We present seismic images of the structure and elastic properties of a subduction interface in three dimensions, showing its evolution in time. A temporary seismograph array recorded an unambiguous forward-scattered P-P phase from the Hikurangi subduction interface of New Zealand. P-P scattering was most prominent within days of the Feb. 19 1990 Weber I M6 normal-faulting intraplate earthquake below the 20-km-deep interface, and less prominent after three months. S-P phase conversion at the sediment-laden interface was absent although it had been widely prominent earlier. Dilatation of fluid-filled pores by the normal-faulting earthquake caused mineral precipitation by depressurization, sealing thin cracks, making the pores more spherical, and drastically lowering Poisson's ratio in the interface. Our seismic imaging technique using earthquakes for energy sources could locate the P-P ``bright spot'' in 3-d and demonstrate its lack of shear-wave reflectivity. The bright spot arose in a wedge of sediment, 3-5 km thick, bulldozed into the interface by a subducted fault offset. An imaged duplex-thrust complex has accreted some of the sediment to the upper plate, raising a small mountain range above. Activation of this duplex by the temporary sealing and locking of the plate interface below may have triggered the M6 Weber II event.

As the agents of continental formation and accretion, Earth's subduction zones exhibit a wide variety of geometric configurations and appear to have erosional as well as accretional modes. The erosional processes of frontal seamount impact and basal scraping have been inspected by deep seismic-reflectivity imaging (Ranero and von Huene, 2000), and by seismic-velocity tomography (Kodaira et al., 2000). Previous evidence of seamount impacts, scraping, and tunnelling, as well as erosive basal ``megalenses;'' help reveal the properties of subduction zones in an erosive mode. Accretionary processes imaged by these techniques include tectonic underplating (Karig, 1983; Hashimoto and Kimura, 1999; Kopp et al., 1999) and accretion folding (Fisher et al., 1999).

Scraping and seamount pinning may produce stiff barriers or asperities that limit the rupture processes of giant subduction earthquakes (Kodaira et al., 2000). Such barriers may also cause uplift and large earthquakes within both the upper and lower tectonic plates. Within 50 km of the surface, some subduction interfaces may be too weak to produce giant interplate earthquakes (Bilek and Lay, 1998; 1999), with trapped water hydrating mantle minerals to weak serpentinite (Kirby, 2000). If such water is brought down with thick sediments atop the subducting plate it may soften and heat the shallowest parts of the overlying accretionary wedge (Reyners et al., 1999), triggering smaller but frequent and still dangerous intraplate thrust earthquakes. Alternatively water may only be released at great depths, where it ponds against rocks sealed by precipitated minerals (Hyndman, 1988; ANCORP Group, 1999). Aside from forming strong seismic reflectors, the trapped water injects itself into the subducting slab to cause seismicity within the lower plate. Observation of subduction-zone fluids, whether ponded (Hyndman, 1988) or in motion (Aoki et al., 1999) by remote, high-multiplicity geophysical measurements would allow direct constraints on models of both subduction processes and their seismic potential.

We produce, for the first time, a detailed 3-d seismic reflectivity image of a subduction interface that is operating in an accretionary mode. Because our sources of data are intraplate earthquakes of the 1990-1992 Weber sequence (Robinson, 1994), we also get a glimpse of the dynamic processes of fluid entrapment and flow within the interface. The Hikurangi margin east of the North Island of New Zealand (fig. 1) features oblique, west-directed subduction at 4-5 cm/yr (DeMets et al., 1994). Earthquake depths in the lower plate show it is an extremely shallow subduction zone, dipping <4 degrees for up to 200 km west of the plate boundary, where it then dives sharply below 20 km depth (Reyners, 1980; Robinson, 1986; Bannister, 1988). Seismic surveys (Davey et al., 1986; Davey and Smith, 1983; Davey and Stern, 1990; Davey et al., 1997) have confirmed this geometry offshore and revealed that the subducting Pacific plate is anomalously buoyant with 15 km crustal thickness, layered with up to 2.5 km of sediment interrupted by volcanic seamounts. As in the Cascadia subduction zone (Flueh et al., 1998), the buoyant and shallow-dipping slab accretes almost all of its thick sedimentary cover to the leading edge of the upper plate (Davey et al., 1986; Davey and Stern, 1990). The subduction interface usually produces striking conversions of shear (S) waves from events below it to compressional (P) waves (Reyners, 1980; Robinson, 1986; Bannister, 1988; Eberhart-Phillips and Reyners, 1999). The phase conversions and low shear velocities (Reyners et al., 1999) suggest the interface to be unusually rich in sediment and fluid, at least in front of seamounts or other structures on the plate.

