PUBLICATIONS. Geochemistry, Geophysics, Geosystems. Background and delayed-triggered swarms in the central Southern Alps, South Island, New Zealand

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1 PUBLICATIONS Geochemistry, Geophysics, Geosystems RESEARCH ARTICLE 1.1/13GC5171 Key Points: Remote triggering of earthquake swarms has been observed twice within two years Delayed-triggered swarms may be clock-advanced background swarms We infer fluid-diffusion as the preferred cause for the delay in triggering Supporting Information: Readme file with information to Supplementary Material Correspondence to: C. M. Boese, Citation: Boese, C. M., K. M. Jacobs, E. G. C. Smith, T. A. Stern, and J. Townend (14), Background and delayedtriggered swarms in the central Southern Alps, South Island, New Zealand, Geochem. Geophys. Geosyst., 15, , doi:1.1/ 13GC5171. Background and delayed-triggered swarms in the central Southern Alps, South Island, New Zealand C. M. Boese 1, K. M. Jacobs 1, E. G. C. Smith 1, T. A. Stern 1, and J. Townend 1 1 Institute of Geophysics, School of Geography, Environment and Earth Science, Victoria University of Wellington, Wellington, New Zealand Abstract Low-magnitude earthquake swarms (M L.8), consisting of up to 47 events of similar waveforms, have been observed repeatedly in the central Southern Alps, a rapidly uplifting orogen bounded by the transpressive Alpine Fault in the South Island of New Zealand. We compare nine background swarms recorded between November 8 and April 1 with five delayed-triggered swarms that occurred after the M W 7.8 Dusky Sound and the M W 7.1 Darfield (Canterbury) earthquakes. The two types of swarms are similar in terms of the magnitudes, depths, focal mechanisms, and interevent times of the constituent microearthquakes, and appear to both involve the rupture of steeply dipping faults in highly fractured crust in a 1 km 3 1 km area in the center of the SAMBA network. The delayed-triggered swarms occurred at similar epicentral distances (c the rupture length of the mainshocks) to the Dusky Sound and Darfield earthquakes, commenced shortly after the passage of the surface waves, continued for 5 and days, respectively, and were followed in each case by a day long quiescent period, which may suggest clockadvanced of faults in their failure-cycle. Triggering thresholds of.1 MPa proposed elsewhere are similar to the dynamic stress changes computed for the Southern Alps (.9 MPa). However, as 98% of the locatable triggered events occurred several hours after the surface waves had passed, the dynamic stress changes associated with the surface waves themselves are unlikely to have triggered the earthquakes directly. Instead, we suggest that the locations and delays of the triggered swarms are more consistent with triggering by pore pressure diffusion. Received 6 NOV 13 Accepted 5 FEB 14 Accepted article online 9 FEB 14 Published online APR Introduction 1.1. Geological Setting The Southern Alps in the central South Island, New Zealand, are the surface manifestation of continental collision between the Australian and Pacific plates at depth. The plate boundary, the 46 km long, transpressive Alpine Fault, bounds the Southern Alps to the west. The central region between Whataroa and Fox Glacier (Figure 1) has the highest topography and rock uplift rates of the Southern Alps [Beavan et al., 4; Beavan et al., 1a; Norris and Cooper, 1] and is characterized by abundant shallow microseismicity [Leitner et al., 1; Boese et al., 1]. In November 8, the Southern Alps Microearthquake Borehole Array (SAMBA) [Boese et al., 1] was installed to densify coverage in the region monitored by three stations of the national network (GeoNet), with a 1 km station spacing. To date, SAMBA has been recording continuously and was temporarily supplemented between January and April 1 by the DFDP1 network as part of the Deep Fault Drilling Project [Sutherland et al., 1; Townend et al., 13]. These networks have lowered the magnitude detection threshold of microearthquakes in the region significantly with a magnitude of catalog completeness of M c 5 1 for the study area [Boese et al., 1] and M c 5.55 in the center of the SAMBA network (Figure ). The central area (red square, Figure 1) exhibits the highest seismicity rates in the region with 35. events/3 d/1 km for M c 5 1 and 7.8 events/3 d/1 km for M c This rate is twice as high as the seismicity rate in the vicinity of Lake Pukaki where a network of comparable station density recorded between 1975 and 1983 [Reyners, 1988]. After the 15 July 9 Dusky Sound [M W 7.8; Beavan et al. [1b]; Fry et al. [1]] earthquake, that occurred approximately 35 km to the south of the study area, the seismicity rate in the center of the SAMBA network increased significantly. The same was observed following the 3 September 1 Darfield (Canterbury) earthquake (M W 7.1; Gledhill et al. [11]; Quigley et al. [1]), 18 km to the east of the central Southern Alps, but not after the energetic M W 6. (M e 6.7) Christchurch earthquake on 1 February 11 [Holden, 11; Beavan et al., 11; Kaiser et al., 1] or other large events at teleseismic distances from the Southern Alps. BOESE ET AL. VC 14. American Geophysical Union. All Rights Reserved. 945

2 1.1/13GC S 17. E GeoNet SAMBA from Nov 8 DFDP1 Jan Apr 1 M L M L 3 M L 5 Swarm Main E 171. E S 36 mm/yr 46 Whataroa S FRAN Franz Josef Fox Glacier WHYM FOZ LABE S Mt Cook 43.8 S Lake Pukaki Figure 1. Study area in the central Southern Alps, showing the station network and the seismicity recorded between November 9 and April 1. The rectangles mark subregions discussed in the text with the area outlined in red referred to as the center of the SAMBA network. Focal mechanisms are shown for background swarms (blue), delayed-triggered swarms (orange), and mainshocks (green) of mainshock-aftershock sequences as well as the 1984, M W 6.1 Godley Valley earthquake (gray). All focal mechanism solutions shown are lower hemisphere projections with the colored areas representing compressional first arrivals (for details see supporting information Table S). The inset shows the setting of the study area in the central South Island and focal mechanisms of large earthquakes: the 9 M W 7.8 Dusky Sound earthquake approximately 35 km southwest, the 1 M W 7.1 Darfield (Canterbury), and 11 M W 6. Christchurch earthquakes 18 km east of the study area. We analyze earthquake clustering in general and earthquake triggering in particular in the central Southern Alps region by studying the depths and magnitude distributions, durations, interevent times and focal mechanisms of earthquakes recorded between November 8 to April 1 and within time periods of at least 1 days before and after the Darfield, Christchurch, Tohoku and other large earthquakes (as listed in supporting information Table S1). We address differences and similarities between background and remotely triggered events and discuss possible earthquake triggering mechanisms. 