Crust and upper mantle structure of a continental backarc: central North Island, New Zealand

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1 Geophys. J. Int. (2006) 166, doi: /j X x Crust and upper mantle structure of a continental backarc: central North Island, New Zealand W. R. Stratford and T. A. Stern School of Earth Sciences, Victoria University of Wellington, Wellington, New Zealand. wanda.stratford@vuw.ac.nz Accepted 2006 February 22. Received 2006 February 21; in original form 2005 April 19 1 INTRODUCTION Behind active continental margins where the slab is in retreat, the lithosphere can be extended, heavily modified and/or completely replaced by new material. In these continental backarc zones (Uyeda & Kanamori 1979) we have the rare opportunity to observe processes that lead to the generation of new continental lithosphere. The Central Volcanic Region (CVR) of New Zealand is a particularly favourable locality to make the observations as most of the area is above sea level. Moreover, because the CVR is located at the south end of the Lau-Havre Trough, the oceanic backarc basin behind the Tonga-Kermadec subduction zone, we can observe the transition from rifting of oceanic to continental lithosphere. One of the main foci for studies of continental rifting is an understanding of how the crust mantle boundary (Moho) evolves in an extensional regime (Griffin & O Reilly 1987; Wendlendt et al. 1991; Wilshire 1990). Seismic velocities of km s 1 are commonly observed at Moho depths beneath continental rifts (Makris & SUMMARY Near vertical and wide-angle seismic data provide evidence for a gradational crust mantle boundary in a depth range of km beneath the Central Volcanic Region (CVR), New Zealand. This volcanic area includes the Taupo Volcanic Zone and is a direct extension of Tonga-Kermadec oceanic backarc spreading into continental lithosphere. Long-range seismic refraction data show velocities of 6 km s 1 and less within the top 15 km of the crust of the CVR. At a depth of 15 km compressional seismic velocities increase to 6.8 km s 1, and then to 7.4 ± 0.2 km s 1 at 20 km depth. These 7.4 km s 1 seismic wave speeds are interpreted as anomalous upper mantle as beneath this level passive seismic studies show similar Pn wave speeds that increase slowly to 7.8 km s 1 at about 80 km deep. We interpret rocks between 15 and 20 km to be a layer of new crust formed by underplating. The strongest reflection observed, and what might also be interpreted as a reflection Moho, is from the top of the proposed underplated layer at 15 km depth. At 20 km depth no such distinct reflection is observed. Rather, wide-angle reflection data show a continuum of low-level reflectivity between 15 and at least 35 km depth, indicating some heterogeneity and/or structure within the lower crust and upper mantle. Thus the transition from lower crust to upper mantle is broad, and a conventional reflection Moho does not exist beneath the CVR. Buoyancy force calculations based on rock uplift for the central North Island indicate that the subjacent mantle, to a depth of km is 70 kg m 3 or 2 per cent less dense than normal mantle. Best estimates attribute half of this density anomaly to the effects of increased temperature with additional contributions from partial melt ( 1.2 per cent) and melt residuum. Key words: backarc basin, crustal structure, Moho discontinuity, reflection seismology, refraction seismology, underplating. Ginzburg 1987; Mackenzie et al. 2005). These velocities have been interpreted as both intruded lower crust (Jarchow et al. 1993; Thybo et al. 2000; Mackenzie et al. 2005; ten Brink & Taylor 2002) and anomalous upper mantle that contains a small percent of partial melt (Makris & Ginzburg 1987; Parsons & McCarthy 1996; Louie et al. 2004). Answering this question is important as it bears on questions related to melt extraction from the mantle and heat transfer mechanisms. This ambiguity first surfaced in the western USA when the results of reconnaissance seismic surveys became available (Cook 1962). In New Zealand early refraction and wide-angle reflection surveys of the CVR interpreted an unusually thin crust of 16 ± 1 km and low upper mantle velocities of 7.5 km s 1 ± 0.2 (Stern & Davey 1987). These velocities were recorded from four shots on 20 analogue instruments in a profile running approximately north-south through the volcanic area. A new wide-angle reflection and refraction data set was collected as part of the NIGHT project (North Island GeopHysical Transect) in 2001 on two transect lines, north south and west east across the central North Island and CVR (Henrys GJI Volcanology, geothermics, fluids and rocks C 2006 The Authors 469

2 470 W. R. Stratford and T. A. Stern et al. 2003; Stratford & Stern 2004). From preliminary analysis of the NIGHT data Stratford & Stern (2004) also interpreted a 16- km-thick crust, a low (7.4 km s 1 ) upper mantle P wave speed, and identified a prominent reflection (PmP 2 ) from within the mantle wedge. An alternative model, from a subset of the NIGHT data set (Harrison & White 2004), featured a 30-km-thick crust beneath the CVR. This study used wide-angle reflection data and inversion of earthquake traveltimes but was not constrained by velocities from long-range refraction shooting. This study builds on these initial interpretations of the NIGHT data set (Stratford & Stern 2004) and, using additional data, reinterprets the fine structure of the crust mantle transition zone and mantle beneath the CVR, and the crustal structure of the central North Island. In particular, we present new long-range refraction data, with better resolution on velocity with depth, to detail the unusually gradational transition from crust to upper mantle of this continental backarc basin. We also integrate these results with geological observations of rock uplift (Pulford & Stern 2004), and temperature anomaly models (Salmon et al. 2004) to infer mantle perturbation. 