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A back-arc basin formed within continental lithosphere: The Central Volcanic region of New zealand
385
Tectonophysrcs, 112 (1985) 385-409 Elsevier Science Publishers
A BACK-ARC
B.V.. Amsterdam
BASIN
THE CENTRAL
- Printed
FORMED
VOLCANIC
in The Netherlands
WITHIN
CONTINENTAL
LITHOSPHERE:
REGION OF NEW ZEALAND
T.A. STERN Geophysics Division, D.S. I. R., Wellrngton (New Zealand) (Received
December
2, 1983; revised version accepted
August
23. 1984)
ABSTRACT
Stern, T.A., 1985. A back-arc New Zealand. Zones.
Tectonophysics,
The Central
basin formed within continental
In: K. Kobayashi
Volcanic
Region
Island of New Zealand
-the
Havre
possesses
back-arc
basins.
mantle
respectively),
the crust
seismic
is anomalously
represents
of New
velocities
Volcanic
and Processes
continental
of heat is required. spreading
Zealand.
characteristics
thin for its continental
a site of active back-arc
and tensional
Region of
in Subduction
setting,
geothermal
Volcanism
the
basin
with oceanic
of 7.4 and 3.95 km/s
and the heat flow, almost is about
to account
therefore,
within a continental
within
back-arc
is predominantly
often associated
systems,
heat flow). In order It is proposed,
tectonics
of a young oceanic
are low (P, and S, velocities
of hot water through
than normal
flow a large scale mass transfer Region
structure
many of the geophysical
in the discharge
some twelve times greater
volcanism
to be a direct continuation
the continental
yet the region
which is expressed
is an area of Quaternary
that appears
rhyolitic,
Upper
the Central
Structures
112: 385-409.
North
Trough-into
lithosphere:
and I.S. Sacks (Editors),
700 mW/m2
all of (i.e.
for such a high heat
that the Central
Volcanic
lithosphere.
INTRODUCTION
Evidence
from seismicity
(Adams
and Ware, 1977; Reyners,
the subducted Pacific plate lies beneath the North the South Island of New Zealand (Fig. 1). Features
1980) indicates
that
Island and the northern end of typically found at a convergent
plate boundary such as an oceanic trench, andesitic volcanism, dipolar gravity anomalies, and crustal and upper mantle seismicity are all found in the North Island. Figure 2 shows associations that can be made (e.g. Cole and Lewis, 1981) between various geological structures of the North Island and Karig’s (1974) generalised structural model of a western Pacific island arc system. At some convergent margins an area is found behind, or landward of the subduction zone where tensional tectonics and high heat flow are observed. These are marginal (Karig, 1974) or back-arc (Uyeda, 1977) basins. The term back-arc basin is used here. The Central Volcanic Region of the North Island (Figs. 1 and 2) C&to-1951/85/$03.30
0 1985 Elsevier Science Publishers
B.V.
386
I
‘\
179E
SOUTH
1 cop
\
I
I
f/I
’
FIJI BAStN
PLATE
Fig. 1. Locality Havre Trough. assuming
map showing Convergence
the Indian
Havre Trough.
the position rates between
of the Central
plate to be fixed. Double
Eathymetry
in km.
Volcanic
Region
the Pacific plate and the Indian headed
arrows
represent
(C.V.R.)
with respect
plate (in mm/yr) back-arc
spreading
to the
are shown within
the
387
km
L -500
I
I
-400
-300
1
-200
,
-100
-50
0
+50
+100
+200
lwclneters
Fig. 2. Diagram structures
showing
of the North
some of the associations Island
that can be made
(top) and the generalised
island arc system as given by Karig (1974). Upright
triangles
structural represent
between
model
low-potash
have been active within
the past 10,000 years. The area1 extent of the Central
and the Taupo
Zone (T.V.Z.) are shown.
Volcanic
the various
(below)
geological
of a western
andesite Volcanic
Pacific
volcanoes
that
Region (C.V.R.)
388
is an area of crustal appears
extension
(the oceanic back-arc There
and high heat flow, and as Karig (1970) notes,
to be a direct continuation basin associated
is a fundamental
Volcanic
Region,
spreading
back-arc
into New Zealand
difference
however.
Whereas
basin (Karig,
of the young
with the Tonga-Kermadec between
the Havre
the Havre Trough
trench-arc
Trough
is thought
1974) within an oceanic
it
Havre Trough
and
system).
the Central
to be an actively
area, the Central
Volcanic
Region is classified by Karig and Jensky (1972) as a “ volcanic-tectonic rift zone” situated within a continental environment. Karig and Jensky postulate a general progression, with increasing amounts of extension, from a volcano-tectonic rift, which they also describe as the earliest phase of extension in a back-arc basin, to an actively spreading back-arc basin where new lithosphere is created. The purpose of this paper is to examine the response of continental lithosphere to back-arc processes by presenting interpretations data that bear upon the crust and upper mantle
of seismic, gravity and heat flow structure of the Central Volcanic
Region.
