Contrasting modes of supercontinent formation and ...

Gondwana Research 15 (2009) 408–420

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Gondwana Research j o u r n a l h o m e p a g e : w w w. e l s e v i e r. c o m / l o c a t e / g r

Contrasting modes of supercontinent formation and the conundrum of Pangea J. Brendan Murphy a,⁎, R. Damian Nance b, Peter A. Cawood c a b c

Department of Earth Sciences, St. Francis Xavier University, Antigonish, Nova Scotia, Canada B2G 2W5 Department of Geological Sciences, Ohio University, Athens, Ohio 45701, USA University of Western Australia, Tectonics Special Research Centre, 35 Stirling Highway, Crawley, WA 6009, Australia

a r t i c l e

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Article history: Received 19 July 2008 Received in revised form 8 September 2008 Accepted 9 September 2008 Available online 26 September 2008 Keywords: Pangea Geodynamics Pannotia Rodinia Supercontinent

a b s t r a c t Repeated cycles of supercontinent amalgamation and dispersal have occurred since the Late Archean and have had a profound influence on the evolution of the Earth's crust, atmosphere, hydrosphere, and life. When a supercontinent breaks up, two geodynamically distinct tracts of oceanic lithosphere exist: relatively young interior ocean floor that develops between the dispersing continents, and relatively old exterior ocean floor, which surrounded the supercontinent before breakup. The geologic and Sm/Nd isotopic record suggests that supercontinents may form by two end-member mechanisms: introversion (e.g. Pangea), in which interior ocean floor is preferentially subducted, and extroversion (e.g. Pannotia), in which exterior ocean floor is preferentially subducted. The mechanisms responsible remain elusive. Top–down geodynamic models predict that supercontinents form by extroversion, explaining the formation of Pannotia in the latest Neoproterozoic, but not the formation of Pangea. Preliminary analysis indicates that the onset of subduction in the interior (Rheic) ocean in the early Paleozoic, which culminated in the amalgamation of Pangea, was coeval with a major change in the tectonic regime in the exterior (paleo-Pacific) ocean, suggesting a geodynamic linkage between these events. Sea level fall from the Late Ordovician to the Carboniferous suggests that the average elevation of the oceanic crust decreased in this time interval, implying that the average age of the oceanic lithosphere increased as the Rheic Ocean was contracting, and that subduction of relatively new Rheic Ocean lithosphere was favoured over the subduction of relatively old, paleo-Pacific lithosphere. A coeval increase in the rate of sea floor spreading is suggested by the relatively low initial 87Sr/86Sr in late Paleozoic ocean waters. We speculate that superplumes, perhaps driven by slab avalanche events, can occasionally overwhelm top–down geodynamics, imposing a geoid high over a pre-existing geoid low and causing the dispersing continents to reverse their directions to produce an introverted supercontinent. © 2008 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.

1. Introduction The Phanerozoic Eon is dominated by the assembly and amalgamation of Pangea in the Paleozoic (Fig. 1), followed by its breakup and dispersal in the Mesozoic and Cenozoic. Although to a first order, there is a consensus on the paleocontinental reconstruction and timing of these events, (e.g. McKerrow and Scotese, 1990; Cocks and Fortey, 1990; Scotese, 1997; van Staal et al., 1998; Cocks and Torsvik, 2002; Stampfli and Borel, 2002; Veevers, 2004), the mechanisms responsible for the amalgamation of this supercontinent are poorly understood. Furthermore, over the last 20 years, evidence has been amassing that Pangea is just the latest in a series of supercontinents that have formed since the Archean, only to breakup and reform again (e.g. Rogers and Santosh, 2004; Silver and Behn, 2008). Although the causes remain elusive, many geoscientists agree that repeated cycles

⁎ Corresponding author. E-mail address: [email protected] (J.B. Murphy).

of supercontinent amalgamation and dispersal (the “supercontinent cycle” of Worsley et al., 1984 and Nance et al., 1988), have had a profound effect on the evolution of the Earth's crust, atmosphere, hydrosphere, and life (e.g. Worsley et al., 1984, 1986; Nance et al., 1986; Veevers, 1990, 1994; Condie, 1994, 1995; Hoffman et al., 1998; Ross, 1999; Condie, 2002; Knoll et al., 2004; Maruyama et al., 2007; Maruyama and Santosh, 2008; Stern, 2008; Meert and Lieberman, 2008; Rino et al., 2008). In this paper, we review the development of this concept and then use the history of Pangea to gain insights into potential mechanisms that might account for episodic supercontinent formation. We first trace these ideas to show how the concept of supercontinent cycles originated. We point out that oceanic lithosphere created and destroyed during this cycle has different geodynamic properties depending on whether the lithosphere was located around the supercontinent (“exterior” ocean floor) or formed between dispersing continents (“interior” ocean floor). We then review the evidence that supercontinents may have formed by different mechanisms. Example, the late Neoproterozoic supercontinent Pannotia (Powell, 1995;

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J.B. Murphy et al. / Gondwana Research 15 (2009) 408–420

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Fig. 1. Paleozoic reconstructions (modified from Scotese, 1997; Cocks and Torsvik, 2002; Stampfli and Borel, 2002; Murphy et al., 2006; Murphy and Nance, 2008). (A) By 540 Ma, the Iapetus Ocean had formed between Laurentia and Gondwana. By 460 Ma, Avalonia–Carolinia (A–C) had separated from Gondwana, creating the Rheic Ocean. By 370 Ma, Laurentia, Baltica and A–C had collided to form Laurussia, and the Rheic Ocean began to contract, closing by 280 Ma, to form Pangea. The Terra Australis orogen (Cawood, 2005) at 280 Ma was located along the periphery of Pangea, representing the vestige of tectonic activity within the paleo-Pacific Ocean.

Dalziel, 1997), which consists of the continental fragments of Gondwana and Laurentia, was formed by the preferential consumption of the exterior oceanic lithosphere (Mozambique Ocean), whereas Pangea was formed by the preferential subduction of interior oceanic lithosphere (Iapetus and Rheic oceans). The implications of different relative ages of oceanic and continental lithosphere interaction are significant when viewed in the light of geodynamic models used to explain supercontinent formation. For example, the formation of Pangea cannot be explained by most widely accepted geodynamic models, since these models, when applied to the widely accepted paleocontinental reconstructions for the Early Paleozoic, do not yield Pangea in the correct configuration (Murphy and Nance, 2008). Hence, a fundamental disconnection exists between the geologic evidence for supercontinent formation, and the models purported to explain their assembly. Finally, to provide constraints for future geodynamic models and to gain insight into the processes leading to the formation of Pangea, we investigate the geodynamic linkages between the Paleozoic evolution of the interior (Iapetus and Rheic) oceans, as recorded in the orogens (Ouachita, Appalachian, Caledonide, Variscan) produced by their closure (e.g. van Staal et al., 1998; Matte, 2001), and the pene-contemporaneous evolution of the exterior (paleo-Pacific) ocean, as recorded in the 18,000 km Terra Australis orogen, which preserves a record of sub-

duction from ca. 570 to 230 Ma (e.g. Cawood, 2005). In doing so, we show that proxy records for oceanic lithosphere development, such as sea level and isotopic data (Hallam, 1992; Veizer et al., 1999; Condie, 2004; Barnes, 2004; Miller et al., 2005), can be used to identify globalscale changes in plate geodynamics, and so provide a complementary approach to the analysis of supercontinent formation throughout geologic time. 2. Development of concepts The notion of episodic crustal development and orogenic activity actually predates the general acceptance of the plate tectonic paradigm. Holmes (1954) suggested that the development of continents took place through the episodic production of new crust. This concept was further developed by Gastil (1960), who drew attention to radiometric data suggesting that granite production throughout geologic time was episodic rather than continuous, by Sutton (1963), who used these data to define global-scale chelogenic or “shieldforming” cycles that he linked to orogenic activity, and by Armstrong (1968, 1981) who used isotopic evidence to argue for crustal recycling and no net continental growth since the early Archean. The advent of plate tectonics in the 1960s and its application to ancient orogenic belts (e.g. Wilson, 1966; Dewey, 1969) showed that

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crustal production and orogeny were linked to subduction of oceanic lithosphere, accretion of terranes and continental collisions. The increase in the number and precision of radiometric dates during the 1970s and early 1980s confirmed the episodic nature of orogenesis, and led to the hypothesis of a “supercontinent cycle” (Worsley et al., 1984, 1986; Nance et al., 1986) in which Pangea was recognized as the youngest of a series of supercontinents that had formed since the Archean. By implication, each supercontinent fragmented and dispersed, only to reform again at a later time (e.g., Hoffman, 1992). Although the details may be controversial, paleocontinental reconstructions for the past 2.5 Ga, based on a variety of lithotectonic, paleomagnetic, geochemical and faunal data, imply the existence of a quasiregular supercontinent cycle. Several end-member models have been proposed to explain this cycle. For example, Dalziel (1992) pointed out that chance collisions of buoyant continental crust are an inevitable consequence of plate motion and suggested that there may be a stochastic component to supercontinent formation. On the other hand, supercontinent fragmentation cannot have a stochastic component, and the presence of rift-related mafic dike swarms (e.g., Windley, 1977), riftto-drift continental margin successions that are accompanied by major changes in global sea level (e.g., Vail et al., 1977), climate (e.g., Fischer, 1981) and evolutionary biogenesis (e.g., Valentine and Moores, 1970), together with the episodicity of orogenesis (e.g., Condie, 1976), led Worsley et al. (1984, 1986) to propose a semi-regular cycle of supercontinent assembly and dispersal with a ca. 500 million year duration. By the early 1990s, several of these supercontinents had been named and the geological database had expanded to the point where testable reconstructions for the 0.6–0.55 Ga (Pannotia) and the 1.1–0.75 Ga (Rodinia) supercontinents were proposed (e.g. Dalziel, 1991, 1992; Moores, 1991; Hoffman, 1992; Stump, 1992; Powell, 1995). More recent reconstructions of these supercontinents are shown in Figs. 2 and 3 respectively (after Pisarevsky et al., 2003, 2008). During each supercontinent assembly, two distinct types of orogenic belts are developed (Murphy and Nance, 1989, 1991); interior (Murphy

Fig. 2. Palaeogeography at ca. 600 Ma (after Pisarevsky et al., 2008). Rifts between Baltica and Laurentia, and between Baltica and Amazonia are shown with solid lines; failed rift between Laurentia and Amazonia is shown with dotted lines. Subduction zones outboard of Baltica and Amazonia are directed in opposite directions. A-A — Afif-Abas; Am — Amazonia; Au — Australia; Av — Avalonia; Az — Azania; Ba — Baltica; Co — Congo; In — India; K — Kalahari; La — Laurentia; M — Mawson; O — Oaxaquia; P — Pampia terrane; Rp — Río de La Plata; S — Saharan Metacraton; SF — São Francisco; Si — Siberia.