2 BACKGROUND Karig (1970) first recognized that the Lau-Havre Trough, the backarc basin behind the Tonga-Kermadec Trench, has an apparent continuation into the North Island of New Zealand (Fig. 1). Backarc extension in the Lau-Havre Trough began around 5 Ma (Wright 1993) and is evident onshore from 4 Ma in the wedge-shaped CVR; a region defined by extensional tectonics, andesite volcanoes and negative residual gravity anomalies (Figs 1 and 2). Plate reconstructions place the southern end of the Tonga-Kermadec-Hikurangi subduction system off the coast of Northland at 30 Ma (Fig. 1) (Walcott 1987). Slab roll-back and an evolving subduction system are manifest in a clockwise rotation (Walcott 1987; Wallace et al. 2004) and southeast translation of the active andesitic front from the arc position at 4 Ma south of Coromandel to the present-day location at the eastern margin of the CVR (Fig. 1) (Hatherton 1969; Stern & Davey 1987). Such rapid fan-like opening of continental backarc basins is seen elsewhere in the Japan Sea (Otofuji & Matsuda 1987) and the Mediterranean (Faccenna et al. 2001). The CVR is an active extensional zone containing approximately km 3 of siliceous volcanics (Houghton et al. 1995) in a 2-km-deep graben (Stratford & Stern 2004). Present-day extension in the graben at the southern and northern ends, as modelled from GPS data, is 8 ± 2mmyr 1 and 15 ± 2mmyr 1, respectively (Wallace et al. 2004). Natural heat output from geothermal fields is around W (Bibby et al. 1994) implying an effective heat flow of about 850 mw m 2 when averaged over the eastern half of the CVR where the active geothermal fields are found (Stern 1987). Rhyolitic volcanism has been more or less continually active in the North Island since 12 Ma (Carter et al. 2003). Active andesite and rhyolite volcanism is now confined to the eastern half of the CVR in what is more widely referred to as the Taupo Volcanic Zone (TVZ) (Cole et al. 1995; Wilson et al. 1995). Most recent normal faulting is confined to the central TVZ area (Villamor & Berryman 2001). However, onshore and offshore seismic reflection profiles show that, except for a few localized highs, the basement surface of the CVR and offshore (Horgan 2003) is depressed by 2 km (Davey et al. 1995; Stern 1986; Horgan 2003) (Fig. 2). This indicates the region as a whole has undergone similar amounts of extension in its opening history. Despite the extension, subsidence and thin crust in the CVR (Stern & Davey 1987; Stratford & Stern 2004; Stern 1987) the regions immediately to the west and south of the CVR are undergoing both Figure 1. Location and tectonic setting of the central North Island. Shading represents a simplified interpretation of the geology of the study region. Plate boundary locations, dates and migration direction relative to North Island from Walcott (1987) and King (2000). Convergence rate of the Pacific Plate relative to a fixed Australian Plate (insert) from DeMets et al. (1990). Active andesitic arc locations (from Stern 1987) are marked by black lines (with dates in Ma). Rotation rate of the eastern North Island from Wright & Walcott (1986). Western and eastern boundaries of the CVR are approximated by the 4 and 0 Ma andesitic volcanic front lines. Western boundary of the Taupo Volcanic Zone (TVZ) is approximated by the line of the 2 Ma andesitic volcanic front.

3 Lithospheric structure of a continental backarc 471 Figure 2. Geological and geophysical evidence for vertical motions in central North Island. Dashed contour lines indicate rock uplift from exhumation studies (Pulford & Stern 2004). Solid contour lines are residual gravity field of the CVR (Stern 1986). Edges of the CVR coincide with the 0 MGal contour. The 40 MGal anomaly of the CVR indicates 2 km of basement subsidence, as evident in interpretations of offshore seismic reflection data (From Davey et al. 1995). Depth scale on seismic interpretation is two-way traveltime. For 2 km s 1 average velocity, 2stwt= 2 km depth. Basement subsidence (extension of the basement on normal faults) is similar across the whole CVR. surface and rock uplift (Armstrong et al. 1998; Pulford & Stern 2004). Mudstone porosity analysis in these regions shows a long wavelength dome of rock-uplift centred on the southern extent of the volcanic area (Fig. 2) (Pulford & Stern 2004). Onset of the uplift of the dome is timed to be 5 ± 1 Ma, based on unconformities cut in the offshore Taranaki Basin (Stagpoole et al. 1997), and appears to precede the onset of andesite volcanism in the CVR by 1 Ma, and in the TVZ by 3 Ma. Pulford & Stern (2004) attribute the uplift seen in the central North Island to a buoyancy force from a distributed low-density mantle. Temperatures 350 above regular mantle and hence lower mantle densities are also implied by the anomalously low Q values for the upper mantle of the western and central North Island (Salmon et al. 2004). 3 SEISMIC DATA COLLECTION Seismic exploration in the CVR has in the past been hampered by the thick volcanic ash giving rise to usually high levels of scattering and attenuation (Bannister & Melhuish 1997; Stern 1986). To optimize energy propagation in such areas, high gain sites, more powerful shots and different survey techniques must be used. For example, best results from the NIGHT survey were gained by undershooting the volcanic area, whereby shots and receivers were placed outside the CVR and energy was reflected from interfaces beneath the region (Stratford & Stern 2004). This strategy has been used successfully to obtain structure beneath other volcanic areas where surface structure inhibits the propagation of seismic energy (Auger et al. 2003). Active source seismic data were collected on two main transects across the central North Island as part of the NIGHT project (Henrys et al. 2003) (Fig. 3). A total of kg shots in 50-m-deep drill holes were used as sources along two lines. A 250-km-long line was recorded west east across the axis of the CVR (Fig. 4). A second line, 120 km long was recorded north south, parallel to the axis and entirely within the boundaries of the CVR (Fig. 5). 300 single component instruments with 4.5 Hz sensors were used with station spacings of 1 km. 77, three-component instruments with 2 and 4.5 Hz sensors were also interspersed at 3 km spacings along the west east line. Offsets over which continuous first arrivals are recorded are variable, shot dependent and range from 30 to 120 km (Figs 4 and 5). Second arrivals and reflections are scarce. Two deep reflections were, however, recorded from beneath the volcanic area where the shot and instrument geometries were favourable. Refracted arrivals from the crust mantle transition zone were recorded from the northern-most shot on the north south line (Fig. 5). Data were also recorded on a 24-channel seismograph at three different locations (Figs 6a, b and c). With a line spread of only 1 km and a 50 m sensor spacing, this 24-channel system provided vital, high-resolution data. Additional data were also recorded on a

