Fault controls on the geometry and location of the Okataina Caldera, Taupo Volcanic Zone, New Zealand

Journal of Volcanology and Geothermal Research 190 (2010) 136–151

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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

Fault controls on the geometry and location of the Okataina Caldera, Taupo Volcanic Zone, New Zealand H. Seebeck a,b,⁎, A. Nicol b, T.A. Stern a, H.M. Bibby b, V. Stagpoole b a b

School of Earth Sciences, Victoria University of Wellington, Wellington, New Zealand GNS Science, PO Box 30368, Lower Hutt, New Zealand

a r t i c l e

i n f o

Available online 3 May 2009 Keywords: caldera structure faults gravity Okataina

a b s t r a c t Okataina Caldera is located within the Taupo Rift and formed due to collapse following eruptions at 325 and 61 ka. Gravity, seismic reflection, topographic and geological data indicate that active rift faults pass into the caldera and have influenced its location and geometry. The caldera has a minimum gravity anomaly of − 50 mGal, is elongate north–south with an inferred minimum depth to caldera floor of 3 ± 0.5 km at the rift axis, and occupies a 10 km hard-linked left step in the rift. The principal rift faults (55–75° dip) define the location and geometry of the northwest and southeast caldera margins and locally accommodate piecemeal collapse. Segments of the east and west margins of the caldera margin are near vertical (70–90° dip), trend north–south, and are inferred to be faults formed by the reactivation of a pervasive Mesozoic basement fabric (i.e. faults and/or lithological contacts). The fault sets which define the caldera geometry predate it, while the step in the rift across the Okataina Volcanic Centre (OVC) is at least as old as the caldera. Within the OVC displacement on rift faults induced by gravitational caldera collapse at 61 ka exceeds tectonic displacement since this time by at least a factor of two. Collapse along pre-existing rift faults and, in particular, Mesozoic basement fabric are important for caldera formation elsewhere in the Taupo Volcanic Zone. © 2009 Elsevier B.V. All rights reserved.

1. Introduction Rhyolitic calderas typically form through roof collapse of an underlying shallow magma reservoir during explosive eruptions greater than a few cubic kilometres (Lipman, 1997; Cole et al., 2005). Caldera structures in rift environments have a variety of map (i.e. sub-circular to elongate) and cross sectional (e.g. piston, piecemeal, or trapdoor) geometries (Cole et al., 2005). A spectrum of different generic models can be found in the literature to account for the geometries of caldera collapse structures (e.g. Cole et al., 2005 and references therein). End-member models proposed for these caldera collapse geometries (surface and sub-surface) range from purely magmatic (e.g. Bosworth et al., 2003) to dominantly fault controlled (e.g. Ferguson et al., 1994) (Fig. 1). The diverse subsidence patterns and geometries of calderas are inferred to reflect the varying size and depth to the source magma chamber(s), the orientations of the principal stress axes (both local and regional) (Fig. 1A and B), the influence of pre-existing faults and structural fabrics (Fig. 1B and C), and/or prior roof and collapse margin trends (e.g. Wilson et al., 1984; Ferguson et al., 1994; Lipman, 1997; Gudmundsson, 1998; Moore and Kokelaar, 1998; Acocella et al., 2002; Prejean et al., 2002; Bosworth et al., 2003; Cole et al., 2005). Due to the explosive nature of some ⁎ Corresponding author. GNS Science, PO Box 30368, Lower Hutt, New Zealand. Tel.: +64 4 570 1444; fax: +64 4 570-4600. E-mail address: [email protected] (H. Seebeck). 0377-0273/$ – see front matter © 2009 Elsevier B.V. All rights reserved. doi:10.1016/j.jvolgeores.2009.04.011

caldera volcanoes it is important to understand the primary control(s) on caldera development and collapse style in a particular location. While faults have been inferred to influence the geometry of calderas in rifts (e.g. Ferguson et al., 1994; Moore and Kokelaar, 1998), their spatial and temporal relations are often poorly resolved. To understand the interrelations between caldera and rift formation in extending continental crust, the intersection of the Okataina Volcanic Centre and Taupo Rift, Taupo Volcanic Zone, New Zealand, has been examined in detail (Fig. 2). This paper investigates the geometries of the Okataina Caldera complex and rift faults using analyses of surface faulting traces, published 1:50 000 geological maps (Nairn, 1989, 2002), reinterpreted seismic reflection lines (Davy and Bibby, 2005), and modelling of new (this study) and existing gravity data sets (New Zealand Land Gravity Database). These geological and geophysical data sets allow us to examine the interrelationship between late Quaternary faulting and two rhyolitic caldera formation events at ca. 325 and 61 ka (Wilson et al., 1995; Nairn, 2002; Cole et al., 2010-this issue). We suggest the geometry and location of the caldera structure has been strongly influenced by rift faults and tectonic fabric (i.e. pre-existing faults and stratigraphic boundaries) within Mesozoic basement rocks. 2. Geological setting The young Taupo Volcanic Zone (TVZ) (340 ka to present) is a 250 km northeast–southwest trending volcanic rift of predominantly andesitic to rhyolitic arc/back arc volcanism within continental crust of

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Fig. 1. Caldera formation models. (A) Borehole breakout model for the formation of elliptical topographic caldera margins. Independent ambient stress field data indicate caldera elongation parallel to the upper crustal minimum horizontal stress direction (Shmin) (Bosworth et al., 2003). (B) Fault-caldera model where the kinematic interaction of active and pre-existing crustal fabric influences the geometry of the caldera collapse structure (e.g. Acocella et al., 2002). (C) Caldera collapse geometry associated with rift fault orientation (e.g. Ferguson et al., 1994; Moore and Kokelaar, 1998). In all models, maximum caldera subsidence is indicated by grey shading.