Fig. 3. Rodinia at 990 Ma (modified after Pisarevsky et al., 2003; Murphy et al., 2004). Am — Amazonia; B — Barentsia; Ba — Baltica; Ch — Chortis; Gr — Greenland; La — Laurentia; O — Oaxaquia; P — Pampia terrane; R — Rockall; RP — Río de La Plata; WA — West Africa; In — India; Ka — Kalahari; Si — Siberia.

and Nance, 1991) or collisional (Windley, 1993) orogens, which form as a result of continental collisions and are stranded in the interior of the supercontinent after those collisions have taken place, and peripheral (Murphy and Nance, 1991) or accretionary (Windley, 1993; Cawood and Buchan, 2007; Cawood et al., in press) orogens, which develop in the oceanic realm surrounding the assembling continental land masses and end up along the periphery of the supercontinent. The late Neoproterozoic collision between East and West Gondwana to form the East African orogen (Stern, 1994), which was a major event in the formation of Pannotia, is an example of an interior orogen. At the same time, peripheral orogens such as the Avalonian and Cadomian belts, which formed along the periphery of northern Gondwana, developed in the oceanic realm surrounding the converging continents. Likewise for Pangea, the Alleghanian–Ouachita orogen of North America and the Variscan orogen of Europe are examples of interior orogens, whereas the Terra Australis orogen (Cawood and Buchan, 2007), which developed in the oceanic realm surrounding Pangea and ended up along its periphery, is an example of a peripheral orogen (Fig. 1). Theoretical models attempting to explain the supercontinent cycle simulate the geodynamic consequences of differential heat flow from the mantle though oceanic versus continental crust. Anderson (1982, 1994) proposed that the cyclic assembly of supercontinents occurs because supercontinents assemble at sites of convective downwelling in the mantle, and fragment because the insulating effects of stationary supercontinents induce mantle upwelling that uplifts and fragments them, and then disperses the continental fragments towards new areas of mantle downwelling. Such a supercontinent cycle has found support in time-dependent numerical and kinematic modeling (e.g. Gurnis, 1988; Duncan and Turcotte, 1994), which show that supercontinents inhibit cooling of the mantle beneath them, and so overheat and fragment under tension. In this scenario, plumes that accompany rifting and fragmentation are generated in the relatively shallow sub-lithospheric mantle. These geodynamic models explain the Pangea-like, latest Neoproterozoic through Early Cambrian breakup record of Pannotia (e.g., Veevers, 1989), which was accompanied by pene-


contemporaneous rifting and sea level rise (Bond et al., 1984; Hallam, 1992; Miller et al., 2005), dramatic changes in global climate and ocean water geochemistry (e.g., Veevers, 1990; Knoll, 1991; Hoffman et al., 1998; Hoffman and Schrag, 2002), and rapid metazoan diversification (e.g., Cowie and Brasier, 1989; McMenamin and McMenamin, 1990; Narbonne, 1998; Knoll et al., 2004; Narbonne, 2005). Recently, these geodynamic models have evolved to the proposition that plate tectonics is primarily driven by subduction of cold lithosphere, which provides 90% of the force needed to drive plate motions (Anderson, 1994, 2001). This view of “top–down” tectonics holds that subduction of cold lithosphere is indirectly responsible for the upwelling beneath mid-ocean ridges, and that supercontinents break up over sites of mantle upwelling — or geoid highs — and migrate to sites of mantle downwelling — or geoid lows. 3. Geodynamic framework When a supercontinent breaks up, two geodynamically distinct tracts of oceanic lithosphere exist (Fig. 4). The new interior oceans that form between the dispersing continents are underlain by oceanic lithosphere that is younger than the age of supercontinent breakup, whereas the exterior ocean surrounding the supercontinent is underlain by older oceanic lithosphere, the vast majority of which predates the time of supercontinent breakup. Because of their contrasting ages, the interior and exterior oceanic lithospheres at the time of breakup have contrasting average thicknesses and elevations (e.g. Sclater et al., 1980). A further consequence of the breakup geometry is that these contrasting tracts of oceanic lithospheres must come in contact with one another, such that the boundary between them will be tectonically unstable as the continents continue to diverge (Fig. 4). In theory, the geodynamic distinction between the interior and exterior oceans

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diminishes with time as the average age of the interior ocean increases, and subduction preferentially removes the oldest oceanic lithosphere from the exterior ocean. Following breakup, the formation of the next supercontinent can occur by way of two end-member models: (i) extroversion, by which subduction preferentially destroys the older oceanic lithosphere of the exterior ocean, while the interior oceans continue to grow (with or without ongoing subduction), and (ii) introversion, by which subduction develops and preferentially destroys the interior oceans with their geodynamically younger oceanic lithosphere (Fig. 4, Murphy and Nance, 2003, 2005, 2008). Which of these oceanic lithospheres preferentially subducts requires fundamentally different geodynamic forces. Both mechanisms have been advocated for the formation of the next supercontinent. If modern subduction in the Caribbean and Scotia arcs spreads along the Atlantic seaboard, then convergence and destruction of the Atlantic Ocean would result in a supercontinent (Pangea Ultima; Scotese, 2003) that would form by introversion. On the other hand, if no major plate re-organization occurs for the next 75 million years the Pacific Ocean will close, resulting in extroversion and a supercontinent configuration dubbed Amasia by Hoffman (1999). 4. Introversion and extroversion in the geologic record: Sm/Nd evidence The geological record suggests that both introversion and extroversion have occurred at different times in the geologic past. For example, most paleocontinental reconstructions imply that the breakup of the Late Mesoproterozoic supercontinent, Rodinia, and subsequent assembly of the Late Neoproterozoic supercontinent, Pannotia, occurred by preferential subduction of the exterior ocean and, hence, is an example of extroversion (e.g. Hoffman, 1992; Dalziel, 1992, 1997; Murphy and

Fig. 4. (A–C) Schematic diagrams (modified from Murphy and Nance, 2003, 2008) showing stages in the breakup of a supercontinent, the fate of relatively old oceanic lithosphere that surround the supercontinent before breakup (exterior ocean), and the creation of relatively new oceanic lithosphere between the dispersing blocks (interior ocean). The interior oceanic lithosphere is shown in the lighter shading, the exterior oceanic lithosphere is shown in the darker shading. In the case of Pangea, the paleo-Pacific represents the exterior ocean and the Iapetus and Rheic Oceans are examples of interior oceans. In B and C, note subduction commences along the boundaries between the interior and exterior oceans. (D) Configuration of a supercontinent formed by introversion (i.e. preferential subduction of the interior oceanic lithosphere). (E) Configuration of a supercontinent formed by extroversion (i.e. preferential subduction of the exterior oceanic lithosphere). (F) An intermediate case in which one ocean is closed by introversion, and the other by extroversion.

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Nance, 2003, 2005). In contrast, the latest Neoproterozoic (Ediacaran) breakup of Pannotia, followed by the late Paleozoic assembly of Pangea, occurred by the preferential subduction of the interior Iapetus and Rheic oceans between Laurentia, Baltica and Gondwana (e.g. van Staal et al., 1998; Stampfli and Borel, 2002), and so exemplifies introversion (Murphy and Nance, 2003, 2005). The validity of these reconstructions is born out by the Sm/Nd isotopic characteristics of vestiges of oceanic lithosphere that were accreted to continental margins during the two supercontinent cycles (see Murphy and Nance, 2003, 2005 for details). At the time of supercontinent breakup (TR), there is a clear distinction between the Sm/Nd isotopic characteristics of the oceanic lithosphere in the exterior and interior oceans (Fig. 5). Vestiges of the exterior ocean have depleted mantle model ages (TDM, DePaolo, 1981, 1988) that predate the age of supercontinent breakup (TDM N TR), whereas those of the interior ocean have depleted mantle model ages that postdate breakup (TR N TDM). 4.1. Assembly of Pannotia The late Neoproterozoic assembly of Gondwana was the principal event in the formation of Pannotia. Collision between the various continental blocks of East and West Gondwana (e.g. Hoffman, 1992; Stump, 1992; de Wit et al., 2001, 2008) was largely accomplished by closure of the Mozambique Ocean (Dalziel, 1992, 1997) and resulted in the formation of the ca. 5000 km-long East African Orogen (EAO; Stern, 1994, 2002). The EAO consists of a number of orogenic belts ranging in age from ca. 950 to 550 Ma. In the northern EAO, the Arabian-Nubian Shield is made up of peripheral orogens that were located along the northern margin of Gondwana. The shield is dominated by ca. 850– 650 Ma arcs and ca. 2.7–2.0 Ga continental microplates that accreted at various times during the Neoproterozoic (Pallister et al., 1988; Kröner et al., 1992). The arc complexes are separated by ophiolitic bodies that trace the sutures between them. Both the arc terranes and ophiolites typically plot within +1εNd of the depleted mantle curve (Stern, 2002). TDM model ages range from 0.66 to 1.26 Ga, with a mean of 0.85 Ga (Stern, 2002). The Sm/Nd isotopic data indicate that most of these complexes were formed from oceanic lithosphere older than the 0.76 Ga breakup age of Rodinia (i.e., TDM N TR), and so were part of the exterior ocean (Fig. 6A). Similar characteristics occur elsewhere within the EAO. In the southern EAO, the Mozambique Belt is an interior orogen consisting of highly deformed and metamorphosed rocks that are either vestiges of juvenile crust, or reworked Archean crust that