Figure 1: Along the east side of New Zealand's North Island the oceanic Pacific tectonic plate subducts westward at 4-5 cm/yr (DeMets et al., 1994) below a continental fragment on the Australian plate at the Hikurangi Trough. The Hikurangi is the southernmost end of the giant and very active Tonga-Kermadec subduction zone, and the plate boundary becomes a transform (the Alpine fault) on the South Island. We examine the properties of the 5-degree NW-dipping subduction zone plate interface where it is buried at 20 km depth, about 150 km inland of the Hikurangi Trough.

Data for reflectivity imaging of the subduction interface arises with the 500 ``passive'' sources of the Weber I and Weber II intraplate earthquake sequences. Shown are Robinson's (1994) lower-hemisphere focal mechanisms for the mainshocks. All the Weber I aftershocks (grey diamonds) were located below the 20 km deep interface, illuminating it from below with forward-scattered seismic energy.

The inset shows example vertical-component seismograms from a small Weber I aftershock recorded at two stations, both in the dilatational quadrant of the normal-fault mechanism. The clear downward first P-wave motions are followed by a simple pulse, then by scattered P-P, P-S, and S-P energy to the onset of the S arrival. The very strong S-P conversions noted by Bannister (1988) and Eberhart-Phillips and Reyners (1999) to the north do not appear in the Weber recordings. The WTA record, along with hundreds of other seismograms passing through the ``bright spot'' area, also shows a clear repeat of the P-wave pulse. The phase labeled Pp is reversed in sign and arises from forward scattering off an interface within a few kilometers of the direct P path, with a scattering coefficient of about -30%. The arrivals recorded at TEU did not pass through the bright spot.

The Weber II event and all but a few of its aftershocks (yellow squares) are above the interface, illuminating it from above with back-scattered energy. Robinson's (1994) deployment of 10 portable seismographs (red triangles) along with nearby permanent seismographs of the New Zealand Network provide the vertical-component seismograms we used for reflectivity imaging. We used all recordings of the 500 quality-A events where the seismograph and the image point were separated by horizontal distances less than 70 km. None of the seismograms showed amplitude clipping by the seismometers or recorders. Fig. 2 shows imaged seismic reflectivity cross sections along the NW-SE orange line, and fig. 3 shows reflectivity volumes below the area outlined by the white square.
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We image the structure of the interface in detail at the Weber sequence (Robinson, 1994) near Cape Turnagain where the plate boundary begins to turn to its most oblique subduction to the south (fig. 1). Our seismic reflectivity images (fig. 2) show how a subducted normal-fault offset of the sea floor entraps a thick wedge of sediment. The continental subduction wedge is then underplated with this sediment through a process of duplex thrusting (Karig, 1983; Hashimoto and Kimura, 1999), effectively accreting oceanic sediment to the bottom of the continental wedge. This process operates where the upper plate is about 20 km thick. Duplexing thickens the plate landward of the thrust stacking (Davey et al., 1986) that extends the very thin upper plate seaward by accretionary outgrowth. Kopp et al. (1999) describe similar accretionary-wedge thickening in the Celebes Sea from thickened low-velocity crust.