GEOLOGICAL
PROVINCES WITHIN THE CENTRAL
NORTH ISLAND
The area Karig (1970) referred to as the Taupo Volcanic Zone closely resembles what is called here the Central Volcanic Region. Figure 2 shows the area1 extent and location of what is commonly termed by New Zealand geologists (e.g., Healy, 1962; Cole, 1979) the Taupo Volcanic Zone (T. V.Z in Fig. 2). This distinction in terminology is important; past geological and geophysical studies have tended to concentrate on the Taupo Volcanic Zone, for here the economically important geothermal activity occurs and the impressive exposures of Quaternary rhyolitic volcanics
can
be seen. About
12,000 km’ of Quaternary
rhyolitic
volcanics
are
contained within the Taupo Volcanic Zone (Cole, 1979) the average heat flow of which is estimated to be about 700-800 mW/m’ (Studt and Thompson, 1969; Allis. 1979).
This
continental extension
value
is approximately
norm of 60 mW/m’. in historical
twelve
The Taupo
times (Walcott,
to fourteen Volcanic
1978; Sissons,
times
greater
than
the
Zone has also been an area of 1980). Analysis
of repeated
geodetic surveys spanning the past 70 years, shows significant shear strain rates that are interpreted by Sissons to be due to (back-arc) spreading in a NW-SE azimuth at a rate of 7 f 4 mm/yr. The Taupo Volcanic Zone forms the eastern half of a larger area of Quaternary volcanism in the North Island-the Central Volcanic Region-with which it shares a common eastern boundary. The western boundary of the Central Volcanic Region separates extrusions of Quaternary rhyolite to the east from Mesozoic greywacke and Tertiary rhyolites to the west (Fig. 3). So defined, the western boundary of the Central Volcanic Region is based on surface geology. However, the boundary also receives emphasis from gravity and seismic data.
Greywacke
RhyoMe domes (Q=Quaternary
T=Ter&y
Andesite volcanoes -active in past 18yrs 0,
,
5;Okm
BAY OF PLENTY
White Island A .:*
TAFiANAKl
1750
Fig.
3. Generalised
greywacke
ranges
the area within basalt about
200-300
cone-shaped
map
and immediately
and andesite
C.V.R. is a broad
geological
of the Central
that flank the C.V.R., outcrops
occur
North
and the Quatemary
adjacent
to the C.V.R.
in only minor
Island
an ignimbrite
slice of the axial greywacke
converging
the C.V.R.
Much of
is mantled
by extensive
ignimbrite
deposits;
volumes.
sheet. The plateau ranges
southwardly
domes within
The Kaingaroa
area where drilling and seismic work (Hochstein m beneath
showing
rhyolite
and Hunt,
has been
that has subsided
towards
Plateau
to the east of the
1970) show greywacke
interpreted the C.V.R.
(Healy,
to lie
1962) as a
390
CRUSTAL
STRUCTURE
Upper crustal structure Rock samples from drill holes (Grindley, 1965, 1982) and seismic refraction studies (Stern, 1982; Robinson et al., 1981) (Fig. 4) indicate that the upper 2 km of crust for the Central Volcanic Region consists of low-density, low-velocity volcanic rocks (the terms density and velocity used here refer to saturated density and compressional seismic velocity respectively). At a depth of 1.8 to 2.2 km a “basement” rock unit of velocity 4.8-5.5 km/s has been identified at three locations within the Central Volcanic Region (Fig. 4). The rock type represented by this basement
layer is debatable.
In drill holes at Wairakei
and Mokai (Fig. 3) andesite
and welded ignimbrite with densities of 2.55-2.65 Mg/m’ were encountered depth of about 2.1 km (Grindley, 1982; Stern, 1982). Rocks in this density would satisfactorily
account
for the observed
basement
At Broadlands (Fig. 3), however, Mesozoic greywacke depths of l-2 km (Hochstein and Hunt, 1970).
velocities
of 4.8-5.5
at a range km/s.
rocks were encountered
at
It has generally been assumed by New Zealand geologists and geophysicists (e.g., Grindley, 1965; Hochstein and Hunt, 1970, Rogan, 1982) that a continuous, albeit extensively faulted, greywacke basement extends beneath the Central Volcanic Region. An alternative view has, however, been put forward by Calhaem (1973) and Evison et al. (1976). They consider that the Central Volcanic Region is an area of active back-arc spreading and that a new lithosphere is in the process of being created. Under this proposal basement rocks will predominantly consist of intruded volcanic and igneous rocks that are in various stages of cooling. Furthermore, Evison et al. note that the drill holes at Broadlands are situated within 5 km of the eastern boundary of the Central Volcanic Region, and thus contend that the severely faulted greywacke surface found at depth here represents blocks that have been spalled from the eastern greywacke ranges during previous rifting episodes. This question of what rock type constitutes basement for the Central Volcanic Region is an important one in terms of unravelling the tectonic history of the region; it will be taken up again later in this paper when heat flow data are considered. Geophysical studies by Rogan (1982) and Stern (1982) point to the possible existence of molten rhyolite magma within the upper crust of the Central Volcanic Region. For example, consider the residual gravity anomaly map shown in Fig. 4b. These residual anomalies were obtained by removing a regional field from the observed gravity anomaliesthe regional field was numerically derived by fitting a polynomial surface of low-order to gravity values situated on greywacke outcrops found each side of the Central Volcanic Region (Fig. 3) (Stern, 1979). The residual field of Fig. 4b, reflects, therefore, the anomalous density distribution in the upper crust of the Central Volcanic Region, with respect to the density distribution in the greywacke crust elsewhere within the North Island. The residual anomalies are
391
0
3/TAUPO
2/MANGAKINO
I /ROTOEHU
Oi8
0.8
3.0 km/:
0.8
4 3.2
km/s
3.2
kmls
5.5
km/s
0.5 3.2 km/s
1.0
i.0 km/s
3.2 km/s
1.5
2.0
5.5 km/s 4.8 km/s
2.5
RESIDUAL
3
GRAVITY
ANOMALIES
contour
20
Fig
4. a. Interpretations
Volcanic
Region.
