Geologic and kinematic constraints on Late Cretaceous to mid Eocene plate boundaries in the southwest Pacific
Introduction
The southwest Pacific (Fig. 1) has had a complex tectonic history since final Gondwanaland dispersal began, dominated by multiple episodes of marginal and back-arc basin formation, subduction and trench rollback (e.g. Crawford et al., 2003, Sdrolias et al., 2003, Schellart et al., 2006, Whattam et al., 2008, Cluzel et al., 2012a, Cluzel et al., 2012b). Understanding the history of this complicated plate boundary activity in the southwest Pacific is of regional and global significance, including for understanding basin subsidence and hydrocarbon formation along the Lord Howe Rise (LHR) and around New Zealand, understanding the mechanisms accommodating Tasman Sea spreading (e.g. Schellart et al., 2006), and modeling mantle plumes within global plate kinematic models. However, less than 10% of the continental crust now in the South Pacific that rifted away from Australia and Antarctica is presently sub-aerially exposed and readily accessible for field exploration (Mortimer, 2008). Sub-aerial exposures are largely complicated by Cenozoic crustal thickening events in New Zealand and New Caledonia. Furthermore, data coverage in several offshore locations (e.g. Coral Sea) is sparse, and partial or complete basin subduction has destroyed large swaths of ocean crust. The magnetic isochrons created at the Tasman, Amundsen and Bellinghausen ridge systems provide the earliest robust constraints on the position of the Pacific plate relative to the rest of the global plate network — prior to this time the Pacific was completely surrounded by subduction zones (Seton et al., 2012). However, while spreading in the Tasman, Amundsen and Bellingshausen seas is reasonably well-constrained, alternative models exist for the early spreading history between Australia and Antarctica (Royer and Rollet, 1997, Tikku and Cande, 1999, Tikku and Cande, 2000, Whittaker et al., 2007, Whittaker et al., 2013). Together, these data gaps have made it difficult to build well-constrained regional plate reconstruction models, and there remain many unresolved and hotly debated problems relating to southwest Pacific evolution. Questions remain concerning the timing, location and polarity of different subduction episodes, the driving mechanism for obduction events in New Caledonia and New Zealand, the mechanism for Tasman Sea opening and widespread rifting in Zealandia (Fig. 1), seafloor ages and the orientation of spreading in various basins (e.g. d'Entrecasteaux Basin), and the origin of several submerged tectonic features (e.g. South Rennell Trough). Conversely, the existence of Tonga–Kermadec subduction (Bloomer et al., 1995) and plate boundary activity within New Zealand (Sutherland, 1995) from at least 45 Ma is well established. Many of the unresolved issues related to southwest Pacific evolution are related to the poorly understood timing and type of plate boundary activity to the east of the LHR during the Late Cretaceous to mid Eocene. This problem will, therefore, be the focus of our study.
A robust reconstruction model for the southwest Pacific has wider implications for global geodynamic studies. Plates in the Pacific realm can only be linked with the motion of plates in the Indo-Atlantic realm via a southwest Pacific spreading corridor. The ability to link the Pacific plate to a relative plate motion circuit connected to the Indo-Atlantic realm enables tighter constraints to be placed on its relative and absolute motion history, and that of other plates in the Pacific ocean basin, such as the Kula and Farallon plates. This in turn affects computation of subduction budgets around circum-Pacific active margins. Predictions of the angle of convergence and amount of slab material subducted in the western, northern and eastern Pacific are influenced by how the Pacific plate is reconstructed (Sutherland, 2008). Furthermore, determining if and how Pacific hotspots have moved relative to Indo-Atlantic hot spots is a long-standing problem in plate tectonics and is of paramount importance for constructing absolute plate motion models (e.g. Tarduno and Gee, 1995). Global models of absolute plate motions based on the time-progressive volcanism along hot spot trails (Steinberger et al., 2004, Cuffaro and Doglioni, 2007, Doubrovine et al., 2012) rely on accurately linking observations between the Indo-Atlantic and Pacific realms. Steinberger et al. (2004) proposed that the ~ 50 Ma bend in the Hawaiian–Emperor seamount chain is best predicted using a plate model assuming no plate boundary between the Pacific and the LHR before chron 20 (43 Ma - timescale of Gee and Kent, 2007 is adopted). However, the assumptions within this plate-circuit are controversial (Schellart et al., 2006, Tarduno et al., 2009). A better understanding of the history of west versus east-dipping subduction in the southwest Pacific and how these intervals are related to episodes of opening versus closure of back-arc basins may contribute to our understanding of the effects of geographic polarity on subduction kinematics (Doglioni et al., 2007).
