Geophys. J. Int. (2006) 166, 957–969
doi: 10.1111/j.1365-246X.2006.03040.x
Three-dimensional tomographic imaging of the Taranaki volcanoes,
New Zealand
Steven Sherburn,∗ Robert S. White and Mark Chadwick†
Bullard Laboratories, Department of Earth Sciences, University of Cambridge, Madingley Rise, Madingley Road, Cambridge CB3 0EZ, UK
SUMMARY
3-D models of the P-wave velocity (Vp), the ratio of P- to S-wave velocity (Vp/Vs), and the
P-wave quality factor (Qp) are determined for the crust in Taranaki, New Zealand, using local
tomography and data from a dense network (average spacing 5 km) of 68 three-component,
broad-band seismographs. Vp and Vp/Vs models were determined by jointly inverting P traveltimes and S–P traveltime intervals, and a Qp model by inverting t ∗ (t/Qp) observations derived
from modelling the velocity amplitude spectrum of P wave arrivals. An approximately 5 km
diameter region of high Vp and low Vp/Vs beneath the Taranaki volcanoes extends from the
surface to a depth of about 10 km and images the roots of the volcanoes formed by successive magmatic intrusions. At Mt Taranaki this will probably be the path through which future
magma intrusions will reach the surface. We are unable to image any magma storage within the
upper 16 km of the crust as the volumes involved are probably smaller than the 5 km resolution
of our models. The volcanoes sit in the Taranaki basin, a large sedimentary basin characterized
by low Vp (c. 4 km s−1 ) and high Vp/Vs (≥1.9), attributed to unconsolidated, water-saturated
sediments, and these are underlain by basement rocks with Vp ≥ 5 km s−1 and Vp/Vs ≤ 1.7.
There is a c. 20 per cent contrast in Vp across the Taranaki fault at the eastern boundary of
the Taranaki basin, and a small, shallow, high Vp/Vs (≥2.0) anomaly above the seismically
active Cape Egmont Fault Zone, west of the Taranaki volcanoes. Although the Qp model is
of significantly lower resolution than the Vp and Vp/Vs models, it shows a low Qp anomaly
(c. 100) between 4 and 10 km depth that corresponds to the deepest onshore part of the Taranaki
basin.
Key words: attenuation structure, local earthquake tomography, sedimentary basin, Taranaki
volcanoes, velocity structure.
1 I N T RO D U C T I O N
The study of past eruption deposits at active and potentially active
volcanoes can indicate the size, style and frequency of eruptions, but
it is only rarely that these data can be used to infer anything about
geophysical or geochemical eruption precursors (Sherburn & Nairn
2004). At a volcano with an existing monitoring program, a model
of the subsurface structure of the volcano and surrounding region
will help in understanding physical pre-eruption processes and may
improve the likelihood of successfully predicting eruptions from
those observations. This may be particularly useful at volcanoes with
no historical record of eruptions where the correct interpretation of
precursory activity could be difficult.
∗ Now at: GNS Science, Wairakei Research Centre, Private Bag 2000, Taupo,
New Zealand.
†GNS Science, Avalon Research Centre, PO Box 30-368, Lower Hutt,
New Zealand. E-mail: [email protected].
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2006 The Authors
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Journal compilation Mt Taranaki (also known as Mt Egmont) is a 2518 m high andesite cone-volcano in the western part of the North Island of New
Zealand that last erupted in 1755 (Neall et al. 1986). At this time
New Zealand was only sparsely populated and there is no written
record of the eruption so nothing is known of precursors to this or any
earlier eruptions. Future eruptions will probably be characterized by
dome emplacement and subsequent collapse, by ash eruption, and
by pyroclastic flows and lahars, as these have all occurred during
previous eruptions (Neall et al. 1986). Any eruption would be a significant hazard to the local community as more than 100 000 people
live within 50 km of the volcano, and would have a strong impact
on the local economy, particularly on dairy farming and oil and
gas production (Hull 1996). Mt Taranaki is therefore an excellent
example of a volcano at which an understanding of the subsurface
structure might be helpful in interpretation of monitoring data, and
ultimately in forecasting eruptions.
In this paper we determine, for the crust in the Taranaki region, three-dimensional (3-D) models of P-wave velocity (Vp),
P- to S-wave velocity ratio (Vp/Vs ) and P-wave quality factor
(Qp). These are three fundamental rock properties, and in an area
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GJI Volcanology, geothermics, fluids and rocks
Accepted 2006 April 12. Received 2006 April 10; in original form 2005 August 31
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S. Sherburn, R. S. White and M. Chadwick
et al. 1986); Mt Taranaki first erupted at ca. 120 ka. The Taranaki
Volcanoes have well-defined positive residual gravity anomalies (Mt
Taranaki 350 g.u.; Pouakai 350 g.u.; Kaitake 240 g.u. and Paritutu
170 g.u.), that are believed to be the result of andesite dyke/stock
complexes formed by successive intrusions of dykes into Taranaki
basin sediments over the lifetime of the volcanoes, and to extend
from the surface to at least 6 km depth (Locke et al. 1993; Locke &
Cassidy 1997). Since European settlement of Taranaki in 1841, there
have been no confirmed reports of volcanic activity or unrest from
Mt Taranaki; there are no fumaroles or hot ground on the volcano,
though there is one warm spring (maximum temperature 30◦ C) and
one cold spring with dissolved CO 2 (Allis et al. 1995).
