A Shallow-Source Plate Model for Intraplate Volcanism in the Panthalassan And

A Shallow-Source Plate Model for Intraplate Volcanism in the Panthalassan and

Pacific Ocean Basins

Alan D. Smith

Department of Geological Sciences, University of Durham, Durham, DH1 3LE, UK

ABSTRACT

The record of intraplate volcanism in the Panthalassan and Pacific basins since the mid

Paleozoic can be explained by tapping of shallow sources with the location of volcanism controlled by localised mantle convection, the stress field acting on the plate, and lithospheric architecture. The source components include recycled subducted oceanic crust remixed into the depleted mantle by mantle convection, volatile-bearing mineral assemblages formed by fluxing of fluids into the mantle at convergent margins, and continental mantle/lower crust incorporated into the shallow mantle during the break-up of eastern Gondwana or the assembly of east Asia in the mid Paleozoic-Early Mesozoic. The latter components migrated eastwards across the basin as a result of lithospheric lag at an average rate of 3 cm yr-1 and are now associated with the South Pacific Isotope and Thermal (SOPITA) anomaly and the source of Hawaiian volcanism. In the Carboniferous to Early Jurassic, ocean island volcanism occurred in the west of the Panthalassan basin accompanying rifting of continental blocks from eastern Gondwana. The crustal fragments and islands were subsequently transported to

Asia on the Izanagi plate and to North America on the Farallon plate. In the Early Jurassic to

1 mid Cretaceous, oceanic plateaus formed on ocean ridge systems around the margins of a growing Pacific plate as a result of entrainment of low-melting point heterogeneities into the ridge upwelling and focusing of melt by triple junction geometry. Seafloor spreading was curtailed by jamming of subduction in the southwest of the basin leading to formation of the

Ontong Java plateau by ponding of melt over the Pacific-Phoenix ridge. Diffuse ocean island volcanism in the Marcus-Wake, Magellan and Marshall islands occurred in the centre of the

Pacific plate as a result of stresses transmitted from the Pacific-Phoenix, Pacific-Farallon and

Pacific-Izanagi ridges. Extensive ocean island volcanism also occurred as a result of asthenospheric shearing beneath the westerly-subducting Izanagi plate, generating many of the ocean islands fragments now preserved as terrane material in east Asia. The Pacific plate began to subduct under Asia in the Late Cretaceous, but the stress field was still dominated by the

Pacific-Kula and Pacific-Antarctic ocean ridges, such that volcanism in the Marshall-Gilbert and Line Island chains may have been oblique to plate motion. The Emperor chain is ascribed to fracturing propagating from the Kula-Pacific ridge following plate re-organisations at ~82

Ma which prematurely halted triple-junction volcanism on Meiji seamount. In the mid

Eocene, the cessation of spreading on the Kula-Pacific and North New Guinea-Pacific ocean ridges subjected the west of the Pacific plate to tensional stresses orthogonal to the direction of plate movement, causing generation of Hawaiian chain as a fracture propagating from the tip of the Emperor chain. Inter-chain correlations of volcanic output along the Hawaiian,

Louisville and Cook-Austral-Marquesas chains with Late Oligocene and Miocene-Pliocene plate re-organisations around the margins of the basin, similarly indicate control by lithospheric architecture and the stresses acting on the plate. Three plate settings with corresponding styles of intraplate volcanism are thus recognised: (1) intra-oceanic, marked by

2 generation of oceanic plateaus and diffuse ocean island volcanism (2) plate bounded predominantly by ocean ridge systems, marked by island chains having no systematic age progression (3) plate bounded by predominantly by convergent margins, characterised by age progressive island chains. Intraplate volcanism on the Pacific plate shows a transition through the above styles, whereas the Izanagi/Kula and Farallon plates were always bounded by convergent margins and were likely characterised by the latter style of volcanism.

Keywords: Pacific basin, ocean island chains, oceanic plateaus, streaky mantle, propagating fractures

INTRODUCTION

Before the mantle plume model was widely invoked, intraplate volcanism in the Pacific basin (Fig. 1) was attributed to shallow mantle convection or propagating fractures induced by stresses acting on the plate (Shaw, 1969, 1973; Richter, 1973; Jackson et al., 1972; Jackson and Shaw, 1975; Bonatti and Harrison, 1976). One of the most advanced explanations was the model of Jackson and Shaw (1975) who noted that volcanism along the Hawaiian chain followed a series of right-lateral en echelon loci and suggested this reflected injection of magma into the lithosphere as the plate underwent rotation due to either changing forces along its boundaries, or variations in coupling between the lithosphere and asthenosphere. Similar loci were suggested to be present on other island chains of comparable age, but at the time the bathymetry of the ocean floor was poorly known and the concepts could not be further developed. With increasing popularity of the hotspot model, emphasis shifted to interpreting

3 the volcanic record as the result of mantle plume activity. However, few examples of intraplate volcanism have been shown to conform to the predictions of the plume model

(Anderson, 2005a). Only three examples in the Pacific basin (Easter-eastern Mid Pacific

Mountains, Louisville-Ontong Java, Marquesas-Hess/Shatsky Rises) have been suggested to fit the plume head-tail model (Clouard and Bonneville, 2001). Doubt is cast on such plateau- hotspot correlations from paleomagnetic evidence (Rissager et al., 2003) which suggests the

Ontong Java plateau formed 200 north of its inferred position in hotspot models, requiring either non-fixed hotspots or true polar wander. In the Cook-Austral-Marquesas, Marshall-

Gilbert, and Line Island chains, non-linear age progressions (Bonatti et al., 1977; Jarrard and

Clague, 1977; Turner and Jarrard, 1982; Schlanger et al., 1984; Duncan and Clague, 1985; Okal and Batiza, 1987; McNutt et al., 1997) have been a persistent problem, and have led to many ad hoc varieties of plume shape and plumbing arrangements (e.g. McNutt and Fischer, 1987;

Sleep, 1992). Even in the Hawaiian-Emperor chain, which has been considered the product of the strongest plume (Sleep, 1990), the volcanic record conflicts with most of the predicted features of the plume model. The volcanism lacks an associated plateau, eruption rates have increased rather than declined over time (Clague and Dalrymple, 1989), and the Emperor chain lacks a swell (Davies, 1992) which is one of the principal features used as evidence that hotspots are hot. Paleomagnetic evidence also indicates formation of the Emperor seamounts as much as 23 degrees north of the present site of volcanism (Tarduno and Cottrell, 1997;

Sager, 2002) such that any hotspot could only have been stationary for approximately half its history.

4 The commonly held view of Pacific intraplate volcanism with superplumes forming oceanic plateaus followed by plume tails generating ocean island chains (e.g. Larson 1991a,b; Tatsumi et al., 1998) has nonetheless been maintained on account of beliefs that other models could not explain the large volumes of melt, eruption rates, and basalt geochemistry (e.g. Mahoney,

1987, Tarduno et al., 1991; Coffin and Eldholm, 1993; Coffin and Gahagan, 1995; Neal et al.,

1997). However, the plume model has been shown to be inadequate for explaining the subsidence history (Ito and Clift, 1998) and lack of uplift associated with the Ontong Java plateau (Korenaga, 2005), which as the largest oceanic plateau, is the example supposedly most requiring a plume origin. In consideration of the lack of predictive ability of the plume model, uncertainties in the number of hotspots, and the failure of the model to account for many aspects of intraplate volcanism, the question should be whether the emphasis on the plume model has been justified and whether shallow-source models alone can offer a comprehensive explanation (Anderson, 1999; 2001; Favela and Anderson, 2000; Smith 2003a;

Natland and Winterer, 2005). The aim of this study is to examine concepts which could be included in a shallow-source model from present day intraplate volcanism, and then to investigate how such concepts could apply to earlier intraplate volcanism in the Panthalassan-

Pacific basin from a reconstruction of the volcanic record from the mid Paleozoic through

Recent. The results concur with, and build on the model of Natland and Winterer (2005) whereby volcanism on the Pacific plate is divided into three stages reflecting the tectonic setting of the developing plate, and illustrate how the history of the basin can be explained without mantle plumes.

5 CONCEPTS

Distinct Reservoirs versus Sampling Effects

The geochemical basis of the plume model is the belief that distinct isotopic signature in ocean island basalts (OIB) and mid ocean ridge basalts (MORB) requires derivation of these two types of magmatism from chemically distinct sources. In the most common version, the mantle has a layered structure, with a depleted reservoir which serves as the source of

MORB, overlying a layer of primitive composition (e.g. Hofmann, 1997). The depleted reservoir is usually equated with the upper mantle, and is considered to be the geochemical complement of the continental crust from calculations suggesting derivation of the latter from approximately one third of the mantle. Subducted oceanic crust is stored at some depth, usually the base of the upper mantle or core-mantle boundary, before rising as plumes to generate intraplate volcanism. Heterogeneity in OIB reflects variations in the proportions of recycled components, whereas heterogeneity in MORB reflects variable degrees of melt extraction from the depleted mantle (Gast, 1968) plus varying amounts of refertilisation by recycled oceanic crust (e.g. Zinder and Hart, 1986; Saunders et al., 1988). Recycling of oceanic crust into the depleted mantle has been considered the result of contamination with unmelted residues of plumes, with the entirety of the reservoir originating by this way in some models

(e.g. Morgan et al., 1995).

In other models, a large percentage of subducted oceanic crust (87% according to

Christensen and Hofmann, 1994) is recycled into the depleted mantle by convective stirring as

6 in the marble-cake mantle model of Allègre and Turcotte (1986). The latter hypothesis was based on interpretations of banding in Alpine peridotites as oceanic crust which had been stretched out and thinned by several orders of magnitude by mantle convection (Polvé and

Allègre, 1979). The marble-cake model was combined with the plume model by Allègre and

Turcotte (1986) because of a belief that noble gas isotope signatures, particularly high 3He/4He ratios in OIB, mandated a contribution from a primitive reservoir, although later work

(Hamelin and Allègre, 1988) demonstrated the isotopic suitability of marble-cake mantle itself as a source for intraplate volcanism. Other geochemical evidence, such as Pt-Os isotope systematics, was subsequently used to support the plume model (e.g. Brandon et al., 1998).

However, high 3He/4He ratios can result from a low abundance of the isotope 4He produced in the decay series of 238U, 235U and 232Th, rather than high abundances of the primordial isotope

3He (Anderson, 1998a,b). Correspondingly, Os isotope signatures in OIB can be interpreted as a pyroxenite rather than core-signature (Smith, 2003b), an interpretation which is supported by tungsten isotopic evidence which shows no core contribution in OIB (Scherstén et al., 2004). Derivation of MORB and OIB from a common reservoir has thus been reconsidered, as in the SUMA model of Meibom and Anderson (2003), where differences between these melt types reflect statistical sampling of a heterogeneities as in plum-pudding models (Zindler et al., 1984) for the derivation of alkaline seamount basalts in the ocean ridge environment.

