Shear wave velocity variation across the Taupo Volcanic Zone, New Zealand, from receiver function inversion

Transcription

1 Geophys. J. Int. (2004) 159, doi: /j X x Shear wave velocity variation across the Taupo Volcanic Zone, New Zealand, from receiver function inversion S. Bannister, 1 C. J. Bryan 2 and H. M. Bibby 1 1 Institute of Geological and Nuclear Sciences, PO Box 30368, Lower Hutt 6315, New Zealand. s.bannister@gns.cri.nz 2 United States Geological Survey, Vancouver, USA Accepted 2004 June 10. Received 2004 May 31; in original form 2003 July 21 1 INTRODUCTION The Taupo Volcanic Zone (TVZ) of New Zealand (Fig. 1) is a zone of active extension associated with the oblique subduction of the Pacific Plate beneath the Australian Plate and is believed to mark a southward continuation of the backarc extension of the Havre Trough (in the north) into the continental crust of New Zealand (Stern 1985). The TVZ is characterized by extremely active volcanism. More than 90 per cent of the Late Pliocene to Quaternary volcanism known in New Zealand has occurred in this region (Wilson et al. 1995; Hochstein 1995). Magma has been erupted from the central rhyolite-dominated portion of the TVZ at an average rate of 0.3 m 3 s 1 over the last 0.3 Myr making this one of the most active rhyolitic systems on Earth (Wilson et al. 1995). Volcanism within the TVZ has been characterized by caldera-forming rhyolitic eruptions that have been the sources of the numerous ignimbrite flows which dominate the surface geology. Associated with the active volcanism are more than 20 high-temperature geothermal systems that together discharge more than 4000 MW of heat (Bibby et al. 1995), indicating an average heat flow through the region of about 700 mw m 2. The bounds of the TVZ are defined geologically by Wilson et al. (1995) as the envelope of volcanism active within the last 2.0 Myr. 1.1 Geophysical definition of the TVZ Geophysically, the eastern margin of the TVZ is marked by a steep linear gradient in the residual gravity anomaly (Stern 1979) associated with the westward downfaulting of the Mesozoic greywacke SUMMARY The Taupo Volcanic Zone (TVZ), New Zealand is a region characterized by very high magma eruption rates and extremely high heat flow, which is manifest in high-temperature geothermal waters. The shear wave velocity structure across the region is inferred using non-linear inversion of receiver functions, which were derived from teleseismic earthquake data. Results from the non-linear inversion, and from forward synthetic modelling, indicate low S velocities at 6 16 km depth near the Rotorua and Reporoa calderas. We infer these low-velocity layers to represent the presence of high-level bodies of partial melt associated with the volcanism. Receiver functions at other stations are complicated by reverberations associated with nearsurface sedimentary layers. The receiver function data also indicate that the Moho lies between 25 and 30 km, deeper than the 15 ± 2kmdepth previously inferred for the crust mantle boundary beneath the TVZ. Key words: crustal structure, receiver functions, Taupo Volcanic Zone, waveform inversion. basement rocks. Within the TVZ more than 2 km of low-density volcanic infill overlies the basement and the deepest drill-hole (2.7 km deep) encountered no greywacke, only sequences of pumice, ash flows, breccia and ignimbrites erupted from a multitude of centres within the TVZ. Electrical studies (Bibby et al. 1998) suggest that the thickness of the volcanic infill may be as great as 4 km in places. Present-day deformation determined from GPS measurements (Darby & Meertens 1995) and earthquake focal mechanisms (Hurst et al. 2002) indicate northwest southeast extension in the central TVZ. This is also reflected by active surface faulting parallel to the eastern margin (Villamor & Berryman 2001). There has been much debate about the processes that control both the large-scale caldera volcanism and the geothermal heat sources beneath the TVZ (e.g. Wilson et al. 1995). Caldera eruptions require accumulation of magma at a relatively shallow depth prior to eruption. Similarly the magnitude of the geothermal heat flow output suggests a magmatic heat source at relatively shallow depths. Geophysical studies can provide possible indications of magma within the crust. Offshore, in the Bay of Plenty, deep multichannel seismic reflection data show strong bright spot reflections at 5 s two-way traveltime, which have been interpreted to be associated with magma or volcanic sills (Davey & Lodolo 1995). Onshore, recently collected magnetotelluric data indicate high-conductivity bodies at depths of km in the central TVZ (Ogawa et al. 1999), consistent with the presence of connected melt or fluids. Such magnetotelluric data can be complemented by the use of seismic experiments which specifically utilize shear (S) waves. S-wave data can be used to infer the presence of fluids, as S-wave velocities are highly sensitive to fluids (while typical seismic GJI Seismology C 2004 RAS 291

