Contrasting Modes of Supercontinent Formation and the Conundrum of Pangea

Abstract

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.

Full text

Contrasting modes of supercontinent formation and the conundrum of Pangea
J. Brendan Murphy
a,
⁎, R. Damian Nance
b
, Peter A. Cawood
c
a
Department of Earth Sciences, St. Francis Xavier University, Antigonish, Nova Scotia, Canada B2G 2W5
b
Department of Geological Sciences, Ohio University, Athens, Ohio 45701, USA
c
University of Western Australia, Tectonics Special Research Centre, 35 Stirling Highway, Crawley, WA 6009, Australia
abstractarticle info
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
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
87
Sr/
86
Sr 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 amalga-
mation 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 respon-
sible for the amalgamation of this supercontinent are poorly under-
stood. 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
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 des-
troyed during this cycle has different geodynamic properties depend-
ing 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;
Gondwana Research 15 (2009) 408–420
⁎Corresponding author.
E-mail address: bmurphy@stfx.ca (J.B. Murphy).
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doi:10.1016/j.gr.2008.09.005
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Dalziel, 1997), which consists of the continental fragments of
Gondwana and Laurentia, was formed by the preferential consump-
tion of the exterior oceanic lithosphere (Mozambique Ocean),
whereas Pangea wasformed by the preferential subduction of interior
oceanic lithosphere (Iapetus and Rheic oceans).
The implications of different relative ages of oceanic and conti-
nental 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 global-
scale 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 para-
digm. 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 “shield-
forming”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
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.
409J.B. Murphy et al. / Gondwana Research 15 (2009) 408–420

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 dis-
persed, only to reform again at a later time (e.g., Hoffman, 1992).
Although the details may be controversial, paleocontinental recon-
structions for the past 2.5 Ga, based on avariety of lithotectonic, paleo-
magnetic, geochemical and faunal data, imply the existence of a quasi-
regular 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), rift-
to-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 super-
continent 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
and Nance, 1991)orcollisional (Windley, 1993)orogens,whichform
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)oraccretionary (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 devel-
oped 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 frag-
ment 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-
Fig. 2. Palaeogeography at ca. 600 Ma (afterPi sarevsky et al., 2008). Rifts between Baltica
and Laurentia, and between Baltica and Amazonia are shown with solid lines; failed rift
betweenLaurentiaand Amazoniais shown with dottedlines.Subduction zonesoutboard of
Balticaand Amazoniaare directedin oppositedirections.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 terran e; 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.
410 J.B. Murphy et al. / Gondwana Research 15 (2009) 408–420

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 platetectonics 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 under-
lain 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 themwill be tectonically
unstable as the continents continue to diverge (Fig. 4). In theory, the
geodynamic distinction between the interior and exterior oceans
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 subduc-
tion 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 byintroversion. 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 byHoffman (1999).
4. Introversion and extroversion in the geologic record:
Sm/Nd evidence
The geological record suggests that both introversion and extrover-
sion 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 assem-
bly 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 inwhich one ocean is closed by introversion, and the other byextroversion.
411J.B. Murphy et al. / Gondwana Research 15 (2009) 408–420

Nance, 2003, 2005). In contrast, the latest Neoproterozoic (Ediacaran)
breakup of Pannotia, followed by the late Paleozoic assembly of Pangea,
occurredby 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 (T
R
), 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 (T
DM
,DePaolo, 1981, 1988) that predate the age of
supercontinent breakup (T
DM
NT
R
), whereas those of the interior ocean
have depleted mantle model ages that postdate breakup (T
R
NT
DM
).
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 +1ε
Nd
of the depleted mantle curve (Stern,
2002). T
DM
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 oceaniclithosphere older than the
0.76 Ga breakup age of Rodinia (i.e., T
DM
NT
R
), 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
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 Shieldsuggests that the Mozam-
bique belt was generated by collisional tectonics (Stern and Dawoud,
1991; Stern, 2002).
The transition zone between the peripheral orogens of the Arabian-
Nubian Shield and the interior orogen of the Mozambique Belt occurs in
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 (T
I
)areyounger than the time of supercontinent breakup (T
R
), whereas the
depleted mantle model ages for the Exterior Ocean (T
E
) 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).
Fig. 6. Summary diagrams of (ε
Nd
)
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 ε
Nd
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.
412 J.B. Murphy et al. / Gondwana Research 15 (2009) 408–420

