three collisions, detailed study of the timing of tectonic events and the arrangement of lithotectonic belts has revealed a number of commonalities: A) separation of the continental margin and arc terrane by an intervening accretionary prism, B) metamorphism and volcanism which precedes or is coeval with the beginning of collisional orogenesis, C) development of an orogenic clastic wedge coeval with subduction of the continental slope-rise and arc-continent collision, D) followed by cessation of arc volcanism. In the region of the Java trench/Timor trough, Australian margin material is actively accreting to the overriding Eurasian and Bismarck plates in a ~100 km wide belt more than 1000 km in length. Subaerial rocks of the Bobonaro mélange on the island of Timor occur as discontinuous bodies, some greater than 15X50 km, over the 500+ km length of the island, and have been tectonically emplaced against forearc rocks of the Banda terrane (Harris et al., 1998). In the Central Range of western New Guinea, the >500 km long Derewo Metamorphic belt ranges from 10 to 30 km in width and represents offscraped, accreted, and partially subducted Australian slope-rise sediments of accretionary prism (Cloos et al., 2005; Warren and Cloos, 2007) The Kenting and Lichi mélanges in southeastern Taiwan, at the leading edge of the accretionary prism, represent subduction and collision related complexes (Chang et al., 56 2000; Huang et al., 2006) and partially comprise an accretionary prism complex 35 to 50 km in width spanning the length of the entire island. In all three cases, low temperature-high pressure (LT-HP) rocks are found as isolated, discontinuous bodies within chaotic mélange deposits scraped off the oceanic and distal continental margin during subduction. The location of accretionary prism complexes between the arc on the overriding plate and subducting continental margin deposits in arc-continent collisions is not surprising, considering that subduction of the leading edge of a continent results in extensive frontal accretion of materials scraped off the downgoing slab (subduction accretion), especially where sediment thicknesses exceed 1 km (von Huene and Scholl, 1991; Le Pichon et al., 1993). Therefore, the presence of a thick accretionary prism complex, complete with mélange and LT-HP rocks, between rocks of a continental foreland basin and obducted arc terrane should be a hallmark of arccontinent collision. These same modern arc-continent collisions also have similar histories regarding the timing relationships between arc volcanism, subduction, age of metamorphism, uplift, and foreland basin clastic wedge deposits. In New Guinea, a subduction reversal associated with jamming of the Outer Melanesion subduction zone by the Ontong-Java plateau resulted in initiation of northward dipping subduction of oceanic lithosphere attached to the Australian plate just north of the Australian continental margin between 30 and 25 Ma (Cloos et al., 2005). Metamorphism of continental margin sediment deposited on subducting oceanic lithosphere began at 28 Ma, prior to arrival of continental basement at the leading edge of the Australian margin (Warren and Cloos, 2007). Introduction of the thick pile of sediment at the slope-rise into the subduction zone resulted in formation and subaerial exposure of a thick (>20 km) accretionary prism, coeval with clastic wedge deposition in the foreland of the Australian continental margin at ~12 Ma. This “precollisional complex”, developed prior to jamming of the subduction zone by Australian crystalline basement between 8 and 4 Ma and the transition from thin-skinned to thick-skinned deformation, synchronous with cessation of subduction (Cloos et al., 2005). In Taiwan, subduction of oceanic crust in the South China Sea resulted in Luzon Arc volcanism at 16-15 Ma (Yang et al., 1995; Huang et al., 2006). Metamorphism prior 57 to the onset of arc-continent collision is recorded in LP-HT rocks of the accretionary prism (8-14 Ma), which formed during intraoceanic subduction (Jahn et al., 1981), while peak metamorphic ages from basement yields ages between 7.7 and 6.4 Ma (Lo and Onstott, 1995; Beyssac et al., 2007). Coeval clastic wedge deposition and exhumation of the accretionary prism occurred between 6 and 5 Ma (Dorsey, 1988; Lee et al., 2006; Huang et al., 2006), concomitant with waning of arc volcanism during oblique collision and cessation of subduction (Yang et al., 1988; Lo et al., 1994, Huang et al., 2006). In Timor, subduction of the oceanic leading edge of the Australian plate beneath the Banda Arc terrane began between 15 and 10 Ma, roughly coeval with arc volcanism at 12 Ma (Abbot and Chamalaun, 1981), followed by subduction and metamorphism of Australian plate sedimentary rocks between 8 and 5 Ma (Berry and McDougal, 1986; Audley-Charles, 2004; Harris, 2006). Subduction of the leading edge of the Australian slope-rise began prior to 5 Ma, with jamming of the subduction zone at 3.5 to 2 Ma, coeval with waning of arc volcanism, uplift of the accretionary prism, and clastic wedge deposition in the Australian foreland basin (Audley-Charles, 2004). Rollback of the subducting Australian plate prior to collisional orogenesis resulted in backarc spreading (12.5 to 3.5 Ma) in the Banda Sea north of the Banda Arc (Honthaas et al., 1998; Hinschberger et al., 2005; Harris, 2006). In all three settings, volcanism and subduction zone metamorphism associated with subduction of intraoceanic and continental slope-rise sediments precedes collisional orogenesis by as much as 13 m.y., with arc-continent collision marked by exhumation of an extensive accretionary prism, deposition of a clastic wedge in the foreland, and cessation of subduction and arc volcanism. The relative timing of these events, coupled with the arrangement of lithotectonic belts should help in identifying ancient arc-continent collisions in the rock record. 4.2 Northern Appalachian Geology The type locality of the Taconic orogeny in New England and farther to the north, in the Canadian Appalachians, has been attributed to subduction of the Paleozoic North American plate (Laurentia) beneath exotic island arcs or arc terranes built on peri- 58 Laurentian blocks separated from the Paleozoic continental margin by a narrow seaway (Stanley and Ratcliffe, 1985; Karabinos et al., 1998; van Staal et al., 2007). In both the New England and Canadian segments of the Taconic Appalachians a model of arc obduction associated with eastward subduction of Laurentia is argued for based on current geometries of lithotectonic belts and the chronologic, stratigraphic, and geochemical characteristics of the rocks within these belts. The Taconic orogenic belt in New England consists of a western sequence of deformed slope-rise sedimentary deposits which formed along the Iapetan rifted margin of Laurentia (Stanley and Ratcliffe, 1985; Lavoie et al., 2003). These “Humber Zone” rocks are separated from rocks of an accreted arc-forearc terrane by rocks of the RoweHawley belt (Dunnage Zone), interpreted as having formed in an accretionary prism between the Ordovician Laurentian margin and a peri-Laurentian arc terrane (Shelburne Falls Arc, Moretown formation forearc) (Karabinos et al., 1998; Kim et al., 2003). 505 to 468 Ma LT-HP metamorphism is recorded in rocks of the Dunnage Zone from Vermont (Laird et al., 1984), encompassing a period of suprasubduction volcanism (485 to 470 Ma) in the peri-Laurentian Shelburne Falls arc (Karabinos et al., 1998). Peak metamorphism in New England at ~466 Ma (Sutter et al., 1985) was followed by clastic wedge deposition beginning in mid-Caradoc times (Ettensohn, 2004), or ~455 Ma (Gradstein et al., 2004). To the northeast, in the Appalachians of Quebec and Newfoundland, 468-461 Ma post Taconic cooling is recorded in 40Ar/39Ar white mica ages from the internal Humber Zone (Castonguay et al., 1997) and St. Daniel mélange of the Dunnage Zone (Schroetter et al., 2006), while accretionary wedge rocks in northern Newfoundland record ~495 Ma metamorphism (Jamieson, 1988; van Staal et al., 2007). Suprasubduction zone rocks of the Notre Dame arc have a complex history, but earliest phases of igneous activity range from 490 to 477 Ma (van Staal et al., 2007). On the Gaspe peninsula of Quebec, loading of the Laurentian margin and deposition of the Taconic clastic wedge began some time after 461 Ma (early Caradoc time) (Prave et al., 2000), while to the north in Newfoundland, flysch deposits of the Taconic clastic wedge are slightly older than 462 Ma (mid to late Whiterockian) (Knight et al., 1991). 59 The sequence of events and relationships between lithotectonic belts in the northern Appalachians is consistent with observations in modern arc-continent collisional orogenies such as New Guinea, Taiwan, and Timor. First, rocks of the Laurentian Humber Zone are separated from rocks of accreted are terranes by extensive belts interpreted as accretionary prism rocks tectonically eroded from the subducting plate, complete with preserved HP-LT rocks (Thompson and Thompson, 2003; van Staal et al., 2007). Suprasubduction volcanism of the accreted arc, like preserved LT-HP metamorphic rocks, predate peak metamorphism, onset of collisional orogenesis, and deposition of the Taconic clastic wedge. These features, along with the presence of abundant cratonward-verging Taconic thrusts, strongly support subduction of the early Paleozoic Laurentian margin beneath an arc terrane, similar to that observed in modern arc-continent collisions. 4.3 Southern Appalachian Geology Talladega belt (Western Blue Ridge) The Middle Ordovician Hillabee Greenstone , a bimodal sequence of tholeiitic metabasalt and subordinate interstratified calc-akaline metadacite/rhyolite, was emplaced atop the lower greenschist facies Talladega belt along a pre-metamorphic fault (Fig. 4.1) in the southernmost Appalachians (Tull et al., 2007; McClellan et al., 2007). Metasediments of the Talladega belt range from Early Cambrian Chilhoweeequivalent drift strata (Kahatchee Mountain Group) to Devonian-lowermost Mississipian(?) cover strata of the Talladega Group and represent the most outboard shelf sequences of the Laurentian margin preserved in this part of the orogen (Tull, 1982, 2002; Tull et al., 1988; Gastaldo, 1995). Analysis of Hillabee Greenstone mafic rocks using various trace element tectonic discrimination diagrams (Tull et al., 1978; Tull and Stow, 1980; Tull et al., 1998), as well as rare-earth element (REE) patterns of HG mafic and felsic units (Holm-Denoma, 2006; Das, 2006; Tull et al., 2007), suggest formation within a backarc environment. Additionally, εNd and zircon εHf values from Hillabee Greenstone metadacites indicate involvement of ~1.1 Ga (Grenville) continental crust in its genesis (Holm-Denoma, 2006; Tull et al., 2007). Finally, a lack of 60 significant deformation during fault emplacement of the Hillabee onto rocks of the Talladega belt Laurentian shelf argues for limited transport of the Hillabee backarc rocks and suggests an origin near the Middle Ordovician Laurentian continental hinge zone on the attenuated continental margin (Holm-Denoma, 2006; Tull et al., 2007). Zircon analyses from HG felsic units yield ages of 470±4 Ma (206Pb/238U –McClellan et al., 2007) and 468 ± 3 Ma (206Pb/208 – Tull et al., 2007). Ashland-Wedowee belt (Eastern Blue Ridge) The >10 km thick, middle to upper amphibolite facies Ashland-Wedowee belt of Alabama and Georgia (Fig. 4.1) consists dominantly of metapelites with minor orthoamphibolite, quartzite, and metagraywacke, and is interpreted as having formed along a rifted continental margin where rift sediments and continental slope deposits were interlayered with ocean floor basalts or intruded by mafic sills (Muangnoicharoen, 1975; Tull, 1978; Thomas et al., 1980; Whittington, 1982; Stow et al., 1984; Drummond et al., 1988; Drummond, 1986; Allison, 1992; Tull et al, 2007). Inheritance of Mesoproterozoic (Grenville) zircons in synkinematic igneous intrusions (Rockford granitoids and trondhjemitic dikes) of the Ashland-Wedowee belt suggest that riftrelated sediments of the Ashland Supergroup and Wedowee Group sourced Laurentian basement (Drummond, 1986; Drummond et al., 1988; Allison, 1992; Steltenpohl et al., 2005). The palinspastic position of the Ashland-Wedowee belt (outboard of the Suwanne-Wiggins suture) (Ferrill and Thomas, 1988; Thomas, 2004), absence of basement rocks, great stratigraphic thickness, lithologic associations (tholeiitic rift/ocean floor basalts intercalated with deep water turbiditic metasediments), and tectonic position above shelf rocks of the Talladega belt Laurentian margin rocks suggest that the bulk of the Ashland-Wedowee belt formed at or near the Laurentian rifted margin during and following Rodinian rifting (Tull, 1978; Thomas et al.,1980; Stow et al., 1984; Drummond et al.,1988; 1994; 1997; Tull et al., 2007). This interpretation is further supported by detrital zircon and whole-rock Nd isotopic work, which suggests that parts of the eastern Blue Ridge and Inner Piedmont of Alabama, Georgia, North Carolina, and South Carolina have Laurentian affinities, forming along the attenuated margin of that continent during Iapetan rifting (Bream et al, 2004; Das, 2006; Hatcher et al, 2007). 