Potential geodynamic relationships between the ...

Miner Petrol (2013) 107:635–650 DOI 10.1007/s00710-012-0207-9

SPECIAL ISSUE GONDWANA COLLISION

Potential geodynamic relationships between the development of peripheral orogens along the northern margin of Gondwana and the amalgamation of West Gondwana J. Brendan Murphy & Sergei Pisarevsky & R. Damian Nance

Received: 9 January 2012 / Accepted: 17 May 2012 / Published online: 13 June 2012 # Springer-Verlag 2012

Abstract The Neoproterozoic-Early Cambrian evolution of peri-Gondwanan terranes (e.g. Avalonia, Carolinia, Cadomia) along the northern (Amazonia, West Africa) margin of Gondwana provides insights into the amalgamation of West Gondwana. The main phase of tectonothermal activity occurred between ca. 640–540 Ma and produced voluminous arcrelated igneous and sedimentary successions related to subduction beneath the northern Gondwana margin. Subduction was not terminated by continental collision so that these terranes continued to face an open ocean into the Cambrian. Prior to the main phase of tectonothermal activity, Sm-Nd isotopic Editorial handling: T. Abu-Alam J. B. Murphy (*) Department of Earth Sciences, St. Francis Xavier University, P.O. Box 5000, Antigonish, NS B2G 2W5, Canada e-mail: [email protected] S. Pisarevsky School of Earth and Environment (M004), University of Western Australia, 35 Stirling Highway, Crawley, WA 6009, Australia S. Pisarevsky ARC Centre of Excellence for Core to Crust Fluid Systems (CCFS) and The Institute for Geoscience Research (TIGeR), Department of Applied Geology, Curtin University of Technology, GPO Box U1987, Perth, WA 6845, Australia S. Pisarevsky Institute of the Earth’s Crust, Siberian Branch of the Russian Academy of Sciences, 128 Lermontov Str., Irkutsk 664033, Russia R. D. Nance Department of Geological Sciences, Ohio University, 316 Clippinger Laboratories, Athens, OH 45701, USA

studies suggest that the basement of Avalonia, Carolinia and part of Cadomia was juvenile lithosphere generated between 0.8 and 1.1 Ga within the peri-Rodinian (Mirovoi) ocean. Vestiges of primitive 760–670 Ma arcs developed upon this lithosphere are preserved. Juvenile lithosphere generated between 0.8 and 1.1 Ga also underlies arcs formed in the Brazilide Ocean between the converging Congo/São Francisco and West Africa/Amazonia cratons (e.g. the Tocantins province of Brazil). Together, these juvenile arc assemblages with similar isotopic characteristics may reflect subduction in the Mirovoi and Brazilide oceans as a compensation for the ongoing breakup of Rodinia and the generation of the Paleopacific. Unlike the peri-Gondwanan terranes, however, arc magmatism in the Brazilide Ocean was terminated by continent-continent collisions and the resulting orogens became located within the interior of an amalgamated West Gondwana. Accretion of juvenile peri-Gondwanan terranes to the northern Gondwanan margin occurred in a piecemeal fashion between 650 and 600 Ma, after which subduction stepped outboard to produce the relatively mature and voluminous main arc phase along the periphery of West Gondwana. This accretionary event may be a far-field response to the breakup of Rodinia. The geodynamic relationship between the closure of the Brazilide Ocean, the collision between the Congo/São Francisco and Amazonia/ West Africa cratons, and the tectonic evolution of the periGondwanan terranes may be broadly analogous to the Mesozoic-Cenozoic closure of the Tethys Ocean, the collision between India and Asia beginning at ca. 50 Ma, and the tectonic evolution of the western Pacific Ocean.

Introduction The Late Neoproterozoic-Early Cambrian was a pivotal time in the Earth’s evolution characterized by global-scale orogeny,

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rapid continental growth, an explosion in biological activity, and dramatic climate swings (e.g. Hoffman et al. 1998; Knoll et al. 2004; Meert and Lieberman 2008; Maloof et al. 2010). Widespread subduction followed by continent-continent collisions in the Neoproterozoic-early Paleozoic led to the formation of the Pan African/Brasiliano/Mozambique orogenic belts and the amalgamation of Gondwana (Fig. 1; e.g. Hoffman 1991; Dalziel 1997; Wingate et al. 1998; Pisarevsky et al. 2003, 2008; Collins and Pisarevsky 2005; Pankhurst et al. 2008). These collisions resulted in major changes in the Earth’s paleogeography including the formation of orogenic belts within Gondwana’s interior (Murphy and Nance 1991, 2003). Recent syntheses suggest that West Gondwana (i.e. the cratons of South America and Africa) was essentially amalgamated by several continental collisions between 650 and 600 Ma (e.g. Pankhurst et al. 2008) although some orogenic

activity persisted into the Cambrian (e.g. Schmitt et al. 2008; Tohver et al. 2010; Casquet et al. 2011). East Gondwana was assembled in at least two stages, between 750 and 620 Ma (East African Orogen) and at 570–500 Ma (Kuunga orogen) (Meert 2001; Meert and Torsvik 2003; Collins and Pisarevsky 2005). For most of this time, the peri-Gondwanan terranes of eastern Laurentia and western Europe (e.g. Avalonia, Meguma, Carolinia, Ganderia, Cadomia) were located along or adjacent to the Amazonian-West African margin of West Gondwana and continued to face an open ocean into the Early Paleozoic (Fig. 1; Murphy and Nance 1989; McKerrow and Scotese 1990; Nance and Murphy 1994; Keppie et al. 1996; Dalziel 1997; Cocks and Torsvik 2002; Murphy et al. 2004; McCausland et al. 2007). The evolution of these terranes is dominated by subduction-related orogenesis and produced peripheral orogens (accretionary orogens of Cawood et al.