1.. Earthquake Swarms and Proposed Triggering Mechanisms Earthquakes often occur in clusters close in time and space [Mogi, 1963]. Throughout this study, we will refer to temporally related spatial clusters as earthquake sequences. Characteristic of mainshockaftershock sequences is that the largest event occurs at or near the beginning of the sequence and that the aftershock decay-rate follows Omori s law [Omori, 1895; Utsu et al., 1995]. In addition to this temporal behavior, earthquakes in mainshock-aftershock sequences also tend to be more uniformly distributed in space than for earthquake swarms [e.g., Vidale and Shearer, 6]. Earthquake swarms are sequences of often similar-sized earthquakes with no distinctive largest event [e.g., Sykes, 197]. They are commonly observed in geothermal fields, volcanic areas and near mid-ocean ridges [e.g., Sherburn, 199] and are often interpreted to be associated with fluid movement [Hagiwara and Iwata, 1968; Mogi, 1989; Vidale and Shearer, 6] or with aseismic creep [e.g., Roland and McGuire, 9]. We adopt the term remotely triggered from the literature to describe seismicity, which occurs following large earthquakes at distances hundreds of kilometers away from the mainshock s hypocenter as observed worldwide [e.g., Hill et al., 1993; Steacy et al., 5; Velasco et al., 8]. At these distances, the triggered earthquakes are inferred to be caused by the dynamic stresses imposed by the surface waves and predominantly occur during and shortly after the passage of both Rayleigh and Love waves [Prejean et al., 4; West et al., 5; Velasco et al., 8]. This process has been termed waveform-triggering [Brodsky, 6]. BOESE ET AL. VC 14. American Geophysical Union. All Rights Reserved. 946

3 1.1/13GC5171 a) Cumulative sum 4 Days since 1 January Days since 1 January CURATE b) 4 Magnitude Days since 1 January 9 c) log(n) Log. of cumulative number Line representing slope of GR Magnitude distribution 1 5 Number per bin Magnitude Figure. (a) Cumulative sum (blue) and CURATE (black) of earthquakes recorded between November 8 and April 1 for M c Vertical gray lines mark times of the sequences listed in Table 1. The inset shows the cumulative sum of events after the Dusky Sound earthquake with a fit of a modified Omori s law (black line) and a constant linear rate obtained from regression (red line). Within a period of 45 days, Omori s law fits the data better than any linear fit. (b) Magnitudes of earthquakes recorded in the study area versus time. Sequences of events that were identified in Figure a are colored with background swarms in blue, delayed-triggered swarms in orange, and mainshocks in green. Note that some of the mainshock-aftershock sequences are not within the SAMBA network and are therefore not shown in Figure a. The horizontal line marks the cutoff magnitude of M c 5 1. The gray box indicates the period when events were identified on spectrograms of station WHYM. (c) Magnitude distribution and the Gutenberg-Richter magnitude-frequency relation (GR) in the center of the SAMBA network (cf. Boese et al. [1, Figure 8], for the whole study area). Sequences that begin several hours [Hill et al., 1993; Gomberg and Bodin, 1994; Bodin and Gomberg, 1994; Gomberg et al., 1] to several days [Husen et al., 4; Pankow et al., 4] after the surface waves of the triggering event have been referred to as delayed-triggered [Brodsky, 6]. Delayed-triggering has been observed with and without initial waveform-triggering [Mohamad et al., ; Gomberg et al., 1; Pankow et al., 4; Husen et al., 4; Prejean et al., 4; Lei et al., 11]. In the oblique-compressional regime of the central Southern Alps, we observe both background swarms, that occur randomly in time, and delayed-triggered swarms. We observe some waveform triggered seismicity, but the largest numbers of events occur in swarms delayed by several hours to tens of hours relative to the arrival of the surface waves. Although triggering is a common phenomenon, the physical mechanism or mechanisms behind it remain obscure. Local aftershock generation of events triggered by surface waves, aseismic slip, and pore fluid diffusion initiated by the surface waves have each been suggested as processes responsible for the occurrence of delayed-triggered earthquakes [e.g., Brodsky, 6; Gomberg and Bodin, 1994; Syracuse et al., 1]. Comparisons of observed and predicted events using Omori s law and aftershock productivity rates have been used to argue for local aftershock generation [Brodsky, 6], while aseismic slip and pore fluid diffusion mechanisms are often based on migration velocities observed in spatiotemporal event distributions. Few studies have been able to place constraints on the fluid sources involved [e.g., Parotidis et al., 3] or BOESE ET AL. VC 14. American Geophysical Union. All Rights Reserved. 947

4 1.1/13GC5171 causative creep events [e.g., Glowacka et al., ]. However, Parotidis et al. [3] showed that the spatiotemporal distribution of swarms in Vogtland/NW-Bohemia in were consistent with fluid diffusion from a deep fluid source of magmatic origin.. Research Methodology To systematically identify earthquake sequences throughout the Southern Alps region, we apply the CURATE-algorithm developed by Jacobs et al. [13]. CURATE is designed to identify both swarm and mainshock-aftershock sequences by focusing on increased earthquake rates as the main characteristic among all types of earthquake sequences. It is based on a cumulative sum technique (CUSUM) [Page, 1954), which is often used to detect subtle changes in rate from a fluctuating background level. At the time of each earthquake in the catalog, CUSUM uses the average daily number of earthquakes for the entire period to estimate the expected cumulative number of earthquakes from the beginning of