The Taranaki volcanoes sit within the Taranaki basin, a
100 000 km2 Cretaceous–Cenozoic basin on the western margin
of New Zealand. The sediment thickness averages 6 km (King &
Thrasher 1996), and overlies a basement of varying lithology (Mortimer et al. 1997). The eastern boundary of the basin is marked
by the Taranaki fault (Fig. 1), a Miocene reverse fault that offsets basement rocks by about 6 km (King & Thrasher 1996). Just
west of Mt Taranaki is the Cape Egmont fault zone (CEFZ), a
zone of dominantly northeast–southwest trending mainly normal
faults that marks the western limit of deformation associated with
the Pacific–Australian plate boundary in the North Island of New
Zealand (Fig. 1). The zone includes the Cape Egmont fault which
such as Taranaki, where there are sediments, basement rocks, and
volcanic/plutonic rocks, we might expect there to be measurable
changes in these properties that can be used to image geological
structures. Specifically, we examine the expression of Mt Taranaki
and nearby extinct volcanoes in the tomographic models to investigate their subsurface structure, and that of a large sedimentary
basin in which the volcanoes sit. This complements earlier work in
Taranaki that reported the seismicity patterns (Sherburn & White
2005), and focal mechanisms and stress axes (Sherburn & White
2006); together they form part of a larger project whose overall
aim is to understand the structure and background seismicity of the
region and to use this information to help in routine monitoring,
especially in the event of an eruption of Mt Taranaki.
1.1 Previous Work on Mt Taranaki
and the Taranaki Region
PLAT
E
Taranaki lies about 400 km west of the obliquely convergent boundary between the Pacific and Australian plates beneath the North
Island, New Zealand (Fig. 1). Mt Taranaki and the extinct Taranaki
volcanoes, Sugar Loaf-Paritutu, Kaitakei, and Pouakai (Fig. 1),
are dominantly andesitic in composition and represent a south–
southeast migration of activity commencing at 1.74 Ma (Neall
IAN
200 km
NI
Hik
ura
ng
i Tr
PA
en
CIF
ch
IC
PL
ATE
SI
Taranaki Fault
–
65
)
rs
ne
ey
(R
Taranaki
Basin
L
TR
AUS
T
RAL
–
SLP
70
K
–
P
on
e
TRL (Stern)
tF
aul
tZ
MT
Ca
–
pe
Eg
mo
n
–
Taranaki Fault
–
–
0
500
1000
1500
2000
25 km
Elevation (m)
–
173.6˚
173.8˚
174˚
174.2˚
174.4˚
174.6˚
174.8˚
175˚
Figure 1. Map of the Taranaki region showing topography and the 500 m elevation contour surrounding the Taranaki volcanoes, the portable seismograph
network (triangles), and the permanent Taranaki Volcano-Seismic Network (open circles). The Taranaki volcanoes are Mt Taranaki (MT), Pouakai (P), Kaitake
(K), and Sugar Loaf–Paritutu (SLP). The Taranaki-Ruapehu Line (TRL), as proposed by Stern et al. (1987) and (Reyners et al. 2006) are shown. Surface heat
flow contours (in mW m−2 ) show a high heat flow region north of Mt Taranaki (Funnell et al. 1996). The inset shows the location of the study area with the
tectonic plates and plate boundary, North (NI) and South (SI) Islands, the Taranaki basin, and other permanent seismographs from which data were used: the
dashed line marks the approximate position of the Taranaki fault.
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Journal compilation Tomographic imaging of the Taranaki volcanoes
2 D AT A
In 1993, the Taranaki Volcano-Seismic Network (TVSN) began operation to monitor earthquakes in the vicinity of Mt Taranaki so as
to be able to provide near real-time warning of potential eruption
precursors. This network consists of six short-period (1 Hz) seismographs (one three-component and five vertical-component) and
can, with the addition of data from seismographs outside Taranaki,
be used to locate adequately most earthquakes in Taranaki of magnitude ≥2.7 (Sherburn & White 2005). However, with such a small
network there are insufficient data to determine a crustal velocity
model for Taranaki, or to image any 3-D structures. For this a much
larger, denser seismograph network is required.
Data for this study were collected between December 2001 and
September 2002 using 68 portable, three-component, broad-band
(0.03–50 Hz) seismographs (Fig. 1), Guralp CMG-6TD and CMG40T, operated at a sampling frequency of 100 Hz (Sherburn & Allen
2002; Sherburn & White 2005). These were supplemented by data
from the TVSN and from several permanent seismographs in western New Zealand operated by GNS Science. A total of 389 Taranaki
earthquakes were located, using more than 15 000 phase picks, comprising 55 per cent P-phases and 45 per cent S-phases (Sherburn
& White 2005). P-phases were picked only on vertical-component
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Journal compilation seismograms and S-phases only on horizontal-component seismograms, with the first arriving, clearest S-phase picked on unrotated
seismograms. Picks were weighted (0–4) according to their perceived reliability judged on the basis of the signal-to-noise (S/N)
ratio and the estimated uncertainty in pick time. Magnitudes ranged
from 1.4 to 4.0 (Sherburn & White 2005).
3 VP A N D V p /V s T O M O G R A P H Y
Inversions for 3-D Vp and Vp/Vs were made with the computer
program SIMUL2000 (Thurber & Eberhart-Phillips 1999), which
uses a damped, iterative, least-squares algorithm that simultaneously
calculates the velocity model and hypocentral adjustments. We calculate Vp/Vs directly using S–P times rather than calculating Vs
from S arrival times alone because this is the most robust method
of constraining the Vp/Vs ratio (Eberhart-Phillips 1993) and the
Vp/Vs ratio is useful for characterizing rock properties and rheology. The velocity model is specified on a grid of nodes (Thurber &
Eberhart-Phillips 1999). The initial Vp model (Fig. 2) was based on
the minimum 1-D velocity model of Sherburn & White (2005) and
the initial Vp/Vs model was determined from a Wadati plot, and had
a constant value of 1.72 at all nodes.
Vp (km s−1)
2
3
4
5
6
7
8
0
10
Depth (km)
is marked by a 1–5-m-high seafloor scarp and has had up to 0.8 mm
yr−1 of dip-slip movement over the last 225 000 yr (Nodder 1993;
Nicol et al. 2005).