Rotoehu
(I ), Mangakino
4a for Rotoehu
of three,
b. Residual
gravity
reversed, anomaly
(2) and Taupo
and Mangakino
seismic
refraction
(3) refraction
refer to the structure
surveys
km
surveys
map of the C.V.R.
1@mgal)
intew
(after
carried Stem,
out within
the Central
1979). Positions
are shown. The velocity
at the south ends of these lines.
columns
of the in Fig.
392
almost wholly negative - 55 to -60 determined
mGal
2 km thickness
4) is estimated residual deeper
within the Central of volcanic
Region
with minimum
gravity
effect of the seismically
rocks at Taupo
and Mangakino
to be only - 35 to - 40 mGa1. The remaining
gravity than
Volcanic
(Fig. 4b). Yet, the anomalous
anomaly
2 km.
must, therefore,
Interpretations
be explained
of these
(Figs. 3 and
- 15 to - 25 mGal of
by negative
“secondary”
values of
mass anomalies
residual
anomalies
at
Mangakino and Taupo (Stern, 1982) have been made in terms of discrete bodies of molten magma at depths of 3 to 10 km. Although it is acknowledged that other interpretations are possible, the molten rhyolite hypothesis receives support from the geochemical studies of Ewart et al. (1975). They show from phenocryst equilibration studies that the rhyolitic rocks erupted from the Central Volcanic Region had, prior to eruption, equilibrated in magma chambers at depths of 7--8 km. Moreover, their data suggest temperature and density ranges for the rhyolite eruption, of 725”-915°C and 2.0-2.38 Mg/m” respectively.
magma,
prior
to
Deep crustal structure
Studies made on dispersive earthquake waves (Thompson and Evison, 1962) and the average value of Bouguer gravity anomalies in New Zealand (Reilly, 1962) point to a crustal thickness of 30-40 km for New Zealand as a whole. There may be considerable variation from this average, however, particularly in the North Island where both Bouguer and isostatic gravity anomalies vary from - 150 to 60 mGa1. Moreover, from a study of earthquake travel-times, Haines (1979) finds that the velocity of upper mantle, compressional seismic waves (P,) varies from 8.5 to 7.4 km/s
within
between
depth
the North
Island.
to the upper mantle
Woollard
(1970)
and Pn velocity
observes
a linear
for North
America,
relationship thus the P,
velocity of 7.4 + 0.1 km/s that Haines associates with an area that covers much of the Central Volcanic Region, suggests that the crust may be thinner here than elsewhere
in the North
Island.
Figure 5 shows the travel-time data and plane-layer interpretation for a two-way deep seismic sounding survey carried out within the Central Volcanic Region. One 400 kg shot was fired in the western bay of Lake Taupo and three 400 kg shots were fired within the vicinity of White Island. Seismograph stations were distributed between these shot points. The interpretation given in Fig. 5 is a horizontally layered model because the apparent velocity and time intercepts of the straight-line segments for both branches of the time-distance graph are about equal. The velocity of 7.4 rt 0.2 km/s at a depth of 15 km beneath the Central Volcanic Region (Fig. 5) is the same as the velocity observed by Haines (1979) for the region, and also is similar to P,, velocities detected in other parts of the world where magmatic activity high heat flow, extension and/or spreading are occurring. For example, P, velocities of 7.1-7.5 km/s, coupled with a thinner than normal crust,
393
CENTRAL
VOLCANIC
REGION TRAVEL-TIME
DATA
0 first arrival
cl-
+
second arrival
o earthquake
travel time(WTZ-
KRP)
3-
o-
L
0
White Island
~STAN~E(km)
Lake TAUPO
PLANE LAYER SOLUTION
O3.0 km/s
2-
6.3 km/s
42
6-
$
8-
: t
lo-
6.0 km&
1214-
..- 14.8
km
7.4 km/s Fig. 5. Data and plane layered Central
Volcanic
Karapiro
(KRP)
beneath
Whakatane
Region.
solution
for a long-range
The open circle
(Fig. 6) for an eathquake (Dr M. Reyners,
represents
seismic refraction
on the 28th February
N.Z. Seismological
survey carried
the travel time between Observatory,
out within
the
(WTZ)
and
Whalcatane
1983 located
apprbximately
pers. commun.,
1983).
1 km
394
are observed in Iceland Atlantic Ridge (Talwani western
(Bott, 1965) the Africa Rifts (Fairhead, et al., 1965) and the Basin and Range
U.S.A. (Thompson
1976). the Midprovince of the
and Burke, 1974; Scholz et al., 1971). Accordingly,
results of Fig. 5 are interpreted in terms of a 15 km thick crust underlain upper mantle with an anomalously low velocity of 7.4 km/s.