the catalog up to that point in time. The CURATE algorithm identifies periods of elevated seismicity rates within an earthquake catalog and applies two selection criteria to earthquakes within these temporal clusters: a distance-rule (earthquakes must have a spatial relationship) and a day-rule (which concatenates sequences in the same location separated by pauses shorter than specified by the day-rule into a single sequence). This algorithm differs from other link-based clustering algorithms in that it is diagnostic of high-rate sequences within a subset of the original earthquake catalog. We apply the CURATE algorithm to the SAMBA catalog for both magnitude thresholds and identify the same sequences but with different numbers of events. To distinguish a mainshock-aftershock sequence from a swarm, we require that the largest event occurs at the beginning of the sequence (% of the duration), and that the magnitude difference between the largest and the second-largest events exceeds a certain threshold. According to Båth s law, this magnitude difference typically is 1. [Båth, 1965]. However, this value is an average of the broad range ( 3) of observed values worldwide [Felzer et al.,, their Figure ] and is typically determined for larger magnitude sequences than considered here. We adopt a value of.7, which is more similar to values observed for different sequence types in the Central Volcanic Region of New Zealand [Sherburn, 199]. Magnitudes and focal mechanisms were determined as described by Boese et al. [1]. During the aftershock periods following the Dusky Sound and Darfield earthquakes, our standard earthquake identification method based on short-term/long-term averages of the summed raw traces [Boese et al., 1], performed poorly and missed local events within the SAMBA network. However, these events are evident as high-frequency signals visible in spectrograms. To systematically check for local events and to ensure that catalog completeness remains the same, continuous spectrograms of station WHYM (Figure 3 and supporting information Figures S1 and S), were examined visually for 1 day periods before and after the large regional and teleseismic events listed in supporting information Table S1. Station WHYM is one of the two stations closest to the swarm area and has been recording continuously since its installation in November 8. All events detected this way, with discernible P and S-phases at 3 stations, were located and their magnitudes determined. The first arrival times of unlocatable events were also noted. The visual examination of spectrograms has proven effective and even lowered the cutoff magnitude by approximately half a unit after the Dusky Sound earthquake (Figure b). However, the detection was not improved after the Darfield earthquake as many more aftershocks were recorded by the SAMBA network due to the shorter distance of 18 km to the mainshock. Pg and Pn arrivals of the Darfield aftershocks had similar differential times and frequency contents as local P and S-phases of remotely triggered events. However, no meaningful location could be obtained if the two P-phases were confused with local arrivals. The event identification using spectrograms was only adopted for short periods because it is time-consuming and introduces bias in the event detection. The waveforms of the majority of events in the swarms identified with CURATE showed high waveform similarity. For this reason, we searched our cross-correlation catalog of all events in the Southern Alps for possible other events belonging to each family (e.g., with magnitudes below the cutoff magnitude or missed events) assuming that waveform similarity is a characteristic of the swarm events. All events in the catalog that were linked by above-average cross-correlation coefficients (>.5 at a minimum of five stations) to events identified by CURATE were visually examined. Highly similar events were added whereas those CURATE events with significantly different waveforms were marked. Cross-correlation coefficients were BOESE ET AL. VC 14. American Geophysical Union. All Rights Reserved. 948

5 1.1/13GC5171 a) 5 Frequency [Hz] b) low pass filtered x Surface waves Amplitude [counts] c) high pass filtered 15 Amplitude [counts] :5 9:3 9:35 9:4 9:45 9:5 9:55 1: :5 9:3 9:35 9:4 9:45 9:5 9:55 1: Time 9:57:5 9:57:3 9:57:35 Figure 3. (a) Spectrogram and (b and c) low-pass and high-pass filtered seismograms of the vertical component of station WHYM showing the M W 7.8 Dusky Sound earthquake of 15 July 9 and subsequent seismicity. Large aftershocks near the mainshock as listed in the GeoNet catalog are shown by annotated arrows (giving the magnitude) or red dashed lines whereas remotely-triggered local earthquakes are shown by down-pointing arrows indicating high-frequency energy or black dashed lines. Note that not all high-frequency events could be unambiguously identified as remotely triggered earthquakes. The inset in Figure 3c shows a local event on the high-pass filtered trace ( Hz). obtained on band-pass-filtered waveforms (3 15 Hz) with window lengths of.5 s and.35 s for the P and S-phase, respectively, starting.1 s before the corresponding phase pick. We set an upper limit of. s on the lag time, allowing for the identification of mispicked phases, but preventing correlation with converted phases. Figure 4 shows the largest swarm as an example for the combined events from CURATE and those additionally found through cross correlation. It should be noted here that, despite the waveform similarity, these sequences do not stand out from all other background events in terms of particularly high cross-correlation coefficients (with means of.63.76, see Table 1), so they would not have been easily detected without using CURATE. The moderate cross-correlation values may simply reflect the multitude of small faults in the region producing earthquakes or could result from correlated Gaussian noise which lowers the crosscorrelation coefficient [Du et al., 4]. In summary, the final sequence catalog (Table 1) is based on seismicity rate-changes analyzed using CURATE but the individual events in a sequence are selected on the basis of their waveform similarity. We also analyze the average time between earthquakes in the sequence (average interevent time, which is similar to the inverse of the rate) to search for distinguishing features of the sequence types. The duration of each sequence is normalized so that the beginning of the sequence is zero and the end is one. This normalization allows us to look for temporal relations in rate that may occur in sequences with different durations. We compare the observed rates of the different sequence types to those expected from Omori-type decay nðtþ5k ðt1c Þ p. Here n(t) is the number of earthquakes per unit time at time t; k represents a measure of the productivity of the mainshock and depends on the magnitude completeness threshold of the BOESE ET AL. VC 14. American Geophysical Union. All Rights Reserved. 949

