JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 117, B11402, doi: /2012jb009489, 2012

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1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 117,, doi: /2012jb009489, 2012 Simultaneous long-term and short-term slow slip events at the Hikurangi subduction margin, New Zealand: Implications for processes that control slow slip event occurrence, duration, and migration Laura M. Wallace, 1,2 John Beavan, 1 Stephen Bannister, 1 and Charles Williams 1 Received 29 May 2012; revised 12 September 2012; accepted 20 September 2012; published 2 November [1] We document a sequence of simultaneous short-term and long-term slow slip events (SSEs) at the Hikurangi subduction zone during the 2010/2011 period. The sequence of short-term events (each 2 3 weeks in duration) ruptured much of the shallow plate interface (<15 km) at central and northern Hikurangi over a 6-month period, was accompanied by microseismicity and involved patchy, irregular migration of SSE slip. We suggest that the patchy migration of the short-term SSE is due to large-scale ( km 2 ) heterogeneities on the plate interface related to seamount subduction and sediment subduction and/or underplating. This is in contrast to a 2010/2011 long-term SSE at the central Hikurangi margin, which evolved steadily over 1.5 years and ruptured much of the plate interface between 20 and 70 km depth. We suggest that the occurrence of long-term versus short-term SSEs at Hikurangi is related to differences in effective normal stresses and relative heterogeneity of the subduction interface. The long-term SSE sequence began 1 year before the short-term sequence. Coulomb stress change models suggest that the long-term SSE may have triggered initiation of the subsequent short-term SSE sequence. Initiation of the short-term sequence occurred in a region just updip of or within an interseismically locked portion of the plate interface and may be located within the updip transition from seismic to aseismic behavior. Alternatively, it could be characteristic of a region undergoing partial interseismic coupling. This is in contrast to SSEs observed elsewhere in the world that typically occur within the downdip transition from seismic to aseismic behavior. Citation: Wallace, L. M., J. Beavan, S. Bannister, and C. Williams (2012), Simultaneous long-term and short-term slow slip events at the Hikurangi subduction margin, New Zealand: Implications for processes that control slow slip event occurrence, duration, and migration, J. Geophys. Res., 117,, doi: /2012jb Introduction [2] We now understand that subduction interface faults exhibit a great variety of slip behaviors, spanning from stickslip behavior (e.g., earthquakes) through to steady aseismic creep. The recent confirmation of episodic slow slip events (SSEs) as a new class of shear slip at subduction margins is widely acknowledged as one of the most exciting discoveries of the last decade in the Earth Sciences [e.g., Schwartz and Rokosky, 2007; Rubinstein et al., 2010]. Although the location of most subduction interface SSEs worldwide at the down-dip transition from stick-slip (velocity weakening) to 1 GNS Science, Lower Hutt, New Zealand. 2 Now at Institute for Geophysics, University of Texas at Austin, Austin, Texas, USA. Corresponding author: L. M. Wallace, Institute for Geophysics, University of Texas at Austin, Burnet Rd., Austin, TX 6009, USA. (lwallace@ig.utexas.edu) American Geophysical Union. All Rights Reserved /12/2012JB aseismic creep (velocity strengthening) suggest that they occur within a transitional, conditionally stable frictional regime [Dragert et al., 2001; Larson et al., 2004; Ohta et al., 2004, 2006], the physical processes governing SSEs are not well known. From an earthquake hazards standpoint, SSEs are important to monitor and understand, particularly given the possibility that an SSE could trigger a damaging subduction interface earthquake within the nearby velocity weakening portion of the plate interface [e.g., Ito et al., 2012; Kato et al., 2012]. [3] Numerical models incorporating rate and state friction laws that successfully reproduce SSE behavior require that the dimensionless a and b values (which describe ratedependent variations in friction) have values where b is only slightly larger than a (e.g., slightly rate weakening) [e.g., Liu and Rice, 2005, 2007]. In other words, the frictional properties of rocks that host SSEs are conditionally stable, straddling the threshold between velocity/rate strengthening (aseismic behavior, a > b) and velocity/rate weakening behavior (seismic slip behavior, b > a). Models based on rate and state friction also require extremely low effective stress 1of18

2 Figure 1. Tectonic setting, interseismic coupling (in terms of coupling coefficient, see red to blue scale), and cumulative slip in SSEs since 2002 (green contours, labeled in millimeters). Black arrow labeled PAC/AUS shows the relative convergence vector between the Pacific and Australian Plates. Dashed, black contours show the depth to the subduction interface (in kilometers below sea level). to reproduce episodic SSE behavior, suggesting that high fluid pressures are needed to generate SSEs [Liu and Rice, 2005, 2007]. Other theories for SSE occurrence include dilatant strengthening [Segall et al., 2010], rate-dependent changes in velocity weakening versus velocity strengthening behavior [Shibazaki and Shimamoto, 2007], among others (see overview in Rubin [2011]). These theories for SSE occurrence also require very low effective stress (e.g., near lithostatic pore fluid pressures) in the SSE source area. Interpretations of seismic attributes of the interface in the SSE source area in Japan, Cascadia, Mexico, and the northern Hikurangi margin also suggest an abundance of fluids near the interface [Kodaira et al., 2004; Audet et al., 2009; Song et al., 2009; Bell et al., 2010]. [4] Slow slip events and related seismic behavior (tremor, low frequency, and very low frequency earthquakes) span a spectrum of duration and magnitude characteristics [Ide et al., 2007; Peng and Gomberg, 2010]. Slow slip is also observed to migrate great distances along strike (up to 300 km) at some subduction zones (Schmidt and Gao [2010] and Dragert and Wang [2011], among others). In some subduction zones, along-strike migration of SSEs is highly irregular (e.g., this