61 Following Neoproterozoic rifting and formation of the Ashland Supergroup and Wedowee Group rocks, the Ashland-Wedowee was intruded by Cambrian(?) and Ordovician plutons (Fig. 4.1). Rocks of the Elkahatchee Quartz Diorite and ZanaKowaliga gneiss, a tonalite-trondhjemite-granodiorite (TTG) suite, indicates suprasubduction activity, as post-Archean TTG rocks are restricted to arc settings (Condie, 2005). Trace element discrimination of the Elkahatchee and Zana-Kowaliga gneiss overwhelmingly plot in the volcanic arc granite field of Pearce et al. (1984), further supporting interpretation of an arc setting for these plutons (Russell et al, 1987; Drummond et al, 1994; Holm-Denoma, 2006), similar to that of the early Late Ordovician (458 Ma) trondhjemitic Villa Rica Gneiss in the more northeastern Georgia segment of the Ashland-Wedowee belt (Thomas, 2001). Rocks at the stratigraphic/structural top of the Ashland-Wedowee belt in Alabama (Emuckfaw Formation) have been correlated with the New Georgia Group in northwestern Georgia (Hurst, 1973; McConnell and Abrams, 1984; Tull, 1987; HolmDenoma, 2006) and can be traced through the Cartersville transverse zone (Tull and Holm, 2005) and into the Dahlonega Gold belt (Fig. 4.1) (Holm-Denoma, 2006). Rocks of the New Georgia Group, like the Ashland-Wedowee in Alabama, have been intruded by a number of Ordovician plutons (Austell, Villa Rica, and Sand Hill plutons) ranging in age from ~460 Ma to 430 Ma (Higgins et al., 1997; Thomas, 2001). As with the Elkahatchee and Zana-Kowaliga plutons, the 458±3 Villa Rica Gneiss (Thomas, 2001) plots overwhelmingly in the volcanic arc field of Pearce et al. (1984), with REE data falling in the volcanic arc granite field on a Ta-Yb discriminant diagram (Spell and Norrell, 1990; Holm-Denoma, 2006). Unlike the TTG suite, I-type Elkahatchee-Zana-Kowaliga-Villa Rica plutons, the Austell-Sand Hill gneiss has strong S-type affinities and probably represent recycling of Mesoproterozoic Grenville crust in attenuated continental crust (Holm-Denoma, 2006). A similar granite, the ~450 Ma Mulberry Rock Gneiss, also has S-type affinities and is interpreted as being genetically related to the Austell-Sand Hill, although it lies within a structurally distinct panel (Fig. 4.1) above footwall rock of the Talladega belt and hanging-wall rocks of the Ashland-Wedowee belt (Holm and Das, 2005; Holm-Denoma, 2006; Mueller, personal communication). 62 The numerous Cambrian(?) to Ordovician dates determined for granitoid bodies in the eastern Blue Ridge of Alabama and Georgia suggest they were emplaced between latest Cambrian(?) and Late Ordovician (Russell et al., 1987; Drummond et al., 1994; Holm-Denoma, 2006). As metasediments of the Ashland Supergroup and Wedowee Group can be best interpreted as representing the rifted margin sedimentary prism along the Laurentian continent, then Cambrian(?) -Ordovician suprasubduction plutonism within the eastern Blue Ridge indicates the existence of an early Paleozoic, Andean-type margin along this segment of Laurentia (Drummond et al., 1994; Hibbard, 2000; Tull et al., 2007). In the northeasternmost portion of the Ashland-Wedowee belt, bimodal volcanic rocks of the ~460 Ma Pumpkinvine Creek Formation (Thomas, 2001; Holm and Das, 2005; Holm-Denoma, 2006; Das, 2006) have been genetically linked with rocks of both the Dahlonega Gold belt (McConnell and Abrams, 1984; German, 1985; 1989) and New Georgia Group in the Ashland-Wedowee belt of Georgia (Holm-Denoma, 2006). Geochemically, amphibolites and felsic gneiss (Galt’s Ferry Gneiss Member) of the Pumpkinvine Creek exhibit suprasubduction characteristics, with felsic Galt’s Ferry Gneiss rocks plotting in the volcanic arc granite field of Pearce et al. (1984), and REE analysis of Pumpkinvine Creek amphibolites suggesting formation in a backarc setting (Holm and Das, 2005; Holm-Denoma, 2006; Das, 2006). Additionally, detrital zircon analyses of metasediments intercalated with the Pumpkinvine Creek show that they sourced Mesoproterozoic Grenville crust, as they contain abundant ~1.0 to 1.2 Ga zircon grains (Holm and Das, 2005; Das, 2006). It is likely that rocks of the Pumpkinvine Creek are stratigraphically related to metasediments of the Canton and Univeter Formations, although structural-stratigraphic relationships are complex due to deformation associated with Paleozoic orogenesis (McConnell and Abrams, 1984; German, 1985; 1989, Holm-Denoma, 2006). If the Pumpkinvine Creek is in stratigraphic contact with rocks of the Ashland-Wedowee belt, this would indicate that the stratigraphically highest sections of the Ashland-Wedowee belt (Emuckfaw-CantonUniveter Formations) are Ordovician in age and genetically linked to an Ordovician backarc forming proximal to a continental margin containing Grenville basement (Laurentia?) (Holm-Denoma, 2006; Tull et al., 2007). 63 Arc-Continent Collisional Orogenesis – Talladega/Ashland-Wedowee Belts Like the northern Appalachians of New England, Quebec, and Newfoundland, volcanism, metamorphism, deformation, and deposition of foreland basin clastic wedge deposits in the southern Appalachians have been attributed to collision between the subducting Laurentian margin and an intraoceanic or peri-Laurentian arc on the overriding plate in the Early to Middle Ordovician (Shanmugan and Lash, 1982; Hatcher, 1987; Spell and Norrell, 1990; Hatcher et al., 2007). However, unlike the northern Appalachians, the timing of events and assembly of lithotectonic belts in the southernmost Appalachians of Georgia and Alabama are not supportive of this tectonic model. In contrast to the northern Appalachians and modern analogues in New Guinea, Taiwan and Timor, the geometry of assembled terranes in the Alabama-Georgia eastern and western Blue Ridge is not diagnostic of an arc-continent collisional setting. Rocks of the Laurentian margin (Talladega belt) are not separated from rocks of an obducted arc by an accretionary prism. Instead, the lower greenschist facies Talladega belt consists of Paleozoic Laurentian shelf stratigraphy separated from rocks of a backarc volcanic suite (Hillabee Greenstone) across a thin-skinned thrust fault that wasn’t active until the Late Devonian-Early Mississippian (Tull et al., 2007; McClellan et al., 2007). This same relationship is observed in the vicinity of the Pumpkinvine Creek formation in Georgia at the boundary between the eastern and western Blue Ridge. Rocks of the Laurentian shelf in the foot wall of the Allatoona fault are in tectonic contact with rocks of a bimodal volcanic suite (Pumpkinvine Creek and Galt’s Ferry Gneiss) in the hanging wall (Spell and Norrell, 1990; Holm-Denoma, 2006). Unlike rocks of the northern Appalachians, rocks of the southernmost Appalachians in Alabama-Georgia do not include the typical arc-accretionary prism-continent geometry observed in modern arc-continent collisions. In fact, there are no reasonable candidates for a Taconic accretionary prism complex in the Blue Ridge of Alabama-Georgia. The well documented shelf strata of the Talladega belt experienced, at most, mild deformation prior to deposition of the Silurian-Lower Devonian Lay Dam formation (Tull, 1998), while the Hillabee 64 Greenstone, which must have lain outboard of the Talladega Group shelf strata, remained virtually undeformed prior to and during its tectonic emplacement above strata of the upper Talladega Group (Tull et al., 2007). Because the Talladega belt and its Cambrian-Ordovician stratigraphy (Kahatchee Mountain and Sylacauga Marble Groups) must restore to at least the present location of the Pine Mountain belt and likely formed at or just beyond the Laurentian continental hinge zone (Thomas, 2004), it would have been in an ideal position to document Ordovician deformation associated with Taconic arc obduction. If the Hillabee Greenstone and Pumpkinvine Creek backarc metavolcanic rocks, as suggested by some workers (Spell and Norrell, 1990; Hatcher et al., 2007, McClellan et al., 2007), represent the obducted Taconic arc, they do not record any evidence of Taconic deformation or metamorphism and it would have been necessary to tectonically transport both of these terranes across the Taconic arc proper, forearc, accretionary prism, and Laurentian slope-rise, to place them in their current position against rocks of the Laurentian shelf , without being significantly deformed (Tull et al., 2007). The tectonic and kinematic complexities of a scenario such as this effectively preclude it as an option and suggest that the Hillabee was not emplaced as an obducted arc terrane from an opposing plate (exotic or peri-Laurentian), but instead must have formed on the same plate as rocks of the Talladega belt (Laurentia). This is supported by geochemical analysis of felsic units in both the Hillabee and Pumpkinvine Creek, which indicate the involvement of Laurentian basement in magma genesis (Holm-Denoma, 2006; Das, 2006; Tull et al., 2007). If the Laurentian margin slope-rise had been subducted beneath an exotic or peri-Laurentian arc, tectonic imbrication should have produced a thick Ordovician accretionary prism subduction complex similar to that observed at modern A-type subduction boundaries (Clift and Vannucci, 2004). The Ashland-Wedowee belt, which contains strata of the Neoproterozoic and younger Laurentian margin slope-rise, and which should have experienced tectonism associated with Taconic subduction of the Laurentian margin, does not contain the disrupted stratigraphy, exotic blocks, or record of LT-HP metamorphism diagnostic of accretionary prisms (Raymond, 1984; Cowan et al., 1985), and experienced a single middle to upper amphibolite facies, Late Devonian prograde metamorphic event (Russell, 1978; Russell et al., 1987; Moore et al., 1987; 65 Steltenpohl et al., 2005; Tull et al., 2007). Additionally, because the Ashland-Wedowee belt has been intruded by a number of Cambrian(?) – Ordovician suprasubduction granitoid bodies, it could not have acted both as an accretionary prism and host rock for suprasubduction related magma bodies in an evolving Taconic arc, which would have intruded the overriding plate. Again, geologic constraints on the nature, origin, and evolution of the Hillabee Greenstone and Ashland-Wedowee belt make interpretations of Taconic arc obduction at the location of the Alabama promontory in the southern Appalachians an untenable model. Arc-Continent Collisional Orogenesis – GA/NC/SC/TN Blue Ridge Although some workers (Spell and Norrell, 1990; Thomas et al., 2001b; McClellan et al., 2007; Hatcher et al., 2007) have proposed that backarc rocks of the Hillabee Greenstone and Pumpkinvine Creek formation were obducted onto the Laurentian margin as part of an extensive Taconic arc complex, the kinematic complexities of obducting these backarc terranes onto the Laurentian shelf, coupled with the structural and geochronological constraints outlined above, make this a difficult scenario to envision. However, in the southern Appalachians of Georgia, North Carolina, South Carolina, and Tennessee, Middle Ordovician (Taconian) metamorphism (Moecher et al., 2004; Corrie and Kohn, 2007), Early and Middle Ordovician suprasubduction igneous activity (Thomas, 2001; Settles, 2002; Meschter-McDowell et al., 2002; Bream, 2003; Miller et al., 2006) , and a Middle Ordovician synorogenic clastic wedge (Thomas, 2004; Ettensohn, 2004; Bayona and Thomas, 2006), suggest a Middle Ordovician tectonic event which affected the southern Appalachians to the northeast of the palinspastic location of the Talladega and Ashland-Wedowee belts of AlabamaGeorgia. In northern Georgia, western North Carolina and eastern Tennessee, rocks of the western Blue Ridge have been correlated with those of the Talladega belt based on stratigraphic, paleontologic, and geochemical constraints (Crawford and Cressler, 1982; Cressler and Crawford, 1976; McConnell and Abrams, 1984; McConnell and Costello, 1982; Unrug et al., 2000; Tull and Guthrie, 1985; Tull and Holm; 2005; Holm-Denoma, 2006; Repetski et al., 2006), and are considered part of the same allochthon. Northeast 66 of the Cartersville transverse zone (Tull and Holm, 2005), rift-related rocks of the Ocoee Supergroup unconformably overly Grenville aged crystalline basement (Odom et al., 1973; Heatherington et al., 1996). These rocks are in turn overlain by a passive margin sequence of Cambrian to Earliest Mississippian(?) shelf strata which formed on the Iapetan-facing, drifting margin of Laurentia. Drift related, Chilhowee and younger strata in the western Blue Ridge of Georgia-North