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Subsurface structure across the axis of the Tongariro Volcanic Centre, New Zealand
Journal of Volcanology and Geothermal Research 179 (2009) 233–240
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Journal of Volcanology and Geothermal Research j o u r n a l h o m e p a g e : w w w. e l s e v i e r. c o m / l o c a t e / j v o l g e o r e s
Subsurface structure across the axis of the Tongariro Volcanic Centre, New Zealand John Cassidy a, Malcolm Ingham b, Corinne A. Locke a,⁎, Hugh Bibby c a b c
School of Geography, Geology and Environmental Science, The University of Auckland, Private Bag 92019, Auckland, New Zealand School of Chemical and Physical Sciences, Victoria University of Wellington, Wellington, New Zealand GNS Science, Lower Hutt, New Zealand
a r t i c l e
i n f o
Article history: Received 18 June 2008 Accepted 7 November 2008 Available online 24 November 2008 Keywords: Tongariro Volcanic Centre gravity aeromagnetic magnetotelluric rift structure
a b s t r a c t The relationship between structure and volcanism in the Tongariro Volcanic Centre, New Zealand, is largely masked by a mantle of young volcanic deposits. Here we report the results of an integrated geophysical investigation (using gravity, multi-level aeromagnetic and magnetotelluric methods) of subsurface deposits and structures in the upper 1–2 km across the axis of the Tongariro Volcanic Centre. Modelling of these data across the Tama Lakes saddle shows that the outcropping volcanic deposits are up to 800 m thick, underlain by Tertiary sediments (of a few 10's to a few 100 m in thickness) and in turn lying above a basement of probable Mesozoic greywacke. Basement faulting is shown to be concentrated in the centre of the rift, which is 18 km wide at this location, but no vertical offset is resolved at the rift axis. Vertical displacements on basement faults of 250–300 m are modelled giving a minimum total basement subsidence of 650 m. A 5 kmwide, deep low resistivity zone occurs at the axis of the rift which is interpreted as either resulting from extensive fracturing and/or hydrothermal alteration within the basement. Steep-sided volcanic bodies with a high proportion of lavas/dykes coincide with the Waihi fault and the rift axis. Coincidence with the Waihi Fault suggests that this fault system may have provided magma pathways to the surface and a focus for dyke emplacement, which could have contributed to rift extension. The lack of offset at the rift axis may reflect the juvenile nature of faulting at this location, which is consistent with the notion of a migration of faulting towards the centre of the graben, alternatively, rifting may have been entirely accommodated by dyke emplacement. © 2008 Elsevier B.V. All rights reserved.
1. Introduction The Tongariro Volcanic Centre (TgVC) lies at the southern end of the Taupo Volcanic Zone (TVZ) (Fig. 1), a volcanic arc–back-arc system resulting from the westward subduction of the Pacific Plate below the North Island of New Zealand. Unusually, the volcanic arc is disrupted by rifting (Wilson et al., 1995) with significant extension occurring within the arc whereas off-shore, north of the TVZ, a back-arc basin occurs separately to the west of the volcanic arc (Gamble and Wright,1995). The TVZ is the focus of recent volcanism in New Zealand, with dominant rhyolitic and lesser andesitic, basaltic and dacitic activity and is considered to be an exceptionally active area of volcanism and tectonism (Wilson et al., 1995). Whilst the volcanological history of the region is well established (e.g. Wilson et al., 1995; Nakagawa et al., 1998; Price et al., 2005), its structural and deep crustal history is less well known due to the mantling of structural features by the young volcaniclastic deposits (Rowland and Sibson, 2001). Recent seismic tomography and magnetotelluric (MT) studies of the TgVC have been carried out to define 3D upper crustal structure and melt zones below the region (Rowlands et al., 2005; Ingham et al., 2008). ⁎ Corresponding author. Tel.: +64 9 3737599x88857; fax: +64 9 3737435. E-mail address: [email protected] (C.A. Locke). 0377-0273/$ – see front matter © 2008 Elsevier B.V. All rights reserved. doi:10.1016/j.jvolgeores.2008.11.017
However the seismic study, though extensive, resulted in generally poor resolution of structures in the upper 3 km on account of limited ray path distribution and the MT study was focused on Ruapehu volcano, to the south of the present study. Other geophysical data from the TgVC (summarized in Rowlands et al., 2005) are either too regional in nature or geographically sparse to delineate the complex near-surface structure of the region. Consequently there are relatively few published geophysical studies. A knowledge of the subsurface structure below the TgVC, especially basement faulting and the distribution of igneous intrusions at depth, is important in investigating the relationship between volcanism and tectonism in the region i.e. the structural controls on volcanism, the possible contribution of magma intrusion to extension and the geometry of any subjacent magma storage reservoirs. An ideal geophysical transect across the active rift axis is provided by the Tama Lakes saddle (Fig. 1), which is an area of subdued topography lying between the edifices of the Tongariro and Ruapehu volcanoes. This transect avoids the near-surface complexity likely to be associated with the volcanic edifices and in addition, the depth to basement may be less, hence subsurface structures should be more easily resolved by geophysical methods. Here we present the results of multiple geophysical surveys across the Tama Lakes saddle consisting primarily of gravity data and
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Fig. 1. Simplified geological map of the Tongariro Volcanic Centre (after Grindley, 1960; Nairn et al., 1998; Rowland and Sibson, 2001) showing the locations of the Waihi Fault, Rangipo Fault and Mt Ruapehu Graben (after Villamor and Berryman, 2006a), as discussed in text. Inset shows North Island, New Zealand, with area of the main figure boxed, located at the southern end of the Taupo Volcanic Zone (dashed).
aeromagnetic data at two different elevations and correlate these with the results of recent magnetotelluric soundings (Ingham et al., 2008). These data are modelled to delineate details of subsurface structure in the upper 2 km across the rift axis and thereby provide a link between surface mapping in the region and the recent seismic tomography studies. The integrated interpretation of these data sets significantly reduces the usual ambiguities associated with the interpretation of a single potential field data set and thereby allows new insights into the relationship between structure and volcanism below the TgVC. 2. Geological setting of the Tongariro Volcanic Centre Within the TVZ, extension is accommodated by a rift structure known as the Taupo Rift (Acocella et al., 2003; Villamor and Berryman, 2006a) which is associated with extension of ~7 mm yr− 1 (Villamor and Berryman, 2001), shallow earthquakes (Bryan et al., 1999; Hurst et al., 2002) and exceptionally high heat flow (Bibby et al., 1995); its presentday manifestation is also known as the Taupo Fault Belt (Berryman and Villamor, 1999) or Ruaumoko Rift (Rowland and Sibson, 2001). Rifting has been considered either to comprise variably-oriented parallel-sided segments linked by accommodation zones with extension orthogonal to the rift axes (Rowland and Sibson, 2001; Villamor and Berryman, 2006b) or to be oblique, resulting in a wedge-shaped rift (Acocella et al., 2003). The most southerly segment of the Taupo Rift, extending south from
Lake Taupo to the south of Mt Ruapehu, is known as the Tongariro Domain (Rowland and Sibson, 2001) within which the TgVC is located. Within the Tongariro Domain, the rift is a 40 km-wide NNE trending graben with a central dense faulting pattern and a 30° trending graben axis (Villamor and Berryman, 2006a; Rowland and Sibson, 2001). Faulting in this segment is predominantly normal dip-slip (Rowland and Sibson, 2001), with NW facing faults trending more easterly (by 15°) than SE facing faults. Evolution of the Taupo Rift has resulted in narrowing of the Tongariro Domain to the north of Ruapehu (Villamor and Berryman, 2006b). The structurally distinct southern part of the Tongariro Domain, known as the Mt Ruapehu Graben and the best studied, is extending at 2.3 ± 1.2 mm yr− 1 (Villamor and Berryman, 2006a). There are no central faults mapped in this graben, however gravity anomalies are considered to be consistent with 150–200 m of vertical displacement of the basement (Villamor and Berryman, 2006a). Extension of the Mt Ruapehu Graben started b400 ka ago (Villamor and Berryman, 2006b), which together with the less evolved nature of the rift and the timing of the onset of andesite volcanism (Gamble et al., 2003), led Villamor and Berryman (2006b) to consider that this southernmost region is a recent southerly increment of the Taupo Rift. Basement in the region is greywacke metasediments of Mesozoic age which outcrop on both sides of the rift, though their continuation at depth across the rift is still debated (Rowland and Sibson, 2001).