The volume of depleted mantle must be greater in marble-cake/SUMA models than in standard plume models, because the presence of recycled material in the MORB-source would mean that the depleted mantle could not be the complement of the lithospheric crust.

7 Otherwise, the basic concept of mixing of recycled material with a depleted residue, is the same in plume, marble-cake and SUMA models. As illustrated by the Pb-Pb isotopic variation in MORB and OIB in figure 2, if mantle heterogeneity is described in terms of the components identified by Zindler and Hart (1986), mixing would concern introduction of high- m (HIMU) and enriched mantle (EM1, EM2) components into a depleted residue DMM

(Depleted MORB Mantle). The composition of recycled crust is invariant through time in such models. However, if declining thermal gradients through time have lead to a shift from subduction regimes dominated by slab melting to regimes dominated by slab dehydration, the composition of recycled components should have changed. Mixing between HIMU and EM- components would be appropriate for dehydrated slab material, but melted slabs would evolve to unradiogenic isotopic compositions similar to DMM. In the streaky mantle model of Smith (2005) a different picture of crustal recycling emerges, with the depleted mantle being a mixture of recycled dehydrated- and recycled-melted oceanic lithosphere. In such a structure, approximately equal proportions of dehydrated- and melted- recycled oceanic crust are involved in the generation of MORB, whereas recycled dehydrated oceanic crust predominates in the source of OIB. As in SUMA models, the difference between MORB and

OIB in the streaky mantle model is a result of different sampling processes under ocean ridges compared to intraplate regions. The common component between MORB and OIB, labelled

FOZO in plume models where it is equated with lower mantle (Hart et al., 1992) or young recycled basaltic crust (Stracke et al., 2005), is then recycled dehydrated oceanic crust comprising 97-99% basalt and 1-3% sediment (Smith, 2005). Mixing with the melted recycled crustal component explains the difference in slope between the MORB- and OIB- isotopic arrays in figure 2.

8 Additional Heterogeneity from Shallow-Level Contamination

Complexity arises in the geochemical modelling from uncertainties in the nature of the EM1 component. This signature is often associated with EM2 in DUPAL isotopic signatures characterised by higher 208,7Pb/204Pb than the Northern Hemisphere Reference Line of Hart

(1984), and high 87Sr/86Sr. The EM1 and EM2 components can also be seen in trace element signature as in figure 3 where the elements Th, Nb and Ce are used to represent large-ion lithophile elements (LILE), high field strength elements (HFSE), and light rare earth elements

(LREE), respectively, as they are immobile and their ratios are unaffected by fractionation of common mineral assemblages in oceanic basalts. Ocean island basalts carrying EM signatures tend to be enriched in LILE, and have high Th/Nb relative to Ce/Nb, variations in the latter resulting primarily from the greater incompatibility of Nb relative to Ce in a peridotite mineralogy on partial melting. In Pacific ocean ridge volcanism, EM1 has only been suggested to be present at 110N on the East Pacific Rise by Niu et al. (1999), who attributed it to plume material flowing over 5000 km through the asthenosphere from Hawaii. However, the signature is considered uncertain as while the basalts at 110N have high 207Pb/204Pb, they show no elevation in 208Pb/204Pb, and have lower Th/Nb than seamount basalts from elsewhere on the East Pacific Rise (Fig. 3). In Pacific island chains, EM1 is present in basalts in the region of the South Pacific Isotope and Thermal Anomaly (SOPITA; Staudigel et al., 1991), notably in the Pitcairn seamounts (Woodhead and Devy, 1993), and is an endmember in Hawaiian volcanism (e.g. Blichert-Toft et al., 1999) (Fig. 2). An EM1 component is also seen in the Pb

9 isotopic and trace element composition of oceanic plateaus, including Ontong Java, Manahiki and Shatsky Rise (Mahoney et al., 1993; 2005) (Figs. 2,3).

The most common interpretations are that EM1 represents pelagic sediment, metasomatised continental mantle, and lower continental crust (Lustrino and Dallai, 2003). If as suggested by

Weaver (1986), EM2 is turbidite, whilst EM1 is pelagic sediment, no additional causes of mantle heterogeneity to the crustal recycling processes discussed in the previous section are required. A relatively uniform sediment composition termed GLOSS (for Global average

Subductiong Sediment) was proposed by Plank and Langmuir (1998) to result from deposition of turbidites on pelagic sediments as oceanic crust approaches subduction zones. However, while such a model may be applicable to modern plate configurations, there is an increased liklihood of endmember compositions being subducted with ancient supercontinent plate configurations. Pelagic sediment has low U/Pb, Th/Pb, and if entrained in the convecting mantle would undergo little change in Pb isotopic composition. Mixing only a small percentage (~1%) of sediment with basaltic crust would be sufficient to generate the EM1 Pb isotopic signatures or the Th/Nb ratios found in OIB. However, on a global scale, the distribution of EM1 signatures is strongly provincial, being common in volcanism in the South

Atlantic and Indian Ocean (Fig. 2), where Hart (1984) and Castillo (1989) suggested it resulted from an equatorial belt of lower mantle upwelling (Fig. 1). However, DUPAL Pb isotopic compositions are also common in basalts from southeast Asia (Flower et al., 1998; Smith,

1998), the South China Sea (Chung and Sun, 1992), Sea of Japan (Tatsumoto and Nakamura,

1991), and Philippine Sea (Hickey-Vargas et al., 1995), suggesting the signature characterises

10 regions where Gondwanan lithosphere is preserved or has been delaminated into the shallow mantle (Smith and Lewis, 1999a; Chung et al., 2003).

A lithospheric origin for EM1 fits with the widespread occurrence of this signature in continental volcanism (e.g. Menzies, 1989). In western North America and southeast Asia,

EM1 Pb isotopic and trace element signatures are best displayed in potassic volcanism where phlogopite is a likely mineral in the source (e.g. Meen et al., 1989; Dudás, 1991; Liu et al.,

1994). EM1 Pb isotopic signatures have been found in phlogopite in the continental mantle

(Tatsumoto et al., 1992), and this mineral also has high Th/Nb and Ce/Nb ratios (Fig. 3). Both amphibole and phlogopite would be stable in the mantle wedge during subduction, with phlogopite most likely to survive the thermal rebound following detachment of the slab at the termination of subduction, although fluids derived from the breakdown of this mineral may also precipitate amphibole at shallower depths. The geochemical characteristics of lower continental crust (Doe and Zartman, 1979; Ben Othman et al., 1984) also match those of EM1

(Figs. 2,3), and a growing number of studies have suggested such material as a source component for intraplate volcanism (Escrig et al., 2004; Hanan et al., 2004; Anderson, 2005b;

Meyzen et al., 2005). Direct evidence for lower crust in the convecting mantle has come from marble-cake banding in the Ronda peridotite massif (Morishita et al., 2003), although trace element abundances in the bands were up to three orders of magnitude lower than for average continental crust.

11 Distribution of Source Material

In the plume model, geochemically distinct sources for intraplate volcanism are introduced vertically from depth in the mantle. In the SUMA/marble-cake/streaky-mantle models, such sources co-exist. Tomographic studies have indicated the presence of upwelling regions in the mantle under Africa and the Pacific basin (e.g. Li and Romanowicz, 1996; Ritsema, 2005). The location of such upwellings corresponds to the location of positive residual geoid anomalies, which have been correlated with the location of, and suggested to be caused by, hotspots/plumes (e.g. Duncan and Richards, 1991). Subsequent plume models (Davaille, 1999;

Courtillot et al., 2003) considered the upwellings to be superplumes which fed localised hotspots in the shallow mantle. However, in a plate model, the upwellings can be interpreted as convection cells in normal mantle where they would provide a mechanism by which shifts in the geochemical features of intraplate volcanism could occur without invoking plumes.

Alternatively, geochemically distinct domains could be introduced by lateral mantle flow.

Before introdution of the plume model and the hotspot refrence frame, plate motions could be modelled using the Antarctic plate as a reference because it is nearly surrounded by ridges and so can be regarded as stationary, while its position centred about the Earth’s rotation axis allows plate motions to be directly related to the x,y,z components of a triaxial Earth (Lewis and Smith, 1998). Initial modelling using the Antarctic plate as a reference indicated a net westward drift of plates at an average rate of 5 cm yr-1 (LePichon, 1968). The concept was re- introduced and described as a differential rotation of mantle and lithosphere by Doglioni

(1990), who invoked it as a cause of structural differences in arcs associated with east- and

12 west-dipping subduction zones. The mechanism was subsequently suggested as a means for lateral introduction of source materials for intraplate volcanism without invoking plumes

(Smith, 1993; Smith and Lewis 1999a,b; Doglioni et al., 2005), as it would also serve as a means whereby continental mantle or lower continental crust delaminated into the shallow mantle (e.g. Mahoney et al., 1989; Lustrino, 2005) would become displaced from the site of rifting or delamination as the lithosphere lags behind the mantle as the Earth rotates.

The relationship between plate motions relative to a fixed Antarctic plate (RAR: rotation axis reference) and the hotspot reference frame are illustrated in figure 4. A net westward drift is observed in the hotspot frame, but is reduced reduced to 1.7 cm yr-1 relative to averge

“hotspots” (Ricard et al., 1991). The difference in velocities can be explained by the sources of intraplate volcanism lying in a zone of shearing between the mesosphere and lithosphere, i.e. residing in the asthenosphere with no deep roots in the mantle (Smith and Lewis, 1999a,b;

Doglioni et al., 2005). In the plume model, plates move over, and exert a drag on, a near-static mantle, which is required to avoid entrainment of plume conduits by convection. In the differential rotation model, the mantle exerts a drag on the lithosphere as the Earth rotates.

The drag is attenuated by the reduction in viscosity through the asthenosphere, and can either supplement or oppose the action of boundary forces acting on plates.

