Magnetotelluric imaging of the Ohaaki geothermal system, New Zealand: Implications for locating basement permeability

Journal of Volcanology and Geothermal Research 268 (2013) 36–45

Contents lists available at ScienceDirect

Journal of Volcanology and Geothermal Research journal homepage: www.elsevier.com/locate/jvolgeores

Magnetotelluric imaging of the Ohaaki geothermal system, New Zealand: Implications for locating basement permeability E.A. Bertrand a,⁎, T.G. Caldwell a, G.J. Hill a, S.L. Bennie a, S. Soengkono b a b

GNS Science, 1 Fairway Drive, Avalon, Lower Hutt 5010, New Zealand GNS Science, 114 Karetoto Rd., Wairakei 3377, New Zealand

a r t i c l e

i n f o

Article history: Received 18 July 2013 Accepted 20 October 2013 Available online 29 October 2013 Keywords: Magnetotellurics Ohaaki geothermal system Taupo Volcanic Zone Permeability

a b s t r a c t Closely-spaced broadband MT (magnetotelluric) data measured on a profile through the Ohaaki geothermal system show that the basement greywacke directly beneath the geothermal field is resistive. There is no evidence in these MT data for the existence of an intrusive heat source directly beneath Ohaaki. However, these data confirm that a dipping conductive zone links the near surface low resistivity anomaly associated with the geothermal field to a deep region of low resistivity that is offset to the northwest. The resistivity of the dipping conductor is consistent with a zone of fracturing within the basement meta-sediments, and is interpreted to show the location of upwelling high-temperature fluids that supply the Ohaaki geothermal field from a deep magmatic source. © 2013 Elsevier B.V. All rights reserved.

1. Introduction The Ohaaki geothermal system in the North Island of New Zealand has for many years provided the canonical electrical resistivity model for high-temperature liquid-dominated geothermal systems (e.g. Simmons and Browne, 1990; Johnston et al., 1992). In this model a highly conductive ‘clay cap’, formed by hydrothermal alteration of the uppermost layers of young volcanics, overlies a deeper, more resistive geothermal reservoir. The higher resistivity of the deep reservoir is thought to reflect the transition in alteration products to less conductive forms at increasing temperatures; in particular, the transition from smectite to illite which occurs at ~200 °C (Ussher et al., 2000). Using this model, the resistive part of a liquid-dominated geothermal system is often identified with the convective up-flow that supplies heat and fluid to the geothermal field. Ohaaki is located near the southeastern margin of the Taupo Volcanic Zone (TVZ), a young rifted-arc (Wilson et al., 1995) that contains more than 20 liquid-dominated geothermal systems (Bibby et al., 1995). Discovery of the Ohaaki geothermal field resulted from electrical resistivity surveys made during the 1950s and 60s (summarized in Hunt, 1989). Despite having relatively little surficial geothermal expression (Blakeley and O'Sullivan, 1983), these resistivity surveys indicated a single large geothermal system, which was later confirmed by drilling. Ohaaki was then championed as a type-locality to test geophysical methods for geothermal exploration, and numerous surveys were carried out (e.g. Hochstein and Hunt, 1970; Risk et al., 1970, 1977; Risk, 1981; Henrys and van Dijck, 1987; Ingham, 1989; Risk, 1993). Combined with extensive geochemistry, geology, and reservoir modeling studies, Ohaaki has become one of the ⁎ Corresponding author. Tel.: +64 4 570 4288. E-mail address: [email protected] (E.A. Bertrand). 0377-0273/$ – see front matter © 2013 Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.jvolgeores.2013.10.010

most thoroughly investigated geothermal fields in the world; lessons learned have been used to guide subsequent developments in New Zealand, Chile, Indonesia and the Philippines (Hunt, 1989). Located near the rift shoulder (where young volcanic deposits that cover the TVZ are relatively thin), Ohaaki is one the few geothermal systems where the Mesozoic basement greywacke (sandstones and argillites) has been regularly penetrated by drilling (Clotworthy et al., 1995). Numerous intersections provide good constraints on the basement depth, which dips from the southeast (b1 km depth) to the northwest (N2 km depth). The north-westward dip is thought to reflect displacement on at least 4 major normal faults (e.g. Lee and Bacon, 2000), which have been suggested as probable conduits for the up-flow of high-temperature geothermal fluid (e.g. Clotworthy et al., 1995). In 1995, a drilling program was undertaken to explore the deep resource potential and included a deviated well designed to intersect the inferred faults within the basement greywacke. Although high-temperatures were encountered, no significant permeability was found within the basement (Brockbank and Bixley, 2011). The lack of permeability challenged the concept that faults directly beneath the near-surface low resistivity anomaly provided the permeability pathway for up-flowing high-temperature fluids. The low permeability encountered was also inconsistent with the conceptual model of heat transport that links low DC apparent resistivity anomalies associated with the TVZ's geothermal systems (see Fig. 1) to the top of a vertically ascending plume of convecting hydrothermal fluid (Bibby et al., 1995). The lack of vertical displacement observed in the overlying 330 ka Rangitaiki (or Whakamaru group) ignimbrite also raised uncertainty on the expected link between upward fluid flow and permeable fault-zones beneath Ohaaki (Wood et al., 2001; Rissman et al., 2011). Early analysis of the reservoir chemistry from Ohaaki suggested a single deep up-flow zone (Mahon and Finlayson, 1972; Glover and