6 1.1/13GC5171 Event number triple Time [s] ML ML ML ML ML ML.4 1. ML.4.9 ML.4.63 ML ML ML ML..97 ML ML 3. ML ML ML ML ML ML ML 1.8. ML.4 ML.69 ML ML ML ML ML.7.55 ML ML ML ML ML ML ML ML ML.56 ML.4 3. ML.9.3 ML ML ML ML ML ML ML Figure 4. The triggered earthquake swarm on 18 June 9 as recorded on the vertical component of station WHYM ordered by origin time (number on the top left of each trace gives time of occurrence in year, month, day, hour, minute, and second). The number on the right on top of each trace gives the magnitude (if it could be determined) and time in minutes since the previous event. For the first event, the time to the last event is calculated instead. Note that multiplets occur so close in time that events are overlapping (events 7 8, 1, and 4 6). The duration without the last event is 7.9 days. The last event occurred 94.9 days after the swarm at the same location and was identified due to its waveform similarity with the remaining events. given earthquake catalog; c describes a temporal offset that compensates for incomplete data in the earliest part of the aftershock sequence, and p describes the exponential decay of the aftershock rate. 3. Results 3.1. Observed Earthquake Sequences We identify 18 sequences between November 8 and September 1, each consisting of 1 events with similar waveforms (Figure 4). These sequences account for 15.3% of the total number of recorded earthquakes of M c 1.. We exclude five smaller sequences that occur throughout the study area from our analysis because there may be more of these sequences that have not been identified due to detection issues. The occurrence time of these smaller sequences appears to coincide with periods during which recording was nonuniform (e.g., before day 167 in Figure and in the time period spanning the Darfield earthquake, see also supporting information Figure S3). The durations of individual sequences are given in Table 1. All mainshock-aftershock sequences (4/4) and the majority of all swarms (11/14) have durations 1 days. Of the 18 sequences, all except four (three mainshock-aftershock sequences and one swarm) occur in the center of the SAMBA network (Figures 1 and 5), an area of 1 km 3 1 km between stations WHYM and LABE. This zone of abundant, low-magnitude seismicity (M L.8) stands out by having the shallowest hypocenter locations throughout the study area. The earthquakes in this zone are well located with average azimuthal gaps of (1r) and depth uncertainties of (1r) km[boese et al., 1]. BOESE ET AL. VC 14. American Geophysical Union. All Rights Reserved. 95

7 1.1/13GC5171 Table 1. List of Earthquake Sequences in the Central Southern Alps Region Identified Using the CURATE Algorithm and the Cross-Correlation Method a CURATE Algorithm Cross Correlation Start Date (DD/MM/YYYY) Lat. [ ] Lon. [ ] M.55 M 1. Diff. Total Pre Post M.55 Max Mag. Duration (Type) CC- Coeff. Mainshock-Aftershock Sequences 1/3/ (S).63 14/8/ (S).67 9/11/ (S).68 8/8/ (S).68 Background Swarms 4/5/ <1 (S).76 6/7/ (S).69 9/7/ (L).7 1/9/ (L).71 /9/ (L).71 16/1/ (S).67 9/1/ <1 (S).71 3/3/ (S).66 9/8/ (S).73 Delayed-Triggered Swarms 15/7/9 (8) (S).7 16/7/9 () <1 (S).68 18/7/9 (7.6) (S).7 3/9/1 (6.4) (S).68 4/9/1 (7.3) (S).67 a For the CURATE method, number of identified events above the cutoff magnitudes of M c 5 1. and M c 5.55 and the number of different events ( diff. ) as seen from cross correlation for the lower cutoff magnitude are listed. For the cross-correlation method, the total number of events with similar waveforms and those before ( pre ) and after ( post ) the main swarm are given. For each sequence the mean location, the mean cross-correlation coefficient, the maximum magnitude (from cross-correlation method) and the duration in days are given, where L stands for long (3 days) and S for short durations (1 days). For long swarms, only the largest group of events is listed for the CURATE method; later events have been identified as individual smaller sequences. For delayed-triggered swarms, the delay of the swarm in hours relative to the surface wave arrival of the triggering earthquake is specified in brackets. Note that the delayed-triggered swarm on 16 July 9 was not identified by the CURATE algorithm. The only swarm recorded outside the center of the SAMBA network occurs in the aftershock zone of the 1984 Godley Valley earthquake [Anderson and Webb, 1994]. This swarm started 5 days before the Darfield earthquake and continued for 5 days afterwards. Only three events in the first 3 days after the Darfield event could be identified from this swarm, despite the improved detection using spectrograms. We obtain at least one focal mechanism solution for each sequence, most often from the largest event. If multiple focal mechanism solutions exist for swarm events, these are found to be identical within their uncertainty limits [cf. Walsh et al., 9]. The angular uncertainty of each focal mechanism is estimated to be 5 31 as listed in supporting information Table S and discussed by Boese et al. [1]. To further verify whether all events in a sequence have similar focal mechanisms, we also compare the distribution of the polarity picks on the focal sphere directly; in each case it is almost the same, reflecting the waveform similarity of the constituent swarm events. The obtained focal mechanism solutions indicate steeply dipping nodal planes (Figure 6 and supporting information Table S) and predominantly strike-slip faulting. We identify the fault plane as the nodal plane that is more favorably oriented in the prevailing stress field [Boese et al., 1; Townend et al., 1] in an Andersonian sense. For this purpose, we determine the angle between the resolved shear vector and the slip vector for each of the nodal planes in 3-D space. Of the 18 measurements, 15 values lie within a range of 6 of the Andersonian angles (Figure 6). 