Fig. 1 Main tectonic elements involved in the assembly of West Gondwana (after the compilation of Klein and Moura 2008, Moura et al. 2008; see also Tohver et al. 2006) and the distribution of peri-

Gondwanan terranes along the Amazonian and West African margin (see Nance et al. 2002, 2010; Keppie et al. 2003; Linnemann et al. 2007; Waldron et al. 2009)

peri-Gondwanan relationship to West Gondwana assembly

2009), the duration of which spanned the timing of interior (collisional) orogenesis associated with the amalgamation of Gondwana (Murphy and Nance 1991). Although there have been several recent syntheses of these peri-Gondwanan terranes (e.g. Nance et al. 2008 and references therein), these focus either on the tectonic significance of the main phase of orogenic activity (i.e. the widespread 640–540 Ma development of an ensialic arc), or on their subsequent incorporation into Paleozoic orogens of the Appalachian-CaledonideVariscan belt (e.g. van Staal et al. 1998, 2009; Hibbard et al. 2007; Pollock et al. 2012). The early history of these terranes (i.e. their evolution before the onset of the main phase of arc development at ca. 640 Ma) is less well understood. This paper focuses on that early (ca. 760–650 Ma) history and attempts to provide insights into the potential geodynamic connection between the processes implicated by this history and those responsible for the assembly and amalgamation of West Gondwana. We also examine how those interactions may have resulted in the transition to the main arc activity in the peri-Gondwanan terranes. We limit our analysis to the time interval 760–620 Ma because the unresolved controversy over the paleolatitude of Laurentia (and, by implication West Gondwana) between 615 and 575 Ma (see Pisarevsky et al. 2008 for discussion) profoundly affects global reconstructions and precludes detailed analysis of potential geodynamic connections during that time interval.

Early arc evolution of the peri-Gondwanan terranes Vestiges of tectonothermal activity that pre-date the main (640– 540 Ma) phase of ensialic arc magmatism within the periGondwanan terranes are preserved in various parts of Cadomia, Avalonia and Carolinia (Fig. 2) (e.g. Doig et al. 1993; O’Brien et al. 1995, 2001; Egal et al. 1996; Murphy et al. 2000; Samson et al. 2003; Hibbard et al. 2007). They have yet to be unequivocally identified in either Meguma or Ganderia, so these two terranes are not included in the following discussion. Cadomia Cadomia is exposed as Precambrian basement blocks within the Variscan orogen of Europe and includes the Iberian massif, the Armorican massif and the Massif Central in France, and the Bohemian Massif of Germany, Poland and the Czech Republic. Only in Cadomia are peri-Gondwanan basement rocks exposed, which include the 2.1 Ga Icartian Gneiss in Armorica (Guerrot and Peucat 1990; Samson and D’Lemos 1998), and the ca. 2.1 Ga Svetlik and the 1.38 Ga Dobra gneisses in Bohemia (Kröner et al. 1988; Gebauer and Friedl 1994; Linnemann et al. 2007). There is no further record of tectonothermal activity in Cadomia until ca. 755–745 Ma— the age of the Penthièvre Complex (Egal et al. 1996) and Port

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Morvan gneisses (Samson et al. 2003) in northern Brittany— which was followed at ca. 616–540 Ma by voluminous arc magmatism (Samson and D’Lemos 1998; Nagy et al. 2002) (Fig. 2a). Samson et al. (2003) have proposed that the ca. 755– 745 Ma Port Morvan gneisses are part of a discrete terrane characterized by juvenile Sm-Nd isotopic compositions that are similar to the contemporary depleted mantle. Within this same terrane, evidence of a ca. 665–625 Ma event with similarly juvenile compositions is preserved in trondhjemite boulders in the Cesson conglomerate, and in discrete granitoid plutons. Thus it appears that part of Cadomia had an early history of arc magmatism derived from juvenile lithosphere, here termed “proto-Cadomia”. In contrast, in most of Cadomia, Sm-Nd isotopic values for 616–540 Ma crustally-derived felsic igneous rocks associated with the main phase of arc activity are dominated by an important component of ancient cratonic basement. These εNd values are predominantly highly negative (+ 1.6 to −9.9 for t0610 Ma) with depleted mantle model ages (TDM) of 1.0– 2.0 Ga (e.g. Samson and D’Lemos 1998). The wide range in εNd values is attributed to the mixing of juvenile Neoproterozoic components with magmas derived from the Icartian basement (Nance and Murphy 1996; Samson and D’Lemos 1998). The contrast of the dominant cratonic derivation of most arc magmas with the juvenile signature of the ca. 755– 745 Ma Port Morvan gneisses suggests that the latter was a discrete terrane that formed outboard of the Cadomian continental margin (Samson et al. 2003). Inglis et al. (2004, 2005) have documented an important phase of crustal thickening, deformation, metamorphism and syntectonic intrusions between 620 and 608 Ma in northern Brittany that may record the timing of that accretion. These 620–608 Ma events are not regionally extensive in Cadomia, however, in other regions, relatively continuous arc-related sections persist (Fig. 2a). Detrital zircon analyses of Neoproterozoic-Early Paleozoic clastic rocks in Cadomia yield age clusters of ca. 0.60– 0.65 Ga, 2.0–2.2 Ga, 2.4 Ga and 2.6 Ga (Fernández-Suárez et al. 2002; Gutiérrez-Alonso et al. 2005; Samson et al. 2003; Linnemann et al. 2004, 2007). Taken together, these isotopic studies are consistent with the development of the Cadomian arc on 2.0–2.1 Ga crystalline basement like that of the Icartian Gneiss. These events can be correlated with those of similar age in West Africa, which lacks evidence of tectonothermal activity between 0.65 and 2.1 Ga, such that Cadomia is widely viewed to have been attached to West Africa in the late Neoproterozoic (e.g. Quesada et al. 1991; Nance and Murphy 1994, 1996; Samson and D’Lemos 1998; Nance et al. 1991, 2002; Keppie et al. 2003). Avalonia The 640–540 Ma evolution of Avalonia is so similar to that of Cadomia that these rocks are commonly interpreted to

