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JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 113, B01403, doi:10.1029/2007JB005170, 2008

Regional structure and kinematic history of a large subduction back thrust: Taranaki Fault, New Zealand V. Stagpoole1 and A. Nicol1 Received 14 May 2007; revised 8 August 2007; accepted 8 September 2007; published 17 January 2008.

[1] The Taranaki Fault is a back thrust antithetic to the Hikurangi margin subduction

thrust. Subduction back thrusts, like the Taranaki Fault, accrue displacement transferred from the subducting plate, and growth analyses of these structures contribute to an improved understanding of subduction processes. The Taranaki Fault forms the eastern margin of the Taranaki Basin and is part of a system that extends for at least 600 km in continental crust of western New Zealand. The fault is preserved beneath young sedimentary cover and provides a rare opportunity to investigate the geometry and kinematic history of a large subduction back thrust. Two-dimensional seismic reflection lines (2–5 km spacing), tied to recently drilled wells and outcrop, together with magnetotelluric and gravity models are used to examine the fault. These data indicate that the fault is thick skinned with dips of 25–45° to depths of at least 12 km. The fault accommodated at least 12–15 km of dip-slip displacement since the middle Eocene (circa 40–43 Ma). The northern tip of the active section of the fault stepped southward at least three times between the middle Eocene and early Pliocene, producing a total tip retreat of 400–450 km. The history of displacements on the Taranaki Fault is consistent with initiation of Hikurangi margin subduction during the middle Eocene, up to 20 Ma earlier than some previous estimates. Fault tip retreat may have been generated by clockwise rotation of the subduction margin and associated progressive isolation of the fault from the driving downgoing Pacific Plate. Citation: Stagpoole, V., and A. Nicol (2008), Regional structure and kinematic history of a large subduction back thrust: Taranaki Fault, New Zealand, J. Geophys. Res., 113, B01403, doi:10.1029/2007JB005170.

1. Introduction [2] Fold and thrust belts can form within the overriding plate at subduction margins [e.g., Karig and Sharman, 1975; Davis et al., 1983; Moore, 1989; Barnes and de Le´pinay, 1997; von Huene and Klaeschen, 1999]. These structures develop in association with plate convergence [e.g., Forsyth and Uyeda, 1975; Backus et al., 1981] and may provide information about the kinematics and history of subduction. Outside of accretionary wedges, back thrusts, that form antithetic to the subduction thrust, also accrue displacement transferred from the subducting plate. Analysis of displacements on these back thrusts and associated growth strata provide an opportunity to resolve the timing and magnitude of subduction-related deformation and may help improve understanding of subduction processes. However, many of these contractional structures are poorly resolved, because crustal shortening typically results in their uplift and erosion (e.g., in the Andes and the Rocky Mountains). In contrast to these two regions, synsubduction Tertiary strata are widely preserved on the overriding plate of New Zealand’s Hikurangi subduction margin. These 1

GNS Science, Lower Hutt, New Zealand.

Copyright 2008 by the American Geophysical Union. 0148-0227/08/2007JB005170$09.00

growth strata provide important constraints on the timing and kinematics of deformation related to subduction [e.g., Barnes et al., 2002; Nicol et al., 2002, 2007]. [3] In this paper we examine the regional geometry and kinematic history of the Taranaki Fault, a large subductionrelated back thrust in New Zealand’s Taranaki Basin. Analysis of displacements of Late Cretaceous and younger strata, up to 8 km thick in the Taranaki Basin, afford a rare opportunity to better understand the circa 43 – 40 Ma kinematic history of the fault and highlight important relations between fault growth (and death, when the fault becomes permanently inactive) and associated subduction at the Hikurangi subduction margin. This study provides constraints for the development of the New Zealand plate boundary, while also demonstrating useful analysis techniques applicable to subduction zones in other parts of the world. [4] The Taranaki Fault is one of the longest (400 km) and highest displacement (>10 km) contractional structures in New Zealand’s continental crust [King and Thrasher, 1996]. The fault forms the boundary of pre-Miocene rocks in the Taranaki Basin and is part of a larger fault system that extends 600 km northward from the South Island’s Alpine Fault in western New Zealand (Figure 1). This fault system is contained entirely within continental crust of the overriding Australian Plate and includes the Waimea-Flaxmore Fault in the northern South Island and the Manaia Fault and