the central North Island of New Zealand (Cole,1990; Wilson et al.,1995). The young TVZ is characterised by frequent large rhyolitic eruptions (30–500 km3 dense rock equivalent) and a high effective heat flow (~800 mW/m2) (Fig. 2 inset) (Stern, 1985; Cole, 1990; Bibby et al., 1995; Wilson et al., 1995). The entire TVZ formed over the past ~2 Myr in response to the 42–48 mm/yr oblique subduction of the Pacific Plate beneath the Australian Plate along the Hikurangi Margin east of New Zealand (Fig. 2 inset) (Wilson et al., 1995; Beavan et al., 2002). Late Permian to Early Cretaceous basement, dominated by steeply dipping (N60°) greywacke and argillite terranes, underlies much of the North Island (Mortimer, 2004) (Fig. 3). Mesozoic terrane boundaries and bedding gradually swing clockwise in strike from north–northwest adjacent to the Hauraki Rift (Fig. 3) (i.e. divergent from the trend of the TVZ by up to 60°), to north–northeast (i.e. parallel to the volcanic rift zone) south of Lake Taupo (Mortimer, 2004; Nicol et al., 2007). The curvature of the basement terranes is replicated by the strike of the main Cenozoic faults east and west of the young TVZ and, where present beneath the TVZ, may have an important control on the structural geometry and location of rhyolitic calderas. Several young TVZ caldera complexes are spatially coincident with a 15–40 km wide zone of active crustal extension (Fig. 3) (Cole, 1990; Wilson et al., 1995; Rowland and Sibson, 2001; Villamor and Berryman, 2001), which is here referred to as the Taupo Rift (after Nicol et al., 2006; Villamor and Berryman, 2006). The Taupo Rift has been the locus for multiple caldera forming events in the Taupo, Whakamaru and Okataina volcanic centres over the past 340 ka (Wilson et al., 1984, 1995). The sub-surface location and geometry of TVZ calderas have primarily been identified through negative gravity anomalies (Rogan, 1982; Wilson et al., 1984; Stern, 1985; Davy and Caldwell, 1998; Stagpoole and Bibby, 1999; Milner et al., 2002; Stratford and Stern, 2008). Infill of caldera collapse structures by low density volcanoclastic products is interpreted to account for the majority of the observed gravity anomaly associated with each volcanic centre (e.g. Wilson et al., 1984; Davy and Caldwell, 1998), while regional faults have been implicated in the control of major structural caldera collapse margins across the central North Island (e.g. Davy and Caldwell, 1998; Milner et al., 2002; Smith et al., 2006). The Taupo Rift comprises a dense array of normal faults striking northeast–southwest typically dipping 60–80° at the ground surface with scarp heights ranging up to 550 m in volcanic units ≤340 ka in age (Anderson and Webb, 1989; Wilson et al., 1995; Rowland and Sibson, 2001; Villamor and Berryman, 2001, 2006, Nairn, 2002; Acocella et al., 2003; Nicol et al., 2006). The Taupo Rift accommodates present-day extension and crustal thinning induced by the clockwise rotation of the

eastern North Island associated with rollback of the subducting Pacific Plate and/or continental collision at the southern end of the Hikurangi margin (Ballance, 1976; Stern, 1987; Walcott, 1987; Wallace et al., 2004; Mortimer et al., 2007; Nicol et al., 2007). Despite the obliquity of the relative plate motion, the direction of extension in the TVZ is approximately orthogonal to the dominant strike of the rift faults (e.g. Rowland and Sibson, 2001; Hurst et al., 2002; Acocella et al., 2003). The rift generally forms a well defined graben and along-strike has been interpreted to comprise of a series of segments (Rowland and Sibson, 2001; Villamor and Berryman, 2001; Acocella et al., 2003). Segment boundaries mark 10–20° changes in the average strike of the faults and/ or steps in the margins of the active rift, the most prominent of which occurs across the Okataina Volcanic Centre (OVC) (Fig. 2) (Rowland and Sibson, 2001; Acocella et al., 2003). Historical seismicity in the young TVZ is typically clustered near to these rift segment boundaries, the margins of caldera structures and normal faults (Bibby et al.,1995; Bryan et al., 1999; Rowland and Sibson, 2001; Hurst et al., 2002). The OVC is located on the eastern margin of the young TVZ, the active arc front associated with the Hikurangi subduction margin, and is the most recently active of the central rhyolitic eruptive centres (Cole, 1990; Wilson et al., 1995; Nairn, 2002) (Fig. 2). The OVC is one of at least eight caldera complexes formed within the TVZ over the past 1.6 Ma by large explosive rhyolitic eruptions thought to be driven by basaltic flux from the mantle (Wilson et al.,1984; Houghton et al.,1995; Annen and Sparks, 2002; Smith et al., 2004; Shane et al., 2005; Bachmann et al., 2007). The OVC has been active for the past c. 530 kyr (Cole et al., 2009-this issue) and can be described as a complex of coalescing collapse structures formed during at least two large caldera forming events, the 150 km3 Matahina (ca. 325 ka) and the ~100 km3 Rotoiti (61 ka) pyroclastic eruptions (combined volume ~50% of the total erupted from the centre) (Houghton et al.,1995; Nairn, 2002; Wilson et al., 2007; Cole et al., 2010this issue). Volcanic activity within the complex for the last 22 kyr has been associated with post-caldera rhyolitic dome building culminating in a 17 km long basaltic fissure eruption across Mt. Tarawera in 1886 AD (Fig. 2) (Smith, 1886; Cole, 1970; Nairn, 2002). Rhyolitic magma produced during dome building and caldera forming events is inferred to have resided in chambers at depths of ~6–16 km beneath the OVC (Leonard et al., 2002; Nairn et al., 2004; Smith et al., 2004; Bannister et al., 2004; Shane et al., 2005; Bibby et al., 2008). Active faults of the Taupo Rift appear to terminate against the western and eastern topographic caldera margins (Fig. 2) (Nairn, 1989), while both soft (Rowland and Sibson, 2001) and hard fault linkage (associated with a right bend) (Acocella et al., 2003) models have been proposed for the OVC. The presence of Mesozoic basement