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a

b AVALONIA ca. 380 Ma RED BEDS

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ICARTIAN GNEISS ca. 2.1 Ga

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Fig. 2 Interpretive late Neoproterozoic-Palaeozoic tectonostratigraphic column for a Cadomia, b Avalonia, c Carolinia (modified from Nance et al. 2008)

have been contiguous at that time and are collectively referred to as the Avalonian-Cadomian belt (e.g. Murphy and Nance 1989). However, vestiges of earlier tectonothermal activity ranging from ca. 765 to 660 Ma are more widely preserved in Avalonia (Fig. 2b). This phase of activity predates the main phase of Avalonian ensialic arc magmatism and the rocks formed at this time are collectively referred to as the products of “early arc” magmas. The crust upon which this arc developed has been termed “proto-Avalonia” (Murphy and Nance 2002; Murphy et al. 2000, 2008). The oldest known example of these “early arc” rocks occurs on the Burin Peninsula of Newfoundland, where the Burin Group (Strong et al. 1978; O’Driscoll et al. 2001) consists of a greenschist-grade submarine mafic volcano-plutonic succession that has been dated at 765–760 Ma (U-Pb, zircon, Krogh et al. 1988; Murphy et al. 2008). The geochemistry of this succession suggests an juvenile arc-back arc environment (Murphy et al. 2008). Other examples of early

Avalonian magmatism in Newfoundland include: (a) the ca 730 Ma felsic flows and porphyries in the core of the Holyrood Horst near Conception Bay (Israel 1998; O’Brien et al. 2001), (b) the calc-alkalic volcanic and plutonic rocks of the 685–670 Ma Tickle Point Formation and Furbys Cove Intrusive Suite north of Fortune Bay (Swinden and Hunt 1991; O’Brien et al. 1992, 1995), and possibly the ca. 680 Ma Cinq Cerf-Grey River gneiss terrane of the Hermitage Flexure (Dunning and O’Brien 1989; O’Brien et al. 1990, 1992; 1995; 1996; Valverde-Vaquero et al. 2006). In Nova Scotia, igneous rocks related to the development of these early arc complexes are preserved in the ca. 734 Ma calc-alkalic Economy River Gneiss (Doig et al. 1993) in the Cobequid Highlands, and in the 700–630 Ma back-arc volcanic deposits in the Creignish Hills, (Keppie and Dostal 1998) and the ca. 676 Ma Stirling Belt (Bevier et al. 1993) on Cape Breton Island. In addition, the sedimentary Gamble Brook Formation (also in the Cobequids) has a depositional age


c CAROLINIA .

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constrained between 1.0 and 0.6 Ga and is dominated by metamorphosed siliciclastic deposits whose geochemistry indicates deposition in a back-arc basin (Murphy 2002). Similar evidence for an early arc-related event is found in correlative Avalonian rocks in Britain. These include the ca. 711 Ma Stanner-Hanter Complex in the Welsh Borderlands (Schofield et al. 2010) and the ca. 677 Ma calc-alkalic Malverns Plutonic Complex of the British Midlands (Tucker and Pharoah 1991; Pharaoh and Gibbons 1994; Strachan et al. 1996). The Burin Group is the only suite of these Avalonian early arc rocks that is overwhelmingly dominated by mafic magmatism. It is predominantly juvenile in composition, as exemplified by εNd values (t0760) of +4.5 to +7.6, which are close to contemporary depleted mantle values, and by depleted mantle (TDM) model ages between 0.76 and 1.0 Ga (Murphy et al. 2008). Of the remaining early arc igneous rocks, Sm-Nd data are available only for the Economy River Gneiss, and the Malverns Plutonic and Stanner-Hanter complexes. These complexes are foliated, intermediate to felsic in composition, and have calc-alkalic geochemical signatures. The age of the