paper), while in most other cases it appears to occur more steadily (Cascadia [Dragert and Wang, 2011]; Mexico [Franco et al., 2005]). Understanding why these differences in SSE duration and propagation occur from one subduction zone to another may reveal some of the physical conditions that control the large variety of source characteristics that we observe for slow slip events. [5] The North Island of New Zealand overlies the Hikurangi subduction thrust, which accommodates westward subduction of the Pacific Plate beneath the eastern North Island (Figure 1). Owing to the ongoing subduction of the Cretaceous Hikurangi Plateau (a large igneous province, similar to Ontong Java), much of the forearc region of the Hikurangi margin is subaerial. Importantly, the subduction interface lies km beneath most of the east coast of the North Island and deepens westward [Ansell and Bannister, 1996, Figure 1]. This is in contrast to most other well-studied subduction margins such as the Nankai Trough and Cascadia, where the shallow (<10 15 km depth) portion of the interface is more than 50 km offshore. The North Island s position above a large depth range of the subduction interface (Figure 1), combined with the comprehensive seismic and geodetic network (Figure 2) that exists there ( makes the Hikurangi margin uniquely well-suited for study of plate boundary slip processes spanning from the shallow (<15 km) to the deeper (50 70 km) subduction interface. [6] Since 2002, more than 15 slow slip events have been documented by a continuous Global Positioning System (cgps) network at the Hikurangi subduction margin [Wallace and Beavan, 2010]. In most cases to date, these SSEs have occurred at the down-dip transition from strong interseismic locking to steady aseismic creep [Wallace and Beavan, 2010; Wallace et al., 2004, Figure 1], similar to SSEs documented at subduction margins elsewhere [e.g., Schwartz and Rokosky, 2007]. A distinctive bimodal behavior is observed in Hikurangi margin slow slip events. Deep (25 60 km) slow slip is observed adjacent to the deeply locked portion of the Hikurangi subduction thrust; these deep SSEs last years, release moment equivalent to M w 7.0, and occur approximately every 5 years. In contrast, shallow slow slip events (<5 15 km) are observed at the northern and central Hikurangi margin, where much of the subduction interface is dominated by aseismic creep; these events last 1 3 weeks, release moment equivalent to a M w , and occur relatively frequently (every 1 2 years at north Hikurangi) [Wallace and Beavan, 2010]. [7] Here, we analyze cgps data in the North Island during the 2010/2011 period to document the slip distribution and evolution of a long-term, deep slow slip event and a sequence of shorter-term propagating SSEs that ruptured much of the shallow portion of the Hikurangi subduction interface. We suggest that the long-term SSE (lasting from mid-2010 until September 2011) triggered the short-term SSE sequence which began in June 2011, just updip of the area of long-term, deep slow slip. Using the cgps displacements and microseismicity that accompanied the short-term SSE sequence, we demonstrate that the along-strike migration of the short-term SSE occurs in patches, and that this may be related to largescale ( km 2 ) heterogeneities in plate interface structure and properties [e.g., Barker et al., 2009; Bell et al.,2010]. We show evidence for the initiation of the East Coast SSE sequence just updip of the interseismically locked portion of the subduction interface. This is in distinct contrast to most SSEs observed worldwide, which are typically observed at the downdip transition from interseismic coupling to aseismic creep. We also discuss the implications that the bimodal behavior of deep versus shallow SSEs have for the physical properties within the SSE source areas at the Hikurangi margin 2of18

3 Figure 2. Map showing current (as of November 2011) configuration of the continuous GPS (cgps) network in New Zealand ( Black triangles denote a cgps site, and site names for most stations are denoted by a four-letter code. and assess SSE duration and propagation characteristics of Hikurangi SSEs in the context of other SSEs worldwide. 2. Continuous GPS Data Analysis [8] The cgps and seismic network in New Zealand is installed and operated by the New Zealand GeoNet project (Figure 2). Further information on the network and the basic data are available at GPS data processing methods are outlined on the GeoNet Web site and described more fully in Wallace and Beavan [2010]. We analyze the cgps data using Bernese v5.0 [Dach et al., 2007] software holding IGS orbits and Earth orientation parameters fixed. The processing has been updated from Wallace and Beavan [2010] in that we now use IGS08 orbits (or IG1 reprocessed orbits and IGS05 final orbits prior to the start of IGS08), and IGS08 antenna and satellite phase patterns (see for further information on these data sources). The resulting daily coordinate time series are regionally filtered [Wdowinski et al., 1997; Beavan, 2005] using a set of NZ stations that have close to linear behavior over the time period. Regionally filtered time series for each site are available to view at www. geonet.org.nz/resources/gps/timeseries, though these are not identical to the time series used in this analysis. The approach we use to estimate the displacements in the time series due to the SSEs is described in Wallace and Beavan [2010]. In brief, the three coordinate time series for each station are examined by eye to select a set of time periods when no slow slip is affecting that station (see Figure 3). A linear regression line is fit to each of these sections and the weighted mean slope of the lines is defined as the inter-sse velocity. This velocity is subtracted from the time series, which are then examined by eye to identify the start and end times of slow slip events. A semiautomated algorithm smooths the time series and picks their values at the start and end of each event for each station affected by that event. For some events we split the event into sections based on changes in the character of the time series and pick the amplitudes at these times. This enables us to roughly follow the time evolution of the event. [9] We estimate slip on the subduction interface during the SSEs using a program called DEFNODE [McCaffrey, 1995, 2002] to invert horizontal and vertical GPS site displacements for slip on the subduction interface. We assume 3of18