A GRAVITY
INTERPRETATION
ADJOINING
MODEL
FOR THE CENTRAL
VOLCANIC
the by an
REGION
AND
AREAS OF THE NORTH ISLAND
A summary
of heat flow (Pandey,
and this paper),
crustal
thickness
1981)
estimates
seismic sounding
data (Garrick,
from S to P conversion
studies
1968; (Smith,
1970 and W.D. Smith, pers. commun., 1983) and seismic attenuating properties of the upper mantle (Mooney, 1970) is given in Fig. 6. East and southeast of the Central Volcanic Region, heat flow is lower than normal, the crustal thickness is about 35 km, and no attenuating region within the upper mantle is observed. West of the Central Volcanic Region heat flow is higher than normal. the upper mantle attenuates seismic waves, and the crust is thinner than normal continental crust (apart from the S to P conversion estimate of 28 km at TUA (Fig. 6), Eiby (1958) reports a crustal thickness estimate of 23 km just to the north of Auckland (AUC in Fig. 6) based on unpublished seismic reflection data). Thus on the basis of the data given in Fig. 6 three distinct tectonic provinces within the North Island can be identified: the Northwest North Island, the Central Volcanic Region, and the East Coast-Wanganui Basin area. This partition of the North Island, as first noted by Hatherton (1970), is also reflected in both the Bouguer and isostatic gravity anomalies of the North Island (Fig. 7). The area of the North
Island
associated
marked
by attenuation
within
with positive gravity anomalies,
the upper
the Central
mantle
Volcanic
and
thin
crust
is
Region is an area of
relatively low gravity flanked by gravity highs, and the East Coast-Wanganui Basin area is dominated by an intense gravity low. The cross-sections AA’ and BB’ of Fig. 7 are shown in Fig. 8. Of note in these cross-sections
is that
markedly
different
gravity
anomaly
profiles
are generated
above different portions of the subducted Pacific Plate. Profile AA’ shows the characteristic dipolar gravity anomaly pattern commonly observed across convergent margins (Hatherton, 1969); a gravity low over the point of major downturn of the subducted plate coupled with a broad positive gravity anomaly over the back-arc area. The gravity anomaly of profile AA’ does, however, differ from the dipolar anomalies observed over other convergent margins like, say, Japan. In the Japanese situation the gravity high peaks within the vicinity of the volcanic arc (e.g. Yoshii, 1979), whereas for the North Island the maximum in the gravity high is reached some 50-100 km westward, or landward, of the volcanic arc. Hatherton (1970) interpreted a profile across the North Island slightly to the
395
f%km)-S
to P conversion
ECZ 0 =
depth
SEISMOGRAPH STATION
= UPPER MANTLE ATTENUATING REGION ~~11
l-y---l 6. 2
- 15 km
7.0 km/s
NORTHWEST NORTH I Mea,fbv,=80”22
mWi&53)
NI
heat
flow(convective)
(49km)
fo
Ir SOUTH WANGANUl i TARANAK$S”JMNG~~~~,,,>
1OOkm
/;F -10
-20
35 km 118’)
I Fig. 6. Geophysical
data bearing
flow data from Pandey depths Garrick
(Smith,
upon the crust and upper
(1981), upper mantle
attenuating
1970 and W.D. Smith, pers. commun.,
(1968) and this paper.
mantle
regions
structure
of the North
from Mooney
(1970), S-P
1983), and seismic refraction
I Island.
Heat
conversion
interpretations
from
TATIC
GRAVfTY
ANOM
area
beneath
subducted
which
Pacific
is present
<
-50
the
plate il.l
mgal
NORTH ISLAND
Fig. 7. isastatic
{Airy-Heiskanen,
1965). Area beneath direction
which
of dip for the subducted
refer to profites in Fig. 8.
T-
30 km) gravity
the subducted
Pacific
plate isnorthwest,
anomaly
plate
map of the
is inferred
i.e. in the direction
North Island (after Reilly,
to be present
is shown.
Note
the
of A’-,4 and B’- 8. AA’ and BB’
AA’
in gravity
anomalies
and BE’ of Fig. 7 showing
Note the gross dissimilarity
Fig. 8. Profiles
between
the position
of the subducted the two profiles.
-d)
studies, and associated
PROFILEtB
from seismicity
WANGANUI
Pacific plate, as inferred
250
250
50
0
200
Bay
2co
Hawke
PROFILECA-h
150
,
BAY
150
C.V.R.
HAWKE
100
I
-
too
Waikato
WAIKATO
gravity
anomalies.
5
north of AA’ (Fig. 7). He ascribed a mass deficiency
within
component
to a mass excess within
BB’
S), however,
(Fig.
seismological present North
evidence
the negative
the upper exhibits
(Adams
here. This observation
component
of the dipolar
80 km of the subducted
to
the plate for the depth range 8OG250 km. Profile no such
dipolar
gravity
and Ware, 1977) indicates suggests
Island are not so much controlled
anomaly
even
the subducted
that the principal
gravity
by mass anomalies
the subducted Pacific plate, but are more influenced mantle inhomogeneities within the overriding Indian Figure 9 shows an interpretation
anomaly
plate. and the positive
directly
by crustal plate.
though
plate to be
anomalies associated
thickness
of the with
and upper
of the 280 km long profile AA’ of Fig. 7. In this
interpretation a model for the suducted plate is included that used by Sager (1980) for the Mariana Arc structure.