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foliation is not known for the Economy River Gneiss. For the Malverns Plutonic Complex, 40Ar/39Ar (hornblende) mineral ages of ca. 650 Ma are interpreted to date cooling following amphibolite facies metamorphism (Strachan et al. 1996). For the Economy River Gneiss, εNd values (t0734 Ma) range from +1.3 to +4.1, with TDM model ages ranging from 1.0 Ga Ma to 1.2 Ga. For the Malverns Plutonic Complex, εNd values (t0677 Ma) are less radiogenic, ranging from −0.1 to −2.0, but the TDM model ages are similar (1.0 Ga to 1.1 Ga, Murphy et al. 2000). εNd values for the StannerHanter Complex (−0.3 to −1.2 at t0711 Ma) are similar to those of the Malverns, but the TDM model ages of these rocks (1.4–1.5 Ga) suggest a greater involvement of an older crustal component (Schofield et al. 2010). With the exception of the crustally contaminated samples of the Stanner-Hanter Complex, the TDM model ages of the “early arc” magmas compare favourably with those of felsic rocks produced during the main phase of Avalonian arc magmatism, suggesting that felsic magmas ranging from 0.76 Ga to at least 0.54 Ga were derived from a similar juvenile crustal source (e.g. Murphy and Dostal 2007; Murphy et al. 2008). Whereas Sm-Nd isotopic analyses from igneous rocks provide data on the composition of crustal sources, U-Pb analyses of detrital zircons constrain the age of basement rocks of contiguous land masses. The Gamble Brook Formation in mainland Nova Scotia consists of greenschist facies interbedded quartzites, psammites and pelites, and is widely viewed as a platformal deposit. The basement, which must have underlain this sequence at the time of its deposition, has not been identified. The youngest group of detrital zircon ages from the quartzites cluster at ca. 1.0 Ga. The analyzed zircons yielded Mesoproterozoic, Paleoproterozoic and Neoarchean ages (Keppie et al. 1998; Barr et al. 2003). Although the total number of analyses (22 zircons in two studies) is insufficient for a rigorous provenance analysis, these data suggest proximity to Amazonian basement. SmNd isotopic studies (Murphy 2002) indicate that the quartzites were derived predominantly from ancient cratonic basement. At 650 Ma, the quartzite samples have negative εNd values (−3.05 and −8.36) and high TDM model ages (1,542 and 1,859 Ma). On the other hand, the pelites (εNd 0+5.9 to −5.9; TDM 01.1 Ga to 1.7 Ga) and psammites (ε Nd 0+7.71 to −1.81; T DM 00.9 to 1.5 Ga) have very wide ranges consistent with a mixed (cratonic-juvenile) source. The Gamble Brook data are collectively interpreted to reflect deposition within a rifted arc setting adjacent to a cratonic basement. Carolinia Carolinia (Fig. 2c) is characterized by Late Neoproterozoic– early Paleozoic arc-related igneous and sedimentary assemblages (Hibbard et al. 2002, 2007). The oldest known

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rocks are those of the Roanoke Rapids complex in North Carolina, which is a juvenile low-grade volcanicsedimentary sequence intruded by ca. 672 Ma granodiorite (Coler et al. 2000). This complex is in fault contact with 630–610 Ma juvenile arc rocks (Virgilina sequence, Samson et al. 1995; Wortman et al. 2000) and a voluminous, 560– 532 Ma volcanic and sedimentary assemblage (Albemarle sequence, Hibbard et al. 2002). Volcanic and volcaniclastic rocks of the Virgilina sequence show positive εNd values (+2 to +7 at t=600 Ma) and TDM model ages of 0.8 to 1.1 Ga and contain zircons whose cores yield ages of ca. 0.97–1.23 (Samson et al. 1995; Wortman et al. 2000). U-Pb data from detrital zircons in clastic rocks yield clusters at 0.6–0.65 Ga, 0.8–1.3 Ga, 1.5–1.8 Ga and 2.2–2.6 Ga (Pollock et al. 2009). Taken together, these data indicate that Carolinia developed in proximity to Early Neoproterozoic to Mesoproterozoic basement, here termed “proto-Carolinia”.

Assembly of West Gondwana Although there are a plethora of models for the amalgamation of West Gondwana, there is general consensus that amalgamation involved the collision of a number of continental blocks between 650 Ma and 600 Ma, which resulted in the Pan African/Brasiliano orogens (Fig. 1; Meert et al. 1995; Cordani et al. 2003; Tohver et al. 2006; Collins and Pisarevsky 2005; Cawood and Buchan 2007; Klein and Moura 2008; Meert and Lieberman 2008; Pisarevsky et al. 2008). Although many of these fold belts have been investigated in detail, isotopic studies monitoring their evolution are rather uneven and will be included in the discussion below where available. Collisions between the Amazonian-West African, Congo/ São Francisco and Paranapanema blocks began at ca. 650 Ma and continued until 600 Ma (Pimentel et al. 1991, 2000; Valeriano et al. 2004, 2008). The BorboremaDahomide belt, located adjacent to the TransbrasilianoKundi lineament (Fig. 1), is an expression of the 650– 600 Ma collision of the northern margin of the Congo/São Francisco craton with the West African craton (e.g. Arthaud et al. 2008; Avigad et al. 2012). The 0.64–0.61 Ga collision between the Amazonian and Congo/São Francisco cratons produced several fold belts in the Tocantins region, including the Araguaia and Paraguay belts to the north and west (e.g. Moura et al. 2008). Before ca. 650 Ma, the relationship between the Congo/ São Francisco and Amazonian cratons is controversial (De Waele et al. 2008). Resolution of this controversy is important in evaluating potential geodynamic connections with the evolution of the peri-Gondwanan terranes. The Li et al. (2008) consensus reconstructions of the 1.1–0.6 Ga time interval show the Congo/São Francisco craton positioned adjacent to Amazonia as part of the supercontinent Rodinia