Weber is mid-way between the earthquake locations of Bannister (1988) to the north, and of Robinson (1986) to the south. Davey et al. (1986) imaged the plate interface offshore from Cape Turnagain, and Davey and Stern (1990) could trace interface reflections to 14 km depth offshore. Duplex thrusting and wedge thickening by subduction of normal-faulted ``washboard'' oceanic crust was imaged by Ranero and von Huene (2000) along the more steeply dipping subduction zone off Nicaragua. Where Ranero and von Huene (2000) observed seamount subduction, they interpreted erosion and wedge thinning. Reyners et al. (1999), using velocity tomography, saw thickly underplated sediment at 20 km depth pushed in by a high-velocity seamount (as in Kodaira et al., 2000) and interpreted it as accretionary. Our result shows structural details (fig. 2) most consistent with accretion and thickening, but resulting from a subducted normal-fault offset rather than a seamount.

Beyond the structural details, the fact that our sources of seismic energy strongly emit shear waves as well as compressional waves means that our images can reveal the elastic properties of the subduction interface. We find that the plate interface at Weber scatters P waves to P waves much more efficiently than it scatters S waves to S waves, or converts P to S or S to P waves. Long experience in petroleum exploration (Castagna et al., 1985; Pickett, 1963) shows that this type of scattering is characteristic of rocks with increased pore fluid content in contrast to surrounding rocks, where the pores have at least a 0.1 thickness:width aspect ratio (Hyndman, 1988). Previous images of seismic reflectivity in subduction zones (Hyndman, 1988; ANCORP Group, 1999) could not substantially assess shear-wave reflectivity. Chavez-Perez and Louie (1998) used earthquake sources but assessed only P reflectivity of a seismic bright spot in Southern California, imaged as well by Ryberg and Fuis (1998) and interpreted as a dilatational zone of fluid accumulation.

In addition, we find that the strongest P-P scattering, shown in fig. 2b, may be limited to the few days' period following an earthquake in the lower plate. Reyners (1980) had found strong S-P conversions from the plate interface near Weber decades before the 1990 earthquakes, as Bannister (1988) and Eberhart-Phillips and Reyners (1999) currently observe farther north along the Hikurangi. But such conversions do not appear in the 1990 data (fig. 1 shows example seismograms), and are not mentioned by Robinson (1994). This result suggests the 1990 earthquake altered the physical conditions of the interface by dilating the wedge of trapped sediment and temporarily changing the geometry of the pores. Dilation decreased and the properties reverted over the following 3 months. The P-P bright spot is not so prominent when we image it again in fig. 2a from later earthquakes in the upper plate. The rapid dilation of rocks within several kilometers of large earthquakes and the injection of fluid into them over periods of just a few days has been observed through earthquake propagation (Sohn et al., 1998) and geodetic arrays (Aoki et al., 1999). The time scale of a few days for propagation of hydrologic pressure pulses that diffuse away after a few months also fits collected observations of earthquake-triggered, near-surface hydrological events (Muir-Wood and King, 1993).

Inspecting the seismogenic regions of subduction zones, above 100 km depth, has heretofore required a concerted and expensive seismic survey campaign (Davey et al., 1986; Flueh et al., 1998; ANCORP Group, 1999; Fisher et al., 1999; Kopp et al., 1999; Saffer et al., 2000; Ranero and von Huene, 2000). The resulting seismic reflectivity and velocity images can be formed only on 2-d cross sections, because the active explosion or airgun sources and the seismograph receivers are deployed along lines. Our imaging uses earthquakes as sources of seismic energy, which are tightly arrayed above and below the 20 km deep subduction interface (fig. 1). Our seismograph receivers, being on land, are loosely arrayed over a 2-d region about 70 km wide (Robinson, 1994). This geometry allows 3-d reflectivity imaging, akin to medical sonogram imaging, in a cone below the receiver array that narrows with depth (Von Seggern, 1994; Lumley et al., 1994; Bostock and Rondenay, 1999). Using earthquakes instead of controlled sources, we achieve imaging that is inexpensive yet provides 3-d geometries. Cold subducted slabs deeper in the mantle can be located in three dimensions by velocity tomography (Van der Voo et al., 1999; Bunge and Grand, 2000) or from phase conversions (Kosarev et al., 1999). Previous velocity tomography of the Hikurangi margin employed swarms of earthquakes and temporary deployments of seismographs (Bannister, 1988; Reyners et al., 1999).

