November 2017 LIP of the Month

Printer-friendly versionPrinter-friendly version

Impact of a LIP with a subduction zone: the Hikurangi Plateau story.

Martin Reyners1, Donna Eberhart-Phillips1,2, Stephen Bannister1, Phaedra Upton1 and David Gubbins3


1GNS Science, PO Box 30368, Lower Hutt 5040, New Zealand. E-mail

2Department of Earth and Planetary Sciences, University of California Davis, Davis, CA95616, USA

3School of Earth and Environment, University of Leeds, Leeds LS2 9JT, United Kingdom

The following is based on a series of recent publications that include, but are not limited to, Reyners et al. (2011), Reyners (2013), Love et al. (2015), and Reyners et al. (2017a, b). For full details, please refer to these papers.


Introduction to the Hikurangi Plateau LIP

The Hikurangi Plateau was originally part of Ontong Java – the largest LIP on Earth. Ontong Java Plateau formed ca. 122 Ma ago (Neal et al., 1997). Seismic reflection and refraction measurements indicate that the plateau has a crustal structure resembling that of normal Pacific oceanic crust, but each layer is abnormally thickened by up to a factor of five (Furumoto et al., 1976; Hussong et al., 1979), with total thickness 35–42 km. Also, unusually high P-wave velocities of 8.4–8.6 km/s, appropriate for eclogite (Saunders et al., 1996), have been detected at the base of the northwest and southwest portions of the plateau (Furumoto et al., 1976).

Shortly after formation of the plateau (at ca. 120 Ma), the Hikurangi and Manihiki plateaux rifted from it, and subsequently the Hikurangi Plateau rifted from the Manihiki Plateau at the Osbourn Trough and drifted south (Taylor 2006). The Hikurangi Plateau then impacted the Gondwana convergent margin at ca. 105 Ma, and plateau crust can be traced 50-100 km southward beneath the Chatham Rise (Fig. 1) on seismic reflection profiles (Davy et al., 2008). By 100 Ma the plateau choked the Gondwana margin and subduction ceased. Currently, the western edge of the plateau is subducting anew, because of east-west convergence between the Pacific and Australian plates.

We have now been able to estimate the shape of the subducted parts of the Hikurangi Plateau, using a variety of seismological techniques. Relocation of seismicity using a nationwide 3-D seismic velocity model derived from seismic tomography has allowed us to identify the edge of the plateau in the Hikurangi subduction zone, based on 1) the distribution of seismic velocities in the top of the subducted plate; 2) changes in the thickness of the dipping seismic zone, and 3) steepening of the subducted plate downdip of the buoyant plateau (Reyners et al. 2011). P-wave velocities (Vp) > 8.5 km/s are imaged at the base of the thickest part of the plateau. These are similar to those seen in the base of the parent Ontong Java Plateau. Also, many earthquakes within or close to the New Zealand land mass show fast, high frequency P-wave precursors at North and South Island seismograph stations. These are explained by propagation through a dipping layer of order 10 km thick with Vp around 8.5 km/s (Love et al. 2015). If we associate this layer with the eclogitic base of the Hikurangi Plateau, this means that the subducted plateau is much more widespread beneath New Zealand than previously supposed.

Figure 1. Map of the Hikurangi Plateau LIP, with its currently subducted western edge unrolled to the surface to show its original extent (Reyners et al. 2011, 2017a). Edges shown dashed are less certain. The inset shows the fit of the subducted portions of the plateau (light grey) into a schematic reconstruction (Davy et al. 2008) of the Ontong Java Plateau (OJP – olive), Manihiki Plateau (MP – orange) and Hikurangi Plateau (HP – dark grey) seen at the surface. Note that the southern part of the OJP subducted beneath the Solomon Islands has not been defined in this reconstruction.

Cretaceous impact of the Hikurangi Plateau with the Gondwana Margin

Seismic tomography from earthquakes recorded by a dense seismograph network in the southern South Island have now allowed us to delineate the leading edge of the subducted plateau in this region, and thus refine the original model of Reyners et al. (2011). Three-dimensional images of Vp and Vp/Vs reveal the southwestern part of the plateau was a relatively narrow salient, and the first part to be subducted at ca. 105 Ma (Reyners et al., 2017a). The plateau then rotated clockwise about this salient until its southern edge was parallel to subduction strike and subduction ceased at 100 Ma (Fig. 2).