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until 0.75 Ga, when it is shown to have separated from Amazonia, only to re-unite with Amazonia by continental collision at ca. 0.63 Ga (Fig. 3). If so, the collisional belt would represent a simple, virtually symmetric, Wilson-cycle orogen. In contrast, Kröner and Cordani (2003) and Collins and Pisarevsky (2005) claim that the Congo/São Francisco craton was not part of Rodinia at ca. 1.0 Ga. Instead, they maintain that a wide ocean existed between the Congo/São Francisco craton and Amazonia at that time, which narrowed throughout most of the Neoproterozoic, culminating in 0.65 Ga collision. The geological record preserved in the mobile belts between the Amazonia and Congo/São Francisco cratons supports the latter scenario. In the central portion of the Tocantins region, the Goias massif records the convergence between these two cratons and is dominated by two ages of arc activity represented by: (a) 0.9–0.85 Ga arc-related metamorphosed volcanic and plutonic rocks that evolved in an oceanic island setting, and (b) ca. 0.76–0.60 Ga intermediate to felsic volcanic assemblages (Pimentel and Fuck 1992; Pimentel et al. 1997, 2000; Laux et al. 2005). Sm-Nd isotopic data for both assemblages indicate derivation from juvenile lithosphere. Initial εNd values (calculated for the age of crystallization) range from +0.2 to +6.9, whereas TDM model ages fall between 0.9 and 1.2 Ga (Pimentel et al. 1999). Pimentel et al. (1999) propose that subduction was directed away from Amazonia and towards the Congo/ São Francisco craton for more than 100 Ma. If so, the longevity of subduction without arc-continent collision implies that either the arc lay at a considerable distance from Amazonia (i.e. the oceanic tract outboard of Amazonia must have been extensive), or an intra-oceanic plate boundary was situated between Amazonia and the arc. The Li et al. (2008) reconstruction, however, shows Amazonia and the Congo/São Francisco craton as being contiguous during the earliest phase of arc magmatism, implying that the oceanic lithosphere subducted prior to the ca. 650 Ma collision between Amazonia and the Congo/São Francisco craton was younger than 750 Ma. However, TDM model ages for 750–650 Ma arc rocks in the Goias massif igneous rocks uniformly indicate formation by recycling of 0.9–1.1 Ga oceanic lithosphere. Paleomagnetic data (Meert et al. 1995; Wingate et al. 2010) suggest that between ∼800 and 750 Ma, the Congo/ São Francisco craton remained near the equator, but rotated by about 90°. Li et al. (2004) suggested that this rotation was a part of the rotation of the whole of Rodinia during an episode of inertial interchange true polar wander (IITPW, Kirschvink et al. 1997), and the same model has been applied in the Rodinia reconstructions of Li et al. (2008). The Li et al. (2004) model is supported by paleomagnetic data from South China (Evans et al. 2000) and India (Radhakrishna and Mathew 1996), but is undermined by new

peri-Gondwanan relationship to West Gondwana assembly Fig. 3 Continental reconstructions for the 1.1–0.6 Ga time interval (after Li et al. 2008)

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geochronology from the Harohalli dykes in the Dharwar craton of southern India (Pradhan et al. 2008). There is also no paleomagnetic support for the ∼800 Ma IITPW episode from other major continental blocks such as Laurentia. Collectively in our view, the Congo/São Francisco craton is unlikely to have been part of Rodinia, and a wide ocean probably separated it from Amazonia in the early Neoproterozoic. We have therefore modified the Li et al. (2008) reconstruction accordingly. In the reconstruction used for this analysis (Fig. 4), the evolution of the Goias massif and Rio Negro arcs are compatible with ongoing convergence between the Congo/São Francisco and Amazonian cratons during the Neoproterozoic (e.g. Heilbron et al. 2008; Valeriano et al. 2008). Although less well exposed, similar arguments can be made along the eastern margin of the West African craton where the Hoggar suture zone, adjacent to the KundiTransbrasiliano lineament, records 800–690 Ma oceanic island arc magmatism followed by ophiolite emplacement and collisional orogenesis from 690 Ma to 605 Ma (e.g. Caby 2003; Henry et al. 2009; Avigad et al. 2012). The terranes within this suture zone share similar characteristics to those of the Tocantins province of Brazil. For example, in Fig. 4 Continental reconstructions (a–d) for the 750–600 Ma time interval, modified from Li et al. (2008) to show an oceanic tract between Congo/São Francisco and Amazonia/West Africa. A-A Afif-Abas; Am Amazonia; Au Australia; Av Avalonia; Az Azania; Ba Baltica; C Cadomia; Co Congo; G Ganderia; In India; K Kalahari; La Laurentia; M Mawson; O Oaxaquia; P Paraná; pG peri-Gondwanan terranes (Cadomia, Avalonia, Ganderia); Pm Pampean; Rp Rio de La Plata; SF São Francisco; Si Siberia. The shaded areas in a show (very schematically) the inferred locations of juvenile (0.8–1.1 Ga) lithosphere, which may have been the sources for some of the periGondwanan terranes (pG; Avalonia, Carolinia, part of Cadomia), Arabian-Nubian shield (A) and the Goias arc (G)

