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From ocean subduction to ocean island

Philip Heron and colleagues discuss the links between supercontinents, subduction and mantle dynamics

Words by Philip J. Heron
28 February 2024
An illustration of how Earth may have looked during the late Paleozoic and early Mesozoic, when the continents assembled to form the supercontinent Pangaea

An illustration of how Earth may have looked during the late Paleozoic and early Mesozoic, when the continents assembled to form the supercontinent Pangaea

Our planet has experienced several supercontinent configurations during its history. In tandem with supercontinent formation, the geological archive also preserves a record of mass outpourings of lava and magmatic emplacement – both within supercontinental interiors and exterior to the landmass, in neighbouring ocean basins. These large igneous provinces, or LIPs, are potentially linked to thermal plumes upwelling from the deep mantle. Thermal insulation of the mantle by an overlying blanket of thick supercontinental lithosphere provides an intuitive mechanism for the generation of hot plumes beneath continental interiors, but a mechanism for generating thermal plumes beneath thin oceanic lithosphere has been elusive. However, our recent numerical modelling efforts (Heron et al., 2024) now show that upwelling mantle, stirred by the subduction of oceanic lithosphere deep into the mantle, could be responsible for the formation of both intra-supercontinental and oceanic LIPs.


A graphic showing four stages of the Pangaean supercontinent assembly and break up.

Figure 1: Pangaean supercontinent assembly and break up. The continents assembled to form Pangaea about 320 Ma (A,B) and broke up from about 175 Ma onwards (C) to form the present-day continental configuration (D). (Graphic by Fabio Crameri, published open access (CC BY-SA 4.0) and made available via the open-access s-Ink repository at s-ink.org/paleotopography, based on the palaeo-digital elevation models and tectonic plate reconstructions from Scotese & Wright (2018) PALEOMAP Paleodigital Elevation Models (PaleoDEMS) for the Phanerozoic [Data set]. Zenodo, doi.org/10.5281/zenodo.5460860)


When we think of a supercontinent, most of us conjure an image of all or most of Earth’s continental pieces coming together – a vast landmass surrounded by wide oceans. What often surprises people is that there is no formal definition of what constitutes a supercontinent – is it 100% of all land? Less? If less, how much less?

Pangaea is believed to be the youngest of a number of supercontinents that assembled and dispersed over the course of Earth’s history. The supercontinent formed during the Carboniferous around 320 million years ago (Ma) with its break-up from the Jurassic-Cretaceous (~175 Ma) onwards (Fig. 1). However, Pangaea didn’t catch all the available landmass, with a few stragglers not completely contiguous.

These supercontinents, which varied in size, are defined more by their impact on our world than their continental extent

Although their ages of amalgamation and break-up are still being refined, a number of pre-Pangaean supercontinents have been proposed. Rodinia was formed by 1,100–900 Ma and is believed to have rifted apart in two separate episodes (c. 850–700 Ma and c. 620–540 Ma; Li et al., 2008, 2013; Li & Evans, 2011). Earlier supercontinents include: Nuna (also known as Columbia) whose existence is most recently dated at c. 1,600–1,400 Ma (e.g., Pehrsson et al., 2015); Kenorland thought to have existed during the interval c. 2,700–2,500 Ma (Williams et al., 1991); and possibly Ur (c. 3,000 Ma), although the latter is better described as a supercraton (i.e., transient, late-Archean landmasses that broke up to form cratons; Bleeker, 2003).

These supercontinents, which varied in size, are defined more by their impact on our world than their continental extent. Each supercontinent has left an enduring trace in the geological record – ‘geomarkers’ that record some abrupt change as a result of the congregation of continental material. These geomarkers include global-scale mountain building from continent collisions, the formation of metamorphic belts, rapid climate swings, major atmospheric events, profound sea-level change, and substantial volcanic activity.

A graphic showing the large igneous provinces (LIPs) generated since the formation of Pangaea and their correlation with extinction rates through time.