Fig. 6. Summary diagrams of (εNd)t vs. time (Ga) for inferred peri-Rodinian crust preserved in Neoproterozoic interior orogens. (A) East African Orogens (Stern et al., 1991; Kröner et al., 1992; Küster and Liégeois, 2001; Stern 2002, Stein, 2003); (B) Borborema and Tocantins provinces of the Brasiliano orogenic belt, Brazil (Pimentel and Fuck, 1992; van Schmus et al., 1995; Pimentel et al., 1997); (C) Trans-Saharan belt of West Africa (Mali from Caby et al., 1989; Algeria from Dostal et al., 2002). In each diagram, the evolution of εNd with time for peri-Rodinian oceanic lithosphere is shown in stipples and is defined by assuming a depleted mantle composition for the oceanic lithosphere formed between the time of amalgamation (A) and breakup (B) of Rodinia at ca. 1.0 Ga and 0.75 Ga respectively. The evolution assumes a typical crustal Sm/Nd ratio of 0.18. In (B), the stippled region is the Sm/Nd isotopic envelope for the Brasiliano orogens.

Fig. 5. Schematic representation of the Sm/Nd isotopic evolution of oceanic lithosphere from the interior and exterior oceans. Depleted mantle model ages for the Interior Ocean (TI) are younger than the time of supercontinent breakup (TR), whereas the depleted mantle model ages for the Exterior Ocean (TE) are older than the time of supercontinent breakup. Relative to the breakup of Rodinia, the Mozambique Ocean is an exterior ocean. Relative to the breakup of Pannotia, the Iapetus and Rheic Oceans are interior oceans (see Fig. 6 and text).

may form the basement to the Arabian-Nubian shield. The occurrence of highly dismembered ophiolites in suture zones that can be traced southward from the Arabian-Nubian Shield suggests that the Mozambique belt was generated by collisional tectonics (Stern and Dawoud, 1991; Stern, 2002). The transition zone between the peripheral orogens of the ArabianNubian Shield and the interior orogen of the Mozambique Belt occurs in


Yemen, Somalia and the Sudan (e.g. Stern and Dawoud, 1991; Lenoir et al., 1994; Windley et al., 1996; Küster and Liégeois, 2001; Stern 2002) and has similar isotopic characteristics to those of the Arabian-Nubian Shield. For example, Sm/Nd isotopic data from oceanic arc-back arc complexes in the Bayuda desert of Sudan reveal mafic and sedimentary sequences characterized by high εNd (+3.6 to +5.2 at t = 806 Ma) with TDM model ages of 0.78 to 0.90 Ga (Fig. 6A), and ca. 740 Ma granulites and charnockites interpreted to have formed as a result of collisional orogenesis with εNd data ranging from +2.9 to +3.4 (t = 740 Ma) and TDM model ages of 0.96 to 0.98 Ga. The EAO is consequently characterized by mafic complexes with TDM ages that predate the age of supercontinent breakup (TDM N TR), implying that the complexes represent vestiges of the exterior ocean. Other late Neoproterozoic orogens yield similar results. The Tocantins and Borborema provinces of Brazil, which are interior “Brasiliano” orogens lying between the Amazonian, West African–Sao Luis and Sao Francisco–Congo–Kasai cratons (e.g. de Wit et al., 1988), were highly deformed and metamorphosed by collision between these cratons between ca. 0.50 and 0.60 Ga. Accreted complexes in these provinces include ca. 0.95–0.85 Ga mafic meta-igneous complexes that originated in an ensimatic arc setting, and younger (ca. 760–600 Ma) arc-related rocks (Pimentel et al.,1997). Initial εNd values for these suites (calculated for the age of crystallization) range from +0.2 to +6.9, whereas TDM model ages lie between 0.9 and 1.2 Ga (Pimentel et al., 1997; Fig. 6B). The Trans-Saharan belt of West Africa (e.g. Caby et al., 1989) is thought to record arc–arc, arc–continent, and continent–continent collisional orogenesis in the Neoproterozoic that reflects convergence and collision between the East Saharan Shield and the West African craton (Caby et al., 1989; Black et al., 1994; Dostal et al., 2002). In Mali, ca. 730 Ma mafic to intermediate volcanic and plutonic complexes with calc alkalic and island arc tholeiitic affinities, formed above an east-dipping subduction zone and were accreted to the margin of the West African craton during late Neoproterozoic collisional orogenesis. εNd values for two of these complexes are high (+ 6.3 to +6.6 and +4.4 to 15.8 at t = 730 Ma) with TDM ages of 760–710 Ma and 940– 840 Ma, respectively (Fig. 6C; Caby et al., 1989). Sm/Nd isotopic data for inliers of Trans-Saharan mafic rocks in southwestern Algeria and southern Morocco have εNd values ranging from + 1.0 to +5.0 that yield TDM model ages of 0.95 to 1.20 Ga (Dostal et al., 2002). Taken together, these data suggest that mafic terranes with TDM model ages older than 0.76 Ga (the time of Rodinia breakup), and ranging up to 1.2 Ga, are widespread within Neoproterozoic interior and peripheral orogens. These model ages imply that much of the oceanic lithosphere that was subducted to yield these complexes was formed before the ca. 830–750 Ma breakup of Rodinia and, hence, must have formed within the peri-Rodinian (Mirovoi) ocean. 4.2. Assembly of Pangea Paleocontinental reconstructions show that Pangea was formed by the sequential closure of the Iapetus and Rheic oceans. The Iapetus Ocean formed in stages from ca. 610 to 530 Ma (Cawood et al., 2001) and its closure by ca. 420 Ma is attributed to collision between Gondwanandervied continents (e.g. Avalonia, Carolinia), Baltica and Laurentia (van Staal et al., 1998; Hibbard et al., 2002; Keppie et al., 2003). Vestiges of the Iapetus Ocean are preserved in the Early Ordovician (ca. 480 Ma) mafic complexes of the Dunnage Zone in Central and Western Newfoundland, Canada, which contain ophiolites, island arc tholeiites and boninites formed in a variety of supra-subduction zone environments (e.g. Elthon, 1991; Jenner et al., 1991; Kurth et al., 1998; van Staal et al., 1998). Uncontaminated mafic complexes have juvenile εNd values ranging from +4 and +7 (Fig. 7), with TDM model ages between 0.50 and 0.60 Ga, indicating that these complexes were generated from Iapetus oceanic lithosphere (e.g. Swinden et al., 1997; MacLachlan and Dunning, 1998) and, hence, are relicts of an interior ocean.

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Fig. 7. Summary diagram of (εNd)t vs. time (in Ga) with compilation of Sm/Nd isotope compositions from Iapetan and Rheic ocean complexes in the Paleozoic AppalachianCaledonide orogen. Field for peri-Rodinian oceanic lithosphere defined by time of amalgamation (a), and breakup (b) of Rodinia. Field for oceanic lithosphere of Iapetus and Rheic oceans are respectively defined between their times of initial opening (o) and closure (c). Isotope evolution of peri-Rodinian, Iapetan, and Rheic oceanic lithosphere calculated by assuming a typical crustal Sm/Nd ratio of 0.18. Data from Camiré et al. (1995), Bedard and Stevenson (1999), Swinden et al. (1997), Jenner and Swinden (1993), MacLachlan and Dunning (1998), Pin and Paquette, (1999), Nutman et al. (2001), Sandeman et al. (2000) and Castro et al. (1996).

Vestiges of the Early Ordovician–Carboniferous Rheic Ocean are preserved in several ophiolitic complexes in western Europe. Sm/Nd data are available from Devonian complexes (Fig. 7) such as the Lizard Complex of Britain (Davies, 1984), the Brevenne metavolcanics in the Massif Central (Pin and Paquette, 1999), the Aracena Metamorphic Belt in the Ossa-Morena zone of the southwest Iberian Massif (Castro et al., 1996), and various complexes in NW Iberia (e.g. Pin et al., 2002, 2006). The Lizard complex is characterized by juvenile MORB to OIB chemistry with εNd values between + 8 and +11. Uncontaminated Brevenne metavolcanics also have juvenile MORB to OIB chemistry with εNd values of + 5 and +8. The Aracena amphibolites have MORB-like geochemistry and yield εNd values of + 7.9 to +9.2. In northern Iberia, several mafic complexes have a supra-subduction zone chemistry (e.g. Arenas et al., 2007) and collectively yield εNd values of +6.4 to + 9.2. In all these complexes, εNd values lie close to, or above, the typical value for the contemporary depleted mantle, indicating that they have a juvenile composition. They are therefore derived from Rheic oceanic lithosphere and are vestiges of this interior ocean. In summary, Sm/Nd isotopic data from mafic complexes within Neoproterozoic orogens are consistent with their derivation from the exterior ocean, and yield TDM model ages ranging from 0.8 to 1.2 Ga, which clearly predates the age of the breakup of Rodinia. In contrast, mafic complexes formed during subduction of the Iapetus and Rheic oceans have juvenile compositions and TDM model ages that are clearly younger than the breakup of Pannotia, implying their derivation from interior oceans. Taken together, these data support the implications of paleocontinental reconstructions that Pannotia formed by extroversion, whereas Pangea formed by introversion, and that supercontinents can therefore form by different geodynamic mechanisms. 5. Geodynamic conundrum of Pangea The “top–down” geodynamic models of Anderson (1982, 1994, 2001) and Gurnis (1988), in which supercontinents breakup over geoid highs and migrate away from those highs to reassemble over geoid lows (represented by subduction zones, Fig. 8), predict that