4 Figure 3. Continuous GPS time series from sites showing representative examples of displacement in the Manawatu 2010/2011 SSE. The red, blue, and green traces are the east, north, and up smoothed time series, respectively. Behind each trace in grey is the original regionally filtered time series. The sections with a light blue background are those used for estimating the inter-sse velocity that has been subtracted from the displayed time series (and in most cases, such sections also occur prior to the start of the plotted data). The sections with a light green background show SSEs for which an offset has been estimated. The small circles show the amplitude picks at the start and end of each SSE subepisode. The vertical scales are different on each plot as we wish to utilize the maximum possible resolution. Locations of the sites are shown on Figure 2. that surface displacements are due to dislocations in an elastic half space, using the equations of Okada [1992]. Rather than explicitly solving for the amount of slip at each of the nodes or patches defining the interface, we invert for three parameters describing a Gaussian function along separate down-dip profiles to estimate the depth and magnitude of the maximum slip and the spread of the slip [e.g., Subarya et al., 2006]. The slip on patches between the down-dip profiles is interpolated linearly. This approach captures the details of the event using fewer free parameters per event than if we solve for slip at individual nodes or fault patches. One drawback to this approach is evident when there is more than one distinct patch of slip occurring along the same down-dip profile. However, the GPS displacements in the SSEs indicate that this is not usually the case. Alternative parameterizations of the slip could also be used [e.g., McCaffrey, 2009], but we find that the Gaussian parameterization works well for Hikurangi SSEs [Wallace and Beavan, 2010]. We assume that the slip direction in the SSEs is parallel to the direction of long-term relative motion on the subduction interface estimated from the elastic block model of Wallace et al. [2004]. We use a subduction interface geometry based on Ansell and Bannister [1996]. Further details on the GPS slip inversion approach used in this study can be found in Wallace and Beavan [2010] The 2010/2011 Manawatu SSE [10] Starting in mid-2010, more than a dozen cgps sites began moving eastward relative to their normal sense of motion (and in some cases, upward and northward or 4of18

5 southward) (Figure 3). This slow change in direction of motion continued for a few months until early September 2010, when the rate of eastward displacement increased markedly, particularly at sites in the northern portion of the study area (for example, see sites PNUI, BHST, KERE, GRNG, UTKU), suggesting some northward migration of the slow slip during this time. Interestingly, this onset of more rapid slip coincided with the occurrence of the 4 September 2010, M w 7.1 Darfield earthquake near Christchurch [Gledhill et al., 2011]. The timing of the two events may simply be coincidental, and we also note that the timing of the onset of more rapid SSE slip is difficult to pinpoint precisely during such a long-duration SSE (it is possible that the increased slip rate began as early as late August). However, if there is a temporal relationship between the Manawatu SSE acceleration in early September and the 4 September 2010 earthquake, it is plausible that passing seismic waves from the Darfield earthquake dynamically triggered the onset of the second, more rapid phase of the Manawatu SSE. Tremor episodes can be triggered dynamically by seismic waves [e.g., Peng et al., 2009; Rubinstein et al., 2009; Fry et al., 2011], raising the possibility that SSEs could also be triggered dynamically [e.g., Johnson et al., 2012]. The cgps sites KERE and PNUI saw the largest horizontal (23 and 26 mm, respectively) and vertical (27 and 29 mm, respectively) displacements during this period. This period of most rapid displacement of the cgps sites lasted 3 months until the end of December 2010, when the rate of displacement at the cgps sites decreased again. The decreased rate of displacement continued until late September 2011, with a cumulative duration for the slow slip event of 16 months. However, the decrease in displacement rate at the southern sites was not as marked as at the northern sites during the final stage, indicating possible southward migration of the locus of the SSE slip during this final stage. [11] To assess the slip distribution during these three main stages of the Manawatu SSE, we calculated displacements of all cgps sites in the region for each stage of the event and inverted these for slip on the subduction interface. We find that during the first stage of the event, 3 cm slip is focused on a small patch at 50 km depth on the interface, 120 km due south of Lake Taupo (Figure 4a). During the second, most intense phase of the Manawatu SSE, the slip spreads out toward the north and east, over an approximately 10,000 km 2 area, with slip occurring over a broad depth range of the subduction interface from 60 to 25 km depth (Figure 4b). The maximum slip attained during this phase is 12 cm. This phase of the event affected more than half of the 80 Hikurangi margin cgps network sites, and a number of sites within the central volcanic region network were also displaced substantially; in all, 25 cgps sites were displaced by more than 10 mm, and up to 50 cgps sites were displaced by at least a few millimeters during this period (Figure 4b). Slip during the final (third) stage of the SSE 5of18 Figure 4. Slip distribution on the interface (in millimeters), with modeled (red arrows) and observed (black arrows with 1s error ellipses) horizontal surface displacements during the three main stages of the Manawatu 2010/2011 SSE. See Figure 5 for cumulative slip in the 2010/2011 Manawatu SSE, compared with slip in the 2004/2005 event. Black dots show nodes defining the subduction interface, and the large red arrows in Figure 4b point to the node profiles along which we invert for slip, using a Gaussian parameterization. Subduction interface depth contours for model slip region (black dashed lines) labeled with depth (in kilometers below sea level).