Tapping of Source Material

Other key concepts in plate models are the mechanisms by which melt is generated and how such melting could be linked with the formation of oceanic plateaus and island chains. In the

13 plume model, the emphasis is on melting as a result of thermal anomalies. However, many petrological studies (e.g. Oxburgh, 1964; Flower et al., 1975; Bonatti, 1990; Francis and

Ludden, 1995; Green and Falloon, 1998; 2005; Presnall and Gudfinnsson, 2005) have argued that the sources of ocean island intraplate volcanism are volatile-rich and contain minerals such as amphibole and phlogopite. The low melting point of such minerals (Fig. 5) poses a paradox for hotspot models, as material subducted to the temperature and pressure conditions postulated for plume sources would be expected to have been purged of volatiles. Volatiles must then be incorporated into plumes at shallow depth in the mantle (Wyllie, 1988), in effect invoking material which could be interpreted as a source in its own right, in order to maintain the plume model. In other models (e.g. Cousens, 1996) volatiles are released from subducting slabs in the transition region, and migrate to form minerals such as amphibole and phlogopite at shallow level. However, such slabs will already have lost a large proportion of their volatiles in the subduction zone environment. The greatest accumulations of volatile-bearing minerals are thus expected in the hanging-wall of the mantle wedge at convergent margins, where release of volatiles from the slab is expected to form pargasitic amphibole, phlogopite and K-richterite with increasing depth (e.g. Tatsumi, 1989; Ionov and Hofmann, 1995;

Trønnes, 2002). These assemblages need not be depleted in HFSE nor impart arc-like geochemical signatures, as although HFSE are expected to be largely retained in the subducting slab, a fraction will be mobilised into fluids/melts derived from the slab and become incorporated into phases such as orthopyroxene and amphibole immediately above the slab

(Kelemen et al., 1992; Ionov and Hofmann, 1995).

14 The solidii for pargasite and phlogopite lie to higher temperature than the off-axis and shield geotherms (Fig. 5), and these minerals are expected to survive for long time periods in the continental mantle or shallow oceanic mantle. However, if entrained into upwelling beneath an ocean ridge system, volatile-bearing mineral assemblages are expected to undergo rapid and extensive melting. The similar Th/Nb, Ce/Nb ratios of Ontong Java basalts to mantle amphiboles (Fig. 3), suggests formation of this plateau may be due to entrainment of former mantle wedge material into an ocean ridge system. Estimated temperatures for the formation of highly undersaturated alkali basalt types in ocean island chains (Green and Falloon, 1998) are only 50-1000C greater than the average intraplate geotherm. Possible mechanisms for increasing the intraplate geotherm to the temperature-pressure conditions estimated for such lavas includes shear melting and small-scale mantle convection. As pointed out by Richter

(1973), upwelling associated with transverse rolls would not be stable with fast plate velocities and would be transformed into longitudinal rolls with the axes of convection perpendicular to the ridge system. Support for the existence of convective rolls has come from spectral analysis of bathymetry and heatflow in the region between the East Pacific Rise and

Hawaii (Furlong and Chapman, 1978), and from tomographic studies around ocean island chains including Hawaii (Katzman et al., 1997). However, convective rolls should lead to the formation of “hotlines” of volcanism without linear age progressions (Bonatti and Harrison,

1976), and must be combined with other mechanisms to yield a comprehensive model for ocean island volcanism.

Lateral introduction of source material in the differential rotation model would be one mechanism for explaining an apparent propagation of volcanism, and also has implications for

15 shear melting. The combination of drag induced by the mesosphere and eastward movement of plates by boundary forces (Fig. 4c), is postulated to lead to high plate velocities but low shear stresses (Smith and Lewis, 1999b). In contrast, movement of plates westwards by boundary forces sets up a counterflow regime with a stationary layer characterised by high shear stress but low velocity, as the asthenosphere becomes sheared by drag from the plate opposing drag from the mesosphere (Smith and Lewis, 1999a,b). For lithosphere 100 km thick, underlain by

200 km of asthenosphere, the stationary layer would lie at approximatly 130 km depth (Fig.

4d). Release of melt from such a layer would appear as a group of hotspots, fixed relative to each other, but moving in unison relative to intraplate volcanism on another plate.

Differentials between lithosphere and mesosphere of 10 cm y-1 for counterflow regimes in the differential rotation model (Fig. 4d), would produce a temperature rise of 600C if the viscosity of the material being sheared was 2x1020 Pa s, as would be reasonable for continental mantle/lower crust delaminated into the mantle (Doglioni et al., 2005). The effect would be similar to in the thermal feedback mechanism of Shaw (1973), whereby shearing causes melting, which facilitates further slip and hence further melting. Shear melting would therefore be sufficient to raise an average intraplate geotherm to conditions for the generation of olivine leucitites and olivine melilitite melts, whereas a combination of rolls and shear heating may raise the geotherm to the temperature-pressure conditions for the formation of other alkali basalt types (Fig. 5).

Lithospheric architecture is an important consideration in plate models (Anderson, 2001), and as noted by Natland and Winterer (2005), the range of intraplate features suggests a range effects. Melt focussing at ridge-ridge-ridge triple junctions has been suggested to lead to the

16 generation of enhanced thicknesses of oceanic crust (Georgen and Lin, 2002). Abandoned transform faults, troughs and ridges will all constitute lines of weakness which may potentially be reactivated (Natland and Winterer, 2005). Seamounts along the Kodiak-Bowie chain have been considered to have formed as a result of a leaky transform or by ocean ridge processes (Dalrymple et al., 1987; Allen et al., 1993), and a relationship with the Eltanin fracture zone has been suggested for the Louisville chain (Hayes and Ewing, 1971; Larson and

Chase, 1972; Smoot, 1999). At convergent margins stresses may be generated from variations in dip or strike of the subducting slab (Solomon and Sleep, 1974). The relationship of volcanism to convergent margin geometry is well illustrated in the eastern Pacific (Favela and

Anderson, 2000; Smith 2003a): the most prominent intraplate line on the Nazca plate is the

Sala y Gomez chain which extends from the Easter microplate (Fig. 1). Although this chain was originally interpreted as a hotline (Bonatti and Harrison, 1976), subsequent studies sought to fit a plume model to the age progression (O´Connor et al., 1995). However, the line extrapolates to the boundary between the northern and central Chile subduction zones where slab dip decreases from 300 to 150 (Isacks and Barazangi, 1977; Uyeda and Kanamori, 1979).

Similarly, the Juan de Fernandez ridge extrapolates to the boundary between the central and southern Chilean subduction zones where slab dip increases to 300, suggesting the lines are related to plate tearing induced by convergent margin geometry (Smith, 2003a). A further example of control by lithopsheric structure is volcanism along the Samoan chain, which lies in a region of deformation of the Pacific plate opposite where the margin changes from subduction along the Tonga trench to a transform boundary along the Fijian margin (Natland and Winterer, 2005).

17 Lithosphere will also undergo flexure over topgraphic anomalies such as the Earth’s equatorial bulge, and be subject to fracturing from thermal stresses (Turcotte and Oxburgh,

1973, 1976). The importance of the latter is illustrated by nearly half of proposed hotspots occurring on lithosphere less than 30 million years old (Favela and Anderson, 2000). Cooling of the plate as it moves away from riges has been suggested to lead to the formation of Puka-

Puka ridges which extend hundreds of kilometres from the ridge (Sandwell et al., 1995;

Natland and Winterer, 2005). The fracturing associated with the formation of Puka-Puka ridges is orthogonal to the direction of seafloor spreading, and is thus different from suggested in the propagating fracture models of Turcotte and Oxburgh (1973, 1976) for ocean island chains. The latter authors suggested that brittle fracture would occur near the surface of the plate, whereas plastic fracture would occur at depth, with the mechanisms producing fractures at 450 and 350 to the direction of tension, respectively. Volcanism along the Hawaiian, Cook-

Austral, Tuamotu-Pitcairn, and Kodiak-Bowie chains lies between 310 and 420 to spreading on the East Pacific Rise and Juan de Fuca ridges, fitting the propagating fracture model. The fracture model appeared to be negated by volcanism along the Marquesas chain which lies at

600 to the East Pacific Rise (Clague and Dalrymple, 1989). However, such objections are negated by the interpretation of Natland and Winterer (2005) of the Marquesas chain as a series of en-echelon Puka-Puka ridges. Plate fracturing as a result of lithospheric loading (e.g.

Walcott, 1976; Hieronymus and Bercovici, 2000) is also of potential importance for accounting for progressions of volcanism. Stress may be imposed on the plate from the weight of the volcanic edifice, or by the development of a high density root of cumulates. The latter was suggested for the Hawaiian chain by Natland and Winterer (2005) as a modified version of the gravitational anchor mechanism of Shaw and Jackson (1973). In such models, sinking of a

18 root of cumulates distorts the asthenospheric flow regime at the base of the plate, leading to upwelling currents which cause propagation of the chain.

APPLICATION TO THE VOLCANIC RECORD SINCE THE MIDDLE PALEOZOIC

Carboniferous to Early Jurassic: The Panthalassan Stage

A record of intraplate volcanism in the Panthalassan Ocean which existed before formation of the Pacific plate, can be found in accreted terranes around the margin of the Pacific basin

(Fig. 1). Topographic features such as plateaus, ocean island chains and seamounts were originally considered candidates for accretion (e.g. Ben-Avraham et al., 1981). However, the accretion of oceanic plateaus has been the subject of debate, with opinion ranging from accretion (Saunders et al., 1996), partial accretion (Kimura et al., 1994), to subduction of such features (Cloos, 1993). The latter argument, based on buoyancy considerations, predicts only plateaus the size of the Ontong Java plateau would be large enough to survive. Nonetheless, conjugates of the Shatsky Rise have been suggested to include the Vatyn terrane (also called

Olutorovsk and Olytorka-Kamchatka terranes) of northeast Russia (Zonenshain et al., 1987), the Sorachi terrane of Japan, and the Caribbean plateau (Kimura et al., 1994). However, the

Caribbean plateau is too young and/or formed in the Central American region (e.g. Donnelly,

1989; James, 2006), while in the correlation by Zonenshain et al. (1987), the age of the

Shatsky Rise was underestimated by 15-35 million years. Basaltic volcanism in the Sorachi terrane occurred at around 150 Ma from paleontological evidence, which would require a very high rate of plate movement of 25 cm yr-1 for even the youngest accretion age of 119 Ma for it

19 to be the conjugate of Shatsky Rise (Kimura et al., 1994). Identification of the belt as a plateau remnant is also debatable. The pre-subduction volume of the Sorachi plateau was estimated by

Kimura et al. (1994) at 26 x 106 km3 which would be comparable to the Mid Pacific

Mountains (21.5 x 106 km3; Schubert and Sandwell, 1989), but the volume of accreted material in the terrane was only estimated at 0.54 x 106 km3, which would be equivalent to a large seamount chain. Basalts from the Sorachi terrane also have low Th/Nb and Ce/Nb, similar to non-DUPAL seamounts rather than oceanic plateaus (Fig. 3). In contrast, shearing and obduction of at least the upper part of seamounts/ocean islands has been observed at consuming plate margins (e.g. Taira et al., 1989), and for such features the terrane record can be considered more complete. Because the fragments are mixed with sediments of the fore-arc wedge, distinction between large island chains and axial seamounts depends largely on the presence of limestones which could be interpreted as capping reefs. Ocean island chains rising above the calcite compensation depth would be expected to capped by such features, whereas seamounts associated with ocean ridge activity would be too low. Both types of edifice are present in the terrane record, and an advantage of the accretion of island chains is the faunal evidence which may also be preserved.