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STAGPOOLE AND NICOL: TARANAKI FAULT, NEW ZEALAND

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Figure 1. Regional map of Taranaki Fault. Map shows the location and geometry of the Taranaki Fault, Manaia Fault, Waimea-Flaxmore Fault, and Tarata Thrust. Contours (2, 4, 6 km) indicate depth below sea level to top basement (data modified from King and Thrasher [1996]). Locations of wells (circles) are shown. Inset shows plate boundary setting. the Tarata Thrust in Taranaki (Figure 1). Displacement within the fault system was driven by subduction of Pacific Plate along the Hikurangi margin and collision of continental crust along the Alpine Fault [Stern et al., 1993]. Currently, the relative motion between Australian and Pacific plates is oblique along the Hikurangi margin [DeMets et al., 1994; Beavan and Haines, 2001] (Figure 1); however, during the Eocene to Miocene relative plate motion was dominated by

convergence [King, 2000; Cande and Stock, 2004; Nicol et al., 2007]. Most of this convergence (i.e., >80%) accrued on the subduction thrust, with the remainder accommodated on upper plate faults, including the Taranaki Fault [Nicol et al., 2007]. [5] Displacement on the Taranaki Fault strongly influenced the location and geometry of the Taranaki Basin, New Zealand’s only producing oil and gas region. Petroleum

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Figure 2. Detailed map of Taranaki Fault and cross sections. Cross sections constructed at a high angle to the fault (modified from Thrasher et al. [1995]) show variations in fault geometry and Taranaki Basin architecture from north (A-A0) to south (D-D0). Map indicates locations of cross sections (bold lines), seismic lines (thin lines), wells (circles), and the line of magnetotelluric measurements (bold dashed line). All sections are at true scale. exploration in the basin has supplied a significant body of subsurface information on the Taranaki Fault, including seismic reflection and well data. The Taranaki Fault is not exposed at the surface, being overlain by up to 3 km of early Miocene and younger sedimentary rocks (Figure 2). It is, however, visible on many seismic reflection lines, in the regional gravity signature and in petroleum exploration wells [e.g., Mills, 1990; King and Thrasher, 1992, 1996; Palmer and Andrews, 1993]. We use two-dimensional (2-D) seismic reflection lines tied to petroleum exploration wells and outcrop to examine the Taranaki Fault where it displaces Late Cretaceous and younger strata. These data have been augmented with magnetotelluric and gravity profiles which provide constraints on the fault geometry at depth (e.g., 5 –12 km). Collectively these data sets indicate that the fault is a thick-skinned thrust which, south of Mount Taranaki, probably extends down to the base of the crust. The fault accrued displacement prior to the Oligocene, with the northern tip of the fault stepping southward over the last 30 Ma. This episodic migration is consistent with increasing isolation of the Taranaki Fault from the driving subducting Pacific Plate, perhaps induced by clockwise rotation

of the subducting plate, and with subduction commencing during the middle Eocene.

2. Data Sources and Analyses 2.1. Seismic Reflection Data [6] Seismic reflection lines provide information on the fault to depths of 4 – 7 km. Forty-three seismic reflection lines (including Figures 3, 4, and 5), which cross the fault and are distributed along 300 km of its length, have been interpreted (see Figure 2 for locations) with the aid of tie lines (not shown in Figure 2). The total interpreted line length is in excess of 1000 km, and the data are primarily from open file petroleum industry sources (available from New Zealand Ministry of Economic Development, http://www.crownminerals. govt.nz/cms/petroleum/technical-data). The seismic lines are all migrated, while some are prestacked depth-migrated sections (e.g., Figure 3). The age and quality of these seismic lines are variable, with Figures 3, 4, and 5 providing representative examples of the data. Most seismic lines are interpretable down to two-way traveltimes of 3 – 5 s. Up to eight reflectors were interpreted in each seismic line, including the top of seismic basement and up to six intra-Tertiary