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Fig. 2. Simplified volcanic stratigraphy of the Okataina Volcanic Centre (OVC), the northern-most rhyolitic volcanic centre in the TVZ, North Island, New Zealand (Data from Nairn, 1989; New Zealand Active Faults database; Leonard and Begg (unpublished QMAP data, 2008); Begg and Mouslopoulou, 2010-this issue). Location of main map shown in inset. Recent post-caldera dome building (b 22 ka) situated within the topographic margin of the caldera delineates the Haroharo and Tarawera linear vent zones. The OVC is partially defined by the distribution of pre-caldera rhyolite lavas up to ca. 530 ka in age (Nairn, 1989, 2002; Cole et al., 2010-this issue).

at depths of ~ 1 km in geothermal drill-holes and in outcrop within 20 km of the OVC to the north, south, and east are consistent with the caldera being underlain by greywacke basement (e.g. Nairn and Beanland, 1989; Nairn, 2002; Bibby et al., 2008) (Fig. 2). Receiver function analysis adjacent to, and beneath, the OVC suggests a crustal thickness of 25–30 km (Bannister et al., 2004), while further to the southwest in the central TVZ, wide-angle seismic reflection data suggests a quartzo-feldspathic crust as thin as 15 km (Harrison and White, 2006; Stratford and Stern, 2006).

3. Methods and data The results presented here primarily draw upon gravity modelling, mapping of faults using topographic traces, and reinterpretation of seismic reflection lines from Davy and Bibby (2005). These new data are used in conjunction with geological information (Nairn, 1989, Villamor and Berryman, 2001; Nairn, 2002) to constrain the geometries of faults in the rift, the Okataina Caldera and the relationship between the two.

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Fig. 3. Mesozoic basement terranes of the central North Island. Mesozoic basement terrane boundaries after Mortimer (2004, 2006). Mesozoic and Tertiary bedding strikes (Edbrooke, 2001, 2005) indicated by rose diagrams collated in bins of 50 km of north–south width and extending from the east and west boundaries of the TVZ to the coast and map boundary. Note the north–south fault and fabric orientation at the latitude of the OVC east–west across the central North Island (New Zealand Active Faults Database; Leonard and Begg unpublished QMAP data, 2008). Abbreviations: OVC, Okataina Volcanic Centre; NIFS, North Island Fault System.

3.1. Gravity Residual gravity anomaly data (after Stagpoole and Bibby, 1999; Stern, 1979) have been used to estimate the minimum depth to the top of Mesozoic basement in the region of the OVC (Fig. 4) (here referred to as top basement). Gravity models in the TVZ have often been used to estimate the depth to top basement utilising the difference in density between basement (2670 kg/m3) and overlying volcanoclastic infill (typically 1800–2200 kg/m3) (e.g. Modriniak and Studt, 1959; Stern, 1979; Rogan, 1982; Stern, 1986; Stagpoole, 1994; Stagpoole and Bibby, 1999; Seebeck, 2008). The density of volcanoclastic infill can approach basement densities particularly at depths of N1.5 km and within geothermal areas (e.g. Stern, 1986). Thus, gravity modelling can produce a reasonable estimate for the geometry of the Okataina Caldera margins (i.e. the main locus of collapse), but because the caldera floor is deeper than 1.5 km may not provide a robust estimate

of its depth. For further discussion of the limitations of gravity studies in defining TVZ basement refer to Stern (1986). The Okataina Caldera covers an area of ~ 400 km2 with the geometry of the top basement modelled using approximately 250 gravity measurements (Fig. 4). Of these measurements 108 were collected by Seebeck (2008) with the remainder from the New Zealand Land Gravity Database (Appendix). These new data are tied to the New Zealand Primary Gravity Network to ensure consistency between the two datasets. The gravity data modelled here exclude 116 gravity stations of Rogan (1982) collected within and adjacent to the OVC (Fig. 4) as these data were not tied to a New Zealand Primary Gravity Network station and therefore are not compatible with our data. However, these excluded data do provide additional information on gravity gradients around the caldera margins and have been utilized for this purpose. Despite these differences in methodology, our first order results are comparable to those of Rogan (1982) (e.g.

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Fig. 4. Residual gravity anomaly map of the Okataina Volcanic Centre (OVC). New gravity measurements (filled circles) and the New Zealand Land Gravity Database (crosses) reduced to Bouguer anomalies using the method of Reilly (1972). Note gravity data of Rogan (1982) (open squares) not included in map construction (refer to text). Residual gravity anomalies derived by the removal of the regional gravity field of Stern (1979). Mesozoic greywacke basement outcrops to the north and east of the OVC (Fig. 2), has been intersected at depths of ~ 1 km in the Kawerau geothermal field, and underlies the volcanic cover of the Kaingaroa Plateau (Healy et al., 1964; Bibby et al., 1995). Location of rift perpendicular (NE–SW) and rift parallel (NW–SE) gravity profiles in Fig. 6 are shown.

the size and shape of −45 mGal contour associated with the caldera are similar). Simple one and two and a half-dimensional gravity models can be used to estimate the depth to basement and the geometry of the caldera margins (e.g. Cordell, 1973; Rogan, 1982; Ferguson et al., 1994; Stagpoole, 1994; Davy and Caldwell, 1998). Multi-one dimensional gravity modelling, where constrained, provides a smoothed firstorder depth estimate to top basement within and adjacent to the OVC (Fig. 5). Two and a half-dimensional gravity modelling can estimate subsurface structure, however this type of gravity model can over or

under estimate basement depth and exaggerate subsurface structure (e.g. Rogan (1982) estimated 5 km of subsidence associated with the caldera using a single layer 2D model). By using both types of gravity model in tandem a more robust caldera geometry can be established than would be possible using each type of model in isolation. A series of 1D gravity models calculated using an exponentially decreasing density contrast Δρ(z) with depth of Cordell (1973) and the gravity contribution of an infinite Bouguer sheet after Stern (1986) (refer to Appendix for details) were examined against an independent dataset of basement depths (with associated residual gravity anomalies)