The record of seamount/ocean island volcanism in accreted terranes in the southeastern

Pacific is quite sparse, with only a few fragments tentatively identified in Columbia (Mégard,

1989). In the southwestern Pacific, much of New Zealand is composed of accreted oceanic material, with seamounts/ocean islands prominent in the Waipapa terrane (Spörli and Ballance,

1989). The principal record of terrane accretion is found along the northwestern and northeastern margins of the basin (Fig. 1), reflecting plate motions and the source of many

20 allochthonous fragments in the southwest of the basin. In Asia, three terrane belts of subduction assemblages may be recognised: (1) A Jurassic belt, including the Bazdahl, Taukha and Yamaguchi terranes, extending from northeastern Russia, through Sikhote-Alin, Honshu, the Ryukuyus, Taiwan, and along the continental shelf off Hainan, to Borneo (Howell, et al.,

1985). The Jurassic belt contains Carboniferous to middle Jurassic seamounts and ocean islands which accreted during the assembly of Asia, and is characterised by Tethyan faunas

(Maruyama et al., 1989; Nockleberg et al., 2000). (2) A Cretaceous belt, including the Yuli,

Hidaka, Sorachi, Tokoro, Anvina, Nabilsky, and Vatyn terranes, extending from Borneo, through Taiwan, Hokkaido and Sakhalin, through eastern Kamchatka. The Cretaceous belt contains material accreted from the Izanagi plate before and after formation of the Pacific plate

(Isozaki et al., 1990). (3) A Cenozoic belt containing material accreted from the Kula/Farallon plates (Kronotskiy terrane), occurring in eastern Kamchatka (Watson and Fujita, 1985;

Zonenshain et al., 1990). In western North America, Carboniferous-Permian seamounts and/or plateau remnants are found in the Cache Creek terrane of the Canadian Cordillera and its equivalent, the Baker (or Blue Mountains) terrane of the northwestern USA, which were accreted in the Early Jurassic (Monger, 1977; Silberling et al., 1987; Nelson and Mihalynuk,

1993; Tardy et al., 2001). An important feature of the Cache Creek terrane is the presence of equatorial Tethyan faunas (Monger, 1977; Cordey and Schiarizza, 1993; Orchard et al., 2001).

Transport of the Cache Creek terrane across the Pacific basin requires high plate velocities, and suggests a Panthalassan plate configuration which was the opposite of that today, with ocean ridges in the west of the basin and a large plate (Farallon) subducting in the east of the basin (Fig. 6a). Potential younger examples of intraplate volcanism are found in the Chugach terrane of Alaska (Nockleberg et al., 2000), Bridge River terrane of British Columbia (Cordey

21 and Schiarizza, 1993), and Guerrero terrane of Mexico (Ortiz-Hernández and Martinez-

Reyes, 1993), although considering the amount of ocean floor which must have been consumed, the record of seamounts/ocean islands accreted in the east of the basin is sparse compared to the record preserved along the western Pacific rim.

The faunal record also indicates the transport of several Paleozoic microcontinents/crustal blocks from eastern Gondwana to Asia and North America. These include the Insular terrane of western Canada (Gehrels and Saleeby, 1987; Coney, 1990), microcontinents (e.g. Reed

Bank) now bordering the South China Sea, the Kurosegawa and Abukuma-Kitakami belts of

Japan, and the Tahin terrane of Sikhote-Alin (Howell, et al., 1985; Maruyama et al., 1989;

Taira et al., 1989; Metcalfe, 1996), all of which were accreted in the Cretaceous. The Sea of

Okhotsk has also been postulated to have been underlain by a terrane, Okhotia (distinct from the Okhotsk terrane of Nockleberg et al., 2000), which accretion trajectories suggest may have derived from a similar position to the microcontinents accreted to Asia (Zonenshain et al.,

1987). Okhotia accreted to northeastern Russia in the Late Cretaceous to Paleocene, and has been suggested as an oceanic plateau (Watson and Fujita, 1985). However, as its size is only a tenth that of the Ontong Java plateau, it should have been subducted if comprised of oceanic material, and may therefore include crustal material of Gondwanan origin. Plate reconstructions have also placed the North China block to the east of Australia rather than along the southern margin of Tethys in the Paleozoic (Li et al., 1996; Smith et al., 1997; Li,

1998). The Omolon block (also known as the Koryakia terrane: Zonenshain et al., 1987) of northeastern Russia was also postulated to have have been derived from northeastern

Gondwana (Maruyama et al., 1989), although more recent correlations have suggested this

22 block has an autochtonous origin (Nockleberg et al., 2000). The implication is nonetheless, that a substantial continental mass “Pacifica” existed in the Pacific to the east of Gondwana, before it was fragmented in the mid Paleozoic (Fig. 6a)

In the differential rotation model, the continental fragmentation event becomes the cause of geochemical heterogeneity now in the south-central Pacific mantle (Smith and Lewis, 1999a).

Loss of more than 120 km of continental mantle from under eastern North China was proposed from the occurence of continental basalts derived from shallow asthenospheric sources in the Late Cenozoic, near where kimberlites had eruptioned in the Early Paleozoic

(Menzies et al., 1993). The loss of such material has generally been assumed to take place by back-arc convection in the Late Mesozoic after the accretion of Asia (Griffin et al., 1998;

Menzies and Xu, 1998; Zhang et al., 1998), under which circumstances the eroded material would have been blocked from entering the Pacific mantle by westwards-directed subduction under Asia. However, an alternative view is that the loss of lithosphere took place either as a result of convection during breakup of eastern Gondwana during the mid Paleozoic, or by delamination on continental collision in the Permian and Triassic (Smith and Lewis, 1999a)

(Figs. 6a,b). If continental mantle or lower crustal material eroded during such events remained in the shallow mantle, westward drift of the lithosphere would have introduced the material into the east-central Pacific basin by the time of formation of the Pacific plate (Fig. 6c). The extensive record of arc terranes in the northwest of the Panthalassan-Pacific basin (Nockleberg et al., 2000) suggests shallow enriched mantle domains could also have been created by subduction processes in the Late Paleozoic-Early Mesozoic.

23 Jurassic to mid Cretaceous: The Intra-Oceanic Evolution of the Pacific Plate

The Pacific plate originated in the central paleo-Pacific basin in the Middle Jurassic, where it was surrounded by the Izanagi, Farallon, and Phoenix plates (Bartolini and Larson, 2001) (Fig.

6c). Although the western part of the plate has now been lost to subduction, it is likely that the plate had a rhombic outline bounded by ocean ridge systems. The terrane record indicates intraplate volcanism at the time was confined largely to the Izangi plate (Fig. 6c). The volcanic record on the Pacific plate from 170 Ma to approximately 90 Ma includes oceanic plateaus and seamounts (Clouard and Bonneville, 2005; Natland and Winterer, 2005). The earliest magmatism includes the Shatsky Rise and Mid Pacific Mountains plateaus, and seamount volcanism in the Marshall-Gilbert group (Fig. 6d). The Shatsky Rise formed in conjunction with a triple junction between the Pacific, Izanagi and Farallon plates (Hilde et al., 1977;

Nakanishi et al., 1999; Sager, 2005), which would fit the melt focusing model of Georgen and

Lin (2002). Although rapid eruption has been suggested in order to fit a plume model (e.g.

Sager and Han, 1993), the position of the Shatsky Rise relative to magnetic lineations

(Nakanishi et al., 1999), indicates formation over a time period of 16 million years between

143 and 127 Ma. The Mid Pacific Mountains formed at a junction between the Pacific-

Phoenix-Farallon plates, but has a more complex morphology (Natland and Winterer, 2005).

Radiometric ages for the Mid Pacific Mountains range from 129 to 83 Ma (Winterer and

Sager, 1995) indicating a protracted evolution, and as only the central and eastern parts have been dated, volcanism may have overlapped formation of the Shatsky Rise.

24 Formation of the Ontong Java plateau in the late Early Cretaceous also likely also occurred in conjunction with a ridge system (Hussong et al., 1979). The position depicted for this plateau in figure 6e is based on the paleomagnetic latitude of 240S reported by Riisager et al.

(2003) for the northern tip of the plateau. Lineations on the plateau have been suggested to correspond to both southwest-northeast (Winterer and Nakanishi, 1995) and northwest- southeast (Neal et al., 1997) trending ridge axes. Age data are available for three drill sites, and indicate episodes of volcanism from 124 to 121 Ma and at 90 Ma (Mahoney et al., 1993).

However, as none of the sites penetrated more than 150 metres of basalt, the existence of any age progression is unknown (Sager and Han, 1993), and arguments for rapid eruption (e.g.

Tarduno et al., 1991; Coffin and Gahagan, 1995) rely on correlations with sites in nearby basins and sections on Malatia and Santa Isabel islands. If the plateau is considered to have formed away from a ridge, or with the aid of a ridge system orientated northeast-southwest, the wide spacing between the sites would be consistent with eruption over a time period of only a few million years. However, if the plateau formed in conjunction with a ridge system orientated northwest-southeast, which would provide a better fit with the orientation of the

Pacific-Phoenix ridge in figure 6e, the drill sites where similar ages have been found and also the sections on Malatia and Santa Isabel Island, would all lie close to the ridge axis, in which case it would not be surprising to find similar ages. The terrane accretion record in New

Zealand shows a hiatus between 130 and 100 Ma which led Spörli and Ballance (1989) to suggest subduction along the southwest margin of the Pacific basin became choked by the arrival of arrival of a microcontinent in the middle Cretaceous. Spreading on the Pacific-

Phoenix ridge may have become very slow, and it is proposed that entrainment of a large streak of fertile mantle, such as an amphibole-bearing assemblage suggested in the previous

25 section, into the ridge upwelling led to generation of large volumes of melt which ponded over the ridge system to generate the plateau. The final phase of plateau volcanism involved formation of the Manahiki plateau and the Hess Rise which have ages of 123-110 and 110-100

Ma, respectively (Mahoney, 1987). The Hess Rise formed along a fracture on the Pacific-

Farallon ridge before it jumped eastward (Windom et al., 1981; Rea and Dixon, 1983; Norton, this volume) (Fig. 6e). The tectonic setting in which the Manahiki plateau formed is unknown, but it has been depicted as forming at a triple junction as for the other plateaus in figure 6e.