Carolina-Tennesse correlate with similar rocks of the Talladega belt, although extensive Ocoee rift-related strata are not found south of the Cartersville transverse zone (Alabama promontory) due to its upper plate configuration inherited during Neoproterozic rifting of Rodinia (Thomas, 1993; Tull and Holm, 2005). Metamorphic ages from rocks of greenschist to granulite facies in the western and central Blue Ridge suggest peak metamorphic ages of 450 to 460 Ma (Moecher et al., 2004; Moecher et al., 2005; Corrie and Kohn, 2007; Merschat et al., 2008), coeval with rare eclogite facies metamorphism in rocks of northwestern North Carolina (Miller et al., 2000; Page et al., 2003; 2005; Moecher et al, 2005; Miller et al., 2006). Ordovician igneous rocks of a volcanic arc proper are largely missing from the southern Appalachians (Thomas et al., 2001a; Meschter-McDowell et al., 2002; Tull et al., 2007), but instead are dominated by ~480 to 460 Ma bimodal and mafic-ultramafic complexes of mid-ocean ridge basalt (MORB) to backarc basin affinity (Spell and Norrell, 1990; Berger et al., 2001; Settles, 2002; Ryan et al., 2005; Swanson et al., 2005; HolmDenoma, 2006). These have been interpreted as ophiolitic ocean floor and suprasubduction complexes accreted to the Laurentian margin during Taconic orogenesis (Spell and Norell, 1990; Berger et al., 2001; Settles, 2002; Hatcher et al., 2007). Finally, the synorogenic Blount clastic wedge, which sourced Taconic highlands southeast of the palinspastic location of the southern Appalachian Ordovician foreland fold and thrust belt, formed during the late Llanvirn (just prior to 464 Ma) and sourced Laurentian provenance rocks, but not rocks of an accreted arc terrane (Finney et al., 1996; Ettensohn, 2004; Bayona and Thomas, 2006; Hatcher et al., 2007). The timing of events in the southern Appalachians north of the Cartersville transverse zone in Georgia is loosely compatible with arc-continent collision in the Middle Ordovician when compared with modern analogues in New Guinea, Taiwan, and 67 Timor. Suprasubduction rocks as old as 482 Ma (Dahlonega Gold belt Cane Creek Gneiss of the Sally Free Mafic Complex) predate the metamorphic peak (~460 to 450 Ma), which roughly coincides with the onset of synorogenic sedimention of the Blount clastic wedge, just prior to 464 Ma(Settles, 2002; Bream, 2003; Moecher et al., 2004; Ettensohn, 2004; Corrie and Kohn, 2007). Relict eclogite facies metamorphism in the Tugaloo terrane (Hatcher et al., 2007) of north Carolina has been dated at 459 Ma, but this age may reflect zircon growth during retrograde decompression (Miller et al., 2006). There are, however, some notable and important issues of timing that are not generally compatible with an arc-continent collision model for the Taconic orogeny in the southern Appalachians. These include the presence of 463 Ma bimodal backarc volcanics of the Pumpkinvine Creek formation (Galt’s Ferry Gneiss felsics), 466 Ma Barlow Gneiss, and 468 Ma Lake Burton Complex; the 458 Ma trondhjemitic Villa Rica Gneiss; a 456 Ma age for a biotite-tonalite phase of the Persimmon Creek Gneiss; 454 to 448 Deicke and Millbrig k-bentonites of the Appalachian foreland basin; and the ~444 Ma Kennesaw granitoid of the Laura Lake mafic-ultramafic complex (Tucker, 1992; Thomas, 2001; Min et al., 2001; Meschter-McDowell, 2002; Bream, 2003; Miller et al., 2006; Holm-Denoma, 2006). Middle Ordovician (late Llanvirn) synorogenic loading of the Laurentian margin has been inferred from the age of the Blount clastic wedge in the southern Appalachians (Ettensohn, 2004). In an arc-continent collisional model, this would suggest uplift of the accretionary prism and the onset of collisional orogenesis prior to 462 Ma. In New Guinea, synorogenic foreland deposition preceded cessation of slab breakoff/delamination related volcanism in the New Guinea Highlands, and the end of subduction, by less than 10 m.y. (Cloos et al., 2005), while in Taiwan and Timor, the delay between synorogenic sedimentation in the foreland basin and cessation of arc volcanism and subduction was less than 6 m.y. (Audley-Charles, 2004; Huang et al., 2006). However, in the southern Appalachians, arc volcanism seems to have continued unabated for >20 m.y. (>464 Ma Blount wedge to ~444 Ma Kennesaw granitoid), following the onset of inferred foreland loading by Taconic thrust sheets. Even more puzzling is the fact that backarc volcanism was also active (Galt’s Ferry ~463 Ma, Barlow Gneiss ~466 Ma, Lake Burton ~468 Ma, Hillabee dacites ~468 Ma) during the 68 period in which the Laurentian margin was entering the inferred Taconic subduction zone. In the collision between the Banda Arc and the northern margin of the Australian continent, backarc spreading ceased at ~3.5 Ma with the onset of subduction of Australian continental crust (Hinschberger et al., 2005). In fact, backarc spreading should cease at the onset of continent-arc collision as buoyant continental crust begins to uplift the arc and subduction complexes, transmitting compressional stresses into the backarc region. The fact that abundant backarc volcanism in the southern Appalachians occurs simultaneously with synorogenic clastic wedge deposition in the foreland is not compatible with observations from modern arc-continent collisions. Additionally, geologic constraints based on the arrangement of Blue Ridge lithotectonic belts in northern Georgia, North Carolina, South Carolina, and Tennessee present other difficulties when used to support a model of Middle Ordovician arccontinent collision between Laurentia and an exotic or peri-Laurentian arc on the overriding plate (Holm et al., 2006; Barineau, 2008; Holm-Denoma et al., 2008; Barineau et al., 2008). First, and foremost, with the exception of the the 468 Ma (Meschter-McDowell et al., 2002) to 456 Ma (Miller et al., 2006) Persimmon Creek Gneiss and 466 Whiteside pluton (Miller et al., 2000), the apparently missing rocks of a Middle Ordovician volcanic arc proper have to have been tectonically exhumed and eroded, subducted during collisional orogenesis, or buried by advancing thrust sheets for most of the 500 km strike length of the Blue Ridge northeast of the Cartersville transverse zone in a traditional Taconic arc-continent collision scenario. While modeling of arc-continent collisions suggests that thin oceanic arcs might be largely subducted during continental collision, thick arcs and those built on continental lithosphere would be largely accreted at shallow levels to the subducting continent (Boutelier et al., 2003). Since numerous geochemical analyses of Ordovician suprasubduction volcanic-plutonic rocks in the southern Appalachians indicate the involvement of Mesoproterozoic basement in their genesis (Thomas, 2001; MeschterMcDowell et al., 2002; Bream, 2003; Holm-Denoma, 2006; Das, 2006; Miller, 2006; Tull et al., 2007 ), any model of arc-continent collision requires that these arc rocks be built atop Grenville-aged (Laurentian or peri-Laurentian?) crust. An arc of this nature would be nearly impossible to completely subduct, and thus rocks of a Taconic arc sensu 69 stricto should have been obducted and preserved on the Laurentian margin following Middle Ordovician orogenesis in an arc-continent collisional model. Thus, models of arc-continent collision for the Taconic orogeny in the southern Appalachians must call for exhumation and erosion of the arc proper along most of its length, or for nearly complete burial beneath thrust sheets associated with later orogenesis (Alleghanian orogeny). Finally, rocks of the central and eastern Blue Ridge often interpreted as having formed in an accretionary prism complex on the subducting Laurentian plate, are themselves intruded by or intercalated with Early to Middle Ordovician suprasubduction plutonic-volcanic suites (Drummond, 1986; Hatcher et al., 2007). Since suprasubduction arc magmas should have intruded the overriding plate and not subduction complexes on the subducting Laurentian plate, interpretation of these rocks as both accretionary prism complexes and host rocks to suprasubduction magmas creates a problematic tectonic model for a Taconic arc-continent collision. 4.4 Accretionary Orogens Southern Appalachian tectonic models which call for subduction of the Laurentian margin beneath an exotic or peri-Laurentian arc during the Middle Ordovician (collisional orogenesis) result in timing, tectonic, and kinematic complexities that are difficult to reconcile in the context of an arc-continent collision. An alternate model of accretionary orogenesis, however, fits well with the arrangement of lithotectonic belts and timing of orogenic events in the Middle Ordovician of the southern Appalachians. Accretionary orogens result from extensional and contractional phases in backarc basins, which alternatively result in the creation of sedimentary basins floored by backarc volcanics and the formation of compressional features including thrust belts and synorogenic clastic wedges (Coney, 1992; Collins, 2002a). Convergent plate boundaries are dominated by accretionary orogens, as collisional orogens due to the subduction of thick, buoyant crustal bodies (island arcs, oceanic plateaus, microcontinents, etc.) are necessarily short lived (Cloos, 1993; Stern, 2004). Following 70 a non-terminal collisional orogen, subduction is generally transferred into nearby ocean basins, resulting in the subduction of oceanic lithosphere. Continental bodies around the Pacific, for example, have faced an open ocean basin since the early Paleozoic and circum-Pacific orogenic belts are all attributed to accretionary orogenesis (Collins, 2002a). The Cambrian to Carboniferous Lachlan orogen of southeastern Australia provides an excellent example of an extensional accretionary orogen along a continental margin, as its geologic history was not obscured by subsequent terrane accretion or a continent-continent terminal collision (Foster and Gray, 2000; Collins, 2002a). The almost 1000 km wide Lachlan orogenic belt consists of extensive turbidite deposits floored by MORB to backarc basin volcanic-plutonic rocks, with marine chert and other sedimentary cover, that developed during periods of roll-back in the subducting oceanic lithosphere and extension on the overriding plate (Foster and Gray, 2000). Turbidites are dominantly derived from the Australian craton and were intruded by I and S-type granitoid bodies during extensional phases (Collins, 2002a; Collins and Richards, 2008). Many granites in the Lachlan formed from melting at granulite facies conditions associated with crustal underplating during extensional phases (Collins, 2002a; Collins and Richards, 2008), similar to the environment of formation for granulites of the Hidaka metamorphic belt in Japan (Kemp et al., 2007). Tectonic “switching” (Collins, 2002b) events in the Lachlan resulted in oceanward migration of the arc and backarc system, as shallow subduction, contraction of the backarc, and crustal thickening was followed by rollback of the subduction hinge, extension on the overriding plate outboard of the thickened crustal region, and migration of the arc and backarc system away from the Australian margin (Foster and Gray, 2000; Collins, 2002b; Collins and Richards, 2008). Contractional events typically lasted less than 10 m.y., and at least one may have involved convergence with a microcontinental block in the backarc region (Cayley and Taylor, 1999; Foster and Gray, 2000). Blueschistbearing mélange zones are associated with major fault systems within the Lachlan backarc and probably represent accretionary prism complexes associated with shortlived subduction of backarc crust during phases of orogenic contraction (Foster and Gray, 2000; Spaggiari et al., 2002; 2004; Fergusson, 2003). 