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Figure 3. Prestacked depth migration (shown in two-way time) of (a) uninterpreted and (b) interpreted seismic line TRV-434 from the northern end of the fault (see Figure 2 for location). Locations and ages of interpreted horizons are indicated. Interpretations are based on data from Stagpoole [1997]. events. The interpreted reflectors range in age from about 2 to 80 Ma and have been tied to 20 nearby wells along the fault length. Our interpretations are broadly consistent with previously published data [Thrasher et al., 1995; King and Thrasher, 1996; Stagpoole, 1997]. 2.2. Magnetotelluric Data [7] A magnetotelluric (MT) survey was undertaken across the Taranaki Fault in 2003 [Stagpoole et al., 2006]. The resistivity contrast between Cretaceous and Cenozoic strata (3 – 30 ohm m) and the older Mesozoic basement rocks (50 – 1000 ohm m) make the fault a suitable target for investigation using the MT technique. The survey comprised 13 broadband soundings, measured using 3 Phoenix MTU-2000 instruments and a single four component recorder, along a 25 km west-east profile on the Taranaki Peninsula (Figure 2). Data were recorded for about 40 h at each site and modeled using the 2-D code of Rodi and Mackie [2001]. Modeling constraints were provided by seismic reflection data and borehole resistivity logs from

nearby exploration wells. Analysis of the MT phase tensor [Caldwell et al., 2004] indicates the data show 2-D character for periods up to 60 s, corresponding to depths of at least 12 km [Stagpoole et al., 2006]. 2.3. Gravity Data [8] There are over 10,000 onshore and offshore (marine track line) gravity observations in the region (Figure 6a). All onshore data (available at http://maps.gns.cri.nz/website/ gravity/) have been reduced to anomalies as described by Reilly [1972]. Marine track line data are available from the U.S. National Geophysical Data Center (http:// www.ngdc.noaa.gov/mgg/geodas/trackline.html). A map of the combined onshore Bouguer anomaly and offshore freeair anomaly data (Figure 6a) is dominated by a large NESW minimum, related to the Pacific Plate subducting beneath the North Island [Bannister, 1989; Stern et al., 1993], and positive anomalies in the north and west that relate to lithospheric structure [Stern et al., 1987] and the lack of subsidence following deposition of recently depos-

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Figure 4. Migrated (a) uninterpreted and (b) interpreted seismic line tp-09 (see Figure 2 for location on the peninsula). Locations and ages of interpreted horizons are indicated and based on data from the Toko-1 and Huinga-1 wells. ited sediments offshore in the Taranaki Basin [Holt and Stern, 1991]. In the vicinity of the Taranaki Fault, these regional long-wavelength effects (>50 km) tend to mask the shorter wavelength anomalies associated with upper crustal structure, and have been separated by computation of regional and residual gravity fields (Figures 6b and 6c) to reveal the Taranaki Fault. [9] The regional gravity field was computed using gravity data more than 25 km from the Taranaki Fault (Figure 6b) that were either observed on basement rocks or where the thickness of the sedimentary succession was known from seismic [Thrasher et al., 1995] and well data. A correction for the density deficiency of the sedimentary succession (based on the compaction curve for mudstone of Funnell et al. [1996] assuming a grain density of 2.7 Mg/m3) was applied to gravity values over the sedimentary basins on both sides of the fault (see Figure 6b for location of data points). Additional corrections have been made for the water depth of offshore data and for departures from the standard compaction curve due to uplift and erosion [Armstrong et al., 1998]. These data were then averaged