Contemporaneous with the formation of the plateaus was seamount volcanism in the

Marcus-Wake, Magellan, and northern Marshall-Gilbert groups (Fig. 6e). Linear east-west, and northwest-southeast trending age progressions were suggested in these islands by

Koppers et al. (2001); however, the volcanism lacks a consistent paleolatitudinal pattern as would be expected from a plume origin (Sager et al., 1993), and other studies (Natland and

Winterer, 2005) have argued that there is no systematic age distribution. The debate is compounded by only a partial record being preserved, as volcanism on the western half of the

Pacific plate at this time would have been lost to subduction. The current position of the

Marcus-Wake, Marshall, and the slightly younger, Japanese seamounts (Fig. 1), indicates such volcanism would have been accreted along southwest Japan and the Ryukus, and it is suggested here that ocean islands immediately west of the Japanese seamounts in the

Cretaceous may now be in the Shimanto belt of southeast Japan. The suggestion of linear age progressions in the volcanism conflicts with other plume models in which the Pacific plate is considered to have been stationary as a result of being pinned by plumes in the Early

Cretaceous (Tarduno and Sager, 1995). In a shallow-source model, the Pacific plate would

26 have been stationary as a result of being bounded by ridges. Comparable rates of convergence around the margins of the Pacific basin (e.g. Rea and Duncan, 1986) suggest spreading rates were similar on the ridges, and the diffuse distribution of ocean island volcanism may have resulted from the lack of any well defined convection regime in the shallow mantle (Natland and Winterer, 2005). Membrane stresses may also have been significant as correcting to a paleomagnetic position for the Ontong Java plateau, the volcanism can be seen to have occurred in the equatorial region (Fig. 6e). The location of volcanism within the centre of the

Pacific plate may also indicate a role for thermal stresses, as at the onset of volcanism in the

Marshall Islands, the Pacific plate would have been 30 million years old. The Pacific plate may also be envisaged as subject to incipient northwest-southeast and east-northeast - west- southwest fracturing (dotted lines; Figs. 6d,e) between the Izanagi-Pacific/Pacific-eastern

Phoenix, and Farallon-Pacific/Pacific-western Phoenix ridge pairs (Smith, 2003a). The volcanism may thus have followed lines of weakness from stressing of the plate, but appears random because of interactions between the two competing stress directions and because only a partial record is now preserved.

Late Cretaceous to mid Eocene: A Subducting Pacific Plate Bounded by Ocean Ridges

The Late Cretaceous was a time of global plate reorganisation, including events such as the separation of South America from Africa (Favela and Anderson, 2000). In the Pacific basin, the Farallon plate split into northern and southern halves along the Mendocino fracture zone, and the Kula plate separated from the Farallon North plate (Woods and Davies, 1982; Rea and

Dixon, 1983) (Fig. 6f). Volcanism in the Marshall-Gilbert islands became more linear and

27 progressed southwards, and began in the Line Islands (Fig. 6f). A second phase of volcanism also occurred on the Ontong Java plateau, forming the eastern lobe around 90 Ma (Neal et al.,

1997). As in the Early Cretaceous, the majority of volcanism occurred in the equatorial belt. In hotspot models, ocean island chains prior to 83 Ma are used to infer plate motion to the northwest, with the direction of Pacific plate motion changing to north-northwest after 82 Ma to satisfy the trend of the Emperor chain (e.g. Gordon and Jurdy, 1986; Coney, 1990;

Lithgow-Bertelloni and Richards, 1998). While a change in plate motion might be attributed to northward subduction of the Kula plate faster than spreading on the Pacific-Kula ridge imparting the motion of the Kula plate to the Pacific plate (e.g. Lonsdale, 1988a), other ocean island chains including the Musicians, Line Island, and Marshall-Gilbert islands, show little or no change in strike at 82 Ma. Plate motion then has to revert back to northwesterly in the mid

Eocene in hotspot models to account for the bend between the Hawaiian and Emperor chains, which has been estimated between 43 and 47 Ma (Sharp and Clague, 1999). The Louisville chain also shows a bend at 46 Ma, but it is more subdued than in the Hawaiian-Emperor chains (Figs. 7,8). The mid Eocene change in Pacific plate motion was originally attributed to a plate re-organisation caused by the closure of Tethys and the collision of India with Asia

(Dalrymple and Clague, 1976; Clague and Dalrymple, 1989). However, other studies have argued the Tethyan events would have had little effect on motion of the Pacific plate

(Richards and Lithgow-Bertellini, 1996). Instead, no significant change in plate direction in the mid Eocene is supported by a lack of corresponding tectonic events around the basin margin

(Herron, 1972; Norton, 1995), continuity of magnetic lineations along the Pacific-Farallon ridge (Foulger and Anderson, 2005), and the consideration that slab pull and ridge push forces

28 are unlikely to change rapidly enough to cause changes in plate direction (Favela and

Anderson, 2000).

Plate motion paradoxes are avoided in the plate model, as changes in the stress field may occur independently of plate motion (Jackson and Shaw, 1975). The shift to quasi-linear island chains, albeit marked by considerable splay in features and non-linear age progressions, suggests the development of a preferred principal stress direction, possibly resulting from the onset of subduction of the Pacific plate beneath the continental margin of Asia (Natland and

Winterer, 2005). Continued expansion of the Pacific plate from the Early Cretaceous reconstruction in figure 6e, would have resulted in the Izanagi-Pacific ridge colliding with, and moving northwards along the margin of Asia as shown in figure 6f. Correspondingly, the

Pacific-Phoenix ridge may have collided with continental blocks in the southwest of the basin, and in several reconstructions (Zonenshain et al., 1987; Maruyama et al., 1989; Coney, 1990;

Smith, 2003a), an ocean ridge is depicted close to the northeast margin of Australia. Other models (Osazawa, 1992; Hall, 1996) have considered the Pacific plate to have been separated from Asia until the Early Eocene by the North New Guinea plate, which was invoked from the tectonics of the Philippine plate (Seno and Maruyama, 1984). A combination of these models is suggested in figure 6f, with the Pacific-North New Guinea ridge equated with the

Pacific-Phoenix ridge, and the Pacific plate subducting along the margin of Asia. However, irrespective of the ocean ridge configuration chosen, the length of any convergent margin would only have been 10-20% of the total margin of the Pacific plate, compared to approximately 40% for the present day Pacific plate. The stress regime of the Pacific plate in the Late Cretaceous would therefore be expected to be quite different to that today, and is

29 proposed to have been dominated by the ocean ridges rather than by the convergent margin.

The transition to a preferred stress direction (SH; figure 6f) in the Late Cretaceous may reflect faster spreading on the Izanagi-Pacific and Antarctic-Pacific ocean ridges compared to on the

Pacific-North New Guinea and Pacific-Farallon ridges. The north-northwest trend followed by volcanism in the Line Islands, Marshall-Gilberts and Louisville chain may thus have resulted from incipient fracturing of the Pacific plate between the Izanagi-Pacific and Antarctic-Pacific ridges, irrespective of the overall direction of plate movement, which is not constrained in the plate model.

Volcanism along the Emperor chain lies along a more northerly trend of N110W, and the origin of this chain may be related to plate re-organisations in the northern Pacific basin around 82 Ma when the Kula plate separated from the Farallon North plate (Smith, 2003a;

Norton, this volume). Paleomagnetic evidence has indicated Emperor volcanism occurred significantly further north than the present location of Hawaii, leading to speculation that the chain was formed by an unstable migrating plume (Norton, 1995; Tarduno and Cottrell, 1997;

Sager, 2002). Dating of Detroit seamount at the northern tip of the Emperor chain (Keller et al., 1995) suggests an age in excess of 81 Ma for Meiji seamount which represents the oldest volcanism attributable to the Hawaiian magma-source. The latitude for Meiji reconstructed in figure 6f is based on the paleomagnetic position of 430+130/-80 for Detroit seamount reported by Sager (2002), and lies close to the position of a triple junction depicted on the Pacific-

Farallon ridge at this time in the reconstruction by Rea and Dixon (1983). The proximity of initial Emperor magmatism to the Kula-Pacific ridge has also been suggested from plate reconstructions (Mamerickx and Sharman, 1988), guyot morphology (Caplan-Auerbach et al.,

30 2000), and the geochemical signatures of the oldest lavas in the Emperor chain (Keller et al.,

2000; Regelous et al., 2003). Meiji seamount may thus have been generated in a triple junction setting between the Pacific, Izanagi and Farallon North plates (Smith, 2003a), with volcanism interrupted by a jump of the Kula-Farallon-Pacific triple junction to immediately north of

Hess Rise (see arrow; Fig. 6f) (Norton, this volume), such that the edifice did not reach the size attained by other oceanic plateaus.

The central location of the Emperor chain in the north Pacific basin, and its proximity to the

Kula plate which was subducting to the northwest under Asia and to the northeast under western North America, allows interpretation of the chain as the product of a propagating fracture generated as a result of Kula-Pacific ridge geometry tearing the Pacific plate. Left- lateral offsets noted between loci along the Emperor chain by Jackson and Shaw (1975) can be attributed to clockwise rotations of the Pacific plate in response to events on the Kula plate to the north. Following the model of Byrne (1979), the cessation of activity on the Kula-Pacific ridge has generally been correlated with capture of the Bering Sea plate and initiation of the

Aleutian arc around 55-50 Ma (e.g. Norton, 1995). Offsets along the Emperor chain between

Ojin and Koko seamounts which have ages of 55 and 49 Ma, respectively (Clague and

Dalrymple, 1989), can then be attributed to changes in the stress field caused by plate interactions along the Aleutian margin (Fig. 7). Intraplate volcanism also occurred on the Kula plate as indicated by Late Cretaceous to Oligocene alkali basalts in the Kronotskiy terrane

(Watson and Fujita, 1985). The Kronotskiy terrane now lies at the same latitude as Meiji but accreted around 10 Ma (Watson and Fujita, 1985), and hence must have lain approximately

900 km to the west of Meiji from current plate velocities and convergence trends (Fig. 6f).

31 The Kronotskiy terrane might therefore represent an equivalent of the Emperor chain, or by analogy with the present day Nazca plate, fracturing as a result of plate tearing at the Aleutian margin.

Mid Eocene to Recent: The Pacific Plate Bounded by Convergent Margins

The Hawaiin chain has been considered as strong evidence for the existence of plumes (e.g.