E.A. Bertrand et al. / Journal of Volcanology and Geothermal Research 268 (2013) 36–45

37

a)

b)

M L K J I H G F E D C B A

Fig. 1. a) Map showing MT measurements through the Ohaaki (OH) geothermal system using the NZMG (New Zealand map grid) coordinate system (labels in km). The hashed border shows the resistivity boundary of the Ohaaki geothermal field from dc-resistivity measurements (Risk et al., 1977). The rose diagram shows the geoelectric strike direction (α–β) from phase tensor analysis. Red circles are MT sites collected in 2008, and the yellow circles show select MT sites from Bertrand et al. (2012a). A, B and C show the locations of MT data included in Fig. 6. b) Black circles show the complete array MT dataset of Bertrand et al. (2012a). The background digital elevation model is overlain by interpolated (using a 1 km grid interval) apparent resistivity values (Bibby et al., 1995). Areas of low apparent resistivity mark the surface extent of geothermal systems: Rotokawa (RT), Ngatamariki (NM), Orakei–Korako (OK), Ohaaki (OH), Te Kopia (TK), Reporoa (RP), Waiotapu–Waikite (WW), Waimungu (WM), Atiamuri (AM), Mokai (MK), Wairakei (WK) and Tauhara (TH). The pink dashed line shows the boundary of the TVZ (Wilson et al., 1995) with the Kaingaroa Plateau (KP) to the east. Active faults (including the Paeroa Fault — PF) are shown as orange lines, the arrow labeled WHS marks the Whangairorohea hot springs, and the Waikato River is outlined in blue. The black box in b) marks the map borders of a).

38

E.A. Bertrand et al. / Journal of Volcanology and Geothermal Research 268 (2013) 36–45

Ohaaki

SE

Fig. 2. Phase tensor ellipses filled with color indicating the skew-angle (β). The phase tensor ellipses are plotted in map view, with vertical and horizontal axes corresponding to north and east, respectively. Blank areas indicate where poor quality data were removed prior to inversion modeling.

Hedenquist, 1989). However, compositional differences between the fluids observed in wells located west of the Waikato River (Fig. 1) and in those located to the east of the river (Giggenbach, 1989) suggested that two regions of deep up-flow with distinct chemistries might exist. In particular, the 3He/4He ratios and chloride concentrations were highest on the west bank, while the B/Cl, N2/Ar and CO2/He maxima were recorded on the east bank. Hedenquist (1990) suggested that boiling and dilution processes could explain these differences, while Lovelock (1990) and later Christenson et al. (2000, 2002) invoked a model with two distinct up-flows. Recently, using a regional MT array in the southeastern TVZ, Bertrand et al. (2012a) imaged low-resistivity zones that link the near-surface expression of the geothermal fields at Rotokawa Ngatamariki and Ohaaki (see Fig. 1) to an inferred magmatic heat source located below the brittle–ductile transition at ~ 7–8 km depths (Bibby et al., 1995; Bryan et al., 1999). At Rotokawa and Ngatamariki, these lowresistivity zones are vertical and lie directly beneath the geothermal fields. However, at Ohaaki, the low-resistivity zone dips to the northwest and is offset from the geothermal field. Therefore, rather than being a type-example for the resistivity signature of a liquiddominated geothermal system, the deep resistivity structure at Ohaaki may be atypical. In this paper, we combine some of the regional-scale MT array data described in Bertrand et al. (2012a) with a profile of closely-spaced MT measurements through Ohaaki to better identify the source of the up-flowing geothermal fluid. We also resolve more clearly the dipping conductive zone that is offset northwest of the geothermal field, as delineated by DC resistivity data (Risk et al., 1977). 2. Data collection and analysis In 2008, broadband MT data (0.01–1000 s) were recorded using Phoenix MTU instruments on a northwest–southeast profile that transects the Ohaaki geothermal system (red dots in Fig. 1a). MT data were recorded for 2 night's duration at locations spaced ~500 m apart through the center of the geothermal field. These MT measurements extend ~12 km southeast from Ohaaki onto the Kaingaroa Plateau, and ~3 km