J. Cassidy et al. / Journal of Volcanology and Geothermal Research 179 (2009) 233–240
Immediately west of the Rangipo Fault (see Fig. 1), in the Mt Ruapehu Graben, basement greywacke is modelled at 100 masl (i.e. implying a ~ 750 m displacement of the Rangipo Fault) based on a line of seismic refraction data (Sissons and Dibble, 1981). Tertiary sediments, which overlie basement on the western margin of the rift, were also modelled along this line at 700 masl below a 250 m thick layer of recent lahar and pyroclastic deposits. However, the occurrence and depth of both Tertiary sediments and basement across the rift are considered to be poorly constrained (Villamor and Berryman, 2006b). The TgVC comprises two large recently active andesite volcanoes, Ruapehu and Tongariro (which includes the Ngauruhoe cone), as well as a number of old eroded centres (Cole et al., 1986); these volcanoes are surrounded by ring-plain deposits which fill the rift. Tongariro volcano is a complex of cones and craters that comprises 17 volcanostratigraphic units; compositionally it is mostly andesite with lesser basaltic-andesite and dacite components (Hobden et al., 1996). The oldest units of Tongariro volcano are exposed in the Tama Lakes area and date from 275 ka (Hobden et al., 1996). Although no systematic migration of the locus of activity is apparent across the volcano, the oldest vents are inferred to have a SW-NE alignment within a corridor 13 km long and 5 km wide (Hobden et al., 1999). Geochemical studies indicate that magmatic pathways have been short-lived and limited in extent during the last 1000 years of activity (Hobden et al., 1999). Ruapehu volcano has had four major cone building events (Hackett and Houghton, 1989) over the last 250 ka involving both central and flank vents and last erupted in 1995 and 1996 (Nakagawa et al., 1999; Gamble et al., 1999). The volcano is characterized by periods of high activity punctuated by periods of less vigorous activity (Gamble et al., 1999). Compositionally, the volcano is similar to Tongariro; fluctuations in magma composition over the short term are interpreted by Price et al. (2005) as indicating a complex subsurface plumbing system of small, often disconnected magma pockets. Eruptive events within the TgVC are usually from single or adjacent vents, however a unique period of intense activity occurred in the TgVC at about 10 ka during which time multiple vents (known as the PM vents) on both Ruapehu and Tongariro were active along a 20 km long, NNE trending 2–3 km wide linear zone (Nairn et al., 1998). Whilst these vents do not lie on any known fault trace, Nairn et al. (1998) consider that other faults in the region ruptured during this eruptive sequence. This, and the pattern of fault and vent zones in the wider rift is
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considered by Nairn et al. (1998) to indicate that an episode of accelerated extension occurred, with extension taken up by a) graben margin rifting on major regional faults and b) intrusion of dykes beneath the graben floor. Magmas feeding the 10 ka eruptions are thought to have been quenched at N1 km depth where abundant groundwater was available, possibly within Tertiary/Mesozoic sediments (Nairn et al., 1998). 3. Previous geophysical studies Previous geophysical studies of the TgVC have been mostly seismological and are summarized by Rowlands et al. (2005). Most relevant here is the seismic tomography study of Rowlands et al. (2005) which identified pronounced sub-vertical low velocity zones extending to 6–9 km depth below the Ruapehu and Tongariro volcanoes, which were attributed to either remnant partial melt, hydrothermal alteration or thickened volcanics. A more laterally extensive (15 km wide), but less pronounced low velocity zone from 3–5 km depth, was modelled below the Tama Lakes region and attributed to either fracturing or alteration (after the work of Ingham and Zeng, 1993). Regional gravity data (Woodward and Ferry, 1974) over the TgVC is dominated by a NNE-trending regional gradient, but shows little effect over the volcanoes due to sparse data distribution. The regional aeromagnetic data (Hunt and Whiteford, 1979) flown at 1050 m, however, show positive anomalies associated with the volcanic centres, reaching maxima of up to 600 nT over the volcanic peaks. A reconnaissance gravity profile across the Tama Lakes saddle was carried out by Zeng and Ingham (1993) and was interpreted (without constraints from other geophysical data) in terms of a basement high below the saddle. MT data collected along the same profile identified an anomalously low resistivity zone which was initially attributed to alteration/mineralization (Ingham and Zeng, 1993). A more recent, and more extensive, MT survey using modern measurement and processing techniques, and aimed at investigating the deep electrical structure of Ruapehu volcano, has been reported by Ingham et al. (2008). 4. Gravity and magnetic surveys A total of 38 new gravity stations established across the TgVC through the Tama Lakes saddle (Fig. 2) were integrated with 8 existing regional
Fig. 2. Geological map of the Tama Lakes saddle area (from Fig. 1) showing the location of the geophysical surveys; magnetotelluric sites are indicated by station numbers and triangles.
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stations (Woodward and Ferry, 1974). Elevations of the new gravity stations were determined either by optical levelling or GPS to an accuracy of ±0.5 m or better, equivalent to an uncertainty of ±1 g.u. in the gravity data (note: 1 g.u. = 0.1 mGal). Gravity data were corrected with a density of 2670 kg m− 3 (i.e. that of the greywacke basement; Reilly,1972) and fully terrain corrected. The most significant sources of error in the corrected gravity data are in the elevation and terrain corrections, especially where the topography is severe; the total error in the gravity data is estimated to be b10 g.u. The regional gravity field for the area was determined from 56 existing gravity measurements (Woodward and Ferry, 1974; Ferry et al., 1974) on basement greywacke surrounding the Tongariro region. This regional field was approximated to a third-order polynomial surface (similar to that of Zeng and Ingham,1993) which decreases to the southsoutheast with a gradient of 12 g.u. km− 1. The resulting residual gravity data (Fig. 3) projected onto the Tama Lakes saddle profile shows an asymmetric negative gravity anomaly across the volcanic axis, reaching a minimum of −190 g.u. Gravity gradients on the eastern limb of this anomaly are steeper (~18 g.u. km− 1) than on the western limb (~13 g.u. km− 1). Whilst the data are smooth across most of the profile, short wavelength variations 2–3 km wide and of about 15 g.u. in magnitude occur in the central section, probably reflecting near-surface variations in the density of the volcanic deposits. Aeromagnetic data (at elevations of 1500 m asl and 2500 m asl) were collected approximately coincident with the projected gravity profile at
a 2-second sampling interval (equivalent to a nominal 100 m spacing). Location and elevation were determined by differential GPS to an accuracy of 3–5 m. Diurnal variations were recorded at a local base station and subtracted from the survey data together with the International Geomagnetic Reference Field (IGRF). The resultant total field anomaly data at both elevations include a similarly subdued long wavelength regional component. A third-order polynomial surface was fitted to data (at 2500 m elevation) in the area remote from the volcanoes to approximate the regional field. This surface, which varies between 50 and 90 nT along the length of the profile, was subtracted from the aeromagnetic data at both elevations to give the residual anomaly data shown in Fig. 3. The residual anomaly data at 1500 m elevation are complex, with a central positive (325 nT) anomaly flanked by negative anomalies of about −100 nT. The width of the anomalous region is about 12 km. The residual anomaly data at 2500 m elevation show an asymmetric central positive anomaly about 8 km wide, reaching a maximum value of 230 nT. 5. Modelling gravity and magnetic data Published density and magnetisation data from the TVZ are somewhat scarce. Mean density values of 2730 kg m− 3 have been measured for littlealtered samples of Rotokawa andesite (Hunt and Harms, 1990) and 2720 kg m− 3 for andesites from the Tihia volcano (Sissons, 1981); similar values have been obtained for Taranaki andesite (2400–2830 kg m− 3;
Fig. 3. Gravity and magnetic models across the Tama Lakes saddle profile. (a) Observed residual gravity anomaly (diamonds) and calculated gravity effect (line) of the model below. (b) Cross-section through 2.75D gravity model. (c) Observed residual magnetic anomaly data at 2500 m elevation (diamonds) and calculated magnetic effect (line) of the model shown below. (d) Observed residual magnetic anomaly data at 1500 m elevation (diamonds) and calculated magnetic effect (line) of the model shown below. (e) Cross-section through 2.75D magnetic model (note that a slight difference in topography occurs between the two models as a consequence of positional differences).