71 Igneous volcanic and plutonic rocks in the Lachlan consist of Cambrian oceanic arc-backarc affinity low to medium-K basalts, boninites, and low-K tholeiites near the Australian craton (Collins, 2002a). Ordovician arc rocks developed farther outboard of the craton and include low to high-K calc-alkaline to shoshonitic basalt, basaltic andesite and andesite similar to that found in modern day Fiji (Collins, 2002a; Glen et al., 2007). These arc rocks were tectonically dismembered and isolated from one another during extensional phases of the Lachlan (Glen et al., 1998). Early Silurian tholeiitic, calcalkaline, and transitional mafic rocks are found in the central and eastern Lachlan, outboard of the Cambrian backarc region (western Lachlan) and are coeval with voluminous S-type granitic magmatism. Late Silurian to Middle Devonian granitoids occur mainly in the eastern Lachlan and show a transition from S to I-type magmatism with the onset of a new extensional phase in the Lachlan backarc (Collins, 2002a; Collins and Richards, 2008). Collins and Richards (2008) argue that S-type granitoids in the Lachlan and other accretionary orogens form from underplating of thickened crust by mafic magma, and thus indicate extensional phases following contraction during tectonic switching. As extension progresses and sedimentary material on thinned crust contributes less and less to magma genesis, S-type granitoids gradually give way to I-type granitoids, a process they note has also occurred in backarc basins in Japan and New Zealand (Collins, 2002a; Collins and Richardson, 2008). Thus, the Lachlan orogenic belt of southeastern Australia consists of a western metamorphic belt, adjacent to the Cambrian Australian craton, in which the earliest backarc volcanic rocks are generally exposed in the hanging walls of major reverse faults which soled at the base of turbidite or mafic sequences. Progressively younger portions of the backarc formed more outboard (east) of the Australian craton and include a diverse group of arc rocks and younger S-type plutons. The youngest arc rocks occur in the easternmost part of the orogen and formed during a period of prolonged lithospheric extension. Periods of contraction resulted in major reverse faulting, deformation, and synorogenic sedimention, whereas periods of extension resulted in mafic and felsic volcanism and the formation of younger rift basins, floored by oceanic lithosphere, more outboard of older portions of the orogen. Contractional 72 phases mark the arrival of buoyant lithosphere (e.g. oceanic plateaus) and shallowing of the subducting plate, while extensional phases indicated rollback of the subduction hinge and oceanward migration of the active arc. Extensional faults were commonly reactivated during contractional phases (inversion) and exhumed deep seated metamorphic rocks, while partial subduction of lithosphere within the backarc resulted in subduction complexes complete with LT-HP bearing mélange. Portions of the central Lachlan were affected by wide mylonitic strike-slip fault zones with periods of both dextral and sinistral movement. Although the Lachlan Orogen can be loosely confined to two major tectonic episodes, time transgressive and localized orogenic events occur between these major events, complicating the magmatic and metamorphic history of the orogen (O’Halloran and Cas, 1995; Gray and Foster, 1997; 1998; Foster and Gray, 2000; Collins, 2002a, Fergusson, 2003; Spaggiari et al., 2004). The timing of tectonic events and arrangement of lithotectonic terranes in the Lachlan share many features with the lower Paleozoic history and lithotectonic elements of the southern Appalachians. 4.5 Lithotectonic Elements of the Southern Appalachians In the southernmost Appalachians of Alabama and Georgia, rocks of a 470 to 460 Ma backarc influenced by Mesoproterozoic (Grenville) continental material have been tectonically emplaced on rocks of the Laurentian shelf at the eastern-western Blue Ridge boundary (Holm-Denoma, 2006; McClellan et al., 2007; Tull et al., 2007). As discussed previously, the tectonic, kinematic, and timing constraints for these Middle Ordovician backarc rocks do not support a model of arc-continent collisional orogenesis for their formation and emplacement. In fact, the inferred proximity of these backarcs to the early Paleozoic Laurentian margin, coupled with the tectonic history of the AlabamaGeorgia Talladega and Ashland-Wedowee belts, suggests formation on the attenuated margin of the Laurentian Alabama promontory (Tull et al., 2007). Consequently, the Cambrian(?)-Ordovician history of the Alabama promontory must involve initiation of a cratonward (westward) dipping subduction zone and subsequent arc-backarc evolution on the Laurentian margin. The northeastern extent of this subduction zone must lie 73 somewhere between the Appalachians of Alabama-Georgia (Alabama promontory) and type area of the Taconic collisional orogen in New England (New York promontory). At this boundary a flip in subduction polarity similar to that observed in modern Taiwan (Huang et al., 2006) allows a transition from lower Paleozoic accretionary orogenesis to the south, to collisional orogenesis in the north. Delineation of this boundary should be evident from the arrangement of lithotectonic belts coupled with the timing of tectonism in the southern Appalachians. Dahlonega Gold Belt Along strike of the Ashland-Wedowee belt and in a similar structural position (in the hanging wall of the Eastern-Western Blue Ridge boundary) lies the Dahlonega Gold belt (Fig. 4.1). Bimodal volcanics of the Dahlonega Gold belt have been interpreted as having formed in a backarc setting by a number of workers (McConnell and Abrams, 1984; Burnell and Cook, 1984; German, 1989; Spell and Norrell, 1990; Settles, 2002). The 466 ± 6 Ma Barlow Gneiss, a possible Pumpkinvine Creek formation (Galt’s Ferry Gneiss) equivalent (Thomas, 2001), and 482 ± 7/482 ± 4 Ma Cane Creek Gneiss of the Sally Free mafic complex (Settles, 2002; Bream, 2003) are interlayered with amphibolites and associated metasediments of the Otto Formation, as well as maficultramafic complexes farther to the northeast in Georgia and North Carolina (Settles, 2002). Chromite cores from rocks of the 468 ± 6 Ma Lake Burton mafic complex (Thomas, 2001) dominantly plot in the arc and suprasubduction fields on a TiO 2 /Al 2 O 3 discrimination diagram of Kamenetsky et al.,(2001) (Swanson et al., 2005) and are interpreted as having formed in an island-arc setting based on Pearce and Cann (1973) Ti-Zr-Y plots (Settles, 2002). Hatcher et al. (2007) suggest that metasedimentary rocks of the Dahlonega may be distal Laurentian facies intercalated with arc and MORB-like metaigneous rocks. Cowrock Terrane The Cowrock terrane (Fig. 4.1), structurally above the Hayesville fault and below (in part) the Dahlonega Gold belt, consists dominantly of metasediments, but also contains mafic-ultramafic rocks within the Coleman Formation and Dicks Creek-Lake Chatuge 74 complex (Swanson et al., 2005; Hatcher et al., 2005; Hatcher et al., 2007). Amphibolites of the Dicks Creek and Lake Chatuge complex have been interpreted as having mid-ocean ridge or backarc basin affinities, based on REE patterns and lithologic character (Shaw and Wasserburg, 1984; Kolsrud et al., 2001; Swanson et al., 2005). Coleman River metasandstones of the Coweeta Group, however, contain Grenville zircons and are interpreted as distal Laurentian sediments (Bream et al., 2004). Additionally, the Coweeta Group is intruded by the 468 ± 3 Ma (Meschter-McDowell et al., 2002) or 455.7 ± 2.1 Ma Persimmon Creek Gneiss (Miller et al., 2006), which consists of mafic, granitic, and TTG rocks, and is interpreted as having formed in a suprasubduction environment with a significant Grenville crustal component (Meschter et al., 2001; Meschter-McDowell et al., 2002). Cartoogechaye Terrane The Cartoogechaye terraned (Fig. 4.1), northeast of the Cowrock terrane and structurally above the Hayesville-Soque River fault, consists of Grenville basement, metasandstone, pelitic schists, mafic-ultramafic complexes, and uncommon quartzite hypothesized to have a chert or felsic tuff protolith (Hatcher et al., 2007). Maficultramafic rocks of the Cartoogechaye terrane include the Buck Creek, Webster-Addie, Balsam Gap, Dark Ridge, and associated unnamed amphibolite bodies, which have been metamorphosed as high as granulite facies (Swanson et al., 2005). Major, trace, and rare-earth elemental analyses suggest MORB to suprasubduction affinities for these rocks (Peterson et al., 2004; Swanson et al., 2005; Ryan et al., 2007), and amphibolites have both calc-alkaline and tholeiitic metabasaltic protoliths (Ryan et al., 2005). Gneissic rocks associated with the Balsam Gap metadunite may have a metagraywacke protolith (Swanson et al., 2005). Like the Cowrock terrane, the Cartoogechaye is interpreted to have a distal Laurentian affinity, possibly forming on oceanic crust with associated sediment derived from the Laurentian margin (Bream et al., 2004; Hatcher et al., 2007). 75 Western Tugaloo Terrane (eastern Blue Ridge) Structurally above the Chattahoochee–Holland Mountain fault and Central Blue Ridge (Dahlonega, Cowrock, and Cartoogechaye terranes), lies the western Tugaloo terrane, which is bounded on the southeast by the Brevard Zone (Fig. 4.1). The western Tugaloo is dominated by the Ashe-Tallulah Falls formation, which contains evidence for relict LT-HP metamorphism southwest and northeast of the Grandfather Mountain Window (Page et al., 2003; Page et al., 2005; Miller et al., 2006), but also has a number of mafic-ultramafic bodies (Greer Hollow, Hoots, Day Book, Laurel Creek, unnamed amphibolites) with variable MORB to arc-backarc affinities (Swanson et al., 2005). The western Tugaloo has also been intruded by a number of plutonic granitoid bodies, including the Whiteside (466 Ma), unnamed dikes (420-410 Ma), Looking Glass (380 or 337 Ma), Pink Beds (393 or 370 Ma), and Rabun pluton (335.1 ± 2.8 Ma), Stone Mountain (335.6 ± 1 Ma), and Mt. Airy granite (334 ± 3) all of which have varying trondhjemitic-granodioritic compositions (Miller et al., 2001; Stahr et al., 2005; Miller et al., 2006; Varnell et al., 2008). The Kennesaw granitoid of the Laura Lake maficultramafic complex is a 444±11 Ma, but may be exposed through a window into the Dahlonega Gold belt (Bream, 2003). Isolated bodies of Grenville basement, 1.2-1.3 Ga detrital zircons from metasedimentary rocks, and Grenville-aged xenocrystic components in granitoids suggest proximity of the western Tugaloo terrane outboard of the Laurentian margin during its formation (Miller et al., 2001; Hatcher et al., 2007). Eastern Tugaloo Terrane (western Inner Piedmont) The eastern Tugaloo terrane, southeast of the Brevard Zone, has been interpreted as distal slope-rise deposits formed in a deep water basin floored by oceanic crust which sourced rock of the Laurentian margin (Bream et al., 2004; Hatcher et al., 2007). Metaigneous rocks include the Henderson Gneiss (470 Ma), Dysartville Tonalite (468 ± 8 Ma), Caesars Head (460 Ma), Lenoir Quarry migmatite (450 ± 6 Ma), and Toccoa (449 ± 4 Ma) granitoids, which generally plot (Pearce et al., 1984) as volcanic arc rocks or syncollisional granites and were influenced by Grenville crustal material (Ranson et al., 1999; Bream, 2003; Stahr et al., 2005). A metatuff(?) quartzite (459 or 445 Ma) of the Poor Mountain formation, which is dominated by mafic-ultramafic rocks 76 with MORB to volcanic arc affinity and includes intercalated metasedimentary sequences, plots in the ocean ridge or within plate granite field (Bream, 2003). Cat Square Terrane (eastern Inner Piedmont) The mixed provenance Cat Square terrane is separated from the eastern Tugaloo terrane across the Brindle Creek fault, and is interpreted as a deep water basin formed atop oceanic crust receiving detritus from both Laurentia and peri-Gondwanan rocks (Bream et al., 2004; Hatcher et al., 2007). Middle Silurian (430 Ma) detrital zircons in the Cat Square provide a maximum age of sediment deposition, which was at least 4 km in thickness (Merschat and Hatcher, 2007). Mafic and ultramafic rocks in the Cat Square are uncommon, but abundant Devonian and younger plutons including the Anderson Mill (415 Ma), northeastern Walker Top (407), Toluca (380 Ma), Pelham (364 Ma), southwestern Walker Top (355), Gray Court (357 Ma), and Reedy River (325 Ma) granitoids are interpreted mostly as anatectic melts (Mapes, 2002; Merschat and Hatcher, 2007; Gatewood, 2007; Hatcher et al., 2007). 