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Mali, ca. 720 Ma juvenile arcs (Tilemsi belt) comprise mafic to intermediate volcanic rocks, interbedded metagreywackes and coeval plutonic rocks and are inferred to have formed above an east-dipping subduction zone (Caby 2003). Available Sm-Nd isotopic data from the metagreywackes and the plutonic rocks are characterized by high εNd values (+4.4 to +5.8 and +6.3 to +6.6 at t0730 Ma) with model ages of 840 Ma to 940 Ma and 710 Ma to 760 Ma respectively (Dostal et al. 1994). These values are comparable with those from the Goias massif, and indicate that the plutonic rocks were derived from depleted mantle. Similar values for the metagreywackes imply predominant derivation from juvenile crust, and that any contribution from continental crust was minor. To the north, inliers of mafic rocks include shoshonites with arc-related geochemical signatures and high εNd values (+1.0 to +5.0) that yield model ages of 950 Ma to 1.20 Ga (Dostal et al. 2002). Although there are some important examples of orogeny continuing into the Cambrian (e.g. Schmitt et al. 2008; Tohver et al. 2010, 2012), collectively, the reconstructions indicate that most of West Gondwana was essentially amalgamated between 650 and 600 Ma. The extent of Cambrian orogenic


activity indicates that telescoping of more restricted oceanic tracts produced younger orogenic events that are not recognized in the paleomagnetic data (e.g. Thompson et al. 2007). The overall setting may be analogous to the Cenozoic evolution of the eastern Mediterranean region (e.g. Schmitt et al. 2008; Tohver et al. 2010). These reconstructions also show, however, significant separation between East and West Gondwana at ca. 600 Ma, that continued into the Cambrian, with final amalgamation of East and West Gondwana by ca. 530 Ma (Collins and Pisarevsky 2005; Cawood and Buchan 2007; Meert and Lieberman 2008; Li et al. 2008; Pisarevsky et al. 2008). The Late Neoproterozoic orogenic belts of West Gondwana are examined in the paleogeographic context implied by these reconstructions. Brasília Belt Along the southeastern margin of the Congo/São Francisco craton, the Brasília Belt was formed by the craton’s collision with continental blocks to the west, such as the Paranapanema block (e.g. Valeriano et al. 2008; Reno et al. 2009), which is largely hidden beneath the Paleozoic and Mesozoic rocks of the Parana basin (Mantovani and Brito Neves 2005). These collisions resulted from the 650–600 Ma closure of the Brazilides Ocean (e.g. Unrug 1996; Campos Neto and Caby 1999, 2000; Reno et al. 2009). The Brasília Belt comprises metasedimentary successions derived from the São Francisco craton that contain an increasing juvenile component up-section and are thought to record the destruction of an ocean basin that flanked the São Francisco craton (e.g. Valeriano et al. 2008). These successions were deformed between 0.8 Ga and 0.6 Ga, and are characterized by thrusting towards the São Francisco craton, which at that time was located on the lower plate (e.g. Campos Neto et al. 2010). U-Pb (zircon) geochronological studies indicate that peak metamorphic conditions occurred at 678±29 Ma, although high-grade metamorphism producing granulite continued to ca. 620 Ma (Reno et al. 2009). Ribeira and São Gabriel belts The Ribeira Belt (Campos Neto 2000; Trouw et al. 2000; Heilbron et al. 2004) occurs to the east of the São Francisco craton and consists of a collage of terranes that accreted to the craton at various times between 620 and 530 Ma. According to Heilbron et al. (2004), collision with the Angola craton began at about 620 Ma and, by ca. 530 Ma, the Kalahari craton had collided. The Ribeira Belt truncates the Brasília Belt at a high angle. Although the original relationship between these belts is controversial, ca. 590– 560 Ma deformation documented in the Ribeira Belt apparently overprints earlier structures in the Brasília Belt implying a linked history by that time. The Ribeira Belt is divided

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into several terranes. One of these, the Oriental terrane, contains belts of ca. 790–620 Ma juvenile arc rocks (Serra du Bolivia and Rio Negro arcs) that are interpreted to reflect intra-oceanic to peri-oceanic settings (Heilbron and Machado 2003; Heilbron et al. 2008). The rocks in the Rio Negro arc are predominantly calc-alkalic plutons with shoshonitic affinities. Sm-Nd data (Tupinambá et al. 2012) indicate very wide ranges in εNd (+3.0 to −14.5) and TDM model ages (0.9 to 2.1 Ga with peaks at 1.1, 1.4 and 1.7 Ga). The younger model ages are thought to reflect derivation from juvenile mantle, whereas the older ages reflect contamination from older crustal and mantle sources. The São Gabriel belt is exposed in southern Brazil and records the ca. 735–720 Ma development of juvenile terranes that were accreted to the Rio de la Plata craton by 680 Ma (Hartmann et al. 2010). This was followed by the formation of the Dom Feliciano belt at 650–550 Ma, which represents either the 650–550 Ma collision between the Rio de la Plata and Kalahari cratons (Hartmann et al. 2000) or piecemeal accretion of outboard terranes (Bossi and Gaucher 2004).

Discussion There is a growing consensus (e.g. Li et al. 2008) that the late Mesoproterozoic and most of the Neoproterozoic were dominated by the assembly, amalgamation and breakup of the supercontinent Rodinia. Rodinia amalgamated between 1.1 and 0.9 Ga, and began rifting apart at ca. 0.8 Ga. This rifting led to its break up, which began at ca. 0.75 Ga and continued until the 0.63–0.55 Ga separation of Laurentia, Baltica and Amazonia to form the Iapetus Ocean (Dalziel 1997; Cawood et al. 2001; Pisarevsky et al. 2008). The potential geodynamic connection between the evolution of the peri-Gondwanan terranes and the amalgamation of West Gondwana is evaluated below in this global context. Late Mesoproterozoic-Early Neoproterozoic igneous rocks formed in the 1.1 to 0.9 Ga collisional belts as Rodinia amalgamated (e.g. the Grenville orogen in eastern North America) have TDM model ages that are considerably older than the age of the orogens themselves (typically >1.5 Ga) because they were produced predominantly by recycling of older crust (e.g. Dickin et al. 1990; Daly and McLelland 1991; Dickin 2000). In contrast, igneous rocks in both the peri-Gondwanan terranes and the orogenic belts formed during the amalgamation of West Gondwana have TDM model ages that are coeval with the time of Rodinia amalgamation. The 1.0 to 1.2 Ga tectonothermal event in Avalonia and Carolinia implied by the Sm-Nd data is interpreted as the age of formation of the juvenile lithosphere (protoAvalonia, proto-Cadomian, proto-Carolinia) from which the early (760–650 Ma) arcs were derived. The oldest arcs

