Download Low seismic-wave speeds and enhanced fluid

Low seismic-wave speeds and enhanced fluid pressure beneath the
Southern Alps of New Zealand
Tim Stern
 School of Earth Sciences, Victoria University of Wellington, P.O. Box 600, Wellington, New Zealand
Stefan Kleffmann* 
David Okaya Department of Earth Sciences, University of Southern California, Los Angeles, California 90089, USA
Martin Scherwath School of Earth Sciences, Victoria University of Wellington, Wellington, New Zealand
Stephen Bannister Institute of Geological and Nuclear Sciences, Lower Hutt, New Zealand
ABSTRACT
A region of low seismic-wave speed is detected beneath the central Southern Alps of
New Zealand on the basis of traveltime delays for both wide-angle reflections and P-waves
from teleseismic events. Respective ray paths for these P-waves are mutually perpendicular, ruling out anisotropy as a cause of the delays. The low-speed region measures about
25 km by 40 km, has a speed reduction of 6%–10%, and is largely above the downward
projection of the Alpine fault. The most likely cause of the low-speed zone is high fluid
pressure due to excess water being released by prograde and strain-induced metamorphism into the lower crust. Because enhanced fluid pressure reduces the work required
for deformation, the existence of the central Southern Alps low-speed zone implies that
this part of the Australian-Pacific plate boundary is relatively weak.
Keywords: fluids, seismic velocity, seismic reflection, Southern Alps, Alpine fault.
faults. It is difficult to steer seismic waves up
through a vertical fault zone or record reflections from a vertical boundary (Hole et al.,
1996). The inclined geometry of the Alpine
fault (Fig. 1C) and the setting of the South
Island as a continental island present a more
favorable opportunity to explore geophysically the internal structure of a continental transform fault. The South Island Geophysical
Transect (SIGHT) program was able, for example, to utilize marine air-gun shots from
both sides of the South Island to image the
subsurface structure beneath the island (Davey
et al., 1998).
INTRODUCTION
Fluids in the crust influence mineralization
(Fyfe et al., 1978), faulting (Rice, 1992), and
crustal strength (Hubbert and Rubey, 1959).
Whereas exhumed ore, quartz veins, and fault
zones give insight into past fluid activity in
the crust, geophysical methods provide evidence of either interconnected fluid or high
fluid pressure in currently active regions of the
crust.
This study presents geophysical evidence
for fluids and high fluid pressure beneath the
Southern Alps of New Zealand. Both the
Southern Alps and the Alpine fault (Fig. 1)
are surface manifestations of the active and
transpressive Pacific-Australian plate boundary in the central South Island of New Zealand (Wellman, 1979). The Alpine fault is exposed on the western margin of the Southern
Alps, but dipping reflectors (Fig. 2D) suggest
that the fault dips southeastward at an angle
of ;408 (Kleffmann et al., 1998). Because the
Alpine fault is a continental transform, similar
to the San Andreas fault (e.g., Hatherton,
1968), this study contributes to and broadens
the debate on the strength of continental transform faults (Lachenbruch and Sass, 1992;
Scholz, 2000; Zoback, 2000).
SEISMIC IMAGING OF TRANSFORM
FAULTS
One of the challenges in seismically imaging faults that are dominantly strike slip, like
the San Andreas, is that they are vertical
*Present address: Department of Earth Sciences,
University of Western Australia, Perth, Australia.
Figure 1. A: Location of shots (SP25–SP27 and BP2–BP7) and Scott’s Hut seismic
array used in this study. Shading indicates generalized geology. B: Plate location
and SIGHT lines 1 and 2, simplified geology for central South Island, and relative
plate vector motion. Refraction shots 25–27 were recorded on 400 seismographs
placed from coast to coast along line 2. C: How crust (Wellman, 1979) and mantle
(Stern et al., 2000) are thought to shorten in plate boundary region.
q 2001 Geological Society of America. For permission to copy, contact Copyright Clearance Center at www.copyright.com or (978) 750-8400.