Figure 2: Large igneous provinces (LIPs) generated since the formation of Pangaea and their correlation with extinction rates through time. Continuous line/blue field, extinction rate through time; red columns, LIP eruption ages. (From Coffin et al. (2006) Oceanography 19(4), 150-160, doi.org/10.5670/oceanog.2006.13, CC BY-SA 4.0, based on extinction data from Sepkoski (1996) and eruption data from White & Saunders (2005))

Large igneous provinces

The supercontinent Pangaea had a significant effect on the thermal evolution of our planet, expressed as large amounts of magmatism. Voluminous emplacements of magma often occurred over relatively short geological timescales (typically 1–5 million years) and are referred to as large igneous provinces (LIPs).

For over 30 years, researchers have investigated the temporal, spatial, and geodynamic relationships between LIPs and plate tectonics. It seems that relatively little LIP activity is associated with the amalgamation stage of the supercontinent cycle. However, after a supercontinent has existed for a period of time (100–200 million years), the number of LIPs increases on a global scale – creating another geomarker. For example, approximately 26 LIPs have been emplaced globally since the formation of Pangaea (with a selection of them shown in Fig. 2).

Several studies have linked LIP events to past mass extinctions

It isn’t clear what causes large igneous provinces to form, but they have been linked to thermal plumes of material upwelling from the deep mantle. Alternative interpretations of the origin and source of LIPs have been suggested but are contested, whereas a deep mantle source for some of the LIPs seems to be supported by tomographic images of the mantle beneath modern hotspots (essentially an ultrasound of the current mantle structures; e.g., French & Romanowicz, 2015). Furthermore, the geochemistry of such LIPs often shows a signature related to the deep mantle, including recycled oceanic lithosphere (e.g., Sobolev et al., 2007; Williams et al., 2019).

Why are LIPs important? Well, several studies have linked these LIP events to past mass extinctions (Fig. 2), so it is useful to understand the processes by which they form.


A graphic showing the supercontinent cycle for Pangaea and oceanic large igneous provinces (LIPs) (left column) and schematic illustrations of what might be going on in the mantle (right column).

Figure 3: The supercontinent cycle for Pangaea and oceanic large igneous provinces (LIPs) (left column) and schematic illustrations of what might be going on in the mantle (right column). Blue lines, ancient oceanic ridges; red lines with teeth, subduction zones; grey, present-day continental interiors; Orange, LIPs (CAMP, Central Atlantic Magmatic Province; OJP, Ontong Java Plateau; KR, Karoo Ridge; BU, Bunbury Basalts; CLIP, Caribbean large igneous province; SR, Shatsky Rise). (Image from Heron et al. (2023) Geological Society, London, Special Publications 542; https://doi.org/10.1144/SP542-2023-12, CC BY 4.0 DEED, based on the global plate reconstructions of Matthews et al., 2016))

The supercontinent cycle

Although there is no formal definition of what constitutes a supercontinent, the processes by which continents amalgamate and disperse are well established as part of the life cycle of an ocean (e.g., the Wilson Cycle), and involve the following steps:

  1. The subduction-induced closure of ocean basins causes large-scale continental collisions (Fig. 3a). Any interior subduction that yields the assembly of the newly amalgamated landmass terminates (Fig. 3b).
  2. The cessation of the subduction that brought the continents together changes the dynamics both within and below the supercontinent. The initiation and development of circum-supercontinent exterior subduction, and the related descent of oceanic lithosphere into the mantle, disrupts the regional mantle dynamics and produces a return flow that can potentially trigger deep mantle plumes.
  3. After an interval of time (typically, on the order of hundreds of millions of years), coupled lithosphere and mantle dynamics ultimately break the landmass up, leading to continental dispersal. These dynamics may be related to plate-driving forces, including a mantle-push force from newly formed plumes beneath the supercontinent and/or slab-pull from circum-supercontinent subduction zones, as well as mantle traction and insulating temperature effects.

The link between deep mantle plumes and plate tectonics has often involved the subduction of oceanic lithosphere. In particular, previous studies have shown that sinking slabs of oceanic lithosphere have the power to stir mantle flow and control the thermal evolution of the mantle, which is of note during amalgamation and advanced breakup phases of the supercontinent cycle, when subduction patterns change on a global scale.