The first event in the Weber sequence was the M6.2 Weber I in Feb. 1990, which faulted the slab steeply down to the northwest. That was followed 3 months later by the M6.4 Weber II event, which was a southeast-directed thrust entirely above the interface. For imaging we use 214 Weber I aftershocks in the lower plate, the Weber II main shock, and 285 Weber II aftershocks in the upper plate. All hypocenters we used were those most accurately relocated by Robinson (1994). Robinson also found that the data could determine a 1-d velocity model quite well that included a 3% P-wave velocity low within the plate interface, but did not contain 3-d velocity constraints, as the data set from East Cape did (Reyners et al., 1999). Robinson (1994) showed that the aftershocks of the Weber I event were similar normal-fault ruptures and that the Weber II aftershocks were all thrusts similar to that of the mainshock. As well, the relocated aftershock positions are very consistent with the commonly made assumption that aftershocks are distributed over the rupture plane of the main shock.

Our reflectivity imaging technique is derived from the simple 3-d prestack Kirchhoff-summation aftershock migration of Chavez-Perez and Louie (1998). It is similar in effectiveness to the inversion of scattered teleseismic arrivals of Bostock and Rondenay (1999). Two-dimensional prestack migration of active-source survey data has previously allowed imaging of the Hikurangi interface to 14 km depth (Davey and Stern, 1990) and to almost 20 km elsewhere (ANCORP Group, 1999; Ranero and von Huene, 2000). Prestack migration is the only imaging technique meeting the demands of a data set having sources and receivers distributed arbitrarily through three dimensions, with both forward- and back-scattered energy used. The imaging condition for such complexly scattered arrivals is simply the travel time from an earthquake source, to an arbitrary reflector candidate, to a surface seismograph, computed with Robinson's (1994) 1-d velocity model.

To the basic aftershock migration we add amplitude scaling, effectively equalizing source energy, receiver sensitivity, and distance. As the seismograms from the very small quake in fig. 1 show, we are looking at amplitudes of scattered phases relative to those of the direct P and S arrivals. We do not apply the amplitude scaling with depth or time of Dellinger et al. (2000) because we want to emphasize imaging with only the most prominent relative amplitudes and not boost noise. Because our geometric coverage of reflection points is extremely limited by the small number of receiver stations, the Kirchhoff migration draws the edges of reflectors up into elliptical artifacts, very clear on the P-S images of fig. 2. We employ a mild migration-operator antialiasing criterion (Lumley et al., 1994) and a prestack statistical coherency enhancement (Harlan et al., 1994) to mitigate the artifacts.

Tests that imaged synthetic seismograms show this process yields accurate placement of back-scattering structures in a narrowing cone below the receiver array. Forward-scattering structures are detected, although their lateral extent is not well determined and they are smeared out over a 5 km depth range. Analyses of complex elastic-wave synthetics (Operto et al., 2000) show that such a simple migration procedure can accurately recover relative reflectivities. Our imaging condition tracks first-arriving energy and sums in refractions and diffractions. Tests with non-refracting velocity models, and by Operto et al. (2000), show no difference in the results of tracking first-arriving or maximum energy waves.

Figure 2a: Images of back-scattered reflectivity using the Weber II events above the interface. These images are along the subduction-zone dip section line (fig. 1). Blue colors show negative reflectivity, red positive reflectivity, and white low reflectivity in these 3-d prestack Kirchhoff migrations of 2-8 Hz bandpass-filtered seismograms. Overlaying the event locations on the sections leads to an interpretable structure for the plate interface from the back-scattered images.