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Locke et al.,1993). It should be noted of course, that the density of andesite depends significantly on the degrees of vesicularity and alteration. The density of volcaniclastic material is even more variable and therefore a bulk density is more difficult to quantify. Studies of volcanoes in New Zealand and elsewhere yield values of 2000–2400 kg m− 3 for volcaniclastics (Budetta et al., 1983; Brown et al., 1987; Williams et al., 1987; Fedi et al., 1991; Deplus et al.,1995); in the present modelling, an average value of 2200 kg m− 3 has been assumed. Late-Tertiary sediments in New Zealand have densities of about 2200 kg m− 3 (Whiteford and Lumb, 1975) and a density of 2670 kg m− 3 has been attributed to the greywacke basement (after Reilly, 1972); these values are incorporated in the models. Magnetic remanence intensities of andesites are typically very variable. For example, values for young fresh andesite flows on Egmont Volcano range from 0.2 to 40 A m− 1 (Downey et al., 1994). Koenigsberger (q) ratios are typically high, about 20 according to Cox (1970) and Downey et al. (1994), hence induced magnetisation is likely to be a minor component. Similar remanence intensities have been deduced for andesites elsewhere, e.g. 1–10 Am− 1 with q-ratios of 1–40 (Blakely and Christiansen, 1978; Flanagan and Williams, 1982). Modelling of the gravity and two aeromagnetic data sets, carried out in 2.75 D using Interpex MagixXL, was performed iteratively in order to find the optimum model that best fitted all three data sets. The calculated effects of this model, together with the observed data and a cross-section through the model, are shown in Fig. 3. The RMS discrepancies between calculated and observed values for the three data sets are: 8.5 g.u. for gravity (b5% of the maximum anomaly), 40 nT for the low-level aeromagnetics (10% of the maximum anomaly) and 14 nT for the high-level aeromagnetics (6% of the maximum anomaly). The optimum model (Fig. 3) comprises an extensive upper layer with density 2200 kg m− 3 and maximum thickness of 1000 m and of effectively infinite extent orthogonal to the profile. A single magnetisation for this layer cannot account for the observed data, hence a lower layer of 300 m thickness and zero magnetisation has been modelled below an upper layer of 1.3 A m− 1 magnetisation. The asymmetry of the gravity anomaly requires an abrupt depth change (modelled as 300 m) in this lowest interface at ~9 km profile distance, further abrupt changes in this interface of 300 m and 250 m are modelled at ~6 and 25 km profile distance, respectively, in order to fit the observed anomaly. The gravity data cannot be accounted for by assuming that the entire upper volcanic layer has a low bulk density (in the region of 2200 kg m− 3), there has to be denser material towards the centre of the profile. This is consistent with the aeromagnetic data which indicates the occurrence of highly magnetised rocks (usually associated with high densities) central to the profile; in addition, having magnetic data at two elevations significantly reduces the modelling ambiguity. The magnetic anomalies show that these central high-magnetisation/denser rocks are unlikely to be due to elevated greywacke basement as this is essentially nonmagnetic. At the surface, between profile distances 4.5 and 13 km, a dense (2400 kg m− 3) and magnetic (4 A m− 1) body is modelled that extends to the basement interface at about 9 km and 12 km profile distance. This body has been modelled as extending 4 km orthogonally either side of the profile. The low-level aeromagnetic data are the most sensitive to modelling and hence the most diagnostic; extensive trial-and-error modelling showed that a shallow magnetic body is required to fit the data. The observed short wavelength variations require that the lower surface of the more magnetic body be undulating i.e. the observed anomaly cannot be accounted for by a single deep steep-sided body. The magnetic lows flanking the positive anomaly are largely a consequence of the subvertical magnetic boundary within the low density layer. 6. Magnetotelluric soundings Broadband magnetotelluric soundings made at a total of 37 sites around Ruapehu were reported by Ingham et al. (2008). Measurements were made using 5-component Phoenix MTU5-2000 systems and were
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remote referenced, providing generally excellent data quality. Although the Tama Lakes saddle is close to the northern boundary of the MT survey area, six of the sites (marked by triangles in Fig. 2) lie close to the gravity and aeromagnetic profiles. The electrical resistivity structure derived across the Tama Lakes saddle from modelling of data from these six sites is considered here. As reported by Ingham et al. (2008), the MT data were subjected to phase tensor analysis (Caldwell et al., 2004) in order to both identify the dimensionality of resistivity structure at each site, and to remove galvanic distortions from the data in the manner proposed by Bibby et al. (2005). As is discussed by Ingham et al. (2008), for periods less than about 10 s the data are 1-dimensional and at each site may be adequately interpreted in terms of a layered resistivity structure. As a result, following removal of galvanic distortion effects, 1-D models which satisfy the short period data at each site were derived from the determinant impedance. Shallow features identified by these models were confirmed with a forward 3-D numerical model (Cairns, 2006) which incorporated topography. A compilation of the 1-dimensional models for the sites across the Tama Lakes saddle are shown in Fig. 4, superimposed on the gravity/ magnetic model. Note that a few sites lie up to 3 km from the gravity/ magnetic profile and at higher elevations (sites 111 and 112 in particular), therefore some caution is required in correlating the MT section, however it is believed to be broadly representative of the resistivity structure across the Tama Lakes saddle. At all of the sites except 115 (the furthest NW), the basic resistivity structure at depth is resistor–conductor–resistor. Whereas the depth to the conductive layer is generally well constrained, the thickness and resistivity of the conductor are non-unique due to equivalence (i.e. a model with the same conductivity-thickness product will also fit the data). Consequently, the depth to the basal resistive layer is not well constrained. This is possibly highlighted at the south-east end of the profile where it is likely (as shown in the gravity/magnetic model) that the observed low resistivity layer thins out as outcropping basement is approached. The two most striking features of the compilation of 1-D models are: (i) a layer of anomalously low resistivity (10–20 Ωm) beneath the central four sites which thickens sharply to almost 1 km thickness beneath site 111, and (ii) a layer of intermediate resistivity lying above the low resistivity layer at sites 114 and 111 and immediately below the high density, high magnetisation surface cover. It is also interesting to note that higher surficial resistivities appear to be associated with these sites. 7. Interpretation Modelling of the different data sets has defined a number of significant geophysical structures which can be correlated. The greywacke basement modelled from gravity data correlates with the lowermost high resistivity layer in the MT model; similar resistivities have been reported for greywacke elsewhere (e.g. Ingham et al., 2001). A relative deepening of the basement surface from west to east is common to both models and depths broadly agree (within probable uncertainties) except notably in the centre of the profile. The low density layer in the gravity model broadly correlates with the low-intermediate resistivity middle layer, modelled from the MT data. Furthermore, the magnetic model differentiates two components within this layer, consisting of an upper relatively magnetised layer and a lower non-magnetised layer. These are interpreted as an upper layer of mixed volcanic deposits (including lahar and pyroclastic deposits) as mapped over much of the profile (Grindley, 1960), together with a lower layer of Tertiary sediments. The combined modelling shows that the volcanic deposits are up to 800 m thick and that the Tertiary sediments vary from a few 10's to a few 100 m in thickness. Three discontinuities in the basement–Tertiary interface, at approximately 6, 9 and 24 km profile distance, have been modelled from the gravity data. The easternmost of these correlates with the mapped
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Fig. 4. Magnetotelluric model compilation superimposed on gravity/magnetic model (from Fig. 3). The individual 1D models showing depths to layers (as short bars) with modelled resistivity values are plotted below the respective MT stations (marked by arrows with station numbers, as referred to in text); tie lines (dotted) and hatching are used to outline areas of similar resistivity. The projected locations of the mapped surface traces of the key faults are also shown for comparison with the model.
surface trace of the Rangipo fault (Grindley, 1960; Rowland and Sibson, 2001). The central discontinuity (at 9 km profile distance) correlates with the Waihi Fault (see Fig. 1) and the western discontinuity correlates with an unnamed mapped fault set (Rowland and Sibson, 2001). Vertical displacements of 250–300 m are modelled on each of these discontinuities, which is less than the 800 m displacement deduced for the Rangipo Fault further south (Villamor and Berryman, 2006a) but of a similar magnitude to that expected on faults in the centre of the graben to the south of Ruapehu (Villamor and Berryman, 2006a). The resolution of the gravity model, however, does not preclude these basement dislocations being accommodated by a series of smaller offsets rather than single fault. The near-surface, high density/high magnetisation body in the centre of the profile correlates with the high resistivities modelled beneath sites 114 and 111 and is interpreted as volcanic deposits with a high proportion of lavas and/or dykes, consistent with geological mapping in the region (Grindley, 1960). The greater apparent lateral extent of the high resistivity layer is probably a consequence of the offline location of the MT sites which lie on or closer to mapped lavas. This high density/high magnetisation body is modelled with two locally steepened structures extending to greater depth, one coincides with the basement discontinuity at 9 km profile distance and the other with the axis of the 10 ka volcanic vents (Nairn et al., 1998). Both of these features penetrate through the zone of intermediate resistivity into the thickened anomalously low resistivity layer and may reflect magma conduits extending into the basement but cannot be resolved as separate features by the sparse MT site distribution. The striking 5 km-wide, deep low resistivity zone at the centre of the profile (noted earlier) may result from enhanced fracture porosity within the basement. Equally, the occurrence of magma feeder systems might suggest that this low resistivity zone (and its extension at shallower levels) results at least in part, from conductive clays produced by hydrothermal alteration of the overlying volcanics. Such alteration could of course have been facilitated by the occurrence of related enhanced permeability. Similarly, conductive clays have been reported by Ingham et al. (2008) and Jones et al. (2008), for example, to be widespread on the higher slopes of Ruapehu volcano. Any occurrence of enhanced porosity or alteration appears to be insufficiently intense to be resolved by the gravity or magnetic data.