that is similar in concept to A negative gravity effect is
ascribed to a 6 km thick oceanic crust layer, that caps the subducted plate, being taken down to a depth of 80 km. At this depth the oceanic crust is assumed to transform to eclogite of density 3.56 Mg/m3. The eclogite layer continues as the upper 6 km of the subducted plate to a depth of 150 km, thus constituting part of the positive mass anomaly directly associated with the subducted plate. The remainder of the positive mass anomaly is due to that part of a 80 km thick, relatively cool subducted plate, less the eclogite layer, in the depth range of 100 to 250 km. A density contrast of to.02 MS/m3 was found to give the best fit for this part of the subducted plate. The Central Volcanic Region is represented in Fig. 9 by three isodensity layers. The top layer represents the 2 km thick section of low-density, low-velocity volcanic rocks as detected
in drill holes and by seismic refraction
data (Fig. 4). Between 2 and
13 km is a 2.60 Mg/m3 layer. The choice of this lower than normal crustal density, with respect to the 2.75 Mg/m’ density used to represent the crust elsewhere in the North Island (Fig. 9), was found necessary to ensure a satisfactory fit between observed and calculated gravity anomalies. However, some support for this choice of density
comes from the observed
seismic
velocities
in this layer of 5.3-6.0
km/s
(Fig. 5), which are also somewhat lower than normal for continental crust. Profile AA’ passes through the Central Volcanic Region on a line just north of Lake Taupo where a deep gravity low of short wavelength is observed (Fig. 4b). In order to accommodate this gravity low a negative mass anomaly of 2.35 Mg/m’ is included within the 2.60 Mg/m’ layer, and is interpreted to represent a volume of molten or partially molten rhyolite (see discussion in previous section). The 3.18 Mg/m3 upper mantle density for the Central Volcanic Region is chosen on the basis of the low (7.4 km/s) upper mantle seismic velocity (Fig. 5), although there is no data that controls the depth to which the 7.4 km/s layer extends. The reduction of the upper mantle density for the area to the west of the Central Volcanic Region from 3.40 to 3.30 Mg/m3 is chosen on the basis of the low (7.887.9 km/s) P, velocities for this area, relative to the more “normal” P, velocities of 8.1-8.5 km/s observed for the area to the east of the Central Volcanic Region
COAST
’
I
side.
left-hand
/
are calculated
II
/-
COAST
Coast
with respect to the standard
density
_
_
Mg/m’
-
basin
is model shown on the
of +0.02
_
sedimentary
plate with a density
1
EAST
of the subducted
~~______~________~~~~_________~__~~_~~________,___
2p0
AA’ of Fig. 7. The portion
*c’
c:
;(_d ’0
to a depth of 250 km. Crust. and upper mantle mass anomalies
field for the profile
f
3.18
of the Bougucr
anomaly
LK”5
as extending
’ //
-’
modelled
150 km
L.IP
(2.35)
C.V.R.
Fig. 9. Interpretation
MODEL
DENSITY
ll STANDARD
3.3
1
3.4
Y/A V//‘/.
WEST
%
(Haines, 1979). The crustal thickness and mantle densities of the western and eastern North Island, and the Central Volcanic Region have all been adjusted in such a fashion
that
pensation
between
achieved Region
local isostatic
compensation
the western
at a depth
North
is maintained.
Island
and the Central
of 30 km and compensation
and the eastern
North
Island
On the East Coast of the North and Lewis (1981) consider
is achieved
Island
More
between
specifically, Volcanic
the Central
com-
Region
is
Volcanic
at a depth of 100 km.
is the East Coast Sedimentary
this basin to be part of an accretionary
Basin. Cole
prism containing
sediments whose ages span from early Tertiary to Quaternary. The exact depth of the basin is unknown but is modelled in Fig. 9 as a 5 km deep basin with a density contrast of -0.25 to - 0.35 Mg/m3. Compensation for this basin is presumably achieved in a regional sense by the excess mass of the subducted plate. That is, the basin is maintained by “Trench flexure” as defined by Davies (1981). The calculated gravity anomaly curve in Fig. 9 is comprised of two principal components: (1) The effect of mass anomalies directly associated with the subducted Pacific plate. (2) The gravity effect of anomalous upper mantle and crustal structures within the overlying Indian Plate. These two components are shown plotted separately in Fig. 10. Of note in Fig. 10 is that the calculated effect of the upper mantle and crustal structure of the Indian plate accounts for much of the character of the observed gravity anomaly curve (Fig. 9). In particular, it explains about 40 of the 50 mGa1 back-arc gravity highs observed west of the Central Volcanic Region. This finding, therefore, concurs with that of Watts
and
Talwani
(1974,
1975) who
argue,
on the
basis
of an elastic
plate
approximation to a subducted plate, that some back-arc gravity highs are due not so much to the excess mass of the subducted plate, but to gravity anomaly effects associated
with back-arc
extension.