Most of the discussion on the supercontinent cycle and formation of plumes has centred on the formation of LIPs and/or the break-up of a supercontinent. In particular, the discussion has focused on the formation of the Central Atlantic Magmatic Province (CAMP) and Paraná-Etendeka LIPs, which may have impacted the opening of the Atlantic, and the Deccan Traps LIP, which may have influenced India’s separation from Madagascar. While today these LIPs are mostly preserved within the ocean (Fig. 4), at the time of their formation, these LIPs were largely emplaced within the interior of Pangaea, and so were first manifested on Earth’s surface within Pangaea’s previous position, bordered by subduction.

The focused, long-term subduction of oceanic lithosphere on the margin of the supercontinent may have been capable of forcing a return flow within the mantle beneath Pangaea, thereby producing sub-supercontinental plumes as discussed in a number of previous studies (e.g., Zhong et al., 2007; Yoshida, 2010; Li & Zhong, 2009).

However, a number of LIPs that formed after supercontinent amalgamation do not lie within this continental interior framework. Those in the Pacific Ocean (oceanic LIPs or plateaux), for example, formed outside the previous site of Pangaea (i.e., in areas that were exterior to the supercontinent) on plates that were actively subducting beneath it. LIPs such as the Ontong Java Plateau, Caribbean LIP, and Shatsky Rise all formed in provinces exterior to Pangaea (Fig. 3), yet investigations into their formation in the context of circum-supercontinent subduction is limited.


A graphic showing the age of the oceanic lithosphere.

Figure 4: Age of the oceanic lithosphere, from Straume et al. (2019), with superposed oceanic large igneous provinces (LIPs) from Torsvik & Cocks (2016) coloured in light blue. NAIP, North Atlantic Igneous Province; HALIP, High Arctic Large Igneous Province. (Graphic by Eivind O. Straume, published open access (CC BY-SA 4.0) and made available via the open-access s-Ink repository at s-ink.org/oceanic-large-igneous-provinces, based on data from Straume, et al. (2019) G3 20, 1756-1772; https://doi.org/10.1029/2018GC008115)

The Pacific Ocean

The Pacific Ocean LIPs are difficult to study, given their remote locations. However, the Ontong Java Plateau is the world’s largest oceanic plateau and is linked to a massive volcanic event that significantly altered global climate and ocean oxygen levels, leading to a mass extinction event.

The position and geometry of the Ontong Java Plateau has changed significantly over time due to drifting of the Pacific Plate. The three currently separated plateaux of Ontong Java, Manihiki, and Hikurangi were originally formed in a single complex known as the greater Ontong Java Plateau. While the greater Ontong Java Plateau is thought to have been emplaced by two major magmatic events at about 122 Ma and 90 Ma, the cause of the magmatism is still debated. Two main hypotheses relate its formation to a mantle plume (i.e., melting anomalies arising from hot diapirs ascending from the deep mantle: e.g., Tarduno et al., 1991; Isse et al., 2021) or to a plate boundary interaction (i.e., melting anomalies fuelled by shallow mantle processes associated with plate tectonics: e.g., Nakanishi et al., 1999; Frey et al., 2000; Sager et al., 2019).

The Ontong Java Plateau is the world’s largest oceanic plateau and is linked to a massive volcanic event

Although a number of studies have assessed the viability of a mantle plume origin, including studies that use seismic evidence, there is little discussion of the mechanism that would have generated the upwelling. The Louisville hotspot, for example, is often discussed as a potential source of the Ontong Java Plateau, but the mechanism that caused the Louisville plume to form at its specific time and location is not known.

The Shatsky Rise, Earth’s third largest oceanic plateau, formed exterior to Pangaea at approximately 145 Ma. Like the Ontong Java Plateau, two main hypotheses exist for its formation: a plume origin or plate-boundary interaction (seafloor spreading).

The Caribbean LIP is among the most recently formed, having largely been emplaced between 95 and 88 Ma, and has a similar chemical composition to the Ontong Java Plateau. Although the suggested formation mechanisms for the Caribbean LIP vary, the Galápagos hotspot has been proposed as a potential deep mantle plume source.

One proposed mechanism for generating the LIPs that occurred beneath the Pangaean landmass is continental insulation, where the thick supercontinent insulates the mantle resulting in significant volumes of melt. This process would not be appropriate for the Pacific Ocean, however, because the thin oceanic lithosphere would not trap the heat in an effective manner.

It is possible that both the continental and oceanic LIPs (those interior and exterior to Pangaea) formed via the same mechanism, one whereby the subduction of ancient oceanic lithosphere can generate a return flow to produce a deep mantle plume (Fig. 5).