2 292 S. Bannister, C. J. Bryan and H. M. Bibby Figure 1. The Taupo Volcanic Zone (TVZ). The locations of broadband seismometer sites are marked as triangles, with their site names. Major rhyolitic volcanic centres and calderas in the TVZ (after Wilson et al. 1995) are shown in a light-grey shading. The larger Calderas are Ka, Kapenga; Ma, Mangakino; Ro, Rotorua; Rp, Reporoa; Wh, Whakamaru; Ok, Okataina; Mr, Maroa, along with Taupo caldera. The inset shows the location of the TVZ in North Island, New Zealand. refraction and reflection experiments often only utilize P-wave seismic data, and so cannot easily image mid-crustal melt bodies). Melt bodies have been inferred elsewhere using seismic wave data, for example beneath Socorro (Sheetz & Shlue 1992), South America (Chmielowski et al. 1999), Japan (Nakamichi et al. 2002) and Iceland (Darbyshire et al. 2000). 1.2 Previous seismic studies in the TVZ Land-based seismic reflection studies in the TVZ have been largely unsuccessful in imaging below 2 3 km depth, because of energy scattering and reverberation within the highly variable near-surface volcanics (Bannister & Melhuish 1997). In the top few hundred metres the P-wave velocity can be as low as 300 m s 1 in the surface ash layers (Bannister 1993), and about 1100 m s 1 for shallow layers of volcanic tuff (Stagpoole 1994). Early refraction studies (Robinson et al. 1981; Stern & Davey 1987) showed consistent structure on the large scale, with low P-wave velocity ( km s 1 )inthe upper 2 km overlying materials with velocities of about 5.5 km s 1. The P-wave velocity of the greywacke metasediment basement to the east of the TVZ is typically 5.5 km s 1 (Stagpoole 1994). Inside the TVZ, P-wave velocities of km s 1, taken together with drill-hole information, indicate a continuous metasediment basement beneath the thick volcanic sediment and pyroclastic infill material. Stern (1985), however, has pointed out that the basement rocks could also consist of andesite, on the basis of density and P-wave velocity. Detailed studies of the P-wave velocity distribution in the upper crust of the central TVZ show lateral variations in the crustal structure. Sherburn et al. (2003) found P-wave velocity anomalies as large as 10 per cent in the top 4 km of the crust, using 3-D tomographic inversion of earthquake P- and S-wave traveltimes. Low P-wave anomalies were found at depths of 0 4 km beneath the young rhyolitic caldera structures at Okataina, Rotorua and Reporoa. The locations of the anomalies closely correspond to large negative residual gravity anomalies of at least 65 mgal, which probably represent thick sequences of low-density volcanic deposits beneath the calderas (Stern 1982; Rogan 1982; Stagpoole & Bibby 1999). Recent seismic refraction studies, carried out as part of the active-source NIGHT experiment (Stratford et al. 2002; Henrys et al. 2003), indicate an average P-wave velocity of 5.7 km s 1 for the crust in the central TVZ, with the velocity no higher than 6.0 km s 1 at km depth. Similar velocities were also found by the 3-D tomographic inversion work, although P-wave velocities in the mid-crust vary laterally within the region by up to 15 per cent (Sherburn et al. 2003).

3 Seismic refraction data collected in the NIGHT experiment suggest P velocities of km s 1 at a depth of km beneath the TVZ (Stratford et al. 2002). Separate analysis of earthquake Pn and Sn traveltimes indicate Pn- and Sn-wave velocities of 7.4 ± 0.1 km s 1 and 3.95 km s 1 respectively beneath the region (Haines 1979). Such velocities, although low, are close to the upper-mantle velocities found beneath some continental rifts (e.g. Keller et al. 1994). 1.3 Receiver function analysis and low-velocity structures Information on S-wave velocity structure can be derived using the broad-band teleseismic receiver function method (e.g. Langston 1977; Owens & Zandt 1985; Ammon et al. 1990; Ammon 1991). The method aims to eliminate source-related and mantle-path effects, enhancing Ps conversions and reverberations associated with crustal and mantle structure near the receiver. The derived receiver function (RF) (Burdick & Langston 1977) primarily consists of phases associated with discontinuities (such as the Moho) beneath the receivers (e.g. Owens & Zandt 1985). The amplitudes of these phases depend on the angle of incidence of the impinging P wave, the velocity contrast across the discontinuities and the gradient of the velocity changes. The relative arrival times of the converted phases and multiples depend on the depth of the velocity discontinuities, and the P and S velocity structure between the discontinuity and the surface. Radial RFs primarily provide information about changes in the S-wave velocity in the crust and upper mantle, whereas transverse RFs can provide information on anisotropy and help to identify dipping structures. A thorough numerical investigation of effects on receiver functions is described by Cassidy (1992). Receiver functions are often derived from the P-wave coda of teleseismic events using damped spectral division (e.g. Ammon 1991). In this study we utilize a variant, the multiple-taper correlation (MTC) technique of Park & Levin (2000), which computes the RFs in the frequency domain. It also allows estimates of RF uncertainties, which become useful when stacking RFs from different seismic events. The relationship between the derived receiver functions and the structure of the Earth beneath any receiver is non-linear, as the RFs are dependent on changes in S-wave velocity (Vs), Poisson ratio, density and attenuation. It is necessary to use non-linear inversion techniques to solve for the variation of S velocity and Poisson ratio with depth to avoid dependency on any assumed starting velocity model. The range of velocity depth profiles which can satisfy observed RF data can be derived using, for example, the non-linear neighbourhood algorithm (NA) of Sambridge (1999a,b), which is a fully non-linear approach involving an ensemble-based Monte Carlo search technique. Stochastic sampling is used in NA to search a multidimensional parameter space for the range of acceptable velocity models, using geometrical constructs known as Voronoi cells to assist in the search and appraisal stages. Details of the background behind the technique, together with examples, are given by Sambridge (1998, 1999a,b). The inversion procedure involves computing synthetic RFs for more than alternative 1-D models, calculating waveform misfits with an appropriate objective function, and determining a set of acceptable velocity models which fit the observed RF data to a prescribed level. Interpretation of the Moho discontinuity from velocity depth profiles derived from RF inversion is often made by identifying the transition from crustal to typical mantle velocities. However, receiver functions are primarily sensitive to the gradients in the shear wave velocity rather than to absolute velocity, and trade-offs Shear wave velocity structure, Taupo Volcanic Zone 293 can occur between absolute velocities and the depth of an interface (Ammon et al. 1990). We use the absolute velocity only as a rough indicator that mantle-like velocities have been reached, and focus on the nature of the velocity gradients in the model ensemble. Separately we note that upper-mantle velocities beneath the TVZ region are actually quite low (Haines 1979). These low velocities are likely to extend to a depth of at least 80 km (Stratford & Stern 2004). In this paper we examine teleseismic data recorded by a broadband seismic array deployed in the TVZ, to constrain the local S- wave velocity structure, using non-linear inversion of teleseismic receiver functions. 2 RECEIVER FUNCTION DATA AND METHODS We examine P-to-S (Ps) converted seismic waves generated at the Moho (and at other velocity discontinuities within the crust) from teleseismic earthquake waves. The waveform data were recorded by broad-band seismometers that were temporarily deployed for 5 months in the central TVZ during 1995, as part of the TVZ95 experiment (described in full by Bryan et al. 1999). The recorded events included local shallow crustal earthquakes, Benioff-zone earthquakes in the underlying subducted Pacific Plate and the teleseismic earthquakes used in this RF study. The broad-band seismometers straddled the TVZ, with station spacings varying between 10 and 25 km (Fig. 1). Suitable teleseismic earthquakes (Fig. 2) were selected by searching the IRIS-DMC catalogues for earthquakes with epicentral distances between 25 and 100, and magnitudes greater than 5.5. This distance range was used to avoid problems with regional and core phases. Only a few suitable events were recorded, primarily because of the relatively short duration of the broad-band seismometer deployment. Some earthquakes to the northeast of the TVZ were rejected because the observed and expected backazimuths of the initial P-wave arrival were notably different, probably due to the effect of refraction along the Kermadec and Hikurangi subduction zones (e.g. Ansell & Gubbins 1986; Galea 1992). The term backazimuth used throughout this study refers to the direction from the station towards the earthquake epicentre, measured positive clockwise from north. The locations of the selected events, and the range in backazimuths and incident angles of the events, are detailed in Table 1. Event locations are illustrated in Fig. 2. Nearly all of the events lie in the northwest quadrant, and the expected angles of incidence at the base of the crust range between 18 and 40. The first 40 s of the teleseismic P-wave coda was used to determine each individual RF, using the method of Park & Levin (2000). Once determined, individual RFs were then combined into composite RFs in bins, with 5 spacing between bins and with 10 overlapping intervals of backazimuth. The spread of event epicentral distance for any one bin was small. Such averaging means that individual RFs intentionally contribute to the composite RFs in two adjacent bins (Park & Levin 2000). Each individual RF was weighted by its inverse variance during bin formation. Such weighting significantly improved the bin RFs. We determined RFs for each seismic event using several different low-pass frequency filters, so as to subsequently examine frequency dependences in the observed RFs (e.g. Owens & Zandt 1985; Cassidy 1992). The analyses discussed below have an effective frequency cut-off of roughly 0.8 Hz, unless specifically mentioned otherwise.

