TY - JOUR
AU1 - Antriasian,, Anson
AU2 - Harris, Robert, N
AU3 - Tréhu, Anne, M
AU4 - Henrys, Stuart, A
AU5 - Phrampus, Benjamin, J
AU6 - Lauer,, Rachel
AU7 - Gorman, Andrew, R
AU8 - Pecher, Ingo, A
AU9 - Barker,, Dan
AB - Summary We present 96 new seafloor heat flow determinations, made with a 3.5-m violin-bow probe and collocated with seismic reflection profiles, from the northern Hikurangi margin on the east coast of New Zealand's North Island. Here the Hikurangi Plateau on the Pacific Plate is subducting under the Australian Plate. The background heat flow is 58 ± 8 mW m−2, consistent with and within the variability of globally observed heat flow (56 ± 15 mW m−2) for oceanic crust 90–120 Ma, the age of the Hikurangi Plateau. Seaward of the deformation front, we find evidence for advective fluid flow associated with basement relief. Landward of the deformation front, we use a 2-D steady-state finite-element model to quantify the thermal regime. Despite corrections for the effects of bottom water temperature change, bathymetry and sedimentation, there is considerable scatter in the heat flow data including a local and sharp increase in heat flow of up to 35 mW m−2 observed over the outermost wedge. Variability in the heat flow data is likely due to complex, unmodelled 3-D fluid flow. We augment our heat flow measurements with estimates from a bottom-simulating reflection and continental bottom hole temperatures and conclude that the effective coefficient of friction, μ*, is approximately 0.06 in the region of observed slow-slip events and increases to 0.18 approximately 50 km landward of the deformation front. This transition in μ* may be marking the downdip edge of overpressures along the subduction thrust, suggesting that slow slip is enabled by overpressure. New Zealand, Heat generation and transport, Large igneous provinces, Rheology and friction of fault zones, Subduction zone processes 1 INTRODUCTION The northern portion of the Hikurangi forearc, east of New Zealand's North Island (Fig. 1), has become a key area to understand shallow slow-slip events (SSEs). At this margin, shallow SSEs occur at depths of <2–15 km every 18–24 months, last 1–2 weeks, have a cumulative slip of several to tens of centimetres and generate equivalent moment magnitudes of 6.0–7.0 (Wallace et al. 2004, 2012, 2016). Globally, shallow SSEs are thought to lie in the transition between velocity weakening and strengthening behaviour and occur in environments characterized by low effective stress and fluid pressures that exceed hydrostatic (Saffer & Wallace 2015). Figure 1. View largeDownload slide Tectonic setting of the North Island, New Zealand and location of heat flow data. (a) North Island, New Zealand showing Hikurangi Trough and Channel. The magenta areas show the location of previous slow-slip earthquakes. Blue circles show location of heat flow measurements on land. White arrow shows convergence direction and rate (DeMets et al. 1994). Figure modified from Barnes et al. (2010). (b) Seismic lines offshore Gisborne. Red circles show location of heat flow measurements. Yellow triangles show location of bottom water temperature sensors deployed during HOBITTS campaign (Wallace et al. 2016). (c) Seismic reflection line 05CM-04. CDP is common depth point. Figure 1. View largeDownload slide Tectonic setting of the North Island, New Zealand and location of heat flow data. (a) North Island, New Zealand showing Hikurangi Trough and Channel. The magenta areas show the location of previous slow-slip earthquakes. Blue circles show location of heat flow measurements on land. White arrow shows convergence direction and rate (DeMets et al. 1994). Figure modified from Barnes et al. (2010). (b) Seismic lines offshore Gisborne. Red circles show location of heat flow measurements. Yellow triangles show location of bottom water temperature sensors deployed during HOBITTS campaign (Wallace et al. 2016). (c) Seismic reflection line 05CM-04. CDP is common depth point. One important aspect to understanding the mechanics of fault-slip behaviour is knowledge of the thermal regime and hydrogeology of the plate boundary and adjacent plates (Wang et al. 1995a; Ranero et al. 2008; Saffer & Tobin 2011; Saffer & Wallace 2015). Although shallow SSEs do not appear to be directly tied to a particular temperature window (e.g. Saffer & Wallace 2015), knowledge of the thermal regime provides important information on the background heat flow, the location of dehydration reactions, areas of advective fluid flow and, in some cases, the magnitude of frictional heating, which can be tied to estimates of the effective coefficient of friction, μ*, along the plate boundary fault (e.g. Wang et al. 1995a; Gao & Wang 2014). Heat flow data along the Hikurangi margin are relatively scarce and, prior to our survey, limited to areas landward of the deformation front. Prior heat flow determinations include values inferred from bottom simulating reflections (BSRs; Townend 1997; Henrys et al. 2003) and bottom hole temperature measurements from exploration boreholes (Field et al. 1997). BSR estimates of heat flow decrease landward from ∼50 mW m−2 near the deformation front to ∼35 mW m−2 over a distance of 100 km (Henrys et al. 2003). Heat flow values from 14 boreholes distributed along the east coast of the North Island are generally between 40 and 80 mW m−2 with a trend to lower values in the south. Heat flow values north of Hawke Bay vary between 50 and 80 mW m−2, and south of Hawke Bay are 40–45 mW m−2 (Field et al. 1997). Thermal models of the Hikurangi subduction zone are consistent in showing a relatively cold margin due to the subduction of the old Pacific plate (McCaffrey et al. 2008; Fagereng & Ellis 2009; Gao & Wang 2014; Yabe et al. 2014). These models estimate the intersection of the plate boundary and the 100 °C isotherm at depths between 8 and 15 km. In this study we report and interpret new heat flow measurements made across the northern Hikurangi margin. Our goals are to: (1) characterize the thermal regime of the incoming plate, deformation front and margin slope; (2) link these observations with heat flow estimates based on BSR observations; and (3) develop new thermal models to better understand the thermal and hydrologic regime of the Hikurangi subduction zone. 2 TECTONIC SETTING The Hikurangi Trough marks the westward subduction of the Pacific Plate beneath the Australian Plate and lies at the southern end of the Tonga-Kermadec subduction zone. The Hikurangi Trough extends from the Raukumara Peninsula (northern North Island) to the Cook Strait (Fig. 1). In this region, the subducting Pacific Plate includes the Hikurangi Plateau, a large igneous province that formed in the early Cretaceous (ca. 118–96 Ma; Davy & Wood 1994; Wood & Davy 1994; Taylor 2006; Davy et al. 2008; Hoernle et al. 2010). Subduction began 20–25 Ma (Ballance 1976; Walcott 1987). The convergence rate along the Hikurangi Trough decreases from 6 cm yr−1 to 3 cm yr−1 from north to south (Wallace et al. 2004). In the area of our heat flow measurements, 80 per cent of the ∼40 mm yr−1 plate motion is accommodated on the subduction thrust with the remainder accommodated by margin-normal shortening, strike-slip faulting and vertical-axis rotations (Nicol & Beavan 2003). The Hikurangi Plateau is a remnant of the larger Ontong-Java-Manihiki-Hikurangi large igneous province (Taylor 2006; Davy et al. 2008). This plateau is thought to have been emplaced upon pre-existing oceanic lithosphere that was young and thin (Kroenke et al. 2004). The onset of breakup between the Hikurangi and the rest of the plateau occurred ≥115 Ma (Mortimer et al. 2006). Despite more recent stages of seamount volcanism, 40Ar/39Ar ages show that Hikurangi basement lavas are 118–96 Ma such that volcanism continued after the initiation of breakup (Hoernle et al. 2010). The lithospheric thickness of the Hikurangi Plateau is ∼75 km (Stern et al. 2015), and its subducted western margin has been imaged to a depth of ∼65 km (Reyners et al. 2006). Subduction of the relatively buoyant Hikurangi Plateau uplifted the trough and forearc such that east of New Zealand the