There are, however, many speculative aspects of the model in Fig. 9. In particular, the depth at which oceanic crust transforms to eclogite is debatable; some authors (e.g., Grow, 1973) place the transformation as shallow as 30 km. Moreover, the model for the crust and upper mantle of the North Island is essentially a static, locally compensated one. In a dynamically active area such as that found at a subduction zone, one cannot rule out the possibility that some lithospheric mass anomalies may be supported by dynamic flow in the asthenosphere rather than by the finite strength of the lithosphere (McKenzie, 1967; Lambeck, 1972). Nevertheless, the import of the gravity anomaly analysis presented in Figs. 9 and 10 is that the gravity anomaly field of the North Island appears to be more dominated by those effects associated with extension within the overlying Indian plate, rather the gravity effect of mass anomalies within the subducted Pacific plate.
than
\d
calculated
(i) the calculated
50
gravity effect of mass anomalies
Fig. 10. A graph showing:
0
u
coast
---
V Vest
-40
-20
o-
20
40
100
effect
the
subducted
150
directly
assockted
plate
200
with the subducted
plate.
uppermantle structure
Jc= C.V.R. =$I
Y__;.
of
gravity effect of the crust and
gravity
1
gravity
Centrasl
1. \ the
I Y
for
combined
crust
250
North
3001km
Island,
Island
structure
North
mantle
western
upper
the
coast II
and
and
and northwestern
East
Region
of
of the western
Volcanic
effect
and (ii) the g
402
HEAT
FLOW AND LITHOSPHERIC
The principal
EVOLUTION
feature of the structural
model developed
Region
in the previous
section,
mantle.
Such a structure
could be reconciled
(1) A greywacke
is the thin crust
for the Central
underlain
Volcanic
by low-velocity
upper
with either:
crust that has been stretched
and thinned
by back-arc
sional forces; or, (2) a newly emplaced lithosphere arising from active spreading. One may possibly distinguish between these two models by examining
exten-
heat flow
data from the Central Volcanic Region. One of the most impressive geophysical features of the Central Volcanic Region is its high heat output. The total natural output, almost all of which arrives at the surface via hydrothermal systems on the eastern side of the Central Volcanic Region 17PE
177O
0 L
Fig. 11. Map showing potash
andesites
area within the Central considered Volcanic
the asymmetrical
within
I I
distribution
and immediately Volcanic
Region.
adjacent
Region
coincides.
at its northern volcanism.
,
for both the K-Ar to the Central
ages (after CaIhaem,
Volcanic
Region,
The down flow region, which includes
in the text as the area of the present
B.P. axis of active andesite
50km
day heat source.
end, with Calhaem’s
The western
(1973) proposed
The vector arrow of 3 cm/year
and rate (at the Bay of Plenty coast) of migration
represents
for the axis of active andesitic
1973) of low
and the geothermal
the geothermal boundary position
for the 4 m.y.
the apparent volcanism.
region. is
of the Central direction
403
(Fig. 11) has been estimated
to be (3.3-3.5).
Allis, 1979). The hydrothermal about
systems
20 km2 in area, where
300°C is found.
upflowing
In the areas separating
tions in bore holes show the geothermal to be due to downflowing fields (Studt and Thompson,
meteoric
10’ W (Studt
consist
hot water
as high as
fields, temperature
observa-
to be close to zero. This is thought
water providing
1969; Thompson,
1969;
fields, each of
with a temperature
the geothermal gradient
and Thompson,
of small geothermal
the recharge
for the geothermal
1977).
Figure 11 shows the area of the Central Volcanic mal fields and the area of downflow as interpreted
Region containing the geotherfrom the distribution of bore
holes where a zero geothermal gradient was found. The assumed downflow region shown in Fig. 11 is roughly the same as that suggested by Allis (1979) and has an area of about 5000 km’. Taking the heat output of the geothermal systems to be 3.5.10’ W, the average heat flow for the eastern half (the downflow area) of the Central Volcanic Region is estimated at 700 mW/m’. Insufficient measurements do not allow an estimate of the average heat flow for the whole of the Central Volcanic Region. A heat flow of 700 mW/m’ 80-110 mW/m’ for continental
is considerably rifts (Morgan,
higher than the average value of 1982). It is also higher than values
reported for many oceanic back-arc basins (e.g. Anderson, 1975). Comparisons with oceanic back-arc basins may not be valid, however, as obtaining an estimate of the average heat flow from spot measurements in these basins effects of hydrothermal convection within the oceanic
is difficult because of the crust (Sclater, 1972). A
volcanic spreading centre that is above sea level, and thus has a readily available heat output estimate, is Iceland. Although Iceland occupies a different plate tectonic setting to that of the Central Volcanic Region, comparison of the pertinent thermal and kinematic properties of both regions is straightforward and is therefore useful. In Iceland the active volcanism and associated geothermal activity (Palmason and Saemundsson, 1974) as for the Central Volcanic Region, are concentrated in an elongated volcanic zone several tens of kilometres wide. Therefore, comparison of volcanic productivity and heat flow output values between the Central Volcanic Region
and Iceland
are normalised
to volcanic
productivity
and heat output
per km
(of strike) of volcanic zone; these data are given in Table 1. From Table 1 it can be seen that by summing the heat output of both the geothermal fields and that represented by volcanic eruptions, the total heat output for the Central Volcanic Region (30 MW/km) and Iceland (35 MW/km) are similar. This suggests that the mode of extension within the Central Volcanic Region is more akin to that of an active spreading centre rather than that of a continental rift where average heat flow values are typically 80-110 mW/m2 (Morgan, 1982). One qualitative proposal to explain the heat output of the Central Volcanic Region is the lithospheric spreading model of Calhaem (1973). He noted a decrease in age of low-potash andesites from west to east across the Central Volcanic Region (Fig. 11). Adopting an observation of Dickinson and Hatherton (1967) that the
404
potash content Zone,
of andesite
Calhaem
andesites
volcanoes
postulated
was accompanied
the Benioff Central
Zone,
Volcanic
is related
that the apparent
by a corresponding
and the eastern Region
to the depth of the subjacent migration
portion
accordingly
the last 4 m.y. by back-arc
southeastwards
of the North
is envisaged
spreading.