To better constrain the broad scale tectonic evolution of the southwest Pacific during the Late Cretaceous to mid Eocene period, and so provide a better understanding of how the Pacific plate should be reconstructed with respect to a relative plate motion chain, we review geologic and kinematic data from the southwest and southern Pacific. Our paper is divided into two parts. In Part A we review geologic data from the southwest Pacific and West Antarctic Rift System to better constrain the timing, types and locations of plate boundaries that were active during the Late Cretaceous to mid Eocene. In Part B we investigate the kinematic consequences of adopting alternative Pacific plate motion circuits for motion in the West Antarctic Rift System and between the LHR and Pacific, where relative motion histories are poorly constrained. This process allows us to identify which plate circuits produce realistic or unrealistic amounts of motion and implied deformation, based on a comparison with geologic observations and documented histories of relative motion and quiescence.
It is well established that subduction along the eastern margin of Gondwanaland was long-lived. Subduction related magmatism preserved in the New England Fold Belt (Leitch, 1975, McPhie, 1987) is evidence that a convergent margin paralleled Eastern Gondwanaland from at least the Carboniferous. Arc related products are also preserved in the Téremba Terrane in New Caledonia (Paris, 1981; Campbell, 1984, Adams et al., 2009) and the Murihiku Terrane and Median Batholith in New Zealand (Ballance and Campbell, 1993, Mortimer et al., 1999, Roser et al., 2002). Subduction along Eastern Gondwanaland waned in the mid Cretaceous at about 105–100 Ma (Veevers, 1984, Bradshaw, 1989, Laird and Bradshaw, 2004), coinciding with a major plate boundary reorganization event (Veevers, 2000, Matthews et al., 2012). Convergence gave way to strike-slip motion along the margin (Veevers, 1984, Sutherland and Hollis, 2001, Siddoway, 2008), and ultimately widespread continental extension and fragmentation (Gaina et al., 1998, Sutherland, 1999, Rey and Müller, 2010), represented by a regional-scale unconformity (‘Eastern Gondwana Composite Surface’) in submerged portions of Zealandia (Bache et al., 2014), and a major near continent-wide angular unconformity in New Zealand (Laird and Bradshaw, 2004). Observations from Marie Byrd Land and southernmost New Zealand (Kula et al., 2007, Siddoway, 2008, McFadden et al., 2010, Saito et al., 2013) indicate widespread continental deformation, with a transition from wrench to transtensional regimes before intitation of seafloor spreading within the Tasman and Amundsen seas at around chron 34y time (83 Ma), although Vry et al. (2004) suggested that subduction may have continued along the New Zealand portion of the margin until ~ 85 Ma. Evidence for subduction continuing after 105–100 Ma (Worthington et al., 2006) comes from the South Island from calc-alkaline activity at ~ 89 Ma in the Canterbury region (Smith and Cole, 1997) and high-grade metamorphism in the Alpine Schist at ~ 86 Ma (Vry et al., 2004). Elsewhere around the circum-Pacific subduction continued beneath eastern Asia, North and South America, and the Antarctic Peninsula.
Another well accepted aspect of southwest Pacific evolution is that from at least 45 Ma to the present one, or multiple plate boundaries have separated the Pacific plate from the LHR. Tonga–Kermadec subduction was active from at least 45 Ma, as supported by dated arc tholeiites from the Tonga forearc and Tongan island of 'Eua (e.g. Duncan et al., 1985, Bloomer et al., 1995), and basin subsidence events (e.g. Sutherland et al., 2010, Bache et al., 2012, Hackney et al., 2012) of a similar age. At ~ 45 Ma it has also been proposed that rifting initiated between the Challenger and Campbell plateaus and a plate boundary propagated through New Zealand which has been active ever since (Sutherland, 1995).