4 472 W. R. Stratford and T. A. Stern A Vp decrease from 6.0 km s 1 to say 5.6 km s 1 in such a layer would reduce the Pn refractor depth to 18 km depth. These are two possible end member models and both give depths for the Pn refractor within its estimated total uncertainty bounds of 20 ± 2 km. Based on regression analysis of traveltime picks, velocities are estimated to have errors of ± 0.1 km s 1 for the crust and ± 0.2 km s 1 for deeper phases. Figure 3. White lines are the west east, and north south seismic transect lines of the NIGHT (North Island GeopHysical Transect) project. Numbered stars are shot locations. Black dots are the deployment locations for the R24/48-channel seismograph. Example shots for the west east and north south lines are in Figs 4 and 5, respectively. The shaded ellipses show the approximate geographic surface extents of the mantle reflector (PmP 2 )on the west east and north south lines. A partial shot gather showing the PmP 2 arrivals on the west east line is shown in Fig channel seismograph, with the same 50 m sensor spacing, on a 2 km line at the western edge of the CVR (Fig. 6d). 4 SEISMIC PROCESSING, MODELLING AND INTERPRETATIONS Processing of the seismic data included basic frequency filtering (bandpass Hz for crustal arrivals and Hz for upper mantle phases) and trace balancing. Forward modelling of traveltimes was carried out using the ray-tracing software of Luetgert (1992). When modelling wide-angle seismic data both model-dependent and data-dependent uncertainties must be considered (Zelt 1999). Data-dependent errors arise from the RMS misfit of the model to the traveltime picks. Uncertainties must also be assigned to the traveltime picks to prevent over or under-fitting of the data (Zelt 1999). For forward modelling, the pick uncertainties and RMS model misfit are combined in a calculation of the total data-dependent error given by A 2 + B 2. Where A is the pick error and B is the RMS misfit of the model to the picks. From this, data-dependent uncertainties for the crustal phases of s and mantle phases of s are estimated. Model-dependent errors occur when the assumed velocity increase with depth is invalid. Hidden low-velocity or blind layers can cause errors in depth determinations and where less traveltime data is available, model errors may also occur where the assumption of isovelocity layers falls down. Without prior knowledge of the velocity structure it is impossible to directly assess these uncertainties. Given that the velocities in the upper crust are well constrained, what can be done is an assessment of the two end-member models of crustal structure by manipulating the velocities in the less well constrained lower crust. First, a hidden layer model where we may have missed a high-speed layer at a depth of say km of velocity 6.5 km s 1 (e.g. as proposed by Harrison & White 2004). Forward modelling shows such a layer would increase the Pn refractor depth to 21 km. The second scenario is that we have missed a low-velocity region between km, as suggested by very low S wave speeds in the central TVZ reported by Bannister et al. (2004). 4.1 Volcanic fill/sedimentary section Near surface deposits of low-density, low-velocity, volcanics fill the 2-km-deep depression in the basement surface of the CVR. Pg arrival velocities increase with depth from 0.5 km s 1 at the surface (Alder 1989) to 3kms 1 at km depth (Figs 4 and 5). From previous small scale refraction experiments, (20 m receiver spacing) velocities as high as 3.9 km s 1 have been recorded within the top 200 m on the western edge of the CVR (Alder 1989). Such high velocities represent layers of welded ignimbrite as found by drilling in the area (Houghton et al. 1986), and are not seen on the NIGHT data as the 1 km receiver spacing averages any high-frequency velocity variations. Outside the CVR boundaries, the basement is of Permian to Jurassic greywacke and schist with velocities of km s 1 in the top 2 km. (Sissons & Dibble 1981). West of the CVR, greywacke basement velocities of km s 1 extend to 4km depth (Fig. 4). At the eastern end of the west east line data from shots 1 and 24 show, in the East Coast Basin, velocities increase from 2.5 km s 1 at the surface to 4.5 km s 1 at 5 km depth (Fig. 4c). 4.2 Upper crust Along the north south line buried volcanic structures and/or basement blocks (Fig. 5b) produce variation in the thickness of the volcanic fill. A greywacke outcrop at the Bay of Plenty coast (Fig. 2) produces a local gravity high and is interpreted as an isolated block or sliver of rifted basement rock (Stern 1986). On the western margin of the CVR, seismic surveying over another gravity high is interpreted to show a basement high coming within 600 m of the surface (Figs 5b and 6d). Seismic velocities of km s 1 are observed from basement here. However, both greywacke and andesite/diorite have velocities in this range and we cannot distinguish the rock type for basement on velocity alone. The deepest drill holes in the CVR end in rock of variable composition; andesite, diorite and greywacke have been found at depths of more than 1.5 km beneath the ash (Stern 1987). Throughout most of the CVR we observe velocities of km s 1 between 2 7 km depth (Figs 4c and 5b). These velocities suggest a mixture of felsic-intermediate volcanic rock and/or greywacke (Cole et al. 1995). 4.3 Lower crust of the CVR Previous studies (Stern & Davey 1987) obtained a maximum velocity of 6.1 km s 1 for the lower crust from long offset shots along the north south axis of the CVR. Resolution from this study is higher with many more seismographs available to record the shots. Out to offsets of 80 km, the highest velocity observed is 5.9 km s 1 and these velocities are interpreted to persist to a depth of 15 km (Figs 4c and 5). In contrast, lower crustal velocities to the east (Reyners et al. 1999) and west of this region (Stern & Davey 1987), are 6 per cent higher (Fig. 4c). In this study, we interpret velocities of km s 1 in the lower crust west of the CVR, and km s 1 for the crust