to the northwest. In 2009 and 2010, an additional 220 broadband MT measurements were made on a 2 km grid in a 25 × 35 km area in the southeast TVZ (Fig. 1b; Bertrand et al., 2012a). These MT data were also recorded using Phoenix MTU instruments for 2 night's duration, and were made to investigate links between shallow hydrothermal systems and their inferred deep magmatic heat source. 2-D and 3-D modeling of subsets of the array MT data that transect the Ohaaki geothermal field showed a dipping low-resistivity zone that is offset to the northwest and extends from the near-surface down to a low-resistivity layer at ~7–8 km depths (Bertrand et al., 2012a). This low-resistivity zone was inferred to represent a convection plume of hot-fluid within fracturepermeability. However, detailed structures were not well resolved in these regional-scale models as measurement sites were ~2 km apart. Here we combine the closely spaced MT data collected in 2008, with some of the array MT data measured northwest of the geothermal field. By integrating these local and regional datasets, we obtain a wide aperture (~30 km) required to image the deep (~10 km) resistivity structure of the geothermal system (Bertrand et al., in press), and maintain a close station spacing through the center of the geothermal field, needed to resolve the shallow structure. All MT time-series data were processed using the remote reference technique (Gamble et al., 1979) to minimize the effects of uncorrelated signals measured simultaneously at the local and remote sites. In general, high-quality MT data were recorded at most sites occupied; a success that can in part be attributed to acquiring data for ~40 h at each measurement site. As a final processing step, obvious outliers in the sounding curves were removed prior to data analysis and inversion. To assess the dimensionality of these MT data we use properties of the magnetotelluric phase tensor (Caldwell et al., 2004), which makes no a priori assumptions regarding the regional resistivity structure that we seek to determine. Booker (2013) suggests that the phase tensor skew-angle (β) is a robust measure to identify the existence of 3-D regional structure in MT data, and argues that skew-angles should be less than 3° to justify a 2-D interpretation. Fig. 2 shows that most phase tensor skew-angles at periods less than 300 s are within ±3°. A rose diagram of the orientation of the maximum phase tensor principal

Fig. 3. Induction vectors plotted using the Parkinson (1962) convention at different periods (T) superimposed on a map of high-voltage power-lines (red) that intersect the MT survey area. Missing arrows indicate where poor quality data were removed.

E.A. Bertrand et al. / Journal of Volcanology and Geothermal Research 268 (2013) 36–45

39

40

E.A. Bertrand et al. / Journal of Volcanology and Geothermal Research 268 (2013) 36–45

Ohaaki

SE

Depth (km)

a)

Depth (km)

b)

Depth (km)

c)

Fig. 4. a) 2-D resistivity inversion model of the TE and TM mode MT data shown in Fig. 1a. b) 3-D resistivity inversion model of the MT impedance tensor data for the same single-profile MT sites included in the 2-D model. c) 3-D resistivity inversion model of the impedance tensor data for 5 profiles (I–M) from the array MT dataset of Bertrand et al. (2012a) (Fig. 1b). Solid gray lines show well tracks projected onto the models, and the dashed rectangle outlines the area of these models shown enlarged in Fig. 7. Whangairorohea hot springs (WHS).

axes (α–β values) for all periods and sites (inset in Fig. 1a) shows that the geoelectric strike direction is well defined and is collinear with the main northeast–southwest structural features in the TVZ (i.e. ~N45°E). This strike direction is also consistent with previous regional MT studies in this area (e.g. Heise et al., 2007; Bertrand et al., 2012a). At periods greater than ~10 s, Fig. 2 shows a consistent trend of increasing β values. This trend indicates that the MT data at longperiods are sensing the presence of 3-D electrical structure, consistent with plots of the induction vector data (Fig. 3), which show an increasing northeast component at periods greater than ~100 s. For a 2-D approach to be valid, the real components of the induction vectors should point orthogonal to the geoelectric strike direction (e.g. Simpson and Bahr, 2005). Thus, while the signature of this MT dataset is 2-D at short periods, given the indications of 3-D behavior at longer periods, both 2-D and 3-D inversion modeling of this profile data has been carried out to ensure that significant model features are robust.

3. Modeling and interpretation The 2-D inversion algorithm of Rodi and Mackie (2001) (contained within the WinGLink software package) was used to generate smooth resistivity models from the measured impedance data between periods of 0.01–300 s. After generating models using a wide range of control parameters, a final model was chosen that balanced the requirements for a spatially smooth model with an acceptably small data misfit. The 2-D model shown in Fig. 4a achieved a normalized r.m.s. (root mean square) misfit of 1.5 after 200 iterations. This model included the profile topography and was set to solve for static shifts at late stage iterations. Error floors of 15% for the apparent resistivity and 7.5% for the phase were used. The 3-D inversion algorithm (WSINV3DMT) of Siripunvaraporn et al. (2005) was used to generate an inversion model of the full impedance tensor data (i.e. both diagonal and off-diagonal tensor components).