4 294 S. Bannister, C. J. Bryan and H. M. Bibby Figure 2. Hypocentres of all teleseismic earthquakes used in this study (left), and plotted as a function of the event backazimuth and the P-wave incident angle at the base of the crust (right). The term backazimuth refers to the direction from the station towards the earthquake epicentre, measured positive clockwise from north. Table 1. Parameters for the teleseismic earthquakes used in this study. Event Date Time Lat. Long. Depth Delta Backaz. Slowness (yr/month/day) (km) ( ) ( ) (skm 1 ) /02/13 15: S E /03/12 12: S E /03/19 23: S E /03/25 22: S E /03/26 02: S 28.2 W /03/31 14: S E /04/07 22: S W /04/13 02: S E /04/17 01: S E /04/17 23: N E /04/18 03: S E /04/20 08: N E a 1995/04/21 00: N E b 1995/04/21 00: N E c 1995/04/21 00: N E d 1995/04/21 05: N E Features of the observed RFs are described below for each of the broad-band stations. Figs 3, 4, 5 and 6 show binned receiver functions for different receiver stations across the TVZ, plotted as a function of backazimuth. The RF bins lie primarily in the northwest quadrant, although a single event observed at several stations from backazimuth 24 contributes information in the northeast quadrant. Note that the averaging procedure used in bin formation places this single event into two adjacent backazimuth bins (as described above). 2.1 Stations within central TVZ Radial and transverse receiver functions from stations TAWV, PDVV, HIHV, PKMV, and BTRV, all within the central part of TVZ, are shown in Figs 3 and 4. Receiver functions derived for stations RKIV and SPCV were too poor for use in this study, and are not shown. A phase delay of 0.5 s is observed for the initial P pulse on the radial RFs for some of the stations (e.g. TAWV, BTRV). Delay of the pulse (from zero time) indicates low velocities in the near-surface layer beneath the station. This delay results from the superposition of the direct P wave and P-to-S converted (Ps)waves generated at the basement-sediment interface, and is often apparent for stations lying above low-velocity sediments, such as present in the TVZ. Similar delays have been observed, for example, in receiver function studies in Japan (Nakamichi et al. 2002) and Iceland (e.g. Darbyshire et al. 2000). The effect of such surface layers is demonstrated in Fig. 7, using forward synthetic reflectivity calculations. Synthetic radial RFs were calculated for a range of 1-D horizontally-stratified velocity models, and a ray parameter of s km 1, (corresponding to an epicentral distance of 60 ). The models all involve a simple two-layer crust, with a crustal thickness of 16 km, and P-wave velocities of 6.2 km s 1 and 7.4 km s 1 in the crust and mantle respectively. This model was modified to include a low-velocity layer at the surface, with a P-wave velocity of 3.2 km s 1, and a thickness varying between 0.5 and 3 km (Fig. 7). A Poisson ratio of 0.25 was assumed for all layers. The calculated synthetic RF for the simple two-layer crust shows a clear Moho Ps phase at 2 s (Fig. 7a). The synthetic RF changes when a low-velocity surface layer is introduced for the same background crustal model (Fig. 7b). As the thickness of the sedimentary layer is increased further (Figs 7b g) the effect of the sedimentary layer on the RF also increases. Strong pulses