4.6 The Southern Appalachian Taconic Orogeny Like the Lachlan and other accretionary orogenic belts, the tectonic history of the southern Appalachians can be explained through a model of tectonic switching between contractional and extensional modes of deformation, possibly associated with the arrival of oceanic or microcontinental terranes on subducting Iapetan oceanic lithosphere. Extensional phases resulted in attenuation of lithosphere, magmatic underplating and high crustal heat flow, plutonic-volcanic activity, the formation of successor basins atop older tectonic elements, and creation of backarc lithosphere with MORB to suprasubduction affinities. Short-lived compressional phases (~10 m.y.) result in shortening and thickening of previously attenuated lithosphere, basin inversion along older extensional faults, synorogenic sedimentation, and dynamothermal metamorphism. Subsequent extensional-contractional phases result in general outboard migration of the active arc, which may become dissected by faulting over time and difficult to identify as a once continuous feature, while anatectic S-type granitoids 77 may alternate in time with suprasubduction I-type granitoids. Tectonic events may be locally isolated, resulting in complex temporal and spatial metamorphic patterns over the life of the orogen. With this in mind, a generalized tectonic history of the southern Appalachians is presented below in the context of accretionary orogenesis (Plate 5). Latest Cambrian(?) to earliest Ordovician (>500 to ~485) Failure of oceanic lithosphere outboard of the Alabama promontory results in initiation of a subduction zone at the continent-ocean boundary and cratonward (Andean-type) subduction (Barineau and Tull, 2006; 2007). Subduction of Iapetan lithosphere outboard of the Alabama promontory continental hinge zone results in latest Cambrian plutonism preserved as xenocrystic components in younger igneous bodies (McClellan et al., 2007) and as detrital zircons in younger sedimentary basins (Miller et al., 2000). The Elkahatchee Granodiorite may be the only preserved plutonic body of this age in the southern Appalachians, but will have to await more thorough geochronologic work to resolve conflicting geochronologic work (Mueller, personal communication). Miller et al. (2000) note that zircons of this age in the Persimmon Creek Gneiss may represent the actual crystallization age of the pluton, but could not resolve the conflict between concordant ~490 U-Pb ages and probability plots showing a dominant 475-470 Ma zircon population. Mild tectonism in both the Talladega belt and Alabama foreland may be a result of initiation of subduction during this time interval. Earliest Ordovician to middle Late Ordovician (<485 to ~455 Ma) Roll back of the oldest, densest part of subducting Iapetan lithosphere (closest to the Laurentian margin) causes extension of the overriding lithosphere, in this case the attenuated continental margin outboard of the Alabama promontory and within the Tennessee embayment. An extensive backarc basin develops outboard of the Laurentian margin, just as the oldest parts of the Lachlan developed outboard of the Australian margin (older Delamerian orogenic belt) (Foster and Gray, 2000). A zone of widespread backarc igneous activity stepped from the transitional-oceanic lower plate Tennessee embayment to continental-transitional upper plate Alabama promontory. 78 The oldest parts of the backarc (Cane Creek, ~482) show strongly negative ε Nd values (Bream, 2003) due to the influence of continental basement and continentally derived sediment in the early phases of lithospheric attenuation. Laurentian margin derived sediment (turbidites?) filled in the widening and deepening rift basins, atop the newly created backarc basement and intercalated older Laurentian basement outboard of the continental margin. Underplating of pre-existing (Iapetan rifting) and newly deposited (backarc rifting) sedimentary rocks by mafic magmatism in the attenuating lithosphere resulted in static granulite facies and lower metamorphism (Moecher et al., 2005; Clemons, 2006). This high grade metamorphic event was not ubiquitous to the entire backarc region, resulting in complex relationships between interpretations of peak metamorphic age. High grade metamorphism from this time period was restricted to rocks of the Tennessee embayment, as thicker continental lithosphere and proximal rifted prism rocks in the Alabama promontory segment of the backarc would have limited both extension of the overriding lithosphere and heat flow from magmatic underplating. This would also explain the greater abundance of mafic-ultramafic complexes in the Dahlonega Gold belt and central Blue Ridge, as thin transitional-oceanic lithosphere in the Tennessee embayment should have preferentially thinned compared to the thicker crust of the Alabama promontory and attenuated margin. Intrusion of Early-Middle Ordovician silicic suprasubduction plutons into the older basement and old-new rift sequences is evident from the abundant granitoids in the central-eastern Blue Ridge and western Inner Piedmont. The continental arc affinity Deicke and Millbrig K-bentonites of the Appalachian foreland, which thicken toward a source on the Alabama promontory (Kolata et al., 1998) could be part of this tectonic phase, depending on the validity of older 457 to 453 Ma ages (Tucker, 1992; Tucker and McKerrow, 1995; Samson et al., 1989) versus more recent 450 to 448 ages (Min et al., 2001). If the younger dates are correct, this would place the eruption of these voluminous ash deposits in the next phase of tectonism. In either case, the presence of continentally influenced arc volcanics sourcing from the Alabama promontory is supportive of accretionary orogenesis for the southern Appalachians. 79 LT-HP rocks of the eastern Blue Ridge Tugaloo terrane (Hatcher et al., 2007) may have formed during a short contractional phase near the end of this time interval, resulting in partial backarc closure north of the Alabama promontory and brief subduction of backarc oceanic crust under the Laurentian margin. Alternatively, the extending backarc may have exhumed LT-HP rocks associated with early phases of subduction of Iapetan crust, with retrograde decompression and exsolution of zirconium resulting in younger 459 Ma zircon growth (Miller et al., 2006). Similar arguments could be used to explain the Middle Ordovician Taconic (Blount) clastic wedge, which might have formed from basement-cover extensional uplifts on the cratonward edge of the backarc, but which might also have formed from a short contractional phase associated with the arrival of buoyant material on the subducting Iapetan lithosphere and the development of a restricted thrust belt in the vicinity of the Tennessee embayment. Evidence to distinguish between the two is not conclusive; however, compressional deformation features from this time interval are supportive of a short phase of contraction and crustal thickening between 465 and 455, a 10 m.y. time frame similar to the duration for contractional phases in the Lachlan orogenic belt. Additionally, intrusion of S-type granites in the next phase of accretionary orogenesis (see below) is also supportive of crustal thickening at the end of this tectonic phase. Late Ordovician and Younger Events (<455 Ma) Following widespread extension in the Early to Late Ordovician and a brief contractional phase of crustal shortening in the late Middle to early Late Ordovician, steepening of the subducting slab resulted in a new phase of extension, which stepped outboard of the older backarc region. Increased heat flow and underplating of older rocks resulted in initial anatectic, S-type granitoids (Mulberry Rock, Kennesaw, Lenoir Quarry, Toccoa) derived in large part from the recently thickened sedimentary package in the Tennessee embayment and the initial, thicker crust and subsequent backarc sediment outboard of the Alabama promontory. Continued extension increased the contribution of mantle wedge to magma genesis and the resumption of more I-type igneous activity (Collins and Richards, 2008). This phase of extension moved the locus of backarc activity farther outboard of the earlier backarc phase and may be 80 represented by the ~445 Ma Poor Mountain Formation, which contain metafelsic volcanics and metamafic rocks of MORB or suprasubduction affinity (Bream, 2003). Timing constraints for subsequent tectonic switching phases younger than Ordovician are less restrictive and the duration of Iapetan subduction beneath the Laurentian plate is unknown, however, at least one other tectonic phase may be evident. Detrital zircons in the Cat Square terrane suggest a maximum age of 430 Ma for sedimentation within this sedimentary basin (Bream et al., 2004; Merschat and Hatcher, 2007). The ~415 Anderson Mill (Mapes, 2002), ~406 Walker Top Granite (Gatewood, 2007), and 400 Ma metamorphic rims on zircons (Bream, 2003; Merschat and Hatcher, 2007) constrain a thermal event during this time period. The collective geochronological record for the Cat Square and position outboard of the eastern Tugaloo terrane suggests that continued backarc rifting migrated farther outboard from the time interval <430 to 400 Ma. Whether or not this event is separated from older 455-445 extension in the western Tugaloo by a contractional phase is unknown. Extension in this younger (<430) time interval is further evidenced by destabilization of the most outboard Laurentian shelf sequences (Talladega and Murphy belt), which contain latest Silurianearliest Devonian successor basin rocks (Tull and Groszos, 1990; Tull, 2002), suggestive of rifting. Formation of the Cat Square during this time period is consistent with rapid Silurian-Devonian extension outboard of the older backarc regions, which may have reactivated older normal faults near the continental hinge zone. If, however, the Cat Square terrane formed farther to the northeast, between the Virginia and New York promontory as suggested by Merschat and Hatcher (2007) and was transported into its current position by significant transpressional or strike-slip tectonics, then these two events may be unrelated. Like the long-lived backarc mobile belt of the North American Cordillera, arrival of exotic terranes on the subducting plate (e.g. Carolina terrane) would have resulted in deformation and metamorphism which may not have affected the orogenic belt as a whole, depending on the size of the docking terrane. However, the stresses generate by these events would have been transmitted into the hot and weaker portions of the backarc (Hyndman et al., 2005), variably affecting regions more inboard of the subducting lithosphere. Arrival of terranes, tectonic switching events, and high heat 81 flow into the southern Appalachian backarc region would also explain the diverse array of metamorphic events suggested by U-Pb dating of zircon rims (Bream et al., 2004; Merschat and Hatcher, 2007). 82 CHAPTER 5 SUBDUCTION INITIATON ALONG THE EARLY PALEOZOIC LAURENTIAN MARGIN 5.1 Models of Subduction Initiation The dynamic nature of global tectonics indicates that the nature of plate boundaries can change through time. In recent decades, geophysicists have made significant progress in understanding how subduction zones nucleate in intraoceanic settings. The nucleation of a subduction zone along a passive continental margin, however, has been one of the more controversial topics in research of the mechanisms for subduction zone initiation. Failure of lithosphere along continental margins has traditionally focused on the presence of earlier rift-related faults and lithospheric flexure due to hinge zone sediment loading. Arguments have been presented both for (Erickson and Arkani-Hamed, 1993; Erickson, 1993; Zheng and Arkani-Hamed, 2002) and against (Cloetingh et al, 1984; Cloetingh et al, 1989) exploitation of these features as potential sites of lithospheric failure. Modeling of failure along a hypothetical Atlantictype margin suggests that under a 10 km sedimentary load, ‘wet’ lithosphere could fail through a thermal-mechanical feedback, resulting in a narrow, through-going lithospheric shear zone (Regenauer-Lieb et al, 2001). However, the absence of Cenozoic initiation of subduction along the well-developed passive Atlantic margins has led to a general skepticism regarding spontaneous nucleation of subduction in a passive margin setting (Stern, 2004). The geologic history of North America’s southern Appalachians, however, suggests that the trailing, Cambrian passive margin along the Laurentian side of Iapetus was transformed from a passive margin to an Andean-type active margin, at least along the Alabama promontory and Tennessee embayment (Tull et al., 2007; Barineau, 2008; Holm-Denoma et al., 2008, Barineau et al., 2008). Because there is currently no evidence to suggest that an older subduction system was transferred along the Laurentian margin or from an adjacent intraoceanic subduction zone onto the passive margin of the Alabama promontory, then a plausible mechanism for nucleation of a subduction zone is needed to explain both the likely Laurentian margin Hillabee backarc 83 volcanic sequence and intrusion of the Laurentian rifted-to-passive margin by suprasubduction plutons in the southern Appalachians. Inference of transition from a passive margin to active subduction tectonics has commonly been attributed to spontaneous collapse of oceanic lithosphere at the edge of a continent, where sedimentary loading and lithosphere-scale, rift-related fracture zones might encourage failure of oceanic lithosphere and nucleation of a subduction zone (Erickson and Arkani-Hamed, 1993; Erickson, 1993; Zheng and Arkani-Hamed, 2002). This scenario, however, is not generally supported by geophysical models and the fact that mature continental margins of the Atlantic basin do not exhibit characteristics that would suggest they have or will spontaneously fail (Stern, 2004). Instead, the formation of most modern subduction zones fall into one of two categories. In the first type of subduction initiation, referred to by Stern (2004) as “induced nucleation”, the attempted subduction of buoyant crust results in either a jump in the locus of subduction (transference), or a change in subduction polarity. Induced subduction nucleation has resulted in incipient subduction along the Yap Trench at the eastern