deformation front lies at a depth of 3 km, much shallower than the >9 km deep trench characterizing the Tonga-Kermadec system to the north (Fig. 1). On the incoming plate north of Hawke Bay, seamounts are common and basement relief commonly exceeds several hundred metres. The Tūranganui Knolls are a prominent feature with relief of up to 860 m. Offshore of the Raukumara Peninsula, north of where the Hikurangi channel turns east, sediments are 1–2 km thick on average (Lewis et al. 1998; Lewis & Barnes 1999; Mountjoy et al. 2009; Pedley et al. 2010). These sediments are Pliocene-Recent turbidites and mudstones that overlie sedimentary rocks of Cretaceous–Paleogene age (Plaza-Faverola et al. 2012). The Hikurangi margin north of Hawke Bay consists of a narrow frontal prism, an outer accretionary margin composed of mainly Plio-Pleistocene turbidites, and an inner margin consisting of a highly deformed Cretaceous–Paleogene sequence overlain by deformed Miocene-Recent basins (e.g. Barker et al. 2009; Barnes et al. 2010; Bell et al. 2010; Pedley et al. 2010). Locally the margin shows embayments where seamounts may have subducted (Collot et al. 2001). The taper is ∼13°, and splay faults that sole into the main detachment are common in the outer wedge. Beneath the northern Hikurangi margin, the subduction thrust is thought to develop at or close to the top of the Cretaceous/Paleogene sequence (Barker et al. 2009). Landward of the deformation front, the dip of the subducted plate increases from <3° near the front to <8° at depths of 10–15 km (Barker et al. 2009), then remains relatively constant until the depth of the plate interface reaches about 40 km (Ansell & Bannister 1996; Reyners 1998). The gentle dip is attributed to the subduction of the unusually thick (10–15 km) and buoyant crust of the Hikurangi Plateau (Wood & Davy 1994). The subducting slab is well defined by seismicity. Seismic tomography reveals a high P-wave velocity (Vp) region extending to a depth of >300 km, corresponding to the relatively cold and dense subducting Pacific slab (Reyners et al. 2006). Below the subduction thrust and ∼40 km landward of the deformation front, Bell et al. (2010) interpret high-amplitude reflectivity in seismic reflection data as indicating fluid-rich underplated sediments. Immediately seaward of this high-reflectivity zone, low reflectivity and a seaward dipping reflection is interpreted as a subducting seamount (Bell et al. 2010). 3 THERMAL DATA Marine heat flow data were collected between 2015 May 20 and June 15 from the R/V Roger Revelle, cruise RR1508, and between 2016 June 19 and 30 from the R/V Tangaroa, cruise TAN1607. A total of 96 new heat flow determinations were collected across the northern Hikurangi margin (Fig. 1). Heat flow data were acquired using a 3.5-m, 11-thermistor, violin-bow style heat flow probe (Hyndman et al. 1979). The heat flow probe consists of a weight stand containing a data logger, pressure sensor and acoustic transponder. The acoustic transponder sends enough data from the probe to the ship to monitor probe performance. The 3.5-m lance supports a thermistor string containing 11 equally spaced thermistors and a heating element. This design, a robust lance with an offset and tensioned thermistor string, provides the mechanical strength to withstand repeated insertions into the seafloor and the sensitivity needed to make precise measurements. The probe is deployed in a multipenetration ‘pogo’ mode whereby transects of measurements are made with a single lowering of the probe through the water column. Each measurement starts by piercing the sediment with the probe. A sensor monitors probe tilt and no reported measurements have an excessively high tilt (>30°). Inserting the probe into the bottom creates a frictional heat pulse, whose decay is monitored for approximately 7 min. Although the thermistors do not reach equilibrium in the 7-min period, it is long enough to extrapolate the data to derive the in situ temperature based on a line source model of radial heat conduction and a known thermal conductivity. Knowledge of the in situ temperatures as a function of subseafloor depth (including consideration of instrument tilt) yields a single estimate of the thermal gradient. Following the 7-min decay, a 20-s calibrated heat pulse is generated along the heating wire and monitored for another 7 min so that in situ thermal conductivity can be determined. The thermal conductivity is proportional to the decay rate of the heat pulse. The mathematical details of each step are given in Villinger & Davis (1987) and Hartmann & Villinger (2002). In practice, nonlinear inversion techniques are used to calculate the equilibrium temperature and thermal conductivity at each thermistor (Hartmann & Villinger 2002). The conductive heat flow through the seafloor is calculated using summed thermal resistance (Bullard 1939), \begin{eqnarray*} T(z) = {T_0} + {q_0}\sum\limits_{i = 1}^n {\frac{{\Delta {z_i}}}{{k{{(z)}_i}}}}, \end{eqnarray*} (1) where T(z) is the equilibrium temperature at depth z, T0 is the temperature at the seafloor, q0 is the surface heat flow, and k(z) is the thermal conductivity measurement over the ith depth interval Δzi. The summation is performed over n depth intervals. q0 and T0 are estimated by plotting temperature as a function of summed thermal resistance. Heat flow data were analysed and corrected for environmental factors including bottom water temperature (BWT) variations, bathymetry and sedimentation (Supporting Information S1). BWTs were recorded for approximately 1 year prior to our measurements during the HOBITSS experiment (Wallace et al. 2016) and indicate that heat flow measurements below a depth of 1500 m are not adversely affected by BWT variations. Bathymetric corrections are estimated using a finite-difference algorithm that computes conductive perturbations due to seafloor relief (Phrampus et al. 2014). These corrections have a mean of 1 per cent but are as large as 46 per cent in areas of large relief. Sedimentation corrections are made using a 1-D finite-difference algorithm (Hutnak & Fisher 2007) based on sediment thickness and stratigraphic interpretations of Orpin et al. (2006), Davy et al. (2008), Pedly et al. (2010), Plaza Faverola et al. (2012) and Henrys et al. (2013). These corrections are as large as 20 per cent. 4 RESULTS We collected heat flow data coincident with parts of seismic lines TAN1114–01, TAN1114–24 and 05CM-04 (Fig. 1) with the objectives of: (1) estimating the background thermal state of the Hikurangi Plateau, (2) testing if advective fluid flow is associated with basement relief and (3) determining the heat flow across the deformation front and along the 2018 IODP Expeditions 372 and 375 drilling transects. Measurement spacing along these transects is nominally 400 m but decreases in areas of special interest. Fig. 2 shows the heat flow transect along seismic line TAN1114–01. This transect is the most seaward transect and was located to minimize the effects of bathymetry and any advective fluid flow while still being collocated with a seismic reflection profile. Corrected heat flow values vary systematically between 45 and 71 mW m−2 with a mean and standard deviation of 56 ± 10 mW m−2. At one location we made two measurements; their difference is 1 mW m−2. Figure 2. View largeDownload slide Heat flow values along seismic line TAN1114–01 to estimate background heat flow. (a) Heat flow data values. Open circles show observed values and red circles show correction for sedimentation. No bathymetry correction is applied because of flat seafloor. Open circle with ‘X’ shows unsuccessful penetration. Average background heat flow (58 ± 8 mW m−2) is shown as dashed line and the standard deviation is indicated by grey band. (b) Seismic reflection profile TAN1114–01 plotted as a function of common depth point (CDP). Figure 2. View largeDownload slide Heat flow values along seismic line TAN1114–01 to estimate background heat flow. (a) Heat flow data values. Open circles show observed values and red circles show correction for sedimentation. No bathymetry correction is applied because of flat