The P-P reflectivity in particular is strong and coherently shows the structural details of the plate interface. The upward-curving migration artifacts appear in areas with poor reflection-point coverage and can be ignored. The double-reflector interface is 3-5 km thick, and dips 5 degrees NW except where it is offset above the Weber I event by 3-4 km downdip. The frontal ramp of the lower-plate step has raised a thrust duplex (Karig, 1983; Hashimoto and Kimura, 1999) out of the upper 2-3 km of the interface. The frontal ramp of the duplex is visible in the P-P back-scattering reflectivity. A previous duplex, aligned with the frontal ramps, could explain the mechanism of the Weber II event. Fault-bend thrusting above the duplex is another possible explanation of that event.

Each of the images, computed for a separate scattered phase, has had its amplitude scaled individually for display. The rms amplitudes of the P-P and S-S images are similar; the S-P image has about half the amplitude of the P-P reflectivity; and the P-S is an order of magnitude lower. The arcing character of the P-S image is due to poor reconstruction of very little P-S energy.
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Figure 2b: Individually scaled images of different reflectivity components, showing mainly forward-scattering reflectivity from Weber I aftershocks below the plate interface. Superimposed are northeastward views into the mainshock's focal sphere. The refracted nodal plane of the Weber I events is clear between the blue negative and red positive values in the P-P image. We have not corrected for phase shifts around the focal sphere. The back-scattered P-P and S-P reflectivity below the Weber I sequence suggests a Moho in the Pacific plate at 35 km depth. The oceanic lower plate throughout the Hikurangi has been thickened to 15 km total by Chatham Rise plateau basalts or by magmatic underplating (Bannister, 1988; Davey and Stern, 1990).

The P-P forward scattering is stunning, especially considering that an order of magnitude less S-S or P-S forward scattering takes place. The S-P scattering magnitude is about half that of the P-P. To avoid direct arrivals we set a one-second minimum time between image points and the sources. Forward P-P scattering is strongest in a wedge above and just northwest of the Weber I sequence, in the headwall of the NW-dipping normal fault. Scattering of other phases seems strongest at the edges of the wedge. The forward-scattered images suggest high P-wave scattering potential and low S-wave potential on the downdip side of the Weber I fault's offset. Our images cannot resolve the downdip extent of this ``bright spot''. The strong P-P scattering (both forward and back) without strong S-P or S-S scattering suggests substantial porosity variations without proportionate rigidity variations.
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With the exception of the 1990 Weber sequence (Robinson, 1994), most intraplate slab earthquakes recorded on the Hikurangi margin show an S-P conversion at the plate interface, appearing prominently between the P and S arrivals (Reyners, 1980; Robinson, 1986; Bannister, 1988; Eberhart-Phillips and Reyners, 1999). The seismograms in fig. 1 are representative of the type of travel path that had in the past shown a strong arrival at 8 or 9 seconds time, including this part of the margin (Reyners, 1980). Instead, the forward-scattered P-P arrival (noted as Pp on fig. 1), following the P arrival by just 0.2 sec, is prominent on most of the 164 seismograms recorded at station WTA. Very few seismograms suggest possible S-P conversions. Our ``Pp'' phase is similar in origin to the ``Xp'' and ``Xs'' forward-scattered late phases described by Ohkura (2000) from the Philippine Sea plate below Japan. Unlike Ohkura, our forward-scattered S-S image (fig. 2b) shows that we observe a much weaker ``Ss'' or ``Xs'' phase.