Sleep, 1990; 1992; DePaolo et al., 2001), and it critical for non-plume models to address why volcanism should have been initiated along this chain. Options are to attibute the chain to mantle processes or to explore how and why the stress field should have changed at the time of the Hawaiian-Emperor bend. Explanations do not need to involve plumes, and Doglioni et al. (2005) have shown how the Hawaiian-Emperor chains could be explained by melting of a streak of fertile mantle, whose shape is the mirror image of the island chains, and which moves eastwards as a result of plate lag, with melts released as a result of lithospheric fracturing. A problem for the model, however, is that basalts along the Emperor chain are more MORB-like

(Keller et al., 2000; Regelous et al., 2003) than along the Hawaiian chain, such that any streak could not be homogeneous. Other models have noted correlations between the intiation of volcanism along the Hawaiian chain and the demise of spreading on the Kula-Pacific ridge

(Smith, 2003a; Natland and Winterer, 2005). Activity on the ridge may have ceased as late as

42 Ma from magnetic anomalies between the northern Emperor seamounts and the Aleutian trench (Lonsdale, 1988a), thereby overlapping the younger 43 Ma age for the Hawaiian-

Emperor bend. However, the anti-clockwise plate rotation suggested by the right-lateral loci of the Hawaiian volcanism suggests volcanism responded to changes in the stress field originating

32 in the south of the Pacific basin (Jackson and Shaw, 1975), and it is pertinent to consider what event in this region corresponded to the transition from Emperor to Hawaiian volcanism. Near contemporaneous with the demise of the Kula ridge, was the cessation of spreading on the

North New Guinea-Pacific ridge bounding the Caroline plate, as a result of collision of the ridge with the Izu-Bonin-Mariana arc (Hall, 1996) (Fig. 6h). The timing of this event is dated around 45 Ma from the earliest boninitic magmatism along the arc (Casey and Dewey, 1984;

Stern and Bloomer, 1992). The collision resulted in subduction of the Pacific plate beneath the

Philippine plate, and was suggested by Seno and Maruyama (1984) as the cause of the supposed change in motion of the Pacific plate at 43 Ma. However, in the stress field model, no change in plate motion is required. Instead, the stress field in the west of the Pacific plate is postulated to have changed as a result of the minimum horizontal shear stress (sh; Fig. 6h), changing from compressional to extensional as a result of the transition from the plate being bounded by the Kula-Pacific and North New Hebrides-Pacific spreading ridges, to being subducted under the Aleutian and Izu-Bonin-Marianas arcs. It is further postulated that fracturing started at the tip of the Emperor chain as a result of lithospheric loading causing localised weakening the plate south of Koko seamount, and then as the Hawaiian chain lies at

340 to the Pacific-Farallon ridge (Fig. 6i), propagated toward the East Pacific Rise as a result of plastic failure of the lithosphere as in the model of Turcotte and Oxburgh (1973; 1978).

Volcanism along the Hawaiian chain is sparse until Midway island at 28 Ma, and is then followed by a hiatus around 25 Ma where the chain crosses the Murray fracture zone

(Norton, 1995) (Fig. 7). Some non-linearity in age progression is observed, with ages of 20 Ma and 27 Ma recorded on both sides of the fracture zone (Clouard and Bonneville, 2005).

33 Volcanism then shifts to slightly different trend with a different pole of rotation (Epp, 1984;

Koppers et al., 2001) until approximately 5 Ma. The older age for the Hawaiian-Emperor bend suggested by Sharp and Clague (1999) gives an approximately 20 million year time interval before establishment of significant volcanism, which would be consistent with re- orientation of roll cell convection from sub-parallel to the Kula-Pacific, to sub-parallel to the

Pacific-Farallon ridges. The spacing between the Marshall-Gilbert and Line Island chains in the Late Cretaceous is similar to between the Hawaiian and Cook-Austral chains in the Late

Cenozoic (Fig. 6j), although any significance is uncertain as the wavelength between chains would be double the aspect ratio predicted for convective rolls. The absence of volcanism as the Hawaiian chain crosses the Murray fracture zone could indicate structural control by the lithospheric section; however, there is no appreciable change in volcanic output where the chain crosses the Mendocino or Molokai fracture zones, although a change in chemistry has been suggested to accompany the latter (Basu and Faggart, 1996). The Late Oligocene hiatus also coincides with basin-wide tectonic events including re-organisations along the Antarctic-

Pacific ridge (Kamp, 1991), the division of the Farallon into Cocos and Nazca plates

(Lonsdale and Klitgord, 1978), and collision of the New Guinea passive margin with arc systems bounding the eastern margin of the Philippine plate (Hall, 2002). Other tectonic events at this time included opening of the Japan Sea (28-11 Ma; Jolivet et al., 1995), and initiation of faulting along the San Andreas fault system (Kamp, 1991). Similarly, the bathymetry of the Hawaiian chain shows a further break between the islands of Nihoa (7 Ma) and Niihau (5 Ma), which corresponds to the timing of plate re-organisations along the Fiji margin (Cox and Engebretson, 1985). The ages have similar significance in other island chains with volcanism in the Cook and Foundation chains initiated during the Late Oligocene event,

34 and volcanism in the Austral and Marquesas islands initiated during the Miocene-Pliocene event (Natland and Winterer, 2005). Plate reorganisations may reactivate lines of weakness in the plate fabric (Favela and Anderson, 2000), hence the correlations between ocean island volcanism and events around the basin margins are strong evidence for control of volcanism by the basin-wide stress field.

Volcanism along the Louisville chain also shows a marked decrease in output around 25 Ma

(Géli et al., 1998) although there is limited age data in this part of the chain. Relationships with the Eltanin fracture zone are refuted in hotspot models, with volcanism on the oldest part of the chain considered to pre-date formation of the fracture zone, and the most recent volcanism to curve away from the fracture zone. The current location of the plume was suggested by Lonsdale (1987, 1988) to be marked by a seamount at 500 S 1390 W, northwest of the Eltanin fracture zone (Fig. 8). The Hollister Ridge to the south of the Eltanin fracture zone, has also been proposed as the site of the hotspot (Wessel and Kroenke, 1997), but was found to have a different geochemical signature (Vlastic et al., 1998), and the original location was preferred by Géli et al. (1998). However, a single seamount should not be considered evidence for a plume, and although the seafloor topography map of Smith and Sandwell

(1997) shows few seamounts in the region, there are others at comparable distances from the

Pacific-Antarctica ridge such as the seamount at 600 35’ S 1290 W south of the Udintsev fracture zone (Fig. 8), which have no relationship to any island chain in the region. But even if the seamount at 500 S 1390 10’W is the youngest expression of the chain, it does not detract from observations that Late Cretaceous volcanism along the Louisville chain is paralleled by several small seamounts which may have formed in Puka-Puka ridge-like settings (e.g.

35 seamounts following N210W trend; Fig. 8), while volcanism from 46 to 25 Ma extrapolates into trace of the Eltanin fracture zone as defined by magnetic anomalies. The magnetic anomalies indicate formation of the Eltanin fracture zone after 84 Ma, likely as part of the

Late Cretaceous plate reorganisations, in which case it would not post-date the oldest (77 Ma;

Clouard and Bonneville, 2005) volcanism on the chain. Displacement along the fracture zone increases around 48 Ma coinciding with the mid Eocene change in stress field of the Pacific plate. The correlation between the decline in volcanism and Late Oligocene plate reorganisations in the region, supports control by plate tectonic processes. Indeed, volcanism along the Louisville chain may be considered to have been quintessentially terminated by the

Late Oligocene plate reorganisations, with the sporadic formation of small seamounts and ridges since then being related to ocean ridge processes or intermittent reactivation of lines of weakness in the lithosphere.

DISCUSSION

The history of the Pacific basin outlined in the previous section and Natland and Winterer

(2005), suggests intraplate volcanism may be divided into events before formation of the

Pacific plate, events during the intra-oceanic history of the Pacific plate, and events following subduction of the Pacific plate beneath Asia. The relationship of the plume, SUMA, and streaky-mantle models to these stages as represented by events in the Paleozoic, Late

Jurassic-Early Cretaceous, and Cenozoic, is summarised in figure 9. In the plume model, the distribution of intra-plate volcanism is controlled by events in the deep mantle, with the large- scale mantle upwellings considered to be superplumes (Figs. 9a-c). The concentration of

36 intraplate volcanism on the Izanagi plate in the Paleozoic could then reflect instabilities in thermal boundary layer conditions beneath the western part of the basin. Plume activity must then shift to under the centre of the basin with the arrival of plume heads to generate oceanic plateaus in the Late Jurassic-Early Cretaceous. However, the association of plateaus with ocean ridges requires fortuitous plume arrival or localisation of ridges over plume heads.

Similar problems exist with regard to multiple occurrences of volcanism in the same location.

In ocean island chains, non-linear age progressions require special pleading, either with regard to numbers of plumes, melt plumbing arrangements, or plume shape (e.g. McNutt and Fischer,

1987; Sleep, 1992), while plume splitting has been speculated for repeat episodes of volcanism on oceanic plateaus (Bercovici and Mahoney, 1994). Further fortuitous events are required to explain why different ocean island chains should undergo simultaneous changes in volcanic output in the Late Oligocene and Miocene-Pliocene (Smith, 2003a; Natland and

Winterer, 2005). Why superplumes should currently feed ocean islands under the south- central Pacific via narrow conduits (Courtillot et al., 2003) (Fig. 9c), but generate oceanic plateaus in the mid Cretaceous (Coffin and Eldholm, 1993) (Fig. 9b), has also not been explained in plume models. In the SUMA (Figs. 9d-f) and streaky mantle models (Figs. 9g-i), intraplate volcanism results from preferential tapping of fertile streaks in the convecting mantle, with the volumes of melt reflecting the size and/or percentage of streaks, and the focusing of melts by lithospheric architecture. The large-scale mantle upwellings represent convection cells, which could arise from lack of cooling of the mantle from an absence of subduction in the central region of the basin following the model of Anderson (2001).

37 Differential rotation concepts are not essential in the SUMA and streaky mantle models, but help explain several features regarding the location of volcanism. The concentration of intraplate volcanism on the westward subducting Izanagi and Pacific plates compared to the eastward subducting Farallon plate, can be ascribed to greater shear heating from the counterflow regime set up under the Izanagi and Pacific plates. The migration of shallow material in figure 6 is simplified as enriched domains would be expected to be disrupted by interaction with ocean ridge systems, and only lateral migration has been depicted although plate lag may follow a sinusoidal pattern (Doglioni, 1990). Nonetheless, the differential rotation model explains the distribution of geochemical signatures, with the eastern part of the basin largely devoid of EM1 sources. The Cache Creek terrane contains the oldest

Panthalassan intraplate volcanism for which significant geochemical data exists. Basalts in this terrane are olivine tholeiites similar to those of the Hawaiian islands, but isotopic compositions are similar to the Magellan seamounts (Smith and Lambert, 1995; Tardy et al.,

2001). Such volcanism may thus represent an even earlier expression of the SOPITA anomaly than the Early Cretaceous seamounts on the Pacific plate suggested by Staudigel et al. (1991).

By the Late Jurassic, the Pacific plate would have been centred over the South Pacific domain, hence the EM1 signature in basats from the Shatsky Rise. Volcanism in the Magellan,

Marcus-Wake and Marshall islands would have overlain the South Pacific domain in the Early

Cretaceous, with entrainment of streaks into the ocean ridge systems explaining the volume of melt and EM1 signatures in the Ontong Java and Manahiki plateaus. Delamination of continental mantle/lower crust during the amalgamation of Asia occurred some 100 million years after the breakup of eastern Gondwana, such that a North Pacific domain would have lagged behind the South Pacific domain by some 3000 km. Intraplate volcanism derived from

38 the North Pacific domain may include the islands in the Shimanto belt and the Japanese seamounts as well as Hawaiian volcanism (Figs; 6e-j). The migration rate for such sources from eastern Gondwana in the mid Paleozoic to the present day central Pacific corresponds to an average lateral drift of 3 cm yr-1. The drift rate is less than the current average of 5 cm yr-1 because the net rate is dependent on plate configuration and the action of boundary forces.