3.. Remotely Triggered Seismicity We observe remotely triggered seismicity following two large earthquakes in New Zealand (supporting information Table S1). The M W 7.8, 9 Dusky Sound earthquake occurred approximately 35 km to the south of the study area and ruptured a c. 8 km long section of the subduction zone interface between the subducting Australian plate and the overlying Pacific plate [Beavan et al., 1b]. It caused about.6 cm of permanent westward coseismic displacement at Haast [Mahesh et al., 11], approximately 85 km southeast of the center of the SAMBA network. This earthquake stands out from other earthquakes of comparable size because of its waves large low-frequency (.1.1 Hz) and relatively small high-frequency contents (>5 Hz)[Fry et al., 1]. To calculate the peak static and dynamic stress changes caused by this event (see sections 4. and 4.3), we model the mainshock rupture using the Coulomb static stress triggering software and use available BOESE ET AL. VC 14. American Geophysical Union. All Rights Reserved. 951

8 1.1/13GC5171 Number of events per day Number of events per hour Number of events per hour 6 Dusky Sound earthquake Mar 4 May 6/ 9 Jul 1/ Sep 17 Dec 9 Jan Apr Darfield earthquake not processed Days since 1 January 9 Dusky Sound earthquake 15 Jul 8 hr 16 Jul hr 18 Jul 7.6 hr nd part of 15 Jul Hours since Dusky Sound earthquake Darfield earthquake 3 Sep 6.4 hr 7.3 hr c) 5 1 Hours since Darfield earthquake Y (km) information on peak ground accelerations (PGA) and peak ground velocities (PGV) in the study area. We determine PGA and PGV from the seismograms of SAMBA station FRAN, a borehole sensor installed at 1 m depth near Franz Josef and GeoNet site FOZ, a broadband surface sensor near Fox Glacier. Absolute PGA values are influenced by the local geology at the site and other factors (signal period, source dimensions of the earthquake, radiation characteristics, etc.), so peak ground velocities, which are more strongly dependent on lowfrequency components of ground motion than PGA [e.g., Boatwright et al., 1], provide a more representative measure of the intensity of ground shaking [Tso et al., 199]. The Dusky Sound earthquake caused PGV of.7 cm/s and PGA of.3 g at 1 s period on the vertical and horizontal components of station FRAN (Table ). Similar values were observed on strong motion accelerometers in the area [cf. Fry et al., 1, Table 1]. Within the first 4 h after the Dusky Sound earthquake, 146 microearthquakes (with 51 events of M c 1) occurred in the central Southern Alps region. This is the highest number of events per day recorded since installation of the SAMBA network began in November 8 (Figure 5). Almost all of the detected events could be located (93.% recorded by 3 stations). The triggered seismicity commenced shortly after the passage of the surface-waves (9:4:5 UTC; first identifiable event 16 min later at 9:4:19 UTC) and continued for approximately 5 days. Three delayed-triggered swarms occurred in the center of the SAMBA network: The first and most vigorous delayed-triggered swarm (M L.8) occurred 8 h after the passage of the surface waves of the Dusky Sound earthquake. This swarm continued for several days, with a second burst occurring 4 days later in the same location. The second swarm started h after the Dusky Sound earthquake, 8 km northeast of the first swarm. A third swarm started 7.6 h after the passage of the surface waves at a distance of 3.7 km SE and 7.7 km SW of the first and second swarms, respectively (Figures 5 and 7). At these distances between the delayed-triggered swarms, static stress changes, which are expected to affect only the area within several source dimensions (on the order of a few hundred meters for small earthquakes according to Sibson [1989]), become insignificant compared to, for example, tidal stress changes [e.g., Hill et al., 1993]. The increased microseismic activity was followed by days of quiescence (Figure 5b). a) b) 8 8 hr 8. km d) hr km 7.7 km hr X (km) Figure 5. (a) Seismicity versus time in the center of the SAMBA network (red rectangle in Figure 1) where the majority of the swarms occur. (b and c) Hourly distributions of events after the M W 7.8 Dusky Sound earthquake and the M W 7.1 Darfield earthquake that caused remotely triggered microseismicity in the central Southern Alps. Red crosses indicate events of M L. (d) Geometry and delay of the delayed-triggered swarms after Dusky Sound (labeled stars) and Darfield earthquakes (unlabeled) with the symbol size representative of the magnitude of the largest event. BOESE ET AL. VC 14. American Geophysical Union. All Rights Reserved. 95

9 1.1/13GC S 43.5 S 43.6 S 17.3 E L S Hmax preferred fault plane 17.4 E Number Number Angle between slip vector and shear vector Elevated seismicity rates in the study area were also observed after the M W 7.1, 11 Darfield earthquake [Gledhill et al., 11; Quigley et al., 1], which produced a complex c. 44 km long surface rupture 18 km due east of the center of the SAMBA network [Stramondo et al., 11; Elliott et al., 1]. This earthquake caused similar to slightly stronger shaking at station FRAN to the Dusky Sound earthquake, with PGA and PGV values of.6 g and cm/s on the horizontal components for waves with periods of.6 s and 3 s, respectively (Table ). The triggered events were not as well recorded as following the Dusky Sound earthquake (only 53.4% of the events could be located), because 3 out of 1 SAMBA stations were not operating at the time following a particularly snowy austral winter. The triggered seismicity commenced shortly after the arrival of the surface waves (16:36:3 with first identifiable event 11 min later at 16:47:5 UTC, as shown in supporting information Figure S1) and continued for at least days. SAMBA recorded 94 events, of which 38 exceeded magnitude M L 1. Two delayed-triggered swarms occurred 6.4 h and 7.5 h after the passage of the surface waves in the same area where delayed-triggered swarms had occurred after the Dusky Sound earthquake (Figure 8). The M W 6. (M e 6.7) Christchurch earthquake on 1 February 11 [Holden, 11; Beavan et al., 11; Kaiser et al., 1] did not trigger discernible seismicity in the central Southern Alps (supporting information Figure S). This event caused lower PGA and