The CEFZ is marked by a 150-km-long lineation of crustal seismicity (Anderson & Webb 1994), with earthquakes confined to the
upper crust (maximum depth ca. 22 km), and many occurring in
earthquake swarms (Sherburn & White 2005). A broad zone of
seismicity in eastern Taranaki marks the Taranaki-Ruapehu Line
(TRL) (Anderson & Webb 1994) which is thought to mark the juxtaposition of a thin crust (25 km) to the north against a normal
thickness crust (36 km) to the south (Stern et al. 1987). Recent tomographic results (Reyners et al. 2006) suggest that this boundary
trends southeast–northwest, parallel to the dip of the subduction
zone beneath the North Island, rather than approximately east–west.
Earthquakes here occur almost exclusively in the lower crust, terminating at ca. 35 km, about the depth of the Moho (Stern et al. 1987;
Sherburn & White 2005). Beneath the summit of Mt Taranaki and
for about 10 km to the east of it, the level of seismicity is significantly
lower than that in both the CEFZ and the TRL, and is restricted to the
upper ca. 10 km of the crust. This is thought to result from a higher
than normal temperature gradient associated with the Taranaki volcanoes (Fig. 1), although the region of shallow seismicity does
not exactly match a region of high heat flow (Allis et al. 1995;
Funnell et al. 1996; Sherburn & White 2005). Earthquakes forming
the Benioff Zone beneath the North Island occur to 250 km depth
beneath southeast Taranaki and beneath northeast Taranaki there are
occasional earthquakes at ca. 600 km depth (Adams & Ferris 1976;
Anderson & Webb 1994; Boddington et al. 2004).
Reyners et al. (2006) have determined detailed 3-D Vp and Vp/Vs
models for a 300 × 200 km area of the central North Island to a
depth of 300 km. The Taranaki region is on the boundary of their
modelled volume and velocities here were calculated on a grid with
a horizontal spacing of 30–40 km and a vertical spacing of 7–10 km.
The node spacing and depth distribution of the seismicity in their
study are such that their resolution in the crust, and of the Taranaki
volcanoes in particular, is very limited. This region is the focus of
our more detailed study.
959
20
30
–
free nodes
fixed nodes
40
–
Figure 2. The initial 3-D Vp model, derived from a minimum 1-D model
(Sherburn & White 2005), that was used as the starting point for the 3-D
tomography. Circles represent the nodes at which the 3-D model was defined,
with open circles at the top and bottom of the model indicating depths at
which all nodes were fixed at the velocity of the initial model, and filled
circles the depths at which the velocity of nodes was free to vary. The initial
model used by SIMUL2000 is a linear interpolation between nodes and gives
a relatively smooth velocity-depth function. The initial Vp/Vs ratio was 1.72,
determined from a Wadati plot.
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S. Sherburn, R. S. White and M. Chadwick
A multistep, graded inversion strategy was used (EberhartPhillips 1993) starting from the initial 1-D model (Fig. 2) and progressing through finer and finer grids. This gives a smooth, but
realistic, 3-D model for poorly-resolved areas and a fine grid in the
well-resolved areas. The final horizontal node spacing was 5 km
in the best-resolved parts of the study area, and the vertical node
spacing was 4–6 km (Fig. 2).
Phase picks were used only from earthquakes that had an azimuth
gap of ≤180◦ , and at least 10 P-phases with an estimated uncertainty
of ≤0.25 s. For the seismically active region west of Mt Taranaki,
at least 15 P-phases were required as too many earthquakes from
this area would have given very uneven resolution. The data set
comprised 178 earthquakes having 4692 P-phases and 3802 S–P
observations. Observations are weighted depending on pick quality,
source–receiver distance and residual size.
Damping was introduced to the inversion to reduce the effect of
non-uniqueness in the solution. It gives a conservative solution with
few artefacts that significantly reduces the data variance without a
significant increase in model complexity.
We inverted jointly for both Vp and Vp/Vs and calculated station
corrections at sites outside the main modelled area (most sites in the
inset in Fig. 1), so that regional velocity variations outside Taranaki
did not significantly affect velocities assigned to nodes within the
modelled area. We did not invert for station corrections within the
modelled area because the high station density gave shallow velocity
models that agree closely with the geology and the final RMS misfit
value for all hypocentres was close to the estimated phase picking
uncertainty in the data. For the 3-D Vp and Vp/Vs models, the data
variance was 24 per cent of that of the initial 1-D model.
4 Qp TO M O G R A P H Y
The computer program SIMUL2000 was also used for the Qp tomography. The Qp model used the same grid of nodes used to determine
Vp and Vp/Vs, but with a larger horizontal node spacing of 10 km
because there are fewer Qp rays. Locations and ray paths were not
perturbed in the Qp inversion. Qs was not determined because there
were insufficient data.
We used t∗ values to invert for Qp using the same procedure as that
used by Eberhart-Phillips & Chadwick (2002) which is based on that
adopted by Haberland & Rietbrock (2001) and Rietbrock (2001).
On an event by event basis we determined the best-fit low-frequency
spectral level, corner frequency ( f c ), and t∗ value from velocity
spectra and a model in which spectral amplitude falls off as f 2 at
high frequencies. The quality of the t∗ estimate is measured by the
fit of the modelled spectrum to the observed spectrum, and is used
to weight the t∗ value in the inversion in a way analogous to that used
with arrival time picks. Although several studies have shown that
Q may be weakly frequency dependent, a frequency-independent Q
was assumed. The use of a frequency-independent Q is not thought
to have a significant effect on the final interpretation of our results as
tomographic models determined using both frequency-dependent Q
and frequency-independent Q have similar features and would be
interpreted in the same way (Lees & Lindley 1994).
Each t∗ observation was calculated using the velocity spectrum
from a 2.56 s window starting at the time of the P arrival. Only
3 per cent of S–P intervals were less than 2.56 s so any contamination of the P-wave spectra by S-wave energy is very minor. A
noise spectrum calculated for the equivalent length window preceding the P arrival was used to determine the S/N ratio of each P wave
spectrum. t∗ was calculated only if the S/N ratio was above 2 over
a frequency range of at least 10 Hz. Spectra were also checked to
ensure that the fit covered frequencies both above and below the calculated f c . An example of fitting t∗ for one event is shown in Fig. 3.