over 12 km grid cells and smoothed to generate the regional gravity map (Figure 6b). The residual gravity field, representing the gravity effect of upper crustal structure, was compiled by subtracting the regional gravity from the observed gravity. [10] Bouguer and free-air gravity anomaly data, with no regional-residual separation, were used for two-dimensional (2-D) models across the fault. The 2-D gravity models are relative to a standard density model with 28 km thick crust and 5 km of sediment (Figure 7), similar to the structure interpreted to occur at the western end of each line where the gravity anomaly is close to zero. The models are constrained by seismic, magnetotelluric and earthquake data. Information for the general sedimentary structure in the Taranaki, Wanganui and East Coast basins are from seismic reflection lines [Anderton, 1981; King and Thrasher, 1996; Field et al., 1997]. For the southern profile, deeper crustal structure is based on interpreted seismic data from southern Taranaki [Stern and Davey, 1990; Holt and Stern, 1994]. For the northern profile, crustal structure is

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Figure 5. Migrated (a) uninterpreted and (b) interpreted seismic lines P81-23 and P95-509 spliced together (see Figure 2 for location south of the peninsula). Locations and ages of interpreted horizons are indicated and based on data from the Kupe South-4 and Kupe-1 wells. based on the seismic interpretation of Harrison and White [2006] and Stratford and Stern [2006]. On both sides of the Taranaki Fault, the model is locally isostatically compensated (compensation depth of 80 km) with respect to the standard density model. There is some uncertainty in the crustal structure beneath the Taupo Volcanic Zone [Stratford and Stern, 2006; Harrison and White, 2006] and this region has not been modeled in detail. Earthquake hypocenters, relocated as part of the 3-D tomographic inversion [Reyners et al., 2006], which occur within 30 km of the lines, are shown on the models and correlate with the position of the subducting Pacific Plate beneath the North Island. Models are constructed in a similar manner to Bannister [1989] and Holt and Stern [1994], incorporating both the overriding and subducted plate structure beneath the North Island.

3. Fault Geometry [11] The Taranaki Fault strikes north-south from 40°30’S to 38°S a distance of about 250 km (Figure 1). At latitude 38°S (about 25 km north of the Te Ranga-1 well) there is a

change in strike of 25° to north-northwest. The fault can be traced on seismic reflection lines to at least latitude 37°S (a further 125 km). Northward from the Te Ranga-1 well fault displacements decrease, the fault is progressively buried by the sedimentary succession and can be masked by submarine volcanic cones. Locally the fault may change in strike by as much as 15– 25°. Notable changes in fault strike occur immediately west of Rotokare-1 and about 15 km south of Awakino-1 wells (Figure 1). These changes in strike coincide with locations where a thrust splays into the footwall of the main fault; the Tarata Thrust is one of these footwall splays. [12] Westward displacement on the fault in the Cenozoic has resulted in uplift and emplacement of a wedge of Mesozoic basement rocks over Late Cretaceous and Tertiary strata of the Taranaki Basin (e.g., Figures 3 – 5). The basement wedge comprises greywacke and argillite of the Murihiku Terrane, which is Late Permian to Late Jurassic in age [Raine et al., 2004]. Outcrop, exploration wells and seismic reflection lines indicate that the Taranaki Fault is approximately parallel to the strike and dip of bedding in the

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Figure 6. (a) Contours (10 mGal interval) of the onshore Bouguer and offshore free-air gravity anomalies and the data locations (dots), (b) contours of the calculated regional field and data points (small crosses) used for its compilation, and (c) shaded residual gravity field. See text for method used to create regional and residual gravity fields.

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Murihiku Terrane. West of the fault, basement rocks are inferred to be Permian age Brook Street Terrane [Spandler et al., 2005], and the boundary between terranes, which formed during Mesozoic accretion and faulting [Thrasher, 1990], probably lies at the Taranaki Fault [Mortimer et al.,

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1997; Stagpoole et al., 2006]. The location and geometry of the Taranaki Fault could therefore be strongly influenced in the upper crust by the location and orientation of the preexisting Murihiku-Brook Street terrane boundary.