6 Figure 5. Cumulative slip in the 2010/2011 Manawatu SSE, shown by yellow to red colors (see scale) and white contours (millimeters). Slip in the 2004/2005 Manawatu SSE shown with dashed green contours, labeled in millimeters. Black dashed lines show depth of subduction interface (in kilometers below sea level). Black triangles show the locations of cgps sites in the area that were operating during the 2004/2005 Manawatu SSE. appears to focus just to the south of the second stage main slip patch (Figure 4c), and is generally 3 6 cm in the slipping areas. Although the spatial migration of slip in the Manawatu SSE is not as marked as that observed in places like Cascadia, there is significant spatial and temporal evolution of the event, both in the along-strike and along-dip direction. [12] The slip that occurred throughout the entire event (Figure 5) released moment equivalent to M w 7.1. Most (56%) of the moment release occurred during the 3-month long second phase (Figure 4b). A total of 13,400 km 2 of the subduction interface underwent slip in this event, over a region >100 km wide (in the along-dip direction). In the total event, slip ranged from a few cm to more than 15 cm in the main slipping patch (Figure 5) Comparison With the 2004/2005 Manawatu SSE [13] In 2004/2005, a slow slip event occurred in a similar location to the 2010/2011 Manawatu SSE (Figure 5, green dashed contours). The evolution and distribution of slip in this earlier event is not as well constrained as for the 2010/ 2011 SSE, due to fewer cgps sites during that time. The 2004/2005 Manawatu SSE occurred during the early stages of the GeoNet cgps network rollout, and only seven cgps sites capable of detecting Manawatu SSE surface displacement were deployed by that time [Wallace and Beavan, 2006, 2010]. By 2010/2011 there was a tenfold increase in the number of cgps sites in the affected region. Despite the sparse network during the 2004/2005 Manawatu SSE, the slip distribution during the 2004/2005 and 2010/2011 events is remarkably similar, though the 2004/2005 slip distribution is rougher than that estimated for 2010/11. This roughness in 2004/2005 is not well constrained by the sparse cgps network at that time, and it is possible that the actual slip distribution is smoother and more similar to the 2010/2011 SSE than is apparent from Figure 5. The overall moment release in the 2004/2005 and 2010/2011 SSEs is reasonably similar: M w 7.2 [Wallace and Beavan, 2006, 2010] and 7.1, respectively. [14] Between the end of the 2004/2005 SSE (in June 2005 [Wallace and Beavan, 2006, 2010]) and the start of the 2010/2011 SSE in mid-2010, there was a 5-year gap when no SSE slip appeared to occur on this portion of the subduction interface, suggesting a 5-year recurrence interval between Manawatu SSEs at the central Hikurangi margin on this occasion. The long-term relative plate motion on this portion of the subduction interface (i.e., the relative motion between the rotating Hikurangi forearc and the subducting Pacific Plate), as determined from elastic block modeling, is mm/yr [Wallace et al., 2004, 2012], implying an accrual of cm of plate motion during the intervening inter-sse period assuming 100% inter-sse coupling. The GPS inversions suggest 15 cm of slip in the main locus of slow slip during the 2010/2011 SSE. This is generally consistent with a 5-year period between Manawatu SSEs. However, Wallace and Beavan [2010] showed that between SSEs, the portion of the subduction interface that undergoes slip in the Manawatu SSEs is only partially coupled (50 80%), which suggests that some of the plate motion budget on the interface within the SSE source area may be accommodated by steady, aseismic creep. Whether or not the 5-year period between Manawatu SSEs is persistent through time will remain to be seen in future SSEs in this location Geodetic Evidence for the 2011 East Coast SSE Sequence [15] During the last week of June 2011, several cgps sites on the east coast of the North Island located near Castlepoint (Figure 6; see GPS sites CAST, BIRF, and AKTO) began moving eastward by up to 15 mm over a period of 2 weeks. Over the following months, until the end of December 2011, sites further north along the east coast were displaced eastward in discrete episodes usually lasting 1 3 weeks, suggesting a northward migrating SSE sequence along the shallow portion of the Hikurangi subduction thrust. To assess the distribution of slip during each of these episodes, we invert the cgps site displacements for slip on the subduction interface (Figure 7). [16] The first episode of slip occurred from 13 June to 16 July 2011 and appears to have involved a northward migration of slip from offshore Castlepoint into the offshore southern Hawkes Bay region. During this first stage of the SSE, a more detailed evaluation of the cgps timeseries show that cgps sites in the south (CAST, TEMA, TRAV) are displaced eastward approximately a week earlier than those sites further north (BIRF, AKTO, PORA), suggesting a northward propagation of the locus of slip over the 3-week long SSE episode (Figures 6 and 8). We have broken this first episode of SSE slip into two subepisodes (Figures 7a 7b), in order to assess the along-strike migration of the event. During the first 2 weeks of July, there also appears to be some slow slip occurring further north, beneath Mahia Peninsula (see eastward motion at MAHI). The June/July 2011 SSE episode initiated up-dip of the locus of slip in the Manawatu SSE, which was still ongoing at the time. [17] There was a 1-month hiatus in slow slip activity from mid-july to mid-august, and slip occurred again from mid- August until early September (Figures 6 and 7c); displacements at cgps sites including MAHI, CKID, KOKO, and 6of18

7 Figure 6 7of18

8 Figure 7. Slip distribution on the interface (in millimeters), with modeled (red arrows) and observed (black arrows with 1s error ellipses) horizontal surface displacements during the five stages of the 2011 East Coast SSE sequence. Black dots show nodes defining the subduction interface, and the larger red arrows point to the node profiles along which we invert for slip, using a Gaussian parameterization. Subduction interface depth contours for model slip region (black dashed lines) labeled with depth (in kilometers below sea level). PARI suggest slip beneath Mahia Peninsula and Hawke Bay during that time (Figure 7c). Another 2-week hiatus was observed in early September and then slip initiated further to the north offshore Tolaga Bay (see site CNST) from mid- September until early October (Figures 6 and 7d). There was another hiatus in SSE activity for more than 2 months until 12 December, when cgps sites near Gisborne (GISB, PARI, among others) began moving eastward until approximately 23 December (Figure 7e). This final stage of the sequence ruptured the gap in slip between the August/September and Figure 6. Continuous GPS time series from sites on the east coast of the North Island, showing representative examples of displacement during the 2011 East Coast sequence. Also visible in some of the timeseries is an earlier SSE sequence during January April 2010 (especially see Figures 6g 6n [Wallace and Beavan, 2010]). The red, blue, and green traces are the east, north, and up smoothed time series, respectively. Behind each trace in grey is the original regionally filtered time series. The sections with a light blue background are those used for estimating the inter-sse velocity that has been subtracted from the displayed time series (and in most cases such sections also occur prior to the start of the plotted data). The sections with a light green background show SSEs for which an offset has been estimated. The small circles show the amplitude picks at the start and end of each SSE subepisode. The vertical scales are different on each plot as we wish to utilize the maximum possible resolution. Locations of the sites are shown on Figure 2. 8of18