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topography along each profile was obtained from the altitude at each gravity station. Rock densities used for 2 1/2 D gravity modelling were 2200 kg/m3 for rhyolitic dome complexes, 2150 kg/m3 for near surface volcanic material (ignimbrite and tuff), 2350 kg/m3 for older volcanic material, and 2670 kg/m3 for basement. The initial surface density used for 2 1/2 D modelling differs from that of the 1D analysis as the density layers here represent depth averaged densities. The simple two-layer density structure is sufficient to model the depths to basement consistent with the Kawerau geothermal drillholes and resistivity model (i.e. 3 ± 0.5 km) (Fig. 6B and C), accounting for the observed gravity anomaly of the caldera. Simple gravity modelling of magma accumulation beneath the OVC capable of producing a large ignimbrite eruption (i.e. a sphere or cylinder with a top depth of 6 km, volume 113 km3, and density contrast of − 300 kg/m3), has a gravity contribution of b−3 mGal and cannot be resolved as it is within observational error (i.e. ±3 mGal). 3.2. Fault mapping

Fig. 5. Preferred 1D gravity model calculated with an exponentially decreasing density contrast with depth (λ) for initial an surface density contrast of 670 kg/m3 (after Cordell, 1973). Refer to Appendix for derivation. Independent constraints for depth to basement (density of 2670 kg/m3) derived from TVZ geological and geophysical data constrain the most suitable initial density contrast and depth decrement (i.e. λ). A third order polynomial fitted to the curve provides a minimum depth to basement verses residual gravity estimate (see Appendix). Interpolation of residual gravity data using this function provides a first order estimate of basement throughout the OVC region (not shown).

from drill-hole and geophysical data (seismic reflection, refraction, and magnetotelluric studies) collected outside of geothermal areas (Seebeck, 2008). The preferred relation between modelled gravity and depth is shown in Fig. 5. This gravity model uses an average TVZ surface density contrast relative to basement (i.e. −70 kg/m3), predicts the residual gravity anomaly for the nearest subsurface density data to the OVC to a depth of 1 km (i.e. Kawerau geothermal drillholes in Fig. 2), and is consistent with depths to basement derived from Vp tomography (i.e. transition from 4 to 5 km/s) (Sherburn et al., 2003) and magnetotelluric studies (Bibby et al., 2008) in the Okataina region (e.g. 2–4 km). The interpretation of a preliminary magnetotelluric model (Bibby et al., 2008) allows the composition, density structure, and depth to caldera floor to be interpreted with greater confidence than using gravity modelling alone. The preliminary inversion of magnetotelluric data across the Okataina Caldera (Fig. 6A) show conductive and/or porous material to depths of 3 km, underlain by a highly resistive basement (Bibby et al., 2008). The low resistivities observed to 3 km depth are interpreted here as conductive volcanoclastic infill material and thus provides an independent estimate for the shape and depth of the caldera floor. This thickness of volcanic infill has been observed elsewhere in the TVZ. In the Mangakino Caldera, for example, welded ignimbrite has been intersected by a drillhole at depths of up to 3.2 km (Spinks et al., 2005a) and supports the interpretation of conductive volcanoclastic compositions at these depths within some TVZ calderas. Gravity models along two profiles (Fig. 6B and C) using a simple two-layer density structure (developed using ENCOM Modelvision pro software) were constructed for the Okataina Caldera using twodimensional bodies to represent surface volcanic rocks and simplistic three dimensional shapes to represent deeper volcanic rocks. The

Active faults of the Taupo Rift intersect Okataina Caldera along its western margin (Nairn, 1989; Villamor and Berryman, 2001; Nicol et al., 2006; Berryman et al., 2008). The locations and geometries of these faults have been mapped using a combination of topographic analysis and reinterpretation of seismic reflection data acquired in Lake Tarawera (Davy and Bibby, 2005) (Fig. 7). Topographic analysis derived from aerial photographs and digital elevation models (DEM) provides detailed information on the locations, lengths, and displacements of active normal faults exposed at the ground surface along with general information on the topographic margin of the caldera. Active normal faults within the Taupo Rift form scarps that range in age up to ~340 ka (e.g. Villamor and Berryman, 2001; Nairn, 2002; Nicol et al., 2006). Fault scarps higher than ca. 1–1.5 m and longer than 200 m have been mapped across the study area towards the topographic caldera margin using orthorectified aerial photographs at approximately 1:17 000 scale. Faults recorded in the topography are biased towards those areas were volcanic and fluvial surfaces have not been significantly modified by erosion or burial. Scarp burial may, for example, account for some of the decrease in density of fault traces from southwest to northeast within several kilometres of the topographic margin of the caldera (Fig. 7). In order to assess whether rift faults terminate against the western topographic caldera margin or continue into the caldera complex we have reanalysed seismic reflection data from Lake Tarawera. Six seismic reflection lines delineate faults and caldera structures post240 ka at depths of up to ~ 900 m within the OVC (Fig. 7). The seismic lines are sufficiently closely spaced to produce a crude structure contour map of key reflectors and to correlate the larger faults between lines and with lake shore geology (Nairn, 1989, 2002). Details of the seismic data (acquisition and processing) and interpretation are given in Davy and Bibby (2005) and Seebeck (2008), respectively. Seismic reflection data indicate the presence of many normal faults beneath Lake Tarawera and within the topographic margin of the caldera. Sub-lake faults, some of which have been correlated with onshore faults southwest of the lake, appear to form a central graben which bounds a zone of subsidence highlighted by the geometry of the top Mamaku Ignimbrite (Fig. 7 inset) (‘M’ horizon of Davy and Bibby, 2005). The faults beneath Lake Tarawera, including the central graben, are interpreted to represent the northeast continuation of the rift into the caldera complex. No seismic reflection lines were collected perpendicular to the rift within the primary caldera collapse structure (Fig. 4); therefore it is not possible to determine whether rift faults extend into the primary caldera margin of the 61 ka Rotoiti eruption (Nairn, 2002; Davy and Bibby, 2005). Seismic reflection lines and topographic analysis of TOPSAR digital elevation models enable the measurement of fault throws in, and adjacent to, the OVC. Throws have been used to measure displacement

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Fig. 6. Rift perpendicular and parallel cross-sections of the Okataina Caldera (refer to Fig. 4 for location). (A) Preliminary interpretation of a resistivity model across the OVC (Bibby et al., 2008). The MT profile coincides with a section of the rift perpendicular gravity profile shown in (B). Dashed line approximates the depth extent of low resistivity volcanoclastic infill. (B) Rift perpendicular gravity profile across the OVC. (C) Rift parallel gravity profile across OVC. The low density volcanic infill (2150 kg/m3 and 2350 kg/m3) is modelled as an infinite Bouguer sheet. Densities below 2 km are modelled as simple 3-dimensional objects (e.g. out of plane 12 km strike length) to incorporate the 3D nature of the caldera. Caldera margins are consistent with dips between 55–75° and 70–90° in the rift perpendicular and parallel profiles respectively. Arrows denote the mapped topographic caldera boundaries and the Tarawera and Haroharo linear vent zones of Nairn (1989, 2002). Horizontal and vertical scale 1:1.