Figure 2.Schematic of Hikurangi Plateau (HP)impact with the Gondwana margin (GM), after Reyners et al. (2017a). For the HP, dark grey is the part currently at the surface, and light grey is the part subducted during the Cretaceous and current subduction episodes. Dashed lines on the GM denote isobaths on the subducted oceanic crust. (a) At ca. 105 Ma the western end of the thick, buoyant plateau begins to subduct (with difficulty) beneath the GM, and seafloor spreading continues at the Osbourn Trough (OT). (b) The HP rotates clockwise around its pinned western end (black curved arrow) until the southern edge is parallel to the GM and subduction stops at ca. 100 Ma. This rotation is accommodated by an anticlockwise rotation of the OT (red curved arrow) and its associated transform fabric (Davy, 2014). Cessation of subduction at the GM causes extinction of OT spreading (EOT), with extension (straight red arrows) due to slab pull now enhanced in subducted oceanic crust at the GM.

An important consequence of the choking of Gondwana subduction is that there was now enhanced slab pull in the oceanic crust surrounding the southern edge of the Hikurangi Plateau, and especially around the southwestern salient which had undergone flat subduction (Fig. 2). This slab pull was enhanced because cessation of spreading at the Osbourn Trough meant that the component of extension previously accommodated at the trough must now have been taken up within the subducted oceanic crust at the Gondwana margin. Also, extensional stresses would already have existed at the change in crustal thickness (e.g. Turcotte & Schubert 2002). Herein lies an explanation for the opening of the Great South Basin (Carter, 1988), Bounty Trough (Grobys et al., 2007) and its extension into the Canterbury Basin (Fig. 1). All these major basins lie above oceanic crust surrounding the southern edge of the plateau. Similarly, widespread pulses of intraplate magmatism around the southern edge of the plateau occurred at ca. 101, 97 and 88-82 Ma (Tulloch et al., 2009). In contrast, there was little extension or intraplate volcanism where the subducted plateau was present at shallow depth.

We interpret the basin formation and intraplate magmatism in oceanic crust surrounding the plateau as a period of failed rifting, prior to the wholesale rifting of the New Zealand continent (Zealandia) from Gondwana at ca. 85 Ma following detachment of the subducted oceanic crust. Numerical modelling (van Hunen and Allen, 2011) indicates that the delay time between first continental collision and slab break-off depends mostly on the strength of the previously subducted oceanic plate, and ranges from 10 Ma (for young, weak slabs) to >20 Ma (for old, strong slabs). A 15 Ma period between the plateau choking the Gondwana margin and slab break-off would be consistent with this modelling. Prior to slab detachment, wholesale rifting will be inhibited due to a sea anchor force exerted on the slab that resists its lateral motion (Scholz and Campos, 1995), leading to the episode of failed rifting and intraplate volcanism that followed LIP collision.

The central role of the Hikurangi Plateau in the Cenozoic tectonics of New Zealand and the Southwest Pacific


The Cenozoic tectonics of New Zealand and the southwest Pacific has been controlled not only by the resistance to subduction of the buoyant Hikurangi Plateau, but also by the shape of its western edge and changing angle of attack of this edge at the plate boundary (Reyners 2013). From ca. 45 Ma, the westernmost tip of the plateau controlled the transition at the plate boundary from subduction to the north to Emerald Basin spreading to the south. By ca. 25 Ma, extensive trench rollback and backarc opening had occurred to the north of New Zealand (Fig. 3a). Given the difficulty of subducting the thick, buoyant Hikurangi Plateau, continued trench rollback to the north would have led to an increasingly curved termination of the Tonga-Kermadec subduction zone against the western edge of the plateau, resulting

Figure 3.Snapshots of the tectonic setting before (a) and after (b) Alpine Fault formation as a STEP fault (after Reyners 2013, incorporating new information on the shape of the Hikurangi Plateau from Reyners et al. 2017a). The background Southwest Pacific plate reconstruction is that of Schellart et al. (2006), with an Australia-fixed reference frame. The Hikurangi Plateau (HP) is shown in purple, with its currently subducted part light, and the portion yet to be subducted dark. The fossil trench where the plateau partly subducted under Gondwana at ca. 105-100 Ma is marked by the grey barbed line. The Alpine Fault is shown in red, and the strike-slip fault that formed at the Emerald Basin (EB) spreading centre upon Alpine STEP fault formation is shown in blue. TKSZ denotes the Tonga-Kermadec subduction zone, which drove STEP fault formation.

