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Microseismicity in theTararua—Wairarapa area: depth-varyingstresses and shallow seismicity in thesouthern North Island, New ZealandWalter J. Arabasz a b & M. A. Lowry aa DSIR , Geophysics Division , Wellington , New Zealandb Department of Geology and Geophysics , University of Utah , SaltLake City , Utah , U.S.A.Published online: 02 Feb 2012.

To cite this article: Walter J. Arabasz & M. A. Lowry (1980) Microseismicity in the Tararua—Wairarapaarea: depth-varying stresses and shallow seismicity in the southern North Island, New Zealand, NewZealand Journal of Geology and Geophysics, 23:2, 141-154, DOI: 10.1080/00288306.1980.10424202

To link to this article: http://dx.doi.org/10.1080/00288306.1980.10424202

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N.Z. Journal of Geology and Geophysics Vol. 23, No.2 (1980): 141-54

Microseismicity in the Tararua-Wairarapa area: depth-varying stresses and shallow seismicity in the southern North Island, New Zealand

WALTER J. ARABASZ* AND M. A. LoWRY

Geophysics Division, DSIR, Wellington, New Zealand

ABSTRACT

Results from previously unpublished microearthquake field studies in the Tararua-Wairarapa area during 1971-72 demonstrate important differences between crustal seismicity shallower than 20 km and intense, regionally extensive, underlying seismicity in the 20-40 km depth range. As earlier observed in the Marlborough region, and more recently in the southernmost North Island, composite focal mechanisms in the Tararua-Wairarapa area indicate markedly different stress orientations above and below 20 km. Compressional transcurrent faulting in the upper crust con­trasts with extensional normal faulting at depths of 20-40 km. The layer 20-40 km deep appears to be far more seismically active than the upper crust and is interpreted to mark the uppermost part of the subducted Pacific plate. Accordingly, the data here provide valuable control for locating the plate interface, for modelling elastic deformation at a convergent plate boundary, and for assessing fault activity and shallow seismicity in the southern North Island.

KEYWORDS Wairarapa, Tararua, seismicity, microseismicity, fault geometry, stress, depth-varying stress, Benioff zone, crustal plate, subduction zone

INTRODUCTION

During the last several years, micro earthquake field studies have been carried out in various parts of New Zealand to elucidate aspects of crustal seismicity, in particular, its depth distribution. its correlation with geologic structure, and the nature of focal mechanisms involved (e.g., Scholz et al. 1973; Robinson & Arabasz 1975; Arabasz & Robinson 1976; Rynn & Scholz 1978). Here we summarise previously unpublished results of two early field studies that are particularly germane to understanding shallow seismicity in the southern North Island: (1) a microearthquake reconnaissance of the southern North Island carried out in 1971, and (2) a detailed follow-up study of the Tararua-Wairarapa area in 1972.

This paper documents evidence from the 1971-72 studies for important differences between seismicity above 20 km and that in the 20-40 km depth range. Focal mechanism from the Tararua-Wairarapa area indicate markedly different stress orientations above and below 20 km, with compressional transcurrent faulting in the upper crust, and extensional normal faulting at depths of 20-40 km. A second important result of the 1971-72 studies is evidence that micro seismicity in the upper crust is less intense than that below 20 km, which is abundant and diffusely distributed in the southern North Island.

Data in this report extend the observations of Arabasz & Robinson (1976) in the northern South Island, where differences in the characteristics of seismi­city above and below about 20 km were first docu­mented. They also complement subsequent results re-

Received 8 August 1979, revised 24 October 1979

cently published by Robinson (1978) from a local tele­metry network of high-gain seismographs that has operated in the Wellington region since 1975. The data are valuable as a reference to which space-time patterns of loea: seismicity or stress orientation can be com­pared for evidence of change.

Regional Seismicity

As part of New Zealand's Main Seismic Region, the southern North Island (Fig. 1) is characterised by a high level seismic activity, particularly at crustal depths (e.g., Hatherton 1970; Adams & Hatherton 1973). A Benioff zone dips north-westward beneath this region (Hamilton & Gale 1968; Adams & Ware 1977), with subcrustal earthquakes occurring chiefly to the north-west, beneath northern Cook Strait. Figure 1 shows an 8' 6-year sample of crustal seismicity (deter­mined by the Seismological Observatory, Wellington) for the area of interest. The sample extends from 1 January 1964 to 30 August 1972, and thus includes the period of the field studies reported here. The epicentral pattern is a representative one inasmuch as it illustrates a diffuse distribution similar to that seen in compila­tions for the earlier years 1942-53 (Eiby 1964) and for 1955-66 (Adams & Hatherton 1973).

Figur~ 1 illustrates the map distribution of earthquake locations restricted to depths of 12 km (12R) and 33 km (33R), respectively, where earthquakes in both groups are assumed to occur within a 33-km-thick continental crust (see New Zealand Seismological Re­ports). Their patterns do not significantly differ, but 12R events are clearly more numerous in the sample;

*Present address: Department of Geology and Geophysics, University of Utah, Salt Lake City, Utah 84112, U.S.A.

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142 N.Z. JOURNAL OF GEOLOGY AND GEOPHYSICS, VOL. 23, 1980

MAG.

• 5 - 5.9

• 4 - 4.9

• ~ 3.9

DEPTH

• 12 R

0 33 R

Long.

• • 0 •

0

•• • •

O •

(5)

0 0-

0

• •

•

50km !

• •

FIG. I-Shallow seismlCJty (depth « 33 km), 1 January 1964 to 30 August 1972, determined by the Seismolo­gical Observatory, Wellington, for the southern North Island, New Zealand. WEL, MNG, and CAZ are stations of the New Zealand seismograph network. Stars indicate historical earthquakes of magnitude 7 or greater. Fault pattern adapted from the "Geological Map of New Zealand 1 :250 000".

they make up 70% of the 225 crustal earthquakes. For the same area and for the time period 1955-63, 76% of 190 crustal earthquakes were assigned a "shallow" depth, an earlier designation for upper crustal earth­quakes that is roughly equivalent to the assignment of 12R depth. The data of this report and those of Robinson (1978) suggest that this apparent predo-nin­ance in the historical catalogue of upper crustal over lower crustal seismicity in the southern North Island is probably erroneous.