The P-P arrival is not prominent on Weber I recordings from other stations such as TEU (fig. 1), where the wave path does not pass through the plate interface at the bright spot just west of the Weber I events (grey diamonds on fig. 1). Fig. 3, at the bottom, shows the location of the bright spot on a horizontal slice through the imaged reflectivity at 21.4 km depth. Tests individually migrating the 164 Weber I seismograms from WTA demonstrate that the bright spot arises from the P-P forward-scattered arrival (such as Pp on fig. 1). The inadequate depth-point coverage of the bright spot is exhibited by the smearing of the reflectivity anomaly 4 km above the plate interface (fig. 3, top and center). Bostock and Rondenay (1999) noted similar problems. However, because Robinson's (1994) temporary seismograph array did record alternate paths from the same events (as in fig. 1), our coverage is sufficient to recognize the imaged bright spot as a true reflectivity anomaly (Operto et al., 2000).

Figure 3: Three-dimensional views of the Weber I forward-scattered P-P reflectivity volume. The top of the volume is 10 km below the surface, extends to 30 km depth, and is 25 km on a side (fig. 1). Warm colors indicate large-magnitude positive or negative reflectivity; cool colors and transparency (in the top view) indicate low reflectivity. The views are all toward the northeast along the strike of the subducting slab, which dips down to the left. Synthetic tests show that the reflective anomaly is artificially stretched above the interface.

The middle view slices the volume along the NW-SE dip section (fig. 1), and reproduces the view of the P-P reflectivity seen in fig. 2b. Because of the distributions of earthquake sources as well as seismograph receivers over regions instead of along lines (fig. 1), imaging quality within most parts of the volume is just as good as in most parts of this section. The bottom view shows that the P-P forward scattering arises in a 10-by-10 km area of the plate interface at 21-22 km depth, west of the Weber I normal fault as defined by Robinson's (1994) aftershock relocations. The location of the bright spot is thus related to the westward motion of subduction thrusting, and not to the northwest dip of the Weber I fault. This relation suggests tectonic erosion of the accretionary wedge (Ranero and von Huene, 2000) and entrapment of reflective sediment by a few kilometers of preexisting offset on the Weber I normal fault.
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The P-P back-scattered reflectivity in fig. 2 suggests a specific geometry for the plate interface at the Weber sequence. Elsewhere both north (Bannister, 1988; Davey et al., 1997; Reyners et al., 1999) and south (Robinson, 1986; Davey and Smith, 1983; Davey and Stern, 1990) along the Hikurangi margin the <5-degree-dipping shallow subducting plate bends down and becomes steeper between 20 and 25 km depth. The bend leads to extension and normal faulting as in the Weber I event (Robinson, 1994). However, because fig. 2a shows a 3-4 km normal offset of the plate interface, the offset must have existed before it was subducted 3-4 million years ago. Volcanic seamounts dominate the Hikurangi plateau east of Cape Turnagain (Davey et al., 1986), but the bathymetry suggests the washboard topography of plate normal faults at some locations.

The preexisting normal-fault offset has protected a 3-5 km section of marine sediment from accretion to the overthrust complex's leading edge (Flueh et al., 1998) and pushed or ``bulldozed'' it to 20 km depth below the North Island. This hypothesis accounts for the plate interface here being thicker than its 1-2 km thickness 300 km to the north (Eberhart-Phillips and Reyners, 1999) and for the lack of a thickened section of high-velocity basaltic crust due to seamount subduction (Kopp et al., 1999; Reyners et al., 1999). In front of the Weber I fault the entire 2-3 km offshore sedimentary section (Davey et al., 1986) is being subducted (as in Saffer et al., 2000). Locally, the 15-km-long Weber I fault offset has formed the frontal ramp of a duplex thrust (fig. 2a; Karig, 1983; Hashimoto and Kimura, 1999) and pushed some of the sediment wedge up to accrete it to the base of the upper plate.