seafloor. Open circle with ‘X’ shows unsuccessful penetration. Average background heat flow (58 ± 8 mW m−2) is shown as dashed line and the standard deviation is indicated by grey band. (b) Seismic reflection profile TAN1114–01 plotted as a function of common depth point (CDP). Fig. 3 shows heat flow across the deformation front along seismic line TAN1114–24, the northern crossing of the deformation front. Seaward of the deformation front, the most conspicuous feature is a sharp peak in heat flow with a maximum value of 103 mW m−2 that decreases to a value of 39 mW m−2 at the eastern end of the transect. Figure 3. View largeDownload slide Heat flow data along seismic line TAN1114–24. (a) Heat flow values. Open circles show observed values, grey shaded circles show values corrected for bathymetry, and red circles show values corrected additionally for sedimentation. Open circles with ‘X’ show unsuccessful penetrations. Average background heat flow (58 ± 8 mW m−2) is shown as dashed line and the standard deviation is indicated by grey band. (b) Seismic reflection line 1114–24 plotted as a function of common depth point (CDP). Dashed line shows sediment basement interface. Figure 3. View largeDownload slide Heat flow data along seismic line TAN1114–24. (a) Heat flow values. Open circles show observed values, grey shaded circles show values corrected for bathymetry, and red circles show values corrected additionally for sedimentation. Open circles with ‘X’ show unsuccessful penetrations. Average background heat flow (58 ± 8 mW m−2) is shown as dashed line and the standard deviation is indicated by grey band. (b) Seismic reflection line 1114–24 plotted as a function of common depth point (CDP). Dashed line shows sediment basement interface. Heat flow measurements near the Tūranganui Knolls are used to assess if advective fluid flow is associated with exposed basement (Fig. 4). Heat flow determinations between common depth point (CDP) values of 7300 and 7400 along seismic line 05CM-04 are on the Knolls and adjacent to a pinnacle. This station was cut short to fix an electrical issue with the probe. Corrected heat flow values vary between 65 and 82 mW m−2. Figure 4. View largeDownload slide Heat flow data along seismic line 05CM-04. (a) Heat flow values. Open circles are observed values, grey shaded circles show values corrected for bathymetry and red circles show values corrected additionally for sedimentation. The average background heat flow (58 ± 8 mW m−2) is shown as dashed line and the standard deviation is indicated by grey band. (b) Simplified interpretation seismic profile 05CM-04 (Barker et al. 2009) plotted as a function of common depth point (CDP). Red line shows subduction interface, dashed red lines show major splay faults and dashed black line is inferred top of subducting seamount. BSR, bottom simulating reflector. Blue shaded region denotes area of high-amplitude reflectivity (HRZ-2) discussed in Bell et al. (2010). (c) Depth converted interpretation based on velocity model in Barker et al. (2009). Panels (b) and (c) are modified from Ellis et al. (2015). Figure 4. View largeDownload slide Heat flow data along seismic line 05CM-04. (a) Heat flow values. Open circles are observed values, grey shaded circles show values corrected for bathymetry and red circles show values corrected additionally for sedimentation. The average background heat flow (58 ± 8 mW m−2) is shown as dashed line and the standard deviation is indicated by grey band. (b) Simplified interpretation seismic profile 05CM-04 (Barker et al. 2009) plotted as a function of common depth point (CDP). Red line shows subduction interface, dashed red lines show major splay faults and dashed black line is inferred top of subducting seamount. BSR, bottom simulating reflector. Blue shaded region denotes area of high-amplitude reflectivity (HRZ-2) discussed in Bell et al. (2010). (c) Depth converted interpretation based on velocity model in Barker et al. (2009). Panels (b) and (c) are modified from Ellis et al. (2015). The second set of values between CDPs 5800 and 6500 abut the Tūranganui Knolls. Immediately adjacent to the Knolls, heat flow is 28 mW m−2 which then increases to a relatively uniform value of ∼ 62 mW m−2 over a distance of about 4 km. Farther west, between CDPs 4750 and 5500, and similar to the heat flow transect along seismic line TAN1114–24, a buried basement high is present seaward of the deformation front and heat flow appears inversely correlated to sediment thickness (Fig. 4). The correlation is not as robust as along seismic line TAN1114–24. Landward of the deformation front, heat flow drops from ∼55 to ∼ 40 mW m−2, but then increases over the first anticlinal ridge before showing a linear decrease to values of approximately 33 mW m−2 (Fig. 4). At CDP 4250, heat flow jumps to a value of about 65 mW m−2 before declining to a little less than 25 mW m−2 at CDP 4009. This local maximum coincides with splay faults that intersect the seafloor just behind the first anticlinal ridge. Landward of this minimum, heat flow increases to about 40 mW m−2. Inland of CDP 4000, heat flow values are low with a slight landward increase. The most landward measurements are coincident with the location of subsequent IODP Expedition 375 drilling (Fig. 4). Here BWT variations perturb the shallow sub-bottom and their removal is described in Supporting Information S1. We estimate that the heat flow in this area is 32 ± 5 mW m−2. 5 THERMAL MODEL OF THE SHALLOW SUBDUCTION ZONE Heat flow data landward of the deformation front show a broad westward decline in values consistent with landward thickening of the margin (Wang et al. 1993) and advection of cold material downward (Fig. 4). Superimposed on the broad landward decline are short-wavelength variations. To better understand the long-wavelength heat flow variations and the thermal regime we constructed a 2-D finite-element model using the source code developed by Wang et al. (1995b). This algorithm solves the steady-state heat conduction-advection equation described by \begin{eqnarray*} \nabla \cdot \left( {k\nabla T} \right) - \rho cv \cdot \nabla T + A = 0, \end{eqnarray*} (2) where k is thermal conductivity, T is temperature, ρc is volumetric heat capacity, v is the convergence rate of the subducting slab, and A is heat generation. Heat is advected with the subducting plate at a rate of 32 mm yr−1 (Wallace et al. 2004). Flow in the mantle wedge results from viscous coupling between the mantle and subducting plate. We use an analytic corner flow solution (Batchelor 1967) with an isoviscous rheology to model flow in the mantle wedge (Peacock & Wang 1999). The details of the mantle rheology have only a small effect on the shallow portion of the subducting thrust (e.g. Currie et al. 2004), which is the focus of this study. Because of these simplifications our model is not appropriate for detailed thermal analysis of the mantle wedge. The 2-D thermal model (Fig. 5) is positioned along seismic line 05CM-04, which is approximately normal to the trench (Fig. 1) and extends from the deformation front to 300 km landward. The landward boundary is sufficiently far from the forearc that the boundary condition has a negligible effect. The model uses a horizontal grid spacing of 1 km for the first 30 km landward of the deformation front. Landward of this point the horizontal grid spacing is 5 km. The bottom boundary is at a minimum depth of 95 km and the temperature is set to 1450 °C. These parameters correspond to the plate cooling model for old (>55 Ma) oceanic lithosphere (Stein & Stein 1992). The top boundary is set to 1 °C at the deformation front, increased to 15 °C over the margin consistent with temperature-depth values of conductivity-temperature-depth (CTD) casts, and is set to 15 °C over subaerial portions of the domain. The seaward boundary condition is based on a conductive geotherm appropriate for 120 Ma seafloor using parameters from GDH1 (Stein & Stein 1992). The landward boundary condition is defined by a geotherm that gives a surface heat flow of 80 mW m−2, appropriate for a backarc setting (Currie & Hyndman 2006). The geotherm