The lithosphere currently shows a net westward lag mainly as a result of boundary forces on the Pacific plate (Knopoff and Leeds, 1972), but a configuration with a large plate moving eastwards as depicted in the Carboniferous through Jurassic reconstructions (Figs. 6a-d) would be expected to be characterised by a drift of less than 5 cm yr-1, thereby slowing down the lateral migration rate.

The antipodal positions of the current large-scale upwelling under the Pacific and western

Africa, may be explained by the most stable distribution of mass being about the axes of a triaxial Earth (Lewis and Smith, 1998). Such a system could never reach equilibrium, as it would be self-disrupting by the mass distrubution being upset by subduction of slabs. If the upwellings are long lived as suggested for under Africa (Ni and Helmberger, 2002), westward plate lag would result not only in eastward migration of geochemical domains, but also of the large-scale convection cells. Fragmentation of eastern Gondwana in the mid Paleozoic could then be explained by lithospheric lag moving this region of the supercontinent over a convection cell which is now under the south-central Pacific (Fig. 9g). The corresponding antipodal cell would have been situated under the eastern Panthalassan basin in the mid

Paleozoic, before moving under Pangea in the Permian-Triassic, and causing the opening of the

Atlantic in the Cretaceous. Such upwellings might be modified, rather than being caused by

39 lithospheric insulation, becoming stronger when supplemented by insulation under a continental mass or not cooled by subducting slabs. The South Pacific domain may thus be a thermal and geochemical anomaly, whereas the North Pacific domain may be predominantly geochemical, hence the greater volume of magmatism in the equatorial and South Pacific regions over time.

The differential rotation model also provides a link between tectonic events and magnetic reversals. In plume models, convection patterns in the outer core are influenced by insulation at the core mantle boundary (Vogt, 1975; Larson, 1991b; Larson and Olson, 1991).

Observations on the magnetic field have shown the non-dipole moment of the magnetic field to be drifting westwards of at a rate of 0.20 yr-1 relative to the mantle (Jault et al., 1988), whereas the inner core has been shown to be rotating eastwards at a rate of 3.30 yr-1 (Su et al.,

1997). The Earth can thus be envisaged as a series of shells with the lithosphere and convection pattern in the outer core showing a net westward drift relative to the rotation axis, while the mesosphere/deep mantle and inner core move eastwards relative to the rotation axis

(Fig. 10). The no-net rotation assumption required to avoid the lithosphere acting as a brake on Earth rotation in standard plate tectonic models, is unnecessary in the differential rotation model (Cuffaro and Doglioni, this volume), as changes in drift in one shell may be compensated in another. Previous studies have argued for a balance of angular momentum between the core and mantle system and noted that torques exerted by atmospheric effects can lead to short term pertubations in the Earth’s rotation (Hide, 1993; Jackson et al., 1993).

Under such circumstances it is reasonable to expect changes in lithospheric plate configuration to affect the balance of angular momentum. As the total angular momentum of the Earth must

40 remain constant in the absence of any changes in external torques (Jackson et al., 1993), balancing angular momentum changes in the lithosphere by changes in core convection patterns would provide a link between magnetic reversals and plate tectonics without having to invoke mantle plumes (Smith and Lewis, 1999a).

CONCLUSIONS

The record of intraplate volcanism in the Panthalassan and Pacific ocean basins shows a profound influence from plate tectonic events and lithospheric architecture, suggesting an origin related to stress fields and surface processes rather than mantle plumes. A shallow- source plate model for such volcanism can be constructed from the following concepts:

(1) Remixing of subducted oceanic crust into the convecting mantle as in the SUMA model of

Meibom and Anderson (2003) and streaky-mantle model of Smith (2005). In both models,

imperfect remixing of subducted material produces a greater size range of heterogenities

than in earlier marble-cake models, but differ in that in the SUMA model subducted oceanic

crust is remixed into a conventional depleted mantle matrix, whereas in the streaky mantle

model the depleted endmember in MORB is recycled oceanic crust which underwent

melting in ancient subduction zone environments.

(2) Low melting point sources for intraplate volcanism. These are equated with volatile-bearing

minerals which may be formed in the hanging-wall of the mantle wedge during subduction

events.

(3) Differential rotation of lithosphere and mantle, whereby the lithosphere shows a westward

lag relative to the deep mantle as a result of attenuation of drag through the low viscosity

41 asthenosphere. The mechanism provides a means for the lateral introduction of source

materials, in particular delaminated/eroded continental mantle or lower crust from eastern

Gondwana, and may lead to shear heating of viscous heterogeneities under westward-

subducting plates. The differential rotation mechanism also provides a link between

magnetic reversals and tectonic events as a net rotation of the lithosphere may be

compensated by convection in the core.

(4) Interpretation of large-scale upwellings delineated in tomographic profiles, as mantle

convection cells rather than superplumes.

Reconstructing the Pacific basin from the current features of the basin and the terrane accretion record around the basin margins, and using base maps based on paleomagnetic data rather than plate motions in the hotspot frame, indicates three phases of volcanism in addition to the Panthalassan history of the basin:

(1) An intra-oceanic stage of the Pacific plate from the Mid Jurassic to mid Cretaceous,

characterised by formation of oceanic plateaus along ocean ridge systems and diffuse ocean

island volcanism in the centre of the Pacific plate. Plateau volcanism is ascribed to focusing

of melts at triple junctions and/or entrainment of low melting point heterogeneities into

ridge upwelling. The lack of age progression in the ocean island volcanism (Marcus-Wake,

Magellan, Marshall islands) reflects stressing of the plate between four ridge systems.

(2) A post- collision of the Pacific plate with the margin of Asia phase from the Late

Cretaceous to mid Eocene, where the plate undergoes subduction along its western margin,

but is still bounded by ridge systems in the north, south and east of the basin. Ocean island

42 volcanism on the Pacific plate was not necessarily parallel to plate motion at this time, and

was controlled largely by stress trajectories between ocean ridge systems.

(3) A mid Eocene to Recent phase where volcanism is dominated by stress fields arising from

convergent margins in the north, south and west of the basin following the demise of the

Kula-Pacific and North New Guinea-Pacific ocean ridges. Volcanism during this phase is

influenced largely by plate tearing and thermal stresses.

Volcanism on the Izanagi and Farallon plates is inferred to have been similar to the third phase for the Pacific plate as these plates were bounded predominantly by convergent margins throughout the time period considered in this study. The model does not require the presence of mantle plumes, indeed such features are superfluous as the geochemical components attributed to such are already present in the convecting mantle and can be tapped by shear heating and localised convection, with eruption of melt controlled by lithospheric architecture.

This feature is underscored by explanations of the Ontong Java and Hawaiian volcanism, hitherto considered strong evidence for the existence of plumes, as a result of plate tectonic processes. Formation of the Ontong Java plateau is attributed to entrainment of former mantle wedge material containing hydrous minerals into ocean ridge upwelling, coinciding with a decrease in spreading rate which resulted in ponding of melt over the ridge system.

Reconstruction of plate configurations for the mid Eocene suggest the stress field in the western Pacific underwent a change from compressional to extensional as a result of the contemporaneous demise of the Pacific-Kula and Pacific-North New Guinea ridges, resulting in formation of the Hawaiian chain as a propagating fracture. The three principal categories of

43 volcanism, intraplate, arc and ocean ridge, all arise from plate tectonic processes, unlike in the hotspot model where intraplate volcanism is superimposed on plate tectonics by plumes.

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Figure Captions

1. Map of the Pacific basin showing the location of ocean island volcanism and oceanic

plateaus on the present day seafloor and in accreted terranes around the basin margin.

Abbreviations: plates: AN Antarctica, CO Cocos, NZ Nazca, PA Pacific, PH Philippine;

ocean ridges: EPR East Pacific Rise JdF Juan de Fuca; fracture zones: EFZ Eltanin fracture

zone; seamounts, island chains, plateaus and rises with ages in Ma: AU Austral-Cook (0-

29), CA Caroline (1-14), CE Cobb-Eickelberg (2-26), CN Carnegie Ridge (0-15), CS Cocos

64 Ridge (0-20), EA Easter (9-0), EM Emperor (47-81), FO Foundation (2-22), HA

Hawaiian (0-47), HR Hess Rise (88-111), JF Juan Fernandez (<30), JS Japanese

seamounts (71-108), KB Kodiak-Bowie (0-25), LI Line islands (35-128), LV Louisville (0-

67), MG Marshall Gilbert (56-138), MJ Meiji (>81), ML Magellan Rise (100-135), MN

Manihiki Plateau (110-123), MP Mid-Pacific Mountains (83->129), MQ Marquesas (1-

9), MS Magellan seamounts (74-129), MU Musicians (64-98), MW Marcus Wake (78-

120), OJ Ontong Java (90, 121-124), PG Pitcairn-Gambier (1-25), PP Pukapuka ridge (5-

28), SA Samoa (5-14), SG Sala y Gomez (0-22), SO Society (0-5), SR Shatsky Rise (129-

143), TO Tokelau (43-80), TU Tuamotu (<43). Accretionary terranes containing ocean

island/plateau material: AN Anvina, BD Badzhal, BR Bridge River, CC Cache Creek, GU

Guerrero, HI Hidaka, KR Kronotskiy, NB Nabilsky, NK North Kitakami, SM Shimanto,

SY Sorachi-Yezo, TK Tokoro, VA Vatyn, WP Waipapa, WU Wuho block, YB Yuli belt,

YM Yamaguchi belt. Compiled from Ben-Avraham et al. (1981), Watson and Fujita

(1985), Coney (1989), Taira et al. (1989) and Nockleberg et al. (2000). Age dates are from

Duncan and Clague (1985), Nakanishi et al. (1999), Clouard and Bonneville (2001; 2005)

and Koppers et al. (2001).