The smoothness of the spectra comes from the multitaper method
used to calculate the spectra and helps in fitting a curve to the data.
A range of problems were found with some earthquake spectra
and they were eliminated from the data set. These included a significant loss of high frequencies that is thought to be due to the analogue
telemetry system at some permanent sites, and large spectral peaks
thought to be due to a resonance in shallow soil layers. Spectra were
also discarded if there were not at least six valid spectra for each
earthquake. This was done to ensure that there were sufficient welldetermined spectra to calculate a robust f c for each event. Within
SIMUL2000, t∗ observations which were close to zero are not used
because they gave an unrealistically high path average Qp > 1500.
The total number of t∗ observations was 678 from 55 earthquakes.
The low S/N ratio of many of the P wave spectra was the main factor
which contributed to the low number of observations. In particular,
there were few t∗ observations from sites in west and southwest
Taranaki (Fig. 4), the area most exposed to strong southwest winds
and sea noise.
To estimate the absolute uncertainty in the t∗ data we compared t∗
values for neighbouring ray paths on a site-by-site basis, comparing
values for all earthquake-pairs that were separated by a distance of
≤3 km. The mean t∗ difference for all data was 0.012 s and this
was taken to be the overall uncertainty in the t∗ data set. When
considered as a percentage uncertainty, to account for the effect of
distance, the uncertainty in our t∗ data is significantly larger than
those estimated by Haberland & Rietbrock (2001) and Rietbrock
(2001) for their data sets, illustrating that the Taranaki data have
relatively large uncertainties.
The initial Qp model used as the starting point for the 3-D inversion was the best-fit half-space model to the data, with a Qp
of 400. The overall RMS misfit with the half-space model was
0.013 s, about the same as the estimated uncertainty in the data of
±0.012 s.
5 M O D E L R E S O LU T I O N
Simple plotting of ray paths shows that the resolution of the Vp and
Vp/Vs models are likely to be similar, and are much better than the
resolution of the Qp model (Fig. 4). To examine the spatial variation of resolution we use the combined results for the spread function and resolution contours (Toomey & Foulger 1989; Michelini &
McEvilly 1991; Eberhart-Phillips 1993; Reyners et al. 1999). The
spread function is a measure of how strong and peaked is the resolution at each node, while the resolution contours show the averaging
contribution from nearby nodes and give the dominant direction of
any smearing.
Spread values and resolution contours are shown in plan view in
Fig. 5, and in east–west cross-sections in Fig. 6, and they confirm
the impression from the ray paths (Fig. 4) that the resolution of Vp
and Vp/Vs models are similar, and, despite the larger grid spacing, that the Qp spread values are larger than those of both Vp and
Vp/Vs.
Those areas with a spread value of ≤2.5 were considered well
resolved as they included most nodes at which the smearing did not
extend beyond adjacent nodes, but did not include many nodes with
greater smearing. For Vp and Vp/Vs this was further constrained
by results from the inversion of synthetic models (Spakman & Nolet 1988; Sudo & Matsumoto 1998; Haslinger et al. 1999). The
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Journal compilation Tomographic imaging of the Taranaki volcanoes
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Figure 3. Example of fitting t∗ for one event; only some of the spectra for this event are shown. The event has a magnitude of ML 3.7, so is an example of data
with a high S/N ratio. Ten seconds of the waveform are shown for each station (5 s before the P arrival and 5 s after). The observed velocity spectrum (solid
line), the noise spectrum (dashed line), and the fit to the velocity spectrum (bold line) are shown. Vertical grey lines mark the corner frequency ( f c ). The station
name, P phase pick information, and pick time are shown above each waveform. Individual t∗ estimates and quality estimates (fit) are also shown.
well-resolved areas are marked in subsequent Vp, Vp/Vs, and Qp
maps and cross-sections and no attempt is made to interpret models
outside those areas. These areas undoubtedly represent a smoothed
view of the well-resolved parts of the models, and there will still be
some artefacts within these areas due to non-uniform resolution.
For Vp and Vp/Vs at 0 km depth, the resolution is strongly limited
by the near vertical ray paths and a consequent lack of crossing rays.
There is little smearing in horizontal x and y directions (Figs 5a
and b), but there is generally strong smearing from the layer below
(Figs 6a and b). However, there is less vertical smearing close to
the Taranaki volcanoes and this area is considered sufficiently well
resolved to be interpreted. At 4 km depth, the well-resolved area
covers almost the whole of the Taranaki Peninsula and extends as
far east as the boundary of the portable network (Figs 5d and e), with
significant smearing only near the coast and at the boundaries of the
network. The resolutions at 10 and 16 km depths (Figs 5g, h, j and k)
are similar to that at 4 km, except for the appearance at 16 km depth
of a small region with east–west smearing beneath and just east of
Mt Taranaki, because the earthquakes here are less than 10 km deep
(Sherburn & White 2005). At 22 km depth, we reach the limit of
the resolution beneath and to the west of the Taranaki volcanoes
(Figs 5m and n) as this is the depth of the deepest earthquakes in
this region (Sherburn & White 2005). In eastern Taranaki the wellresolved region extends ca. 100 km east of Mt Taranaki, but is only
25–30 km wide, reflecting the relatively narrow band of earthquakes
in eastern Taranaki (Sherburn & White 2005). The resolution at
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2006 The Authors, GJI, 166, 957–969
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Journal compilation 28 km depth is further reduced, and displaced east, and in synthetic
tests the amplitudes of the recovered Vp anomalies drop to less than
half the amplitudes of the anomalies in the input model. Nothing
could be resolved at 34 km or deeper, because few rays sampled
these depths (Sherburn & White 2005). The only area where Vp/Vs
resolution is significantly poorer than Vp resolution is in eastern
Taranaki as a result of the absence of three-component seismometers
east of the main study area.