through the mantle wedge follows an adiabatic gradient of 0.4 °C km−1. Figure 5. View largeDownload slide Model grid and thermophysical model units. The model consists of 80 km thick oceanic lithosphere, a sedimentary prism, upper and lower crust, and a mantle wedge. The seaward temperature boundary condition is set by the plate cooling model for 90 Ma crust (Stein & Stein 1992) with 1 km of sediment cover. The landward temperature boundary condition consists of 80 mW m−2 geotherm through the upper plate crust (corresponding to a temperature gradient of 24 °C km−1 for uniform thermal conductivity of 2.5 W m−1 K−1) and an adiabatic mantle gradient of 0.4 °C km−1. Figure 5. View largeDownload slide Model grid and thermophysical model units. The model consists of 80 km thick oceanic lithosphere, a sedimentary prism, upper and lower crust, and a mantle wedge. The seaward temperature boundary condition is set by the plate cooling model for 90 Ma crust (Stein & Stein 1992) with 1 km of sediment cover. The landward temperature boundary condition consists of 80 mW m−2 geotherm through the upper plate crust (corresponding to a temperature gradient of 24 °C km−1 for uniform thermal conductivity of 2.5 W m−1 K−1) and an adiabatic mantle gradient of 0.4 °C km−1. The geometry of the modelled shallow subduction thrust between the deformation front and 50 km landward is guided by seismic reflection data from seismic lines 05CM-04 and 05CM-05 (Bell et al. 2010). Between 50 km and the Taupo Volcanic Zone the plate interface is based on relocated earthquake hypocentres (Williams et al. 2013), and west of the Taupo Volcanic Zone (∼220 km landward of the deformation front), the subduction thrust is assigned a constant slope of 43°. Bathymetric and topographic data define the top of the margin and subaerial region (Ryan et al. 2009). We divide the thermal model into five thermal physical units (Fig. 5). Each thermal unit is assigned homogenous thermophysical parameters that represent bulk values (Table 1). The incoming sediment is 1200 m thick at the deformation front. We assume the Hikurangi Plateau has properties similar to oceanic lithosphere. Heat production measured on core samples from industry wells predominantly lies between 1.07 and 1.95 μW m−3 with a mean of about 1.7 μW m−3 (Pandey 1981). The margin extends from the deformation front to about 170 km landward where it is juxtaposed against the indurated Cretaceous age sediments of the Raukumara ranges. Table 1. Material parameters of the 2-D finite-element model. Model unit Thermal conductivity Volumetric heat capacity Heat generation (W m−1 K−1) (MJ K−1 m−3) (μW m−3) Sediment 1.6a 2.5 0.7b Margin 1.6a – 1.7c Upper continental crust 2.5b – 1.8b Lower continental crust 2.5b – 0.4b Mantle wedge 3.1b 3.3 0.02b Lithosphere 3.1e 3.3 0.02b Model unit Thermal conductivity Volumetric heat capacity Heat generation (W m−1 K−1) (MJ K−1 m−3) (μW m−3) Sediment 1.6a 2.5 0.7b Margin 1.6a – 1.7c Upper continental crust 2.5b – 1.8b Lower continental crust 2.5b – 0.4b Mantle wedge 3.1b 3.3 0.02b Lithosphere 3.1e 3.3 0.02b a Derived from Supporting Information eqs (S3) and (S4). b Gao & Wang (2014). c Shipboard Scientific Party (1999a). d Shipboard Scientific Party (1999b). e Stein & Stein (1992) View Large Table 1. Material parameters of the 2-D finite-element model. Model unit Thermal conductivity Volumetric heat capacity Heat generation (W m−1 K−1) (MJ K−1 m−3) (μW m−3) Sediment 1.6a 2.5 0.7b Margin 1.6a – 1.7c Upper continental crust 2.5b – 1.8b Lower continental crust 2.5b – 0.4b Mantle wedge 3.1b 3.3 0.02b Lithosphere 3.1e 3.3 0.02b Model unit Thermal conductivity Volumetric heat capacity Heat generation (W m−1 K−1) (MJ K−1 m−3) (μW m−3) Sediment 1.6a 2.5 0.7b Margin 1.6a – 1.7c Upper continental crust 2.5b – 1.8b Lower continental crust 2.5b – 0.4b Mantle wedge 3.1b 3.3 0.02b Lithosphere 3.1e 3.3 0.02b a Derived from Supporting Information eqs (S3) and (S4). b Gao & Wang (2014). c Shipboard Scientific Party (1999a). d Shipboard Scientific Party (1999b). e Stein & Stein (1992) View Large We incorporate shear heating along the plate interface following Wang et al. (1995b). This formulation assumes that the long-term plate convergence and frictional properties at any point do not vary as a function of time. Importantly, we are assuming that the surface heat flow used to estimate the frictional heat, Af, depends on the long-term average rate of heat generation. Individual events, even very large ones, only increase heat flow close to the fault for short periods of time (e.g. Fulton et al. 2010, 2013). Frictionally generated heat builds over long periods and diffuses toward the seafloor. Thus our results are insensitive to individual events or when in the seismic cycle heat flow is measured. In the brittle regime, Af, can be expressed as \begin{eqnarray*} {A_\mathrm{ f}} = {\tau _b}\frac{v}{w} = \mu^*{\sigma _n}\frac{v}{w}, \end{eqnarray*} (3) where τb is shear stress calculated using Byerlee's Friction Law (Bylerlee 1978), w is the fault width, μ* is the effective coefficient of friction and σn is the normal stress. In the ductile regime, frictional heating can be expressed as \begin{eqnarray*} {A_\mathrm{ f}} = {\tau _d}\dot{\varepsilon } = {\tau _d}\frac{v}{w}, \end{eqnarray*} (4) where |$\dot{\varepsilon }$| is the strain rate. Here τd is calculated using a power-law rheology with parameters appropriate for diabase (Caristan 1982). The ductile regime is the region where the τd < τb. Shear heating is turned off near 150 km, where the plate boundary enters the mantle. In eqs (3) and (4) the width of deformation is difficult to assign with confidence. In our models we follow the approach of Gao & Wang (2014) and assign a thickness of w = 500 m. 6 DISCUSSION 6.1 Seaward of the deformation front Heat flow measurements were made seaward of the deformation front to satisfy two objectives. The first was to provide an estimate of the background heat flow and the second was discern whether or not advective fluid flow in the oceanic basement of the incoming Hikurangi Plateau is significant. Heat flow was measured along seismic line TAN1114–01 to estimate the background heat flow (Fig. 2). Unfortunately, there is significant unexplained variation yielding a mean and standard deviation of 56 ± 10 mW m−2 in corrected values. Along seismic line 05CM-04 (Fig. 4) between CDPs 5889 and 6305 heat flow is relatively uniform with a mean and standard deviation of 61 ± 5 mW m−2. Although it is interesting that the heat flow west of the Tūranganui Knolls is about 5 mW m−2 greater than to the east, this difference is within the observed variability of the measurements. As discussed below, although we document evidence for fluid flow we do not see evidence for large-scale advective heat transport within the upper oceanic crust. The mean and standard deviation of these values is 58 ± 8 mW m−2 that we use to represent background. Although the Hikurangi Plateau does not necessarily have the same thermal structure as a normal oceanic plate of the same age, it is useful to compare our value with globally observed heat flow values and conductive cooling models of oceanic lithosphere. Global compilations of heat flow data show that heat flow changes little on lithosphere older than about 80 Ma. The global average heat flow between 90 and 120 Ma (the age range of Hikurangi basement lavas) is 56 ± 15 mW m−2 (Stein & Stein 1993). For seafloor ages of 90–120 Ma, the GDH1 cooling plate model (Stein & Stein 1992) predicts heat flow values of 56–51 mW m−2 whereas RT1 (Grose 2012) predicts heat flow values of 62–55 mW m−2. The similarity of our measured values to the global mean and conductive models suggests that the thermal effects of plateau emplacement, relative to normal oceanic lithosphere, have diffused away. In areas of thin or patchy sediment, large variability in heat flow values, not attributed to heat refraction, changes in BWT, sedimentation or erosion, can often be due to discharge of fluids from a basement aquifer or recharge of fluids from the ocean (e.g. Hutnak et al. 2006; 2007). Heat flow data associated with the Tūranganui Knolls