8. Conclusions and implications Integrated interpretation of multiple geophysical data sets across the Tama Lakes saddle has successfully delineated a number of key features of the subsurface structure of the Tongariro Volcanic Centre to 2 km depth. These being: 1. A minimum basement subsidence of 650 m at the centre of the rift relative to the margins of the rift. 2. Significant offsets in the basement interface that correlate with the mapped surface traces of the Waihi Fault and an unnamed fault (mapped 3 km to the west). However, no offset at the rift axis (i.e. the locus of the PM vents) is apparent from the geophysical data. 3. A deeper region of enhanced fracture porosity and/or alteration within the basement at the centre of the rift. 4. Steep-sided bodies of volcanic rocks with a high proportion of lavas/ dykes, that are coincident with both the Waihi Fault and the rift axis. Basement faulting is concentrated in the centre of the rift (apart from the Rangipo Fault at the eastern margin), there is no geophysical evidence for any significant faulting at the western margin. Across the Tama Lakes saddle, the rift is shown to be 18 km wide and therefore the N–S segmentation from a 15 km- to 40 km-wide modern Taupo Fault Belt as suggested by Villamor and Berryman (2006b) must occur at a more southerly location (i.e. closer to Ruapehu volcano) than they proposed. The zone of anomalously low resistivity (at ~1–2 km depth) interpreted as resulting from enhanced fracture porosity and/or alteration in the basement may relate to the deeper, more extensive low-velocity zone identified from seismic tomography which was also attributed to possible fracturing (Rowlands et al., 2005). This low resistivity zone coincides with the estimated depth of N1 km at which the PM magmas were quenched, based on vesicularity studies (Nairn et al., 1998). Groundwater for this magma–water interaction was suggested by Nairn et al. (1998) to be sourced from either andesite intrusives in Tertiary sediments or basement greywacke at depths of N1 km; the geophysical modelling indicates that the latter is more probable. There is no geophysical evidence for any large volume high-density bodies that might indicate relict magma chambers associated with
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volcanism at shallow levels below this region (c.f. the Taranaki volcanoes; Locke and Cassidy, 1997). This is consistent with the notion that the PM magmas rose rapidly from deep levels such that large volumes of magma did not reside for any lengthy period at high levels at this location (Nakagawa et al., 1998). The coincidence of one of the zones of denser volcanic rocks with both the Waihi Fault and a significant basement offset, suggests that this fault system may have provided magma pathways to the surface and a focus for dyke emplacement which could have contributed to extension of the rift. In contrast, there is no resolvable offset in the basement interface at the rift axis, where the PM vents and the second dense volcanic body occur. The PM eruptions are considered to be the result of a regional rifting episode which tapped small magma bodies (Nairn et al., 1998; Nakagawa et al., 1998), therefore it is perhaps surprising that no substantial basement offset is apparent at the rift axis from the geophysical data. This may reflect the juvenile nature of faulting at this location and be indicative of the migration of faulting towards the centre of the graben, as has occurred further north (Villamor and Berryman, 2006b). Alternatively, rifting across the Tama Lakes saddle during the PM event could have been entirely accommodated by dyke emplacement. Acknowledgements We thank The University of Auckland Research Committee, Earthwatch Centre for Field Research, and Environment Waikato for financial support; the Department of Conservation for allowing access to the area; our pilot the late Alan Smallfield; Colin Yong for technical assistance; University of Auckland graduate students and Earthwatch volunteers for assistance with gravity surveys; Stacey Dravitzki, Yasuo Ogawa and Stewart Bennie for help with MT measurements; Louise Cotterall for assistance with manuscript preparation and Julie Rowland for providing information for Fig.1. John Gamble and Glyn Williams-Jones are thanked for helpful reviews. References Acocella, V., Spinks, K., Cole, J., Nicol, A., 2003. Oblique back arc rifting of Taupo Volcanic Zone, New Zealand. Tectonics 22 (4), 1045. doi:10.1029/2002TC001447. Berryman, K., Villamor, P., 1999. Spatial and Temporal Zoning of Faulting in the Taupo Volcanic Zone, New Zealand. Geological Society of New Zealand Miscellaneous Publication, vol. 107A, p. 15. Bibby, H.M., Caldwell, T.G., Davey, F.J., Webb, T.H., 1995. Geophysical evidence on the structure of the Taupo Volcanic Zone and its hydrothermal circulations. J. Volcanol. Geotherm. Res 68, 29–58. Bibby, H.M., Caldwell, T.G., Brown, C., 2005. Determinable and non-determinable parameters of galvanic distortion in magnetotellurics. Geophys. J. Int 163, 915–930. Blakely, R.J., Christiansen, R.L., 1978. The magnetisation of Mt Shasta and implications for virtual geomagnetic poles determined from seamounts. J. Geophys. Res 83 (B12), 5971–5978. Brown, G.C., Rymer, H., Thorpe, R.S., 1987. The evolution of andesite volcano structures: new evidence from gravity studies in Costa Rica. Earth Planet. Sci. Lett 82, 323–334. Bryan, C.J., Sherburn, S., Bibby, H.M., Bannister, S.C., Hurst, A.W., 1999. Shallow seismicity of the central Taupo Volcanic Zone, New Zealand: its distribution and nature. N.Z. J. Geol. and Geophys 42, 533–542. Budetta, G., Nunziata, C., Rapolla, A., 1983. A gravity study of the island of Vulcano, Tyrrhenian Sea, Italy. Bull. Volcanol 46, 183–192. Cairns, P.J., 2006. 3-D modelling of the electrical resistivity structure of Mount Ruapehu. Unpublished MSc Thesis, Victoria University of Wellington, Wellington, New Zealand. Caldwell, T.G., Bibby, H.M., Brown, C., 2004. The magnetotelluric phase tensor. Geophys. J. Int 158, 457–469. Cole, J.W., Graham, I.J., Hackett, W.R., Houghton, B.F., 1986. Volcanology and petrology of the Quaternary composite volcanoes of Tongariro Volcanic Centre, Taupo Volcanic Zone. In: Smith, I.E.M. (Ed.), Late Cenozoic Volcanism in New Zealand Royal Soc. N.Z. Bull. 23, pp. 224–250. Cox, A., 1970. Remanent magnetisation and susceptibility of late Cenozoic rocks from New Zealand. N.Z. J. Geol. Geophys 14, 192–207. Deplus, C., Bonvalot, S., Dahrin, D., Diament, M., Harjono, H., Dubois, J., 1995. Inner structure of the Krakatoa volcanic complex (Indonesia) from gravity and bathymetry data. J. Volcanol. Geotherm. Res 64, 23–51. Downey, W.S., Kellett, R.J., Smith, I.E.M., Price, R.C., Stewart, R.B.,1994. New Palaeomagnetic evidence for the recent eruptive activity of Mt Taranaki, New Zealand. J. Volcanol. Geotherm. Res 60, 15–27.
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Fedi, M., Nunziata, C., Rapolla, A., 1991. The Campania–Campi Flegrei area: a contribution to discern the best structural interpretation from gravity interpretation. J. Volcanol. Geotherm. Res 48, 51–59. Ferry, L.M., Woodward, D.J., Whiteford, C.M., 1974. Sheet 7, Taranaki (1st ed.). Gravity map of New Zealand, 1:250000, Bouguer Anomalies. Dept. of Scientific and Industrial Research, Wellington, New Zealand. Flanagan, G., Williams, D.L., 1982. A magnetic investigation of Mt Hood, Oregon. J. Geophys. Res 87 (B4), 2804–2814. Gamble, J.A., Wright, I.C., 1995. The southern Havre Trough geological structure and magma petrogenesis of a back arc rift complex. In: Taylor, B. (Ed.), Back Arc Basins: Tectonics and Magmatism. Plenum Press, New York, pp. 29–62. Gamble, J.A., Wood, P., Price, R.C., Smith, I.E.M., Waight, T.E., 1999. A fifty year perspective of magmatic evolution on Ruapheu volcano, New Zealand: verification of open system behavior in an arc volcano. Earth. Planet. Sci. Letts 170, 301–314. Gamble, J.A., Price, R.C., Smith, I.E.M., McIntosh, W.C., Dunbar, N.W., 2003. 40Ar/39Ar geochronology of magmatic activity, magma flux and hazards at Ruapehu volcano, Taupo Volcanic Zone, New Zealand. J. Volcanol. Geotherm. Res 120, 271–287. Grindley, G.W., 1960. Sheet 8 Taupo (1st edition) Geological map of New Zealand. Dept. of Scientific and Industrial Research, Wellington, New Zealand. Hackett, W.R., Houghton, B.F., 1989. A facies model for a Quaternary andesitic composite volcano: Ruapehu, New Zealand. Bull. Volcanol 51, 51–68. Hobden, B.J., Houghton, B.F., Lanphere, M.A., Nairn, I.A., 1996. Growth of the Tongariro volcanic complex: new evidence from K–Ar age determinations. N.Z. J. Geol. Geophys 39, 151–154. Hobden, B.J., Houghton, B.F., Davidson, J.P., Weaver, S.D., 1999. Small and short-lived magma batches at composite volcanoes: time windows at Tongariro volcano, New Zealand. J. Geol. Soc. (London) 156, 865–868. Hunt, T.M., Harms, C., 1990. Gravity Survey of the Rotokawa Geothermal Field. Proceedings of the 12th New Zealand Geothermal Workshop. University of Auckland, Auckland, pp. 91–92. Hunt, T.M., Whiteford, C.M., 1979. Sheet 5 (Rotorua) Magnetic map of New Zealand 1:250000 Total force anomalies. Dept. of Scientific and Industrial Research, Wellington, New Zealand. Hurst, A.W., Bibby, H.M., Robinson, R.R., 2002. Earthquake focal mechanisms in the central Taupo Volcanic Zone and their relation to faulting and deformation. N.Z. J. Geol. Geophys 45, 527–536. Ingham, M., Zeng, Y., 1993. AMT soundings to determine electrical structure in Tongariro National Park. Proceedings of the 15th New Zealand Geothermal Workshop. University