Starkly contrasting kinematic interpretations have been proposed for the tectonic evolution of the southwest Pacific in the intervening period (~ 85–45 Ma). Schellart et al. (2006) proposed that from 82–45 Ma there was 1500 km of subduction of the Pacific plate, while Steinberger et al. (2004) proposed that the LHR was part of the Pacific plate, with no intervening plate boundary. Amongst published tectonic reconstructions for the Late Cretaceous to Eocene evolution of the southwest Pacific, these two scenarios can be considered as end-member cases (Mortimer et al., 2007) (Fig. 2). We are going to consider four alternative reconstruction scenarios for the evolution of the southwest Pacific which explore different plate boundary configurations to the east of the LHR (Fig. 2). The first end-member, Model 1, involves continuous subduction in the southwest Pacific to the north of New Zealand, with one or several subduction polarity reversals (Crawford et al., 2003, Sdrolias et al., 2003, Sdrolias et al., 2004, Schellart et al., 2006, Whattam et al., 2008, Cluzel et al., 2012a, Cluzel et al., 2012b, Meffre et al., 2012). In striking contrast is end-member scenario Model 2, of tectonic quiescence east of the LHR until 45 Ma (Steinberger et al., 2004, Mortimer et al., 2007, Sutherland et al., 2010, Doubrovine et al., 2012). Alternative reconstruction models that incorporate subduction to the east of the LHR have been presented for this timeframe that vary in subduction polarity and timing of subduction, these are presented as models 1a–c.
In the southwest Pacific plate reconstruction model of Schellart et al. (2006) west-dipping subduction occurs to the east of New Caledonia driven by convergence between the Pacifc and LHR. Eastward slab roll-back drives opening of the South Loyalty Basin to the east of New Caledonia from ~ 85–55 Ma. Schellart et al. (2006) further suggested that back-arc basin extension associated with the west-dipping subduction zone accommodated opening of the New Caledonia Basin to the west of New Caledonia from 62–56 Ma (after Lafoy et al., 2005), and partially accommodated opening of the Tasman Sea. In this model west-dipping subduction is continuous in the southwest Pacific to present-day, and occurs contemporaneous with east-dipping subduction that initiates at ~ 50 Ma in the South Loyalty Basin. This east-dipping subduction zone had consumed the South Loyalty Basin and subduction ended diachronously with New Caledonia obduction at ~ 38 Ma in the north and Northland oduction at ~ 25 Ma in the south.
Model 1b is primarily based on the reconstructions of Crawford et al. (2003) and Whattam et al. (2008). In this model west-dipping subduction to the east of an unnamed continental ribbon, and eastward slab roll-back accommodate opening of the South Loyalty Basin from 85–55 Ma. A subduction polarity reversal occurs at 55 Ma, with the initiation of northeast-dipping subduction in the South Loyalty Basin at the site of the recently extinct spreading ridge. Reinitiation of west-dipping subduction occurs at ~ 50 (Whattam et al., 2008) or 45 Ma (Crawford et al., 2003, Meffre et al., 2012). In contrast to the model of Schellart et al. (2006) west-dipping subduction ceases during the Cretaceous, and east and west-dipping subduction is only contemporaneous for a short interval of time towards the end of the period during which the South Loyalty Basin is closing. In the model of Crawford et al. (2003) the subduction polarity flip at 55 Ma occurs across the unnamed continental ribbon to the east of New Caledonia. In the model of Whattam et al. (2008) an intra-oceanic arc is established by northeast-dipping subduction within the South Loyalty Basin that is separate from this continental ribbon further west. Relative motion between the Pacific and LHR is not specified in these models.
Eissen et al. (1998) also support a model of opening of the South Loyalty Basin from at least 85–55 Ma, including subduction initiation at ~ 50 Ma within the basin. However due to a lack of evidence for a Late Cretaceous arc to the east of New Caledonia their model does not incorporate subduction prior to ~ 50 Ma, rather the South Loyalty Basin opens as a marginal basin.