A graphic showing a proposed mechanism to generate oceanic and continental LIPs.

Figure 5: A proposed mechanism to generate oceanic and continental LIPs. Subducting oceanic lithosphere (blue) may force deep mantle plumes (red) to upwell both below the Pangaean supercontinent (thick grey) and neighbouring oceanic basin (what is now the Pacific). (Image from Heron et al. (2023) Geological Society, London, Special Publications 542; https://doi.org/10.1144/SP542-2023-12, CC BY 4.0 DEED)

Numerical modelling

We set about testing the hypothesis of both continental and oceanic LIP formation via subduction-induced return flow in the deep mantle using numerical modelling. The full details of our modelling approach are available in Heron et al. (2024). We essentially studied global mantle convection patterns using two-dimensional experiments that simulate the formation of a supercontinent and compressible convection in Earth’s mantle. Our models also account for the anomalously warm material present at the base of the mantle beneath the Pacific and African plates in the present-day Earth (Garnero et al. 2016), with the latter lying below the reconstructed site of the latest supercontinent, Pangaea (e.g., Torsvik et al. 2010; Burke et al. 2008; Heron et al. 2021; Murphy et al. 2021).

We simulated models over sufficient time to allow for plumes, slabs, and supercontinents to form from the homogenous initial conditions, as well as time for these components to then influence mantle flow. Our 2D and 3D model results confirm that subduction related to the supercontinent cycle and the subsequent deep mantle return flow can indeed reproduce both the location and timing of the Ontong Java Plateau, Caribbean LIP and potentially the Shatsky Rise. This mechanism simplifies the need for different processes to create oceanic and continental LIPs, and instead relies on mantle return flow to create the dynamics for post-supercontinent formation.


A graphic showing potential future supercontinents.

Figure 6: Potential future supercontinents. (A) Novopangaea; (B) Pangaea Ultima; (C) Aurica; (D) Amasia. (Images modified from Green et al. The Conversation (2018). Copyright of and supplied with kind permission from Hannah Davies and Joao Duarte.)

Future supercontinents?

The hard-to-reach mantle, located hundreds to thousands of kilometres beneath Earth’s surface, is difficult to understand. In their study of supercontinents, Pastor-Galan and colleagues (2018) defined them as ‘single continental plates with a size capable of modifying or controlling mantle dynamics and core-mantle boundary processes, altering convection cells and enhancing thermal activity’, implying that perhaps we should define supercontinents on their ability to trigger fundamental changes within Earth’s interior – a colossal shifting of the material beneath our feet.

The geological community’s crystal ball is divided on where our continents are moving to next – will we have to wait many millions of years to figure out where our next supercontinent will be? If the Pacific Ocean continues to shrink, we may end up with the Novopangaea supercontinent (Fig. 6a) as postulated by Roy Livermore in the 1990s. If the Atlantic Ocean reverses its growth and closes to reform Pangaea, Pangaea Ultima (also known as Pangaea Proxima) will be next (Scotese, 2003; Fig. 6b). However, if both major ocean basins collapse, we may end up with Aurica (Fig. 6c; Duarte et al., 2018). A final scenario could be continuation of the northward movement of several continents to create a northern hemisphere supercontinent called Amasia (Mitchell et al. 2012; Fig. 6d).

Or perhaps other continental configurations will form in between those proposed that could be deemed ‘super’? Will any of these future supercontinents create the disruption caused by Pangaea?

What is clear is that any future supercontinent will be defined by its impact on Earth’s surface and mantle evolution, rather than the continental configuration.


Philip J. Heron

University of Toronto Scarborough, Canada

Erkan Gün

University of Toronto Scarborough, Canada

Grace E. Shephard

University of Oslo, Norway, and Australian National University, Australia

Juliane Dannberg

University of Florida, USA

Rene Gassmöller

University of Florida, USA

Erin Martin

IGO Limited, Australia

Aisha Sharif

University of Toronto Scarborough, Canada

Russell N. Pysklywec

University of Toronto, Canada

R. Damian Nance

Ohio University, USA, and Yale University, USA

J. Brendan Murphy

St Francis Xavier University, Canada

Further reading

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