of Auckland, New Zealand, pp. 373–379. Ingham, M., Whaler, K., McKnight, D., 2001. Magnetotelluric sounding of the Hikurangi Margin, New Zealand. Geophys. J. Int 144, 343–355. Ingham, M.R., Bibby, H.M., Heise, W., Jones, K.A., Cairns, P., Dravitzki, S., Bennie, S.L., Caldwell, T.G., Ogawa, Y., 2008. A magnetotelluric study of Mount Ruapehu volcano, New Zealand. Submitted to Geophys. J. Int. Jones, K., Ingham, M., Bibby, H., 2008. A high frequency MT study of the hydrothermal vent system of Mount Ruapehu, New Zealand. J. Volcanol. Geotherm. Res 176, 591–600. Locke, C.A., Cassidy, J., 1997. Egmont Volcano: three-dimensional structure and its implications for evolution. J. Volcanol. Geotherm. Res 76, 149–161. Locke, C.A., Cassidy, J., MacDonald, A., 1993. Three-dimensional structure of relict stratovolcanoes in Taranaki, New Zealand: evidence from gravity data. J. Volcanol. Geotherm. Res. 59, 121–130. Nairn, I.A., Kobayashi, T., Nakagawa, M., 1998. The ~ 10 ka multiple vent pyroclastic eruption sequence at Tongariro Volcanic Centre, Taupo Volcanic Zone, New Zealand: part 1. eruptive processes during regional extension. J. Volcanol. Geotherm. Res 86, 19–44. Nakagawa, M., Nairn, I.A., Kobayashi, T., 1998. The ~ 10 ka multiple vent pyroclastic eruption sequence at Tongariro Volcanic Centre, Taupo Volcanic Zone, New Zealand. Part 2. Petrological insights into magma storage and transport during regional extension. J. Volcanol. Geotherm. Res 86, 45–65. Price, R.C., Gamble, J.A., Smith, I.E.M., Stewart, R.B., Eggins, S., Wright, I.C., 2005. An integrated model for the temporal evolution of andesites and rhyolites and crustal development in New Zealand's North Island. J. Volcanol. Geotherm. Res 140, 1–24. Reilly, W.I., 1972. New Zealand Gravity Map Series. N.Z. J. Geol. Geophys 15, 3–15. Rowland, J.V., Sibson, R.H., 2001. Extensional fault kinematics within the Taupo Volcanic Zone, New Zealand: soft-linked segmentation of a continental rift system. N.Z. J. Geol. Geophys 44, 271–283. Rowlands, D.P., White, R.S., Haines, A.J., 2005. Seismic tomography of the Tongariro Volcanic centre, New Zealand. Geophys. J. Int 163, 1180–1194. Sissons, B.A., 1981. Densities from surface and subsurface gravity measurements. Geophysics 46, 1568–1571. Sissons, B.A., Dibble, R.R., 1981. A seismic refraction experiment south-east of Ruapehu volcano. N.Z. J. Geol. Geophys 24, 31–38. Villamor, P., Berryman, K.B., 2001. A late Quaternary extension rate in the Taupo Volcanic Zone, New Zealand, derived from fault slip data. N.Z. J. Geol. Geophys 44, 243–269. Villamor, P., Berryman, K.B., 2006a. Late Quaternary geometry and kinematics of faults at the southern termination of the Taupo Volcanic Zone, New Zealand. N.Z. J. Geol. Geophys 49, 1–21. Villamor, P., Berryman, K.B., 2006b. Evolution of the southern termination of the Taupo Rift, New Zealand. N.Z. J. Geol. Geophys. 49, 23–37. Whiteford, C.M., Lumb, J.T., 1975. A catalogue of physical properties of rocks. Volumes 1– 3: Geophysics Division Report No 106. Dept. of Scientific and Industrial Research, Wellington, New Zealand.
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Williams, D.L., Abrams, G., Finn, C., Dzurisin, D., Johnson, D.J., Denlinger, R., 1987. Evidence from gravity data for an intrusive complex beneath Mt St Helens. J. Geophys. Res 92, 10207–10222. Wilson, C.J.N., Houghton, B.F., McWilliams, M.O., Lanphere, M.A., Weaver, S.D., Briggs, R.M., 1995. Volcanic and structural evolution of the Taupo Volcanic Zone: a review. J. Volcanol. Geotherm. Res 68, 1–28.
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Contents lists available at ScienceDirect
Journal of Volcanology and Geothermal Research j o u r n a l h o m e p a g e : w w w. e l s e v i e r. c o m / l o c a t e / j v o l g e o r e s
Subsurface structure across the axis of the Tongariro Volcanic Centre, New Zealand John Cassidy a, Malcolm Ingham b, Corinne A. Locke a,⁎, Hugh Bibby c a b c
School of Geography, Geology and Environmental Science, The University of Auckland, Private Bag 92019, Auckland, New Zealand School of Chemical and Physical Sciences, Victoria University of Wellington, Wellington, New Zealand GNS Science, Lower Hutt, New Zealand
a r t i c l e
i n f o
Article history: Received 18 June 2008 Accepted 7 November 2008 Available online 24 November 2008 Keywords: Tongariro Volcanic Centre gravity aeromagnetic magnetotelluric rift structure
a b s t r a c t The relationship between structure and volcanism in the Tongariro Volcanic Centre, New Zealand, is largely masked by a mantle of young volcanic deposits. Here we report the results of an integrated geophysical investigation (using gravity, multi-level aeromagnetic and magnetotelluric methods) of subsurface deposits and structures in the upper 1–2 km across the axis of the Tongariro Volcanic Centre. Modelling of these data across the Tama Lakes saddle shows that the outcropping volcanic deposits are up to 800 m thick, underlain by Tertiary sediments (of a few 10's to a few 100 m in thickness) and in turn lying above a basement of probable Mesozoic greywacke. Basement faulting is shown to be concentrated in the centre of the rift, which is 18 km wide at this location, but no vertical offset is resolved at the rift axis. Vertical displacements on basement faults of 250–300 m are modelled giving a minimum total basement subsidence of 650 m. A 5 kmwide, deep low resistivity zone occurs at the axis of the rift which is interpreted as either resulting from extensive fracturing and/or hydrothermal alteration within the basement. Steep-sided volcanic bodies with a high proportion of lavas/dykes coincide with the Waihi fault and the rift axis. Coincidence with the Waihi Fault suggests that this fault system may have provided magma pathways to the surface and a focus for dyke emplacement, which could have contributed to rift extension. The lack of offset at the rift axis may reflect the juvenile nature of faulting at this location, which is consistent with the notion of a migration of faulting towards the centre of the graben, alternatively, rifting may have been entirely accommodated by dyke emplacement. © 2008 Elsevier B.V. All rights reserved.
1. Introduction The Tongariro Volcanic Centre (TgVC) lies at the southern end of the Taupo Volcanic Zone (TVZ) (Fig. 1), a volcanic arc–back-arc system resulting from the westward subduction of the Pacific Plate below the North Island of New Zealand. Unusually, the volcanic arc is disrupted by rifting (Wilson et al., 1995) with significant extension occurring within the arc whereas off-shore, north of the TVZ, a back-arc basin occurs separately to the west of the volcanic arc (Gamble and Wright,1995). The TVZ is the focus of recent volcanism in New Zealand, with dominant rhyolitic and lesser andesitic, basaltic and dacitic activity and is considered to be an exceptionally active area of volcanism and tectonism (Wilson et al., 1995). Whilst the volcanological history of the region is well established (e.g. Wilson et al., 1995; Nakagawa et al., 1998; Price et al., 2005), its structural and deep crustal history is less well known due to the mantling of structural features by the young volcaniclastic deposits (Rowland and Sibson, 2001). Recent seismic tomography and magnetotelluric (MT) studies of the TgVC have been carried out to define 3D upper crustal structure and melt zones below the region (Rowlands et al., 2005; Ingham et al., 2008). ⁎ Corresponding author. Tel.: +64 9 3737599x88857; fax: +64 9 3737435. E-mail address: [email protected] (C.A. Locke). 0377-0273/$ – see front matter © 2008 Elsevier B.V. All rights reserved. doi:10.1016/j.jvolgeores.2008.11.017
However the seismic study, though extensive, resulted in generally poor resolution of structures in the upper 3 km on account of limited ray path distribution and the MT study was focused on Ruapehu volcano, to the south of the present study. Other geophysical data from the TgVC (summarized in Rowlands et al., 2005) are either too regional in nature or geographically sparse to delineate the complex near-surface structure of the region. Consequently there are relatively few published geophysical studies. A knowledge of the subsurface structure below the TgVC, especially basement faulting and the distribution of igneous intrusions at depth, is important in investigating the relationship between volcanism and tectonism in the region i.e. the structural controls on volcanism, the possible contribution of magma intrusion to extension and the geometry of any subjacent magma storage reservoirs. An ideal geophysical transect across the active rift axis is provided by the Tama Lakes saddle (Fig. 1), which is an area of subdued topography lying between the edifices of the Tongariro and Ruapehu volcanoes. This transect avoids the near-surface complexity likely to be associated with the volcanic edifices and in addition, the depth to basement may be less, hence subsurface structures should be more easily resolved by geophysical methods. Here we present the results of multiple geophysical surveys across the Tama Lakes saddle consisting primarily of gravity data and
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Fig. 1. Simplified geological map of the Tongariro Volcanic Centre (after Grindley, 1960; Nairn et al., 1998; Rowland and Sibson, 2001) showing the locations of the Waihi Fault, Rangipo Fault and Mt Ruapehu Graben (after Villamor and Berryman, 2006a), as discussed in text. Inset shows North Island, New Zealand, with area of the main figure boxed, located at the southern end of the Taupo Volcanic Zone (dashed).