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[13] The density contrast between basement rocks on the up-thrown eastern side of the fault and lower density Cretaceous and Cenozoic sediments in the west means the Taranaki Fault is marked by a strong residual gravity gradient that extends along the west coast of the North Island to latitude of 37° 300S (Figure 6c). North of 37° 300S the fault does not have a strong gravity expression, because of the greater sediment thickness over the fault and the smaller vertical throw seen on seismic data. Thicker sediment cover is also indicated by the lower gravity gradient between the negative anomalies associated with thick (3 – 7 km) low-density sediments in the Wanganui and Taranaki Basin. The fault relay zone between the Manaia and Taranaki faults is clearly visible on the residual gravity map (Figure 6c). [14] In cross section the fault defines the base of the basement wedge, while the top of the basement wedge comprises a ridge (variously referred to as the PateaTongaparutu High and Herangi High). Seismic reflectors and beds dip away from the crest of the basement high or ridge and indicate that it forms an anticline with a hinge 15 – 20 km east of the upper tip of the Taranaki Fault (Figure 2). Seismic reflection lines and well information indicate that, in many cases, the Taranaki Fault comprises multiple slip surfaces which may splay from the main fault surface within 2 – 6 km vertical distance of the upper tip. These thrust splays can occur entirely within Tertiary strata (e.g., Tarata Thrust cross section C-C0 in Figure 2 and the thrust at the Awakino-1 well in cross section B-B0 in Figure 2), produce interfingering basement and CretaceousTertiary strata (e.g., Te Ranga-1 well cross section A-A0 in Figure 2) or are confined to the basement wedge. In the Pukearuhe-1 well, for example, four thrust-bound blocks of basement occur in mixed stratigraphic order separated by 5 thrusts [Raine et al., 2004]. Such fault zone complexities are typical of thrusts or thrust systems with significant displacement (e.g., >10 km) [e.g., Boyer and Elliott, 1982] and appear to occur widely along the length of the Taranaki Fault. However, individual splays tend not to extend further than a few 10s of kilometers along strike; for example, the Tarata Thrust comprises segments ranging from about 10 to 40 km in length. [15] The average dip of the principal Taranaki Fault surface to depths of 3 – 7 km ranges from 20 to 50°. These dips were estimated along approximately 300 km of the fault length using depth-converted seismic lines crossing the fault (Figure 8). They take account of compaction of the Late Cretaceous-Tertiary succession beneath the fault, which resulted in an average reduction in fault dip of about 4°, and of the occasional obliquity between seismic line and fault strike. The resulting range of fault dips may be partly

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due to uncertainties in the measurements, including those arising from the seismic velocity model, which are collectively estimated to be ±7°. These uncertainties are, however, insufficient to account for all of the variability in fault dip (up to 20 –30°) over strike distances of 5 – 80 km. Therefore the data also capture local and regional variations in dip along strike (Figure 8) but do not include up-fault changes in dip along individual sections, such as the apparent shallowing of dip at the upper fault tip (e.g., Figure 3). There is a spatial coincidence between low fault dips and low altitudes of the western tip of the thrust wedge (Figure 8), which is consistent with the suggestion that some of the along-strike variability in dip may be limited to the upper crust (Figure 8 inset diagram). For example, using 2 km for the variation in the altitude of the tip of the thrust wedge and a general change in fault dip from 30° on the structural lows to 45° on the structural highs, we estimate that corrugations on the fault surface could extend to depths of 6 –7 km. [16] Fault dips in the upper crust are consistent with those estimated for depths of 5 – 12 km using MT and gravity models. Two dimensional models of MT data [Stagpoole et al., 2006] show a zone of low resistivity (3 – 30 ohm m), corresponding to Late Cretaceous and Tertiary sediments overlying higher resistively (20 – 300 ohm m) basement rocks (Figure 7). A large change in sediment thickness (5 km) between western and eastern sides of the model occurs in the region of the Taranaki Fault. Models also predict a discontinuity in basement resistivity at the fault. The discontinuity extends at a dip of about 45° to the base of the models (12 km), well beyond the penetration depth of seismic reflection data (Figure 7). This resistivity contrast between basement rocks on eastern and western sides of the fault indicates a difference in electrical properties, possibly relating to basement lithology, and supports the interpretation of Mortimer et al. [1997] that the Taranaki Fault is developed on the Murihiku-Brook Street terrane boundary. [17] Two 2-D gravity models across the North Island (Figure 7) are also consistent with the interpretation of an eastward dipping fault. The prominent (100 to 150 mGal) NE-SW minimum that dominates the gravity anomaly distribution over the North Island [Robertson and Reilly, 1958; Bannister, 1989] occurs at the eastern end of each gravity profile (Figure 7) and follows the trend of seismicity in the subducting plate. Gravity modeling indicates that no more than about 60 mGal of this negative anomaly can be attributed to the low-density sediments in the East Coast and Wanganui basins (Figures 6 and 8), with the remaining 50 to 100 mGal regional gravity anomaly inferred to be related to mass deficiencies beneath the sedimentary section. Uncertainties in the subsediment density distribution