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known in the Tocantins province (Mara Rosa, Porangatu, Arenopolis; Pimentel et al. 1999) formed soon after Rodinia amalgamation and their juvenile composition is indicated by the narrow time interval between their crystallization and TDM model ages. Since these model ages reflect the time of formation of juvenile lithosphere, the proto-Avalonian, proto-Cadomian, proto-Carolinian and Tocantins crust must have formed somewhere within the Panthalassa-like Mirovoi Ocean that surrounded Rodinia during the Neoproterozoic (McMenamin and McMenamin 1990). The beginning of Rodinia breakup at 0.75 Ga (Wingate and Giddins 2000; Harlan et al. 2003; Li et al. 2008; Pisarevsky et al. 2003) is broadly coincident with the onset of juvenile arc magmatism in Avalonia (Burin Group, Murphy et al. 2008), Cadomia (Port Morvan gneisses, Samson et al. 2003) and possibly Carolinia (Hibbard et al. 2007; Pollock et al. 2012), as well as renewed arc magmatism in the oceanic lithosphere between the converging continental blocks of West Gondwana (Pimentel et al. 1999). Indeed, the isotopic similarities between the Tocantins Province of central Brazil (Pimentel and Fuck 1992; Pimentel et al. 1999) and both Avalonia and Carolinia have been noted (Nance and Murphy 1994, 1996; Murphy et al. 2000; Pollock et al. 2012) and suggest their derivation from oceanic lithosphere of similar age and composition. The same inference can be extended to the juvenile Port Morvan gneisses of Cadomia. In each case, the Sm-Nd data imply that the arcs were formed from recycled 0.9–1.1 Ga juvenile lithosphere between converging continental blocks in the Mirovoi Ocean. More generally, these events are coeval with other juvenile arc assemblages with similar Sm-Nd isotopic characteristics such as those of the Arabian Shield (Stern 2002; Johnson et al. 2011) and the Bou Azzer and coeval ophiolites of Morocco (Thomas et al. 2002; D’Lemos et al. 2004; Samson et al. 2004), suggesting that Rodinia breakup and the opening of the Paleopacific Ocean were compensated by widespread intra-oceanic subduction in the Mirovoi Ocean, which resulted in the development of a series of island arcs built upon the 0.9 to 1.1 Ga juvenile lithosphere that was prevalent in the Mirovoi Ocean. This situation is broadly analogous with the Mesozoic breakup of Pangea and the opening of the Atlantic Ocean, which resulted in compensatory subduction in Panthalassa. The main phase of magmatism in Avalonia is continental in nature, implying that Avalonia had accreted to the northern margin of West Gondwana by ca. 650 Ma. The ca. 650 Ma high-grade metamorphic events recorded in parts of Avalonia (e.g. Strachan et al. 1996, 2007) are interpreted to reflect this accretionary event. Detrital zircon studies (e.g. Keppie et al. 1998; Barr et al. 2003) suggest that Avalonia collided with the Amazonian margin of West Gondwana. The ca. 620–608 Ma documentation of crustal thickening, deformation, metamorphism and syntectonic intrusions in

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Cadomia (Inglis et al. 2004, 2005) may represent accretion of the outboard juvenile terrane documented by Samson et al. (2003). However, the 630–610 Ma juvenile arc assemblage in Carolinia (Hibbard et al. 2007) suggests that Carolinia still lay outboard of the Gondwanan margin at that time. The timing of the accretion of Carolinia to Gondwana is less well constrained since continental magmatism does not begin until ca. 560 Ma (Hibbard et al. 2002). But it presumably occurred after the accretion of Avalonia and probably Cadomia. These accretionary events may also be a far-field response to Rodinia breakup, analogous with the Mesozoic-Cenozoic accretion of outboard terranes to the western margin of North America as the Atlantic Ocean widened and the Pacific Ocean contracted following Pangea breakup. Hence terrane accretion along the northern Gondwanan margin may have occurred in piecemeal fashion, again broadly analogous with accretionary events in the North American Cordillera during the Mesozoic-Cenozoic. Following the accretion of Avalonia and Cadomia, the main phase of arc magmatism began when renewed subduction occurred along the northern Gondwanan margin outboard of the accreted terranes. Subduction resulted in the recycling of this juvenile crust yielding igneous complexes with similar TDM model ages to their predecessors (Murphy et al. 1996, 2008). Hence, accretion and the onset of the main phase of Avalonian-Cadomian arc magmatism are broadly synchronous with the onset of the continental collisions that resulted in the amalgamation of West Gondwana. There are several striking features of the reconstructions (Fig. 4) that provide insights into the potential geodynamic relationship between the peri-Gondwanan terranes and the amalgamation of West Gondwana. First, LaurentiaAmazonia-West Africa undergoes clockwise rotation between 720 and 650 Ma, but after collision with the Congo/ São Francisco craton this rotation is reversed, beginning at ca. 620 Ma. The clockwise rotation is related to two poles from Laurentia, the 723 Ma Franklin dykes and Natkusiak Formation poles (Heaman et al. 1992; Palmer et al. 1983; Park 1994; Denyszyn et al. 2009) and the 615 Ma Long Range pole (Murthy et al. 1992; Kamo and Gower 1994; Hodych et al. 2004), the clockwise rotation being the shortest and smoothest movement between the two. Since Avalonia and Cadomia had collided with the Gondwanan margin by 650 and 620 Ma, respectively, the oceanic tract between Gondwana and these terranes must have been subducted by that time. The record of arc activity coeval with this subduction occurs in the peri-Gondwanan terranes but is not recorded in cratonic West Africa-Amazonia. Hence, the polarity of subduction must have been directed oceanward, and the rate of northward motion of the arc complex must have exceeded that of Amazonia-West Africa. The amalgamation of West Gondwana would have led to the progressive elimination of subduction between several