5 Lithospheric structure of a continental backarc 473 Figure 4. Shot 24 (a) and shot 23 (b) from the west east line. traveltime picks for the different phases are marked with small black horizontal lines. Pg = crustal refraction, PmP 1 = reflection from the crust mantle transition zone. Note the attenuation of the Pg arrivals through the volcanic area (west east extent of the CVR is marked with arrows). (c) Interpretations from forward modelling of seismic refraction and reflection arrivals. Shot numbers are labelled with reference to Fig. 3. Note the arched shape of the crust mantle boundary beneath the central North Island. Ray paths where the crust mantle boundary and lower crust are sampled from shots in Figs 3, 6(b) and (d) are shown. Other boundaries without ray paths are interfaces required for modelling crustal refractions (Pg). Dashed lines are inferred boundary continuations. Velocities quoted are in km s 1 and for the crust mantle zone, are from refracted arrivals along the north south line (Fig. 5). to the east from analysis of wide-angle reflection data. Resolution of velocity from these reflections is, however, poor by comparison to the resolution from long-range refraction data. From a near vertical seismic profile at the very western edge of the CVR (Fig. 6d), near Tokoroa, reflectivity in the crust extends down to 6.0 ± 0.5 s two-way traveltime. From a global review of deep seismic reflection data Klemperer (1989) observe that the Moho in regions of extension, as defined by seismic refraction, is approximately coincident with the region where lower crustal reflectivity terminates. On this basis, and using an average crustal velocity for the CVR of 5.6 km s 1 (i.e. 2 km of 3.2 km s 1 and 15 km of 5.9 km s 1 ), the reflection Moho is interpreted to be at a depth of 17 ± 2km (Fig. 4c). 4.4 Crust mantle transition zone beneath the CVR Two sets of refractions and one major reflection event are recorded from the crust mantle transition zone beneath the CVR. Along the north south line refracted arrivals from the top of this zone have an apparent velocity of 7.3 ± 0.2 km s 1 and are interpreted to come from a 15 ± 1 km depth (Fig. 5b). A velocity of 6.8 ± 0.2 km s 1 is obtained for this layer when corrections are made for the pinch-out of near surface ash layers to the south (Fig. 5b). Earlier Pn arrivals are also identified in the shot gather but are of lower frequency, and fan away from the 6.8 km s 1 layer, arriving >0.5 s before it at offsets >100 km (Figs 5a, 6c and 7a). Similarly, the Pn arrivals have a high apparent velocity of 8.1 ± 0.2 km s 1 but corrections

6 474 W. R. Stratford and T. A. Stern Figure 5. (a) Shot 20 from the north south line. traveltime picks for different phases are marked with small black horizontal lines. Pg = crustal refraction, Pn = upper mantle refraction, P = crust mantle transition zone refraction and, PmP 2 = mantle reflection. Note the attenuation of the Pg arrivals at offsets greater than 60 km due to the travel-path being through geothermal areas around the city of Rotorua. Insert is a 24-channel seismograph record recorded at 108 km offset (position is marked with an arrow on main shot gather). These arrivals are also shown in an expanded section, with a low-frequency filter, in Fig. 7(a). (b) Interpretation model from forward modelling of seismic refraction and reflection arrivals assuming horizontal interfaces deeper than 8 km. Ray paths where the crust mantle boundary is sampled are drawn on (from shots gathers in Figs 5 and 6(a), (c) Other boundaries without ray paths are interfaces required for modelling crustal refractions (Pg). Shot numbers are labelled with reference to Fig. 3. North south line skirts over the eastern edge of a basement high near shot 11, at 1 km depth with Vp = 4.8 km s 1. for the variation in near surface velocities gives an apparent velocity of 7.4 ± 0.2 km s 1. This same event is seen on the R24 system with an apparent velocity of 8.1 km s 1 (Fig. 6c). From forward modelling of traveltimes these Pn arrivals are inferred to come from a depth of 20 ± 2 km (Fig. 5b). The Pn velocity of 7.4 ± 0.2 km s 1 is apparent because the profile is unreversed. Nevertheless, some confidence that it represents a true wave speed comes from the fact that a similar velocity for Pn comes from wave propagation in the opposite direction from previous studies (Haines 1979; Stern & Davey 1987). Moreover, more recent studies also infer low Pn velocities (7.4 ± 0.2 km s 1 ) beneath the CVR (Seward et al. 2005) but these increase with depth to 7.8 km s 1 above the slab (Harrison & White 2004). A layer is inferred with velocities increasing from km s 1 in the depth range of km (Fig. 5b). Alternatively, the data are also consistent with a constant velocity layer of 6.8 km s 1 that extends to 19 km depth where the velocity would then jump to 7.4 km s 1. This is a significant velocity change, for which no prominent reflections are recorded. A layer where the velocity gradually increases from the intruded lower crust into the mantle is, therefore, preferred. The primary reflection from the crust mantle region is interpreted as wide-angle arrivals from the 15-km-deep interface. These are recorded on the north south line on the 24-channel record (Fig. 6a), and on the west east line as arrivals with a steep moveout on shot 24 (Fig. 4). Energy on the 24-channel record (Fig. 6a) shows that some diffuse reflectivity extends to depths of at least 35 km (9 s reduced traveltime) beneath the CVR, but the 15 km interface (at 3 s reduced traveltime) produces the only coherent reflected arrivals from this crust mantle zone. We also use the Wiechert-Herglotz-Batement inversion for traveltimes (Lay & Wallace 1995; Meissner 1986) to obtain a continuous

7 Lithospheric structure of a continental backarc 475 Figure 6. Locations for shot gathers given in Fig. 3. (a) 24-channel record recorded in CVR (location 1). Shot gather shows strong Pg arrivals and a reflection from the top of the crust mantle transition zone at 15 km depth. Below this interface reflectivity continues to 9 s reduced traveltime (or 35 km depth). (b) 24-channel record recorded in the western North Island (location 2). Shot gather shows weak Pg arrivals and strong PiP (crustal reflection) and PmP (Moho) reflections. These arrivals are from outside the CVR and show a different pattern of reflectivity to the record recorded from within the CVR boundaries (a). (c) 24-channel record recorded at the southern end of the CVR (location 3) from shot 20. Shot gather shows first arrival Pn, crust mantle refraction (P ) and a PmP 2 (mantle) reflection. Apparent velocities are labelled. This is the same shot gather as the 24-channel record as that shown in Fig. 5. (d) A 48-channel near-vertical (50 m receiver spacing) shot record recorded on the western margin of the CVR (location 4). From 100 kg dynamite shots crustal reflectivity is recorded to 6 s TWTT. Assuming an average crustal velocity of 5.6 km s 1, this gives an approximate depth to the top of the crust mantle zone of 17 km. velocity model with depth (Fig. 8), and hence an alternative estimate for the depth to mantle velocities. The advantage of this method over the plane-layer method used in the previous section is that it does not require the subjective choice of discrete isovelocity layers, but rather assumes a continuous increase in velocity with depth. A numerical function of the form Y = ax b, (1) is fit to the observed traveltimes for refracted arrivals, for shot 20 on the north south line where Y = calculated traveltime and x = offset.