E.A. Bertrand et al. / Journal of Volcanology and Geothermal Research 268 (2013) 36–45

This model included the same off-diagonal impedance data used to generate the 2-D inversion model, but down-sampled to 4 periods per decade. Following Tietze and Ritter (2012), who explored factors that influence 3-D inversion of a dominantly 2-D regional structure, the model grid and data were rotated 45° to parallel the geoelectric strike direction. Error floors of 20% for the diagonal and 10% for the off-diagonal impedance tensor components were used. The larger error floor for the diagonal components avoids over-fitting the small amplitude impedances in the strike coordinate system, which have a low signal-to-noise ratio. The model mesh comprised 200 m cell widths between stations (the same as used for the 2-D inversion modeling) and coarsened outside of the area of interest. The final 3-D model (Fig. 4b) yielded a normalized r.m.s. misfit of 1.1 after 8 iterations. Comparison of Fig. 4a and b shows strong similarity between resistivity structures imaged in the 2-D and 3-D models within the upper ~3 km. The main differences between these models are variations in the depth extent and resistivity of equivalent features; an observation made previously in a comparison of the same inversion codes (Bertrand et al., 2012b). These differences are explained by the fact that MT data resolve conductance and gradients rather than resistivity (e.g. Simpson and Bahr, 2005), and by the additional degrees of freedom available to the 3-D inversion algorithm to change along-strike model resistivity. In essence, Fig. 4b shows the result of inverting a 2-D profile of MT data with a 3-D inversion algorithm and model space. With this configuration, the 3-D algorithm is poorly constrained along-strike, and is free to change the model resistivities away from the cross-section shown, in order to fit the observed data. However, despite the differences in the resistivity values below an ~3 km depth in Fig. 4a and b, the key observation is that the resistivity gradients are similar; it is the contrasts that are smaller in Fig. 4b. To further examine the deep resistivity structure northwest of Ohaaki, an additional 3-D inversion model (Fig. 4c) was generated using 5 profiles of data (I–M in Fig. 1b) that encompass the Ohaaki

41

geothermal system (see Bertrand et al., 2012a for a description of the array MT dataset). This inversion model used the same 3-D control parameters as described above, but owing to the larger station spacing of these array data (~2 km), the model mesh had a minimum cell width of 500 m. After 8 iterations, the final 3-D model produced an r.m.s. misfit of 1.1. The resistivity model in Fig. 4c necessarily lacks detailed resolution at shallow depths (compared with Fig. 4a and b), owing to the wider station spacing and cell widths used. However, Fig. 4c clearly shows the deep conductor northwest of the Ohaaki geothermal system, and model resistivity values agree very well with Fig. 4a. The more pronounced deep conductor in Fig. 4c can be attributed to the greater along-strike constraints from to the 3-D configuration of the array MT data used. To assess the reliability of these inversion results it is important to identify features in the model and measured data that correspond. Fig. 5 shows pseudosections of the TE (transverse electric) and TM (transverse magnetic) mode impedance phases, and phase tensor ellipses filled with color indicating the geometric mean of the maximum and minimum phase tensor axes (Φ2); the latter being unaffected by galvanic distortion. All of the pseudosections show phase values N 45° (indicating that resistivity is decreasing with depth) at short periods (i.e. shallow depth) that dip to the northwest beneath the Ohaaki geothermal field. These high phase values are underlain by lower phases, indicating the presence of increasing resistivity beneath the center and to the east of the geothermal field, in good agreement with the overall resistivity structure of the inversion models. To assess the data fit of these inversion models, Fig. 6 shows the measured data and model responses of the TE and TM mode apparent resistivity and phase at 3 measurement locations (sites A, B and C indicated in Fig. 1a). These sites were selected as they are included in each of the inversion models shown in Fig. 4 and sample different regions of the study area (e.g. site A northwest of Ohaaki, site B above the deep

Ohaaki

SE

a)

b)

Fig. 5. a) Pseudosections of the TE mode impedance phase, the TM mode impedance phase, and b) phase tensor ellipses filled with color indicating the invariant parameter Φ2. These phase tensor ellipses are plotted in map view, with the vertical and horizontal axes corresponding to north and east, respectively. Blank areas indicate where poor quality data were removed prior to inversion modeling.

42

E.A. Bertrand et al. / Journal of Volcanology and Geothermal Research 268 (2013) 36–45

a)

b)

c)

Fig. 6. Columns show the TE (red triangles) and TM (blue circles) mode apparent resistivity and phase data for the three sites labeled A, B and C in Fig. 1a. Rows show the inversion model fits to these data (TE — solid gray lines, TM — solid black lines) for the 2-D, 3-D profile and 3-D inversions shown in Fig. 4. The phase panels also include the forward modeled TE (red dashed lines) and TM (blue dashed lines) responses after replacing the modeled resistivity values with a 500 Ω m layer between 3 and 7 km depths.