The Qp model resolution is poorer than that of the Vp and Vp/Vs
models and because of it’s larger horizontal node spacing the minimum size of structure that can be imaged is larger.
In constant latitude sections (Fig. 6), smearing is again restricted
to the boundaries of the modelled volume and reflects the ray path
distribution resulting from the occurrence of deep earthquakes in
the east and more shallow earthquakes in the west and particularly
in the centre of the modelled volume (Sherburn & White 2005).
This is particularly clear in a region about the position marked by
x = −10, z = 25 km, where there is smearing upwards and to the
west (Figs 6a and b).
6 TOMOGRAPHIC MODELS
Map views of the Vp, Vp/Vs, and Qp models are shown in Fig. 7
and true-scale east–west cross-sections in Fig. 8.
At 4 km depth there is a strong east–west contrast in Vp that
coincides almost exactly with the Taranaki fault (Fig. 7d). West of
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174˚
-38.5˚
174.5˚
175˚
175.5˚
0 km depth (Figs 8b and e), and which in cross-section appears to
be linked to more widespread low Vp/Vs in the basement rocks.
Qp is generally at or below the half-space value of 400 at 4 km
depth, indicating that the attenuation at this depth is generally higher
than the average for the volume sampled by the Qp rays. The most
prominent Qp feature is a region of relatively high attenuation (Qp
of 100–200) that extends from beneath Mt Taranaki to the east and
then to the south towards the eastern margin of the Taranaki basin
(Fig. 7f).
Nodes at 10 km depth correspond, everywhere except the Manaia
sub-basin (Fig. 1), to basement rocks beneath and east of the
Taranaki basin. Within the Taranaki basin Vp averages 5.5 km s−1
(Fig. 8a), and is lowest (ca. 4.8 km s−1 , −15 to −20 per cent relative
to the initial model) along the southern part of the Taranaki fault
where sediments are thicker and the model at this depth samples the
base of the sedimentary sequence. The southern part of the basin
and particularly the eastern parts close to the Taranaki fault still
have high Vp/Vs (≥1.85), reflecting the thick sediments here. Other
areas west of the Taranaki fault have low Vp/Vs (1.6–1.7) (Figs 7h,
8b) implying that the Vp/Vs of basement rocks is significantly lower
than that of basin sediments. In the eastern and southeastern parts
of the Taranaki basin the low Qp at 4 km depth persists at 10 km
depth. A 15–20 km wide discontinuous band with Vp 1–4 per cent
higher than the 1-D model at 10 km depth extends from northeast
of Mt Taranaki, beneath the mountain towards the south, and this is
also a region of relatively high Qp (400–700) (Figs 7g and i).
At 16 km depth, Vp is low to the north and west and high between
Mt Taranaki and the Taranaki fault to the east (Fig. 7j). The corresponding Vp/Vs is relatively high (≥1.8) close to the Taranaki fault,
with lower values (≤1.7) further west (Fig. 7k). By 22 km depth
the resolved area is only 30–35 km wide and restricted to central
and eastern Taranaki. However, there are contrasts in both Vp and
Vp/Vs across the Taranaki fault. Vp has a 5–10 per cent contrast,
low to the west and high to the east, while Vp/Vs is ≤1.7 west of
the fault and ≥1.8 east of the fault, the same pattern as at 16 km
depth.
175˚
175.5˚
7 DISCUSSION
Vp rays
-39˚
-39.5˚
-40˚
50 km
-38.5˚
Vp/Vs rays
-39˚
-39.5˚
-40˚
-38.5˚
Qp rays
-39˚
-39.5˚
-40˚
174˚
174.5˚
Figure 4. Ray paths used in determining Vp, Vp/Vs, and Qp models. Note
that the distribution and density of Vp and Vp/Vs ray paths are similar, but
that there are significantly fewer Qp ray paths. There are 4692 P-phases,
3802 S-P observations, but only 678 t ∗ observations. This means that the
resolution of the Qp model is significantly poorer than that of both Vp and
Vp/Vs .
the fault, within the Taranaki basin, Vp is almost everywhere low
(average ca. 4 km s−1 , −15 per cent relative to the initial model)
and east of the fault Vp is +15 per cent relative to the initial model
(ca. 5 km s−1 ). Within the Taranaki basin there is a region, 15–
20 km in diameter, of higher velocity northeast of Mt Taranaki and
a small area of high Vp associated with the Taranaki volcanoes
(Fig. 7d). Vp/Vs does not exhibit the same sharp change across the
Taranaki fault. It is almost uniformly high (≥1.9) within the Taranaki
basin (Fig. 7e), with high Vp/Vs extending outside the basin to the
south, but not to the north where it decreases to 1.65–1.75 in the
northeast part of the basin. West of the Taranaki volcanoes there
is a small, shallow region of very high Vp/Vs (≥2.0) (Figs 7e and
8b), almost directly above the western cluster of earthquakes. There
is also a small region of relatively low Vp/Vs that coincides with
a more pronounced low Vp/Vs beneath the Taranaki volcanoes at
The reasons for a Vp or Vp/Vs anomaly may not always be obvious because Vp and Vp/Vs depend on many factors including
rock type, porosity, the presence of fractures, clay content, in situ
stress, pore pressure, fluid saturation, and the type of fluid (EberhartPhillips et al. 1995). At shallow depths in Taranaki (≤10 km),
the rock type is likely to be a significant factor because there
are strong contrasts between Taranaki basin sediments and both
basement rocks and volcanic/plutonic rocks of the Taranaki volcanoes. Within the Taranaki basin itself the porosity, fluid saturation,
and pore pressure might affect Vp and Vp/Vs, especially at shallow depths. If a significant volume of magma exists beneath the
Taranaki volcanoes then it might be detected by Vp and Vp/Vs
tomography. At greater depths, variations in basement lithology
(Mortimer et al. 1997) might be visible in the observed velocity
structure.