Geology; August 2001; v. 29; no. 8; p. 679–682; 4 figures.
679
DATA
On the southern, 160-km-long SIGHT line
2, seven large dynamite shots were detonated
and recorded on 400 seismographs placed
coast to coast (Fig. 1B). Onshore seismographs also recorded offshore air-gun shots,
which effectively allowed us to extend shot
gathers to offsets of 150–250 km (Fig. 2A).
Teleseismic signals were recorded on lines 1
and 2 (Stern et al., 2000).
Shot 25, Mid–South Island
Shot 25 in the center of the South Island
(Fig. 1) provides excellent control with nearly
continuous lower crustal and Moho reflections
(PiP and PmP phases, respectively), from vertical incidence out to an offset of 100 km east
of the shot (Fig. 2, A and B). A Pn phase
(refracted wave through the mantle), can be
seen for offsets of 105–150 km. A simple velocity structure for the crust is derived from
these arrivals where the P-wave velocity in the
graywacke and schists varies from 5.0 to 6.1
km/s at 5 km depth, then 6.3 km/s in the lower
crust (Fig. 2D). Below is a 6–9-km-thick layer
of velocity 6.9–7.2 km/s that is thought to be
oceanic crust (Kleffmann et al., 1998).
Shot 27, West Coast
For the same simple velocity model shown
in Figure 2D, calculated PmP and PiP arrivals
for shot 27, west of the Alpine fault (Fig. 1A),
are about 0.8–1.0 s earlier than observed arrivals (Fig. 2C). Structure beneath the shot is
fixed by close-in seismographs, so we isolate
the region for the anomalous delay to the
down-going path from shot 27; the up-going
path is in the region defined by data from shot
25.
Figure 2. A: Data from shot 25 reduced at 6 km/s [time 2 (distance/6)]. Note
strong PiP phase (reflection from lower crust layer) recorded to near zero offset.
Ray paths are shown in D as three arrows at bottom left. Beyond 105 km offset,
we append receiver gather for seismograph located at shotpoint 25 that recorded air-gun shots from an offshore profile. Pn (refraction from mantle) becomes evident beyond offset of 130 km. B: Observed picks for data shown in
A and calculated arrivals for velocity model shown in D. Note that Pn and lower
crustal refraction (Pi) are recorded by receiver gather and appear continuous
with our picks for PmP (reflection from upper mantle) and PiP. C: Shot gather
for refraction shot 27, located west of Alpine fault. Solid lines are calculated
traveltimes for simple velocity model of D. Note that observed Pg fits for offsets
,70 km, then becomes progressively delayed with respect to calculated Pg.
Calculated PmP and PiP phases are too early by as much as 1.0 s to observed
data. D: Crustal structure model and traced rays for shots 25 and 27. Overall
crustal structure from Scherwath et al. (1998). Also shown are rays from shot
27 that bounce off parts of lower crust and mantle defined by shot 25. Shaded
oval region is area where anomalous delays in traveltime are isolated for both
teleseismic waves and wide-angle reflections. Shaded rectangle on surface represents location of seismographs that recorded residual teleseismic delays of
;0.3 s (Stern et al., 2000). Migrated reflection segments are shown that define
subsurface extent of Alpine fault.
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Wide-Angle Reflections From the Eastern
South Island
Wide-angle reflections from a high-resolution reflection experiment located east of the
Southern Alps were recorded on two 48-channel seismograph arrays in the Copland and
Karangarua Valleys west of the Southern Alps
(Fig. 1A). Data from Burkes Pass (BP) shots
2 to 7 (Fig. 3) recorded on the Scott’s Hut
array (Fig. 1), 3.5 km east of the Alpine fault,
show a low-amplitude Pg (direct arrival
through the upper crust), but strong wideangle reflections from the lower crust (PiP
phase) and the Moho (PmP phase). When rays
are traced for these two shots through the simple crustal velocity model of Figure 2D, they
are seen to be 0.8 6 0.1 s early with respect
to the observed data (Fig. 3). Because data
from shot 25 give good control on crustal
structure beneath shots BP2–BP7 east of the
Southern Alps, the part of the ray path to
which we ascribe the 0.8 s delay is, as for shot
GEOLOGY, August 2001
Figure 3. Multichannel, wide-angle reflection records for shots BP2, BP4, BP6, and BP7
(Fig. 1) recorded on 48-channel seismograph at Scott’s Hut in lower Karangarua River (Fig.