8 476 W. R. Stratford and T. A. Stern Figure 7. Velocity with depth profiles for the Wiechert Herglotz Bateman (WHB) integral and forward modelling solutions. Both velocity profiles are for the arrival traveltimes recorded on the north south line (Fig. 5, shot 20). The ray parameter (p) with offset (x) is obtained by differentiation of eq. (1). The velocity (V) with depth (z) profile for the traveltimes is given by V (z) = 1 π z 0 cosh 1 (V 1 /V ) dx, (2) where V = 1/p and V 1 is the highest velocity sampled by the ray. From eq. (2) a depth to the Pn velocity of 7.4 km s 1 of 22 ± 2km is calculated (Fig. 8). This concurs with the forward modelling solution depth of 20 ± 2 km, the small difference being due to the WHB integral method requiring lateral homogeneity, whereas the raytracing solution allows arbitrary variation of structure along the profile. 4.5 Lower crust and crust mantle transition zone of the western/eastern North Island A strong PiP reflection (lower crust), from beneath the western North Island, is recorded on the R24 system (Fig. 6b), as well as on the main NIGHT data (shot 9, location in Fig. 3). We interpret this arrival as a reflection from a lower crustal interface at 18 km depth (Fig. 4c). On the R24 shot gather a PmP reflection (Moho) is also seen at about 3 s reduced traveltime, which for the offset of km corresponds to a depth of 26 km for the velocity with depth variation shown in Fig. 4(c). This is consistent with results from receiver function analysis (Horspool 2004) and from previous long-range refraction results (Stern et al. 1987). From shot 23 on the western margin of the volcanic region, the primary reflecting horizon is an interface at km depth (Fig. 4b).

9 Lithospheric structure of a continental backarc 477 Depth km Velocity km s 1 Crust-mantle transition zone depth Pn arrival depth WHB integral Forward modelling solution Figure 8. (a) Seismic traces from an expanded section of the long-offset arrivals on the north south line from shot 20. P or crust mantle transition zone refractions and Pn upper mantle refractions are labelled. traveltime picks for the different phases are marked with small black horizontal lines. Arrivals in grey are additional data from shot 10, recorded on instruments on the west east line, that have ray paths close to the north south line. At offsets greater than 100 km the Pn arrivals come in 0.5 s earlier than the P. (b) Pg (crustal refraction) and Pn arrivals from shot 10 as fan arrivals into the west east line. Arrivals are divided into four ray path segments (A, B, C, D) for 2-D modelling of crust and upper mantle structure (see Fig. 9). The irregular shape of this wide-angle reflection indicates it is from a surface that shallows to the east (Fig. 4c). Together, the reflections from shots 9, and 23 and the near vertical record recorded at location 4 (Figs 3, 4c and 6d), describe an arched crust mantle surface beneath the western north and western CVR (Fig. 4c). East of the CVR, crustal thickening is evident from seismic tomography results (Reyners et al. 1999) and a Moho depth of 35 km has been interpreted. 4.6 Crust and upper mantle structure of the western North Island Pn arrivals were recorded within the western North Island and southern CVR in a fan-like array from the Bay of Plenty shot (shot 20, Figs 3, 7 and 9). The array profile extends west of, and includes, the north south line, giving traveltimes for Pn in an offset range of km (Fig. 7). In the following interpretation crustal structure beneath the north south line is fixed for ray paths in the CVR (Figs 5b and 9 insert). Full reversed ray-path coverage of the western North Island is not available. There are, nevertheless, enough data to construct a testable crustal structure model if we use constraints from this study and previous seismic work in the area (Haines 1979; Stern et al. 1987). These earlier studies had approximately north south ray paths that cut the fan array ray paths (Fig. 9 insert) at high angles. The fan array is divided into four sections, A, B, C and D. Crust and mantle structure and P wave speeds are assumed to be common for all arrivals within a section and each section is modelled in 2D. The surface position of the western boundary of the CVR (Fig. 9) is estimated from the gravitationally defined edge of the extended crust (Fig. 2). Where each ray path crosses the CVR boundary, the crustal structure changes from the CVR to western North Island model (Fig. 9b). Arrival times for ray paths entirely within the boundaries of the CVR (ray path D) give a depth and velocity for the Pn refractor of 20 ± 2 km and 7.4 ± 0.2 km s 1, respectively, which are consistent with the interpretation of the north south line (Fig. 5b). Where the upper-mantle ray paths sample the western margin of the CVR (ray path C), the mantle interface is required to be deeper at approximately 21 km but the model is still consistent with low 7.4 km s 1 mantle velocities. A depth of 21 km for the crust mantle interface is consistent with wide-angle reflections recorded right on the western boundary of the CVR on the west east line. The traveltimes of the northern and most westward propagating ray paths (A and B on Figs 7 and 9) are consistent with higher mantle velocities of 7.8 km s 1 beneath the western North Island (Haines 1979; Stern et al. 1987). However, keeping the low Pn velocities of 7.4 km s 1 along the whole CVR section of these ray paths (A and B) will not fit the arrival times. Best fits for the arrival times on paths A and B (Fig. 9) are modelled by increasing mantle velocities to km s 1 under the northwestern section of the CVR (Fig. 9). The alternative, a crustal thickness decrease of at least 4 km coupled with increasing the mantle velocity under the northwestern North Island, gives a poor fit to the slope of the arrivals. Moreover, slightly higher Pn wave speeds under the northwestern section of the CVR is in keeping with this sector being the oldest and, therefore, the coolest part of the CVR, and also concurs with lower attenuation values in the mantle (higher Qp) here than farther east beneath the TVZ (Salmon et al. 2004). Alternatively, anisotropy in the upper mantle may effect our traveltimes. If so, the difference between the 7.4 km s 1 velocities of the north south line and the 7.7 km s 1 velocities in the northwestern corner of the CVR may be due to travel path orientation. Audoine et al. (2004) show short period S-wave anisotropy with approximately west east fast directions in the western CVR at depths <100 km. Where we measure Pn beneath the CVR, the ray paths do not exceed 30 km depth (Fig. 9). An increase in velocity with increasing west east orientation of the ray paths cannot, however, be ruled out as an effect of some shallow fast polarization. Fan arrivals to the east of the north south line could not be seen as these ray paths pass through the highly attenuating geothermal areas and active volcanic centres of the TVZ. Attenuation beneath receivers within the CVR also prevented shots from the western and eastern ends of the west east line being recorded as fan-like arrivals on the north south line. 5 DISCUSSION 5.1 Crust and upper mantle structure Heat flow, extension and intrusion Evidence is equivocal as to what rock type the basement is beneath the CVR. A basement comprised, in part, by greywacke is indicated by lithic fragments found in ignimbrites (Krippner et al. 1998).