Seismic wave attenuation is caused by a combination of intrinsic
absorption and scattering loss. Intrinsic attenuation occurs through a
loss of energy when seismic waves pass through rocks and is mainly
a result of viscous damping by local pore fluids and frictional sliding on grain boundaries. Intrinsic attenuation therefore increases
with increasing permeability, porosity, and crack density (EberhartPhillips & Chadwick 2002). High pore pressure also causes a
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Figure 5. Spread and resolution contours for Vp, Vp/Vs, and Qp. Nodes are marked by small dots on the z = 0 km layer. Small spread values (red colours)
indicate nodes at which smearing is small. Resolution contours are marked by thin black lines. They are drawn to extend to 60 per cent of the value of the
resolution diagonal element at each node and are shown only where they extend beyond adjacent nodes. Nodes that have low ray coverage and were not included
in the inversion have no spread value. Because Vp and Vp/Vs used a 5 km grid spacing and Qp a 10 km spacing, the Qp spread values cannot be compared
directly with those for Vp and Vp/Vs . Although Qp has a larger grid spacing, the spread values are larger than those of both Vp and Vp/Vs, indicating that the
resolution of Qp is significantly poorer.
significant increase in attenuation and the presence of more viscous fluids will cause higher attenuation. Scattering attenuation is
a redistribution of wave energy in the medium and occurs when
some of the coherent energy in the direct wave is scattered into
incoherent energy by reflection, conversion, or refraction by smallscale features (Eberhart-Phillips & Chadwick 2002). Attenuation
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Journal compilation measurements of direct waves, such as those used here, give a total
attenuation value (Sato et al. 2002), so the factors influencing Qp
include permeability, porosity and crack density, particularly that associated with Taranaki basin sediments and the sediment–basement
boundary, and any sources of significant scattering. In principle, attenuation tomography in Taranaki may also be able to image volumes
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Vp
0
10
20
30
a
Vp/Vs
0
10
20
30
b
Qp
Depth (km)
0
10
20
30
c
30
20
10
0
-10
-20
-30
-40
-50
-60
-70
-80
-90
-100
Profile Position (km)
-0.5
0.0
0.5
1.0
1.5
2.0
2.5
3.0
3.5
4.0
spread value
Figure 6. Spread and resolution contours for east–west true-scale cross-sections through the summit of Mt Taranaki for Vp, Vp/Vs, and Qp. Inverted triangles
mark the position of the coastline. Above each section is the topography, with a vertical exaggeration of 4:1. Refer to Fig. 5 for more explanation.
that contain partial melt as this has been successful for large-scale
subduction-related volcanoes in the Andes (Haberland & Rietbrock
2001) and northern Japan (Tsumura et al. 2000).
7.1 Taranaki basin and Taranaki fault
Most of the Taranaki basin has substantially lower Vp than the initial
model (Fig. 7d), and within the basin there are significant differences
in Vp (Fig. 9). The low Vp along the eastern boundary of the basin
and south of Mt Taranaki coincides with regions of relatively thick
sediments and the 15–20 km diameter region of relatively high Vp
northeast of Mt Taranaki coincides with relatively shallow (5.5–
6 km) sediments (Fig. 9). The low Vp at 4 km depth west of Mt
Taranaki does not seem to reflect overall sediment thickness.
The eastern boundary of the low Vp region at 4 km depth coincides almost exactly with the eastern boundary of the Taranaki
basin, marked by the Taranaki fault (Figs 7d and 8a). Based on
seismic reflection sections, the fault has a 6 km vertical offset
here (King & Thrasher 1996), which is clearly visible in the Vp
cross-section (Fig. 8a). The juxtaposition of rocks of different seismic velocities across the fault results in a velocity contrast of 20–
25 per cent (Fig. 7d), or ca. 1.0 km s−1 (Fig. 8a).
The low Qp anomaly east and southeast of Mt Taranaki at depths
of 4 and 10 km approximates the eastern part of the Taranaki basin
close to the Taranaki fault, and especially the Manaia sub-basin
(Fig. 1), where the sediment thickness is greatest. This is because
intrinsic attenuation increases with permeability and porosity, and
is higher in the basin sediments than in the basement rocks.
The discontinuous high Vp and high Qp region beneath Mt
Taranaki at 10 km depth (Figs 7g and i) may reflect differences
in basement lithology. Mortimer et al. (1997) do not recognize any
change in basement lithology here, but there are no wells in this area
that penetrate basement rocks.
High Vp/Vs in the eastern Taranaki basin at 16 km depth may
reflect crustal thickening beneath the eastern margin of the basin,
with the lower Vp/Vs in the west reflecting the undisturbed basement at that depth. Downwarping of basement rocks close to the
eastern margin of the Taranaki basin may be an extension of crustal
thickening observed in this area (Stern & Davey 1987).
The presence of velocity contrasts across parts of the Taranaki
fault at 16 and 22 km depths (Figs 7j, m and 8) suggests that the
fault remains a significant structural boundary at these depths, and
supports the occurrence of a major change in basement lithology
here (Mortimer et al. 1997; Sherburn & White 2005). At 32 and
40 km depths Reyners et al. (2006) report a strong Vp contrast close
to the Taranaki fault, with low velocity (thicker crust) to the east and
high velocity (thinner crust) to the west, suggesting that the fault
extends to the base of the crust.
7.2 Structure of the Taranaki volcanoes
Improving our knowledge of the structure of the Taranaki volcanoes
is a specific goal of this study because it allows better interpretation
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Figure 7. Map views of Vp, Vp/Vs, and Qp. Vp is shown as percentage differences from the initial 1-D model, and Vp/Vs and Qp are shown as absolute values.
Nodes (small dots) are shown on the z = 0 km layer. The Taranaki volcanoes, indicated by the 500 m elevation contour, and the Taranaki fault are shown.