Because: less than 15% of the epicentres in Fig. 1 have an estimated accuracy better than 10 km, the reader is cautioned against interpreting fine detail in their spatial pattern. Indeed, the limitations of this type of compilation, particularly for correlation with mapped geology, are demonstrated in this paper.

Temporal clustering of epicentres in Fig. 1 is mini­mal. There are only three cases where more than three locatable shocks occurred within 50 km of each other during any 30-day period. A magnitude 5·2 earthquake on 21 September 1967 (40·8°S, 175'1°E) was followed by six close events, and another magnitude 5·2 earth­quake on 11 November 1967 (40·9°S, 175·8°E) was followed by nine close events. Four earthquakes, in­cluding shocks of magnitude 5·3 and 5·0 occurred near 40·3°S latitude and 175·4°E longitude on 18 December

1970, and were preceded by two events within 20 km on 5 December 1970.

Several crustal shocks in the magnitude 5 range, but none larger, occurred during the time sample of Fig. 1. Since 1940, only one shock of magnitude 6 or greater has occurred in this area, the magnitude 6·0 Wairarapa earthquake of 2 December 1942, located at 41'I°S, 175·7°E, with a focal depth of 10-15 km (see Gibowicz 1973a). Just outside the area of Fig. 1, shallow earth­quakes in the magnitude 6 range occurred in 1951 (40·2°S, 177·0 0 E), in 1959 (41·0 0 S, 174·2°E), in 1961 (41·7°S, 176·7°E), and in 1977 (41·7°S, 174·5°E).

Also shown in Fig. 1 are the epicentres of four de­structive earthquakes of magnitude 7 or greater that have occurred in the southern North Island since 1843 (Eiby 1972). The 1855 earthquake, presumed to be of magnitude 8, is notable for its accompanying surface faulting, which extended more than 100 km to the north-east of the epicentre (Ongley 1943). In contrast no surface faulting accompanied the magnitude 71 Pahiatua earthquake of 1934 (Hayes "937), and only short surface breaks over a distance of 2 km were found after the magnitude 7·0 Wairarapa earthquake of 24 June 1942 (Ongley 1943 )-but these breaks may not have been fault related (Lensen 1968). Both the

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ARABASZ & LOWRY - TARARUA-WAIRARAPA SEISMICITY 143

Pahiatua earthquake and the June 1942 Wairarapa earthquake occurred at what would presumably be crustal depths (Gibowicz 1973a). On the other hand, the magnitude 7' 1 Wairarapa earthquake of 1 August 1942 had a focal depth of approximately 55 km (Eiby 1968), as determined from PKP observations (G. A. Eiby pers. comm.). An understanding of depth variations in seismicity in this area is clearly important to evalua· ting fault activity and seismic hazard.

MICROEARTHQUAKE RECONNAISSANCE

In a preliminary study, microearthquakes were sys­rematically sampled at 15 sites in the southern North Island between May and September 1971, representing the first micro earthquake study in the area. Rates of occurrence (Table 1, sites a-o) can be directly com­pared with those from a survey of the northern South Island (Arabasz & Robinson 1976) where identical instruments and techniques were used. The events-per­day statistics based on S-P intervals of 3' 0 s or less, apply to events originating within about 24 km radial distance of a site.

Figure 2 summarises results of the 1971 reconnaissance. In general, the micro earthquake rates indicated low to moderate microseismicity shallower than 24 km. The highest rate of 4· 9 events per day (site d) is com­parable to the highest rates determined in the vicinity of the Marlborough faults, but is low, for example, compared with 30·4 events per day measured in the Inangahua aftershock zone, 3·6 years after the magni­tude 7'1 main shock (Arabasz & Robinson 1976).

Relatively higher rates of activity were found at the latitude of Masterton in the vicinity of West Wairarapa Fault (Fig. 2), the major active fault that broke in 1855. Recordings at three sites occupied during 1972 in the same general area of sites band d yielded slightly lower rates of shallow activity (see Table 1, W AI, HOL, TEM). The counts of microearthquakes are particularly significant in this respect. More events were systematic­ally recorded in the 3-6 s S-P range than in the 0-3 s range (see Table 1), despite greater instrumental sensitivity to closer events. Factors other than attenua­tion and the spatial location of foci, such as instrumen­tal frequency response, can affect S-P distributions (e.g., Asada 1957), but a likely explanation for the greater frequency of 3-6 s S-P intervals was suggested by the focal depths of 25 microearthquakes of magnitude l' 9-2' 9 that were located during the reconnaissance. The more reliable focal depths (plotted as larger circles in Fig. 2) were computed with good depth control and indicated substantial activity in the 25-40 km depth range, below the level sampled by the 3' 0 s S-P hemi­spheres of 24-km radius.

A preliminary composite focal mechanism for micro­earthquakes located during the 1971 reconnaissance implied normal faulting with sinistral slip on a north­east-trending modal plane, at variance with the regional geology. Dextral transcurrent slip would be expected on such a plane. More detailed results from the 1972 follow-up study clarified this discrepancy by indicating that a change in stress orientation occurs below about 20 km. Extension had to be predominant within the 25-10 km depth range during the 1971 observation period.

TABLE I-Sample rates of microearthquake occurrence.

Site lat. S long. E

a. Wrights Hill 41° 17.8' 174° 44.3' b. Waiohine Valley 41° 01.8' 175° 23.9' c. Burnside Farm 40° 48.5' 175° 52.6' d. Mikimiki Valley 40° 50.7' 175° 33.8' e. Belmont Hill 41° 09.9' 174° 53.2'

f. G1enva1e Farm 40° 54.9' 175° 07.3' g. Mangahao 40° 37.1' 175° 28.9' h. Cross Creek 41° 10.1 ' 175° 14.1 ' i. Papatahi Stream 41° 20.9' 175° 03.8' j. Blue Rock Stream 41° 19.9 ' 175° 23.5'

k. Hastwells 40° 42.3' 175° 42.7' 1. Puketotara 40° 59.3' 175° 46.4' m. Huru 40° 39.8' 175° 42.7' n. Marima 40° 30.4' 175° 40.2' o. Jury Hill 41° 07.8' 175° 31.0'

WAr, Waiohine Valley 41 ° 01.6 ' 175° 24.0' HOl, Mt Holdsworth 40° 54.1 ' 175° 28.8' TEM, Te Mara Bush 40° 46.7' 175° 34.6'

• Based 011 time sample of less than 48 hr.