10 478 W. R. Stratford and T. A. Stern (a) A B C D X (b) Western North Island CVR (c) A Crust CVR boundary X 0 km 20 B Mantle km s km Crust 20 C Mantle Crust km 20 Mantle D 40 0 km Crust 20 Mantle km Figure 9. (a) Model ray paths for fan arrivals. Arrivals are grouped into the four ray paths A, B, C, and D for 2-D modelling. (b) Crustal structure models used in the 2-D ray-path modelling. Ray paths that cross from the CVR into the western North Island have sections of both crustal structure models. (c) 2-D models of arrivals on the north south line (ray path D), and fan arrivals on the west east line in the western North Island (ray paths A-C). Arrivals for each ray-path segment are shown in Fig. 7. Mantle velocities beneath each ray-path segment are changed until the observed Pn arrivals times (Fig. 7) are matched. However, if the present-day heat output of W is a steady state phenomenon then a significant amount of the basement rock must be igneous material in varying stages of cooling (Stern 1987; Weir 1998). An intruded layer at least 15 km thick is required at the highest extension rate of 15 mm yr 1, measured at the north end of the TVZ, (Wallace et al. 2004) to explain the observed heat output (Stern 1987). However, the more difficult problem is extracting the heat efficiently. Weir (1998) considers a more elaborate model that involves high levels of heat conduction through a ductile region of exponential geothermal gradient. Weir is able to reproduce the heat flux of the TVZ with appropriate extension rates for the whole region by complete replacement of the crust by cooling intrusives and combining circulation by meteoric water in the top 8 km and the zone of high conduction between depths of 8 15 km. Analyses of the chemical and isotopic composition of rhyolites from the CVR also indicate a mostly igneous parent rock, with a small component of greywacke (Graham et al. 1990). Thus although the volumes of intrusive and original greywacke basement rock beneath the CVR cannot be quantified by current geochemical and geophysical data, evidence points to the crustal rock beneath the CVR being composed of a mix of older greywacke fragments in a predominately intruded crust.

11 Lithospheric structure of a continental backarc 479 Figure 10. Summarized crust and upper mantle structure of the central North Island. P wave speeds are given in km s 1. Note the arched shape of the crust mantle transition zone across the central North Island. Mantle velocities under the western North Island are from Stern et al. (1987) and slab velocities from Reyners et al. (1999). Approximate extent of the new crust/underplating and the mantle reflector at depth beneath the CVR are shown. Velocity structure of the crust is simplified and, for the CVR, is based on the interpretation of seismic P wave speeds Lower crust upper mantle One of the key issues in the CVR is whether the Pn velocity of 7.4 ± 0.2 km s 1 (Fig. 10) represents fast lower crust, or perturbed upper mantle containing a few percent partial melt. This problem of interpreting km s 1 Pnspeeds has beset geophysicists working in areas of continental extension for a long time (Cook 1962; Nicolas 1985). We argue that such high velocities for lower crust are unreasonable in a region of high heat flow and volcanic activity. For example, if we apply velocity with temperature derivatives (Wiens & Smith 2003) of δvp/δt = 1.0 m s 1 / k then when these 7.4 km s 1 rocks cool to more normal lower crustal conditions (350 ± 50 temperature anomaly estimated from attenuation studies (Salmon et al. 2004)), they will have increased in wave speed to about 7.8 km s 1, which is too high for lower crustal rocks (Christensen & Mooney 1995). Similarly, if temperature derivatives are applied to the velocities of km s 1 in the km depth range, cooling of these rocks to more normal lower crustal temperatures would give velocities of km s 1. These velocities straddle the boundary between crustal and mantle velocities and a layer of crust/mantle mix or underplating is inferred (Furlong & Fountain 1986). Moreover, a magmatically intruded and underplated layer is in keeping with the extensional tectonic setting, the high volcanic output and heat flow of the CVR, and with crustal structure reported from other continental rifts (Thybo et al. 2000; Jarchow et al. 1993). The primary reflection event and the only clear one we see at wide incidence angles is from the 15-km-deep interface; the boundary between the felsic/intermediate composition intrusions, what remains of the original greywacke basement and the underplate (Figs 4c and 5b). The strong arrivals from this 15 km interface could be interpreted as Moho reflections. However, true, albeit perturbed, mantle velocities are found at 20 ± 2 km depth. These Pn arrivals are not associated with any major reflecting boundary, rather they are interpreted to be from the depth within a continuum of lesser reflectivity where the P wave speeds reach the upper mantle velocity for the CVR of 7.4 km s 1, as calculated by earthquake studies (Seward et al. 2005; Haines 1979; Harrison & White 2004). Similar low Pn wave speeds have been reported from other continental rifts such as the Kenya Rift (7.5 km s 1 ) (Keller et al. 1994) and the Afar Triangle (7.4 km s 1 ) (Makris & Ginzburg 1987). A km variation in the crust mantle boundary depth beneath the central North Island (Figs 4c and 10) is evident. The variation depends on where the Moho interface is taken beneath the CVR, 15 or 20 km. Nevertheless, the broad scale picture of the interface is that of an arched surface shoaling from 25 km under the western North Island, to 21 km at the western margin of the CVR, km beneath the central axis and deepening again to 35 km under the eastern North Island (Figs 4c and 10). Such a quasi-symmetrical arched surface is similar to that seen beneath the Rio Grande Rift (Wilson et al. 2005) and is suggestive of a component of pure shear in the extension of the CVR (Buck et al. 1988). Regardless of where Pn velocities are reached within the crust mantle zone of the CVR, it is the 15-km-deep interface between the sialic upper crust and the intruded/underplated lower crust that is the major seismological, petrological and density boundary in this region. After the lithosphere has cooled and volcanic activity has abated, we might expect a more distinct reflection Moho or Moho