Benioff
with time of the low-potash migration
Island.
The wedge-shaped
to have developed
From the distribution
of both
entirely
within
and ages of the various
geothermal fields, in addition to the age distribution of the low-potash andesites. Calhaem suggests the spreading to be asymmetric with intrusion presently occurring along the eastern boundary of the Central Volcanic Region, and the newly emplaced lithosphere being in various stages of cooling to the west. A brief order of magnitude attempt can be made to quantify details and assumptions
inherent
this model. Specific
in the model are:
(1) The geothermal systems are driven by a heat source that consists of molten rock that has been intruded at the eastern margin of the Central Volcanic Region and is in progressively greater stages of cooling to the west. Thus the heat source is envisaged to be migrating southeastwards with time, as an accompaniment to the migration
TABLE
of the low-potash
andesites,
so that
as new geothermal
fields
on the
1
A comparison
of spreading
rates and thermal Central
Spreading
rate
7+4
Volcanic
mm/yr.
20-30 0.9’10
Rate of production of eruptives
data from the C.V.R. and Iceland Region
Iceland * Up to 20 mm/yr.
(Sissons. 1980) ***
mm/yr.
(Calhaem.
1973)
(double drift rate)
4 km3/yr. (Ewart
1.4.10
4 km’/yr.
et al., 1975) **
per
km of strike of volcanic zone Heat output
22 MW/km
from
geothermal
15 MW/km
**
activity
per km strike of volcanic
zone
Equivalent output
8 MW/km
heat
20 MW/km
(Stem, 1982) **
of
eruptives Sum of heat outputs
30 MW/km
35 MW/km
**
from geothermal activity and volcanic
extrusives
* All Icelandic
data from Palmason
** These calculations
*** S&son’s rate is based Calhaem’s Volcanic
and Saemundsson
(1974).
assume a strike length for the C.V.R. of 160 km.
rate is based
on an analysis
of repeated
on the apparent
migration
Region in the last 4 m.y.
geodetic rate
observations
of low-potash
that span andesites
the past 70 yrs.
across
the Central
405
eastern
boundary
become
active older fields to the west will shut down.
(2) The area of the present
day heat
source
is taken
downflow
area in Fig. 11, i.e. 5000 km2. The thickness
parameter
to be determined
(3) The considers
typical
lifetime
that the Wairakei
from the following
calculation.
of a geothermal
system
geothermal
to be the same
as the
of the heat source
is the
is 1 m.y.
Grindley
(1965)
field has been active for at least the past 0.5
m.y. on the basis of hydrothermally altered ignimbrite fragments found in a 0.5 m.y. old conglomerate. Given that Wairakei is now situated approximately mid-way between the eastern and western boundaries of the geothermal lifetime estimate for Wairakei of 1 m.y. appears reasonable.
area (Fig. 11) a total
(4) The average heat flow from the source over the past 1 m.y. is the same as the present day average of 700 mW/m2. With these assumptions the total heat required per unit area of source (Eo) to maintain the geothermal systems for 1 m.y. is: E, = 2.2. 1013 J/m2
The heat available heat = Vp(K+
(1)
from molten
rock is given by Jaegar (1968) as:
L)
(2) where V = volume; p = density; B = temperature drop; C = specific heat; L = latent heat. Taking a density for the intruded, molten rock of 2.6 . lo3 kg/m3, its temperature drop 1000°C and its specific and latent heat to be lo3 J kg-’ ‘C-’ and 4. lo5 J/kg respectively, it can be calculated from eqn. (1) and (2) that the volume of molten rock required to maintain the geothermal systems for 1 m.y is about 3.0. lo4 km3. Given that the area of the heat source is assumed fixed at 5000 km2 its thickness must be about 6 km. In the above account no consideration is given to the heat required to produce the extensive rhyolite eruptives. The prevailing view, with regard to the genesis of the rhyolites
erupted
from the Central
Volcanic
Region,
is that they arise from a partial
melt of greywacke rocks at a shallow level (Ewart and Stipp, 1968; Reid, 1983). If so, the heat source would have to be about 9 km thick, as the heat associated with the erupted rocks from the Central Volcanic Region amounts to about one third of the heat discharged by the geothermal systems (Table 1). With the above model, the observed heat output can be accounted for if about two-thirds of the 13 km (15 km less the 2 km of superficial pyroclastic cover) thick crust consists of cooling intruded rock. But the result rests on the initial assumptions, and other more complex models to explain the large heat output may be possible. CONCLUSIONS
A major contribution Region displays many
of this study has been to show that the Central Volcanic of the characteristics commonly associated with actively
spreading oceanic back-arc basins-i.e. a thin crust, low seismic velocities crust and upper mantle, possible evidence for upper crustal magma tensional
tectonics
to extension
and a very high heat flow. The mechanical
within
the Central
extension
manifested
spreading
accompanied
Volcanic
by the stretching
Region
by the intrusion
process that gives rise
is, nevertheless,
of an old continental of a new lithosphere
within the chambers,
still contentious.