Start date

(1971 )

17 MAY 31 MAY 31 MAY 1 JUN

10 JUN

11 JUN 28 JUN 3 AUG 4 AUG 4 AUG

25 SEP 25 SEP 4 OCT 5 OCT

18 OCT

--(1972)--

14 AUG 14 AUG 14 AUG

Sample time

(hrs)

474.1 133.0 68.2 49.5 58.8

69.5 89.9 72.4

104.5 96.6

12.5 20.8 46.9 35.8 49.5

256.5 180.4 227.0

Atten. (dB)

24,30,36 30,36,42 36,42 30,36 36,42

36 36 36,42 36 36

36,42 42 36,42 42 36

36 30,36 30,36

No. of events recorded S-P(s)

< 3~0 3-6 >6.0

9 42 165 13 18 38 2 11 8 8 12 30 3 10 13

1 10 14 0 16 43 1 4 18 3 2 22 3 9 36

2 5 8 0 5 7 2 5 12 0 0 13 3 4 12

19 64 130 15 66 196 8 94 200

Events/day (S1 ~.3.0s) §

Obs., Norm.

0.5 0.6 2.3 4.3 0.7 0.7 3.9 4.9 1.2 4.1

0.3 0.7 0 0

0.3 0.9 0.7 0.9 0.8 1.5

3.8 3.8* 0 0*

1.0 2.0* 0 0*

1.5 2.9

1.7 2.5 2.0 2.0 0.8 1.1-

t Observed number: simply takes into account the number of de tee ted shocks. regardless of attenuata! setting or amplitude. § Normalised number: adjusted to attenuator setting of 30 dB (see text).

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144

EVENTS/DAY

A 3.0- 5.0

.A.. 1.0 - 3.0

... < 1.0

COOK STRAIT

N.z. JOURNAL OF GEOLOGY AND GEOPHYSICS, VOL. 23, 1980

o I

PACIFIC OCEAN

50km [

FIG. 2-Reference map of summary of results of 1971 microearthquake reconnaissance of southern North Island. Solid triangles indicate normalised counts of microearthquakes per day with S-P < 3.0 s, as in Table 1; open triangles indicate short time samples of less than 48 hr. Each numbered circle indicates the epicentre and focal depth (km) of a microearthquake (ML ;;;. 3·0) located during the survey; large circles imply reliable focal depths; small circles, approximate depths. Faults as in Fig. 1.

TABLE 2-Station data for Tararua-Wairarapa microearthquake study, August 1972.

Station lat. S long. E Elevation Dates Contribution t (~50m) (1972) (%)

MAN 40° 43.55' 175° 15.98' 250 14-19 AUG 57

OTK 40° 51.86' 175° 13.90' 150 14-19 AUG 31

TEM 40° 46.68' 175° 34.59' 450 14-24 AUG 92

HOl 40° 54.10 ' 175° 28.84' 350 14-24 AUG 97

WAI 41° 01.60' 175° 24.04' 150 14-24 AUG 82

RYE 40° 55.39' 175° 35.51 ' 200 15-17 AUG 11

BUC 41° 04.17' 175° 41.57' 150 18-24 AUG 54

TRF 41° 00.12' 175° 38.91' 100 21-24 AUG 6

JRY 41° 07.83' 175° 30.97' 250 21-24 AUG 13 MNG* 40° 37.12' 175° 28.92' 400 14-24 AUG 41 WEL* 41 ° 17.17 ' 174° 46.10' 100 14-24 AUG 6 CAZ* 40° 54.25 176° 13.57' 50 14-24 AUG 2

• Permanent station of New Zealand seismograph network. t Percentage of 142 hypocentre. for which a P·wave arrival WIJ used.

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ARABASZ & LOWRY - TARARUA-WAIRARAPA SEISMICITY 145

DETAILED MICROSEISMICITY:

TARARUA-W AIRARAPA AREA

Field Experiment and Data Analysis During the period 14-24 August 1972, six portable

high-gain seismographs were operated on the flanks of Tararua Range and in the Masterton-Lake Wairarapa Depression (Kingma 1967). A changing network-incor­porating seismograph stations MNG, WEL, and CAZ (Fig. 1) of the New Zealand seismograph network­involved a total of 12 recording sites, as indicated in Fig. 3 and Table 2. The portable seismographs have been described by Arabasz & Robinson (1976); all were operated with vertical-component I-Hz seismo­meters. Drum speeds of 1 mm/s were used, and all clocks were calibrated daily with radio time signals to assure a timing accuracy of ± O· 1 s.

During 11 complete days that the Tararua-Wairarapa network was in operation, more than 550 earthquakes were recorded, of which 142 were located. Numerous events were clearly outside rather than beneath the network. Consequently, limited S-P ranges were search­ed for locatable events such that monitoring beneath

the network systematically extends only to a depth of about 60 km. At least 75% of all the detected shocks that could have originated above this level and in the vicinity of the network defined by the temporary stations were located.

Hypocentres were computed with an iterative, least­squares location program described by Robinson et al. (1975) that accounts for both direct and critically­refracted arrivals. All but 11 small earthquakes were located using a minimum of 6 phases, including at least one reliable S arrival; an average of 8-9 phases were used. We conservatively accepted free-depth solutions only if the focal depth was less than the epicentral distance to the closest station; otherwise the focal depth was restricted to one of a range of trial depths. The standard New Zealand crustal model con­sisting of two layers, with layer thicknesses of 12 and 21 km, over a half space, and respective P velocities of 5' 5, 6· 5, and 8' 1 km/s was adopted. A mean P to S velocity ratio of 1· 73 determined from the 1972 data set was applied. A complete list of the located earthquakes is on file at the Seismological Observatory, Wellington.

Long. 175°00'E 30' 176°00'E 30' .. ___ IIIIii ... __ ... __ ... ____ .... iliillllll ___ ~ ... Lat.

MAG_

• 4.2

• 2- 2.9

• I - 1.9

• 0- 0.9

0 I

50 km I

/

FIG. 3-Epicentres for earthquakes with depths shallower than 20 km, 14-24 August 1972. Solid circles represent well-located earthquakes; open circles, less reliable ones (see text). Crosses indicate seismographic sites, as in Table 2. Traces of the Wellington, West Wairarapa, and Carterton Faults (from Fig. 1) shown for reference.