5 Shear wave velocity structure, Taupo Volcanic Zone 295 Figure 3. Radial (R) and transverse (T) receiver functions for central TVZ stations TAWV, PDVV, HIHV and PKMV. Note the strong negative pulse at around 1 2 s time. The composite RFs shown here were formed by binning unmigrated individual RFs, weighted by their inverse variance (following Park & Levin 2000). develop at 1 2 s, and the initial pulse becomes complex, with an apparent phase delay. These calculated effects become more marked if the velocity of the surface layer is reduced further, from 3.2 km s 1 down to 2.5 km s 1 for example. Such low P-wave velocities have been determined in the central TVZ beneath some calderas (Sherburn et al. 2003). It is apparent from comparing Figs 7(a) (c) and (g) that conversions associated with the surface sedimentary layers may potentially overwhelm the arrivals from Ps conversion at the Moho discontinuity. Negative pulses with large amplitudes are observed 1 2 s after the initial pulse on the radial RF bins at stations TAWV, PDVV HIHV and PKMV (Fig. 3). These negative pulses are very clear, for example, on the radial RFs for station PDVV (situated on the southern edge of Rotorua caldera, Fig. 1) and for RFs from station HIHV (situated on the northern edge of the Reporoa caldera, Fig. 1). The RFs for station HIHV show little to no time delay for the initial 0-s pulse, indicating that the near-surface layers beneath this station do not have a low velocity, and that the RFs are not strongly

6 296 S. Bannister, C. J. Bryan and H. M. Bibby Figure 4. Radial (R) and transverse (T) receiver functions for central TVZ station BTRV, plotted against backazimuth. There is only one contributing event for the northeastern bins. Figure 5. Radial (R) and transverse (T) receiver functions for stations TWHV and KAIV, which lie to the east of the TVZ. The RFs for both stations are derived from only one contributing event, from a northeast backazimuth. The initial Pp pulse on both radial RFs is slightly delayed from 0stime, due to the effect of low velocities in the near-surface layers. influenced by the surface layers. We also note that the energy levels on the transverse RFs for HIHV are small compared with the energy on the radial RFs, implying that the structure is reasonably 1-D beneath HIHV and that an alternative explanation is required to explain the strong negative pulses at 1 2 s. We now investigate the possibility that these strong negative pulses represent Ps conversions from a mid-crustal discontinuity, in particular the top interface of a low S-wave velocity layer (LVL). Such a LVL may also generate multiple reverberations, noticeable at 2 10 s in many of the RFs, from surface reflections of the incident P-wave which reflect and convert at the top interface of the LVL. The effect of such a low-velocity layer can be demonstrated using forward synthetic reflectivity calculations. Synthetic radial RFs were calculated as described above, for models involving a crustal thickness of 16 km, and P-wave velocities of 6.2 km s 1 and 7.4 km s 1 in the crust and mantle respectively. Fig. 8(a) shows the synthetic radial RF calculated just using this crustal model; the arrival near 2 s is the Ps conversion from Moho. In comparison, Fig. 8(b) shows the synthetic radial RF for a velocity model involving the same crustal thickness but including a thin, 0.2 km thick, low-velocity layer (with an S velocity of 1.7 km s 1 ) placed at a depth of 8 km. The presence of the LVL changes the RF. The response changes further as the layer thickness is increased, to 1 km thickness (Fig. 8d), 2kmthickness (Fig. 8e) and 4 km thickness (Fig. 8f). As the LVL thickness is increased the Moho Ps conversion becomes obscured by the conversions from the top and bottom of the LVL (e.g. Figs 8d f). The majority of the arrivals between 3 and 10 s are likely to represent multiples, reverberating between the surface and the top (and bottom) of the LVL. The characteristics of the synthetic RFs shown in Fig. 8 are quite similar to the RFs derived for some of the stations (e.g. PKMV, shown in Fig. 8g), with negative pulses at 1 s,and other multiple arrivals between 2 6 s.

7 Shear wave velocity structure, Taupo Volcanic Zone 297 Figure 6. Radial (R) and transverse (T) receiver functions for western TVZ stations MAMV and CAMV, plotted against backazimuth. The RF for CAMV has only one contributing event. Figure 7. Synthetic receiver functions calculated for velocity models involving a near-surface low-velocity layer. (a) Moho at 1-km depth, with crustal P-wave velocity of 6.2 km s 1 and no surface layer. (b) (g) Moho at 16 km depth, with a low-velocity surface layer (P-wave velocity 3.2 km s 1 ), varying in thickness from 0.5 km (b) to 3 km (g). A ray parameter of s km 1 was used for the synthetic calculations. Figure 8. Synthetic receiver functions calculated for velocity models involving (a) a 16 km thick crustal layer, with a P-wave velocity of 6.2 km s 1 and a mantle velocity of 7.4 km s 1.Aray parameter of s km 1 was used for the synthetics. (b) (f) The same crustal model, but including a thin low-velocity layer (LVL) in the mid-crust, with a P-wave velocity 3.0 km s 1. The top interface of the LVL was positioned at 8 km depth, and the thickness of the LVL was varied between (b) 0.2 and (f) 4.0 km. (g) The observed RF at station PKMV, for comparison with the synthetics. These examples provide an indication of the range of expected RF response for different models. The presence of any LVL in the midcrust has the greatest influence on the RF. Secondary changes in the RFs occur with changes in the thickness of the LVL (depending on the wavelength of the incoming teleseismic wave), and (although not demonstrated here) the nature of the change in the physical properties (e.g. P- and S-wave velocities) across the top and bottom boundaries of the layer. The inferred depth of the mid-crustal discontinuity associated with the negative pulse is dependent on the S-wave velocity structure and the Poisson ratio in the upper crust. 3-D tomographic analysis of local earthquake arrival times in this region (Sherburn et al. 2003) indicates a P-wave velocity between 3.5 and 4.2 km s 1 in the surface layers beneath station HIHV, a P-wave velocity gradient between 2 and 5 km depth and P-wave velocities ranging between 5 and 5.8 km s 1 between 5 and 7 km depth. Using these velocities,