show compelling evidence for fluid flow (Fig. 4). Near the sediment pinch out, at CDP 6550, values are 48 per cent below the mean and then climb to mean values toward the west. This pattern is consistent with recharge into exposed basement. Extrapolated temperatures along the sediment–basement interface vary between 4 and 85 °C. The lower temperatures are consistent with cool seawater recharging the upper basement. With these data it is not possible to determine the extent to which fluids are flowing to the west, east or in and out of the plane of the transect. To the east, the higher than average heat flow associated with the pinnacle is probably due to fluid discharge (near CDP 7400). These observations are consistent with the results of fluid flow models through pillow basalts forming oceanic layer 2A, which predict recharge into exposed basement will depress temperatures locally and discharge through basement highs will increase temperatures. Although it might be tempting to infer a connection between the recharging and discharging areas, it is not demanded by the data and several different flow paths may be present, including flow into and out of the plane of our measurements. In areas of more continuous sediment cover, inverse correlations between heat flow and sediment thickness suggest vigorous fluid flow within oceanic basement. This condition occurs when the seafloor is isothermal and fluid flow is sufficiently vigorous to maintain the top of the permeable basement at a nearly constant temperature so that the thermal gradient is inversely proportional to sediment thickness (Davis et al. 1989). Evidence for fluid flow within oceanic basement comes from buried basement highs just seaward of the deformation front. The most dramatic example is along the heat flow transect collocated with seismic line TAN1114–24 (Fig. 3). The seafloor slopes gently to the east and the total relief is less than 80 m. We converted sediment two-way traveltime to thickness (Supporting Information eq. S2) and find that it thins from about 1060 to 310 m over the top of the basement high. Heat flow varies from 40 to 103 mW m−2 and values are inversely correlated with sediment thickness suggesting that this anomaly reflects fluid flow in the upper basement. A lateral contrast in thermal conductivity can preferentially refract heat through the higher thermal conductivity material causing a heat flow anomaly. Assuming a contrast in thermal conductivity between the basement and sediments of 2.5:1 leads to a maximum heat flow anomaly of ∼60 per cent in contrast to the observed anomaly of 164 per cent. Estimated basement temperatures are remarkably uniform between 30 and 40 °C (Fig. 6a). The mean heat flow for this portion of the data is 62 mW m–2, a little lower than the estimated mean of the basal heat flux over the field area, indicating that although flow is vigorous enough to maintain nearly isothermal conditions at the sediment–basement interface, little heat is being transported into or out of this region. Figure 6. View largeDownload slide Inverse relationship between heat flow and sediment thickness. Open circles show observed data and red circles show data corrected for both bathymetry and sedimentation. Lines are computed assuming an isothermal sediment–basement interface. The effects of thermal refraction are small. (a) Heat flow values along seismic line TAN1114–24 between common depth points (CDPs) 3000 and 4500. (b) Heat flow values along seismic line 05CM-04 between CDPs 4800 and 5500. Figure 6. View largeDownload slide Inverse relationship between heat flow and sediment thickness. Open circles show observed data and red circles show data corrected for both bathymetry and sedimentation. Lines are computed assuming an isothermal sediment–basement interface. The effects of thermal refraction are small. (a) Heat flow values along seismic line TAN1114–24 between common depth points (CDPs) 3000 and 4500. (b) Heat flow values along seismic line 05CM-04 between CDPs 4800 and 5500. Along seismic line 05CM-04, a basement high similar to that seen on TAN1114–24 is observed just seaward of the deformation front (Fig. 4), but the inverse correlation between sediment thickness and heat flow is less compelling (Fig. 6b). Sediment thickness along this line varies between about 1200 and 500 m, so the scale of the buried basement relief here is similar (700 m at 05CM-04 versus 750 m at TAN1114–24). Heat flow varies between 45 and 83 mW m−2 with a mean value of 61 mW m−2. Extrapolated basement temperatures vary between 28 and 49 °C showing greater variability than along line TAN1114–24 (Fig. 6b). Therefore, advective fluid flow appears to be less vigorous here than along line TAN1114–24. The correlation between heat flow and sediment thickness along line TAN1114–01 (Fig. 2) is opposite to that observed in Figs. 3 and 4. There is a slight decrease in the depth to the sediment–basement interface associated with low heat flow. Although we have not found a satisfactory explanation for the observed variability in heat flow it may be due to out-of-plane effects. In these examples, the mean heat flow value is consistent with global cooling models of oceanic lithosphere (Stein & Stein 1992) indicating that significant quantities of heat are not being advected out of the system. We suggest that these observed heat flow anomalies are locally, rather than regionally, significant. Global heat flow data show that hydrothermal circulation associated with ridge flanks plays an important role in advecting heat from the lithosphere until ∼60–70 Ma (Stein & Stein 1994), at which time the global average heat flow anomaly approaches zero. The agreement between heat flow data and plate cooling models for crust older than ∼60–70 Ma is taken as evidence that advective fluid flow is not removing heat from the crust, and this age is often referred to as the ‘sealing age’. Our finding of advectively significant fluid flow is anomalous but not unique. Advectively significant hydrothermal circulation has been documented in oceanic crust much older than the reported 60–70 Ma ‘sealing age’. In a re-assessment of heat flow values on old oceanic crust, Von Herzen (2004) argued that heat flow variations resulting from heat refraction due to basement or seafloor relief and contrasts in thermal conductivity between basement and sediment are unlikely to be more than about 25 per cent of the regional value. That is, observed heat flow variations greater than about 25 per cent of the regional value are probably due to hydrothermal circulation. Of 58 heat flow surveys on old seafloor re-assessed by Von Herzen (2004), a significant portion remains hydrothermally active. More recently, advective fluid flow has been documented on the 135 Ma crust entering the Japan trench (Kawada et al. 2014; Yamano et al. 2014). Thus, the finding that fluid flow plays a significant role in heat advection may not be surprising given the presence of exposed basement around the Tūranganui Knolls. The picture that emerges from these studies suggests that old seafloor contains sufficient permeability to host fluid flow and that fluid flow will occur if driving forces are sufficient. Areas with bathymetric gradients and exposed basement are documented as favourable environments for hosting significant advective fluid flow (e.g. Lowell 1980; Harris et al. 2004). 