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146 N.Z. JOURNAL OF GEOLOGY AND GEOPHYSICS, VOL. 23, 1980

175°00'E 30' 176°00'E 30' •• __ ._..-_..,. ___ ';-ip-iiliii.,.-_..,._. Lat.

MAG.

• 3 - 3.5

• 2- 2.9

• I - 1.9

• 0-0.9 ML =5.5 h=22km

30 JAN 1976 o ML =5.1

h=33R 030 JULY 1972

FIG. 4-Epicentres for earthquakes with depths equal to 20 km or more, 14-24 August 1972. Solid cjrcles represent well·located earthquakes; open circles, less reliable ones (see text). Epicentres of two magmtu~e 5 earthquakes discussed in text shown by stars. Regions enclosed by dashed lines were sampled for composIte focal mechanism (Fig. 6). Line X-X' keyed to Fig. 10. Major axes of the ellipse labelled A-A' and B-B' keyed to Fig. 5. Crosses and faults as in Fig. 3.

To estimate the precision or relative accuracy of the earthquake locations, the data have subsequently been processed with the computer program HYPOELLIPSE (Lahr 1979). The hyprocentres computed with HYPO­ELLIPSE are in excellent agreement with our original locations, which are used in this paper. For example, the mean difference for the unrestricted focal depths is 0·1 ± 1· 2 km, and differences for the well located epicentres (solid circles, Figs 3-4) are less l' 0 km.

For each hypocentre, HYPOELLIPSE computes a one-standard·deviation confidence ellipsoid specified by ERH, the major semi·axis of the ellipsoid's horizontal projection, and ERZ, the largest vertical deviation within the ellipsoid measured from its centre. Reading error of 0·1 s was assumed. Epicentres plotted as solid circles (Figs 3, 4) have a mean value of ERH := 2' 3 ± 1· 3 km; for epicentres plotted as open circles ERH := 3' 4 ± l' 6 km. As usual, epicentres far outside a station network are unreliable.

The precision of focal depths is more critical to this study than the precision of the horizontal coordinates. Free depth solutions plotted as solid circles in the cross­sections (Figs 5A-B, 9) have a mean value of ERZ :=

2·1 ± o· 9 km. The restricted· depth hypo centres plot-

ted as open circles in the cross-sections are variable in quality; those beneath the area of the stations, (e.g., Figs 5A-B) have depths that approximate free depths computed with HYOELLIPSE, with a mean difference of o· 6 ± 2' 2 km (hence we estimate ERZ := ± 2-5 km); those outside the network should be considered as having depths less reliable than ± 5 krn.

Magnitudes were systematically determined from measurements of trace amplitude and signal duration to provide estimates of ML (see Robinson et al. 1975), calibrated with respect to Wood-Anderson seismographs at Wellington. More recent studies indicate that such magnitude scales extrapolated below about ML := l' 5 are unreliable without extraordinary calibration (Bakun & Lindh 1977; Suteau & Whitcomb 1979). Thus sym­bols in various figures of this paper that indicate smaller magnitudes imply only relatively small size.

Earthquake Locations

Epicentres are plotted in Fig. 3 and Fig. 4 for foci shallower and deeper than 20 km respectively. Epi­centres for 29 earthquakes shallower than 20 km shown in Fig. 3 include a magnitude 4· 2 shock, 8 km deep,

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ARABASZ & LOWRY - TARARUA-WAIRARAPA SEISMICITY 147

A A'

10 f- o 00 0° • • .... - 10

o

• 20 t-- - 20

o c

• 30 .. 30

• • o

.40 r- - .40

HORIZONTAL SCALE = VERTICAL SCALE

FIG. 5 - A, B - Vertical cross-sections of seismi­city located during field study of 14-24 August 1972, including all earth­quakes shown in Figs 3 and 4 beneath the elliptical region with major axes A-A', B-B' (Fig. 4). Solid circles represent hypocentres located with free depths to an esti­mated vertical accuracy of ± 2-3 km; open circles, restricted locations with a vertical accuracy of ± 2-5 km (see text). C­Frequency histogram of earthquake depths com­puted from August 1972 study.

~

B

10 f-

20 -

30

40 f-

0

4!,

•

MAG.

0 4 .2

A

0 2 .0 - 2 .4

o 1.5-1.9 o 1.0-1..4 o 0.5 - 0.9

o <0.5

10

~ :I: 20 .... a. w C

30 ... .. u o ... 40

50

0° • • •• o

• .. • 0 .. , ..

• ••••• 0 • • •

• • 1'00 -o· •

NO. OF EARTHOUAKES

20

B'

o 0 .... 10

.0 0 - 20 0 o "'i-

~ O~ 0 CP • •• o· • • .0 30

- 40

30

MAG.

0 42

c

B

0 2 .0 - 2 .4 01.5-1.9 o 1.0-1.4 o 05-0.9

o <0.5

HORIZONTAL SCALE = VERTICAL SCALE

that occurred on 15 August, 1 day after our network was installed, and 9 related events. Two foreshocks and six of its aftershocks clustered east of Carterton (inset, Fig. 3). Our location for the main shock was solved using arrival times at five temporary stations west of the West Wairarapa Fault, together with arri­vals at MNG, WEL, and CAZ. Using the latter three stations plus others of the standard seismograph net­work, the Seismological Observatory determined an epicentre 2 km north-north-east of ours and assigned a restricted depth of 12 km.

The clustered epicentres east of Carterton, away from West Wairarapa Fault, were located by different sub­sets of our stations, with and without azimuthal con­trol from the south-east at BUC. The relation of this seismicity to active faulting becomes clearer when the focal mechanisms are considered later.

In general, Fig. 3 shows scattered seismicity on the eastern side of Tararua Range extending into the Mas-

terton-Lake Wairarapa Depression. On the western side of the Tararuas, shallow earthquakes were located only offshore. Perhaps the most notable aspect of the shallow seismicity is the small number of events in the sample, even counting an aftershock series, in contrast to the number of microearthquakes deeper than 20 km (Fig. 4).

fhe number of recorded micro aftershocks following the magnitude 4· 2 earthquake on 15 August was an order of magnitude lower than that expected for a main shock of this size (e.g., Arabasz & Robinson 1976). The average magnitude of 4· 2 assigned to the main shock by the Seismological Observatory may have been overestimated by a half magnitude unit Of more, as suggested by measurements of signal duration on tl).e portable seismographs. Alternatively, a very short-lived aftershock sequence could have been due to an unusually low stress drop or an unusually large fault area (Gibowicz 1973b).