PGV values than the Dusky Sound and Darfield events. No remotely triggered seismicity following large distant earthquakes (from 3 to 13,6 km distance, M W as listed in supporting information Table S1) between 9 and 11 has been observed either. All Dip preferred fault plane both nodal planes Figure 6. Preferred fault plane (red) and auxiliary nodal plane (blue) of focal mechanism solutions of swarm events in the central Southern Alps in map-view (left). By determining the angle between the slip vector and the shear vector for both nodal planes, we identify the fault plane as the plane which is more consistent with Andersonian faulting (values listed in supporting information Table S). Faults are adopted from Cox and Barrell [7]. Note that faults shown are Mesozoic, nearsurface features and their traces reflect intersection with the topography. Their Cenozoic motion is more consistent in orientation than the map implies (S. Cox, personal communication 14). Panels on the right show the (top) dip and the (bottom) angle to the shear vector of all nodal planes (blue) and the preferred fault planes (red). The gray boxes illustrate the range of angles within of the optimal orientation of the shear vector. Table. Observed Peak Ground Accelerations (PGA, Relative to 9.81 m/s ), Peak Ground Velocities (PGV), and Dynamic Stress (r D ) at Sites of the SAMBA (FRAN) and GeoNet Networks (FOZ) for Earthquakes That Triggered (Marked by *) and Did Not Trigger Seismicity in the Central Southern Alps Region a PGA [31 3 g] PGV (31 3 m/s) Date (DD/MM/YYYY) Earthquake M W Dist. (km) Site Code V H1 H T (s) V H1 H T (s) r D (MPa) 15/7/9 Dusky Sound* FOZ FRAN /9/1 Darfield (Canterbury)* FRAN //11 Christchurch FRAN /11/9 Fiji 7. 3 FOZ FRAN //1 Chile FOZ /3/11 Tohoku FOZ FRAN a V, H1, and H mark the vertical and the two horizontal channels of these sensors. T indicates the period of the wave for which the PGA and PGV were determined. BOESE ET AL. VC 14. American Geophysical Union. All Rights Reserved. 953

10 1.1/13GC5171 Background seismicity Triggered seismicity a) Background swarm events b) after Dusky Sound earthquake after Darfield earthquake 43 3' 43 3' c) 17 18' 17 4' d) 17 18' 17 4' Depth [km] 5 1 Depth [km] 5 1 Pre Dusky Sound to 4 days after Post Dusky Sound to 4 days after Darfield e) f) 43 3' 43 3' 17 18' 17 4' 17 18' 17 4' Time [hr] after Dusky Sound earthquake Time [hr] after Darfield earthquake Figure 7. Background seismicity and triggered seismicity in comparison. (a and c) Background seismicity recorded between 11 November 8 and 15 July 9, and July and 1 May 1 in map view and on a depth section. Those events identified to form background swarms are highlighted in blue. (b and d) Triggered seismicity after the Dusky Sound (15 19 July 9, orange) and Darfield earthquakes (3 7 September 1, green) shown with mapped secondary faults (gray) from the QMAP Aoraki 1:5 by Cox and Barrell [7]. (e) Background seismicity before (black) and triggered seismicity following the Dusky Sound earthquake (colored by time of occurrence). The gray line shows the location of a possible fault from which fluid diffusion is consistent with the spatiotemporal distribution of the triggered events after Dusky Sound. (f) Background seismicity after the Dusky Sound and before the Darfield earthquake (black) and triggered seismicity following the Darfield earthquake (colored by time of occurrence). The same gray line as in Figure 7e is for reference. of these distant events caused significantly less shaking in the central Southern Alps than either the Dusky Sound or the Darfield event (Table ) Long-Term Record of Triggered Seismicity Although the swarms analyzed in this study occur predominantly in the center of the SAMBA network, two large swarms known as the Mt. Cook and Fox swarms have previously been recorded in the study area [Leitner et al., 1; O Keefe, 8]. To extend the search for sequences and potential triggered events, we analyzed the seismicity in the GeoNet catalog between 1993 and 8 for different subregions in the study area. An upgrade of the GeoNet stations was undertaken in 3, so apparent changes in the catalog BOESE ET AL. VC 14. American Geophysical Union. All Rights Reserved. 954

11 1.1/13GC5171 Fluid diffusion scenarios a) b) Depth distribution Fluid reservoir hr 43 3' 8hr Fault Z(km) 5 1 background seismicity only 7.6hr background swarms only 15 delayed triggered swarms only 17 18' 17 4' Events c) Distance from reservoir [km] d) Distance from fault [km] Fluid diffusion from reservoir 1 m/s 3 m/s 1 m/s Fluid diffusion from fault 1 m/s 3 m/s 1 m/s Time [hr] Time [hr] after Dusky Sound earthquake Time [hr] after Darfield earthquake Figure 8. Possible interpretation of fluid sources from which diffusion starts and causes triggered seismicity. (a) Triggered seismicity after the Dusky Sound (15 19 July 9, orange to blue colors) and Darfield earthquakes (3 7 September 1, green colors) shown with possible sources from where fluid diffusion would be plausible (gray). Note that only the southern boundary of the fluid reservoir is resolved. Circles and diamonds indicate the mean locations of the delayed-triggered swarms after the Dusky Sound and Darfield earthquake, respectively. Dashed circles reflect distances that fluids could propagate by diffusion plotted with respect to the mean swarm location. (b) Depth distributions of all events not in sequences (black), the background swarms (blue), and the triggered seismicity only (orange) in comparison. (c) Fluid diffusion envelope for the fluid reservoir as a source of fluids. (d) Fluid diffusion envelope for the fault as a fluid source. around this time are an artifact of changes to the network. This analysis reveals past swarms in the center of the SAMBA network and in the area of the Fox swarm (Figure 9). We also analyzed the PGA values recorded at Haast (approximately 85 km southwest of the network) for large earthquakes similar to the Dusky Sound and Darfield earthquakes (Figure 9 and supporting information Table S3). There is no indication of remotely triggered seismicity commencing shortly after a large earthquake, but the triggered swarms recorded by our network in 9 and 1 are not listed in the Geo- Net catalog because of their small magnitudes. Therefore, the cutoff magnitude of the GeoNet catalog of M c.6 [Petersen et al., 11] is insufficient to establish whether triggering occurred previously in this area. This emphasizes the fact that the detection of low-magnitude earthquake swarms strongly depends on the local station network and its detection threshold [e.g., Frankel et al., 198]. BOESE ET AL. VC 14. American Geophysical Union. All Rights Reserved. 955