conductor, and site C southeast of Ohaaki). Fig. 6 shows that at all periods inverted, the measured data are well fit by all of the inversion models. As observed in previous MT studies of the TVZ (e.g. Ogawa et al., 1999; Heise et al., 2007, 2010), young, resistive (~100–300 Ω m) volcanic material near the surface overlie areas of low resistivity (~10–30 Ω m) away from geothermal fields (e.g. C1 in Fig. 4). The low resistivity is caused by diagenetic alteration that forms conductive clays and zeolites (Stanley et al., 1990) within old (i.e. N~700 ka) volcaniclastic layers (Bibby et al., 1995). At the Paeroa Fault (Fig. 1a), a low-resistivity zone (C2 in Fig. 4) is imaged that dips to the southeast, in agreement with interpretations of DC resistivity measurements by Risk and Bibby (1995). At depths below the brittle–ductile transition (corresponding with the depth extent of the seismogenic zone; Bibby et al., 1995) widespread areas of lowresistivity are consistent with a zone of ~4% partial melt (Heise et al., 2007). Beneath the Ohaaki geothermal field and the Whangairorohea hot springs (WHS in Fig. 4), model resistivities of ~3–10 Ω m occur in the upper ~2 km. Low resistivities observed in the shallow parts of

hydrothermal systems can be explained by combinations of the following: 1. Geothermal fluids can have a high-concentration of dissolved salts (e.g. Barnes, 1979), yielding a conductive electrolyte, 2. The resistivity of water is inversely related to temperature up to ~300 °C (Ucock et al., 1980), and 3. Hydrothermal alteration products at temperatures b~200 °C (e.g. smectite) are conductive (e.g. Ussher et al., 2000). Beneath the Ohaaki geothermal field, up-doming of the shallow lowresistivity layer is observed, which is routinely interpreted to indicate the existence of an underlying, high-temperature reservoir in geothermal exploration studies. Although the depth extent of a conductor in regularized MT inversion models will be smeared downwards (e.g. Bertrand et al., 2009) the transition from low-to-higher resistivity beneath Ohaaki occurs within the overlying volcanic material above the greywacke basement (Fig. 7). As previously observed in Bertrand et al. (2012a), a dipping, lowresistivity feature (C3 in Fig. 4) exists northwest of Ohaaki that extends from the near-surface to a deeper low-resistivity zone. This C3 conductor is important for understanding the deep source of heat and fluids that supply the Ohaaki geothermal field. We have therefore tested the

E.A. Bertrand et al. / Journal of Volcanology and Geothermal Research 268 (2013) 36–45

SE

Depth (km)

a)

Depth (km)

b)

Fig. 7. Enlargements of the dashed regions outlined in Fig. 4 of a) the 2-D resistivity inversion model and b) the 3-D profile resistivity inversion model. Solid gray lines show well tracks projected onto the MT models and white dots show the intersection with the greywacke basement. Note that Fig. 4c is not shown enlarged owing to the decreased near-surface resolution from the wider-spaced array data used (Fig. 1b).

data sensitivity to this feature by replacing the model resistivity structure between 3 and 7 km depths with a 500 Ω m layer, comparable to that observed in the southeastern parts of all the models, and representative of basement metasediments. 2-D and 3-D MT data were then forward modeled and the subsequent phase data, which are unaffected by static shifts, are shown using dashed lines at the 3 sites included in Fig. 6. It is clear in Fig. 6 that this sensitivity test caused little change to the phase data at sites A and C, which are located away from the C3 conductor. In contrast, a strong effect is observed at site B located above the C3 conductor; particularly in the TE mode phase at periods in the range 3–30 s that are sensitive to depths of ~ 3–7 km in these models. This sensitivity test shows that the C3 conductor is required by the measured data and is a robust model feature. 4. Discussion Low resistivity observed at depth to the northwest of the Ohaaki geothermal field exists within what is likely to be greywacke, or higher metamorphic grade equivalents. Beneath Ohaaki and to the southeast, these rocks have resistivities N ~300 Ω m (Fig. 4). Therefore, the