9 geonet.org./nz/resources/earthquake/), detected by the New Zealand GeoNet seismometer network. More than 70 events followed, through to 8 July, including several M L 4 events, and a M L 5.0 (M w 4.8) event on 4 July, 08:36 UTC. The earthquake hypocenters clearly migrate northward over the 2-week period, at a rate of 6 9 km/day, tracking the northward migration of the slow slip event determined from the cgps inversions (Figures 8 9). More detailed examination of the continuous seismological data indicate a total of 120 events, over a 2-week period, with a tight cluster of events occurring between 1 July and 4 July located on the southwest side of the second SSE subepisode (Figure 9). Moment tensor solutions were obtained by GeoNet (www. geonet.org./nz/resources/earthquake/) for the largest two events in this subepisode (Figure 9), using the method developed at the University of California, Berkeley Seismological Laboratory [Dreger and Helmberger, 1993; Dreger, 2003; Ristau, 2008]. These solutions indicate thrust faulting, with a small strike-slip component. [19] The cumulative moment produced by microseismicity during the June/July portion of the East Coast sequence is 1.8e+17 Nm, while the equivalent moment released in the June/July SSE (determined geodetically) is 1.8e+19 Nm. The microseismicity only accounts for 1% of the SSE slip, indicating that the vast majority of the geodetically determined slip is aseismic. Although tremor during the 2011 East Coast sequence has not yet been clearly identified, we cannot rule it out at this stage. The mix of microseismicity and slow slip within the southern portion of the East Coast sequence indicates strong heterogeneity of frictional properties and physical conditions on the subduction interface there. In many ways this coexisting slow slip and microseismicity is analogous to similar processes observed on the Japan Trench in the source region of the March 2011 M w 9.0 earthquake [Ito et al., 2012; Kato et al., 2012]. Figure 8. (a) Time evolution of seismicity as a function of latitude, illustrating an apparent northward migration velocity of 6 9 km/day. (b d) Regionally filtered GPS displacements (millimeters) for the east components of GPS stations CAST, AKTO, and PORA for the same time period. Note that the time series are ordered from south to north (from bottom to top) to more clearly show the northward migration of the event with time. Gray arrows highlight the northward migration with time. September/October portions of the sequence, following a similar spatial and temporal pattern to a previous SSE sequence in that area that occurred in early 2010 [Wallace and Beavan, 2010]. 3. Seismicity Related to the East Coast SSE Sequence [18] In addition to marked cgps displacements onshore, an intriguing cluster of earthquakes took place during the June/July slow slip events. On around 21 June, at the start of the June/July 2011 SSE, a sequence of M L 2-M L 3 earthquakes occurred 30 km offshore of the east coast (www. 4. Rupture of the Shallow Plate Interface at Southern Hikurangi in Previous East Coast Sequences [20] We have also searched the time series of the southeast coastal cgps sites to determine if any events similar to the June/July 2011 event had occurred in the past. The first cgps site (CAST) was installed in this region in 2006, so it is not possible to identify SSEs here prior to We have identified one possible candidate for a smaller SSE in December 2009, when 5 mm east to southeast displacement occurred at sites TEMA, CAST, and TRAV (Figures 6 and 10). Inversion for slip in this late 2009/early 2010 event (assuming a subduction interface source) requires slip on the interface just to the south of the June/July 2011 SSE (Figure 10). Intriguingly, this SSE occurred just offshore, in a region where the plate interface currently appears to be locked (see discussion in section 5.2). [21] Seismicity clusters have also been observed in the past in the same location as for June July 2011, for example in September October 2006, and in September October 2001 (Figure 11). The 2006 cluster coincides with (and is slightly preceded by) an SSE in southern Hawke s Bay that occurred from late August until early September 2006 [Wallace and Beavan, 2010]. There is otherwise very little indication of earthquake clusters offshore this section of the east coast of 9of18

10 Figure 9. The SSE slip distribution (contours labeled in millimeters; see also Figure 7) and seismicity ( associated with the June July 2011 SSE sequence, for the time periods (a) 13 June to 1 July and (b) 1 July to 16 July. Moment tensor solutions ( org./nz/resources/earthquake/) are shown for two of the events (M w 4.8 and M w 4.1), indicating thrust faulting with a small component of strike-slip. Dashed gray contours show depth to the subduction interface (below sea level; after Ansell and Bannister [1996]). Figure 10. (a) Slip during the late 2009/early 2010 SSE (green contours, 5 mm intervals) and the June/ July 2011 SSE (phases 1 and 2 of the East Coast Sequence combined, dashed contours, 20 mm intervals). Black arrows show horizontal displacement of cgps sites during the 2009/2010 SSE and white arrows are the vectors predicted by the best-fitting slip models (shown). The slip contours are overlain on the interseismic coupling coefficient determined from campaign GPS data [Wallace and Beavan, 2010]. Heavy yellow dashed line shows the contact between the Paleocene/Cretaceous backstop and the outer, actively deforming accretionary prism [Barnes et al., 2010] in the region of the 2010 and 2011 SSEs. Dashed black line shows depth contours (labeled) to the subduction interface. (b) Cross-sectional schematic of the slip behavior domains suggested by some previous studies of subduction zone interface behavior. The detailed location of these domains will be different for each subduction zone. This schematic is intended to represent the locations of these domains at a subduction margin analogous to the southern Hikurangi margin. The white dashed lines show the location of frictional transition zones where slow slip behavior might be expected. Lower right panels show the detrended time series (east component only) for cgps sites TEMA, TRAV, and CAST during 2009 and Green transparent box highlights the timing of the late 2009/early 2010 SSE. 10 of 18