12 1.1/13GC5171 Peak Ground Velocity (mm/s) Magnitude Magnitude Magnitude Magnitude 1 a) M > b) c) d) e) This study Brodsky and Prejean Temporary Deployment M > 6.5 M > 7. M > Year Network Upgrade MC FX SAMBA Figure 9. Time series of seismicity in the central Southern Alps region since (a) Maximum horizontal peak ground velocities (PGV in mm/s) of major New Zealand earthquakes and their aftershocks (see supporting information Table S3) recorded at Haast c. 85 km southeast of the center of the SAMBA network. The horizontal dashed line indicates the threshold proposed by Brodsky and Prejean [5] above which triggering is independent of the waves amplitude, duration, and energy-density but dependent on the frequency of the waves. (b d) Seismicity versus time for the subregions as shown in Figure 1. Black circles indicate sequences detected in this study, annotated arrows show the Mt. Cook (MC) and Fox swarms (FX), and changes in the station network. The daily seismicity rate is (b).86, (c) 1.66, and (d).45 events/3 d/1 km for the GeoNet catalog (M c.6) between 1993 and 7. (e) Seismicity versus time for the whole study area. 4. Discussion In the following, we discuss different scenarios and evidence for and against specific mechanisms for delayed-triggering of earthquakes in the Southern Alps region Likelihood of Random Occurrence To quantify whether the occurrence of the observed sequences following two large regional earthquakes was unusual, and to determine whether they are in fact triggered, we calculate the probability of recording more than a given number of sequences at any given time, following the approach of Lei et al. [11]. For this calculation, we only use the 11 background and mainshock-aftershock sequences recorded inside the center of the SAMBA network with at least six earthquakes above the cutoff magnitude M c The average time between these 11 sequences is 37 days. Note that it is the actual average time between sequences, rather than the total time divided by the number of sequences, which would give a longer mean time as it includes time at the beginning and end of the catalog in which sequences do not occur. The probability of observing more than a given number of sequences n then becomes PðnÞ5 xn exp ðxþ n! where x is the time window of observation over the mean time between sequences (e.g., 5/37). The probability of observing a given number of sequences for n > is then 1½PðÞ1Pð1Þ1Pðn1ÞŠ: The probability of observing at least one sequence in any 5 day window is 1.7%. The probability of observing two swarms or more within 5 days after a large earthquake, but unrelated to it, is.8%. This suggests that there is only a small chance of observing several background swarms by coincidence. Given that delayed-triggered swarms were observed twice following a large regional earthquake, these probability calculations suggest that remote earthquake-triggering is common in the Southern Alps. Shaking from large BOESE ET AL. VC 14. American Geophysical Union. All Rights Reserved. 956

13 1.1/13GC5171 local and regional earthquakes (M 6.) above the observed PGV value of.7 cm/s produced by the Dusky Sound and Darfield earthquakes occurred at least 5 9 times in 18 years in the central Southern Alps (as recorded by the strong motion sensors at Fox Glacier and Franz Josef) and 9 times at Haast (Figure 9). This also suggests that triggering should be commonly observed by dense local networks. 4.. Static Stress Changes We compute static stress changes for the Dusky Sound and Darfield ruptures using Coulomb software version 3.3 of Toda et al. [11]. Mainshock rupture patterns for the events are adopted from Beavan et al. [1b] and Stramondo et al. [11]. The static stress change is determined for both optimally oriented fault planes and for the inferred failure planes of the remotely triggered swarms as listed in supporting information Table S. Given the reported rupture dimensions of the two mainshocks [Beavan et al., 1b; Stramondo et al., 11; Elliott et al., 1], the swarm area in the central Southern Alps lies 4.5 rupture lengths away from both the Dusky Sound and Darfield earthquakes. In both cases, we obtain positive static stress changes of.3.6 MPa in the center of the SAMBA network and often the swarm area lies in the transition zone between areas of stress increase and decrease for the inferred orientations of the triggered swarms Triggering Threshold No remotely triggered seismicity was observed following the M W 6. Christchurch earthquake or other large distant earthquakes, suggesting that there may be a stress threshold, amongst other factors, that has to be exceeded for triggering to occur. We calculate peak dynamic stress changes in the Southern Alps caused by the triggering mainshocks. The peak dynamic stress r D can be calculated from the PGV u D using the relationship r D l S u D =v S ; (1) where l S Pa is the shear modulus and v S km/s the surface wave phase velocity [Jaeger and Cook, 1979]. The calculated peak dynamic stresses corresponding to PGVs of c..1 cm/s of distant earthquakes that did not trigger microearthquakes (Table ) are on the order of MPa. The maximum PGV values of.3 cm/s of the Christchurch earthquake correspond to peak dynamic stresses of.3 MPa. This value exceeded the threshold of. cm/s noted by Brodsky and Prejean [5] above which triggering was observed in the Long Valley Caldera of California. This threshold was exceeded 13 times at Haast between 1993 and 1. As triggering has not been reported previously for the Southern Alps, despite several temporary networks in this area, this may suggest that larger maximum PGV values than. cm/s are required for triggering. Indeed, the ground motion of the Dusky Sound and the Darfield earthquake locally exceeded PGV and PGA values of.7 cm/s and.1 g, respectively. These values correspond to peak dynamic stresses of.9 MPa. The stress threshold therefore is in the range.3.9 MPa, which lies at the lower end of the range of peak dynamic stresses observed worldwide spanning.1.5 MPa [Brodsky et al., ; Prejean et al., 4; Husen et al., 4; Pankow