The basal accretion by duplex thrusting may have significantly affected the local topography. Fig. 4 shows how the 1.5-km-high backstop thrust ridges forming the spine of this part of the North Island are interrupted by a water gap at the Manawatu River, which flows westward from the isolated Puketoi Range. The Puketoi Range is southwest of the Weber I fault and conceivably has its origin in the underplating by duplex thrusting 15 km below. Under this hypothesis the topography of the Puketoi Range should have developed within the last 1-2 million years, and should be progressing westward. Although not erosional, this topographic development would be a large-scale parallel to the uplift of accretionary wedges by underthrust seamounts (Reyners et al., 1999; Kodaira et al., 2000; Ranero and von Huene, 2000).

Figure 4: Regional map of topography and bathymetry related to subduction processes on the Hikurangi margin. East of the dotted line the bathymetry shows the thrust ridges and trench-slope basins of the accretionary wedge (Davey et al., 1986; Davey and Stern, 1990). The indentation ``I'' in the shelf off Cape Turnagain suggests erosion of the wedge by seamount impact (Ranero and von Huene, 2000; Kodaira et al., 2000) as is seen northward along the margin off East Cape (Eberhart-Phillips and Reyners, 1999; Reyners et al., 1999). Davey et al. (1986) and Davey and Stern (1990) show that seamounts east of this map on the thick crust of the Hikurangi plateau are currently localizing the development of thrusting in the 2 km sediment section.

We propose that an offset of the subducting plate at the Weber I fault has led to duplex thrusting that has accreted interface material to the upper plate, with the Weber II sequence (yellow squares) manifesting as the frontal ramp of a duplex. The high topography of the Puketoi Range ``R'' shows the local thickening of the upper plate, just west of the Weber I fault, in a scenario similar to that of Reyners et al. (1999) at East Cape. This high allows the Manawatu River ``M'' to breach the summit of the Rimutaka, Tararua, and Ruahine Ranges that form the main thrust backstop ``B'' to the subduction wedge, and flow through them west to the Tasman Sea instead of east to the Pacific.

Satellite and topographic image derived from Eartha^TM Global Explorer, registered trademark of and copyright DeLorme, Yarmouth, Maine.
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The wedge of sediment bulldozed under the margin by the fault offset appears as a reflectivity ``bright spot'' in all of our image sections (fig. 2). ``Bright spots'' (5-20% reflectivity) are prominent in reflection imaging of subduction zones, leading to hypotheses of trapped pore water at lithostatic pressure (e.g., Hyndman, 1988; ANCORP Group, 1999). Bright spots also appear over magma bodies at spreading ridges (e.g., Kent et al., 1993, 2000), and sometimes in continental decollements (e.g., Louie and Clayton, 1987; Louie, 1990; Ryberg and Fuis, 1998). In all those examples except Ryberg and Fuis (1998), the band-limited and relatively high-frequency nature of the artificial seismic survey energy source limited estimation of the sign of bright-spot reflectivity. Using earthquake sources and a 2-8 Hz frequency band, the sign of the forward-scattered P-P phase (Pp in fig. 1) from the bright spot in the headwall of the Weber I normal fault (fig. 3) is unambiguously negative. Robinson's (1994) high-quality gain-ranged digital recordings make this determination possible. Bannister (1988), who used smoked-paper records, could not filter or interpret the sign of the S-P phases he observed. Chavez-Perez and Louie (1998) examined the bright spot in Southern California of Ryberg and Fuis (1998) with earthquake records that were not gain ranged, and could not interpret the sign of its reflectivity.

The strong P-P phase, absence of any significant S-P phase, and the lower P-S, S-P, and S-S reflectivity of the bright spot in the multicomponent images (fig. 2) all show that the bright spot must have an anomalous P-wave but not S-wave velocity. Assuming a realistic small density range, both velocities depend on rigidity but only P velocity depends on the Lame' elastic parameter lambda. Our multicomponent observations thus suggest that the bright spot is a variation in P velocity and lambda, without significant S-velocity or rigidity variation. Further, lambda variations have a special isotropic property, for which the P-wave forward scattering will have the same sign and strength as the P-wave back-scattering and there is no S-wave scattering (Wu and Aki, 1985). Because the P-P reflection coefficient at the bright spot appears to be negative (fig. 1), the plate interface there must have a lower P velocity than its surroundings. Because any S-velocity variations must be much smaller, we are observing a zone of reduced Poisson's ratio.