1). These data are shown as receiver gather. Offset for first geophone to BP7 is 80.1 km.
Spacing between geophones is 16 m; spread length is 752 m. Near-surface velocity structures beneath both shots and receivers were determined by series of small shots fired into
fixed array. Note that Pg is weak and PiP and PmP are strong. Dashed lines marked Pg*
and PiP* are calculated for simple velocity model of Figure 2D; solid lines marked Pg, PiP,
and PmP are calculated for low-speed model of Figure 4. Pg* is close to prediction, but
predicted PiP* phase is early by 0.8 6 0.1 s on all shot records.
27, the slanted one beneath and west of the
Southern Alps (shaded area of Fig. 2D).
Delayed Teleseismic Events
A 0.3 6 0.1 s delay was observed on lines
1 and 2 (Fig. 1) for three teleseismic events
from the western Pacific (Stern et al., 2000).
Ray paths for these events passed through the
Alpine fault and beneath the Southern Alps
with incidence angles of ;258 (Fig. 2D). The
simplest interpretation of these delays from
the mutually perpendicular teleseismic events
and wide-angle reflections is, therefore, that
there is a low-speed zone in the crust that is
nearly three times longer in the down-dip direction of the Alpine fault zone than perpendicular to it (Fig. 4).
Model Solution
Ray tracing of both wide-angle reflection
and teleseismic data sets was carried out to
obtain a satisfactory fit (Fig. 4). Although the
model is clearly nonunique, control is good as
mutually perpendicular ray paths traverse the
body. Furthermore, the turning Pg (direct refraction through the crust) wave from shot 27
defines the upper limit of the low-speed zone
(Fig. 2, C and D). Beyond a range of 65 km
the misfit for this Pg phase becomes increasingly evident (Fig. 2C), but it is eliminated
when lower speeds are introduced at depths of
5–8 km. Similarly, for the Pg phase from the
GEOLOGY, August 2001
BP2–BP7 shots, with offset ranges ,85 km,
there is only just a perceptible difference
(;0.1 s) between observed Pg and that calculated for a simple velocity model (Fig. 3).
At offsets of 65–85 km, the turning Pg wave
has a maximum penetration of 5–8 km (Figs.
2D and 4), and this is our estimate for depth
to the top of the low-speed zone.
Causes of the Low-Speed Zone
Anisotropy cannot explain the low-speed
zone, because the zone is defined by mutually
perpendicular rays. Fault gouge is not a likely
cause of the observed velocity anomaly, because gouge is typically restricted to the top
10 km of a fault in a region 1–2 km wide
(Wang, 1984). Temperature effects due to advected rocks in the upper crust being hotter
than elsewhere in the crust are also rejected as
a cause of the low velocities. For crustal
rocks, dVp/dT 5 24 3 1024 km·s21·8C21
(Christensen and Mooney, 1995), which, for a
maximum 250 8C crustal temperature (T)
anomaly (Shi et al., 1996) would amount to
only a 0.1 km/s change in P-wave velocity
(Vp).
A 30–300 ohm-m resistivity anomaly has
been discovered beneath the Southern Alps
along SIGHT line 1 (Wannamaker et al.,
1998). When projected 60 km southward onto
the line 2 structure, using the surface trace of
the Alpine fault as a reference line, there is
striking correspondence, in both depth and extent of the low-speed zone, to the core of the
low-resistivity region (Fig. 4). Dipping reflectivity that defines the subsurface extension of
the Alpine fault (Kleffmann et al., 1998) is
also within the low-speed zone (Fig. 4).