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148 N.Z. JOURNAL OF GEOLOGY AND GEOPHYSICS, VOL. 23, 1980

¢= 350° a=66°SW N

• cP

o

•

~ 0 o

T

o 6.

h = 7.0-18.0 km

¢= 03° N a = 38°E

• o •

h=24.0-36.0 km

•

A

o

c

¢ = 60° a = 67°NW

x

T

¢ = 352° N B a = 48°E

¢ = 58° a = 66°NW

-.0 1> 0

'000 • ~ CP '"

(~ .-1 0 0 0 0

• • «> -t . . ~ 00 x 0./ . • 0.." T

.", 0

h= 19.8-23.9 km

N ¢ = 26° o a = 60 0 SE

T ¢= 60° a = 35°NW

h=30.0-40.0 km

FIG. 6-Composite focal mechanism for subsets of microearthquakes (see text and Table 3) located during August 1972. Diagrams are equal-area projections of the upper focal hemisphere. Open symbols (circles, squares, tri­angles) represent dilational P-wave first motions; solid symbols, compressions. All symbols, except squares, are associated with directly upward-refracted rays; squares = critically refracted arrivals projected from the lower hemisphere. Triangles represent arrivals that have a nodal character and a determined first motion; crosses represent arrivals that have a nodal character, but the direction of first motion is not clear. Smaller symbols are less reliable, cp and 8 are the strike and dip respectively of the nodal planes. P and T refer to the compressional axis and the tensional axis respectively.

During the August 1972 field experiment, micro­earthquakes deeper than 20 km were three to four times more abundant than those above 20 km in the upper crust. The deeper seismicity is broadly scattered, as shown in Fig. 4, and chiefly occurs in the 20-40 km

depth range. (The deepest earthquake located· was close to Wellington at a depth of 47 km.) An ellipse with dimensions 70 km X 40 km outlines an area of con­centrated, well-located activity beneath the temporary network. Foci beneath the ellipse (including those

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ARABASZ & LOWRY - TARARUA-WAIRARAPA SEISMICITY 149

shallower than 20 km) were projected onto vertical sections coincident with the major and minor axes of the ellipse, as shown in Fig. SA-B.

A striking feature of the cross-sections is a concen­tration of earthquakes in a layer between the depths of 20 and 35 km, with earthquakes in the upper crust occurring at a discrete higher level. Maximum focal depths appear to gradually deepen to the south-west in section B-B' (Fig. SB). A frequency histogram in Fig. sC summarises focal-depth information for all 142 located earthquakes. Recall that foci plotted as solid circles in Fig. SA-B were computed with free-depth solutions and should have focal depths accurate to . ± 2-3 km; open circles represent restricted-depth hypo­centres with uncertainties in focal depth of ± 2-5 km. The absolute accuracies of the earthquake locations are impossible to evaluate; they depend, for example, on the validity of the velocity model used. Nevertheless, the fortuitous distribution of deep earthquakes, which resulted in :very good focal depth control, and the proven reasonableness of the standard New Zealand velocity model down to at least 20 km, suggest that the spatial distribution of earthquakes located in this study is reliable.

Composite Focal Mechanisms

Using procedures described by Arabasz & Robinson (1976), composite focal mechanisms were determined for four groups of well-located earthquakes from the August 1972 study (Figs 6A-D). Essential data for each mechanism are listed in Table 3. Published focal mechanisms for North Island shallow earthquakes in­clude: single-event mechanisms for the 1966 Gisborne earthquake (Johnson & Molnar 1972) and the 1974 Opunake earthquake (Robinson et ai. 1976), and two composite focal mechanisms for the southern Wai­farapa-Palliser Bay area (Robinson 1978).

The cluster of shallow earthquakes near Carterton (inset, Fig. 3) provided a good restricted grouping

for a composite mechanism (Fig. 6A). Despite incom­plete focal-sphere coverage, nodal information constrains both nodal planes, and the mechanism is reasonably well determined. It implies either dextral strike slip with a reverse component on a plane striking N73°E, or sinistral strike slip with a reverse component on a plane striking N 10 oW. Faulting clearly does not parallel the N3S ° -50 0E trend of the major West Wairarapa Fault (see Fig. 2), but divergent active faulting has been identified east of West Wairarapa Fault in this area (Kingma 1967; Lensen 1968). The sampled events probably do not lie on a single fault plane. Their occur­rence in an area where the regionally important Carter­ton f'ault (Lensen 1968) would project to intersect West Wairarapa Fault (see Fig. 3) suggests that the nodal plane trending east-north-east may be the more important in the focal mechanism. This nodal plane parallels the trend of Carterton Fault and correctly indicates dextral slip.

In Fig. 6A, both the P axis and the slip vector in the N700E nodal plane (Table 3 )-assuming that to be the fault plane-agree with regional observations made by Arabasz & Robinson (1976) from upper crustal focal mechanisms throughout the northern South Island, that is, north-east/south-west crustal compression and slip parallel to plate convergence in an east-north-east direction.

Foci beneath the 70 km X 40 km ellipse in Fig. 3 that had been located with unrestricted depths were used to determine two focal mechanisms for activity deeper than 20 km (Figs 6B-C). Figure 6C clearly documents that the source mechanisms of these deeper earthquakes are radically different from those in the upper crust and are incompatible with surficial tectonics. Normal faulting predominates on the N600E nodal plane, with sinistral strike slip, and normal faulting with dextral strike slip occurs on the N03°E nodal plane. Considering the spatial scatter of the foci, Fig. 6C is remarkably consistent. Discrepancies can be ex-

TABLE 3-Focal mechanism information.

DEPTH TOTAL P, TAXES SLIP VECTt "1 § "3 §

Group RANGE EVENTS (deg) (deg) (deg) (deg) (km) Az. *_ Pl. Az.* Pl. Az.* Pl. Az.* Pl.