11 the North Island in the time period We infer that, as for June July 2011 and September October 2006, the seismicity cluster in 2001 could also have been associated with SSE deformation, before the onshore continuous GPS network was fully established. If so, this suggests an approximately 5-year recurrence interval between SSEs (and associated microseismicity) in the region of the June July 2011 SSE in the southern Hawke s Bay region. Note that the December 2009 SSE further south (Figure 10) does not appear to have associated microseismicity (Figure 11). Figure 11. (top) Time series of seismicity in the vicinity of the 2011 July SSE, as a function of time and latitude (first panel). Transparent, purple vertical bands indicate the time and latitude range of known SSE episodes in September 2006 and in June July The second, third, and fourth panels show the same seismicity in map view, for the most highly active periods (time frames labeled on maps). The seismicity suggests possible earlier SSE episodes in 2001 and mid-2002, before the GPS network was installed. 5. Discussion 5.1. Triggering of the East Coast Sequence by the Manawatu SSE? [22] Numerous studies have noted the relationship between an increase in static Coulomb stress change following an earthquake and the distribution of aftershocks, suggesting that some aftershocks are triggered by these static stress changes (Stein and Lisowski [1983], Toda et al. [1998], and Anderson and Johnson [1999], among many others). Moreover, many studies have documented dynamic triggering of tremor episodes by passing surface waves (Peng et al. [2009], Rubinstein et al. [2009], and Fry et al. [2011], among others). Additional studies have suggested triggering of SSEs by earthquakes due to either static [e.g., Brooks et al., 2008] or dynamic [e.g., Johnson et al., 2012] stress changes. The 2011 East Coast SSE sequence initiated in June/July, just updip of the region of ongoing slip in the Manawatu SSE (Figure 12). The spatial and temporal relationship between the two SSEs suggests the possibility that the redistribution of plate boundary stresses that occurred during the Manawatu SSE actually triggered the initiation of the East Coast sequence, toward the end of the 2010/2011 Manawatu SSE. To test this idea, we calculate the change in Coulomb failure stress (CFS) on the subduction interface in the region of the Manawatu SSE. [23] To perform our analysis, we used the finite element code PyLith [Williams et al., 2005; Williams, 2006; Aagaard et al., 2007, 2008], which can compute stresses on the fault surface for a given slip distribution. We apply the inferred slip distribution for the 2010/2011 Manawatu SSE and then compute the CFS changes on the subduction interface, assuming a shear direction consistent with the plate convergence direction [Wallace and Beavan, 2010]. The results of our Coulomb stress calculations are shown in Figure 13. For our calculations we used a relatively low effective friction coefficient of 0.2, since we assume that pore fluid pressures are relatively high; however, the results do not change significantly when we assume a value of 0.4. Also, since CFS calculations are strongly dependent on small variations in the applied slip, we use a simplified version of the geodetically inferred slip distribution where all slip values less than 1 cm are set to zero. This allows us to focus on the broader-scale features of the stress change patterns. [24] The overall effect of the 2010/2011 Manawatu SSE is to produce CFS increases on the order of 1 kpa near the maximum slip regions of the June/July 2011 phase of the East Coast sequence and on the order of 10 kpa near the downdip end of the June/July 2011 slip area (Figure 13). Since the computed stress is a function of the derivative of the applied slip, the CFS calculations are strongly dependent 11 of 18

12 Figure 12. Cumulative slip on the interface in the East Coast 2011 Sequence and the Manawatu 2010/2011 sequence (yellow to brown colors, see scale, and white contours, labeled in millimeters). Green dashed lines bound the area where most of the subduction interface (between 70 km and <10 km depth) undergoes slow slip behavior. on the details of the assumed fault slip. This is reflected in Figure 13 as small-scale fluctuations between positive and negative values in CFS due to slight gradients in the applied slip. Small-scale slip gradients are not well resolved by the cgps data, so these small-scale variations in CFS are not necessarily real. Despite these uncertainties, there is a clear region of increased CFS surrounding the location of the 2010/2011 Manawatu SSE, and this region of increased CFS coincides with the initiation location of the East Coast sequence near Cape Turnagain. [25] Although the static stress change from the Manawatu SSE within the source area of the June/July 2011 episode is small (1 10 kpa), it is comparable to dynamic stress changes (tens of kpa, due to passing surface waves) associated with many cases of dynamically triggered tremor [e.g., Peng et al., 2009; Rubinstein et al., 2009; Fry et al., 2011]. For tremor to be triggered by such small stresses, the tremor source areas are thought to be very close to failure, possibly due to fluid overpressure [Miyazawa and Mori, 2005; Peng and Chao, 2008]. Numerous modeling and observational studies also suggest that fluid pressures are very high in SSE source areas [Liu and Rice, 2005, 2007; Kodaira et al., 2004; Audet et al., 2009; Song et al., 2009]. Thus it seems plausible that small stress increases on the plate interface, such as that induced by the Manawatu 2010/2011 SSE, could have triggered the onset of the 2011 East Coast SSE sequence. [26] A similar spatiotemporal relationship between the occurrences of long- and short-term SSEs is observed in southwest Japan. There, the recurrence interval of short-term ETS events beneath Shikoku Island becomes even shorter during the nearby long-term SSEs beneath Bungo Channel [Hirose and Obara, 2005], suggesting that the timing of short-term SSEs can be influenced by stress loading from the adjacent (larger) long-term SSEs. Numerical models assuming rate and state friction also suggest that long-term SSEs can lead to a shorter interevent time for nearby shortterm SSEs [Matsuzawa et al., 2010] Initiation of the East Coast Sequence Within the Previously Locked Plate Interface [27] The earliest stages of the 2011 East Coast SSE appears to have begun within the shallow portion of the interseismically locked subduction interface at the southern Hikurangi margin (Figure 10). Given the widespread correlation between well-documented SSEs and the downdip transition zone from stick-slip to aseismic creep behavior, we are surprised to observe SSE behavior within the previously defined strongly locked portion of the Hikurangi subduction interface. We suggest that this SSE is either (1) occurring in a zone of transitional frictional behavior at the updip limit of the seismogenic zone or (2) is occurring in a region of partial interseismic coupling characterized by interseismically locked asperities surrounded by slowly slipping areas. We evaluate these two possibilities and discuss their implications below. [28] Nearly all SSEs observed worldwide are documented to occur at the downdip transition from seismic to aseismic behavior (see review in Schwartz and Rokosky [2007]), but the June/July portion of the 2011 East Coast Sequence is an exception. We suggest that this event may instead be located within the updip transition from seismic to aseismic behavior (Figure 10). The occurrence of episodic slow slip updip of the seismogenic zone has been suggested in Costa Rica based on evidence for hydraulic transients correlated with seismic tremor and slow slip [Brown et al., 2005; LaBonte et al., Figure 13. Computed Coulomb failure stress (CFS) change from the Manawatu 2010/2011 SSE where the geodetically inferred slip has been set to zero for predicted slip values less than 1 cm. Computed CFS changes are shown in red/blue colors, and the contour corresponding to a positive value of 10 kpa is shown in white. Also shown are the East Coast SSE slip distributions in 1 cm contour intervals for 13 June to 1 July 2011 (blue contours) and 1 July to 16 July 2011 (green contours) SSE phases. Stresses near the downdip ends of the June/July 2011 slip distributions correspond to positive CFS changes on the order of 10 kpa. 12 of 18

13 2009; Davis et al., 2011]. Outerbridge et al. [2010] also suggest SSE slip updip of the seismogenic zone from shorebased cgps data, although the cgps network in their study is comparatively sparse, and we contend that Outerbridge et al. s [2010] models for SSE slip updip of the seismogenic zone are not well constrained. The identification of very low frequency earthquakes (like SSEs, these are very low stress drop events) in the accretionary wedge offshore southwest Japan, [Ito and Obara, 2006] that coincide with pore pressure transients in borehole observatories within the Nankai Accretionary wedge [Davis et al., 2006] also raises the possibility that behavior similar to slow slip might occur updip of the interseismically coupled, seismogenic portion of the subduction thrust at the Nankai Trough. However, there is no unambiguous surface deformation evidence in either the Nankai Trough or Costa Rica cases to confirm that slow slip occurred on the subduction interface updip of the seismogenic zone. Our confirmation of a slow slip event located updip of or within the interseismically coupled (stick-slip) portion of the southern Hikurangi subduction interface, makes the occurrence of SSEs on the subduction interface updip of the seismogenic zone in Costa Rica, southwest Japan, and elsewhere seem more likely. [29] A number of mechanisms have been suggested to control the location of the updip seismic to aseismic transition, including mineral phase transformations and metamorphic reactions [Vrolijk, 1990; Moore and Saffer, 2001], temperature [Oleskevich et al., 1999], fluid pressure [Moore and Saffer, 2001], fault gouge lithification [Marone and Saffer, 2007], and the transition from the active accretionary prism (where sediments are weak and less consolidated) to a more rigid backstop [Byrne et al., 1988]. Thermal modeling of heat flow data at southern Hikurangi [McCaffrey et al., 2008] suggests that temperatures on the interface in the region of the East Coast SSEs range from 50 to 150 C, indicating that a thermal control on the updip seismic to aseismic transition is possible. However, acquisition of additional heat flow data in this part of the Hikurangi margin and more sophisticated thermal modeling are required to test this. Following Byrne et al. [1988], an obvious potential control on the updip seismic to aseismic transition at southern Hikurangi is the location of the boundary between the active accretionary prism and the Paleogene/Cretaceous backstop [Barnes et al., 2010], the surface expression of which occurs at the updip (seaward) end of where our inversions place the 2011 East Coast SSE (Figure 10). If we assume that the wedge/backstop boundary dips 45, the transition from the wedge to the backstop on the interface lies within the slow slip source area. [30] We also note that the 2011 East Coast sequence initiated along a portion of the plate interface that is currently undergoing partial (75%) interseismic coupling [Wallace et al., 2004; Wallace and Beavan, 2010]. Thus it is equally possible that rather than being interpreted as an event within the updip seismic to aseismic transition (Figure 10, inset), slow slip on the southern segment of the East Coast sequence may instead be representative of a region with highly heterogeneous plate interface properties; in this case, adjacent to the along-strike transition from deep interseismic locking (southern Hikurangi) to an aseismic creep dominated plate interface (central and north Hikurangi) (Figure 10). We can imagine such a region is composed of a mosaic of asperities that are strongly velocity weakening (i.e., slip in earthquakes), surrounded by regions of the interface with more transitional frictional properties that undergo episodic slow slip. [31] Only a small component of the total plate motion budget (>30 mm/yr) required to be accommodated at the southern Hikurangi interface occurred during the 2011 June/ July SSE and December 2009 SSE. These are the only two SSEs observed to occur at the southern Hikurangi margin since cgps coverage became widespread here in Over the period of campaign GPS data acquisition (approximately 15 years, since the early to mid 1990s), this portion of the interface has been accumulating a slip deficit at a rate of mm/yr [Wallace et al., 2004; Wallace and Beavan, 2010], which is 75 99% of the long-term plate motion there. Of the total slip deficit accrued in the last 15 years ( cm), less than 20% was relieved by slip in the SSEs. Thus a large slip deficit remains offshore the southeast Hikurangi margin that we expect will be recovered in a future major subduction thrust earthquake, perhaps similar to the recent Japan Trench M w 9.0 earthquake where the largest slip along the plate interface occurred near the trench [Ito et al., 2011]. It is certainly possible that SSEs occurred at the shallow, southern Hikurangi interface between the early 1990s and 2006; if so, this would help explain why some of the slip deficit rates are less than the long-term plate motion rates (i.e., coupling coefficients less than 1.0) Irregular SSE Propagation and Heterogeneity of the Shallow Subduction Interface [32] The 2011 East Coast Hikurangi sequence migrated along-strike in a highly irregular and patchy manner that included short, 1 3 week bursts of intense SSE behavior punctuated by weeks of quiescence, over a total period of 6 months. However, we note that within individual SSE episodes of the East Coast sequence, SSE migration can occur at rates of 5 9 km/day (see