et al., 4; Peng et al., 1]. This suggests that triggering thresholds in the compressional regime of the Southern Alps are of the same order as those in transtensional and extensional settings, where triggering has been observed previously. Brodsky and Prejean [5] demonstrated that above a shaking threshold of. cm/s, dynamic triggering is independent of the amount of shaking (amplitude), its duration and the energy-density of the waves but it is dependent on the period of the wave (typically 1 4 s) [Anderson et al., 1994; Prejean et al., 4; Brodsky and Prejean, 5]. In the central Southern Alps, the seismicity response to the Dusky Sound earthquake, which generated more low-frequency energy (.1 Hz) than the Darfield event, was more intense in that it caused more earthquakes despite the larger PGA and PGV of the Darfield earthquake (Table ). The high stress-drop Darfield earthquake [Quigley et al., 1] is unusual because of the dominant periods of 1 3 s. However, the long-period waves may not have been well recorded by the 4.5 Hz short-period seismometer at FRAN, and the broadband GeoNet stations in the vicinity of the SAMBA network were all clipped for long periods Triggering Due to Local Aftershock Generation and Decay One mechanism suggested for the occurrence of waveform-triggering and delayed-triggering is that local aftershocks are generated by the stress changes imposed by the surface waves that in turn generate their BOESE ET AL. VC 14. American Geophysical Union. All Rights Reserved. 957

14 1.1/13GC5171 own sequence of aftershocks [Brodsky, 6]. Brodsky [6] compared numbers of observed and predicted events using Omori s law and aftershock productivity rates to explain long-lasting triggered seismicity in the Long Valley caldera after the M W 7.3, 199 Landers earthquakes as local aftershocks. We can fit the rate-change following the Dusky Sound earthquake with a modified Omori s law. Assuming p 5 1[Utsu et al., 1995], the resulting cumulative number NðtÞ5Klnð11t=c Þ is overlain in Figure (inset). We obtain a best fitting c-value of c 5.35 days, and a corresponding K-value of 14.4 earthquakes ðk5n max =ln ð11t max =c ÞÞ, for t 5 t max, where N max is the total number of events in this time and t max is the duration fitted to this time using a maximum likelihood method for M c 5.55 in the swarm area. To test the fit, we compare the sum of the residuals for the Omori s law decay with the residuals from linear fits using the background rates before and after the Dusky Sound earthquake (.75 and 1. earthquakes/d) and the best linear regression during that time (.71 earthquakes/d). The Omori s law fit is better than all three linear fits for 45 days following the Dusky Sound earthquake. This suggests that a decay process similar to that of a mainshock-aftershock sequence can explain the observed seismicityrate change. However, while the overall number of triggered seismicity can be fit with Omori s law, individual delayed-triggered sequences do not show the same decay pattern as the observed local aftershock sequences (Figure 1). Assuming that local delayed-triggered sequences were aftershocks generated by larger waveform-triggered events, the magnitude distribution of the events during the period of elevated seismicity after the Dusky Sound and Darfield earthquakes should reflect this [Brodsky, 6]. The largest event within 1 h of the arrival of the surface waves of the Dusky Sound earthquake was a M L 1. event. Two M L. earthquakes occurred 1. and 1.4 h after the surface-wave arrivals. All other larger events (M L ) including the largest triggered event of M L.8, which is also the largest earthquake recorded in this region, occurred later than 8. h after the surface-wave arrival (Figure 5). All events that occurred during the passage of the surface waves and could be located have insignificant magnitudes. It is reasonable to assume that the 7% of events we were unable to locate were smaller than the located events during this time. Similarly, after the Darfield earthquake, one M L 1.5 occurred.6 h after the surface wave train and was later followed by a M L. and M L.5 at.8 h and 3.8 h, respectively. All other M L events occurred >11.5 h later. Therefore, we suggest that observed magnitude distributions of the delayed-triggered events are inconsistent with the idea that they are aftershocks of the waveform-triggered earthquakes. Also, given the numbers of events in previously observed sequences (maximum events for mainshock-aftershock sequences and 3 events for background swarms given the cutoff magnitude M c.55), at least three sequences of similar type are needed to explain the number of observed events within 5 days (71 events with 81 of M c.55) Fluid Diffusion Isotope and fluid inclusion studies by Upton et al. [1995] suggest that a fluid-rich facture system extends to depth of 5 6 km in the central Southern Alps. Overpressured fluids at depth 6 km have been inferred from the reduction of seismic wave-speeds in wide-angle reflection and teleseismic P-wave arrival data on SIGHT transect [Stern et al., 1,, 7], while low resistivities indicate interconnected fluids at these depths [Wannamaker et al., ]. A distinct transition between seismic (7.5 km) and aseismic depths can be seen in Figures 7c and 7d. To test whether fluid diffusion is a plausible triggering mechanism for the delayed-triggered swarms, we compute whether the geometry and delay of swarms occurring after the Dusky Sound earthquake can be explained by the expansion of a triggering front of high pore-fluid pressure. Shapiro et al. [1997] showed that for isotropic fluid diffusion as the dominant triggering mechanism, p the induced seismicity around an injection well (point source) follows an expansion radius r given by r ffiffiffiffiffiffiffiffiffiffiffi 4pDt, where D is the hydraulic diffusivity and t is time. We make use of the fact that three swarms were observed, and treat the location of the point source in a similar manner to the earthquake location problem by using linearized methods: t i 5T1 1 h ðx-x i Þ 1ðY-y i Þ 1ðZ-z i Þ i () 4pD Here i describes the number of observations of triggered swarm delay times t, T is the starting time of the fluid diffusion, D is the hydraulic diffusivity, X, Y, Z the coordinates of the fluid point source, and x, y, z are BOESE ET AL. VC 14. American Geophysical Union. All Rights Reserved. 958