edge of the western Pacific Carolina plate (Lee, 2004), Neogene subduction polarity reversal at the Solomon Islands arc along the San Cristobal trench (Cowley et al, 2004), and a mid-Cretaceous polarity reversal along the Greater Antillean arc (Pindell, 1994). A second category of subduction initiation, “spontaneous nucleation” (Stern, 2004), does not require the existence of a pre-existing subduction zone, but instead results from lithospheric convergence across pre-existing fracture zones, especially where juxtaposition of oceanic crust of differing densities occurs (Casey and Dewey, 1984; Stern and Bloomer, 1992). Forced convergence across a fracture zone with mechanically contrasting oceanic lithosphere has been proposed as the mechanism by which the Izu-Bonin-Mariana subduction system originated in the Middle Eocene (Hall et al., 2003), whereas evolution of a transform to a transpressional, and finally oblique subduction boundary, is proposed as the origin of the Puysegur trench south of New Zealand (Sutherland et al., 2006). Thus, a latest Cambrian(?) – Ordovician arc along the trailing margin of the Laurentian Alabama promontory-Tennessee embayment could have been the result of 84 subduction nucleation caused by a polarity switch in a pre-existing subduction zone, transference of a pre-existing subduction zone, or spontaneous nucleation of a subduction zone along a pre-existing fracture or transform boundary. All three of these scenarios are explored below. In the northern Appalachians, the Ordovician Taconic orogeny is attributed to subduction of various elements of the Laurentian margin beneath an exotic arc or arcs built on peri-Laurentian crust (Karabinos et al., 1998; Waldron and van Staal, 2001; van Staal et al., 2007). Work by van Staal et al. (2007) suggests that following Neoproterozoic rifting of the Laurentian margin in Newfoundland and separation of the Dashwoods block from Laurentia by a narrow seaway (Humber Seaway), subduction nucleated in an abandoned oceanic spreading center seaward of the Dashwoods and resulted in the formation of a seaward dipping subduction zone (509-495 Ma), eventually leading to subduction of the outboard Dashwoods block beneath the Lushes Bight oceanic tract, a 490-481 Ma ‘infant’ forearc terrane. Following introduction of the Dashwoods into the subudction zone, transference of Iapetan-directed subduction jumped into the proposed Humber Seaway separating the Dashwoods from the Laurentian margin proper, and resulted in obduction of the Notre Dame arc (built on Dashwoods) onto the Laurentian margin (480-470 Ma) and the traditional Taconic orogeny (van Staal et al., 2007). Seaward subduction of the Laurentian margin beneath an exotic arc is also invoked in explaining the Taconic orogeny in New England (Karabinos et al., 1998), a model also often applied to the southern Appalachians (Hatcher, 1987; Hatcher et al., 2007). In the case of both the New England and Newfoundland Appalachians, the assembly of terranes is consistent with obduction of an arc onto the subducting Laurentian margin. In each case, rocks of the obducted arc terrane (Shelburne Falls in New England and Notre Dame in Newfoundland) are separated from rocks of the Laurentian margin by an extensive mélange zone which formed in an accretionary prism on the subducting Laurentian plate (Karabinos et al., 1998; van Staal et al., 2007). In the southernmost Appalachians, however, the most distal rocks of the Laurentian shelf are in the footwall of a pre-metamorphic thrust which emplaced Middle Ordovician backarc rocks of the Hillabee Greenstone onto uppermost Devonian- 85 lowermost Mississippian(?) platform rocks of the upper Talladega Group (Tull et al., 2007). The collectively metamorphosed Laurentian shelf metasedimentary rocks and Hillabee Greenstone (Talladega belt) are, in turn, in the footwall of the postmetamorphic Hollins Line fault which emplaced the slope-rise sediments and plutonic bodies of the Ashland-Wedowee belt atop the Talladega belt (Fig. 2.1). Notably absent, and key to interpretations of arc obduction, are rocks of an accretionary prism. In fact, the Middle Ordovician Hillabee Greenstone and Ordovician plutons of the Ashland Wedowee belt most likely formed in an arc-backarc system built on attenuated continental crust outboard of the Laurentian margin hinge zone (Barineau et al., 2006; Tull et al., 2007). Thus, evidence for obduction of an exotic volcanic arc or arc built on peri-Laurentian crust is lacking in the southernmost Appalachians. This suggests that the seaward dipping subduction zone proposed for the northern Appalachians did not extend the length of the Laurentian margin. Because the Alabama promontory subduction zone was not preceded by seaward dipping subduction, this makes initiation of the Alabama promontory subduction system as a result of a flip in subduction polarity an unlikely scenario. If subduction initiation was triggered by transference of a pre-existing subduction zone, then evidence for that subduction system was either not preserved or has been obscured by Paleozoic orogenic events affecting the Laurentian margin. A model which incorporates subduction initiation via transference requires that a Middle or late Cambrian subduction system existed in proximity to the southeastern Laurentian margin in order for tectonic stresses sufficient for subduction initiation to be transferred to the latest Cambrian-Early Ordovician Alabama promontory. Paleotectonic reconstructions and tectonic interpretations of Appalachian lithotectonic belts do not suggest the presence of any such subduction system proximal to the southeastern Laurentian margin, thus making a passive-to-active margin conversion via subduction transference unlikely. Because initiation of subduction along the Alabama promontory cannot be readily explained by a subduction transference model, without resorting to models involving Middle to late Cambrian subduction systems that have yet to be identified in the southern Appalachians, then it is likely that initiation of the Alabama promontory 86 subduction zone formed as a result of spontaneous subduction nucleation. Our model for this initiation event overcomes the difficulties associated with conversion of a passive continental margin due to spontaneous foundering at the continent-ocean boundary by exploring the interaction between a continental transform boundary of considerable length and an evolving ocean basin. 5.2 Development of the Alabama Promontory Subduction Zone The subduction model presented here is based upon the plate geometry and sequence of events that defined the Late Proterozoic – Early Paleozoic breakup of Rodinia along the trailing margin of Laurentia. The geometry of the early Paleozoic Laurentian continental margin was a consequence of multistage rifting and the presence of offset rift axes and connecting transform faults defining continental promontories and am a-O kla ho ma Tra ns for m Iap Ouachita Embayment OU AC HI TA RI FT Oceanic Crust Iapetus Oceanic Crust Te xas s Ala b 300 etu 200 Crust Rif ted olcan Ma ic rgin 100 Continental Co ntin en Co tal ntin Ma en rgi t-O n ce a nB ou nd ary 0 N Bi r mi Gr ngha ab en m KILOMETERS A PR LABA OM M ON A No TO RY nv embayments (Thomas, 1976, 1991). Following widespread Neoproterozoic Iapetan PRECORDILLERA Tra nsf orm Figure 5.1 - Evolution of the Ouachita embayment along southern Laurentia following Early Cambrian seafloor spreading at the Ouachita rift. Adapted from Thomas and Astini, 1996. 87 rifting along eastern Laurentia (Cawood et al., 2001; Miller and Barr, 2004), a subsequent Early Cambrian rifting event (530-539 Ma) separated the Argentine Precordillera from the southern Laurentian craton along the Ouachita rift (Thomas, 1991; Thomas and Astini, 1996), forming the Ouachita embayment (Fig. 5.1). Translation of the Precordillera relative to Laurentia was accommodated by the northwest-striking Alabama-Oklahoma transform separating Laurentian continental crust from developing oceanic crust across a narrow (~25 km) transitional zone along the southwestern margin of the Alabama promontory (Thomas, 1991; Thomas and Astini, 1996; Thomas and Astini, 1999). This motion facilitated the Cambrian passage of a mid-ocean ridge (Ouachita rift of Thomas, 1991) and flanking young oceanic lithosphere along the southwestern margin of the Alabama promontory and away from Laurentia into the widening ocean basin (Thomas, 1991; Thomas and Astini, 1996). By late Cambrian (Franconian, ~495 Ma), the Ouachita ridge passed beyond the seaward corner of the Alabama promontory and a passive margin was established around the continental margin of the Ouachita embayment (Thomas, 1991; Thomas and Astini, 1996). Once the Ouachita ridge axis approached the seaward corner of the Alabama promontory (Fig. 5.2), young oceanic lithosphere between the trailing margin of the Precordillera and Ouachita ridge would have been juxtaposed first against transitional crust at the continent-ocean boundary, and then older, thermally mature oceanic crust north of the transform that had formed during initial Iapetan seafloor spreading in this portion of the Laurentian margin (Aleinikoff et al., 1995; Cawood et al., 2001; Miller and Barr, 2004). This contrast in age-density has been shown to be one of the key factors in initiating subduction in intraoceanic settings (Hall et al., 2003; Stern, 2004). Assuming that the eastern side of the Alabama promontory represented a Labrador-like, thick sediment, amagmatic rifted margin discussed previously (Louden and Chian, 1999), the continent-ocean boundary and transitional crust east of the Laurentian margin could have been 200 km or more outboard of the shelf edge. If so, just prior to earliest late Cambrian (~501 Ma), as the Ouachita ridge approached the seaward corner of the Alabama promontory, the young (<30 m.y.) oceanic lithosphere along the trailing margin of the Precordillera would have been juxtaposed against the 40 to 60 m.y. Iapetan lithosphere outboard of the Alabama promontory, which formed 88 following synrift Catoctin volcanism. Modeling of subduction initiation involving oceanic lithosphere age offsets of this magnitude (10 to 30 m.y), coupled with convergence rates of 1-2 cm/yr and fault zone stresses less than 20 MPa, indicate that a self- ifte Co ntin ted M 20 -3 0m .y. 20 m. y. 10 - 0- 10 m. y. Ri f 60 m.y . 300 en t- O cea 200 40 - 100 arg in 0 m d M atic arg in KILOMETERS nB ou nd ary N Laurentian Continental Crust Bir mi Gr ngha ab en m ~501 Ma latest Middle Cambrian PR ALAB OM A M ON A Am TO RY R ag sustaining subduction zone could be generated along such a boundary (Hall et al., Figure 5.2 – Middle Cambrian juxtaposition of contrasting lithosphere across the Oklahoma-Alabama transform. Adapted from Thomas and Astini, 1996. 2003). In fact, the lithospheric age-density contrast conditions necessary for subduction initiation along the Alabama-Oklahoma transform span a period of time from Middle Cambrian, following translation of the Precordillera microcontinent beyond the seaward corner of the Alabama promontory and attenuated margin, to Early Ordovician when elastic resistance in the aging Iapetus oceanic lithosphere would significantly increase the amount of convergence necessary for self sustaining subduction and reduce the likelihood of subduction initiation near the continental margin (Hall et al., 2003). Euler pole migrations associated with plate movement during this time interval (Torsvik et al., 89 1996; Kirschvink et al., 1997) would have reorganized plate boundaries, potentially introducing a component of convergence across the Alabama-Oklahoma transform and providing the slight convergence necessary for sustained subduction at sites of former transform boundaries. Convergence across the Alabama-Oklahoma transform in conjunction with juxtaposed oceanic lithosphere of contrasting age and density resulted in initiation of subduction along the Alabama promontory and subsequent arc plutonismvolcanism in the overlying Laurentian marginal basins. In western Pacific subduction systems the evolution of arc volcanism following Cenozoic subduction initiation indicates a lag of 5-15 m.y. between the initiation event and mature arc volcanism/plutonism. For example, Pliocene or younger andesitic arc volcanism associated with the Puysegur trench developed within ~10 to15 m.y. of subduction initiation (Sutherland et al., 2006), andesitic volcanism above the TongaKermadec-Hikurangi subduction zone developed within 5 m.y. of subduction initiation (Stern et al., 2006), and a mature arc developed within ~15 m.y. of subduction initiation in the Izu-Bonin-Marianas subduction system (Stern and Bloomer, 1992). If the trondjhemite-tonalite-granodiorite bodies of the Ashland-Wedowee belt represent the plutonic roots of a mature suprasubduction system, this would suggest that the age of subduction initiation must be at least 