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Fig. 8. Time-dependent numerical models of supercontinent breakup and dispersal followed by reassembly and amalgamation (after Gurnis, 1988). Supercontinents insulate the mantle beneath them, which overheats (A) and the resulting continents (C) disperse towards subduction zones in the exterior ocean (B), where they amalgamate to form a new supercontinent (C), and the cycle begins again (D). This model predicts that supercontinents form by extroversion. The subduction zone shown in A and in B is the same. It has been centered in B to show the continents converging towards it.

supercontinents form by extroversion, unless there is a fundamental change in the location of the geoid anomalies. Hence, they provide an adequate explanation for the formation of Gondwana/Pannotia from the 0.83–0.75 Ga breakup of Rodinia. However, they do not provide an explanation for the formation of introverted supercontinents, and so cannot explain the formation of Pangea. If these geodynamic models are applied to a 500 million-year-old paleogeography (i.e. before the onset of subduction in the Iapetus Ocean and before the opening of the Rheic Ocean), Pangea would have formed by closure of the paleo-Pacific exterior ocean rather than by closure of the Iapetus and Rheic interior oceans. The Neoproterozoic to late Paleozoic record of tectonic activity in the paleo-Pacific ocean is preserved in the Terra Australis Orogen, which, in a Pangea reconstruction, is a peripheral orogen that forms a continuous belt, 18,000 km in length, along the southern and western periphery of Gondwana (Fig. 1; Cawood, 2005; Cawood and Buchan, 2007). The Terra Australis Orogen contains continental and oceanic basement blocks with peri-Gondwanan and intra-oceanic affinities that were accreted to the Gondwanan margin at various times during the Paleozoic (Cawood, 2005; Cawood and Buchan, 2007; Ramos, 2008). Subduction commenced at ca. 570 Ma, and continued along the margin of the supercontinent until the breakup of Pangea in the Cretaceous (Cawood, 2005; Ramos, 2008). The continuity of the Terra Australis Orogen is apparent on late Paleozoic reconstructions where it stretches from Northern Australia, through Antarctica and the southern tip of South Africa to the western margins of South America. Subduction of the paleo-Pacific Ocean was well established by ca. 570 Ma, which predates the opening of the Iapetus Ocean (Cawood,

2005). Development of the Terra Australis Orogen culminated with the ca. 300–230 Ma Gondwanide Orogeny, which records peripheral orogenic activity pene-contemporaneous with the formation of interior orogens formed by the continent–continent collisions that amalgamated Pangea. As noted by Cawood (2005), this protracted history of ongoing subduction from the late Neoproterozoic to the late Paleozoic contrasts with the record of ocean opening and closing preserved in the interior orogens of Pangea. The 570 Ma record of subduction in the Terra Australis Orogen, when considered in conjunction with late Neoproterozoic–Early Cambrian paleocontinental reconstructions, suggests that when Pannotia broke up, the dispersing continents initially migrated towards the already established subduction zones in the paleo-Pacific, in accordance with top–down geodynamic models, and continued to do so until about 500 Ma. According to these models, however, slab-pull forces should have continued to draw the dispersing continents towards the paleoPacific subduction zones, eventually closing the paleo-Pacific ocean to form an extroverted supercontinent. However, paleocontinental reconstructions from the Early Ordovician through to the Carboniferous (Fig. 1) clearly imply that when subduction first began in the interior oceans at about 500 Ma, the rate of subduction was sufficient to close the interior oceans, resulting in the amalgamation of Pangea by introversion in the late Paleozoic. This conundrum focuses attention on the mechanisms of subduction initiation within the Iapetus Ocean, where subduction of the interior oceans first began. The onset of subduction in the interior Iapetus Ocean is documented by van Staal et al. (1998, 2007) at ca. 510 Ma along the Laurentian and Gondwanan margins. Subduction


directed away from Laurentia resulted in the formation of an ensimatic island arc within Iapetus. Collision between this arc and a Laurentiaderived microcontinent known as the Dashwood terrane, resulted in subduction zone flip and the obduction of ophiolites that formed in an arc environment. These tectonic events are generally assigned to the Taconic (or Taconian) Orogeny in Laurentia and the Grampian Orogeny in the Caledonides of Britain. Modern geodynamic models suggest that the most likely sites for subduction zone formation are adjacent to transform faults that juxtapose oceanic lithosphere of contrasting age (Casey and Dewey, 1984; Stern 2004). This is the case for modern intraoceanic subduction initiation along the Isu–Bonin–Mariana subduction zone, which originated along a transform fault that brought relatively young, buoyant oceanic lithosphere into contact with oceanic lithosphere that was older and denser (Stern and Bloomer, 1992; Stern, 2004). Subduction initiation was then followed by rapid trench retreat and voluminous mafic (including boninitic) magmatism (Fig. 9). At ca. 500 Ma, significant age contrasts across transform faults are less likely within the young Iapetus Ocean than they are along boundaries between the young oceanic lithosphere of Iapetus and that of the older paleo-Pacific (see Fig. 4B,C). Such boundaries between interior and exterior oceanic lithospheres are a geometric requirement of supercontinent breakup (Fig. 4). Moreover, the Iapetus ocean ridge is likely to have been oriented at a high angle to such boundaries, such that spreading along the ridge would have resulted in a strong strike-slip component (Fig. 4B,C). Such settings are therefore favourable environments for the generation of transform faults that separate oceanic lithosphere of contrasting ages. The age of the paleo-Pacific oceanic

Fig. 9. Subduction infancy model (from Stern, 2004), showing the development of a subduction zone where two oceanic lithospheres of differing density are juxtaposed across a transform fault. In the context of this paper, such conditions are most likely to occur along the boundary between interior and exterior oceans after supercontinent breakup (Fig. 4B, C). A and B shows initial condition in cross-section and map view respectively. In C and D, old, dense lithosphere sinks asymmetrically. In the context of the model presented in this paper, this would reflect subduction of exterior oceanic lithosphere beneath the interior oceanic lithosphere. In C and D, the asthenosphere migrates above the subducting lithosphere. In E and F, extension occurs resulting in ocean-floor type magmatism and the development of an infant arc.

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lithosphere adjacent to these transform faults will probably never be known. However, it is almost certain to have been older than the adjoining Iapetus oceanic lithosphere. The age and density contrast between the two oceanic lithospheres would have been most pronounced when the Iapetus Ocean was young. Such contrasts in the ages of oceanic lithospheres also occur across the boundaries between interior and exterior oceanic lithospheres (Murphy and Nance, 2008), which therefore provide a favourable geodynamic scenario for onset of subduction and the origin of the ca. 500 Ma ensimatic Iapetan arc complexes. The evolution of this plate boundary would have primarily depended on the spreading rate at the Iapetus ridge, the drift of the dispersing continents, and the rate of roll-back of the subduction zone. The documentation of boninites and mafic sheeted dikes in Iapetus ophiolitic complexes (e.g. Jenner et al., 1991; Bédard et al., 1998) indicates that fore arc extension and, hence, slab roll-back, did occur (cf. Cawood and Suhr, 1992). In this scenario, the paleo-Pacific lithosphere is subducted beneath that of Iapetus. Hence, the ensimatic arc complexes form in the overriding Iapetus plate, and, as the continents continued to diverge, these complexes would have become stranded within the Iapetus realm. This situation is broadly analogous to the Mesozoic–Cenozoic “capture” of the subduction zones around the Caribbean plate within the Atlantic realm (e.g. Pindell et al., 2006). The strike of an arc complex produced in this fashion would have been oriented at a high angle to the Iapetan ocean ridge, a geometry that would have facilitated its subsequent obduction onto the margins of Laurentia and Baltica during the ca. 500–480 Ma Taconic and Grampian orogenies (Fig. 4). Indeed, the geometry of spreading that generated the Bay of Islands ophiolite in Newfoundland, and the subsequent kinematics of its obduction, require generation at the inside corner of a transform fault with spreading at a high angle to the Laurentian margin and the inferred orientation of the arc complex (Karson and Dewey, 1978; Cawood and Suhr, 1992, Suhr and Cawood, 1993, 2001). This model consequently explains the initiation of subduction in the young Iapetan realm, the generation of supra-subduction zone ophiolitic complexes, and the obduction of these complexes soon after they were formed. As the Iapetus Ocean contracted between 500 and 420 Ma, paleocontinental reconstructions indicate that the Rheic Ocean formed as the result of the separation of terranes such as Avalonia and Carolinia from the northern Gondwanan margin, and expanded as the Iapetus Ocean contracted (Fig. 1). After the Iapetus Ocean closed by collision between Laurentia, Baltica and Avalonia to form Laurussia, northeast-directed subduction commenced beneath the Laurussian margin between ca. 440 and 420 Ma. This subduction ultimately closed the Rheic Ocean (van Staal et al., 1998; Martinez Catalan et al., 2007), resulting in the amalgamation of Pangea. In this scenario, therefore, the age of the Rheic Ocean lithosphere initially subducted was no more than 60 million years. Furthermore, the rate of subduction of Rheic oceanic lithosphere must have been greater than that of the paleo-Pacific, despite the fact that the age of the Rheic oceanic lithosphere initially subducted was significantly younger and probably became younger still as the Rheic Ocean contracted. If events in the paleo-Pacific exterior ocean were geodynamically linked to those in the Rheic interior ocean, then the style of subduction in the paleo-Pacific should have dramatically changed as the interior ocean started to contract. The Terra Australis Orogen preserves vestiges of paleo-Pacific subduction events during this crucial time period (Cawood, 2005), which are best documented in the Lachlan fold belt of eastern Australia (e.g. Collins, 2002, Fig. 10). Here, Ordovician rocks consist of interbedded shallow marine limestone and clastic sediments. However, in the Early to Middle Silurian, there is a dramatic change in tectonic environment and a series of ensimatic arcs associated with extension and roll-back started to form (Gray and Foster, 2004). The cause of this major change at ca. 440 Ma is a subject of much debate (see Gray and Foster, 2004, Fig. 11) and various models have been proposed, including those involving mantle heat input (also known as surge tectonics), intra-plate stress transfer involving convergence, intra-plate