the June/July 2011 SSE; Figures 7a, 7b, and 8). The shallow subduction interface in the region of the East Coast sequence is heterogeneous, and seismic reflection data show that it is impacted by subducted seamounts, and intervening regions of subducting and/or underplated sediments [e.g., Barker et al., 2009; Barnes et al., 2010; Bell et al., 2010;Pedley et al., 2010].Bell et al. [2010] also note a correlation between some SSE locations at the northern Hikurangi margin and zones of high-amplitude reflectivity near the plate interface, which they interpret to be fluid-rich, subducted sediment (Figure 14). It is also likely that subduction interfaces impacted by seamount subduction will be highly fractured and possess a heterogeneous stress field, as a consequence of the seamount subduction process [Wang and Bilek, 2011]. Shallow slow slip at the Hikurangi margin is often accompanied by microseismicity (Delahaye et al. [2009] and this paper), suggesting that this portion of the interface is composed of a complex matrix of velocity weakening and conditionally stable patches. We suggest that the strong spatial variability of frictional properties and fluid pressure regime along the shallow subduction interface in the region of the East Coast sequence is responsible for the observed irregular and patchy migration of slow slip. [33] The scale of the heterogeneities (seamounts and inferred subducted sediments) imaged on the plate interface at north and central Hikurangi is generally km (along- 13 of 18

14 Figure 14. Interface seismic characteristics (high-amplitude reflectivity = HRZ, blue shaded; low amplitude reflectivity = LRZ, pink shaded; subducted seamounts = S, brown shaded) at the northern Hikurangi margin from Bell et al. [2010], and slip on the subduction interface (showing as the 20 mm slip contour only) during portions of the 2011 East Coast sequence (see also Figure 7). strike) by km (down-dip) [Bell et al., 2010, Figure 14]. This is similar in scale to the patches that slipped in each of the subevents during the East Coast sequence (Figures 7 and 14). Portions of the interface where subducted sediment is interpreted to accumulate (e.g., the high-amplitude reflectivity zones (HRZ)) are thought to contain higher fluid pressures compared to the intervening regions of lower amplitude reflectivity and/or seamount subduction. If the HRZ regions are sites of elevated fluid pressure, the interface there will slip more readily, due to the reduced effective normal stress acting on the interface [e.g., Scholz, 1998]. Thus we expect that the patches of the interface where fluid pressures are highest will be the first to undergo slip and that slip could be delayed within the intervening, lower fluid pressure portions of the interface. The irregular and patchy slip observed in the East Coast sequence also suggests that portions of the interface can form barriers to slip, causing slow slip to arrest, and preventing steady, uniform along-strike propagation of the SSE. Similar to what is thought to occur during the arrest of seismic slip [cf. Kaneko et al., 2010], these barriers may be regions of greater friction on the interface compared to surrounding areas, or conversely, they could be velocity strengthening regions where steady, aseismic creep is favored over slow slip (perhaps due to even higher fluid pressures and/or lower friction relative to surrounding areas). [34] SSEs in Cascadia often propagate great distances alongstrike (up to 300 km), over time periods of around 1 month. Although the geodetically resolved migration rates during an individual SSE in Cascadia can vary (e.g., 2 15 km/day), SSE migration often appears to occur steadily throughout the event [Dragert and Wang, 2011], with some exceptions [Schmidt and Gao, 2010]. The comparatively steadier migration of Cascadia SSEs contrasts with the highly irregular and patchy evolution of the East Coast Hikurangi 2011 SSE sequence. Despite this steady migration, there are some instances of gapfilling behavior (similar to Hikurangi s East Coast sequence) during Cascadia SSEs [Kao et al., 2007]. A sequence of Cascadia SSEs occurring over a period of 6 months in 2004 is a good example of this somewhat irregular SSE migration and gap-filling behavior [e.g., Schmidt and Gao, 2010]. Previous studies have also noted jumping and halting migration of tremor patterns during SSEs in Cascadia [Kao et al., 2009; Boyarko and Brudzinski, 2010], which could be considered analogous to the highly irregular SSE migration patterns that we observe in the 2011 East Coast sequence. [35] Overall, however, geodetically detected slow slip in Cascadia does seem to migrate more steadily over long distances compared with the notably irregular and patchy migration patterns observed in East Coast Hikurangi SSEs. SSEs in Cascadia occur at km depth, along the down-dip transition from interseismic coupling to aseismic creep (e.g., McCaffrey [2009], among others). The distribution of interseismic coupling in Cascadia is comparatively uniform in the along-strike direction and is thought to be determined by the thermally controlled brittle-ductile transition [Hyndman and Wang, 1995; Wang et al., 2003]. This is in contrast to the more heterogeneously coupled Hikurangi margin, for which the distribution of interseismic coupling and slow slip cannot be explained by thermal models [McCaffrey et al., 2008, Figure 1]. The subducting Juan De Fuca plate at Cascadia is uniformly young (4 8 Myrs) and covered in a thick (2 3 km) blanket of sediment [Flueh et al., 1998] which helps to shield the interface from basement irregularities, such as seamounts, on the subducting plate. Thus the properties of the subduction interface within the SSE source region in Cascadia are likely to be more homogeneous compared to northern and central Hikurangi, where there is stronger variability in properties and morphology of the incoming Pacific plate (see review in Wallace et al. [2009]). However, we do note that spatial and temporal variability in nonvolcanic tremor behavior at Cascadia [Wech and Creager, 2011; Ghosh et al., 2010] suggest that smallscale heterogeneities (less than a few km 2 ) do exist on the interface, and some large-scale along-strike variations in recurrence behavior of tremor and slow slip have been ascribed to along-strike variations of geological terranes in the upper plate at Cascadia [Brudzinski and Allen, 2007]. It may be that the scale of heterogeneities on the subduction interface at northern and central Hikurangi ( km 2 ) are of the optimal size to produce strongly nonsteady, patchy SSE migration. Moreover, we expect the frictional properties within the source area of deeper SSEs (such as those at Cascadia and in the Kapiti and Manawatu regions of the Hikurangi margin) to be more homogeneous (or be the site of smaller scale heterogeneities) compared to shallow SSEs, as the subduction process will serve to homogenize the rock at the interface as the material is exposed to progressively higher temperatures and extremely large shear strains. 14 of 18