8 298 S. Bannister, C. J. Bryan and H. M. Bibby and assuming a Poisson ratio of 0.25, the inferred depth of the discontinuity beneath station HIHV is 6 8 km. Similar negative pulses are observed on the radial RF bins at several of the other stations, although the pulses vary in their amplitude and relative time. The initial P phase at some of the other stations are delayed (Fig. 3), which suggests that the RFs at these other stations may be affected to some degree by the presence of near-surface sedimentary layers. Teleseismic event data were examined for all central TVZ array stations. However, the data for some of the stations were considered too complex for analysis and interpretation, often involving considerable energy on transverse components. The use of such data would not be appropriate for receiver function analysis and interpretation, given the underlying assumptions of 1-D structure, as the presence of high levels of energy on the transverse components implies crustal heterogeneity beneath the station, lateral variations in structure (e.g. dipping layers) or anisotropic effects. 2.2 East of the TVZ, stations TWHV and KAIV Fig. 5 shows radial and traverse receiver functions derived for stations KAIV and TWHV. There is only one contributing event for the RFs for both of these stations, which is from a backazimuth of 24. Both of these stations were situated to the east of the TVZ, on the Kaingaroa Plateau, where the metasediment basement is known to lie only a few hundred metres deep. The radial RF for station TWHV is not particularly well defined. The large peak near 0 s on the radial RF results from the direct P-wave arrival at the receiver. This initial peak is delayed from zero time by about 0.3 s, indicating the presence of low-velocity surface sediments beneath the station. Secondary pulses are observed at 4.5 and 7.5 s. On the transverse RF, however, there is considerable energy near 0 s, comparable to the energy observed on the radial RFs. The transverse energy possibly results from the incoming energy arriving from directions other than the expected backazimuth. Significant energy on transverse RFs can be an indication of anisotropy (Savage 1998), dipping structure (e.g. Langston 1977; Peng & Humphreys 1997) or other complications such as near-receiver scattering (e.g. Bannister et al. 1990). Multichannel array studies near TWHV, on the Kaingaroa Plateau, suggest that scattered waves result from heterogeneity in the near-surface volcanic layers (Bannister & Melhuish 1997). We do not have the necessary RF data (across a range of backazimuths) to distinguish between the above possibilities. In comparison with TWHV, the RF data from station KAIV are quite clean, with much less energy on the transverse RF compared with the radial RF. The data show relatively subdued phases following the initial direct pulse, suggesting gentle S-wave velocity gradients rather than sharp discontinuities. 3 INVERSION FOR S -WAVE VELOCITY DEPTH PROFILES We have derived velocity depth profiles of S-wave velocity structure using the non-linear neighbourhood algorithm (NA) of Sambridge (1999a,b), which is a fully non-linear approach, involving an ensemble-based Monte Carlo search technique to find the set of velocity models which best satisfy the objective function. We use as our objective function a scaled l2-norm of the difference between the calculated RF (predicted for any one earth model) and the observed RF. The crustal structure was parametrized for the inversions using 24 parameters, four parameters for each of six layers, that describe the S velocity at the top and bottom of each layer, the layer thickness (in km), and the ratio of P to S velocity (Vp/Vs) for that layer. The two shear wave velocity parameters in each layer allow definition of a velocity gradient for that layer, which allows representation of a large number of potential velocity depth distributions. Very loose a priori constraints were placed on the minimum and maximum S-wave velocities, the layer thicknesses, as well as on Vp/Vs. These constraints act as scale factors in the inversion to non-dimensionalize the parameter space (Sambridge 1999a,b). The velocity model parametrization is shown in Fig. 9. Velocities in the 2.3 West of the TVZ, stations MAMV and CAMV Fig. 6 shows the radial and transverse RF bins for stations MAMV and CAMV, which are located west of the TVZ (Fig. 1). The radial RFs for station MAMV vary with the backazimuth, but secondary pulses have relatively low amplitude, with only a slight negative pulse at 2 s, and a secondary positive phase at 3to4s.There is only one contributing event for station CAMV; the binning procedure places this event into two adjacent backazimuth bins. As for MAMV the derived RF for CAMV is subdued, with very little evidence of strong phases after the first 3 4 s. Figure 9. The velocity model parametrization used in the NA inversion. Each model is defined using six layers, each of which has four variables: layer thickness (H), shear wave velocity at the top and bottom of the layer (Vs1 and Vs2) and the ratio between the P-wave and S-wave velocity (Vp/Vs) for that layer.

9 Shear wave velocity structure, Taupo Volcanic Zone 299 Figure 10. Seismic S-wave velocity models (top) and radial receiver functions (bottom), derived for station HIHV from non-linear inversion. All models searched in the inversion are shown by the outline of the light-grey shading. The shaded regions indicate the density of the 1000 best models, the darkness being logarithmically proportional to the number of models, as shown by the scale bar. The solid white line represents the mean of the best 100 models, and the dotted line represents the best of those models. Results are shown for backazimuths (a) and (b) Lower diagrams show the observed (solid line) and predicted (dotted line) RFs for each backazimuth range, low-pass filtered with an effective frequency cut-off of (a, b) 0.5 Hz and (c, d) 0.8 Hz. The predicted RF waveforms plotted for each backazimuth were derived using the best-fitting velocity solution. layers are assumed to be isotropic, and the layers are assumed to be flat-lying. The clear underlying assumption is that the structure beneath each receiver is 1-D, can be approximated by plane layers, and that there is only minimal side scattering from local crustal heterogeneity. Binned subsets of the radial RF data (Figs 3, 4, 5 and 6) (strictly limited in backazimuth and epicentral distance) were used as the input (observed) data for the inversion. Events in any single bin subset varied no more than 20 in backazimuth and 15 in epicentral distance, although they usually varied much less than these maxima. Optimally, given a higher density of events, these bin sizes should be even smaller (e.g ), to avoid attenuating Ps phases (e.g. Cassidy 1992). We applied the NA inversion only to the best of the RF bins, i.e. those RFs which clearly had relatively low levels of energy on the transverse RF (relative to the radial RF). Any phases caused by