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Fig. 7. Active fault map of the southwest margin of the Okataina Volcanic Centre (OVC). (inset) Reinterpreted seismic reflection lines of Davy and Bibby (2005) indicate subsurface faulting within the topographic margin of the caldera. Structural contours (in metres below lake level) for the ca. 240 ka Mamaku Ignimbrite (Davy and Bibby, 2005; Gravley et al., 2007) beneath Lake Tarawera shows a re-entrant subsidence feature extending SW from the 61 ka collapse boundary (Davy and Bibby, 2005) partially bounded by normal faults. The gravity derived primary caldera collapse boundary is consistent with the seismically defined collapse boundary of Davy and Bibby (2005).

variations along the Whirinaki Fault zone (see Figs. 7 and 8 for location) and into the Okataina Caldera. Displacement rates have been estimated for horizons of 61, 188, and 240 ka in age that cross the fault zone by measuring scarp heights and vertical displacement of seismic reflectors along each individual fault. Displacements were measured on 20 transects (each ~ 3 km long) normal to fault strike covering a zone of 22 km along strike southwest of the OVC (Fig. 8). Scarp heights measured in the field (using tape and level or RTK GPS) and from the

TOPSAR DEM typically have a standard error of ±1.31 m. Estimates for depths of seismic horizons are typically ±5–10 m (dependent on the velocity model used). 4. Caldera geometry The primary margins of the Okataina Caldera are well defined by the residual gravity anomaly map (Fig. 4) which shows a 9 × 15 km

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Fig. 8. Whirinaki Fault zone displacement rates along-strike towards Okataina Caldera. (A) Location map showing Taupo Rift active fault traces including Whirinaki Fault zone (black lines) and displacement transects relative to caldera boundaries. (B) Displacement rates NE along-strike for the Whirinaki Fault zone approaching the primary caldera margin for the across-strike total (black line) and southern fault strand (grey line). Errors estimated as the square root of the sum of the squares for horizon displacement and age uncertainties. (C) Schematic model of fault displacement interpretation across caldera structure, here a constant cumulative tectonic vertical displacement rate along the Whirinaki Fault zone increases by N50% upon entering the caldera complex through the reactivation of rift faults during collapse events (Fig. 7).

elongate negative residual anomaly within the topographic margin of the Okataina Caldera (Nairn, 1989, 2002). Both the gravity anomaly and the topographic margin of the caldera are elongate in an approximately north–south direction (long axis 175 ± 5° eccentricity E = L min / L max = 0.6). The minimum residual gravity anomaly (−50 mGal), which is smaller in amplitude than the minimum anomalies associated with the Mangakino and Taupo caldera complexes (−65 mGal) (Stagpoole and Bibby, 1999), is located along the projection (from the southwest of the caldera) of the Taupo Rift axis, and bounded to the southeast and northwest by the Tarawera and

Haroharo linear vent zones, respectively (Fig. 9) (Nairn, 2002). The locus of maximum subsidence for the Okataina Caldera in the southcentral region of the topographic caldera margin is consistent with the caldera collapse history proposed by Nairn (2002). Nairn (2002) infers that the Okataina Caldera formed in response to at least two spatially overlapping collapse events, one rift parallel and one basement fabric parallel (at 325 ka and 61 ka respectively). These observations suggest that the magma chambers formed during each collapse event may have been in slightly different locations (map view) within the caldera complex. The inferred zones of collapse for

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Fig. 9. Structural interpretation of the OVC in relation to the Taupo Rift. Residual gravity contours (Fig. 4) highlight structural trends and are indicative of relative basement altitudes. The presence of north–south and northeast–southwest faults within, and adjacent to the OVC is supported by regional seismicity (NZ Earthquake catalogue 2000–2008) and the double-difference location of earthquakes of Hurst et al. (2008). (A) Right-bend pure-shear model of Acocella et al. (2003) and Spinks et al. (2005b) with direction of principal horizontal extension (arrow) aligned in direction of inferred rift opening. (B) Left-stepping rift model proposed in this study where reactivation of north–south striking basement fabric hard-links the step in the active rift. Direction of principal horizontal extension (arrow) from GPS and stress inversion (Hurst et al., 2002; Beavan et al., 2007; Hurst et al., 2008) is oblique (α 30°) to the caldera long-axis.

these two events overlap in the region of maximum subsidence determined by gravity and along a northeast projection of the rift axis (Fig. 10). Multiple 2 1/2 D gravity profiles (N = 5) with different orientations across the volcanic centre enable the geometry of the primary collapse structure to be estimated, this structure is distinct from the mapped topographic margin with a similar geometry to the inner ring fault inferred by Nairn (1989). This primary collapse structure is approximated by the −45 mGal gravity contour and delineates the location of

maximum subsidence associated with caldera collapse (Fig. 4). These data suggest that within the OVC, residual gravity gradients across the maximum low are steep in a rift parallel direction (northeast–southwest) (−4.4± 0.4 to −15.8 ±0.5 mGal/km) and shallower perpendicular to the rift (northwest–southeast) (−2.0 ±0.5 to −4.7± 0.5 mGal/km). The gravity profile perpendicular to the Taupo Rift (Fig. 6B) is consistent with a southeast primary caldera margin dipping 55–75° across which top basement steps down at least 3 km to the northwest. A second step in the gravity to the southeast of the primary caldera