The area considered well resolved in each layer is outlined and outside this the colours are subdued. The colour bars apply only to the well-resolved areas. All
relocated hypocentres (filled circles) in each depth range are shown on the Vp and Vp/Vs models. Ray tracing extended to the surface, although we did not
determine models above 0 km depth.
of both long-term monitoring data and precursory seismicity. The
region of high Vp and low Vp/Vs coincident with the centre of the
Taranaki volcanoes suggests that the velocity tomography may be
able to image a root or core to the volcanoes. Locke et al. (1993) and
Locke & Cassidy (1997) modelled the gravity anomalies associated
with the Taranaki volcanoes as a ca. 5 km diameter high-density
core between the surface and a depth of 6 km, which was inter
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2006 The Authors, GJI, 166, 957–969
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Journal compilation preted as a dyke/stock complex produced by repeated intrusion of
magma into Taranaki basin sediments over the lifetime of the volcanoes. Eberhart-Phillips (1993) demonstrated that it is possible
to derive the main features of the gravity field from a Vp tomographic model, and Sherburn et al. (2003) showed that Vp is low in
calderas where gravity anomalies are negative. It therefore seems
likely that we are imaging the same structures modelled from the
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S. Sherburn, R. S. White and M. Chadwick
Figure 8. Vp, Vp/Vs, and Qp for an true-scale east–west cross-section through the summit of Mt Taranaki. The contour interval for Vp is 1 km s−1 , for Vp/Vs
is 0.1, and for Qp is 200. Inverted triangles mark the position of the coastline, ‘TF’ is the position of the Taranaki fault, and ‘CEFZ’ the Cape Egmont fault
zone. The area considered well resolved is outlined and outside this the colours are subdued. The colour bars apply only to the well-resolved areas. All relocated
hypocentres within 25 km of the cross-sections are shown on the Vp and Vp/Vs models.
174˚
174.5˚
175˚
174˚
174.5˚
175˚
−
−
20 km
20 km
−
−
Vp 4 km
−
−
−
−
−
−
0
20
–
4
5
6
7
8
Sediment thickness (km)
Figure 9. Vp and sediment thickness in the Taranaki basin. Vp is shown as percentage differences from the initial 1-D model, with a colour scheme that
highlights differences within the Taranaki basin. Refer to Fig. 7 for further explanation.
Taranaki gravity data, and a specific test was carried out in an attempt to confirm this. In a synthetic Vp model the velocity was
assumed to mirror the density contrast in the gravity model, namely
23 per cent higher than that of the surrounding sediments at the
surface (Fig. 10a), decreasing to 4 per cent higher at a depth of
4 km (Fig. 10d). The recovered anomalies (Figs 10b and e) reflect
the shape of the input anomalies well at 0 km depth, though the recovered amplitude is only about one third of the synthetic amplitude
(+7 per cent). At 4 km depth, the recovered anomaly agrees with
the synthetic one in both amplitude and position. This shows that
the velocity tomography has the potential to image the dyke/stock
complex beneath the volcanoes, but that at 0 km depth it will underestimate the magnitude of the velocity contrast by as much as
two-thirds.
The shape of the observed anomaly at 0 km depth (Fig. 10c) is
consistent with a high velocity–density root at this depth beneath all
of the Taranaki volcanoes. At 4 km depth, the observed Vp anomaly
is smaller than that at 0 km depth and is confined to Mt Taranaki
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Figure 10. Plan views of synthetic, recovered, and observed Vp models at depths of 0 and 4 km from a resolution test to assess the ability of the data to resolve
the Locke & Cassidy (1997) gravity model of the Taranaki volcanoes. Crosses show the grid nodes and triangles those seismographs close to the Taranaki
volcanoes. In the synthetic and recovered plots the grey circles represent the nodes at which the synthetic anomalies were placed. The Taranaki volcanoes are
shown by the 500, 1000 and 1500 m elevation contours.
and Pouakai, with nothing seen beneath the smaller, older Kaitake
volcano.
Locke & Cassidy (1997) estimated the volume of the dyke/stock
complexes beneath both Mt Taranaki and Pouakai to be 150 km3 ,
but beneath Kaitake to be only 85 km3 . If the complex beneath
Kaitake extends to the same depth as those beneath Mt Taranaki
and Pouakai, ca. 6 km, then its diameter would be only ca. 2 km,
which may explain why it is not visible at 4 km depth in the observed
data (Fig. 10f). If the high Vp core to the volcanoes represents cooled,
solidified magma, then the velocities should approximately match
those of andesite at these depths. At a depth of 5 km the Vp of
andesite is ca. 5.4 km s−1 (Christensen & Mooney 1995), close to
the 5.0–5.1 km s−1 observed from the tomography at 4 km depth
beneath Mt Taranaki, which supports our interpretation that the high
Vp, low Vp/Vs feature is solidified andesite magma.
Locke & Cassidy (1997) noted that updoming of sediments
near the basement is observed beneath Sugar Loaf–Paritutu (SLP,
Fig. 1) on seismic reflection sections, from which they argue that
the dyke/stock complex there extends into the basement. Our Vp
model shows evidence for updoming at 6 km beneath Mt Taranaki,
and maybe also at 10 km depth (Fig. 8a) suggesting that the
intrusions do extend into the basement, and maybe as deep as
10 km.
At Kirishima (Tomatsu et al. 2001) and Tungurahua (Molina et al.
2005) a high Vp core or pipe-like structure was imaged by velocity
tomography, at Redoubt (Benz et al. 1996) and Mount Spurr (Power
et al. 1998) a low Vp region was seen beneath the crater. Both the
studies that image high Vp regions and those that image low Vp
regions interpret these as magmatic conduits and a source zone for
recharge of the shallow magmatic system. In the studies where the
tomography used data collected during eruptive activity (Benz et al.
1996; Power et al. 1998; Molina et al. 2005), the region of anomalous
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Journal compilation Vp beneath the craters was also the most seismically active. This
supports the inference that those structures are pathways for magma
transport, and importantly for Mt Taranaki, suggests that any future
precursory seismicity is likely to occur within or close to the high
Vp, low Vp/Vs core, directly beneath the summit crater.