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Fig. 10. Time–space diagram (after Collins, 2002) of the eastern Lachlan and New England orogens (eastern Australia), which are part of the Terra Australis Orogen. Extension events are marked by troughs and basins, contraction events are marked by unconformities. See text for discussion. BA — Baldwin Arc; CT — Cowra Trough; D — Dulladerry Rift; ECY — EdenComerong-Yalwal Rift; HET — Hill End Trough; MQ — Macquarie Arc; NB — Ngunawal Basin; NbB — Nambucca Basin; SyB — Sydney Basin; TT — Tumut Trough (A and B are rift-related sequences). Note the major change in tectonic environment between 440 and 420 Ma.

stress transfer involving divergence, and the development of multiple subduction zone systems. 6. Proxy records The dramatic change in tectonic environment recorded in the Terra Australis Orogen at precisely the time that the Rheic Ocean begins to subduct is consistent with a geodynamic linkage between events in the interior and exterior oceans. To examine this potential linkage further, we can look at proxy records that document on a broad scale, events within the global oceanic domain. For example, tectonic events can effect sea level by as much as 100 m on timescales varying from one million to 100 million years (e.g. Miller et al., 2005), with midocean ridge development resulting in a sea level rise, and the onset of subduction resulting in a sea level fall. The extent of these sea level changes depends on a number of factors, including the dimensions, respectively, of the ridge and the oceanic trench. Global sea level curves for the early Paleozoic (e.g. Hallam, 1992, Fig. 12) show a dramatic rise throughout the Cambrian, which is consistent with the opening of the Iapetus Ocean and the development of new oceanic ridge lithosphere. By the end of the Cambrian, however, global sea level fell by about 20 m, which is coincident with the onset of subduction within the Iapetus Ocean and, therefore, the development of trenches in and around the Iapetan realm. After these subduction zone systems stabilized, sea level began to rise once more as the Iapetus

Ocean continued to spread. However, a dramatic change in sea level occurred at ca. 440 Ma coincident with the onset of subduction within the Rheic Ocean and, to a first order, this drop in sea level continued until the amalgamation of Pangea at the end of the Paleozoic. Sea level fall from the Late Ordovician to the Carboniferous implies that the average elevation of the oceanic crust decreased during this time period. This, in turn, implies that the average age of the oceanic lithosphere increased as the Rheic Ocean was contracting. The reconstructions require that the increased sea floor spreading in the paleo-Pacific Ocean that must have accompanied convergence in the interior Rheic Ocean was not compensated by subduction of old lithosphere. Instead, the increased spreading rate in the paleo-Pacific Ocean was compensated by subduction of newer Rheic Ocean lithosphere. Hence, subduction of relatively young, buoyant Rheic Ocean lithosphere was favoured over the subduction of relatively old, thick, and dense paleo-Pacific ocean lithosphere. This scenario is incompatible with subduction-driven top–down tectonics and, instead, suggests the convergence across the interior ocean that led to the closure of the Rheic Ocean was imposed by some other mechanism. Initial 87Sr/86Sr ratios provide proxy records of times of rapid ocean spreading (low initial 87Sr/86Sr) versus enhanced continental weathering (high initial 87Sr/86Sr). Hence, the relatively low initial 87Sr/86Sr deduced for late Paleozoic ocean waters (Fig. 13; Veizer et al., 1999) implies an increase in the rate of sea floor spreading. Taken in tandem, the sea level and Sr isotope data indicate that any increase in sea floor


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Fig. 11. A summary (see Gray and Foster, 2004) of the various models proposed to explain the major change in tectonic environment at ca. 440 Ma in the eastern Lachlan and New England orogens. Models proposed include (A) significant mantle heat input, (B) intra-plate stress transfer associated with convergence, (C) intra-plate stress transfer associated with divergence, and (D) the development of multiple subduction zones.

spreading in the paleo-Pacific Ocean during Pangea assembly was compensated by subduction of relatively new Rheic Ocean lithosphere, rather than by subduction of old paleo-Pacific lithosphere. If, following the breakup of Pannotia, the dispersing continents initially moved from geoid highs towards geoid lows, as implied by top– down tectonic models, then any reversal in their direction of motion suggests that the geoid low towards which the continents were moving became a geoid high. The mechanisms by which this may occur are unclear, but could be fundamental to our understanding of the processes that gave rise to the formation of Pangea. Modelling by Zhong et al.

Fig. 12. Global sea level curve for the Phanerozoic (after Hallam, 1992). A dramatic rise occurs in the Cambrian, consistent with the opening of the Iapetus Ocean. Global sea level drop at the end of the Cambrian is attributed to the onset of subduction within the Iapetus Ocean. The change in sea level at ca. 440 Ma is coincident with the onset of subduction within the Rheic Ocean. The rise in sea level in the Jurassic and Cretaceous coincides with the opening of the Atlantic Ocean.

(2007) shows that the style of long wavelength mantle convection patterns alternates between one with two major antipodal upwellings when a supercontinent is present, to one with major upwelling in one hemisphere and major downwelling in the other when the continents are dispersed. However, the relationship of these models to supercontinent formation by introversion versus extroversion is unclear. Although speculative, some published geodynamic models that propose scenarios for rapid mantle upwelling (e.g. Gurnis, 1988; Tan et al., 2002) may provide clues to the mechanisms involved. Although different in detail, these models have in common a significant role for the subducted slab within the mantle. Recent tomographic images (e.g. Hutko et al., 2005) and geochemical and isotopic data from oceanic basalts (e.g. Tatsumi, 2005) indicate that subduction zones penetrate through the mantle to the core–mantle boundary. Modeling by Zhong and Gurnis (1997) suggests that subducted slabs may initially accumulate at the 670 km seismic discontinuity in the mantle, leading to a potentially global-scale slab avalanche into the lower mantle. The 670 km discontinuity probably corresponds to a solid–solid phase transition (spinel or ringwoodite to perovskite +magnesiowüstite) with a negative Clapeyron slope, and slab accumulation at this discontinuity

Fig. 13. 87Sr/86Sr variations in the Phanerozoic sea water (see Veizer et al., 1999 for the original dataset).

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depends on the degree of negativity of this slope. Pysklywec et al. (2003) point out that the phase transition, coupled with a viscosity contrast across the discontinuity, would act as a temporary barrier to slab penetration into the lower mantle, a scenario consistent with seismic tomographic images of widespread of subducted material within the transition zone (e.g., Fukao et al., 2001). As the subducted slabs approach and accumulate along the core– mantle boundary, hot fluids are pushed aside and plumes form at the periphery of the slab accumulation. According to Condie (1998), such slab avalanches are correlated with juvenile crust formation that is the surface expression of superplumes that rise from the core–mantle boundary in the aftermath of these avalanches. Subducted slabs ponding at the core–mantle boundary may also have a blanketing effect (Tan et al., 2002) by providing thermal insulation that results in the ponding of hot mantle material beneath the slabs. This ponded material overheats and eventually penetrates through the slab accumulation, rising to the surface as a superplume. Irrespective of the causal mechanism, a superplume event would result in a geoid high and enhanced sea floor spreading that, if centred in the paleo-Pacific Ocean, could have been responsible for the reversal in the direction of motion of the continents. Indeed, the present-day Pacific geoid high, which is antipodal to the geoid high that occupies the former position of Pangea (Anderson, 1982), may represent a vestigial record of this superplume. In this scenario, “top–down” tectonics of the early Paleozoic may have given way to “bottom–up” tectonics in the Late Paleozoic. The stratigraphic record in the Ordovician is compatible with such a superplume event, as indicated by a number of plume proxies, including enhance biological activity, global warming events, the formation of ironstones and black shales, and elevated sea level (Condie, 2004; Barnes, 2004). Moreover Sm/Nd isotopic systematics suggest a major addition of juvenile crust at ca. 450 Ma (Condie et al., 2009), an event that is consistent with a superplume. 7. Conclusions Geodynamic models of supercontinent cycles involve continental breakup over geoid highs and the movement and re-amalgamation of the continents over geoid lows (e.g. Anderson, 2001). Such models imply a top–down geodynamic driver in which continental amalgamation is controlled by surface plates and the location of subduction zones, which correspond to geoid lows. These models require that supercontinents form by extroversion in which the exterior ocean surrounding the supercontinent is consumed as interior oceans open and break the supercontinent apart. The breakup of the end-Mesoproterozoic supercontinent Rodinia and formation of the next supercontinent, the end-Neoproterozoic Pannotia, appears to have evolved by this mechanism. However, applying such models to the Earth's Paleozoic paleogeography would not produce Pangea in the correct configuration. Instead, paleocontinental reconstructions indicate that Pangea was assembled by the preferential subduction of new interior oceans rather than the geodynamically less buoyant, older oceanic lithosphere of the paleoPacific Ocean, as predicted by top–down geodynamics. This conundrum highlights potential geodynamic linkages between interior and exterior oceans and the need to look at critical global events within the oceans that record these linkages. It also suggests that superplumes, perhaps driven by slab avalanche events, may occasionally overwhelm top– down geodynamics imposing a geoid high over a pre-existing geoid low, and cause dispersing continents to reverse their directions and close the interior oceans, as was the case for Pangea. Acknowledgments We are grateful to Victor Ramos and an anonymous reviewer for their insightful and constructive comments. JBM acknowledges the continuing support of the Natural Sciences and Engineering Research Council, Canada through Discovery and Research Capacity grants. RDN is supported by