in high strain. At the same time, the Emerald Basin spreading centre was a weak zone in the oceanic crust to the south. This situation led to conditions conducive to the development of a Subduction-Transform Edge Propagator (STEP) fault (Govers & Wortel 2005). In this model, the Alpine Fault forms as the subduction zone rolls back by ripping oceanic crust seaward of the trench from the western edge of the plateau. As the trench rolls back, trench suction takes the overlying plate with it, resulting in varied extension in the overlying plate and strike- slip motion on the Alpine Fault (Fig. 3b). Movement of the overlying plate is facilitated by the weak Emerald Basin spreading centre changing to a strike-slip zone. But because of the presence of the thick and strong Hikurangi LIP, the northern half of the oceanic crust in the Emerald Basin is subducted beneath Fiordland. So the STEP fault model can explain not only the initiation of the Alpine Fault, but also the initiation of subduction beneath Fiordland, with this subduction being of opposite polarity to that of the Tonga-Kermadec-Hikurangi subduction zone. But the key ingredient in this tectonic history has been the Hikurangi Plateau.

At ca. 15 Ma the western edge of the plateau became parallel to the trench, and thus no longer favoured STEP fault formation. Wholesale subduction of the LIP at the Hikurangi subduction zone began at ca. 10 Ma. The development of a subduction décollement above the LIP mechanically favoured deformation within the overlying Australian plate continental crust. This led to inception of the Marlborough fault system in the northern South Island at ca. 7 Ma, and the North Island fault system at 1-2 Ma.

The imprint of Hikurangi Plateau structure and history on current tectonics of New Zealand

Having a LIP underlying most of New Zealand has profound consequences for current tectonics and seismic hazard. We are just beginning to understand these, and here we introduce a recent study, which underlines the effects of the two episodes of subduction that this LIP has suffered.

The southward transition from typical subduction in the North Island to continental collision in the South Island.

We find that this transition is controlled by the Hikurangi Plateau being more dehydrated to the south, as a result of being more deeply subducted at the Gondwana margin (Reyners et al. 2017b). The southward transition from localized slip at the plate interface to distributed upper plate deformation with no active plate interface occurs in Cook Strait and is relatively sharp (Fig. 4). To get localized slip at a plate interface, one needs the pore fluid pressure ratio λ to be > ca. 0.4. This suggests that continental collision in the northern South Island reflects a low value of λ, due to the underlying LIP being drier there. Over time, this situation has led to a transition zone that offsets geological terranes across Cook Strait. The transition zone has on average accommodated ca. 8 mm/a of dextral margin-normal motion since the present episode of plateau subduction began ca. 10 Ma ago (see Reyners et al. 2017b for details). This finding provides context for the 2016 Mw 7.8 Kaikoura earthquake in the northern South Island. This rupture propagated northward for > 170 km along at least 12 major crustal faults, before terminating in Cook Strait (Hamling et al., 2017).

Figure 4.A conceptual model of fluid movement and the nature of plate coupling at the transition zone between the North and South Islands, along the section shown in Fig. 1 (after Reyners et al., 2017b). Crosses are relocated earthquakes for the period 2001-2011, and arrows show fluid movement. The red line shows the region of the plate interface where the slip rate deficit is > 20 mm/a (Wallace et al. 2012). The northern extent of crustal faulting in the 2016 Mw 7.8 Kaikoura earthquake is shown by the red dashed line. The transition zone in upper plate deformation is controlled by the pore fluid pressure ratio λ in the underlying Hikurangi Plateau LIP, which reflects its Gondwana subduction history.


  • The type of impact a LIP has with a subduction zone depends on both the shape of the LIP and its angle of attack at the subduction zone.
  • The choking of Gondwana subduction by the Hikurangi Plateau LIP also led to a concentration of slab pull in the adjacent subducted oceanic crust, explaining the episode of basin opening and intraplate magmatism that occurred at the same time.
  • Such an episode of extension has led to previous models suggesting that the global-scale plate reorganization event at 105-100 Ma was due to subduction of mid ocean ridges. Our results suggest that this event was instead caused by the Hikurangi Plateau LIP choking the Gondwana subduction zone.
  • Once embedded into a continental margin, the strength and thickness of a LIP means it can continue to play an important role in subsequent tectonic history. From ca. 45 Ma, the westernmost tip of the Hikurangi Plateau LIP controlled the transition at the Pacific/Australia plate boundary from subduction to the north to Emerald basin opening to the south. And at ca. 23 Ma, the western edge of the plateau controlled the concomitant development of the Alpine Fault and Fiordland subduction zone.
  • The subaerial part of Zealandia (i.e. the New Zealand land mass) is mostly underlain by the Hikurangi Plateau. The structure and tectonic history of this underlying LIP provide important information on current tectonics and seismic hazard.

Click to open/close ReferencesReferences

Carter, R. M., 1988. Plate boundary tectonics, global sea-level changes and the development of the eastern South Island continental margin, New Zealand, Southwest Pacific, Mar. Pet. Geol., 5, 90-107.