6.2 Landward of the deformation front Fig. 7 shows the modelled thermal regime and surface heat flow. We add heat flow estimates based on an analysis of BSRs to our shallow probe measurements. The conversion of BSR depth to heat flow is described in Supporting Information S2. Continental heat flow determinations in line with our transect come from boreholes Rere-1 and Waitangi-1 (Field et al. 1997). These two determinations are based on bottom hole temperatures and a mean surface temperature of 14 °C, and combined with thermal conductivity measurements give values of 64 and 57 mW m−2, respectively. The thermal field at these boreholes is transiently responding to uplift and erosion at a rate of approximately 0.3 mm yr−1 over the last 8 Ma (Litchfield et al. 2007). An analytic formula (Powell et al. 1988) shows that this rate and duration of erosion increases surface heat flow by approximately 16 per cent. We remove this effect by decreasing the observed heat flow by this amount so that the basal heat flow values at Rere-1 and Waitangi-1 are 54 and 48 mW m−2, respectively. Other values of heat flow along the east coast of the North Island are present but fall farther from our transect. Values within 50 km of the transect are also decreased by 16 per cent and are shown as open triangles in Fig. 7. One additional value from the Global Heat Flow database (http://www.heatflow.org/data, compiled by D. Hasterok) about 250 km from the deformation front is also included. We first explore the landward increase in heat flow across the forearc and then turn to short-wavelength variations near the deformation front. Figure 7. View largeDownload slide Thermal models across the northern Hikurangi margin, New Zealand. (a) Heat flow. Red circles show heat flow determinations, shaded regions show heat flow from bottom simulating reflections (BSR) and triangles show heat flow from continental boreholes. Filled and hollow triangles are within 15 and 100 km of the transect, respectively. Solid lines show modelled heat flow as a function of the effective coefficient of friction, μ’ with all relative motion along the subduction thrust. Var denotes variable effective coefficient of friction, for distances ≤50 km μ* = 0.06, for distances >50 km μ* = 0.18. (b) Preferred thermal model with variable μ*. Thermal model is computed using source code developed by Wang et al. (1995b). Contour interval is 100 °C except for the bottom contour which is 1450 °C. White lines show thermophysical unit boundaries. Vectors show relative motion. Figure 7. View largeDownload slide Thermal models across the northern Hikurangi margin, New Zealand. (a) Heat flow. Red circles show heat flow determinations, shaded regions show heat flow from bottom simulating reflections (BSR) and triangles show heat flow from continental boreholes. Filled and hollow triangles are within 15 and 100 km of the transect, respectively. Solid lines show modelled heat flow as a function of the effective coefficient of friction, μ’ with all relative motion along the subduction thrust. Var denotes variable effective coefficient of friction, for distances ≤50 km μ* = 0.06, for distances >50 km μ* = 0.18. (b) Preferred thermal model with variable μ*. Thermal model is computed using source code developed by Wang et al. (1995b). Contour interval is 100 °C except for the bottom contour which is 1450 °C. White lines show thermophysical unit boundaries. Vectors show relative motion. 6.2.1 Evidence for a landward increase in the effective coefficient of friction Fig. 7(a) shows seafloor heat flow as a function of the effective coefficient of friction along the plate boundary for values between 0 and 0.18 as well as for our preferred model in which the effective coefficient of friction increases from 0.06 to 0.18 at 50 km landward of the deformation front. The values of heat flow from both the shallow probe data and BSRs seaward of the coastline suggest a relatively low effective coefficient of friction whereas the continental values of heat flow suggest a higher effective coefficient of friction (Fig. 7a). Although not shown, we explored the possibility that heat was being transferred from deep in the system towards the deformation front as suggested by Spinelli & Wang (2008) at the Nankai trough. Consistent with our assertion that advective fluid flow is locally but not regionally significant, these models degraded the model fit to the data and are not discussed further. Our preferred thermal model is shown in Fig. 7(b). White lines show boundaries of our thermophysical units (Table 1) overlain on the calculated isotherms. Flow vectors show the downward advection of the oceanic plate leading to the characteristic tongue of cold temperatures in the upper oceanic crust. Modelled heat flow just landward of the deformation front decreases rapidly from ∼ 50 to ∼30 mW m2 before increasing to ∼75 mW m−2 at 250 km (Fig. 7a). Although there are trade-offs between bulk thermal physical rock properties in the margin and the effective coefficient of friction on the plate boundary, varying the heat production between 2.0 and 0.7 μW m−3 in the margin does not produce a significantly better fit between models with low or high values of μ*. We also determined that increasing the thermal conductivity of the inner margin from 1.6 to 2.0 between 50 km from the deformation front and the continent did not explain the misfit between onshore and offshore observations for models with a constant μ*. We conclude that a downdip increase in μ* is the best explanation for the observed change from lower heat flow to higher heat flow across the coastline. It is worthwhile to compare our values for the coefficient of friction with previously estimated values. Both McCaffrey et al. (2008) and Yabe et al. (2014) suggest a low effective coefficient of friction between 0.01 and 0.05 similar to our estimate for μ* on the shallow plate boundary. Note that the McCaffrey et al. (2008) model is constrained by BSR values of heat flow. In contrast, the warmest model by Fagereng & Ellis (2009) assumes hydrostatic pore pressure that implies a relatively high coefficient of friction of about 0.4. Even with enhanced shear heating their model underpredicts continental heat flow values. This discrepancy (> 10 mW m−2) may result from the low value of background heat flow they assumed seaward of the deformation front (∼ 44 mW m−2). Gao & Wang (2014) fit their model to both marine and continental heat flow values and suggest that a value of μ* = 0.13 is most appropriate, which is close to the average of the two values we obtain. We note that both Fagereng & Ellis (2009) and Gao & Wang (2014) rely on heat flow values projected onto their model profiles over a considerable distance (∼150 km) potentially smearing along strike heterogeneity in the measurements. In contrast, our choice of preferred model relies on probe and BSR values of heat flow along a single transect in the outer and mid-slope region of the forearc. If the intrinsic coefficient of friction is 0.4, appropriate for clays (Ikari & Saffer 2011), an effective coefficient of friction of 0.06 implies that the pore pressure ratio is close to 0.9. This high pore pressure ratio is consistent with the presence of fluid overpressures. Support for overpressure also comes from P-wave to S-wave velocity ratios (Vp/Vs) and P-wave attenuation (inverse Qp) values (Eberhart-Phillips et al. 2005, 2008; Reyners & Eberhart-Phillips 2009; Bassett et al. 2014), high-amplitude reflectivity along the fault plane (Bell et al. 2010), overpressures in exploration boreholes (Darby & Funnel 2001; Sibson & Rowland 2003), active mud volcanoes (Ridd 1970; Pettinga 2003) and thermomechanical numerical models (Ellis et al. 2015). The landward increase in the effective coefficient of friction, as indicated by the continental heat flow data, suggests either a decrease in the pore pressure ratio, an increase in the intrinsic coefficient of friction, or some combination of both. Recent friction experiments on sediment comprising the input material for this subduction zone show that the intrinsic coefficient of friction decreases with increasing depth (normal stress) between 0 and 2 km but can be fit with a nearly constant value between 2 and 10 km (Rabinowitz et al. 2018). Our inferred increase in the effective coefficient of friction is therefore consistent with a constant intrinsic coefficient of friction and decreasing pore pressure ratio. 