National Science Foundation grant EAR-0308105 and a Baker Award from Ohio University, and PAC acknowledges the support of the Australian Research Council. This paper is a contribution to the International Geoscience Programme, IGCP Projects 453 and 497 and to the International Lithosphere Program, Taskforce 1; Earth Accretionary Systems. References Anderson, D.L., 1982. Hotspots, polar wander, Mesozoic convection and the geoid. Nature 297, 391–393. Anderson, D.L., 1994. Superplumes or supercontinents? Geology 22, 39–42. Anderson, D.L., 2001. Enhanced: top–down tectonics? Science 293, 2016–2018. Arenas, R., Martínez Catalán, J.R., Sánchez Martínez, S., Fernández-Suárez, J., Andonaegui, P., Pearce, J.A., Corfu, F., 2007. The Vila de Cruces Ophiolite: a remnant of the early Rheic Ocean in the Variscan suture of Galicia (Northwest Iberian Massif). Journal of Geology 115, 129–148. Armstrong, R.L., 1968. A model for Pb and Sr isotope evolution in a dynamic earth. Reviews of Geophysics 6, 175–199. Armstrong, R.L., 1981. Radiogenic isotopes; the case for crustal recycling on a nearsteady-state no-continental-growth Earth. Royal Society of London Philosophical Transactions 301, 443–472. Barnes, C.R., 2004. Was there an Ordovician superplume event? In: Webby, B.D., Paris, F., Droser, M.L., Percival, I.G. (Eds.), The Great Ordovician Biodiversification Event. Columbia University Press, New York, pp. 77–80. Bedard, J.H., Stevenson, R., 1999. The Caldwell Group lavas of southern Quebec: MORBlike tholeiites associated with the opening of the Iapetus Ocean. Canadian Journal of Earth Sciences 36, 999–1019. Bédard, J.H., Lauziere, K., Tremblay, A., Sangster, A., 1998. Evidence for forearc seafloorspreading from the Betts Cove ophiolite, Newfoundland: oceanic crust of boninitic affinity. Tectonophysics 284, 233–245. Black, R., Latouche, L., Liegeois, J.P., Caby, R., Bertrand, J.M., 1994. Pan-African displaced terranes of the Tuareg shield (central Sahara). Geology 22, 641–644. Bond, G.C., Nickerson, P.A., Kominz, M.A., 1984. Breakup of a supercontinent between 625 and 555 Ma: new evidence and implications for continental histories. Earth and Planetary Science Letters 70, 325–345. Caby, R., Andreopoulos-Renaud, U., Pin, C., 1989. Late Proterozoic arc/continent and continent/continent collision in the Trans-Saharan belt. Canadian Journal of Earth Sciences 26, 1136–1146. Camiré, G., Laflèche, M.R., Jenner, G.A.,1995. Geochemistry of pre-Taconian mafic volcanism in the Humber Zone of the northern Appalachians. Chemical Geology 119, 55–77. Casey, J.F., Dewey, J.F., 1984. Initiation of subduction zones along transform and accreting plate boundaries, triple junction evolution, and forearc spreading centres: implications for ophiolitic geology and obduction. In: Gass, I.G. (Ed.), Ophiolites and Oceanic Lithosphere. Geological Society of London, Special Publication, vol. 13, pp. 269–290. Castro, A., Fernandez, C., de la Rosa, J., Moreno-Ventas, I., Rogers, G., 1996. Significance of MORB-derived amphibolites from the Aracena metamorphic belt, southwest Spain. Journal of Petrology 37, 235–260. Cawood, P.A., 2005. Terra Australis Orogen: Rodinia breakup and development of the Pacific and Iapetus margins of Gondwana during the Neoproterozoic and Paleozoic. Earth-Science Reviews 69, 249–279. Cawood, P.A., Suhr, G., 1992. Generation and obduction of ophiolites; constraints for the Bay of Islands Complex, western Newfoundland. Tectonics 11 (4), 884–897. Cawood, P.A., Buchan, C., 2007. Linking accretionary orogenesis with supercontinent assembly. Earth Science Reviews 82, 217–256. Cawood, P.A., McCausland, P.J.A., Dunning, G.R., 2001. Opening Iapetus: constraints from the Laurentian margin of Newfoundland. Geological Society of America Bulletin 113, 443–453. Cawood, P.A., Kröner, A., Collins, W.J., Kusky, T.M., Mooney, W.D. and Windley, B.F., in press. Earth accretionary systems in space and time. In: P.A. Cawood and A. Kröner (Editors), Accretionary Orogens in Space and Time. Geological Society, London, Special Publications, London. Cocks, L.R.M., Fortey, R.A., 1990. Biogeography of Ordovician and Silurian faunas. In: McKerrow, W.S., Scotese, C.R. (Eds.), Geological Society Memoir. Paleozoic Paleogeography and Biogeography, 12, pp. 97–104. Cocks, L.R.M., Torsvik, T.H., 2002. Earth geography from 500 to 400 million years ago: a faunal and palaeomagnetic review. Journal of the Geological Society, London 159, 631–644. Collins, W.J., 2002. Hot orogens, tectonic switching, and creation of continental crust. Geology 30, 535–538. Condie, K.C., 1976. Plate Tectonics and Continental Drift. Pergamon Press, New York, N.Y. Condie, K.C., 1994. Greenstones through time. In: Condie, K.C. (Ed.), Archean Crustal Evolution, Ch. 3. Elsevier, Amsterdam, pp. 85–120. Condie, K.C., 1995. Episodic ages of greenstones: a key to mantle dynamics? Geophysical Research Letters 22, 2215–2218. Condie, K.C., 1998. Episodic continental growth and supercontinents: a mantle avalanche connection. Earth and Planetary Science Letters 163, 97–108. Condie, K.C., 2002. The supercontinent cycle: are there two patterns of cyclicity? Journal of African Earth Sciences 35, 179–183. Condie, K.C., 2004. Supercontinents and superplume events: distinguishing signals in the geologic record. Physics of The Earth and Planetary Interiors 146, 319–332. Condie, K.C., Belousova, E., Griffin, WL., Sircombe, K.N., 2009. Granitoid events in space and time: constraints from igneous and detrital zircon age spectra. Gondwana Research 15, 228–242. doi:10.1016/j.gr.2008.06.001.

J.B. Murphy et al. / Gondwana Research 15 (2009) 408–420 Cowie, J.W., Brasier, M.D., 1989 (Eds.). The Precambrian–Cambrian boundary. Oxford: Clarendon press, 213pp. Dalziel, I.W.D., 1991. Pacific margins of Laurentia and East Antarctica–Australia as a conjugate rift pair: evidence and implications for an Eocambrian supercontinent. Geology 19, 598–601. Dalziel, I.W.D., 1992. On the organization of American plates in the Neoproterozoic and the breakout of Laurentia. GSA Today 2 (11), 238–241. Dalziel, I.W.D., 1997. Overview: Neoproterozoic–Paleozoic geography and tectonics: review, hypotheses and environmental speculations. Geological Society of America Bulletin 109, 16–42. Davies, G.R., 1984. Isotopic evolution of the Lizard Complex. Journal of the Geological Society, London 141, 3–14. DePaolo, D.J., 1981. Neodymium isotopes in the Colorado Front range and crust-mantle evolution in the Proterozoic. Nature 291, 193–196. DePaolo, D.J., 1988. Neodymium Isotope Geochemistry: an Introduction. SpringerVerlag, New York. 187pp. Dewey, J.F., 1969. Evolution of the Appalachian–Caledonian orogen. Nature 222, 124–129. de Wit, M., Jeffrey, M., Bergh, H., Nicolaysen, L., 1988. Geological map of sectors of Gondwana reconstructed to their disposition ca. 150 Ma. American Association of Petroleum Geologists, Tulsa OK, scale 1:10,000,000. de Wit, M.J., Bowring, S.A., Ashwal, L.D., Randrainasolo, L.G., Morel, V.P.I., Rambelson, R.A., 2001. Age and tectonic evolution of Neoproterozoic ductile shear zones in southwestern Madagascar, with implications for Gondwana studies. Tectonics 20, 1–45. de Wit, M.J., de Brito Neves, B.B., Trouw, R.A.J., Pankhurst, R.J., 2008. Pre-Cenozoic correlations across the South Atlantic region: (the ties that bind). Geological Society, London, Special Publications, vol. 294, pp. 1–8. Dostal, J., Caby, R., Keppie, J.D., Maza, M., 2002. Neoproterozoic magmatism in Southwestern Algeria (Sebkha el Melah inlier): a northerly extension of the Trans-Saharan orogen. Journal of African Earth Sciences 35, 213–225. Duncan, C.C., Turcotte, D.L., 1994. On the breakup and coalescence of supercontinents. Geology 22, 103–106. Elthon, D., 1991. Geochemical evidence for formation of the Bay of Islands ophiolite above a subduction zone. Nature 354, 140–143. Fischer, A.G.,1981. Climatic oscillations in the biosphere. In: Nitecki, M.H. (Ed.), Biotic Crises in Ecological and Evolutionary Time. Academic Press, New York, N.Y., pp. 17–37. Fukao, Y., Widiyantoro, S., Obayashi, M., 2001. Stagnant slabs in the upper and lower mantle transition region. Review Geophysics 39, 291–323. Gastil, G., 1960. Continents and mobile belts in the light of mineral dating. 2. International Geological Congress (Pt.9), 162–169 1960. Gray, D.R., Foster, 2004. Tectonic evolution of the Lachlan Orogen, southeast Australia: historical review, data synthesis and modern perspectives. Australian Journal of Earth Sciences 51, 773–817. Gurnis, M., 1988. Large-scale mantle convection and the aggregation and dispersal of supercontinents. Nature 322, 695–699. Hallam, A., 1992. Phanerozoic Sea-level. Columbia University Press, New York. 266p. Hibbard, J.P., Stoddard, E.F., Secor, D.T., Dennis, A.J., 2002. The Carolina Zone: overview of Neoproterozoic to early Paleozoic peri-Gondwanan terranes along the eastern flank of the southern Appalachians. Earth-Science Reviews 57, 299–339. Hoffman, P.F., 1992. Supercontinents. Encyclopedia of Earth Systems Science, vol. 4. Academic Press, London, pp. 323–328. Hoffman, P.F., 1999. The break-up of Rodinia, birth of Gondwana, true polar wander and the snowball Earth. Journal of African Earth Sciences 28, 17–33. Hoffman, P.F., Schrag, D.P., 2002. The snowball Earth hypothesis; testing the limits of global change. Terra Nova 14, 129–155. Hoffman, P.F., Kaufman, A.J., Halverson, G.P., Schrag, D.P., 1998. A Neoproterozoic snowball Earth. Science 281, 1342–1346. Holmes, A., 1954. Principles of Physical Geology. Thomas Nelson & Sons. 1288 pp. Hutko, A.R., Lay, T., Garnero, E.J., Revenaugh, J., 2005. Seismic detection of folded, subducted lithosphere at the core–mantle boundary. Nature 441, 333–336. Jenner, G.A., Swinden, H.S., 1993. The Pipestone Pond Complex, central Newfoundland: complex magmatism in an eastern Dunnage ophiolite. Canadian Journal of Earth Sciences 30, 434–448. Jenner, G.A., Dunning, G.R., Malpas, J., Brown, M., Brace, T., 1991. Bay of Islands and Little Port Complexes, revisited: age, geochemical and isotopic evidence confirm suprasubduction zone origin. Canadian Journal of Earth Sciences 28, 1635–1652. Karson, J., Dewey, J.F., 1978. Coastal Complex, western Newfoundland: an early Ordovician oceanic fracture zone. Geological Society of America Bulletin 89, 1037–1049. Keppie, J.D., Nance, R.D., Murphy, J.B., Dostal, J., 2003. Tethyan, Mediterranean, and Pacific analogues for the Neoproterozoic–Paleozoic birth and development of periGondwanan terranes and their transfer to Laurentia and Laurussia. Tectonophysics 365, 195–219. Knoll, A.H., 1991. End of the Proterozoic Eon. Scientific American 265 (4), 64–73. Knoll, A.H., Walter, M.R., Narbonne, G.M., Christie-Blick, N., 2004. A new period for the geologic time scale. Science 305, 621–622. Kröner, A., Pallister, J.S., Fleck, R.J., 1992. Age of initial oceanic magmatism in the Late Proterozoic Arabian Shield. Geology 20, 803–806. Kurth, M., Sassen, A., Suhr, G., Mezger, K., 1998. Precise ages and isotopic constraints for the Lewis Hills (Bay of Islands Ophiolite): preservation of an arc-spreading ridge intersection. Geology 26, 1127–1130. Küster, D., Liégeois, J.-P., 2001. Sr, Nd isotopes and geochemistry of the Bayuda Desert high grade metamorphic basement (Sudan): an early Pan-African oceanic convergent margin, not the edge of the East Saharan ghost craton. Precambrian Research 109, 1–23. Lenoir, J.-L., Küster, D., Liégeois, J.-P., Utke, A., Haider, A., Matheis, G., 1994. Origin and regional significance of late Precambrian and early Paleozoic granitoids in the PanAfrican belt of Somalia. Geologische Rundschau 83, 624–641.