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Fig. 10. Schematic diagram of the main structural and volcanic elements of the OVC. The axis of the Taupo Rift is defined by fault dip directions and minimum residual gravity anomalies along strike. The intersection geometry of Taupo Rift faults (NE–SW) with basement structural trends (N–S) (i.e. North Island Fault System — NIFS) is highlighted on the SE margin of the caldera. Basement form lines parallel the structural trends in Fig. 3.

margin, which is consistent with magnetotelluric data (Bibby et al., 2008), bounds the Kaingaroa Plateau and is approximately coincident with the southern topographic caldera boundary as indicated by Nairn (1989). The gravity defined intersection of caldera floor with the south-eastern primary caldera margin is estimated to within ±1 km (laterally) approximately 3.5 km vertically beneath the Tarawera linear vent zone (Fig. 6B) (Nairn, 1989, 2002). The primary northwestern margin of the caldera is less well constrained due to the limited number of data points. However, the decrease in gravity values towards the maximum gravity low occurs at least 4 km southeast of the topographic caldera margin. One interpretation of these data is that the northwest margin steps down to the southeast along two or more collapse faults each with a vertical displacement of 1 km or less (e.g. Stagpoole, 1994). Therefore, the northwest and southeast primary caldera margins define an asymmetric graben with the largest displacement along the southeast margin (Fig. 6B) (vertical displacement of 2 km and 3 km respectively). The northwest and southeast gravity-defined inner caldera margins identified on the rift perpendicular gravity profile coincide approximately with the projected locations of the rift faults (e.g. Ngapouri Fault, Figs. 9 and 10) mapped at the ground surface outside the south-western margin of the caldera. The location of the caldera within the rift, the spatial coincidence of the northwest and southeast margins of the caldera with the projected locations of important rift faults, and the inferred 55–75° dips of the northwest and southeast caldera margins all suggest that the geometries of these caldera margins are defined by rift faults (see Section 5 for discussion of relative timing of faulting and caldera formation).

The west and east margins of the caldera can be seen in the rift parallel gravity profile (Fig. 6C). The primary collapse structures defined by gravity in this profile have well constrained locations and are consistent with near vertical faults (70–90°), steeper than the southeast and northwest margins of the caldera and rift faults. The east and west primary margins of the caldera strike north to north– northwest. It is inferred, therefore, that the locations and geometry of the primary east and west margins of the caldera were not controlled by normal faults which strike parallel to the rift. The north–south trending east and west primary margins of the caldera are sub-parallel to some faults in the rift (Fig. 7) and to a zone of seismicity within, and to the north of, the caldera (Fig. 9). North–south trending alignments of well located earthquakes are also observed within the Haroharo linear vent zone crossing the north-eastern margin of the caldera (Hurst et al., 2008). In addition, depth dependent anisotropy (Eberhart-Phillips & Reyners, in press) also supports a model with a north-south and northeast-southwest structural fabric at depths of 4 km on the eastern margin of the TVZ. These data indicate the presence of a subtle, but important, north–south structural trend in the region of the OVC. This trend is approximately parallel to the strike of the basement terranes and bedding east and west of the OVC (Fig. 3). It is proposed here that the orientations and locations of the north to north–northwest striking margins of the caldera have been controlled by Mesozoic basement fabrics (i.e. faults and lithological boundaries) of a similar orientation. It is suggested that both faults within the rift and caldera collapse structures (defined by both gravity and topography) utilised pre-existing planes (or zones) of weakness in the basement (cf. Tommasi and Vauchez, 2001) in a similar manner

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to the North Island Fault System (N20 km east of the OVC). These planes of weakness are parallel to, and inferred to be defined by, faults and lithological boundaries in Mesozoic basement which typically dip between 70 and 90° (i.e. similar to dips of the east and west segments of the caldera margins). 5. Rift and caldera interaction Seismic reflection lines, geomorphic mapping and gravity modelling all suggest that normal faults within the active rift extend into the OVC. The geometry of the Okataina Caldera appears to have been strongly influenced by a combination of rift faults and reactivation of Mesozoic basement fabric. The resulting caldera, which is elongate in an approximately north–south direction, is oblique (α 30°) to the ~ 145° orientation of the regional extension direction (Hurst et al., 2002; Beavan et al., 2007) (Fig. 9). The extent to which the geometry of the caldera was influenced by the geometry of underlying melt accumulation(s) is yet to be determined. However, if both the basement fabric and rifting predated caldera formation the possibility remains that the geometry of the caldera and location of magma chamber(s) were influenced by faulting. Active faults, caldera margin geometries, and minimum basement depths indicate that the Okataina Caldera is located at a segment boundary in the rift across which the margins and axis of the rift step to the left by ~10 km (Fig. 9). The length of the left-step is defined by the distance between the southern most active rift faults on either side of the caldera (Fig. 9). The east and west margins of the step in the rift are inferred to be faults which hard-link the rift across this step (Fig. 10) (cf. Walsh and Watterson, 1992; Rowland and Sibson, 2001) and utilise planes of weakness (e.g. stratigraphic boundaries or faults) in Mesozoic basement. Therefore, the step is likely to be at least as old as the caldera. Maximum caldera subsidence (≤− 45 mGal) coincides with the centre of the left-step and the projected position of the rift axis within the volcanic complex. The left-stepping rift model differs significantly from the right bending rift model of Acocella et al. (2003) (Fig. 9), which is not supported by a detailed analysis of faulting and minimum basement depths adjacent to the OVC. Gravity gradient variations along the caldera margin suggest that its geometry, and therefore the style of collapse, varies with margin orientation. In rift parallel gravity profiles the east and west primary margins of the caldera are steep and are best approximated by piston style collapse (Fig. 6C) (cf. Lipman, 1997). However, the rift perpendicular gravity profile is best described by asymmetric piecemeal collapse (cf. Lipman, 1997) focused along the southeast margin of the caldera (Fig. 6B) (e.g. Nairn, 2002). Piecemeal collapse occurs on several faults and appears to be important where the location and geometry of the margin of the caldera is controlled by spatially distributed rift faulting. Localisation of collapse along the southeast margin of the caldera is consistent with asymmetry of the rift northeast and southwest of the OVC where the Edgecumbe and Paeroa faults (Fig. 10), respectively, accommodate the greatest tectonic displacements (Villamor and Berryman, 2001; Mouslopoulou et al., 2008). The similarity in asymmetry of both the rift and the caldera suggests that the south-eastern margin of the caldera has accommodated tectonic displacement and collapse. Gravity modelling is consistent with the northeast continuations of the Ngapouri and Ngakuru faults forming the boundaries to the inner caldera collapse structure during the 325 ka Matahina eruption (Fig. 10) (Nairn, 2002; Cole et al., 2010-this issue). The rift faults may have propagated into the caldera after 325 ka, nucleated on the caldera margins during (or after) collapse, or predate this caldera forming event. A post-caldera fault model in which the rift faults post-date the caldera would require that the fault and caldera margin relations (i.e. their similar geometries and locations) occurred by chance which seems unlikely given that caldera margins are typically near vertical in the absence of rift faults (cf. Lipman, 1997). A syn-collapse fault model for