A. 7.0-18.0 P :122 07 080 24 109 13 014 26 T:029 28 343 15 316 02 046 29

B. 19.8-23.9 16 P:l96 49 262 42 219 51 128 02 T:300 12 148 24 175 43 289 23

C. 24.0-36.0 33 P: 190 57 273 52 219 62 316 03 T:309 17 150 23 172 47 300 29

D. 30.0-40.0 10 P:254 69 296 30 277 57 135 29 T:130 14 150 55 208 73 305 02

* Azimuth of downward plunge (PI.). t Slip vector (coincident with pole of auxiliary plane) indicate. direction of movement in fault plane. Two possible slip vectors are associated with each mechanism if selection of fault plane is ambiguous. § Ul and U3 equal greatest and least compressive principal stresses, respectively. In plane of P and T axes, and measured from slip vector. al is 30° in direction of P axis. and U3 is 60° in direction of Taxis.

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150 N.Z. JOURNAL OF GEOLOGY AND GEOPHYSICS, VOL. 23, 1980

plained by slight variations in the attitude of the nodal planes for individual events. As for most composite mechanisms where foci are not sampled from a single fault plane, importance is attached to the basic mechan­ism, rather than to the precise attitude of either nodal plane.

Figure 6C is based on earthquakes 24· 0-36· 0 km deep, many sampled from directly beneath the cluster near Carterton that defined the mechanism of Fig. 6A. To check the depth at which the implied change in stress orientation occurs, earthquakes beneath this same area were sampled from the more restricted depth range of 19'8-23'9 km (Fig. 6B). Although not as consistent as that in Fig. 6C, the mechanism in Fig. 6B indicates that earthquakes as shallow as 20 km exhibit the extensional mechanism found for deeper events. Individual events with some inconsistent data points in Fig. 6B are equally inconsistent with the upper crustal mechanism of Fig. 6A, but they can be explain­ed by the same basic mechanism as Fig. 6C with slightly reoriented nodal planes. The shallowest earth­quake found to be incompatible with the upper crustal mechanism was 19' 8 km deep, with an epicentre only about 10 km distant from the Carterton cluster.

A fourth focal mechanism (Fig. 6D) was determined for a group of events 30' 0-40' 0 km deep that lie beneath the elongate dashed area in Fig. 3, on the north-west periphery of the temporary network. The only nodal plane that is well constrained in Fig. 6D

is the one striking N60oE; however, the basic mechan­ism must be extensional and very similar to those of Fig. 6B and Fig. 6C, indicating that extensional fault­ing below 20 km is regionally widespread. A similar focal mechanism obtained by Robinson (1978) for earthquakes deeper than 20 km beneath the southern Wail'arapa-Palliser Bay region confirms this observation.

Relationship of Short-term Microeflt'thquake Sample to Regional Seismicity

Is the brief sample of 1972 microseismicity repre­sentative of the long-term background seismicity? (1) During the 1971 reconnaissance (Fig. 2), numerous earthquakes in the 20-35 km depth range (also with extensional mechanisms) were located in the same area as that for the 1972 study. (2) Earthquake locations in the southern North Island for 1976-77 (Robinson 1978) indicate abundant seismicity at depths of 20-50 km also in the same area. (3) Intense seismicity located during the August 1972 survey coincides with an area of concentrated historical seismicity (Fig. 1).

Frequency histograms of S-P intervals at station MNG (Fig. 1) suggest a relatively stationary pattern of small-magnitude earthquake activity within about 80 km (S-P < 10 s) of station MNG for several years between 1965 and 1972 (W. J. Arabasz unpub­lished data). A persistent modal peak at 5-6 s S-P closely corresponds with a theoretical peak at station

12 .. ----------------------------~----~------------__,

10

~ o a:: 8 w 0-

en I-z ~ 6 w LL o a:: ~ 4 ~ :;)

z

2

MAG 5.1 SHOCK

30 JULY 23h 39m

I

-1 , SHALLOW SHOCKS NEAR CARTERTON 15 AUG

SEPT 1972 ,.( -.;

FIG. 7-Daily rate of occurrence of earthquakes of MLj 2'3 recorded at station MNG with S-P k. 7'5 s, 2 June-30 September 1972, including time period of field study during 14-24 August 1972. Open tnangle indi­cates one-day record gap.

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ARABASZ & LOWRY - TARARUA-WAIRARAPA SEISMICITY 151

MNG for the 142 earthquakes located during the August 1972 study. This evidence is only suggestive, however. because S-P modal peaks at a single station cannot be uniquely correlated with isolated source regions.

One additional point needs to be discussed. A magni­tude 5·1 shock (h=33R) occurred within the study area at 40·86°S, 175·96°E (see Fig. 4) on 30 July 1972, two weeks prior to the microearthquake study, but there is no evidence of an aftershock sequencel dominating the sample of 14-24 August 1972. Seismo­grams ar station MNG (/::,. < 50 km) were carefully analysed for the four-month period, 2 June-30 Sep­tember 1972, to test for anomalously high activity in' the Tararau-Wairarapa area during our study period. Figure 7 summarises a chronoldgy of activity detected at MNG for shock with S-P < 7·5 sand ML ;;;, 2·3 (for sample completeness).

Figure 7 does not support the idea that seismicity during the August 1972 study period is dominated by aftershocks of the 30 July earthquake. A peak on 15 August includes three events of the shallow earthquake sequence near Carterton that was earlier discussed, and a peak on 2 August-more than 48 hrs after the magni­tude 5·1 shock-includes five events which, having S-P = 3·5-4·5 s, are too close to MNG to be aftershocks of the stronger earthquake on 30 July. Neither of these peaks, then, reflects anomalously high activity 20-40 km deep beneath the Wairarapa area. On the contrary, Fig. 7 indicates a relatively constant background level of earthquake activity during the four-month sample. Significantly, during the three months following the magnitude 5·1 shock of 30 July 1972, only one other earthquake was located by the Seismological Obser­vatory within 50 km-the magnitude 4.2 shock of 15 August 1972 near Carterton. On 30 January 1976 an­other magnitude 5 earthquake (ML = 5·5) occurred within the 1972 study area (see Fig. 4). The earth-

Id "-"'-~~-"'-""----I-~'1

Id

0:: W III :;: ::J Z

W ~ I-<t ....J ::J

10 :;: ::J u

MAGNITUDE, ML

FIG. 8-Cumulative frequency of occurrence versus magnitude for: (A) earthquakes located at depths equal to 20 km or greater during August 1972 study, and (B) earthquakes from four-month sample of Fig. 7. The values of the slope coeffcient b at 95 % con­fidence limits were computed using a maximum­likelihood procedure (e.g., Gibowicz 1973a) applied to non· cumulative magnitude intervals of O· 1.