43

decrease in basement resistivity to values of 30–100 Ω m northwest of Ohaaki requires the existence of a conductive phase. The convective model of heat transport described by Bibby et al. (1995) implies that this conductive phase is a high-temperature, saline fluid contained in fractures. At Ohaaki, the molarity of the deep geothermal fluid is ~0.035 M (Barnes, 1979). Using experimental data in Quist and Marshall (1968), and assuming hydrostatic pressure, yields an estimate of the resistivity of the Ohaaki deep aquifer fluid to be ~1 Ω m, at a depth of 3 km and a temperature of ~300 °C. Assuming that within the basement rocks, hot upwelling fluids are contained in fractures, it is most appropriate to estimate porosity using the Hashin–Shtrikman upper-bound (Hashin and Shtrikman, 1962), which assumes complete connectivity. This bound yields 1–5% porosity for resistivities of 30–100 Ω m. An alternative estimate, which assumes less connectivity can be obtained from Archie's Law (Archie, 1942), with a cementation exponent of 1.5. This method gives 4–10% porosity. Since laboratory analysis of greywacke samples from Ohaaki (in the resistive part of the MT model) and at the Kawerau geothermal system indicate near-zero primary porosity (Brathwaite et al., 2002), the low-resistivities observed in the basement northwest of Ohaaki likely result from saline fluid interconnected by fractures, and/or in secondary porosity generated by geothermal fluid transport. It remains to discuss what causes the offset of the near-surface Ohaaki geothermal field, with respect to what appears to be its heat source in the northwest. On the basis of thickening to the northwest, Henrys and Hochstein (1990) suggested that the Ohaaki rhyolite (located beneath the west bank) was derived from an eruptive center located to the west or northwest of the geothermal field. Although long considered to not be located within or on the faulted margins of one of the many large rhyolite calderas that characterize the central TVZ (e.g. Wood, 1995), Soengkono (2012) showed that Ohaaki is in fact located on the margin of a low-gravity anomaly, referred to as the Mihi volcanic depression (Fig. 8). Fractures associated with the margin of this depression may provide a pathway for the lateral migration of hightemperature upwelling fluids, rising from near the brittle–ductile transition northwest of Ohaaki. In support of this interpretation, recent seismic tomography results indicate that a strong decrease in both P-wave and S-wave velocity occurs at 4 km depth northwest of Ohaaki (Bannister et al., in press). We also note that the dip of the low-resistivity zone in the MT models is ~ 50° (Fig. 4), near to the value of 60° predicted by Anderson (1905) for normal faulting. Thus, the dipping C3 conductor is plausibly interpreted to be caused by fluids in fractures related to an active-rift or caldera-collapse fault. At Ohaaki, no significant permeability has been found within the resistive greywacke that underlies the geothermal field (Brockbank and Bixley, 2011). However, a dipping low-resistivity zone in the basement exists to the northwest that is consistent with saline fluids contained in fractures. Since fracture permeability is required to convectively transport heat through the basement in the TVZ, our results suggest that low-resistivity zones (in the basement) indicate permeable regions that may be targets for deep geothermal exploration.

5. Conclusions 2-D and 3-D resistivity inversion models of MT data that transect the Ohaaki geothermal system have shown that a transition from lowresistivity to high-resistivity occurs above the greywacke interface, directly beneath the geothermal field. Further to the northwest, a dipping low-resistivity zone is imaged within the basement that is interpreted to represent the source of deep, high-temperature fluids within fractures. This paper shows that within the basement rocks of the TVZ, conductive rather than resistive zones in the MT models may identify the convective up-flows that supply heat and fluids to the geothermal fields.

44

E.A. Bertrand et al. / Journal of Volcanology and Geothermal Research 268 (2013) 36–45

Fig. 8. Map showing Bouger gravity values in the survey area. Blue colors show the MVD (Mihi volcanic depression) after Soengkono (2012). All other labels as in Fig. 1.

Acknowledgments Cooperation from Contact Energy and the landowners in the survey area is greatly appreciated. We thank Weerachai Siripunvaraporn for the use of his 3-D inversion program, WSINV3DMT. We thank two anonymous reviewers for their constructive comments on this manuscript. This project was supported by public research funding from the government of New Zealand. References Anderson, E.M., 1905. The Dynamics of Faulting. Oliver and Boyd, London, England (206 pp.). Archie, G.E., 1942. The electrical resistivity log as an aid in determining some reservoir characteristics. Trans. Am. Inst. Min. Metall. Pet. Eng. 146, 54–62. Bannister, S., Bourguignon, S., Sherburn, S., Bertrand, E.A., 2013. Seismic imaging of the central Taupo Volcanic Zone using double-difference tomography. Proceedings of the 35th New Zealand Geothermal Workshop, Rotorua, New Zealand (in press). Barnes, H.L., 1979. Geochemistry of Hydrothermal Ore Deposits, 2nd edition. John Wiley and Sons, New York, United States (798 pp.). Bertrand, E.A., Unsworth, M.J., Chiang, C.W., Chen, C.S., Chen, C.C., Wu, F.T., Turkoglu, E., Hsu, H.L., Hill, G.J., 2009. Magnetotelluric evidence for thick-skinned tectonics in central Taiwan. Geology 37. http://dx.doi.org/10.1130/G25755A.1. Bertrand, E.A., Unsworth, M.J., Chiang, C.W., Chen, C.S., Chen, C.C., Wu, F.T., Turkoglu, E., Hsu, H.L., Hill, G.J., 2012a. Magnetotelluric imaging beneath the Taiwan orogen: an arccontinent collision. J. Geophys. Res. 117. http://dx.doi.org/10.1029/2011JB008688. Bertrand, E.A., Caldwell, T.G., Hill, G.J., Wallin, E.L., Bennie, S.L., Cozens, N., Onacha, S.A., Ryan, G.A., Walter, C., Zaino, A., Wameyo, P., 2012b. Magnetotelluric imaging of upper-crustal convection plumes beneath the Taupo Volcanic Zone, New Zealand. Geophys. Res. Lett. 39. http://dx.doi.org/10.1029/2011GL050177. Bertrand, E.A., Caldwell, T.G., Sepulveda, F., 2013. The importance of survey aperture for imaging high-temperature geothermal systems with magnetotellurics. Proceedings of the 35th New Zealand Geothermal Workshop, Rotorua, New Zealand (in press). 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 circulation. J. Volcanol. Geotherm. Res. 68, 29–58. Blakeley, M.R., O'Sullivan, H.J., 1983. A simple model of the Ohaaki geothermal reservoir. Proceedings of the 5th New Zealand Geothermal Workshop, Auckland, New Zealand. Booker, J.R., 2013. The magnetotelluric phase tensor: a critical review. Surv. Geophys. http://dx.doi.org/10.1007/s10712-013-9234-2.