12 480 W. R. Stratford and T. A. Stern zone to form (Griffin & O Reilly 1987; Nicolas 1985; Furlong & Fountain 1986) beneath the CVR. Whether or not such an interface can be identified with modern seismic methods in the contemporary crust mantle zone is, however, questionable. An indicator of the inherent ambiguity in interpreting crust and mantle structure of the central North Island is highlighted by two recent studies using receiver functions, earthquake traveltimes and reflection seismology within the CVR (Bannister et al. 2004). Vs values from 13 receiver functions are used by Bannister et al. (2004) to interpret a 30-km-thick crust. From their receiver functions, a mean S wave speed of 3.5 ± 0.1 km s 1 at 22 km depth is estimated and a Vp/Vs of 1.8 is derived. Such a S wave speed and Vp/Vs ratio would imply a P wave speed of 6.3 ± 0.2 km s 1. However, from earthquake arrival times Harrison & White (2004) interpret Vp/Vs = 2.05 for the depth range of km beneath the CVR. This value for Vp/Vs is more typical of those found in continental rifts (Dugda et al. 2005). A Vp/Vs of 2.05 coupled with the derived S wave speed from Bannister et al. (2004) of 3.5 km s 1 implies a Vp of 7.2 ± 0.2 km s 1 in the km depth range. This is not significantly different from the P wave speed we derive from long-range seismic refraction results presented here. Velocity and depth interpretations differ considerably between these previous studies and the interpretations presented in this paper. This is, in part, a reflection of the differing resolution of the methods used. Receiver functions techniques utilize lower frequency waves and have an inherent trade off between velocity and depth. The low frequencies are an advantage in penetrating the near surface ash, however, in the CVR multiples are generated by low-density, near-surface volcanics (Bannister et al. 2004) and the over all depth resolution of interfaces is lower. Moreover, the trade off between velocity and depth is particularly strong for Vs (i.e. Vp/Vs) compared to Vp and changes in Vp/Vs of only 0.1 can lead to 4 km change in crustal thickness (Zhu & Kanamori 2000). This is especially important in regions of extension and magmagenesis because Vs is effected by partial melts and cracks at about twice the rate of Vp (Hammond & Humphries 2002). The major difference between our P-wave velocity model and that of Harrison & White (2004) is they have P wave speeds reaching 7.4 km s 1 at 30 km whereas we interpret 7.4 ± 0.2 km s 1 at 20 ± 2 km. Their PmP 2 reflections from 30 km depth are modelled as the crust mantle boundary where P wave speeds increase from 7.3 to 7.4 km s 1 (Harrison & White 2004), and Poisson ratio decreases from 0.34 to Using Zoeppritz s equations (Sheriff & Geldart 1995) we calculate that the amplitude ratio of PmP 1 /PmP 2 for the Harrison and White model is 0.5. What is, however, observed is PmP 1 /PmP 2 is 1.5 (Stratford & Stern 2004). Moreover, in this study, upper mantle Pn arrivals with Vp of 7.4 km s 1 have been shown to arrive from depths above the PmP 2 reflector. Our interpretations are based on long-range refraction data where we see the Pn branch on seismographs in the offset range km. Velocity determinations can be made more directly and precisely from refraction analysis, compared to moveout of a wide-angle reflection (Keary & Brookes 1991), so we favour the interpretation of 20 ± 2 km as the depth where the 7.4 km s 1, upper mantle material starts. However, the best test of this model will come from more long-range refraction data, especially a reversal of the north south line (Fig. 5). 5.2 Mantle buoyancy Pulford & Stern (2004) use the rock uplift history provided by mudstone porosity measurements to quantify the buoyancy force in the mantle required to produce the km wavelength, domed surface, and rock uplift of the central North Island (Fig. 2). A density anomaly in the mantle is required not only beneath the thin crust of the CVR, but also the 25-km-thick crust of the western North Island and Northland Peninsula (Stern et al. 1987). Pulford & Stern s (2004) calculations show that a density anomaly distributed over 75 km depth in the mantle of 70 ± 10 kg m 3 under the CVR and 40 ± 10 kg m 3 beneath the western North Island can produce the observed rock uplift. There are two main contributory factors to this reduction in density: thermal expansion due to temperature increase (Turcotte & Schubert 1982) and partial melting of mantle rocks (Lowry et al. 2000). Temperature anomalies (with respect to normal continental conditions) from attenuation studies are estimated at δ350 ± 50 beneath the CVR and δ250 ± 50 beneath the western North Island (Salmon et al. 2004). Applying the equations of thermal expansion (Turcotte & Schubert 1982), we estimate temperature increase contributes about half (40 kg m 3 ) of the density anomaly beneath the CVR and, within error bounds, the whole density anomaly beneath the western North Island. The remaining density reduction required beneath the CVR may be explained by partial melt and melt residuum. We estimate the percent of partial melt by the observed reductions in Pn velocity. In a region of high attenuation like the CVR (Salmon et al. 2004), both the anelastic and anharmonic effects on mantle rheology are important (Wiens & Smith 2003). Wiens & Smith (2003) estimate δvp/δt of 1.0 m s 1 / k for temperature derivatives and based on laboratory studies of melt inclusions (Hammond & Humphries 2002), P-wave velocities decrease at 3.6 per cent per percent melt. For the 7.4 km s 1 Pn velocity of the CVR, a 1.2 per cent melt is estimated, while the km s 1 Pn wave speed of the western North Island can be explained by temperature effects alone. Given these melt percentages, the density drop per percent melt can explain only 20 per cent of the remaining density anomaly beneath the CVR. Melt residuum in the mantle beneath the CVR may also provide an additional component of density reduction Mantle density and rhyolite production Up to km 3 of volcanic material fills the 2-km-deep basement depression of the CVR, and when allowances are made for the pyroclastic deposits outside this region, a total volume of km 3 of erupted material is estimated (Houghton et al. 1995). For every 10 km 3 of sialic volcanic rock found at the surface of the Kenya Rift, five to ten times this volume of parent basalt are required for its production (Latin et al. 1993; Hay et al. 2000). The 5-kmthick layer of underplating beneath the CVR (Fig. 10) is similar in thickness and velocity to the 6.8 km s 1 and 2 5-km-thick lower crustal layer beneath the Kenya Rift (Hay et al. 2000; Latin et al. 1993). As well as thickening the crust, this new layer may play a role in the formation of some of the upper crust, providing a source zone for differentiation of basaltic magmas into the large volume of sialic eruptives seen at the surface (Latin et al. 1993; Hay et al. 2000). At depth, assuming the area of melt producing mantle is the same as the surface expression of extension, and using the relations of Hay et. al and Latin et al., between five and ten times this volume, or a km-thick layer of parent basalt, is expected. Thus the 5-km-thick layer of underplating beneath the CVR (Fig. 10) is not by current estimates thick enough to be the sole producer of the