Yemen, Somalia and the Sudan (e.g. Stern and Dawoud, 1991; Lenoir
et al.,1994; Windleyet 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 ε
Nd
(+3.6 to + 5.2 at t=806Ma)with
T
DM
model age s 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 ε
Nd
data ranging from +2.9 to +3.4 (t=740Ma) and
T
DM
model ages of 0.96 to 0.98 Ga. The EAO is consequently
characterized by maficcomplexeswithT
DM
ages that predate the age
of supercontinent breakup (T
DM
NT
R
), 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 ε
Nd
values forthese suites
(calculated for the age of crystallization) range from +0.2 to +6.9,
whereas T
DM
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.
ε
Nd
values for two of these complexes are high (+6.3 to +6.6 and
+4.4 to 15.8 at t=730 Ma) with T
DM
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 ε
Nd
values ranging from +1.0 to +5.0 that
yield T
DM
model ages of 0.95 to 1.20 Ga (Dostal et al., 2002).
Taken together, these data suggest that mafic terranes with T
DM
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 Gondwanan-
dervied continents (e.g. Avalonia, Carolinia), Baltica and Laurentia (van
Staal etal., 1998; Hibbardet 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 avariety 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 ε
Nd
values ranging
from +4 and +7 (Fig. 7), with T
DM
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.
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 byjuvenile MORB to OIB chemistry
with ε
Nd
values between +8 and + 11. Uncontaminated Brevenne
metavolcanics also have juvenile MORB to OIB chemistry with ε
Nd
values of +5 and + 8. The Aracena amphibolites have MORB-like
geochemistry and yield ε
Nd
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 ε
Nd
values of +6.4 to +9.2. In
all these complexes, ε
Nd
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 T
DM
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 T
DM
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 extrover-
sion, whereas Pangea formed by introversion, and that superconti-
nents 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
Fig. 7. Summary diagram of (ε
Nd
)
t
vs. time (in Ga) with compilation of Sm/Nd isotope
compositions from Iapetan and Rheic ocean complexes in the Paleozoic Appalachian-
Caledonide 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).
413J.B. Murphy et al. / Gondwana Research 15 (2009) 408–420

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 recon-
struction, is a peripheralorogen 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 paleo-
Pacific subduction zones, eventually closing the paleo-Pacific ocean to
form an extroverted supercontinent. However, paleocontinental recon-
structions from the Early Ordovician through to the Carboniferous
(Fig. 1) clearly imply that when subduction first began in the interior
oceans atabout 500 Ma, the rateof subductionwas sufficient to closethe
interioroceans, resultingin the amalgamation ofPangea by introversion
in the late Paleozoic.
This conundrum focuses attention on the mechanisms of subduc-
tion 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
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 shownin A and in B is the same. It has been
centered in B to show the continents converging towards it.
414 J.B. Murphy et al. / Gondwana Research 15 (2009) 408–420

directed away from Laurentia resulted in the formation of an ensimatic
island arc within Iapetus. Collision between this arc and a Laurentia-
derived 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 intra-
oceanic 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 litho-
sphere 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(seeFig. 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 tohave 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 environ-
ments for the generation of transform faults that separate oceanic
lithosphere of contrasting ages. The age of the paleo-Pacific oceanic
lithosphere adjacent to these transform faults will probably never be
known. However, it is almost certain to have been older than the ad-
joining Iapetus oceanic lithosphere. The age and density contrast bet-
ween the two oceanic lithospheres would have been most pronounced
when the Iapetus Ocean wasyoung. Such contrastsin the ages of oceanic
lithospheres also occur across the boundaries between interior and
exterior oceanic lithospheres (Murphy and Nance, 2008), which there-
fore 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 over-
riding 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
atransform 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 millionyears. 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 tothose 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 variousmodels have been proposed,
including those involving mantle heat input (also known as surge
tectonics), intra-plate stress transfer involving convergence, intra-plate
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.
415J.B. Murphy et al. / Gondwana Research 15 (2009) 408–420

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 mid-
ocean 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 litho-
sphere. 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 theclosure of the Rheic
Ocean was imposed by some other mechanism.
Initial
87
Sr/
86
Sr ratios provide proxy records of times of rapid ocean
spreading (low initial
87
Sr/
86
Sr) versus enhanced continental weath-
ering (high initial
87
Sr/
86
Sr). Hence, the relatively low initial
87
Sr/
86
Sr
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
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 —Eden-
Comerong-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.
416 J.B. Murphy et al. / Gondwana Research 15 (2009) 408–420

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.
(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 supercon-
tinent 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 toperovskite+magnesiowüstite) with
a negative Clapeyron slope, and slab accumulation at this discontinuity
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.
Fig. 13.
87
Sr/
86
Sr variations in the Phanerozoic sea water (see Veizer et al., 1999 for the
original dataset).
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.
417J.B. Murphy et al. / Gondwana Research 15 (2009) 408–420

depends onthe 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 accu-
mulation, rising to the surface as a superplume.
Irrespective of the causal mechanism, a superplume event would
result ina 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 representa 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 amalgama-
tion is controlled by surface plates and the location of subduction zones,
which correspond to geoid lows. These models require that super-
continents form by extroversion in which the exterior ocean surround-
ing 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 mechan-
ism. However, applying such models to the Earth's Paleozoic paleogeo-
graphy would not produce Pangea inthe 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 paleo-
Pacific 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 imposinga geoid highover 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 ResearchCouncil, 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 Geo-
science Programme, IGCP Projects 453 and 497 and to the International
Lithosphere Program, Taskforce 1; Earth Accretionary Systems.
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