Brathwaite, R.L., Wood, C.P., Rosenberg, M.D., Faure, K., 2002. Porosity and permeability in the basement rocks at the Kawerau and Ohaaki geothermal fields, New Zealand. Proceedings of the 24th New Zealand Geothermal Workshop, Auckland, New Zealand. Brockbank, K.M., Bixley, P.F., 2011. The Ohaaki deep reservoir. Proceedings of the 33rd New Zealand Geothermal Workshop, Auckland, New Zealand. 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. Geophys. 42, 533–542. Caldwell, T.G., Bibby, H.M., Brown, C., 2004. The magnetotelluric phase tensor. Geophys. J. Int. 158, 457–469. Christenson, B.W., Mroczek, E.K., Stewart, M.K., Lyon, G., Kennedy, B.M., 2000. Ohaaki reservoir chemistry: insights into the nature and location of the heat sources. Proceedings World Geothermal Congress 2000, Kyushu-Tohoku, Japan. Christenson, B.W., Mroczek, E.K., Kennedy, B.M., van Soest, M.C., Stewart, M.K., Lyon, G., 2002. Ohaaki reservoir chemistry: characteristics of an arc-type hydrothermal system in the Taupo Volcanic Zone, New Zealand. J. Volcanol. Geotherm. Res. 115, 53–82. Clotworthy, A., Lovelock, B., Care, B., 1995. Operational history of the Ohaaki geothermal field, New Zealand. Proceedings World Geothermal Congress 1995, Florence, Italy. Gamble, T.D., Goubeau, W.M., Clarke, J., 1979. Magnetotellurics with a remote reference. Geophysics 44, 53–68. Giggenbach, W.F., 1989. The chemical and isotopic position of the Ohaaki field within the Taupo Volcanic Zone. Proceedings of the 11th New Zealand Geothermal Workshop, Auckland, New Zealand. Glover, R.B., Hedenquist, J.W., 1989. A brief history of chemical exploration at OhaakiBroadlands. Proceedings of the 11th New Zealand Geothermal Workshop, Auckland, New Zealand. Hashin, Z., Shtrikman, S., 1962. Variational approach to the theory of effective magnetic permeability of multiphase materials. J. Appl. Phys. 33, 3125–3131. Hedenquist, J.W., 1990. The thermal and geochemical structure of the Broadlands–Ohaaki geothermal system, New Zealand. Geothermics 9, 151–185. Heise, W., Bibby, H.M., Caldwell, T.G., Bannister, S.C., Ogawa, Y., Takakura, S., Uchida, T., 2007. Melt distribution beneath a young continental rift: the Taupo volcanic zone, New Zealand. Geophys. Res. Lett. 34. http://dx.doi.org/10.1029/2007GL029629. Heise, W., Caldwell, T.G., Bibby, H.M., Bennie, S.L., 2010. Three-dimensional electrical resistivity image of magma beneath an active continental rift, Taupo Volcanic Zone, New Zealand. Geophys. Res. Lett. 37. http://dx.doi.org/10.1029/2010GL043110. Henrys, S.A., Hochstein, M.P., 1990. Geophysical structure of the Broadlands–Ohaaki geothermal field New Zealand. Geothermics 19, 129–150. Henrys, S.A., van Dijck, M.F., 1987. Structure of concealed rhyolites and dacites in the Broadlands–Ohaaki geothermal field. Proceedings of the 9th New Zealand Geothermal Workshop, Auckland, New Zealand. Hochstein, M.P., Hunt, T.M., 1970. Seismic, gravity and magnetic studies, Broadlands geothermal field, New Zealand. Geothermics 2, 333–346 (Special Issue).