6.2.2 Effect of slumps and splay faults on heat flow near the deformation front Although we used the thermal model of subduction to estimate the long-wavelength thermal field, there are short-wavelength features that are not fit. Possible explanations for the scatter in heat flow include sharp lateral variations in thermophysical rock properties, slip on splay faults or slumping that puts warmer material on top of cooler material, and fluid flow (Fig. 8). Reasonable variations in thermal conductivity and heat production do not significantly decrease the misfit between distances of about 3 and 8 km landward of the deformation front. Figure 8. View largeDownload slide Temperature along the plate boundary. Solid lines show modelled temperature as a function of the effective coefficient of friction, μ*. Var denotes variable effective coefficient of friction, for distances ≤50 km μ* = 0.06, for distances >50 km μ* = 0.18. The location of SSEs (Wallace et al. 2012). Temperature inversion shows position where rheology changes from brittle to ductile shear heating. Figure 8. View largeDownload slide Temperature along the plate boundary. Solid lines show modelled temperature as a function of the effective coefficient of friction, μ*. Var denotes variable effective coefficient of friction, for distances ≤50 km μ* = 0.06, for distances >50 km μ* = 0.18. The location of SSEs (Wallace et al. 2012). Temperature inversion shows position where rheology changes from brittle to ductile shear heating. Slumping may contribute to the scatter in heat flow. For example, a slump is mapped at a distance of about 7 km from the deformation front (Fig. 9b). The slump headwall exposes previously buried sediment that presumably leads to an increase in heat flow. If the slump block remains coherent, we would expect that the heat flow from its surface would remain approximately the same or decrease. This pattern correlates qualitatively with the heat flow data, but the mapped slump straddles the heat flow offset such that the upper part of the slump deposit is associated with high heat flow values and the lower part is associated with low heat flow values. It is difficult to reconcile the position of the slump deposit with the heat flow offset, but alignment between them cannot yet be ruled out because bathymetry at this location is complex and the scale of the bathymetry is significant compared to the lateral resolution of seismic line 05CM-04. Figure 9. View largeDownload slide Thermal model across the frontal part of the northern Hikurangi margin, New Zealand. (a) Heat flow. Red circles show heat flow determinations, shaded band shows heat flow from bottom simulating reflections (BSR). Solid lines show modelled heat flow as a function of the effective coefficient of friction, μ* with all relative motion along the subduction thrust. (b) Simplified interpretation seismic profile 05CM-04 (Barker et al. 2009). Red line shows subduction interface, dashed red lines show major splay faults and dashed black line is inferred top of subducting seamount. Yellow line shows the base of forearc sediments that are well stratified. Blue region shows position of high-amplitude reflectors suggesting fluid-rich sediments (Bell et al. 2010). (c) Schematic diagram of proposed fluid discharge generating the heat flow highs. In this view compaction of underthrust sediment (region denoted by lower ellipse) generates overpressures (Ellis et al. 2015) driving fluid escape through overlying material. Heat flow highs are interpreted in terms of fluid discharge. Figure 9. View largeDownload slide Thermal model across the frontal part of the northern Hikurangi margin, New Zealand. (a) Heat flow. Red circles show heat flow determinations, shaded band shows heat flow from bottom simulating reflections (BSR). Solid lines show modelled heat flow as a function of the effective coefficient of friction, μ* with all relative motion along the subduction thrust. (b) Simplified interpretation seismic profile 05CM-04 (Barker et al. 2009). Red line shows subduction interface, dashed red lines show major splay faults and dashed black line is inferred top of subducting seamount. Yellow line shows the base of forearc sediments that are well stratified. Blue region shows position of high-amplitude reflectors suggesting fluid-rich sediments (Bell et al. 2010). (c) Schematic diagram of proposed fluid discharge generating the heat flow highs. In this view compaction of underthrust sediment (region denoted by lower ellipse) generates overpressures (Ellis et al. 2015) driving fluid escape through overlying material. Heat flow highs are interpreted in terms of fluid discharge. The misfit (Fig. 9) of and large scatter in the data just landward of the deformation front can be qualitatively attributed to pervasive fracturing and fluid flow in the outer 10–15 km of the wedge. When viewed in detail, the short-wavelength variations in heat flow within about 15 km of the deformation are systematic and can be correlated with shallow structural variations. Fluid flow along shallow faults intersecting the seafloor may be responsible for much of the variability in heat flow observed near the wedge toe. Between CDPs 4375 and 4450 a series of splay faults intersect the surface. The trend of seaward increasing heat flow may be due to the geometry of the faults and warm fluids discharging through them. With this explanation, the heat flow offset seen at CDP 4330 corresponds to the intersection of a splay fault with the seafloor. There are several candidate splay faults interpreted seismically in this region but limits on image resolution and out-of-plane effects make correlation difficult. High-resolution bathymetry data (25 m), however, show a fault cutting across the slump, making it a likely candidate for fluid discharge. Lauer & Saffer (2012) have emphasized the importance of splay faults in the fluid budget of convergent margins. They found that splay faults may be responsible for discharging 10–30 per cent of incoming fluids at rates of ∼0.2–4.4 cm yr−1 for a reasonable range of model parameters. Pore pressures for these simulations were modestly elevated but below lithostatic values. Modelling results of Ellis et al. (2015) along line 05CM-04 support the existence of excess fluid pressure at the updip and downdip ends of a subducted seamount. The overpressure anomaly is at the updip end of the seamount that is 5–8 km deep and 6–13 km from the deformation front. For our preferred model (μ* = 0.06 in this region), this temperature range lies between approximately 10 and 60 km landward of the deformation front (Fig. 9), and the splay faults that surface near the toe may tap this region. We estimate fluid velocity using a solution to the 1-D steady-state advective diffusion equation with a constant temperature upper boundary and a constant background heat flow at depth L (Harris et al. 2010). The Darcy velocity, Vz, can be expressed as \begin{eqnarray*} {V_z} = \frac{k}{{\phi {\rho _\mathrm{ w}}{c_\mathrm{ w}}L}}\ln \left( {\frac{{{q_\mathrm{ b}}}}{{{q_\mathrm{ o}}}}} \right) \end{eqnarray*} (5) where k is the thermal conductivity, ϕ is the porosity, ρw is the density of water, cw is the heat capacity of water, qb is the background heat flow and qo is the observed heat flow at the surface. At CDP 4250 we use the offset in heat flow between 25 and 65 mW m−2 for qb and qo, respectively. Here the depth of the décollement is at 2.5 km as interpreted from the seismic reflection data (Fig. 5), and we assume the porosity is 20 per cent (Ellis et al. 2015) and the thermal conductivity is 1.6 W m−1 K−1. These values suggest an upward Darcy velocity of approximately 20 mm yr−1. This value lies within the range of values modelled by Lauer & Saffer (2012) for splay faults, and that modelled by Ellis et al. (2015) at this margin. Based on a similar analysis of a BSR anomaly higher on the margin, Pecher et al. (2017) calculated a flow rate of 10 mm yr−1. Our simplified model precludes a precise estimate of the Darcy velocity, but suggests that a reasonable flow rate can produce the observed anomaly. 