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152 N.Z. JOURNAL OF GEOLOGY AND GEOPHYSICS, VOL. 23, 1980

quake was located at a depth of 22 km and, in this case also, few aftershocks were detected (Robinson 1978).

Frequency of Occurrence versus Magnitude

Earthquake data from the 1972 Tararua-Wairarapa study were analysed to determine values of the slope coefficient b of the frequency-magnitude relationship. Testing for a significant difference between seismicity above and below 20 km was of particular interest; how­ever, the sparse seismicity above 20 km and the possi­bility of systematic error in determining magnitudes less than 1· 5 precluded a reliable comparison. We examined (1) the sample of all earthquakes located during the August 1972 study that had a focal depth of 20 km or greater, and (2) the four-month sample (June through September 1972) of all earthquakes re­corded at station MNG with S-P ,;;; 7' 5 s. Maximum­likelihood estimates of b for these two samples (Fig. S) indicate values of 1·12 ± 0'30 and 1'11 ± 0·15. These values are comparable to the upper range of mean values (O'SO-l'Ol) determined by Gibowicz (1973a) for the historical Pahiatua and 1942 Wairarapa earthquake sequences, but it must be emphasised that

x

50

E ~

100

CRUST ?

$

t!f+ $

o cPo

o

k---NORTH

/+

b-values are clearly not time invariant (e.g., Gibowicz 1973a; Robinson 1979). The yalues presented here document mean values of seismicity for specific time periods and areas and, in the first case, for a specific seismic volume.

DISCUSSION

The most important observation of the field experi­ments reported here is that of intense s€ismicity be­neath the Tararua-Wairarapa area in the 20-40 km depth range. The regional extent of this seismicity is emphasised by more recent data (Robinson 1978). Arabasz & Robinson (1976) related similar seismicity beneath the Marlborough region to subduction of the Pacific plate and indicated that earthquakes with exten­sional mechanisms were occurring close to a shallow, gently dipping plate interface, either within the upper part of the suducted oceanic lithosphere, or immediate­ly above the plate interface within continental crust. The observed seismicity beneath the Tararua-Wairarapa area can similarly be related to subduction of the Pacific plate.

Figure 9 is a vertical section (keyed to line X-X' of Fig. 4) that is virtually coincident with the plane of

ISLAND--~

50

I

+ " ! I

! ! I

10 , I 100

HORIZONTAL SCALE = VERTICAL SCALE

FIG. 9-North-west/south-east vertical cross-section along line X-X' of Fig. 4 summarising earthquake infor­mation obtained from this study (small, solid, and open circles, as in Figs 5A-B) and from other sources (see text). The large circles are subcrustal foci located by the Seismological Observatory for 1964-72, with stand­ard errors indicated. The uppermost part of the median curve determined by Hamilton & Gale (1969) is also shown. The numbered lines indicate plate-interface profiles (from Karig et al. 1976) for subduction zones in the folJowing regions: (1) eastern Aleutians, (2) Japan (400N), (3) eastern Shikoku, (4) middle America, (5) western Shikoku, (6) South America (23°S), (7) Tonga, (S) central Aleutians, (9) southern New Hebrides, and (10) northern New Hebrides.

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ARABASZ & LOWRY - TARARUA-WAIRARAPA SEISMICITY 153

section D-D' of Hamilton & Gale (1968; 19(9) and the plane of section for the region D of Adams & Ware (1977). Figure 9 includes earthquake foci from Figs 3, 4 that are within 50 km normal distance of the section, together with subcrustal foci located within the same bounds by the Seismological Observatory for 1964-72. Also shown are: (1) offshore bathmetry south­east of the North Island, extending to within several kilometres of the axis of the Hikurangi trench; (2) the standard model of a 33-km-thick New Zealand crust; (3) the uppermost part of the median curve determined by Hamilton & Gale (1969) for subcrustal foci close to this plane; (4) the upper part of 10 plate-interface profiles at circum-Pacific plate margins (from Karig et al. 1976), normalised to the trench axis; and (5) projected nodal planes from the focal mechanisms de­termined in this study.

At convergent plate boundaries, rheological proper­ties of the subducted slab and subjacent mantle, the dip of the down-going slab at the trench, and the load­ing of accreted materials beneath the inner trench slope constrain the geometry of curvature of the down­going slab (e.g., Turcotte et al. 1978; Karig et al. 1976). The superposed plate-interface profiles in Fig. 9, and other lines of reasoning, suggest that the intense seismicity 20--40 km deep beneath the North Island may well define where the plate interface lies. Indeed, where the 100-km isobath of a Benioff zone lies more than 200 km behind a trench, as it does in the southern North Island region, well-located plate interfaces typi­cally have such a gently dipping upper section (pro­files 1-6, Fig. 9).

What about the subcrustai earthquakes 50-100 km deep in Fig. 9? Subcrustal seismicity in this region depicted by Hamilton & Gale (1969) for 1963-65, and by Adams & Ware (1977) for 1971-74, is sparse above 100 km, but abundant below this depth-allow­ing good definition of the deeper Benioff zone, but not its uppermost part. The higher quality data of Adams & Ware (1977) include no earthquakes in this region that are deeper than 50 km trenchward of about the western third of the North Island If the foci 50-100 km deep beneath the North Island in Fig. 9 are accurately located, they conceivably could represent the deeper part of a double Benioff zone reflecting bending stresses within the core of the downgoing plate (e.g .. Engdahl & Scholz 1977). Such double seismic zones merge at depths of about 175 km.

Do the earthquakes 20-40 km deep lie above or below the plate interface? Their extensional focal mechanisms suggest that they occur within the oceanic lithosphere. Such normal faulting within subducted slabs has been invoked as a mechanism to accommodate flexure of the downgoing slab (Lliboutry 1969; Spence 1977). Extensional strain in the upper part of the downgoing slab can variously be explained by elastic bending, by strain propagation through the slab due to episodic locking and subduction (see Spence 1977 for review), by gravitational pulling exerted by the deeper part of the slab, or by some combination of these.