E.A. Bertrand et al. / Journal of Volcanology and Geothermal Research 268 (2013) 36–45 Hunt, T.M., 1989. Geophysical exploration of the Broadlands (Ohaaki) geothermal field: review. Proceedings of the 11th New Zealand Geothermal Workshop, Auckland, New Zealand. Ingham, M., 1989. Audiomagnetotelluric/magnetotelluric soundings of the Broadlands– Ohaaki geothermal field. Proceedings of the 11th New Zealand Geothermal Workshop, Auckland, New Zealand. Johnston, J.M., Pellerin, L., Hohmann, G.W., 1992. Evaluation of electromagnetic methods for geothermal reservoir detection. Trans. Geotherm. Resour. Counc. 16, 241–245. Lee, S., Bacon, L., 2000. Operational history of the Ohaaki geothermal field, New Zealand. Proceedings World Geothermal Congress 2000, Kyushu–Tohoku, Japan. Lovelock, B.G., 1990. The Ohaaki Geothermal System: Fluid Chemistry, Hydrology and Response to Production in 1989–1990. (M.Sc. Thesis) University of Auckland, New Zealand. Mahon, W.A.J., Finlayson, J.B., 1972. The chemistry of the Broadlands geothermal area, New Zealand. Am. J. Sci. 272, 48–68. Ogawa, Y., Bibby, H.M., Caldwell, T.G., Takakura, S., Uchida, T., Matsushima, N., Bennie, S.L., Tosha, T., Nishi, Y., 1999. Wide-band magnetotelluric measurements across the Taupo Volcanic Zone, New Zealand — preliminary results. Geophys. Res. Lett. 26, 3673–3676. Parkinson, W.D., 1962. The influence of continents and oceans on geomagnetic variations. Geophys. J. R. Astron. Soc. 6, 411–449. Quist, A.S., Marshall, W.L., 1968. Electrical conductances of aqueous sodium chloride solutions from 0 to 800 °C and at pressures to 4000 bars. J. Phys. Chem. http://dx.doi.org/ 10.1021/j100848a050. Risk, G.F., 1981. Induced polarization (IP) in the Broadlands geothermal field. Proceedings of the 3rd New Zealand Geothermal Workshop, Auckland, New Zealand. Risk, G.F., 1993. Resurvey of the resistivity boundary of Ohaaki geothermal field, 1975–1992. Proceedings of the 15th New Zealand Geothermal Workshop, Auckland, New Zealand. Risk, G.F., Bibby, H.M., 1995. Resistivity lows near Paeroa fault (TVZ, NZ) caused by topographic effects. Proceedings of the 17th New Zealand Geothermal Workshop, Auckland, New Zealand. Risk, G.F., Macdonald, W.J.P., Dawson, G.B., 1970. DC resistivity surveys of the Broadlands geothermal region, New Zealand. Geophys. J. R. Astron. Soc. 2, 287–294 (Special Issue). Risk, G.F., Groth, M.J., Rayner, H.H., Dawson, G.B., Bibby, H.M., Macdonald, W.J.P., Hewson, C.A.Y., 1977. The resistivity boundary of the Broadlands geothermal field. Geophysics

45

Division Technical Report No 123. Department of Scientific and Industrial Research, Wellington, New Zealand (42 pp.). Rissman, C., Nicol, A., Cole, J., Fairley, J., Christenson, B., Leybourne, M., Milicich, S., Ring, U., Gravely, D., 2011. J. Volcanol. Geotherm. Res. 204, 12–26. Rodi, W., Mackie, R.L., 2001. Nonlinear conjugate gradients algorithm for 2-D magnetotelluric inversion. Geophysics 66, 174–187. Simmons, S.F., Browne, P.R.L., 1990. A three dimensional model of the distribution of hydrothermal alteration minerals within the Broadlands–Ohaaki geothermal field. Proceedings of the 12th New Zealand Geothermal Workshop, Auckland, New Zealand. Simpson, F., Bahr, K., 2005. Practical Magnetotellurics. Cambridge University Press, Cambridge, England (254 pp.). Siripunvaraporn, W., Egbert, G., Lenbury, Y., Uyeshima, M., 2005. Three-dimensional magnetotelluric inversion: data-space method. Phys. Earth Planet. Inter. 150, 3–14. Soengkono, S., 2012. Gravity modeling of Reporoa basin, eastern Taupo Volcanic Zone TVZ. New Zealand. Proceedings of the 34th New Zealand Geothermal Workshop, Auckland, New Zealand. Stanley, W.D., Mooney, W.D., Fuis, G.S., 1990. Deep crustal structure of the Cascade range and surrounding regions from seismic refraction and magnetotelluric data. J. Geophys. Res. 95, 19419–19438. Tietze, K., Ritter, O., 2012. 3D magnetotelluric inversion in practice, Extended Abstract. 21st Electromagnetic Induction Workshop, Darwin Australia. Ucock, H., Ershaghi, I., Olhoeft, G., 1980. Electrical resistivity of geothermal brines. J. Pet. Technol. 32, 717–727. Ussher, G., Harvey, C., Johnstone, R., Anderson, E., 2000. Understanding the resistivities observed in geothermal systems. Proceedings World Geothermal Congress 2000, Kyushu–Tohoku, Japan. 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 Taupo Volcanic Zone, New Zealand: a review. J. Volcanol. Geotherm. Res. 68, 1–28. Wood, C.P., 1995. Calderas and geothermal systems in the Taupo Volcanic Zone, New Zealand. Proceedings World Geothermal Congress 1995, Florence, Italy. Wood, C.P., Brathwaite, R.L., Rosenberg, M.D., 2001. Basement structure, lithology and permeability at Kawerau and Ohaaki geothermal fields, New Zealand. Geothermics 30, 461–481.