2. Pb-Pb isotopic variation in ancient and modern MORB (black), oceanic plateau basalts

(orange), seamounts and intraplate ridges (purple), and ocean islands (green) in the Pacific

basin. The mantle components BE (bulk Earth), DM (depleted mantle), EM1 (enriched

mantle 1), EM2 (enriched mantle 2) and HIMU (high-m) components of Zindler and Hart

(1986), Northern Hemisphere Reference Line (NHRL) of Hart (1984), and Lower Crust

estimate of Doe and Zartman (1979) are shown for reference. The fields for ancient

65 samples are based on measured rations to facilitate comparison of samples of a wide range

of ages. Data sources: (a,b) modern Pacific basin: East Pacific Rise MORB (EPR; Ito et al.,

1987; White et al., 1987), East Pacific Rise seamounts (Niu et al., 2002), Garrett transform

(Wendt et al., 1999), Guadalupe (Sun, 1980), Hawaii (Abouchami et al. (2003), Juan de

Fuca ridge (Hegner and Tatsumoto, 1989), Mangaia and Tubuai (Palacz and Saunders,

1986; Zindler and Hart, 1986; Schiano et al., 2001), Mathematicians Ridge (Sun, 1980),

Pitcairn seamounts (Woodhead and Devey, 1993), Pukapuka ridge (Janney et al., 2000),

Shimada seamount (Gardner et al., 1984), Society islands (White and Duncan, 1996). (c,d)

Ancient samples: 130-151 Ma MORB (Janney and Castillo, 1997), 160 Ma MORB

(Floyd et al., 1991), Alexander island (Mahoney et al., 1998), Cache Creek (Smith and

Lambert, 1995), Detroit and Meiji seamounts (Regelous et al., 2003), Magellan, Marshall,

and Marcus-Wake islands (Staudigel et al., 1991), Manahiki and Ontong Java plateaus

(Mahoney, 1987; Mahoney et al., 1993), Shatsky Rise (Mahoney et al., 2005). Also

shown is the field for Indian Ocean MORB from Mahoney et al. (1998).

3. Plot of Ce/Nb versus Th/Nb illustrating variation in light rare earth, large-ion lithophile and

high field strength elements in (a) modern- (b) ancient MORB, oceanic plateau basalts,

seamounts, and ocean islands in the Pacific basin. Data sources as for figure 2, additionally

N-MORB: Saunders et al. (1988); Mathematicians Ridge: Batiza and Vanko (1985); 170

Ma MORB: Fisk and Kelley (2002); East Pacific Rise seamounts: Niu and Batiza (1997);

Sorachi terrane: Kimura et al. (1994); Wuho block: Sun et al. (1998); Yuli belt: Yui et al.

(1990). Also shown are compositional rages for potassic volcanic rocks from eastern China

(Wudalianchi; Liu et al., 1994) and western North America (Crazy Mountains province;

66 Dudás, 1991), and compositions for mantle phlogopite (phlog) and amphibole (amph)

(Ionov and Hofmann, 1995). The N-MORB, Pitcairn and Mathematicians Ridge fields are

shown as dotted lines in (b) for reference.

4. Asthenosphere flow regimes beneath easterly- and westerly- subducting plates in

conventional plate tectonic models using the hotspot reference frame (HRF) and the

differential rotation model which uses the rotation axis reference (RAR). The

asthenosphere is depicted (following Turcotte and Schubert, 1982) as a channel where

decoupling takes place between the lithosphere and deep mantle (mesosphere)

characterised by large-scale thermal convection.

(a) Nazca/Farallon plates, HRF. Movement of plates by boundary forces exerts a drag on

the top of the asthenosphere which induces simple Couette flow in the asthenosphere.

(b) Pacific/Izanagi plates, HRF. Movement of plates by boundary forces is balanced by

counterflow in the asthenospheric channel.

imparts a drag on the base of the asthenosphere which supplements movement of the

plates boundary forces, resulting in a Couette flow profile with reduced displacement

between lithosphere and deep mantle.

(d) Pacific/Izanagi plates, differential rotation model. Eastward movement of the mantle acts

in the opposite direction to plate movement by boundary forces. The resulting

counterflow regime is characterised by a stationary layer (u=0) as in plume models, but

the relative displacements of lithosphere and deep mantle are again reduced.

67 Average westward migration of the lithosphere of 5 cm yr-1 in the RAR frame relative to 1.7 cm yr-1 in the hotspot reference frame compared to relative to a fixed Antarctic plate can be explained by the sources of intraplate volcanism lying at shallow depths in the asthenosphere.

5. Pressure-temperature stability of hydrous peridotite assemblages relative to slab, shield

and oceanic intraplate geotherms. Geotherms are from Thompson (1992). The mantle

potential temperature is assumed to be 12800C following McKenzie and Bickle (1988).

The dry peridotite solidus is from Trønnes (2002). The fields for pargasite, phlogopite and

K-richterite were constructed from Menner and Dunn (1995), Sato et al., (1997), Trønnes

(2002) and Konzett and Fei (2000). The shaded field illustrates the potential increase in off

axis geotherm for shear melting under the counterflow regime in figure 4d. Temperature-

pressure estimates for the formation of leucitite-picrite melt types are from Green and

Falloon (1998).

6. Plate reconstructions showing the distribution of active intraplate volcanism (orange) in the

Panthalassan-Pacific basin from the Mid Paleozoic to Recent as compiled from present day

features of the ocean floor and evidence from accreted terranes. The position of oceanic

plateaus is shown for reference, although for simplicity, inactive ocean islands and arc

volcanism have been omitted. Ocean ridge positions for the Carboniferous through Mid

Jurassic are speculative. Ocean ridge positions at other times and outlines of continents are

based on Scotese et al. (1979; 1988) with modifications following Rea and Dixon (1983),

Zonenshain et al. (1987), Lonsdale (1988a) Maruyama et al. (1989), Metcalfe (1996), Hall

(1996), Li (1998), Sager (2002), Riisager et al. (2003), and Norton (this volume). Also

68 shown is the eastward migration of mantle domains according to the differential rotation model. Abbreviations are the same as in figure 1. Additionally, crustal blocks and microcontinents from Zonenshain et al. (1987) and Maruyama et al. (1989): AB Abukuma-

Kitakami, IC Indochina, IT Insular terrane, KU Kurosegawa belt, NC North China, OK

Okhotia, RB Reed Bank (representative of microcontinental fragments in the South China

Sea), SC South China, TH Tahin, TR Tarim.

(a) Carboniferous (~340 Ma). The eastern margin of Gondwana is fragmented, eroding

continental mantle/lower crust into the shallow mantle to create a South Pacific

enriched mantle domain (blue) which eventually becomes the SOPITA anomaly.

(b) Late Permian (240-250 Ma). The continental blocks and microcontinental fragments

derived from Gondwana migrate northward across eastern Tethys, accompanied by

extensive ocean island volcanism on the Izanagi plate. Collisions between the

continental blocks, as well as fluxing of fluids into the mantle via arc activity, creates a

North Pacific enriched mantle domain (purple) in the northwest of the basin.

island volcanism occurs largely on the Izangi plate.

(d) Late Jurassic (150-130 Ma). The Shatsky Rise and Mid Pacific Mountains form in

conjunction with triple junctions along the Pacific-Izanagi and Pacific-Farallon ridges.

Dotted lines represent the directions of incipient fracturing of the Pacific plate.

(e) Late Cretaceous (130-100 Ma). The Pacific plate has enlarged but remains intra-

oceanic. Continued formation of the Mid Pacific Mountains and volcanism on the

Magellan Rise can be related to triple junctions along the Pacific-Farallon ridge. The

69 Ontong Java and Hess Rise plateaus form on the Pacific-Phoenix and Pacific-Farallon

ridges, respectively. Dotted lines as for figure 6d.

(f) Late Cretaceous (~83 Ma). Ocean island volcanism immediately prior to the formation

of the Kula plate) occurs in the Marshall-Gilbert, Line Island and Musicians chains.

The position of Meiji seamount corresponds to that of a triple junction on the Farallon

North-Pacific ridge and is interpreted as a failed oceanic plateau.

(g) Paleocene (~59 Ma). The Kula plate forms from the North Farallon plate with

volcanism along the Emperor chain interpreted as a propagating fracture induced by

Kula-Pacific ridge geometry.

(h) Mid Eocene (47-43 Ma). Near contemporaneous abandonment of spreading on the

Kula-Pacific ridge and subduction of the North New Guinea-Pacific ridge, alters the

stress field in the western part of the Pacific plate, resulting in formation of the

Hawaiian chain as a propagating fracture.

(i) Early Oligocene (~38 Ma). The Nazca and Juan de Fuca plates form as remnants of the

Farallon plate as the Pacific-Farallon ridge collides with North America.

(j) Miocene-Recent (20-0 Ma). Intraplate volcanism occurs along the Hawaiian chain and

over the SOPITA anomaly in the south-central Pacific. Volcanism along the Louisville

chain is considered to have largely ceased following plate reorganisations in the Late

Oligocene.

7. Possible morphological relationship of the Hawaiian-Emperor chain (bathymetric outline

according to the 4 km isobath) to events around the basin margin. Age dates are from Keller

70 et al. (1995) and Clouard and Bonneville (2005). Fracture zone trends are from Rea and

Dixon (1983).

8. Map showing the relationship of mid Eocene-Oligocene volcanism along the Louisville

ridge to the Eltanin fracture zone (After Smith and Sandwell, 1997; Vlastelic et al., 1998).

Magnetic lineations are from Müller et al. (1997), age dates are from Clouard and

Bonneville (2005). Abbreviations: AN Antractic plate, PA Pacific plate.

9. Geodynamic models for intraplate volcanism in the Panthalassan-Pacific basin during the

Late Paleozoic, Mid Cretaceous and Cenozoic. Abbreviations: OR ocean ridges, OIV ocean

island volcanism, DM depleted mantle, LM lower mantle, PM primitive mantle. (a-c)

Plume model. In the simplest version, subducted oceanic crust is stored in thermal

boundary layers, whilst heterogeneity in the depleted mantle results from refertilization

with plume residues. (d-f) SUMA model. Heterogeneity in the depleted mantle results

from remixing of subducted oceanic crust (black streaks) along with delaminated lower

crust (orange) into the convecting mantle. (g-i) Streaky mantle model. Subducted oceanic

crust is remixed with the convecting mantle as in SUMA models, but the model

differentiates between subducted dehydrated crust (black) and subducted melted crust

(green). A further level of heterogeneity results from the eastward lateral migration of

mantle domains containing delaminated continental mantle and lower crust (orange) as a

result of westward plate drift.

71 10. Causes of correlation between magnetic reversal frequency and tectonic events (after Smith

and Lewis, 1999a): (a) Plume model. The pattern of convection in the outer core which

gives rise to the magnetic field is controlled by the insulating effect of the D” layer which in

turn is controlled by the release of plumes (b) Differential rotation model. The link with

surface tectonic events is provided by the balance of angular momentum between the

lithosphere, mantle and core. Numbers indicate the vertical components of motion in

degrees per year relative to the hotspot frame (HRF) or relative to the rotation axis

reference (RAR).

72 Figure 1 Figure 2 Figure 3 Figure 4 Figure 5 Figure 6a

Figure 6b Figure 6c

Figure 6d Figure 6e

Figure 6f Figure 6g

Figure 6h Figure 6i

Figure 6j Figure 7 Figure 8 Figure 9 Figure 10