13 Lithospheric structure of a continental backarc 481 calculated volume of sialic rocks and some additional residue must reside in the upper mantle (Hay et al. 2000). Alternatively, some of the sialic rocks may be produced by anatextic melting of the pre-existing crust (Graham et al. 1990); and by direct wet melting of the mantle above the subduction zone (Tamura & Nakamura 1996), thus reducing the volume of parent basalt needed. However, given that a maximum of 40 per cent of the volume of sialic rock can be fractionated from the 5-km-thick layer of underplating (Hay et al. 2000; Latin et al. 1993), a significant amount of basaltic residue is expected in the upper mantle. This finding is consistent with that from the rock uplift analysis (see Section 5.2) of a low upper mantle density (Pulford & Stern 2004). 5.3 Mantle structure Strong, low-frequency reflections (PmP 2 ) are recorded from an interface within the upper mantle beneath the CVR (Figs 5a and 11). Previous analysis, based on amplitude variation with offset (AVO), of these arrivals (Stratford & Stern 2004) interpreted them as coming from the top surface of a 40-km-wide reservoir of partial melt, which is constrained geographically to be beneath the surface expression of the volcanic area (Fig. 3). However, if the Fresnel zones (Sheriff & Geldart 1995) of the low frequency PmP 2 reflections (dominant frequencies are 8 12 Hz) are also considered, then the lateral extent of a reflector required to produce these arrivals could be as small as 16 km. A 16 km width is 4 times wider than the magma bodies found beneath oceanic spreading centres with higher extension rates ( mm yr 1 ) (Collier & Singh 1990, 1997), but is significantly smaller than some magma bodies found in some continental rift environments where extension rates are much less (Sheetz & Schlue 1992; de Voogd et al. 1986). It is unclear if the PmP 2 reflections on the north south line would have the same interpretation offered above for PmP 2 on the west east line. We do, however, note that the north south and west east line occurrences of these PmP 2 reflectors are beneath the active geothermal fields of Rotorua and Wairakei, respectively. Moreover, deep electrical soundings show low resistivities roughly coincident with the PmP 2 reflector on the west east line (Ogawa et al. 1999). These low resistivities are interpreted as as connected melt in the upper mantle (Ogawa et al. 1999). Therefore, one possible interpretation of the PmP 2 arrivals is that they are coming from the top of a body of 6 per cent partial melt at a depth of 35 ± 3 km (Stratford & Stern 2004). However, for this interpretation only the anharmonic effects of melt on S-wave velocities were taken into account. If the anelastic and anharmonic effects on both melt and temperature are considered (Hammond & Humphries 2002; Wiens & Smith 2003) and the mantle temperature is presumed to be above the wet solidus at 35 km depth, then the S wave drop required to produce the amplitude of the PmP 2 reflection can be explained by <4 per cent melt (Vp/Vs = 2.5). There is, however, no control on the thickness and, therefore, volume of the putative melt body. Why would melt pool at this depth? Stratford & Stern (2004) offer two possibilities that bear on the interpretation of Moho and mantle wedge structures beneath the CVR: Figure 11. (a) Mantle reflection (PmP 2 ) arrivals from shot 3, shooting west on the west east line (refer Fig. 3) from Stratford & Stern Stratford & Stern (2004). PmP 2 reflection is from bounce points beneath the CVR. Distance scale in (a) is shot offset (km). Grey bar shows the range of incidence angles, predicted by Zoeppritz s solutions (Sheriff & Geldart 1995), for strong wideangle reflections (WAR) from a standard Moho model where P wave speed increases from 7.4 km s 1 to 8.1 km s 1. Insert shows Zoeppritz s solutions for the Moho model and a negative impedance contrast model, where S- wave velocities drop. An S-wave velocity drop of 40 per cent is required to produce the observed high amplitudes of PmP 2 in the calculated incidence angle range of the reflections. The S wave drop is inferred to be caused by a layer or 4 per cent partial melt. The PmP 2 or mantle reflector is spatially confined to the CVR by shots on the west east line as labelled in Fig. 3 and shown in (b). (i) By chance a rising diapir of low-density melt on its way to the surface has been snapped in situ. Numerical studies suggest such diapers rise quickly (3 m yr 1 ) or yr to travel from the subduction interface to the surface (Hall & Kincaid 2001). Given that the periodicity of ignimbrite eruptions is of this order (Wilson et al. 1995) it is reasonable that at any given epoch such a feature would be somewhere in the mantle wedge. (ii) At a depth of 35 km there is a marked viscosity change that favours the pooling of melt at this depth. In essence this is arguing for the 35 km depth to be the lithosphere asthenosphere boundary. If so, then the lithosphere has effectively thinned by a β factor of 3 (where β = original thickness/extended thickness) (McKenzie 1978) compared to the apparent crustal extension factor of β 2, assuming an original lithospheric thickness of 100 km (Molnar et al. 1999). This is a reasonable result given that non-uniform pure shear is thought to be a common extension mechanism for the axial regions of continental rift zones (Ruppel 1995). We would then interpret the region between 20 and 35 km depth as being representative of an upper mantle zone where convection is inhibited by increased viscosity. An increase in viscosity would allow some solidification of melts in this zone, thus causing the diffusive reflectivity reported in Section 4.5. Finally, that the proposed melt body is spatially associated with the region of highest crustal extension and volcanics suggests that decompression melting may control not only melt production but melt extraction paths beneath the volcanic area (Condor et al. 2002).