5-15 m.y. older than the oldest pluton. As discussed previously, the ages of these plutons are poorly constrained. However, both Rb-Sr whole rock and U-Pb multi-grain zircon analyses (Russell et al., 1987), the presence of inherited latest Cambrian-Early Ordovician zircons in the Hillabee Greenstone metadacites (McClellan et al., 2007), Elkahatchee Quartz Diorite Gneiss, and other plutons of the eastern Blue Ridge and western Inner Piedmont (Miller et al., 2000) as well as latest Cambrian-Early Ordovician zircons in Pleistocene coastal plain sediments of coastal Georgia (Mueller et al., 2008), suggests that plutons of ~500 to 490 Ma may be present in the southernmost Appalachians in general and the AshlandWedowee belt specifically. If ongoing geochemical studies uphold the existence of late Cambrian plutons in the Ashland-Wedowee belt, this would indicate that the Alabama Promontory subduction system may have initiated in the Middle to late Cambrian (513488 Ma) just prior to or during the time in which the Ouachita ridge passed the seaward corner of the Alabama promontory (Thomas and Astini, 1996). However, because of 90 the poor age control on the granitoid bodies of the Ashland-Wedowee belt, timing constraints on subduction initiation based solely on magmatic ages of these granitoids are limited until more detailed geochronologic studies are completed. Other lines of evidence for timing of subduction initiation along the Alabama promontory, however, support interpretation of a Middle to late Cambrian age for this event. εNd and zircon εHf values for the backarc metavolcanic rocks of the Ordovician 470-460 Ma Hillabee Greenstone in the Alabama-Georgia Talladega belt (Russell, 1978; Tull and Stow, 1980; Tull et al., 2007) indicate involvement of ~1.1 Ga continental crust in its genesis. Additionally, a lack of significant deformation during fault emplacement of the Hillabee onto rocks of the Talladega belt (TB) Laurentian shelf argues for limited transport of the Hillabee backarc rocks and suggests an origin on the attenuated Laurentian continental margin associated with backarc rifting (Tull et al., 2007). Both modelled and observed backarc extension in the Izu-Bonin-Marianas subduction system and other Cenozoic arc systems suggest that rollback in newly subducting slabs induces backarc extension within a few tens of millions of years following inception of subduction (Hall et al., 2003, Gurnis et al., 2004; Stern et al., 2006). A 20-40 m.y. gap between Middle to late Cambrian subduction initiation and formation of the Hillabee Greenstone above the attenuated margin matches these observations in Cenozoic systems quite well. Burgess and Moresi (1999) modelled subsidence curves due to dynamic topography on the overriding plate following subduction initiation and found that initial subsidence increased with slab penetration, yielding subsidence curves which began with a concave downward segment. Cambrian tectonic subsidence curves from the Appalachians, Ouachitas, and Precordillera have been used to illustrate synrift and post rift-related subsidence in the southernmost Appalachians (Goodman et al., 1992; Thomas and Astini, 1999). When corrected for the GTS2004 time scale (Gradstein and Ogg, 2004), the concave upward Early and Middle Cambrian rift-related segment of the Birmingham graben tectonic subsidence curve is followed by a concave downward segment in the Late Cambrian to Early Ordovician. This concave downward segment is also reflected in the Cambrian portion of subsidence curves for Kentucky’s Rome trough (Goodman et al., 1992). In the Birmingham graben, probably >600 km from the 91 Alabama promontory subduction zone proposed here, a change from exponentially decreasing subsidence rates typical of rifted margins, to increasing subsidence rates may signal the onset of subduction. While increasing subsidence rates are not unique to dynamic topography above subducting plates, the inflection change within the Birmingham graben and Rome trough tectonic subsidence curves could reflect subduction initiation sometime prior to earliest Late Cambrian (~501 Ma). Modelling by Gurnis (1992) indicates that an overriding continental plate experiences dynamic topography associated with subduction initiation, including large wavelength subsidence close to the trench and a forebulge within a few hundred kilometers of the trench. Gentle uplift of the Alabama promontory shelf strata in the Middle Cambrian could explain the presence of arkosic sandstones within the carbonate-rich Shelvin Rock Church Formation, a distal equivalent of the Conasauga Group in the Talladega belt (Johnson and Tull, 2002). A broad forebulge and shallowing of the Talladega belt carbonate platform could have resulted in local dissection of underlying feldspathic units and recycling of their clastic components. Subduction initiation along the Alabama promontory during this time interval might also be marked by synsedimentary deformation in the Middle to lowermost Upper Cambrian Conasauga Formation of Alabama. A unique Conasauga facies in Alabama includes slumped beds and pseudobreccias suggestive of creep and downslope failure, whereas equivalent units in Tennessee and Virginia suggest gently sloping ramps indicative of low-gradient environments (Astini et al., 2002). This synsedimentary deformation probably resulted from reactivation of boundary faults in the Birmingham graben and has been traditionally attributed to extension associated with passage of the Ouachita rift south of the Alabama-Oklahoma transform (Thomas, 1991; Thomas and Astini, 1996). However, studies of the Romanche Fracture Zone paleotransform on the Atlantic margin of Africa suggests that the influence of rift axes propagating along significant continental margin offsets during seafloor spreading rapidly decreases on the landward side of the transform boundary. Seismic studies within continental margin synrift basins of offshore Ghana indicate that continental extension and transform related deformation occurred within a zone less than 80 km wide paralleling the continental margin transform and perpendicular to the ridge-transform terminus (Attoh et 92 al., 2004). Therefore, using the paleo-Romanche transform as an analogue, it is unlikely that fault activity within the Birmingham graben, 200 km or more northeast of the Alabama-Oklahoma transform, is associated with passage of the Ouachita ridge during the Early and Middle Cambrian. It is more likely that Middle and/or late Cambrian extension (Thomas, 1991; Thomas and Astini, 1996) within the Birmingham graben was associated with dynamic topography within the overlying plate as a function of subduction initiation outboard of the Alabama promontory. Several lines of evidence suggest initiation of the subduction zone proposed here at a location hundreds of kilometers outboard of the continental hinge zone. First, if lithospheric failure had developed within flexural basins and/or rift-related faults immediately adjacent to the Laurentian shelf, subsequent suprasubduction igneous activity would have penetrated shelf rocks of the outermost Alabama promontory (i.e., Talladega belt rocks). Additionally, significant uplift-subsidence associated with dynamic topography in the overriding plate close to the trench would have left a dramatic record in these same rocks had subduction initiated close to the shelf edge (Burgess and Moresi, 1999; Sutherland et al., 2006). Because neither arc rocks nor evidence of dramatic dynamic topography are associated with the Middle Cambrian stratigraphy of the Talladega belt’s Sylacauga Marble Group, subduction initiation along the shelf edge can be precluded. The Alabama promontory shelf stratigraphy (Shelvin Rock Church and Conasauga Formations), instead, records far-field effects of dynamic topography at the shelf edge (Talladega belt) and further inboard on the Alabama promontory (Birmingham graben). 5.3 Evolution of the Alabama Promontory Subduction Zone Conversion of a transform plate boundary to a subduction boundary has been proposed for both the Izu-Bonin-Marianas subduction system in the western Pacific (Hall et al., 2003; Stern, 2004) and the Puyesgur trench south of New Zealand (Lebrun et al., 2003; Sutherland et al., 2006). In each case, migration in Euler poles resulted in convergence across a pre-existing transform or fracture zone. Both the Puysegur and Izu-Bonin-Marianas models of subduction initiation (Fig. 5.3) result in plate convergence 93 approximately perpendicular to the pre-existing transform (i.e., the subduction zone propagates parallel to the transform). In our Paleozoic Laurentian example, however, Figure 5.3 – Evolution of the Puysegur subduction zone. Modified from Lebrun et al., 2003 and King, 2000. an additional component of influence over the developing subduction zone must be included in order to result in plate convergence nearly parallel to the pre-existing transform (i.e., beneath the continental margin). This component is present in the margin-parallel, faulted, weak serpentinized transitional crust at the continent-ocean boundary. Studies of amagmatic rifted margins indicate that limited melt production may be the result of a cooler than normal (<1300°C) sublithospheric mantle (Reston and Morgan, 2004). Crustal thinning under these conditions is accommodated by a system of concave-downward faults and results in the exhumation of serpentinized mantle rock at the ocean-continent transition zone (Whitmarsh et al., 2001; Lavier and Manatschal, 2006). Studies along non-volcanic or amagmatic rifted continental margins suggest a transition zone between attenuated continental crust formed during rift-related extension 94 and oceanic crust formed during the initial stages of true seafloor spreading (Louden and Chian, 1999; Whitmarsh et al., 2001). The exhumed serpentinized mantle of the ocean-continent transitional crust is mechanically significant in terms of potential lithospheric failure, as the presence of serpentinite can reduce lithospheric strength by up to 30% (Louden and Chian, 1999; Escartin et al., 1997). Furthermore, the attenuated margin adjacent to a continent and transitional crust of the continent-ocean boundary can be affected by the faulting which accommodates mantle exhumation (Salisbury and Keen, 1993; Louden and Chian, 1999; Whitmarsh et al., 2001), providing yet an additional component of lithospheric weakness. Should subduction initiate near the continent-ocean boundary along a transform framing a continental promontory, propagation of the subduction zone toward the adjacent continent would be mechanically limited (Fig. 6a), as it would require subduction of thermally buoyant continental lithosphere or attenuated continental lithosphere tied to the downgoing oceanic slab. However the existence of faulted oceanic crust or faulted-serpentinized transitional crust (Whitmarsh et al., 2001) near the locus of subduction initiation might allow trench propagation parallel to the rifted margin. In a scenario similar to the formation of the Puysegur trench, transpression along a transform boundary could result in initial collapse of oceanic lithosphere and subsequent asthenospheric flooding of its upper surface (Lebrun et al., 2003; Sutherland et al., 2006), especially if accentuated by lithosphere of significantly contrasting age and density (Hall et al, 2003), as our model suggests. The direction of trench propagation could vary, depending on the mechanical weakness of the lithosphere involved (Fig. 5.4). Three potential scenarios for the development of this propagating subduction zone are presented here. 1) If the faulted, serpentinized transitional crust at the continent-ocean boundary were mechanically weaker than the dense oceanic lithosphere north of the Alabama-Oklahoma transform (Iapetus), trench propagation would develop parallel to the continental margin (Fig. 5.4b) and perpendicular or highly oblique to the Alabama-Oklahoma transform (depending on the initial geometry of the rifted margin). In this scenario, the Alabama-Oklahoma transform outboard of the Alabama promontory would remain an active transform boundary, developing slab pull would be toward Laurentia, and the resultant arc would develop 95 Figure 5.4 – A) Lithospheric collapse at the continent-ocean boundary and potential paths of trench propagation. B) Trench propagation develops along the continent-ocean boundary with resulting subduction beneath the continental promontory. C) Trench propagation develops along the Oklahoma-Alabama transform with resulting subduction initiation beneath oceanic lithosphere in the embayment. D) Trench propagation develops in two directions resulting in subduction oblique to both the continental promontory and Oklahoma-Alabama transform. 96 along the attenuated continental margin. 