continental blocks beginning at ca. 650 Ma and continuing until ca. 530 Ma. Murphy and Nance (1991) argued that this amalgamation would have been compensated by renewed subduction and voluminous ensialic arc magmatism along the Amazonian (Avalonia) and West African (Cadomia) periphery of Gondwana (Fig. 1). Indeed, subduction zones may have virtually encircled West Gondwana at this time, with the active margin of the Terra Australis orogen lying along the margin of Amazonia (Cawood and Buchan 2007) and subduction occurring along its eastern flank that may have ultimately led to the collision between East and West Gondwana in the late Ediacaran and early Cambrian (e.g. Collins and Pisarevsky 2005). In addition, the change from clockwise to counter-clockwise rotation of the LaurentiaAmazonia-West-Africa-Congo/São Francisco block beginning at ca. 620 Ma would explain the Andean-style of peripheral (or advancing accretionary, Cawood et al. 2009) orogenic activity. In the reconstruction, we apply the paleolatitude (38oS at 597 Ma) derived by Thompson et al. (2007) in the New England part of Avalonia, which they argue constrains the position of West Gondwana at that time (see also McNamara et al. 2001; Pisarevsky et al. 2012). In order not to violate the conclusion of Cawood et al. (2001) about the opening of the western segment of the Iapetus Ocean between Laurentia and Amazonia after 550 Ma, we move Avalonia further north, noting that the position of Baltica between 615 and 555 Ma is unconstrained. Voluminous ensialic magmatism in the peri-Gondwanan terranes ended diachronously between ca. 600 to 540 Ma. With reference to the Cordilleran analogy once again, diachronous termination of magmatism has been interpreted to reflect ridge-trench collision (e.g. Kusky et al. 1997) and the generation of a San Andreas-style intra-continental transform regime (Murphy and Nance 1989; Murphy et al. 1999; Nance et al. 2002; Keppie et al. 2003). If so, the cessation of magmatism may reflect the continuing breakup of Rodinia as Gondwana assembled, as evidence by the coeval generation of the Iapetus Ocean (Cawood et al. 2007). The geodynamic relationship between the closure of the Brazilide Ocean, the collision between the Congo/São Francisco and Amazonia/West Africa cratons, and the tectonic evolution of the peri-Gondwanan terranes may be broadly analogous to the Mesozoic-Cenozoic closure of the Tethys Ocean, the collision between India and Asia beginning at ca. 50 Ma (Searle et al. 2011), and the tectonic evolution of the western Pacific Ocean. Late Mesozoic-Cenozoic reconstructions (e.g. Scotese 2001) show the lithospheres beneath the Tethyan and Pacific oceans to have been juxtaposed along a northward dipping subduction zone. At 50–47 Ma, coincident with the closure of Tethys, there was a major change in the direction of motion of the Pacific plate (Sharp and Clague 2006; Tarduno et al. 2008) and along the boundary between the

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Australian and Antarctic plates (Whittaker et al. 2007). We may never know Neoproterozoic plate geometries in the same detail, but in very general terms, the closure of the Brazilide Ocean would have severed the plate boundary connection between the Brazilide and Mirovoi oceanic lithospheres, and may have been accompanied by changes in direction of motion of adjacent plates, including the plates upon which Avalonian lithosphere resided.

Conclusions The available data indicate important geodynamic connections between the breakup of Rodinia, the amalgamation of West Gondwana and the evolution of the peripheral orogens along the margins of West Gondwana. The breakup of Rodinia, beginning at ca. 0.75 Ga, is coeval with the onset of juvenile arc magmatism occurring in the peri-Gondwanan terranes as well as with renewed subduction that ultimately led to the amalgamation of West Gondwana. Progressive elimination of subduction zones between the converging blocks of West Gondwana was compensated by renewed arc magmatism along its margins. The proposed relationship may be analogous to the Mesozoic-Cenozoic closure of the Tethys Ocean and the evolution of the western Pacific. Acknowledgements We thank Peter Cawood and Tim Kusky for constructive reviews. JBM acknowledges the continuing support of Natural Sciences and Engineering Research Council, Canada. Paleogeographic figures have been plotted with the aid of GPLATES opensource software (http://www.gplates.org/). RDN acknowledges NSF grant EAR-0308105. This is contribution 186 from the ARC Centre of Excellence for Core to Crust Fluid Systems, and TIGeR publication 420. Contribution to IGCP 597.

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