Figure 7. Two-dimensional gravity (top and center) and magnetotelluric (bottom left) models of the Taranaki Fault. In the gravity models Figure 7 (top) shows the Bouguer and free-air gravity data (crosses) and the predicted gravity (line) for the structural cross section in Figure 7 (bottom). The gravity anomaly of the cross section is calculated relative to the crustal model at lower right (densities are in Mg/m3). The location map shows the lines of data used for modeling. Gravity models are constrained by seismic reflection data (in the Taranaki Basin) and earthquake hypocentres [Reyners et al., 2006] within 30 km of each profile (crosses) that mark the position of the subducting slab. The northern gravity profile is also constrained by wide-angle seismic data [Harrison and White, 2006; Stratford and Stern, 2006] that have been used to define the base of the crust east and west of the Taupo Volcanic Zone (bold lines). The magnetotelluric model is overlaid on southern gravity profile in alignment with the position of the fault. 9 of 19

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Figure 8. Variation in dip of the Taranaki Fault from 36.8°S (zero distance) southward to about 40.4°S (400 km distance). Dips are average decompacted values calculated for the upper 5 – 7 km of the crust using depth-converted seismic sections. Velocity information for depth conversion was obtained from seismic stacking velocity and well check-shot data (Awakino-1, Te Ranga-1, Ariki-1, Te Kumi-1) and have an estimated error of ±5%. Decompaction analysis [Sclater and Christie, 1980] on the depthconverted data indicates that the burial depth of tip of the fault has increased by between 400 and 900 m due to compaction, resulting in an average reduction in fault dip of 4°. Errors are estimated to be less than ±7°. The locations of seismic lines in Figures 3, 4, and 5, the Tarata Thrust and Manaia Fault, and four wells are indicated. Grey shaded line approximates the elevation of the western tip of the thrust wedge (see Figure 9) and shows the spatial coincidence between low fault dips and low elevations. These variations in dip may be limited to the upper crust (inset diagram). mean that gravity models cannot be used to accurately predict the fault dip or conclusively identify offsets at the base of the crust (as shown in Figure 7, our preferred model). However, these data are consistent with the moderate to low dips measured from the seismic lines and the MT model. [18] Modeled changes in crustal thickness across the fault are consistent with the view that the fault is thick skinned and extends to the base of the crust at its southern end. From west to east across the fault, the crustal thickness increases by about 11 km (27 – 38 km) in the southern profile (Figure 7). This change in crustal thickness across the fault is in agreement with deep crustal seismic reflection data south of the peninsula from which Stern and Davey [1990] infer a thick skinned fault model in southern Taranaki. By contrast gravity modeling predicts only about 3 km (22 – 25 km) change in crustal thickness across the fault on the northern profile (Figure 7). The crust west of the fault has more-orless uniform thickness between northern and southern profiles, whereas there is a marked difference in crustal thickness on the eastern side of the fault (25 – 38 km). This north-south change in crustal thickness was noted by Stern et al. [1987], and earthquake data [Sherburn and White,

2005] and seismic tomography [Reyners et al., 2006] indicate that it occurs on the eastern side of the fault at about the latitude of Pukearuhe-1 well (Figure 1). The northward decrease in thickening across the fault is consistent with a decrease in fault displacement in the same direction (see next section); however, we cannot presently discount the possibility that it was also influenced by processes unrelated to Taranaki Fault displacements. Stern et al. [2006] argue, for example, that crustal thinning in the central and western North Island was induced by the post Miocene convective removal of mantle lithosphere that was thickened in the early Miocene.