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the 325 ka event, where faults initiated on the caldera margin during collapse, could account for the co-location of faults and caldera margin, but again is likely to produce near vertically dipping faults (rather than the 55–75° dips estimated from gravity). Therefore, we suggest that rift faults may have extended into, and predated, caldera formation at 325 ka, with the faults being utilised within the caldera complex to accommodate collapse. A model in which rift faults predated caldera formation also appears to best account for the fault-caldera margin relations formed in the 61 ka Rotoiti eruption. Outside and along the western topographic margin of the caldera, rift faults (which pass into the caldera) display higher displacements on ~ 180–300 ka horizons than the ~ 61 ka surface, from which we infer that these faults predate 61 ka collapse. Structure contours on the inferred top of the c. 240 ka (Gravley et al., 2007) Mamaku Ignimbrite (‘M’ horizon of Davy and Bibby, 2005) show a small re-entrant along the western margin of the primary collapse structure, which is bounded by rift faults (Fig. 7) that appear to have been locally utilised to accommodate collapse. The notion that caldera collapse at ~61 ka utilised pre-existing tectonic rift faults is further supported by displacement rate variations along the Whirinaki Fault zone (Fig. 8) which enters the caldera along its southwest margin. Cumulative displacement rates along what we interpret to be the Whirinaki Fault zone, for the last ~60–240 ka, are typically in the range of 1 to 1.5 mm/yr outside, and to the southwest of, the topographic caldera margin. Between the topographic caldera margin and the primary collapse boundary cumulative displacement rates increase up to 3–3.5 mm/yr despite a decrease in the number of faults sampled. An example of this abrupt rate increase is observed along the southern-most strand of the Whirinaki Fault zone (Fig. 8), which has vertical displacements of up to 80 m. The southern-most strand has constant 0.25 mm/yr displacement rate over 5 km approaching the topographic margin of the caldera. A sharp increase in the displacement rate on the southern-most strand occurs across the topographic caldera margin towards the inner collapse structure. The topographic caldera margin of Nairn (1989) is offset across the southern-most fault strand (Fig. 8) suggesting that during caldera collapse at ~ 61 ka, greater subsidence occurred in the hanging wall of this fault than the footwall. Both the displacement rates and geometry of the topographic caldera margin indicate that the Whirinaki Fault zone was utilised during caldera collapse at ~61 ka (Fig. 7). Within the caldera, displacement across the Whirinaki Fault zone during gravitational caldera collapse at 61 ka exceeded tectonic displacement since this time by at least a factor of two. As caldera collapse is a geologically instantaneous process, the rates of fault displacement due to gravitational collapse must be several orders of magnitude higher than the tectonic displacement rate, which was approximately constant (i.e. varying by a factor of two or less) over time periods of hundreds of thousands of years (Villamor and Berryman, 2001; Nicol et al., 2006). 6. Discussion and conclusions Gravity, seismic reflection, topographic and geological data indicate that active rift faults pass into the Okataina Caldera complex. The geometry of the Okataina Caldera, which formed during collapse events at 325 and 61 ka, has been influenced by a combination of rift faulting and reactivation of basement fabrics. The caldera, which is elongate in a north–south direction and at least 3 ± 0.5 km deep, is confined to the Taupo Rift. The Okataina Caldera occupies a ~10 km left step in the Taupo Rift which is hard-linked by north–south faults. The left-step in the rift across the OVC is as old as, or predates, the caldera. The northwest and southeast boundaries of the main gravity-defined caldera (i.e. below −45 mGal) are consistent with dips of 55–75° and appear to be controlled by the principal rift faults whereas the east and west margins of the caldera are consistent with near vertical faults (70–90°), striking approximately north–south. These north-south caldera margins are

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inferred to reactivate Mesozoic basement fabric. The geometry of the caldera margin and the style of collapse are variable. Piecemeal collapse is important where the location and geometry of the margin of the caldera is controlled by spatially distributed rift faulting. Within the OVC, displacement on rift faults induced by gravitational caldera collapse at 61 ka exceeds tectonic displacement since this time by at least a factor of two. Outside the OVC, caldera collapse does not appear to have increased fault displacements. The role of rift faults and basement fabric at Okataina also applies to other calderas in the central North Island (Fig. 11). Regional faults are, for example, inferred to bound the margins of major structural collapse for the Taupo, Rotorua and Waihi calderas (Fig. 11) (e.g. Davy and Caldwell, 1998; Milner et al., 2002; Smith et al., 2006) and many authors have suggested a relationship between faulting, volcanism, and/or geothermal activity in the TVZ (e.g. Wilson et al., 1984; Cole, 1990; Wilson et al., 1995; Wood, 1995; Rowland and Sibson, 2001; Wood, 2001; Acocella et al., 2003; Spinks et al., 2005b; Villamor and Berryman, 2006). Preliminary analysis of calderas throughout the TVZ suggests that their geometries and relationship with faults is variable (Fig. 11). The