Excess fluid under high pressure is therefore proposed as the cause of the low-speed
zone. This interpretation is consistent with
laboratory data which show that enhanced
pore pressure will lower the compressional
wave speed by 10% when pore pressure approaches lithostatic for low-porosity rocks
(Jones and Nur, 1984). Moreover, providing
the pores have an ellipsoid shape and the as-
Figure 4. Crustal structure model showing low-speed region in crust that satisfies wide-angle reflections (shots 27, BP2 to BP7) and teleseismic delays
(Stern et al., 2000). Ray paths for shots 27, 25, and BP2 are shown. Solid lines
are contours of wave speed (km/s). Contours of apparent resistivity are superimposed (40, 100, and 600 ohm-m contours shown).
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pect ratio of the pores is ,0.01, a porosity of
1% is sufficient to produce such a drop in
wave speed, an increase in reflectivity, and a
30 ohm-m resistivity low (Marquis and Hyndman, 1992; Warner, 1990).
DISCUSSION
Our proposed low-speed zone is also detected beneath SIGHT line 1, 60 km to the
northeast of line 2 (Fig. 1), by teleseismic delays (Stern et al., 2000) and by an inversion
of wide-angle reflection and refraction data
(Van Avendonk et al., 1999). This zone appears to be more extensive than those linked
to the central parts of the San Andreas fault
(Mooney and Ginzburg, 1986), or continental
transforms elsewhere (Stern and McBride,
1998). A possible reason is that, unlike in central California, crustal rocks of the central
South Island are being shortened, strained, and
then exhumed at one of the highest (;10 mm/
yr) rates documented (Blythe, 1998). Rocks
will move to greater depths as the crust is
thickened in their approach to the plate boundary (Fig. 1C). This ensures a steady supply of
water in the lower crust from graywacke and
schist rocks undergoing prograde or straininduced metamorphism (Koons et al., 1998).
Having a continually replenished source of
water also accounts for some of the problems
raised by Warner (1990) of retaining water in
the lower crust for periods of 106 yr or more.
One of the consequences of enhanced fluid
pressure in the crust is that there will be a
change in the effective stress, and less work
will be required to deform the crust. The condition for sliding to occur on a preexisting
fracture is (Hubbert and Rubey, 1959)
tf 5 F (Tn 2 Pf),
where tf is the shear traction at failure, F is
the coefficient of friction, Tn is the normal
traction on the fault and Pf is the fluid pressure. An increase in fluid pressure can, therefore, cause a corresponding decrease in the
differential stress required for shear failure on
a fault, or an array of faults making up a fault
zone (Rice, 1992).
Beneath the central Southern Alps, the
maximum depth of earthquakes is reported to
be ;8 km (Leitner et al., 2001), which is
roughly coincident with the top of the lowspeed zone (Fig. 4). Enhanced fluid pressures
could, therefore, be a factor in controlling the
shallow limit of seismicity there. If they are,
the question remains open as to whether strain
release on this part of the plate boundary is
by means of large to great earthquakes (Adams, 1990; Leitner et al., 2001), or by other
means, such as slow earthquakes, as docu-
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mented for other plate boundaries (Heki et al.,
1997; Kanamori and Kikuchi, 1993) where
fluid pressures are also known to be high.
ACKNOWLEDGMENTS
This project was jointly funded by the U.S. National Science Foundation (grant EAR-9418530)
and the New Zealand Science Foundation. We thank
R. Walcott, F. Evison, P. Wannamaker, T. Little, M.
Zoback, and P. Koons for constructive comments.
We are grateful for assistance in the field from T.
Harrison, S. Yuan, A. Crossland, A. McKevilly, T.
Haver, and M. Robertson.
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Manuscript received November 2, 2000
Revised manuscript received April 10, 2001
Manuscript accepted April 24, 2001
Printed in USA
GEOLOGY, August 2001