2) If the Iapetus oceanic lithosphere were mechanically weaker than the transitional crust at the continent-ocean boundary, then the developing trench would propagate seaward along the Alabama-Oklahoma transform and perpendicular or highly oblique to the continental margin (Fig. 5.4c). In this scenario, developing slab pull would convert the outboard segment of the AlabamaOklahoma transform to a subduction zone, with the arc developing on the oceanic lithosphere trailing the Precordillera microcontinent. A transform boundary would be required along and parallel to the Laurentian margin, possibly in the weak transitional crust at the continent-ocean boundary. 3) If the mechanical weakness of the transitional crust and oceanic lithosphere were similar, then trench propagation might develop seaward along the Alabama-Oklahoma transform and along the Laurentian margin at the continent-ocean boundary. In this scenario, Iapetus oceanic lithosphere would subduct initially oblique to both the Laurentian margin and Alabama-Oklahoma transform (Fig. 5.4d), while the arc would develop on both the attenuated Laurentian margin and oceanic lithosphere trailing the Precordillera. Although three potential scenarios are presented, the first (Fig. 6b) is more likely for several reasons. First, if trench propagation had developed along the Alabama-Oklahoma transform (Fig. 5.4c), then a Middle to late Cambrian arc should have developed on oceanic lithosphere trailing the Precordillera microcontinent. This arc should have been obducted onto the Laurentian margin during Alleghenian orogenesis, when the Carboniferous age Ouachita arc collided initially with the Alabama promontory, eventually being emplaced onto Laurentian slope and rise facies rocks in the Ouachita embayment (Viele and Thomas, 1989; Keller et al., 1989). While Neoproterozoic-Cambrian arcs of Gondwanan or peri-Gondwanan affinity (Carolina terrane) are known from the southern Appalachians (Hibbard et al., 2002), there are currently no recognized arc terranes which would have occupied a paleotectonic position proximal to Laurentia in the Middle or late Cambrian, thus making scenario 2 unlikely. The same can be said of scenario 3, in which the arc steps from attenuated continental lithosphere outboard of the Alabama promontory onto oceanic lithosphere trailing the Precordillera microcontinent (Fig. 5.4d). The resultant Middle to late Cambrian intraoceanic arc segment should be present in either the southern Appalachians or the Ouachitas. Secondly, if trench propagation had 97 developed for a significant distance along the Alabama-Oklahoma transform, Iapetus oceanic lithosphere should have eventually been subducted beneath and generated an arc on the Precordillera microcontinent. Again, this is an unlikely scenario as the Precordillera experienced passive-margin sedimentation over the time interval during which subduction initiation would have occurred (Thomas and Astini, 1999). It is possible that trench propagation did not initially develop parallel to the Laurentian margin, but instead was transferred to the margin by another Euler pole migration (following that which resulted in Middle to late Cambrian subduction initiation). If this is the case, then both events must have occurred within the time interval between Middle to late Cambrian initiation and development of the suprasubduction granitoid bodies of the Ashland-Wedowee belt. In a simpler model, subduction initiation resulted in margin perpendicular to margin oblique subduction and that trench propagation developed preferentially within faulted, weak transitional crust at the continent-ocean boundary, with the resultant arc developed initially on the attenuated Laurentian margin. This most likely occurred during the Middle to late Cambrian, with development of the mature Alabama promontory arc by late Cambrian to Early Ordovician (Fig. 5.5). Figure 5.5 – Lithospheric collapse along the Alabama-Oklahoma transform, propagation through transitional crust at the continent-ocean boundary and development of the Alabama promontory subduction zone and continental margin arc. Adapted from Thomas and Astini, 1996. 98 CHAPTER 6 CONCLUSIONS In the southern Appalachians Talladega belt, the Middle Ordovician (~470 Ma) Hillabee Greenstone was tectonically emplaced atop rocks of the DevonianMississippian(?) Laurentian shelf Talladega Group prior to metamorphism of the amalgamated shelf-volcanic terrane at ~330 Ma (Middle Mississippian). Lithologic and geochemical characteristics of the Hillabee Greenstone strongly support its formation in a backarc basin in which magma genesis was influenced by Grenville-aged continental basement. A lack of deformational features associated with its pre-metamorphic emplacement suggests a palinspastic location proximal to the Laurentian shelf and limited tectonic transport prior to Alleghanian orogenesis. Ordovician backarc volcanic rocks, similar to the Hillabee, can be found along the western-central-eastern Blue Ridge metamorphic terrane boundaries in Alabama, Georgia, and North Carolina. Additionally, rocks of the central and eastern Blue Ridge, generally interpreted as representing distal Laurentian slope-rise deposits, are intruded by an array of Ordovician suprasubduction plutons. Rocks which might be interpreted as having formed in an accretionary subduction complex (accretionary prism) are completely absent in the southern Appalachian Blue Ridge of Alabama-Georgia, southwest of the Cartersville transverse zone, and no unequivocal accretionary prism rocks are present northeast of the Cartersville transverse zone in the Blue Ridge of North Carolina. Instead, rocks of the Alabama-Georgia-North Carolina eastern Blue Ridge, are dominated by turbidite sequences which formed along a continental margin consisting of Grenville-aged basement. The arrangement of lithotectonic belts in the southern Appalachians is not generally supportive of traditional Ordovician tectonic models which call on subduction of the Laurentian margin beneath an exotic or peri-Laurentian volcanic arc. Modern studies of arc-continent collisions (collisional orogens) suggest separation of subducting continental margin deposits from the accreting arc by extensive accretionary prism complexes, complete with LT-HP rocks – an arrangement of lithotectonic belts not observed in the southern Appalachians. Additionally, comparisons of the timing of 99 Ordovician igneous activity, metamorphism, and foreland basin clastic wedge deposition in the southern Appalachians with timing of analagous events in the latest Cenozoic collisional orogenic belts of Taiwan, Papua New Guinea, and Timor are marked by notable and important differences. First, suprasubduction igneous activity in the southern Appalachians continues for more than 20 m.y. following the onset of inferred Middle Ordovician orogenic loading and clastic wedge deposition in the Taconic foreland (Blount clastic wedge) – a significantly longer period of time than that observed in modern arc-continent collisions. Secondly, backarc extension (Hillabee GreenstonePumpkinvine Creek-Barlow Gneiss-Lake Burton) on the inferred overriding plate during the time in which the Laurentian continental margin was supposedly subducted continues well past the time interval in which compressive stresses should have caused backarc extension to cease, similar to that observed in Timor, where continental margin subduction resulted in cessation of backarc igneous activity in the Banda Sea. Finally, the general absence of arc rocks sensu stricto in the southern Appalachians forces Taconic models of collisional orgenesis to argue for burial of the arc “proper” rocks beneath the central-eastern Blue Ridge allochthons along the vast majority of the “Taconic suture” from Alabama to Virginia, while simultaneously placing a host of backarc terranes in exact juxtaposition with rock that formed along the Laurentian Iapent rifted margin.. While models of Ordovician collisional orogenesis are not generally compatible with the geology of the southern Appalachians, models of accretionary orogenesis largely support both the arrangement of lithotectonic terranes in the southern Appalachians and the timing of Ordovician tectonic events. Accretionary orogens, such as the Lachlan of southeastern Australia, have notable similarities with the southern Appalachian Taconic orogeny, including: A) formation of backarc rocks coeval with inferred orogenic activity on the continental margin, B) emplacement of backarc rocks along major faults at the leading edge of the pre-orogenic continental margin, C) the presence of extensive turbidite deposits intercalated with and outboard of backarc igneous rocks, D) short-lived extensional and contractional tectonic “switching” events complete with deformation, metamorphism, and igneous activity, E) regionally isolated metamorphic and deformational events along strike of the orogen, and F) a lack of 100 extensive accretionary prism deposits separating suprasubduction igneous rocks from rocks of the continental margin. A model of Andean-type subduction of Iapetan lithosphere and backarc extension along the Ordovician Laurentian margin explains the lack of Taconic metamorphism and deformation in the southernmost Appalachians of Alabama-Georgia, intrusion of the rifted Laurentian slope-rise (eastern Blue Ridge) by suprasubduction plutons, the juxtaposition of backarc rocks at the eastern-western Blue Ridge boundary of Alabama-Georgia, and the source of continentally influenced Middle Ordovician backarc volcanic rocks (Hillabee Greenstone) and k-bentonites deposits (Deicke and Millbrig) from the palinspastic location of the Alabama promontory far better than models calling for subduction of the Laurentian margin during the Ordovician. Conversion of the Cambrian passive, drifting Laurentian margin to an active subduction zone was accommodated by passage of the Ouachita rift axis during the Middle to latest Cambrian along the continent bounding Alabama-Oklahoma transform fault. Juxtaposition of young oceanic lithosphere outboard of the Ouachita embayment with mature Iapetan lithosphere across the seaward extension of the AlabamaOklahoma transform, coupled with slight convergence, created the conditions necessary for foundering of Iapetan lithosphere at the continent-ocean boundary within weak, serpentinized transitional crust. Dynamic topography associated with this Middle to latest Cambrian subduction initiation event resulted in tectonic destabilization of the distal Laurentian shelf and deposition of immature, intrabasinal deposits within the Laurentian Cambrian carbonate platform. Full-fledged subduction of oceanic lithosphere outboard of the continental margin hinge zone resulted in mature arc volcanism of the Alabama Promontory arc and intrusion of the Laurentian slope-rise by suprasubduction plutons during the latest Cambrian – earliest Ordovician, followed by backarc extension and the development of backarc volcanic suites proximal to the Laurentian margin. Rollback in the Iapetan subduction hinge allowed the active arc to migrate rapidly away from the Laurentian margin, limiting preservation of arc rocks sensu stricto in the Blue Ridge, but, formation of the suprasubduction igneous complex atop the attenuated Laurentian margin resulted in magmas influenced by Grenville continental basement during their genesis. Later tectonic events (Neoacadian and Alleghanian) resulted in exhumation of deeper portions of this backarc system along 101 younger faults, many of which may have exploited older faults associated with Ordovician (and younger?) tectonic switching phases, and overprinted Ordovician orogenic events with younger metamorphic and deformation phases. 102 APPENDIX Included with this dissertation are five plates. Plates 1 through 4 are a compilation of the twenty-two, 7.5 minute, 1:24,000 scale USGS quadrangles showing the geology in proximity to the Hollins Line fault system and Hillabee thrust. Plate 5 is a detailed chronology of significant geologic events, Neoproterozoic to Permian, in geologic regions of interest important to this study. References for ages of these events are listed in the text. 103 REFERENCES Abbott, L.D., and Chamalaun, F.H., 1981, Geochronology of some Banda Arc volcanics. The Geology and Tectonics of Eastern Indonesia. 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After working full time, Clinton returned to college in 1992, graduating with a B.S. in Geology from Florida State University in 1995. During the completion of Clinton’s doctoral work at Florida State, he taught full time at both Valdosta State University (4 years) and Columbus State University (2 years), in addition to serving as an adjunct instructor at Tallahassee Community College, Thomas University, and Florida State University, and working a number of part time jobs with the Florida Geological Survey and Florida Department of Environmental Protection. Clinton is married to Diedre Lloyd and the proud father of two children. 136