10 300 S. Bannister, C. J. Bryan and H. M. Bibby crustal heterogeneity or anisotropy will not be modelled correctly, given the inherent 1-D assumptions in the modelling. As a second quality criterion we also examined the consistency in the RFs between different backazimuths; large changes in the RFs with small changes in backazimuth imply large lateral variations in crustal structure beneath the receiver. We discuss the consistency in the results below, for each station. The tunable parameters required by the NA (Sambridge 1999a,b) were set after a number of trials. Each inversion run involved 1000 iterations, generating velocity models. Stability of the inversion solutions was tested using a large range of initial random seeds, incidence angles, and velocity model parametrizations. 3.1 Sites within TVZ Station HIHV Station HIHV was situated inside the TVZ (Fig. 1) on the northern edge of Reporoa caldera. In Fig. 10 we present plots of the S-wave velocity model ensembles generated in the inversion, for two backazimuth ranges, , and The lateral extent of the darkest grey areas in the figure provides an indication of how well the velocity structure is constrained by the data. Only a certain number (typically about the best ) of the generated models give an acceptable level of data fit to the observed RF waveforms; we Figure 11. Seismic S-wave velocity models (top) and radial receiver functions (bottom), derived for station TAWV from non-linear inversion. The solid white line represents the mean of the best 100 models, and the dotted line represents the best of those models. Figure details are as shown in Fig. 10. Results are shown for backazimuths of (a, c) and (b, d)

11 Shear wave velocity structure, Taupo Volcanic Zone 301 plot the misfit-weighted mean of the best 100 models in the figure. For the full ensemble of models there is also a single best-fit model that generated the least misfit between the calculated and observed RFs, which is plotted in Fig. 10 as a dotted line for each solution ensemble. However, the useful information is really provided by the full ensemble of solutions, as discussed in detail by Sambridge (1999a,b). The synthetic radial RF derived using the best-fit model from the non-linear inversion is shown in Fig. 10 for the two RF bins, together with the observed RF, low-pass filtered with an effective frequency cut-off of both 0.5 and 0.8 Hz. The best 100 members of the ensembleofsolutions all have similar levels of fit, and so the waveform fit in Fig. 10 can be regarded as representative of the waveform match that can be achieved. The waveform fit differs slightly for different low-pass frequency filters, for the same velocity model. The inversion results for station HIHV (Fig. 10) show S velocities of 2 3 km s 1 near the surface. At 7 12 km depth the inversion indicates a low-velocity layer (LVL) (Fig. 10), with a S-wave velocity as low as 2.0 km s 1. The Vp/Vs ratio derived from the inversion is relatively high in the same depth range as the LVL. Below the LVL there is a gentle positive velocity gradient, down to 30 km depth. Mantle velocities are reached at around km, although the transition is not well defined. The waveform match between the synthetic and observed radial RF is reasonable for both backazimuth bins, although the amplitude of the large negative pulse at 1 sin the observed RF is not Figure 12. Seismic S-wave velocity models (top) and radial receiver functions (bottom), derived for station BTRV from non-linear inversion. The solid white line represents the mean of the best 100 models, and the dotted line represents the best of those models. Figure details are as shown in Fig. 10. Results are shown for backazimuths of (a, c) and (b, d)

12 302 S. Bannister, C. J. Bryan and H. M. Bibby fully matched for either backazimuth, suggesting that the LVL is not fully represented by the ensemble of inversion models. Synthetic modelling (not shown) suggests that a thinner LVL layer, with S velocity less than 2 km s 1,may improve the waveform fit. Our parametrization in the inversion specifically constrained the S-wave velocities to be above 1.6 km s 1 at mid-crustal depths. Any relaxing of that constraint, for example to allow S-wave velocities of 1.0 km s 1 in the mid-crust, would lead to a gross level of nonlinearity, and instability for the inverse problem. The P-wave velocity structure in the crust near station HIHV has been previously derived by tomographic inversion of earthquake traveltimes (Sherburn et al. 2003). That study indicates P-wave velocities between 3.5 and 4.2 km s 1 in the top 2 km, below which were P-wave velocities varying from 5 km s 1 at 5 km depth to 5.8 km s 1 at 7 km depth (Sherburn et al. 2003). No low-velocity layer was delineated by their study at greater depth, although lowvelocity bodies are usually not well resolved with such traveltime tomography studies because of the tendency of seismic waves to refract around low-velocity zones Stations TAWV and BTRV Stations TAWV and BTRV both lie near the eastern margin of the TVZ. Results of analyses for the two stations are shown in Figs 11 and 12 respectively. The inversion results for station TAWV show a S-wave velocity of less than 2 km s 1 at the surface, increasing to between 3.0 and 3.5 km s 1 at 6 km depth. The low velocity of this surface layer leads Figure 13. Seismic S-wave velocity models (top) and radial receiver functions (bottom), derived for station PDVV from non-linear inversion. The solid white line represents the mean of the best 100 models, and the dashed line represents the best of those models. Figure details are as shown in Fig. 10. Results are shown for backazimuths of (a, c) and (b, d)