aeromagnetic data at two different elevations and correlate these with the results of recent magnetotelluric soundings (Ingham et al., 2008). These data are modelled to delineate details of subsurface structure in the upper 2 km across the rift axis and thereby provide a link between surface mapping in the region and the recent seismic tomography studies. The integrated interpretation of these data sets significantly reduces the usual ambiguities associated with the interpretation of a single potential field data set and thereby allows new insights into the relationship between structure and volcanism below the TgVC. 2. Geological setting of the Tongariro Volcanic Centre Within the TVZ, extension is accommodated by a rift structure known as the Taupo Rift (Acocella et al., 2003; Villamor and Berryman, 2006a) which is associated with extension of ~7 mm yr− 1 (Villamor and Berryman, 2001), shallow earthquakes (Bryan et al., 1999; Hurst et al., 2002) and exceptionally high heat flow (Bibby et al., 1995); its presentday manifestation is also known as the Taupo Fault Belt (Berryman and Villamor, 1999) or Ruaumoko Rift (Rowland and Sibson, 2001). Rifting has been considered either to comprise variably-oriented parallel-sided segments linked by accommodation zones with extension orthogonal to the rift axes (Rowland and Sibson, 2001; Villamor and Berryman, 2006b) or to be oblique, resulting in a wedge-shaped rift (Acocella et al., 2003). The most southerly segment of the Taupo Rift, extending south from
Lake Taupo to the south of Mt Ruapehu, is known as the Tongariro Domain (Rowland and Sibson, 2001) within which the TgVC is located. Within the Tongariro Domain, the rift is a 40 km-wide NNE trending graben with a central dense faulting pattern and a 30° trending graben axis (Villamor and Berryman, 2006a; Rowland and Sibson, 2001). Faulting in this segment is predominantly normal dip-slip (Rowland and Sibson, 2001), with NW facing faults trending more easterly (by 15°) than SE facing faults. Evolution of the Taupo Rift has resulted in narrowing of the Tongariro Domain to the north of Ruapehu (Villamor and Berryman, 2006b). The structurally distinct southern part of the Tongariro Domain, known as the Mt Ruapehu Graben and the best studied, is extending at 2.3 ± 1.2 mm yr− 1 (Villamor and Berryman, 2006a). There are no central faults mapped in this graben, however gravity anomalies are considered to be consistent with 150–200 m of vertical displacement of the basement (Villamor and Berryman, 2006a). Extension of the Mt Ruapehu Graben started b400 ka ago (Villamor and Berryman, 2006b), which together with the less evolved nature of the rift and the timing of the onset of andesite volcanism (Gamble et al., 2003), led Villamor and Berryman (2006b) to consider that this southernmost region is a recent southerly increment of the Taupo Rift. Basement in the region is greywacke metasediments of Mesozoic age which outcrop on both sides of the rift, though their continuation at depth across the rift is still debated (Rowland and Sibson, 2001).
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Immediately west of the Rangipo Fault (see Fig. 1), in the Mt Ruapehu Graben, basement greywacke is modelled at 100 masl (i.e. implying a ~ 750 m displacement of the Rangipo Fault) based on a line of seismic refraction data (Sissons and Dibble, 1981). Tertiary sediments, which overlie basement on the western margin of the rift, were also modelled along this line at 700 masl below a 250 m thick layer of recent lahar and pyroclastic deposits. However, the occurrence and depth of both Tertiary sediments and basement across the rift are considered to be poorly constrained (Villamor and Berryman, 2006b). The TgVC comprises two large recently active andesite volcanoes, Ruapehu and Tongariro (which includes the Ngauruhoe cone), as well as a number of old eroded centres (Cole et al., 1986); these volcanoes are surrounded by ring-plain deposits which fill the rift. Tongariro volcano is a complex of cones and craters that comprises 17 volcanostratigraphic units; compositionally it is mostly andesite with lesser basaltic-andesite and dacite components (Hobden et al., 1996). The oldest units of Tongariro volcano are exposed in the Tama Lakes area and date from 275 ka (Hobden et al., 1996). Although no systematic migration of the locus of activity is apparent across the volcano, the oldest vents are inferred to have a SW-NE alignment within a corridor 13 km long and 5 km wide (Hobden et al., 1999). Geochemical studies indicate that magmatic pathways have been short-lived and limited in extent during the last 1000 years of activity (Hobden et al., 1999). Ruapehu volcano has had four major cone building events (Hackett and Houghton, 1989) over the last 250 ka involving both central and flank vents and last erupted in 1995 and 1996 (Nakagawa et al., 1999; Gamble et al., 1999). The volcano is characterized by periods of high activity punctuated by periods of less vigorous activity (Gamble et al., 1999). Compositionally, the volcano is similar to Tongariro; fluctuations in magma composition over the short term are interpreted by Price et al. (2005) as indicating a complex subsurface plumbing system of small, often disconnected magma pockets. Eruptive events within the TgVC are usually from single or adjacent vents, however a unique period of intense activity occurred in the TgVC at about 10 ka during which time multiple vents (known as the PM vents) on both Ruapehu and Tongariro were active along a 20 km long, NNE trending 2–3 km wide linear zone (Nairn et al., 1998). Whilst these vents do not lie on any known fault trace, Nairn et al. (1998) consider that other faults in the region ruptured during this eruptive sequence. This, and the pattern of fault and vent zones in the wider rift is
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considered by Nairn et al. (1998) to indicate that an episode of accelerated extension occurred, with extension taken up by a) graben margin rifting on major regional faults and b) intrusion of dykes beneath the graben floor. Magmas feeding the 10 ka eruptions are thought to have been quenched at N1 km depth where abundant groundwater was available, possibly within Tertiary/Mesozoic sediments (Nairn et al., 1998). 3. Previous geophysical studies Previous geophysical studies of the TgVC have been mostly seismological and are summarized by Rowlands et al. (2005). Most relevant here is the seismic tomography study of Rowlands et al. (2005) which identified pronounced sub-vertical low velocity zones extending to 6–9 km depth below the Ruapehu and Tongariro volcanoes, which were attributed to either remnant partial melt, hydrothermal alteration or thickened volcanics. A more laterally extensive (15 km wide), but less pronounced low velocity zone from 3–5 km depth, was modelled below the Tama Lakes region and attributed to either fracturing or alteration (after the work of Ingham and Zeng, 1993). Regional gravity data (Woodward and Ferry, 1974) over the TgVC is dominated by a NNE-trending regional gradient, but shows little effect over the volcanoes due to sparse data distribution. The regional aeromagnetic data (Hunt and Whiteford, 1979) flown at 1050 m, however, show positive anomalies associated with the volcanic centres, reaching maxima of up to 600 nT over the volcanic peaks. A reconnaissance gravity profile across the Tama Lakes saddle was carried out by Zeng and Ingham (1993) and was interpreted (without constraints from other geophysical data) in terms of a basement high below the saddle. MT data collected along the same profile identified an anomalously low resistivity zone which was initially attributed to alteration/mineralization (Ingham and Zeng, 1993). A more recent, and more extensive, MT survey using modern measurement and processing techniques, and aimed at investigating the deep electrical structure of Ruapehu volcano, has been reported by Ingham et al. (2008). 4. Gravity and magnetic surveys A total of 38 new gravity stations established across the TgVC through the Tama Lakes saddle (Fig. 2) were integrated with 8 existing regional
Fig. 2. Geological map of the Tama Lakes saddle area (from Fig. 1) showing the location of the geophysical surveys; magnetotelluric sites are indicated by station numbers and triangles.