6.2.3 Synthesis of fluid overpressures and flow Understanding the sources of fluid overpressures or flow is important for understanding slow slip and the hydrogeology of this margin. The primary fluid source in outer accretionary wedges is generally thought to result from mechanical compaction of the sediments, with a secondary contribution from dehydration of clays at temperatures of 50–150 °C (e.g. Peacock & Hyndman 1999; Saffer & Tobin 2011), In our model this temperature range occurs between 10 and 50 km from the deformation front (Fig. 8). Indeed, Ellis et al. (2015) modelled fluid production for the northern Hikurangi and concluded that the primary fluid source is compaction of sediments with a contribution from clay dehydration. Deeper in the system, dehydration of the crust and upper mantle does not begin until a temperature of ∼300 °C (e.g. Peacock & Hyndman 1999), which in our model occurs ∼90 km from the deformation front where the plate interface is at a depth of ∼18 km. This thermal structure results in a fluid input gap 50–90 km from the deformation front that may lead to an increase in the effective coefficient along this segment of the plate boundary. Although the pore pressure on the plate boundary may increase again at distances greater than 90 km, the heat flow data are not very sensitive to the effective coefficient of friction landward of ∼90 km (Fig. 7a). We note that Heise et al. (2012) detected a transition from intermediate to high electrical conductivity in this region, which they attributed to a downdip increase in pore pressure at ∼20 km depth, consistent with our temperature model. Our interpretation of overpressure (and consequently a low μ*) along the plate boundary and fluid flow through splay faults in the outer wedge may seem discrepant. Previous studies have inferred overpressures in this region (Bell et al. 2010; Bassett et al. 2014; Ellis et al. 2015) and many studies have shown a link between shallow SSEs and overpressures (Kodaira et al. 2004; Kitajima & Saffer 2012; Saffer & Wallace 2015). At the same time faults in the northern Hikurangi margin have been inferred to represent permeable pathways, facilitating the discharge of overpressured fluids (Ellis et al. 2015). Our interpretation of the heat flow data are consistent with these studies indicating the occurrence of advective fluid flow through the margin along permeable splay faults sourced from a deeper overpressured region. The combination of these phenomena implies that fluid flow along the splay faults is transient. The transient behaviour is most likely to be generated from stress cycling processes at the décollement (Sibson 1992,2018; Sibson & Rowland 2003). The region of inferred fluid flow (0–15 km landward of the deformation front) is seaward of an imaged seamount, whereas evidence for a lower coefficient of friction, and zone of high-amplitude reflectivity, is found to a distance of about 50 km landward of the deformation front. These interpretations are consistent if the seamount marks a change in the fluid pressure regime as suggested by numerical models and models of slow slip (Ellis et al. 2015; Wallace et al. 2016). 7 CONCLUSIONS On the basis of this study we conclude the following: Mean corrected heat flow seaward of the deformation front is ∼58 ± 8 mW m−2. This value is within the variability of globally observed heat flow for lithosphere of this age (56 ± 15 mW m−2) but slightly higher than the mean GDH1 (Stein & Stein 1992) prediction of 56–51 mW m−2 for 90–120 Ma lithosphere, respectively. Advective fluid flow is locally significant seaward of the deformation front. Evidence of recharge is conspicuous at the Tūranganui Knolls. Seaward of the deformation front, where the transects cross basement highs, an inverse correlation between sediment thickness and heat flow is observed. Along seismic line TAN1114–14 the basement–sediment interface is nearly isothermal to within ±5 °C and along seismic line 05CM-04 to within ±10 °C. Mean values of heat flow are similar to the observed background heat flow indicating that a significant quantity of heat is not being removed via advection. Heat flow from 0 to 12 km landward of the deformation front showing large variability is attributed to advective fluid flow along numerous faults, possibly enhanced by effects of slope instability. A simple analytical model of flow suggests Darcy flow rates of a few cm yr−1. Our analysis, which includes explicit corrections for changes in the BWT, bathymetry and sedimentation, illustrates the complexity of the thermal regime in the outer wedge. While a better fit could be obtained with a coupled advective/conduction model, it would be difficult to evaluate the uniqueness of such a solution given the structure in the region. Heat flow values across the accretionary prism are consistent with thickening of the accretionary wedge and downward advection of heat with the subducting plate. Our models indicate μ*, is approximately 0.06 in the region of observed SSEs and increases to 0.18 approximately 50 km landward of the deformation front. The transition in μ* may be marking the downdip edge of overpressures along the subduction thrust, consistent with the hypothesis that slow slip is enabled by overpressure. SUPPORTING INFORMATION Figure S1. Water column temperatures. (a) Circles show mean bottom water temperatures between May 2014 and June 2015 from HOBBITS experiment (Wallace et al. 2016). Red lines show XBT casts made during cruise RR1508, 2015 May 20 to June 15. Black lines show CTD casts from NOAA database (Levitus et al. 2013). CTD casts were made in 1986 and 2012. (b) Standard deviation of bottom water temperature measurements as a function of depth. Below 1500 m the standard deviation is less than 0.1 °C. Figure S2. Sediment accumulation rates based on cores from the Hikurangi trough and plateau. Orange triangles are core locations and red circles are heat flow measurement locations. Triangle size is proportional to the estimated sediment accumulation rate. Data from IODP Sites 1124 and 1125 are indicated by site number (Shipboard Scientific Party, 1999a,b). Other data from Lewis (1973), Carter et al. (1995), Collot et al. (2001) and Pouderoux et al. (2012). Table S1. Sites used for bottom water temperature analysis. Table S2. Thermal parameters used with or sedimentation correction. Table S3. Results of sedimentation analysis. Please note: Oxford University Press are not responsible for the content or functionality of any supporting materials supplied by the authors. Any queries (other than missing material) should be directed to the corresponding author for the article. ACKNOWLEDGEMENTS We thank the officers, crew, and technicians of the R/V Roger Revelle and R/V Tangaroa for their assistance and advice during these cruises. We thank Justin Ball for assistance with recovery of the bottom water thermistors. We thank editor D. Blackman, an anonymous reviewer and A. Fagereng for constructive comments that helped improved this paper. We also thank H. Rabinowitz for discussions that helped improved our thinking about friction. The work was supported by NSF grants OCE-1355878 (RNH and AMT), OCE-1355870 (RL), by public good research funding from the Government of New Zealand to GNS Science (SAH and DB), and by the Ministry of Business, Innovation and Employment contract C05 × 0908 (IAP and ARG). Seismic data collected during TAN1114 was partially supported by the Land Information New Zealand's OS2020 program. Voyage TAN1607 was supported by the New Zealand Ministry of Business, Innovation and Employment funding to NIWA for RV Tangaroa. The 05CM data set was acquired by Ministry of Economic Development and was processed by Fugro Seismic Imaging and GNS Science. Heat flow and seismic data collected during cruise RR1508 are available at http://www.marine-geo.org/tools/new_search/index.php?&output_info_all=on&entry_id=RR1508. REFERENCES Ansell J. , Bannister S. , 1996 . Shallow morphology of the subducted Pacific Plate along the Hikurangi margin, New Zealand , Phys. Earth planet. Inter. , 93 , 3 – 20 . https://doi.org/10.1016/0031-9201(95)03085-9 Google Scholar Crossref Search ADS Ballance P.F. , 1976 . Evolution of the Upper Cenozoic magmatic arc and plate boundary in northern New Zealand , Earth planet. Sci. 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TI - Thermal regime of the Northern Hikurangi margin, New Zealand
JF - Geophysical Journal International
DO - 10.1093/gji/ggy450
DA - 2019-02-01
UR - https://www.deepdyve.com/lp/oxford-university-press/thermal-regime-of-the-northern-hikurangi-margin-new-zealand-fuilWIuxoB
SP - 1177
VL - 216
IS - 2
DP - DeepDyve
ER -