13 Shear wave velocity structure, Taupo Volcanic Zone 303 to the apparent time delay in the initial pulse on the radial RF (Fig. 11). The best-fitting models involve low velocities at km depth, with slower velocities indicated for backazimuth. A velocity gradient is indicated between 14 and 20 km depth, and mantle velocities are reached at a depth of km although the transition is poorly defined. The two radial RF bins for station TAWV (Fig. 11) show high amplitudes over the first 5 10 s, probably representing the combined reverberation effect of the mid-crustal LVL and the effect of the surface sediments. It is probable that the reverberations overwhelm the Moho Ps conversion. The calculated synthetic RFs match the observed RFs fairly well for both backazimuth bins. Similar results are obtained for station BTRV (Fig. 12), which lies to the southwest of TAWV (Fig. 1) and close to the eastern margin of the TVZ. The observed radial RFs for backazimuth are quite complex, and indicate high levels of reverberation. Derived S-wave velocities for that backazimuth bin indicate S-wave velocities less than 3.0 km s 1 in the top few kilometres, increasing to km s 1 between 4 and 8 km depth. A broad low-velocity layer is indicated at km depth, and mantle velocities are reached at around km. The solutions for backazimuth show higher velocities at km depth, with a thin LVL only weakly defined. The waveform fit for both backazimuth bins is quite good, especially in the first 5 8 s, matching the amplitude of some of the negative pulses (e.g. at 5 s). Figure 14. Seismic S-wave velocity models (top) and radial receiver functions (bottom), derived for station PKMV from non-linear inversion. The solid white line represents the mean of the best 100 models, and the dashed line represents the best of those models. Figure details are as shown in Fig. 10. Results are shown for backazimuths of (a, c) and (b, d)

14 304 S. Bannister, C. J. Bryan and H. M. Bibby Station PDVV Station PDVV was situated inside the TVZ (Fig. 1), near the southwest edge of Rotorua caldera. The observed radial RFs (Fig. 3) have a large-amplitude negative pulse at s, as observed for HIHV and BTRV. Generally the radial RF bins for PDVV show considerably less reverberation than for BTRV; amplitude levels at 5 10 s are much less than the amplitude of the initial pulse. The inversion solutions for PDVV, shown in Fig. 13, indicate slow S-wave velocities in the near-surface layer, with velocities less than 2.0 km s 1 in the top 1 km. Results for the two backazimuth bins differ below 6 km depth. Results for backazimuth indicate a velocity decrease below 6 8 km depth, down to 26 km, although the S velocity still remains above 2.0 km s 1. The Vp/Vs ratio increases temporarily, reaching a maximum at km depth. Results for backazimuth indicate much lower velocities at km depth. Mantle velocities appear to be reached at km depth. The waveform match between observed and synthetic RFs is reasonable for both RF bins, although the maximum amplitude of the observed negative pulse is not fully matched for either backazimuth Station PKMV Inversion results for station PKMV (Fig. 1) are shown in Fig. 14 for backazimuths , and The velocity solutions for both backazimuths indicate low-velocity sediments in the top 2 km, and velocities between 3.0 and 3.7 km s 1 below that, down to 7 9 km. A broad low-velocity zone is required between 9 18 km depth, with an S-wave velocity as low as 2 km s 1. The solutions for the two backazimuths are quite similar, although the LVL appears slightly shallower, at 8 16 km depth, for backazimuth. Beneath the LVL the S velocity increases with a gentle gradient, and levels off to mantle velocities at km depth. The waveform fit is good for backazimuth, matching the amplitude of the negative pulse at s. The fit for is only fair, with mismatch of amplitudes for several phases. 3.2 East of the TVZ, stations KAIV and TWHV Results from the non-linear inversion for station KAIV are shown in Fig. 15, for backazimuth The velocity solutions indicate Figure 15. Seismic S-wave velocity models (top) and radial receiver functions (bottom), derived for station KAIV from non-linear inversion. The solid white line represents the mean of the best 100 models, and the dashed line represents the best of those models. Figure details are as shown in Fig. 10. Results are shown for backazimuth The observed and predicted RFs are shown filtered with an effective low-pass filter cut-off of (a) 0.5 Hz and (b) 0.8 Hz.

15 S-wave velocities of less than 2 km s 1 at the surface, with a velocity step at 0.5 km depth. The low velocity of this surface layer leads to the small apparent time delay ( 0.3 s) in the initial pulse on the radial RF (Fig. 15). Below the surface layer the velocities increase to km s 1 by 5kmdepth, and then decrease slightly below 10 km depth, with a minimum velocity of 2.8 km s 1 at km depth. There is a steady increase in the S-wave velocity between 18 and 26 km depth, with S-velocities greater than 4.0 km s 1 below about 35 km. The observed RF and the synthetic radial RF is shown in Fig. 15 for the best of the inversion solutions. The fit between the waveforms is quite reasonable in the first few seconds, although the amplitudes of the observed peaks at 3 and 5.5 s is not fully matched by the synthetics. Results for station TWHV, to the east of KAIV, are shown in Fig. 16, for the backazimuth range Earlier, we noted the high level of energy observed on the transverse RFs for this station (Fig. 5), which indicates structural heterogeneity or anisotropic effects beneath the station. Although we carried out inversion on the Shear wave velocity structure, Taupo Volcanic Zone 305 RF data, we note that the complexity of the structure beneath the station is likely to lead to poor inversion solutions. The solution ensemble indicates a S-wave velocity of 2kms 1 at the surface, increasing to km s 1 at 3 km depth. The velocities fluctuate down to 30 km depth, and then increase to mantle velocities at km depth. However, we note that the waveform fit derived for even the best models is only fair, matching the initial Ps phase, as well as the time of the secondary phase at 4s,but poorly matching the secondary peaks at 2 3 and 6 8 s. Several inversion runs were carried out using a range of incidence angles, and a range of model constraints, in an unsuccessful attempt to improve this waveform fit. The poor waveform fit may well be due to the scattering effects implied by the observed energy on the transverse RFs. 3.3 West of the TVZ, stations MAMV and CAMV Inversion results for stations MAMV and CAMV, situated to the west of the TVZ (Fig. 1), are shown in Figs 17 and 18 respectively. Figure 16. Seismic S-wave velocity models (top) and radial receiver functions (bottom), derived for station TWHV from non-linear inversion. The solid white line represents the mean of the best 100 models, and the dashed line represents the best of those models. Figure details are as shown in Fig. 10. Results are shown for backazimuths of The observed and predicted RFs are shown filtered with an effective low-pass filter cut-off of (a) 0.5 Hz and (b) 0.8 Hz.