419

MacLachlan, K., Dunning, G., 1998. U–Pb ages and tectonomagmatic relationships of early Ordovician low-Ti tholeiites, boninites, and related plutonic rocks in central Newfoundland. Contributions to Mineralogy and Petrology 133, 235–258. Martínez Catalán, J.R., Arenas, R., Díaz García, F., Gómez-Barreiro, J., González Cuadra, P., Abati, J., Castiñeiras, P., Fernández-Suárez, J., Sánchez Martínez, S., Andonaegui, P., González Clavijo, E., Díez Montes, A., Rubio Pascual, F.J., Valle Aguado, B., 2007. Space and time in the tectonic evolution of the northwestern Iberian Massif. Implications for the comprehension of the Variscan belt. In: HatcherJr. Jr., R.D., Carlson, M.P., McBride, J.H., Martínez Catalán, J.R. (Eds.), 4-D Framework of Continental Crust. Geological Society Memoir, vol. 200, pp. 403–423. Maruyama, S., Santosh, M., 2008. Models on snowball Earth and Cambrian explosion: a synopsis. Gondwana Research 14, 22–32. Maruyama, S., Santosh, M., Zhao, D., 2007. Superplume, supercontinent and postperovskite: mantle dynamics and anti-plate tectonics on the core–mantle boundary. Gondwana Research 11, 7–37. Matte, P., 2001. The Variscan collage and orogeny (480–290 Ma) and the tectonic definition of the Armorica microplate: a review. Terra Nova 13, 122–128. McKerrow, W.S., Scotese, C.R. (Eds.), 1990. Palaeozoic Palaeogeography and Biogeography. Geological Society of London Memoir, vol. 12. 435 pp. McMenamin, M.A., McMenamin, D.L., 1990. The Emergence of Animals; The Cambrian Breakthrough. Columbia University Press, New York, N.Y. 217 pp. Meert, J.G., Lieberman, B.S., 2008. The Neoproterozoic assembly of Gondwana and its relationship to the Ediacaran–Cambrian radiation. Gondwana Research 14, 5–21. Miller, K.G., Kominz, M.A., Browning, J.V., Wright, J.D., Mountain, G.S., Katz, M.E., Sugarman, P.J., Cramer, B.S., Christie-Blick, N., Pekar, S.F., 2005. The Phanerozoic record of global sea-level change. Science 310, 1293–1298. Moores, E.M., 1991. Southwest U.S.–East Antarctic (SWEAT) connection: a hypothesis. Geology 19, 425–428. Murphy, J.B., Nance, R.D., 1989. A model for the evolution of the Avalonian–Cadomian belt. Geology 17, 735–738. Murphy, J.B., Nance, R.D., 1991. A supercontinent model for the contrasting character of Late Proterozoic Orogenic belts. Geology 19, 469–472. Murphy, J.B., Nance, R.D., 2003. Do supercontinents introvert or extrovert?: Sm–Nd isotopic evidence. Geology 31, 873–876. Murphy, J.B., Nance, R.D., 2005. Do supercontinents turn inside-in or inside-out? International Geology Review 47, 591–619. Murphy, J.B. and Nance, R.D., 2008. The Pangea conundrum, Geology, 36, 703–706. Murphy, J.B., Pisarevsky, S.A., Nance, R.D., Keppie, J.D., 2004. Neoproterozoic–early Paleozoic configuration of peri-Gondwanan terranes: implications for Laurentia– Gondwanan connections. International Journal of Earth Sciences 93, 659–682. Murphy, J.B., Gutierrez-Alonso, G., Nance, R.D., Fernandez-Suarez, J., Keppie, J.D., Quesada, C., Strachan, R.A., Dostal, J., 2006. Origin of the Rheic Ocean: rifting along a Neoproterozoic suture? Geology 34, 325–328. Nance, R.D., Worsley, T.R., Moody, J.B., 1986. Post-Archean biogeochemical cycles and long-term episodicity in terrane processes. Geology 14, 514–518. Nance, R.D., Worsley, T.R., Moody, J.B., 1988. The supercontinent cycle. Scientific American 256 (1), 72–79. Narbonne, G.M., 1998. The Ediacara biota: a terminal Neoproterozoic experiment in the evolution of life. GSA Today 8 (2), 1–6. Narbonne, G.M., 2005. The Ediacara biota: Neoproterozoic origin of animals and their ecosystems. Annual Reviews of Earth and Planetary Sciences 33, 421–442. Nutman, A.P., Green, D.H., Cook, C.A., Styles, M.T., Holdsworth, R.E., 2001. U–Pb dating of the exhumation of the lizard peridotite and its exhumation over crustal rocks. Journal of the Geological Society, London 158, 809–820. Pallister, J.S., Stacey, J.S., Fischer, L.B., Premo, W.R., 1988. Precambrian ophiolites of Arabia: geologic settings, U–Pb geochronology, Pb isotopic characteristics and implications for continental accretion. Precambrian Research 38, 1–54. Pimentel, M.M., Fuck, R.A., 1992. Neoproterozoic crustal accretion in central Brazil. Geology 20, 375–379. Pimentel, M.M., Whitehouse, M.J., das, G., Viana, M., Fuck, R.A., Machado, N., 1997. The Mara Rosa arc in the Tocantins province: further evidence for Neoproterozoic crustal accretion in central Brazil. Precambrian Research 81, 299–310. Pin, C., Paquette, J.-L., 1999. A mantle-derived bimodal suite in the Hercynian belt: Nd isotope and trace element evidence for a subduction-related rift origin of the Late Devonian Brevenne metavolcanics, Massif Central (France). Contributions to Mineralogy and Petrology 129, 222–238. Pin, C., Paquette, J.L., Santos Zalduegui, J.F., Gil Ibarguchi, J.I., 2002. Devonian suprasubduction zone ophiolite related to incipient collisional processes in the Western Variscan Belt: The Sierra de Careón Unit, Órdenes Complex, Galicia. In: Martinez Catalan, J.R., Hatcher, R.D., Arenas, R., Diaz Garcia, F. (Eds.), Variscan–Appalachian Dynamics: The Building of the Late Paleozoic Basement. Geological Society of America Special Paper, 364, pp. 57–71. Pin, C., Paquette, J.L., Ábalos, B., Santos Zalduegui, J.F., Gil Ibarguchi, J.I., 2006. Composite origin of an early Variscan transported suture: ophiolitic units of the Morais Nappe Complex (north Portugal). Tectonics 25, TC5001. doi:10.1029/2006TC001971. Pindell, J., Kennan, L., Stanek, K.-P., Maresch, W.V., Draper, G., 2006. Foundations of Gulf of Mexico and Caribbean evolution: eight controveries resolved. In: IturraldeVinent, M.A., Lidiak, E.G. (Eds.), Caribbean Plate Tectonics. Geologica Acta, vol. 4, pp. 303–341. Pisarevsky, S.A., Wingate, M.T.D., Powell, C. McA., Johnson, S., Evans, D.A.D., 2003. Models of Rodinia assembly and fragmentation. In: Yoshida, M., Windley, B., Dasgupta, S. (Eds.), Proterozoic East Gondwana: Supercontinent Assembly and Breakup. Geological Society of London Special Publication, vol. 206, pp. 35–55. Pisarevsky, S.A., Murphy, J.B., Cawood, P.A., Collins, A.S., 2008. Late Neoproterozoic and Early Cambrian palaeogeography: models and problems. In: Pankhurst, R.J., Trouw, R.A.J., Brito Neves, B.B., de Wit, M.J. (Eds.), West Gondwana: Pre-Cenozoic