Walcott (1978) shows that the orientations of

geodetic strain axes in the North Island vary with time (implying episodic compression and extension in the cru3t), which he relates to interaction between the New Zealand lithosphere and the subducted Pacific plate across a shallow low-angle boundary. The geodetic data (Walcott 1978) indicate crustal compression in the southern North Island since 1920 with the principal axis of compression at about 115 0

, which is in agree­ment with the single focal mechanism determined here (Fig. 6A, Table 3) for upper crustal earthquakes. If Walcott (1978) is correct in believing that the sub­duction zone beneath the North Island is locked, with compressive strain accumulating in the overlying litho­sphere, then arguments made by Robinson (1978) for the southernmost North Island apply to the Tararua­Wairarapa area. The earthquakes with extensional mechanisms beneath this area (Fig. 9) cannot be occur­ring within the overlying lithosphere, and they must be down-dip from the locked segment of the subducted plate.

ACKNOWLEDGMENTS

R. Martindale, R. Maunder, H. Orr, and R. Robinson of the Geophysics Division, DSIR, participated in the 1972 field experiment. The 1971-72 microearthquake field studies were encouraged by T. Hatherton, DSIR, and were facilitated by F. F. Evison, Victoria Univer­sity of Wellington, who kindly loaned some of the portable seismographs used for the field work.

REFERENCES

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ADAMS, R. D.; WARE, D. E. 1977: Subcrustal earth­quakes beneath New Zealand; locations determined with a laterally inhomogeneous velocity model. N.Z. Journal of Geology and Geophysics 20: 59-83.

ARABAsz, WALTER J.; ROBINSON, RUSSELL 1976: Micro­seismicity and geologic structure in the Marl­borough region, New Zealand. N.Z. Journal of Geology and Geophysics 19: 569-601.

ASADA, T. 1957: Observations of near-by microearth­quakes with ultra sensitive seismometers. Journal of Physics of the Earth 5: 83-113.

BAKUN, W. J.; LINDH, A. G. 1977: Local magnitudes, seismic moments, and coda durations for earth­quakes near Oroville, California. Bulletin of the Seismological Society of America 67: 615--29.

EIBY, G. A. 1964: The New Zealand sub-crustal rift. N.z. Journal of Geology and Geophysics 7: 109-33.

--- 1968: An annotated list of New Zealand earth­quakes, 1460-1965. N.z. Journal of Geology and Geophysics 11: 630-47-.

--- 1972: Space and time trends in New Zealand seismicity. Joltrnal of Geophysical Research 77: 2626-8.

ENGDAHL, E. R.; SCHOLZ, C. H. 1977: A double Beni­off zone beneath the central Aleutians: an unbend­ing of the lithosphere. Geophysical Research Letters 4: 473-6.

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--- 1973b: Stress drop and aftershocks. Bulletin of the Seismological Society of America 63: 1433-46.

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--- 1969: Thickness of the mantle seismic zone be­neath the North Island of New Zealand. Journal of Geophysical Research 74: 1608-13.

HATHERTON, T. 1970: Gravity, seismicity, and tectonics of the North Island, New Zealand. N.Z. Journal of Geology and Geophysics 13: 126-44.

HAYES, R. C. 1937: The Pahiatua earthquake of 1934, March 5: a report on seismological aspects. N.z. Jouma} of Science and Technology 19: 382-8.

HICKS, S. R.; WOODWARD, D. J. 1978: Gravity models of the Wairarapa region, New Zealand. N.z. Journal of Geology and Geophysics 21: 539-44.

JOHNSON, T.; MOLNAR, P. 1972: Focal mec~anisms and plate tectonics of the southwest PaCIfic. Journal of Geophysical Research 77: 5000-33.

KARIG, D. E.; CALDWELL. ]. G.; PARMENTIER, E. M. 1976. Effects of accretion on the geometry of the descending lithosphere. Journal of Geophysical Re­search 81: 6281-91.

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LAHR, J. C. 1979: HYPOELLIPSE: A computer pro­gram for determining local earthquake hypocentral parameters, magnitude, and first motion pattern. U.s. Geological Survey Open-File Report 79-431. 53 p

LENSEN, G. J. 1968: Sheet N158 Masterton (1st. ed.). "Late Quaternary Tectonic Map of New Zealand, 1: 63 360". N.Z. Department of Scientific and Industrial Research, Wellington.

LLIBOUTRY, 1. 1969: Seafloor spreading, continental drift, and lithosphere sinking with an asthenosphere at melting point. Journal of Geophysical Research 74: 6525-40.

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faults recently active. N.Z. Journal of Science and Technology 25B: 67-78.

ROBINSON, R. 1978: Seismicity within a zone of plate convergence-the Wellington region, New Zea­land. Geophysical Journal of the Royal Astro­nomicalSociety 55: 693-702.

--- 1979: Variation of energy release, rate of occur­rence and b-value of earthquakes in the Main Seismic Region, New Zealand. Physics of the Earth and Planetary Interiors 18: 209-20.

ROBINSON, R.; ARABASZ. W. ]. 1975: Micro-earthquakes in the north-west Nelson region, New Zealand. N.z. Jottrnal of Geology and Geophysics 18: 83-92.

ROBINSON, R. ARABASZ, W. J.; EVISON, F. F_ 1975: Long term behaviour of an aftershock sequence: The Inangahua, New Zealand, earthquake of 1968. Geophysical Joumal of the Royal Astronomical Society 41: 37-49.

ROBINSON, R.; CALHAEM, 1. M. THOMSON, A. A. 1976: The Opunake, New Zealand, earthquake of 5 November 1974. N.Z. Journal of Geology and Geophysics 19: 335-45.

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SCHOLZ, C. H.; RYNN, J. M. W.; WEED, R. W.; FROHLICH, C. 1973: Detailed seismicity of the Alpine fault zone and Fiordland region. New Zealand. Geological Society of America Bulletin 84: 3297-316.

SPENCE, W. 1977: The Aleutian arc: tectonic blocks, episodic subduction, strain diffusion, and magma generation. Journal of Geophysical Research 82: 213-30.

SUTEAU, A. M.; WHITCOMB, J. H. 1979: A local earth­quake coda magnitude and its relation to duration, moment Mo, and local Richter magnitude ML. Bulletin of .the Seismological Society of America 69: 353-68.

TURCOTTE, D_ L.; McADOO, D. c.; CALDWELL, J. G. 1978: An elastic-perfectly plastic analysis of the bending of the lithosphere at a trench. T ectono­physics 47: 193-205.

WALCOTT, R. 1. 1978: Geodetic strains and large earth­quakes in the axial tectonic belt of North Island, New Zealand. Journal of Geophysical Research 83: 4419-29.

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