eight sub-circular to elongate rhyolitic TVZ caldera complexes, identified through volcanic mapping, vent location, and residual gravity expression, have variably orientated long axes with respect to the regional extension direction and/or Taupo Rift (Wilson et al., 1984, 1995; Stagpoole and Bibby, 1999; Seebeck, 2008). The Kapenga Caldera complex is elongate parallel to, and located within, the rift, while the southeast and northwest margins of Okataina and Mangakino calderas are also parallel to rift faults. However, perhaps the most striking feature of TVZ calderas is the number with margins parallel to the trend of Mesozoic basement fabric. The volcanic centres of Okataina, Whakamaru, Rotorua, Reporoa and Mangakino (defined by gravity and/or topography) have caldera elongation or topographic margins trending north–south (Nairn et al., 1994; Wilson et al., 1995; Stagpoole and Bibby, 1999; Seebeck, 2008). In addition, residual gravity and magnetic anomalies associated with the western inner collapse margin of Taupo Caldera are aligned north–south (e.g. Davy and Caldwell, 1998). North–south and northeast–southwest aligned post-caldera domes are situated on the margins of, and within the Whakamaru Caldera (Healy et al., 1964; Wilson et al., 1984). These observations suggest that reactivation of Mesozoic basement fabric

Fig. 11. Residual gravity map of the central North Island (including the TVZ) with inferred basement form lines. Mapped caldera complexes of the TVZ (and/or associated gravity contours) are predominantly elongate north–south and/or bounded by active rift faults. Data: Wilson et al. (1995); Davy and Caldwell (1998); Edbrooke (2005); Smith et al. (2006); Gravley et al. (2007); Leonard and Begg (unpublished QMAP data, 2008); New Zealand Active Faults Database; New Zealand Land Gravity Database. Silicic volcanic centre abbreviations: Ka, Kapenga; Ma, Mangakino; Oh, Ohakuri; Ok, Okataina; Ro, Rotorua; Rp, Reporoa; Tp, Taupo; Wi, Waihi; Wh, Whakamaru.

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occurred during the formation of a number (i.e. 6 of 8) of calderas and that the orientation of this fabric had an important control on caldera geometry in the TVZ.

contrasts at depth increments of 100 m from each initial density contrast to a total depth of 5 km. A cumulative 1-D model gravity anomaly, assuming an infinite Bouguer plate configuration, is calculated using a basement density of 2670 kg/m3 defined by:

Acknowledgments This paper is the result of studies completed with financial support from Victoria University of Wellington, the Royal Society of New Zealand Marsden Fund (GNS 303) and the New Zealand Public Good Science and Technology Fund. Bryan Davy, Nick Mortimer, Vasso Mouslopoulou, Ian Nairn, Euan Smith, Pilar Villamor, and John Walsh are thanked for their data and/or discussions. John and Catherine Ford are thanked for their hospitality and continuing support. Graham Leonard kindly provided a review of an early version of this manuscript. Julie Rowland and Roberto Weinberg are thanked for the helpful reviews. Appendix A. Gravity analysis A total of 108 new gravity measurements were collected in the vicinity of the OVC (at ~ 1 km spacing) using a La Coste and Romburg G Meter (Fig. 4). Differential GPS tied to a base station of known elevation provided altitude control for each new gravity station (to an accuracy of ±0.15 m). These new gravity data were tied to the New Zealand Primary Gravity Network and reduced via methods consistent with the New Zealand Land Gravity Database (Reilly, 1972) using a density of 2670 kg/m3 for Bouguer and Terrain corrections. Terrain corrections have been applied to each gravity measurement to a distance of 22.9 km. These new data have been incorporated into an existing onshore Bouguer gravity dataset covering an area of 60 × 103 km2 (13,560 stations) extracted from the New Zealand Land Gravity Database for the central North Island. The estimated accuracy of the calculated Bouguer anomaly data from this study is ±0.1 mGal and is predominantly due to uncertainty in near zone terrain correction estimates (Hammer zone D b 170 m). The accuracy of gravity anomaly data from the New Zealand Land Gravity Database is primarily dependent on the method used for altitude determination. Errors of up to ±3 mGal could be expected for data where altitude has been derived from barometric levelling (i.e. with estimated altitude errors of ±10 m). All stations within the New Zealand Land Gravity Database have been checked against a 10 m grid digital elevation model (DEM) and removed when observed and DEM heights differ by N20 m. A third-order, two-way, regional gravity field derived from Bouguer gravity anomalies observed either side of the TVZ on geologically uniform Mesozoic greywacke basement has been subtracted from the Bouguer gravity anomaly to provide a residual gravity anomaly for the region (Stern, 1979; Stagpoole and Bibby, 1999). One dimensional gravity models have been calculated using the exponentially decreasing density contrast Δρ(z) with depth of Cordell (1973) (e.g. Fig. 5): ΔρðzÞ = Δρ0 e

− λz

Here z is positive vertically downward in kilometres, λ represents the compressibility of the rock in reciprocal length units, and Δρ0 represents the density contrast between the basin infill at the surface and the basement. Cordell (1973) observed that the residual gravity effect of a Bouguer slab with decreasing density contrast with depth tends to a finite limit, placing a limitation on the resolution of the depth estimate over deep structures. Three density contrasts were considered for the range of TVZ surface densities. Two end members and an intermediate (800 kg/m3, 550 kg/m3, and 670 kg/m3 respectively) were based on previous gravity studies and analysis of TVZ rock density data. A range of λ between 0.3 and 0.8 km− 1 were used to calculate theoretical density

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Δg = 2πG

XN i − 1

ðhi − hi − 1 Þðρi − 2670Þ

after Stern (1986) Here G is the gravitational constant, ρi and hi are the density (kg/m3) and depth of the ith sample respectively,Nis the number of samples and h0 is the ground surface. A series of curves for different values of λ can then be produced for each of the initial density contrast. An independent set of constraints (e.g. published drill-hole, seismic refraction, reflection, and tomographic data along with preliminary magnetotelluric models) to basement depth is then used to select the most appropriate curve for each initial density contrast. An initial surface density contrast of 670 kg/m3 with λ = 0.4 km− 1 provided the best-fit (least squares) to the independent constraints (Fig. 5). The residual gravity–depth curve is not adequately approximated by an exponential function and therefore a third order polynomial provides a residual gravity verses depth to basement function: 3

2

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