Most elastic modeling of subduction interfaces (ANCORP Group, 1999; Reyners et al., 1999; Eberhart-Phillips and Reyners, 1999) concludes that free pore water reduces interface rigidity and S velocity and increases Poisson's ratio relative to surrounding rocks. Source durations of subduction-interface earthquakes show quite low rigidity values at depths <50 km (Bilek and Lay, 1998; 1999). However, Hyndman (1988) showed that pore geometry (the thickness:width aspect ratio) can control whether increasing porosity will produce Poisson's-ratio reduction or increase. More spherical pores with aspect ratios of 0.1-1.0 decrease Poisson's ratio while thin, crack-like, pores with ratios <0.03 will increase Poisson's ratio. Hyndman suggested that the sign of Poisson's-ratio variations could not be deduced from rock type and would require shear-wave observations.

Laboratory samples of sandstones and shales have often been known to have Poisson's ratios below 0.25 (Pickett, 1963; Castagna et al., 1985). Poisson's ratio decreases with increasing porosity, due to pore aspect ratios between 0.1 and 1.0. Dense, lithified rocks such as limestones with thin fracture porosity yield Poisson's ratios above 0.25. More recent in-situ studies involving shear-wave well-logging techniques continue to verify such properties in oil reservoirs as deep as 2 km (Jorstad, 1999). Assessment of Poisson's-ratio variations with reflection amplitude-versus-incidence-angle (or versus offset; AVO) have suggested that bright reflectors deeper in the continental crust can appear at both increases and decreases in Poisson's ratio with depth (Louie and Clayton, 1987; Louie, 1990). Hyndman's (1988) models suggest a porosity approaching 5% and a pore aspect ratio of 0.1 to 1.0 (relatively spherical) for the Weber I headwall bright spot, with a clear decrease in Poisson's ratio.

We suggest that the bright spot is a temporary effect of the nearby Weber I rupture, given that no S-P phase was observed (although it was observed near Weber previously by Reyners, 1980), taken together with the unusual reduction in Poisson's ratio and large pore aspect ratio. The bright spot is on the headwall of the Weber I normal fault and above it, so is in the dilatational quadrant of the earthquake's focal sphere. The earthquake occurred on February 19, 1990, and our observations of Weber I seismograms are all within a few days of the magnitude 6.2 mainshock, on February 20-22. Three months later, when we observe the same region using back-scattered P-P reflections from the Weber II events, the bright spot is not nearly as prominent.

Static dilatation of the sediment wedge in the Weber I headwall may have lowered the pore fluid pressure there enough to cause precipitation of minerals out of pore fluids (Hyndman, 1988; ANCORP Group, 1999; Kirby, 2000), sealing cracks and drastically increasing pore aspect ratios. The temporary decrease in Poisson's ratio raised the wedge's S velocity to nearly equal that of the basement rocks bounding the plate interface, preventing the S-P phase conversion (Reyners, 1980; Eberhart-Phillips and Reyners, 1999) usually observed. The simultaneous decrease in P velocity allowed observation of the strong forward-scattered P-P bright spot (figs. 2b and 3). Fluid migration into the wedge over the next three months may have increased fluid pressures gradually, leading to more equal P-P, S-S, and S-P scattering (fig. 2a). We note that the depressurization sealing of cracks and increase in the rigidity of the plate interface would have temporarily locked the main subduction fault immediately after the Feb. 19, 1990 Weber I event. It is conceivable that any aseismic slip deeper on the fault (Robinson, 1994) would have led to the transfer of shear stress around the locked wedge, to trigger the Weber II thrust above.

Correspondence and requests for materials should be addressed to Louie: louie@seismo.unr.edu; http://www.seismo.unr.edu/ftp/pub/louie/weber/.