Davy, B., 2014. Rotation and offset of the Gondwana convergent margin in the New Zealand region following Cretaceous jamming of Hikurangi Plateau large igneous province subduction. Tectonics, 33, 1577-1595.

Davy, B., Hoernle, K. & Werner, R., 2008. The Hikurangi Plateau - crustal structure, rifted formation and Gondwana subduction history, Geochem. Geophys. Geosyst., 9, Q07004, doi:07010.01029/02007GC001855.

Furumoto, A.S., Webb, J.P., Odegard, M.E. & Hussong, D.M., 1976. Seismic studies on the Ontong Java Plateau, 1970, Tectonophysics, 34, 71-90.

Govers, R. & Wortel M.J.R. 2005. Lithosphere tearing at STEP faults; response to edges of subduction zones, Earth Planet Sci Lett. 236, 505-523.

Grobys, J.W.G., Gohl, K., Davy, B., Uenzelmann-Neben, G., Deen, T. & Barker, D., 2007. Is the Bounty Trough off eastern New Zealand an aborted rift? J. Geophys. Res., 112, B03103.

Hamling, I.J. and 28 others, 2017. Complex multi-fault rupture during the 2016 Mw 7.8 Kaikoura earthquake, New Zealand, Science, 356, 6334.

Hussong, D.M., Wipperman, L.K. & Kroenke, L.W., 1979. The crustal structure of the Ontong Java and Manihiki oceanic plateaus, J. Geophys. Res., 84, 6003-6010.

Love, H., LeGood, M., Stuart, G., Reyners, M., Eberhart-Phillips, D. & Gubbins, D., 2015. Fast P-wave precursors in New Zealand: high velocity material associated with the subducted Hikurangi Plateau, Geophys. J. Int., 202, 1223-1240.

Neal, C.R., Mahoney, J.J., Kroenke, L.W., Duncan, R.A. & Petterson, M.G., 1997.The Ontong Java Plateau, Geophys. Monogr., 100, 183-216.

Reyners, M., 2013. The central role of the Hikurangi Plateau in the Cenozoic tectonics of New Zealand and the Southwest Pacific, Earth planet. Sci. Lett., 361, 460-468.

Reyners, M., Eberhart-Phillips, D. & Bannister, S., 2011. Tracking repeated subduction of the Hikurangi Plateau beneath New Zealand, Earth planet Sci. Lett., 311, 165-171.

Reyners, M., Eberhart-Philips, D., Upton, P. & Gubbins, D., 2017a. Three-dimensional imaging of impact of a large igneous province with a subduction zone, Earth planet. Sci. Lett., 460, 143-151.

Reyners M, Eberhart-Phillips D, Bannister S. 2017b. Subducting an old subduction zone sideways provides insights into what controls plate coupling, Earth Planet Sci Lett. 466, 53-61.

Saunders, A.D., Tarney, J., Kerr, A.C. & Kent, R.W., 1996. The formation and fate of large oceanic igneous provinces, Lithos, 37, 81-95.

Schellart, W.P., Lister, G.S & Toy, V.G., 2006. A late Cretaceous and Cenozoic reconstruction of the Southwest Pacific region: tectonics controlled by subduction and slab rollback processes, Earth Sci. Rev. 76, 191-233.

Scholz, C.H., & Campos, J., 1995. On the mechanism of seismic decoupling and back arc spreading at subduction zones, J. Geophys. Res., 100, 22,103-22,115.

Taylor, B., 2006. The single largest oceanic plateau: Ontong Java-Manihiki-Hikurangi. Earth planet Sci. Lett., 241, 372-380.

Tulloch, A.J., Ramezani, J., Mortimer, N., Mortensen, J., van den Bogaard, P. & Maas, R., 2009. Cretaceous felsic volcanism in New Zealand and Lord Howe Rise (Zealandia) as a precursor to final Gondwana break-up, Geol. Soc. London Spec. Publ., 321, 89-118.

Turcotte, D.L. & Schubert, G., 2002. Geodynamics (2nd edition). Cambridge University Press.

van Hunen, J. & Allen, M.B., 2011. Continental collision and slab break-off: a comparison of 3-D numerical models with observations, Earth Plan. Sci. Lett., 302, 27-37.

Wallace, L.M., Barnes, P., Beavan, J., Van Dissen, R., Litchfield, N., Mountjoy, J., Langridge, R., Lamarche, G. & Pondard, N., 2012. The kinematics of a transition from subduction to strike-slip: an example from the central New Zealand plate boundary, J. Geophys. Res., 117, B02405.