16 306 S. Bannister, C. J. Bryan and H. M. Bibby Figure 17. Seismic S-wave velocity models (top) and radial receiver functions (bottom), derived for station MAMV from non-linear inversion. The solid white line represents the mean of the best 100 models, and the dashed line represents the best of those models. Figure details are as shown in Fig. 10. Results are shown for backazimuths of (a, c) and (b, d) The solutions for the two stations are similar. Velocities for station MAMV increase quite rapidly, from km s 1 in the near surface, to km s 1 at 8 10 km depth. Slightly lower velocities, as low as 3.0 km s 1, are present between 15 and 25 km depth, while mantle velocities are reached at between 26 and 30 km depth. The waveform fit is only fair (at best) for backazimuth, perhaps suggesting out-of-plane propagation effects. Velocities derived for station CAMV are similar, with high S velocities in the top 8 km, and slower S-wave velocities, km s 1, between 12 and 25 km depth. The waveform fit for this station is reasonable only in the first 3 5 s, as the amplitudes of later phases are not well matched. 4 DISCUSSION 4.1 TVZ crustal structure An illustrative cross-section of shear wave velocities along a northwest southeast profile (AA of Fig. 1) is shown in Fig. 19. As receiver functions are more sensitive to changes in the velocity with depth than to absolute velocities the vertical derivative of the velocities is also shown in Fig. 19(b), to highlight the depths of velocity contrasts. The cross-section was formed by interpolation between the weighted mean of the best velocity depth inversion solutions, for each station. Inversion solutions derived from

17 Shear wave velocity structure, Taupo Volcanic Zone 307 Figure 18. Seismic S-wave velocity models (top) and radial receiver functions (bottom), derived for station CAMV from non-linear inversion. The solid white line represents the mean of the best 100 models, and the dashed line represents the best of those models. Figure details are as shown in Fig. 10. Results are shown for backazimuths of The observed (solid line) and predicted (dotted line) RFs are shown, low-pass filtered with an effective frequency cut-off of (a) 0.5 Hz and (b) 0.8 Hz. backazimuth bins were used, when available. Where bins with this backazimuth were not present (e.g. stations KAIV and CAMV), the available solutions were used. Although the data do vary with backazimuth the primary features of the crustal structure remain, and are consistent between sites, as shown in Fig. 19. High positive velocity gradients in the top 1 2 km are associated with the base of the sedimentary layers across the region (Fig. 19b). A pronounced low-velocity layer, LVL1 (Fig. 19a,b), is apparent in the mid-crust between stations HIHV and BTRV, the top interface of which stands out as the large negative velocity gradient at km depth in Fig. 19(b), across which the S velocity drops by at least 1 km s 1.A negative velocity gradient of smaller magnitude is also apparent at shallower depth, at 6 8 km, between stations PDVV, PKMV and HIHV (LVL2 in Fig. 19a,b). 4.2 Interpretation of the low-velocity layers Investigations of the earthquake depth distribution in the TVZ indicate that the base of the seismogenic zone within the TVZ is shallow (Bryan et al. 1999), with more than 80 per cent of well-located earthquakes occurring above 6 km depth. The base of this seismogenic zone, where temperatures are inferred to be C, has been interpreted to mark both the boundary between the convective and conductive heat transport systems, and the brittle ductile transition (Bibby et al. 1995; Bryan et al. 1999). The low-velocity layers in Fig. 19 thus lie beneath the inferred convective fluid transport (geothermal) system, as well as below the brittle ductile transition. The top of the shallowest LVL ( LVL2 of Fig. 19), is also close to the expected minimum depth of magma chambers required on the basis of petrology (Wilson et al. 1984). Given the high heat flux, and the volcanic history of the TVZ, we infer that the LVLs found beneath the TVZ stations (LVL1 and LVL2 of Fig. 19) represent bodies of partial melt. These bodies appear to have a lateral extent in the order of km. The physical parameters of the LVLs are relatively weakly constrained because of the small number of data. Improved density in the data sampling (especially in wave slowness) is required to improve these constraints. This interpretation of the LVLs is consistent with the presence of several high-conductivity zones found in the central TVZ by magnetotelluric (MT) measurements (Ogawa et al. 1999), the shallowest of which lies at a depth of km. Ogawa et al. (1999) interpreted these high-conductivity zones as indicative of the presence of connected melt.

18 308 S. Bannister, C. J. Bryan and H. M. Bibby Figure 19. Top: An illustration of the shear wave velocity structure along a northwest southeast cross-section across the TVZ (AA of Fig. 1), formed by interpolation between the weighted mean of the best velocity depth inversion solutions, for each station. Bottom: The derivative of Vs with depth, along the same cross-section, highlighting where the S-wave velocity changes rapidly. Circles show the projections of earthquake hypocentres located by Sherburnet al. (2003) within 5 km of the profile. Similar magma bodies have been inferred elsewhere on the basis of receiver function data, for example beneath Socorro (Sheetz & Shlue 1992), South America (Chmielowski et al. 1999), Japan (Nakamichi et al. 2002) and Iceland (Darbyshire et al. 2000). 4.3 Crustal thickness The RF inversion solutions indicate that mantle velocities are reached at a depth of between 25 and 30 km beneath the TVZ (Fig. 19a). The implied crustal thickness appears to vary within the TVZ, thinning to 25 km toward the eastern boundary of the central TVZ, but it is still greater than the 15± 2 km thickness determined by Stern & Davey (1987) from P-wave seismic refraction data. More recent seismic refraction and reflection data from the NIGHT experiment (Henrys et al. 2003) suggest the presence of a crust mantle transition layer in the depth range of km, with P velocities of km s 1 (Stratford et al. 2002). Our RF inversion solutions show a positive S-wave velocity gradient at km depth beneath