4. Fault Displacements [19] Measurement of fault displacements requires the reliable correlation of horizons across a fault [e.g., Muraoka and Kamata, 1983; Chapman and Williams, 1984; Barnett et al., 1987; Ellis and Dunlap, 1988]. Displacement on thrust faults is typically accompanied by folding, particularly close to the fault tip [e.g., Boyer, 1986; Chester and Chester, 1990; Suppe and Medwedeff, 1990]. Contraction across the Taranaki Fault in the upper 5 – 6 km of crust is

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Figure 9. Vertical separation diagram showing the elevation (with respect to sea level) of the top basement and base Oligocene horizons along the crest of the basement high in the fault hanging wall and in the footwall west of the fault together with the tip of the thrust wedge (see Figure 10 for details of measurements). The elevation of the tip of the basement wedge is from this study, while data for the top basement and base Oligocene are principally from structure contour maps of Thrasher et al. [1995] and King and Thrasher [1996]. accommodated by a combination of fault displacement and folding. For the purposes of this study we measure the vertical separation of horizons deformed by the fault. These vertical separations incorporate displacement on the principal slip surface, displacement on subsidiary faults in the zone around the Taranaki Fault and fault-related folding (Figure 9). Vertical separations for the top basement, top Cretaceous, base Oligocene, top early Miocene and top Miocene have been estimated by measuring the difference in elevation of these horizons in the footwall and hanging wall of the fault (Figure 10). Our estimates of separation assume that each horizon was approximately horizontal prior to faulting, which is consistent with paleoenvironmental interpretations of the Taranaki Basin succession [King and Thrasher, 1996]. Because of uplift, material has been eroded from the top of the basement high in the fault hanging wall. The amount of hanging wall erosion is unknown, in all cases except for the base Oligocene from the Herangi High north (210 km distance in Figure 9) and the top Miocene, and vertical separations are minimum values for the eroded horizons (for further discussion, see Figure 11 caption). The fault dips to the east with the hanging wall block displaced upward relative to the footwall block, indicating that the fault accommodated a component of reverse dip slip. As the fault does not steepen with

depth (as would be expected for a strike-slip flower structure) and bends in the fault do not appear to be associated with local pop-up or basin structures, we infer that the fault did not accommodate a significant component of strike slip. [20] The vertical separation profiles in Figure 11 confirm the large fault displacements of King and Thrasher [1996]. The vertical separation profile for top basement horizon, for example, reaches a maximum of at least 7 km. Given the observed fault dips, a vertical separation of 7 km requires that at least 10 km of shortening and 12– 15 km of dip-slip displacement accrued on the fault since the mid-Cretaceous (i.e., the age of the oldest sedimentary rocks resting on top basement horizon). [21] Vertical separations show a progressive increase in displacement with horizon age from which it can be inferred that the fault was active from prior to the Oligocene until immediately after the end of the Miocene. This interpretation may in part arise because the top Cretaceous, the southern part of the base Oligocene and the top early Miocene are not present on the hanging wall, and therefore vertical separations for these horizons are minimums and could be larger. However, the vertical separations for the base Oligocene from the Herangi High northward and all of the top Miocene are well constrained, because these horizons are present on both the hanging wall and footwall.

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Figure 10. Schematic cross section showing the general thrust geometry and measurement of horizon elevations for Figure 9. From the vertical separation profiles, the minimum preOligocene vertical separation across the fault from the Herangi High north is the difference in separation between the top basement horizons and base Oligocene horizons and ranges from 2 to 5 km (Figure 11). Along this northern section of the fault (