Any magma beneath Mt Taranaki would appear as a region of
low Vp and high Vp/Vs (e.g. O’Connell & Budiansky 1974; Takei
2002). Based on the stability of amphibole and biotite that are found
in Mt Taranaki eruptives, magma has to be within basement rocks at
depths greater than 5–6 km (Bob Stewart, personal communication,
2004), but at these depths we see no distinct regions of low Vp and
high Vp/Vs (Figs 7g, h, j, k, 8a and b). In tomographic studies at
volcanoes similar to Mt Taranaki it is not unusual that evidence for a
discrete magma body cannot be found (Power et al. 1998; Benz et al.
1996; Tomatsu et al. 2001). Price et al. (2005) suggested that at longlived andesite volcanoes, eruptions are fed from separate, unlinked
crustal magma chambers or migrate through the crust through a
series of dispersed dykes or sills, rather than accumulating in the
classical single magma chamber. If this is the case at Mt Taranaki,
then it is not surprising that we cannot detect the direct effects of
any magma in the seismic data.
Allis et al. (1995) modelled the source of a high heat flow anomaly
just north of Mt Taranaki (Fig. 1) as either a mid-crustal intrusion of
ca. 500 m thickness intruded over the last 0.2–0.5 million years, or
as crustal underplating over the last 2–4 million years producing ca.
5 km thickness of magma. We are unable to discriminate between
these models based on our data.
Despite the strong Vp and Vp/Vs contrast between the roots of
the Taranaki volcanoes and the surrounding sediments, it would
not have been possible to delineate accurately these with a seismograph spacing larger than the 5 km we used. Similar studies
have taken advantage of the high level of seismicity preceding and
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S. Sherburn, R. S. White and M. Chadwick
accompanying eruptive activity to provide sufficient ray paths to
image volcanic structures, but the background seismicity is low beneath the Taranaki volcanoes (Sherburn & White 2005) so a 5 km
seismograph spacing was essential.
7.3 Evidence for fluids in the Cape Egmont fault zone
Faults are characterized by highly fractured material, fault gouge,
and elevated pore pressures, which tend to lower Vp and increase
Vp/Vs (Eberhart-Phillips et al. 1995; Johnson & McEvilly 1995;
Zhao et al. 1996; Thurber et al. 1997). The small region of high
Vp/Vs (≥2.0) above the Cape Egmont fault zone (CEFZ) seismicity
may be associated with such a fractured, yet high pore-pressure
region (Fig. 8b). There is low Vp at 4 km (Fig. 9) though it is not
as significant as the high Vp/Vs anomaly. This may be because
Vp/Vs is more sensitive to fluid saturated cracks and pores than is
Vp (Nur 1972; Johnson & McEvilly 1995; Zhao et al. 1996). The
high Vp/Vs region does not continue to the most seismically active
depth (Fig. 8b) as was observed in the source region of the Kobe
earthquake (Zhao et al. 1996), but is restricted to shallow sediments.
This may be because in saturated rocks an increase in the confining
pressure will cause a decrease in Vp/Vs because of a decrease in
pore volume and a consequent increase in Vs (Nur 1972). When
taken together with the frequent occurrence of earthquake swarms
in the Cape Egmont fault zone and a small region of anomalously
high b-value (Sherburn & White 2005), the high Vp/Vs here suggest
that some parts of the Cape Egmont fault zone may be characterized
by elevated pore fluid pressures.
7.4 Earthquake relocation
All 389 earthquakes were relocated with the 3-D Vp and Vp/Vs models. Epicentral differences from those calculated with the minimum
1-D velocity model with station corrections (Sherburn & White
2005) are ≤2 km, and depth differences are typically ≤5 km. These
are mostly random, though in eastern Taranaki some of the shallower earthquakes (≤30 km) were relocated almost 10 km deeper.
This increases the concentration of seismicity in the lower crust in
this area (Sherburn & White 2005), and strengthens the conclusion
that the most seismically active depth here is almost completely
restricted to the lower part of the crust.
8 C O N C LU S I O N S
Using data from a large portable seismograph network we have been
able to determine 3-D Vp, Vp/Vs and Qp models for the crust in the
Taranaki region in the western part of the North Island of New
Zealand. These models show, as high Vp and low Vp/Vs regions,
the 5 km diameter remnant conduits of solidified magma beneath
the Taranaki volcanoes that are probably the paths through which
future magma intrusions will reach the surface. We are unable to
image any magma storage within the upper 16 km of the crust.
We can also image accurately the eastern boundary of the Taranaki
basin, changes in basin sediment thickness, and a small, high Vp/Vs
anomaly above the seismically active Cape Egmont Fault Zone west
of Mt Taranaki.
Although many of the 3-D Vp and Vp/Vs features found in this
study have large velocity contrasts, many more than 10 per cent, the
accurate delineation of these would not have been possible without
such a large network of seismographs at a 5 km station spacing.
AC K N OW L E D G M E N T S
We would like to thank many farmers and landowners, the Department of Conservation, Taranaki Maori Trust Board, and the local
iwi for permission to operate seismographs in Taranaki; SEIS-UK
and the Taranaki Region Council for field support; and many colleagues, family and friends, especially Claudia Allen, for assistance
with fieldwork. Discussions with Regina Lippitsch and Daniel Rowlands were greatly appreciated while this work was being done.
Comments and advice from Donna Eberhart-Phillips on several aspects of the tomography, and comments from Tony Hurst were very
helpful. Stephan Husen’s TOMO2GMT program was used to format
SIMUL2000 output ready for plotting. Diagrams were drawn using
GMT (Wessel & Smith 1998). While at Cambridge University, SS
was funded by a Bright Futures Top Achiever Doctoral Scholarship
from the New Zealand Ministry of Education. Department of Earth
Sciences, University of Cambridge, contribution no. ES.8352.
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