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stations (Woodward and Ferry, 1974). Elevations of the new gravity stations were determined either by optical levelling or GPS to an accuracy of ±0.5 m or better, equivalent to an uncertainty of ±1 g.u. in the gravity data (note: 1 g.u. = 0.1 mGal). Gravity data were corrected with a density of 2670 kg m− 3 (i.e. that of the greywacke basement; Reilly,1972) and fully terrain corrected. The most significant sources of error in the corrected gravity data are in the elevation and terrain corrections, especially where the topography is severe; the total error in the gravity data is estimated to be b10 g.u. The regional gravity field for the area was determined from 56 existing gravity measurements (Woodward and Ferry, 1974; Ferry et al., 1974) on basement greywacke surrounding the Tongariro region. This regional field was approximated to a third-order polynomial surface (similar to that of Zeng and Ingham,1993) which decreases to the southsoutheast with a gradient of 12 g.u. km− 1. The resulting residual gravity data (Fig. 3) projected onto the Tama Lakes saddle profile shows an asymmetric negative gravity anomaly across the volcanic axis, reaching a minimum of −190 g.u. Gravity gradients on the eastern limb of this anomaly are steeper (~18 g.u. km− 1) than on the western limb (~13 g.u. km− 1). Whilst the data are smooth across most of the profile, short wavelength variations 2–3 km wide and of about 15 g.u. in magnitude occur in the central section, probably reflecting near-surface variations in the density of the volcanic deposits. Aeromagnetic data (at elevations of 1500 m asl and 2500 m asl) were collected approximately coincident with the projected gravity profile at
a 2-second sampling interval (equivalent to a nominal 100 m spacing). Location and elevation were determined by differential GPS to an accuracy of 3–5 m. Diurnal variations were recorded at a local base station and subtracted from the survey data together with the International Geomagnetic Reference Field (IGRF). The resultant total field anomaly data at both elevations include a similarly subdued long wavelength regional component. A third-order polynomial surface was fitted to data (at 2500 m elevation) in the area remote from the volcanoes to approximate the regional field. This surface, which varies between 50 and 90 nT along the length of the profile, was subtracted from the aeromagnetic data at both elevations to give the residual anomaly data shown in Fig. 3. The residual anomaly data at 1500 m elevation are complex, with a central positive (325 nT) anomaly flanked by negative anomalies of about −100 nT. The width of the anomalous region is about 12 km. The residual anomaly data at 2500 m elevation show an asymmetric central positive anomaly about 8 km wide, reaching a maximum value of 230 nT. 5. Modelling gravity and magnetic data Published density and magnetisation data from the TVZ are somewhat scarce. Mean density values of 2730 kg m− 3 have been measured for littlealtered samples of Rotokawa andesite (Hunt and Harms, 1990) and 2720 kg m− 3 for andesites from the Tihia volcano (Sissons, 1981); similar values have been obtained for Taranaki andesite (2400–2830 kg m− 3;
Fig. 3. Gravity and magnetic models across the Tama Lakes saddle profile. (a) Observed residual gravity anomaly (diamonds) and calculated gravity effect (line) of the model below. (b) Cross-section through 2.75D gravity model. (c) Observed residual magnetic anomaly data at 2500 m elevation (diamonds) and calculated magnetic effect (line) of the model shown below. (d) Observed residual magnetic anomaly data at 1500 m elevation (diamonds) and calculated magnetic effect (line) of the model shown below. (e) Cross-section through 2.75D magnetic model (note that a slight difference in topography occurs between the two models as a consequence of positional differences).
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Locke et al.,1993). It should be noted of course, that the density of andesite depends significantly on the degrees of vesicularity and alteration. The density of volcaniclastic material is even more variable and therefore a bulk density is more difficult to quantify. Studies of volcanoes in New Zealand and elsewhere yield values of 2000–2400 kg m− 3 for volcaniclastics (Budetta et al., 1983; Brown et al., 1987; Williams et al., 1987; Fedi et al., 1991; Deplus et al.,1995); in the present modelling, an average value of 2200 kg m− 3 has been assumed. Late-Tertiary sediments in New Zealand have densities of about 2200 kg m− 3 (Whiteford and Lumb, 1975) and a density of 2670 kg m− 3 has been attributed to the greywacke basement (after Reilly, 1972); these values are incorporated in the models. Magnetic remanence intensities of andesites are typically very variable. For example, values for young fresh andesite flows on Egmont Volcano range from 0.2 to 40 A m− 1 (Downey et al., 1994). Koenigsberger (q) ratios are typically high, about 20 according to Cox (1970) and Downey et al. (1994), hence induced magnetisation is likely to be a minor component. Similar remanence intensities have been deduced for andesites elsewhere, e.g. 1–10 Am− 1 with q-ratios of 1–40 (Blakely and Christiansen, 1978; Flanagan and Williams, 1982). Modelling of the gravity and two aeromagnetic data sets, carried out in 2.75 D using Interpex MagixXL, was performed iteratively in order to find the optimum model that best fitted all three data sets. The calculated effects of this model, together with the observed data and a cross-section through the model, are shown in Fig. 3. The RMS discrepancies between calculated and observed values for the three data sets are: 8.5 g.u. for gravity (b5% of the maximum anomaly), 40 nT for the low-level aeromagnetics (10% of the maximum anomaly) and 14 nT for the high-level aeromagnetics (6% of the maximum anomaly). The optimum model (Fig. 3) comprises an extensive upper layer with density 2200 kg m− 3 and maximum thickness of 1000 m and of effectively infinite extent orthogonal to the profile. A single magnetisation for this layer cannot account for the observed data, hence a lower layer of 300 m thickness and zero magnetisation has been modelled below an upper layer of 1.3 A m− 1 magnetisation. The asymmetry of the gravity anomaly requires an abrupt depth change (modelled as 300 m) in this lowest interface at ~9 km profile distance, further abrupt changes in this interface of 300 m and 250 m are modelled at ~6 and 25 km profile distance, respectively, in order to fit the observed anomaly. The gravity data cannot be accounted for by assuming that the entire upper volcanic layer has a low bulk density (in the region of 2200 kg m− 3), there has to be denser material towards the centre of the profile. This is consistent with the aeromagnetic data which indicates the occurrence of highly magnetised rocks (usually associated with high densities) central to the profile; in addition, having magnetic data at two elevations significantly reduces the modelling ambiguity. The magnetic anomalies show that these central high-magnetisation/denser rocks are unlikely to be due to elevated greywacke basement as this is essentially nonmagnetic. At the surface, between profile distances 4.5 and 13 km, a dense (2400 kg m− 3) and magnetic (4 A m− 1) body is modelled that extends to the basement interface at about 9 km and 12 km profile distance. This body has been modelled as extending 4 km orthogonally either side of the profile. The low-level aeromagnetic data are the most sensitive to modelling and hence the most diagnostic; extensive trial-and-error modelling showed that a shallow magnetic body is required to fit the data. The observed short wavelength variations require that the lower surface of the more magnetic body be undulating i.e. the observed anomaly cannot be accounted for by a single deep steep-sided body. The magnetic lows flanking the positive anomaly are largely a consequence of the subvertical magnetic boundary within the low density layer. 6. Magnetotelluric soundings Broadband magnetotelluric soundings made at a total of 37 sites around Ruapehu were reported by Ingham et al. (2008). Measurements were made using 5-component Phoenix MTU5-2000 systems and were
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remote referenced, providing generally excellent data quality. Although the Tama Lakes saddle is close to the northern boundary of the MT survey area, six of the sites (marked by triangles in Fig. 2) lie close to the gravity and aeromagnetic profiles. The electrical resistivity structure derived across the Tama Lakes saddle from modelling of data from these six sites is considered here. As reported by Ingham et al. (2008), the MT data were subjected to phase tensor analysis (Caldwell et al., 2004) in order to both identify the dimensionality of resistivity structure at each site, and to remove galvanic distortions from the data in the manner proposed by Bibby et al. (2005). As is discussed by Ingham et al. (2008), for periods less than about 10 s the data are 1-dimensional and at each site may be adequately interpreted in terms of a layered resistivity structure. As a result, following removal of galvanic distortion effects, 1-D models which satisfy the short period data at each site were derived from the determinant impedance. Shallow features identified by these models were confirmed with a forward 3-D numerical model (Cairns, 2006) which incorporated topography. A compilation of the 1-dimensional models for the sites across the Tama Lakes saddle are shown in Fig. 4, superimposed on the gravity/ magnetic model. Note that a few sites lie up to 3 km from the gravity/ magnetic profile and at higher elevations (sites 111 and 112 in particular), therefore some caution is required in correlating the MT section, however it is believed to be broadly representative of the resistivity structure across the Tama Lakes saddle. At all of the sites except 115 (the furthest NW), the basic resistivity structure at depth is resistor–conductor–resistor. Whereas the depth to the conductive layer is generally well constrained, the thickness and resistivity of the conductor are non-unique due to equivalence (i.e. a model with the same conductivity-thickness product will also fit the data). Consequently, the depth to the basal resistive layer is not well constrained. This is possibly highlighted at the south-east end of the profile where it is likely (as shown in the gravity/magnetic model) that the observed low resistivity layer thins out as outcropping basement is approached. The two most striking features of the compilation of 1-D models are: (i) a layer of anomalously low resistivity (10–20 Ωm) beneath the central four sites which thickens sharply to almost 1 km thickness beneath site 111, and (ii) a layer of intermediate resistivity lying above the low resistivity layer at sites 114 and 111 and immediately below the high density, high magnetisation surface cover. It is also interesting to note that higher surficial resistivities appear to be associated with these sites. 7. Interpretation Modelling of the different data sets has defined a number of significant geophysical structures which can be correlated. The greywacke basement modelled from gravity data correlates with the lowermost high resistivity layer in the MT model; similar resistivities have been reported for greywacke elsewhere (e.g. Ingham et al., 2001). A relative deepening of the basement surface from west to east is common to both models and depths broadly agree (within probable uncertainties) except notably in the centre of the profile. The low density layer in the gravity model broadly correlates with the low-intermediate resistivity middle layer, modelled from the MT data. Furthermore, the magnetic model differentiates two components within this layer, consisting of an upper relatively magnetised layer and a lower non-magnetised layer. These are interpreted as an upper layer of mixed volcanic deposits (including lahar and pyroclastic deposits) as mapped over much of the profile (Grindley, 1960), together with a lower layer of Tertiary sediments. The combined modelling shows that the volcanic deposits are up to 800 m thick and that the Tertiary sediments vary from a few 10's to a few 100 m in thickness. Three discontinuities in the basement–Tertiary interface, at approximately 6, 9 and 24 km profile distance, have been modelled from the gravity data. The easternmost of these correlates with the mapped
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Fig. 4. Magnetotelluric model compilation superimposed on gravity/magnetic model (from Fig. 3). The individual 1D models showing depths to layers (as short bars) with modelled resistivity values are plotted below the respective MT stations (marked by arrows with station numbers, as referred to in text); tie lines (dotted) and hatching are used to outline areas of similar resistivity. The projected locations of the mapped surface traces of the key faults are also shown for comparison with the model.