420


Correlations Across the South Atlantic Region. Geological Society, London, Special Publications, vol. 294, pp. 9–31. Powell, C. McG, 1995. Are Neoproterozoic glacial deposits preserved on the margins of Laurentia related to the fragmentation of two supercontinents? Comment, Geology 23, 1053–1054. Pysklywec, R.N., Mitrovica, J.X., Ishii, M., 2003. Mantle avalanche as a driving force for tectonic reorganization in the southwest Pacific. Earth and Planetary Science Letters 209, 29–38. Ramos, V.A., 2008. Patagonia: a Paleozoic continent adrift? Journal of South American Earth Sciences. doi:10.1016/j.jsames.2008.06.002. Rino, S., Kon, Y., Sato, W., Maruyama, S., Santosh, M., Zhao, D., 2008. The Grenvillian and Pan-African orogens: world's largest orogenies through geologic time, and their implications on the origin of superplume. Gondwana Research 14, 51–72. Rogers, J.J.W., Santosh, M., 2004. Continents and Supercontinents. Oxford University Press. Ross, G.M.,1999. Paleogeography: an Earth systems perspective. Chemical Geology 161, 5–16. Sandeman, H.A., Clark, A.H., Scott, D.J., Malpas, J.G., 2000. The Kennack gneiss of the Lizard Peninsula, Cornwall, S.W. England: commingling and mixing of mafic and felsic magmas during Givetian continental incorporation of the Lizard ophiolite. Journal of the Geological Society, London 157, 1227–1242. Sclater, J.G., Jaupart, C., Galson, D.,1980. The heat flow through oceanic and continental crust and the heat loss of the Earth. Reviews of Geophysics and Space Physics 18, 269–311. Scotese, C.R., 1997. Continental Drift, PALEOMAP Project, Arlington, Texas, 7th edition. 79 pp. Scotese, C.R., 2003. PALEOMAP Project. http://www.scotese.com/newpage9.htm. Silver, P.G., Behn, M.D., 2008. Intermittent plate tectonics. Science 319, 85–89. Stampfli, G.M., Borel, G.D., 2002. A plate tectonic model for the Paleozoic and Mesozoic constrained by dynamic plate boundaries and restored synthetic oceanic isochrones. Earth and Planetary Science Letters 196, 17–33. Stein, M., 2003. Tracing the plume material in the Arabian-Nubian Shield. Precambrian Research 123, 223–234. Stern, R.J., 1994. Arc assembly and continental collision in the Neoproterozoic East African Orogen: implications for the consolidation of Gondwanaland. Annual Reviews of Earth and Planetary Sciences 22, 319–351. Stern, R.J., 2002. Crustal evolution in the East African orogen: a neodymium isotopic perspective. Journal of African Earth Sciences 34, 109–117. Stern, R.J., 2004. Subduction initiation: spontaneous and induced. Earth and Planetary Science Letters 226, 275–292. Stern, R.J., 2008. Neoproterozoic crustal growth: the solid Earth system during a critical episode of Earth history. Gondwana Research 14, 33–50. Stern, R.J., Dawoud, A.M., 1991. Late Precambrian (740 Ma) charnockite, enderbite and granite from Moya, Sudan: a link between the Mozambique belt and the Arabian– Nubian shield. Journal of Geology 99, 648–659. Stern, R.J., Bloomer, S.H., 1992. Subduction zone infancy: examples from the Eocene Izu– Bonin–Mariana and Jurassic California. Geological Society of America Bulletin 104, 1621–1636. Stern, R.J., Kröner, A., Rashwan, A.A., 1991. A late Precambrian ~710 Ma high volcanicity rift in the southern Eastern Desert of Egypt. Geologische Rundschau 80, 155–170. Stump, E., 1992. The Ross orogen of the Transantarctic Mountains in the light of the Laurentian–Gondwana split. GSA Today, 2, (2), 25–27 and 30–33. Suhr, G., Cawood, P.A., 1993. Structural history of ophiolite obduction, Bay of Islands, Newfoundland. Geological Society of America Bulletin 105, 399–410. Suhr, G., Cawood, P.A., 2001. Southeastern Lewis Hills (Bay of Islands Ophiolite): geology of a deeply eroded, inside-corner, ridge-transform intersection. Geological Society of America Bulletin 113, 1025–1038. Sutton, J., 1963. Long-term cycles in the evolution of the continents. Nature 198, 731–735. Swinden, H.S., Jenner, G.A., Szybinski, Z.A., 1997. Magmatic and tectonic evolution of the Cambro-Ordovician Laurentian margin of Iapetus: geochemical and isotopic

constraints from the Notre Dame subzone, Newfoundland. In: Sinha, A.K., Hogan, J.P., Whalen, J.B. (Eds.), Magmatism in the Appalachians. Geological Society of America Memoir, vol. 191, pp. 337–365. Tan, E., Gurnis, M., Han, L., 2002. Slabs in the lower mantle and their modulation of plume formation. Geochemistry, Geophysics, Geosystems 3, 1067. doi:10.1067/ 2001GC000238. Tatsumi, Y., 2005. The subduction factory, how it operates in the evolving Earth. GSA Today 15 (7), 4–10. Vail, P.R., Mitchum, R.M., Thompson, S., 1977. Global cycles of relative changes in sea level. In: Payton, C.E. (Ed.), Seismic stratigraphy-application to hydrocarbon exploration. American Association of Petroleum Geologists, Memoir, vol. 26, pp. 83–97. Valentine, J.W., Moores, E.M., 1970. Plate tectonic regulation of faunal diversity and sea level. Nature 228, 657–659. van Schmus, W.R., de Brito Neves, B.B., Hackspacher, P., Babinski, M.,1995. U/Pb and Sm/Nd geochronologic studies of the eastern Borborema Province, Northeastern Brazil: initial conclusions. Journal of South American Earth Sciences 8, 267–288. van Staal, C.R., Dewey, J.F., MacNiocaill, C., McKerrow, W.S., 1998. The Cambrian–Silurian tectonic evolution of the Northern Appalachians and British Caledonides: history of a complex, west and southwest Pacific-type segment of Iapetus. Geological Society of London, Special Publication, vol. 143, pp. 199–242. van Staal, C.R., Whalen, J.B., McNicoll, V.J., Pehrsson, S., Lissenberg, C.J., Zagorevski, A., van Breemen, O., Jenner, G.A., 2007. The Notre Dame arc and the Taconic orogeny in Newfoundland. In: Hatcher Jr., R.D., Carlson, M.P., McBride, J.H., Martínez Catalán, J.R. (Eds.), 4-D Framework of Continental Crust: Geological Society of America Memoir 200, pp. 511–552. doi:10.1130/2007.1200(26). Veevers, J.J., 1989. Middle/Late Triassic (250 + 5 Ma) singularity in the stratigraphic and magmatic history of the Pangean heat anomaly. Geology 17, 784–787. Veevers, J.J., 1990. Tectonic–climatic supercycle in the billion-year plate-tectonic eon: Permian Pangean icehouse alternates with Cretaceous dispersed-continents greenhouse. Sedimentary Geology 68, 1–16. Veevers, J.J., 1994. Pangea: evolution of a supercontinent and its consequences for Earth's paleoclimate and sedimentary environments. Geological Society of America Special Paper, vol. 288, pp. 13–23. Veevers, J.J., 2004. Gondwanaland from 600.570 Ma assembly through 320 Ma merger in Pangea to 160 Ma breakup: supercontinental tectonics via stratigraphy and radiometric dating. Earth Science Reviews 68, 1–132. Veizer, J., Davin, A., Azmy, K., Bruckschen, P., Buhl, D., Bruhn, F., Carden, G., Diener, A., Ebneth, S., Godderis, Y., Jasper, T., Korte, C., Pawellek, F., Podlaha, O., Strauss, H., 1999. 87Sr/86Sr, δ13C and δ18O evolution of Phanerozoic seawater. Chemical Geology 161, 59–88. Wilson, J.T., 1966. Did the Atlantic Ocean Close and then Re-open? Nature 211, 676–681. Windley, B.F., 1977. The Evolving Continents. John Wiley and Sons, Chichester, U.K. 385 pp. Windley, B.F., 1993. Uniformitarianism today: plate tectonics is the key to the past. Journal of the Geological Society, London 150, 7–19. Windley, B.F., Whitehouse, M.J., Ba-Btatt, M.A.O., 1996. Early Precambrian gneiss terranes and Pan-African island arcs in Yemen: crustal accretion of the eastern Arabian shield. Geology 24, 131–134. Worsley, T.R., Nance, R.D., Moody, J.B., 1984. Global tectonics and eustasy for the past 2 billion years. Marine Geology 58, 373–400. Worsley, T.R., Nance, R.D., Moody, J.B., 1986. Tectonic cycles and the history of the Earth's biogeochemical and paleoceanographic record. Paleoceanography 1, 233–263. Zhong, S., Gurnis, M., 1997. Dynamic interaction between tectonics plates, subducting slabs, and the mantle. Earth Interactions 1, 1–18. Zhong, S., Zhang, N., Li, Z.-X., Roberts, J.H., 2007. Supercontinent cycles, true polar wander, and very long-wavelength mantle convection. Earth and Planetary Science Letters 261, 551–564.