In contrast to the southwest Pacific reconstructions described above, the model of Sdrolias et al. (2004) and Sdrolias et al. (2003) incorporates long-lived east-dipping subduction rather than west-dipping subduction. In this model east-dipping subduction beneath the Loyalty arc from ~ 85–45 Ma closes a Cretaceous aged basin that existed to the east of New Caledonia. Obduction onto New Caledonia begins at ~ 45 Ma. According to a later iteration of this model (Seton et al., 2012), this east-dipping subduction zone specifically closes an Early Cretaceous aged back-arc basin that opened as a result of eastward roll-back of the west-dipping Eastern Gondwanaland subduction zone from 140–120 Ma.
An objection to Models 1a–c is that relative motion between the LHR and the Pacific implies a plate boundary running through New Zealand throughout the Late Cretaceous and early Paleocene, whereas most observations suggest this was a period of tectonic quiescence in New Zealand (Sutherland, 2008). Further weighing against Model 1 is the lack of evidence for arc magmatism expected from subduction beneath the LHR or a continental fragment rifted from Eastern Gondwanaland between ~ 100 and 55 Ma (Eissen et al., 1998). Instead of subduction and backarc basin formation, this scenario favours widespread extension in the southwest Pacific during the Late Cretaceous to mid Eocene that was unrelated to subduction zone processes, followed by Tonga–Kermadec subduction initiation at ~ 45 Ma coincident with plate boundary activity in New Zealand (Steinberger et al., 2004, Mortimer et al., 2007, Sutherland et al., 2010, Doubrovine et al., 2012). In this model the ~ 84–55 Ma Tasman Sea mid-ocean ridge formed the plate boundary between the Pacific and Australian plates and elsewhere there was no plate boundary activity, rather the LHR was part of the Pacific plate (Mortimer et al., 2007).
Section snippets
PART A: Geologic constraints on plate boundary activity
In the following sections we review geologic and geophysical data from the southwest Pacific and West Antarctic Rift System that provide information about the nature and timing of plate boundary activity in the region during the Late Cretaceous to mid Eocene. These data are summarized in Fig. 3 and Table 1. Where possible we focus on primary geologic observations rather than tectonic reconstruction models. We also outline the constraints on relative motion between East and West Antarctica, the
Part B: Kinematic constraints on plate boundary activity
In the following sections we investigate the kinematic consequences of using different plate circuits in the southwest Pacific for the period from 74 to 45 Ma. These results will later be discussed in light of the geologic data reviewed above, in order to better understand the broad evolution of southwest Pacific plate boundaries during this timeframe. We have chosen to focus on times younger than 74 Ma as chron 33y (~ 74 Ma) is the oldest magnetic anomaly that has been widely identified in the
Southwest Pacific plate boundaries during the Late Cretaceous to mid Eocene
An important and unresolved question concerning the tectonic evolution of the southwest Pacific is whether or not there was relative motion and plate boundary activity between the Pacific and LHR from the Late Cretaceous to mid Eocene, prior to the widely accepted onset of Tonga–Kermadec subduction from at least ~ 45 Ma (e.g. Bloomer et al., 1995), and following long-lived subduction beneath Eastern Gondwanaland from the Carboniferous to at least the mid Cretaceous (e.g. Leitch, 1975, Paris, 1981
Conclusions
We combined geologic observations from the southwest Pacific with kinematic data to determine how the region evolved from the Late Cretaceous to mid Eocene (~ 85–45 Ma), including the nature and timing of plate boundary activity. This allowed us to place tighter constraints on the time-dependent evolution of the southwest Pacific regional plate circuit so that motion between the plate pairs is consistent with geologic observations and known tectonic regimes.
Based on our analysis of currently
Acknowledgements
K.J.M. was supported by an Australian Postgraduate Award, S.E.W. and R.D.M. were supported by ARC grant FL0992245, M.S. was supported by ARC grant DP0987713 and would like to thank support from Statoil, and G.L.C. was supported by ARC grant A39600827. Discussions on southwest Pacific geology held at the Southwest Pacific New Cruise Results Workshop hosted by the Geological Survey of New Caledonia, June 2013, were extremely helpful during preparation of this manuscript. This manuscript also
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