surface trace of the Rangipo fault (Grindley, 1960; Rowland and Sibson, 2001). The central discontinuity (at 9 km profile distance) correlates with the Waihi Fault (see Fig. 1) and the western discontinuity correlates with an unnamed mapped fault set (Rowland and Sibson, 2001). Vertical displacements of 250–300 m are modelled on each of these discontinuities, which is less than the 800 m displacement deduced for the Rangipo Fault further south (Villamor and Berryman, 2006a) but of a similar magnitude to that expected on faults in the centre of the graben to the south of Ruapehu (Villamor and Berryman, 2006a). The resolution of the gravity model, however, does not preclude these basement dislocations being accommodated by a series of smaller offsets rather than single fault. The near-surface, high density/high magnetisation body in the centre of the profile correlates with the high resistivities modelled beneath sites 114 and 111 and is interpreted as volcanic deposits with a high proportion of lavas and/or dykes, consistent with geological mapping in the region (Grindley, 1960). The greater apparent lateral extent of the high resistivity layer is probably a consequence of the offline location of the MT sites which lie on or closer to mapped lavas. This high density/high magnetisation body is modelled with two locally steepened structures extending to greater depth, one coincides with the basement discontinuity at 9 km profile distance and the other with the axis of the 10 ka volcanic vents (Nairn et al., 1998). Both of these features penetrate through the zone of intermediate resistivity into the thickened anomalously low resistivity layer and may reflect magma conduits extending into the basement but cannot be resolved as separate features by the sparse MT site distribution. The striking 5 km-wide, deep low resistivity zone at the centre of the profile (noted earlier) may result from enhanced fracture porosity within the basement. Equally, the occurrence of magma feeder systems might suggest that this low resistivity zone (and its extension at shallower levels) results at least in part, from conductive clays produced by hydrothermal alteration of the overlying volcanics. Such alteration could of course have been facilitated by the occurrence of related enhanced permeability. Similarly, conductive clays have been reported by Ingham et al. (2008) and Jones et al. (2008), for example, to be widespread on the higher slopes of Ruapehu volcano. Any occurrence of enhanced porosity or alteration appears to be insufficiently intense to be resolved by the gravity or magnetic data.
8. Conclusions and implications Integrated interpretation of multiple geophysical data sets across the Tama Lakes saddle has successfully delineated a number of key features of the subsurface structure of the Tongariro Volcanic Centre to 2 km depth. These being: 1. A minimum basement subsidence of 650 m at the centre of the rift relative to the margins of the rift. 2. Significant offsets in the basement interface that correlate with the mapped surface traces of the Waihi Fault and an unnamed fault (mapped 3 km to the west). However, no offset at the rift axis (i.e. the locus of the PM vents) is apparent from the geophysical data. 3. A deeper region of enhanced fracture porosity and/or alteration within the basement at the centre of the rift. 4. Steep-sided bodies of volcanic rocks with a high proportion of lavas/ dykes, that are coincident with both the Waihi Fault and the rift axis. Basement faulting is concentrated in the centre of the rift (apart from the Rangipo Fault at the eastern margin), there is no geophysical evidence for any significant faulting at the western margin. Across the Tama Lakes saddle, the rift is shown to be 18 km wide and therefore the N–S segmentation from a 15 km- to 40 km-wide modern Taupo Fault Belt as suggested by Villamor and Berryman (2006b) must occur at a more southerly location (i.e. closer to Ruapehu volcano) than they proposed. The zone of anomalously low resistivity (at ~1–2 km depth) interpreted as resulting from enhanced fracture porosity and/or alteration in the basement may relate to the deeper, more extensive low-velocity zone identified from seismic tomography which was also attributed to possible fracturing (Rowlands et al., 2005). This low resistivity zone coincides with the estimated depth of N1 km at which the PM magmas were quenched, based on vesicularity studies (Nairn et al., 1998). Groundwater for this magma–water interaction was suggested by Nairn et al. (1998) to be sourced from either andesite intrusives in Tertiary sediments or basement greywacke at depths of N1 km; the geophysical modelling indicates that the latter is more probable. There is no geophysical evidence for any large volume high-density bodies that might indicate relict magma chambers associated with
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volcanism at shallow levels below this region (c.f. the Taranaki volcanoes; Locke and Cassidy, 1997). This is consistent with the notion that the PM magmas rose rapidly from deep levels such that large volumes of magma did not reside for any lengthy period at high levels at this location (Nakagawa et al., 1998). The coincidence of one of the zones of denser volcanic rocks with both the Waihi Fault and a significant basement offset, suggests that this fault system may have provided magma pathways to the surface and a focus for dyke emplacement which could have contributed to extension of the rift. In contrast, there is no resolvable offset in the basement interface at the rift axis, where the PM vents and the second dense volcanic body occur. The PM eruptions are considered to be the result of a regional rifting episode which tapped small magma bodies (Nairn et al., 1998; Nakagawa et al., 1998), therefore it is perhaps surprising that no substantial basement offset is apparent at the rift axis from the geophysical data. This may reflect the juvenile nature of faulting at this location and be indicative of the migration of faulting towards the centre of the graben, as has occurred further north (Villamor and Berryman, 2006b). Alternatively, rifting across the Tama Lakes saddle during the PM event could have been entirely accommodated by dyke emplacement. Acknowledgements We thank The University of Auckland Research Committee, Earthwatch Centre for Field Research, and Environment Waikato for financial support; the Department of Conservation for allowing access to the area; our pilot the late Alan Smallfield; Colin Yong for technical assistance; University of Auckland graduate students and Earthwatch volunteers for assistance with gravity surveys; Stacey Dravitzki, Yasuo Ogawa and Stewart Bennie for help with MT measurements; Louise Cotterall for assistance with manuscript preparation and Julie Rowland for providing information for Fig.1. John Gamble and Glyn Williams-Jones are thanked for helpful reviews. References Acocella, V., Spinks, K., Cole, J., Nicol, A., 2003. Oblique back arc rifting of Taupo Volcanic Zone, New Zealand. Tectonics 22 (4), 1045. doi:10.1029/2002TC001447. Berryman, K., Villamor, P., 1999. Spatial and Temporal Zoning of Faulting in the Taupo Volcanic Zone, New Zealand. Geological Society of New Zealand Miscellaneous Publication, vol. 107A, p. 15. Bibby, H.M., Caldwell, T.G., Davey, F.J., Webb, T.H., 1995. Geophysical evidence on the structure of the Taupo Volcanic Zone and its hydrothermal circulations. J. Volcanol. Geotherm. Res 68, 29–58. Bibby, H.M., Caldwell, T.G., Brown, C., 2005. Determinable and non-determinable parameters of galvanic distortion in magnetotellurics. Geophys. J. Int 163, 915–930. Blakely, R.J., Christiansen, R.L., 1978. The magnetisation of Mt Shasta and implications for virtual geomagnetic poles determined from seamounts. J. Geophys. Res 83 (B12), 5971–5978. Brown, G.C., Rymer, H., Thorpe, R.S., 1987. The evolution of andesite volcano structures: new evidence from gravity studies in Costa Rica. Earth Planet. Sci. Lett 82, 323–334. Bryan, C.J., Sherburn, S., Bibby, H.M., Bannister, S.C., Hurst, A.W., 1999. Shallow seismicity of the central Taupo Volcanic Zone, New Zealand: its distribution and nature. N.Z. J. Geol. and Geophys 42, 533–542. Budetta, G., Nunziata, C., Rapolla, A., 1983. A gravity study of the island of Vulcano, Tyrrhenian Sea, Italy. Bull. Volcanol 46, 183–192. Cairns, P.J., 2006. 3-D modelling of the electrical resistivity structure of Mount Ruapehu. Unpublished MSc Thesis, Victoria University of Wellington, Wellington, New Zealand. Caldwell, T.G., Bibby, H.M., Brown, C., 2004. The magnetotelluric phase tensor. Geophys. J. Int 158, 457–469. Cole, J.W., Graham, I.J., Hackett, W.R., Houghton, B.F., 1986. Volcanology and petrology of the Quaternary composite volcanoes of Tongariro Volcanic Centre, Taupo Volcanic Zone. In: Smith, I.E.M. (Ed.), Late Cenozoic Volcanism in New Zealand Royal Soc. N.Z. Bull. 23, pp. 224–250. Cox, A., 1970. Remanent magnetisation and susceptibility of late Cenozoic rocks from New Zealand. N.Z. J. Geol. Geophys 14, 192–207. Deplus, C., Bonvalot, S., Dahrin, D., Diament, M., Harjono, H., Dubois, J., 1995. Inner structure of the Krakatoa volcanic complex (Indonesia) from gravity and bathymetry data. J. Volcanol. Geotherm. Res 64, 23–51. Downey, W.S., Kellett, R.J., Smith, I.E.M., Price, R.C., Stewart, R.B.,1994. New Palaeomagnetic evidence for the recent eruptive activity of Mt Taranaki, New Zealand. J. Volcanol. Geotherm. Res 60, 15–27.
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