lithosphere, the principal
Is
or is active means of
extension? The case for the active spreading alternative, as argued in this paper, rests on both the high heat flow and the absence of Mesozoic greywacke of greywacke derived
sediments
in any of the deep bore holes that have been drilled
median area of the Central Volcanic Region. The critical likely come from further deep drilling associated with exploration programme. It is unlikely that extension
or spreading
within
taking place in a simple, symmetrical manner Indeed, one of the consequences of the Central
within
the
test will, however, most the current geothermal
the Central
Volcanic
about a central Volcanic Region
Region
is
spreading ridge. being a back-arc
basin that is above sea level, is that we can observe a strong asymmetry in the pattern of heat discharge. The grouping of the active geothermal fields within the eastern half of the Central Volcanic Region, and on a more detailed scale, the tendency for the individual geothermal fields to become younger to the east (Calhaem, 1973), points towards intrusion now taking place close to the line of presently active andesite volcanoes on the eastern margin of the Central Volcanic Region. Although asymmetric spreading has been claimed for some oceanic back-arc basins on the basis of magnetic anomaly lineations (Barker and Hill, 1980), it is generally regarded that back-arc spreading, like its mid-oceanic counterpart, is symmetric
(Weissel,
1981).
Asymmetric
spreading
within
the
Central
Volcanic
Region should be amenable to testing by microearthquake studies (e.g., Evison et al., 1976). If intrusion is occurring along the eastern margin of the Central Volcanic Region, and the newly emplaced lithosphere is at progressively more advanced stages of cooling to the west, then the thickness of the new lithosphere should also increase
to the west.
ACKNOWLEDGEMENTS
This paper is partly derived from a PhD thesis undertaken at the Institute of Geophysics, Victoria University of Wellington. I thus readily acknowledge the assistance of my supervisors Prof. F.F. Evison, Dr R.R. Dibble and Dr T. Hatherton. In particular, I am indebted to Prof. Evison for many useful discussions with regard to the heat flow problem for the Central Volcanic Region. 1 also thank Drs R.I. Walcott, and F.J. Davey of Geophysics Division, D.S.I.R. for their critical reviewing
of this paper.
407
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A BACK-ARC
B.V.. Amsterdam
BASIN
THE CENTRAL
- Printed
FORMED
VOLCANIC
in The Netherlands
WITHIN
CONTINENTAL
LITHOSPHERE:
REGION OF NEW ZEALAND
T.A. STERN Geophysics Division, D.S. I. R., Wellrngton (New Zealand) (Received
December
2, 1983; revised version accepted
August
23. 1984)
ABSTRACT
Stern, T.A., 1985. A back-arc New Zealand. Zones.
Tectonophysics,
The Central
basin formed within continental
In: K. Kobayashi
Volcanic
Region
Island of New Zealand
-the
Havre
possesses
back-arc
basins.
mantle
respectively),
the crust
seismic
is anomalously
represents
of New
velocities
Volcanic
and Processes
continental
of heat is required. spreading
Zealand.
characteristics
thin for its continental
a site of active back-arc
and tensional
Region of
in Subduction
setting,
geothermal
Volcanism
the
basin
with oceanic
of 7.4 and 3.95 km/s
and the heat flow, almost is about
to account
therefore,
within a continental
within
back-arc
is predominantly
often associated
systems,
heat flow). In order It is proposed,
tectonics
of a young oceanic
are low (P, and S, velocities
of hot water through
than normal
flow a large scale mass transfer Region
structure
many of the geophysical
in the discharge
some twelve times greater
volcanism
to be a direct continuation
the continental
yet the region
which is expressed
is an area of Quaternary
that appears
rhyolitic,
Upper
the Central
Structures
112: 385-409.
North
Trough-into
lithosphere:
and I.S. Sacks (Editors),
700 mW/m2
all of (i.e.
for such a high heat
that the Central
Volcanic
lithosphere.
INTRODUCTION
Evidence
from seismicity
(Adams
and Ware, 1977; Reyners,
the subducted Pacific plate lies beneath the North the South Island of New Zealand (Fig. 1). Features
1980) indicates
that
Island and the northern end of typically found at a convergent
plate boundary such as an oceanic trench, andesitic volcanism, dipolar gravity anomalies, and crustal and upper mantle seismicity are all found in the North Island. Figure 2 shows associations that can be made (e.g. Cole and Lewis, 1981) between various geological structures of the North Island and Karig’s (1974) generalised structural model of a western Pacific island arc system. At some convergent margins an area is found behind, or landward of the subduction zone where tensional tectonics and high heat flow are observed. These are marginal (Karig, 1974) or back-arc (Uyeda, 1977) basins. The term back-arc basin is used here. The Central Volcanic Region of the North Island (Figs. 1 and 2) C&to-1951/85/$03.30
0 1985 Elsevier Science Publishers
B.V.
386
I
‘\
179E
SOUTH
1 cop
\
I
I
f/I
’
FIJI BAStN
PLATE
Fig. 1. Locality Havre Trough. assuming
map showing Convergence
the Indian
Havre Trough.
the position rates between
of the Central
plate to be fixed. Double
Eathymetry
in km.
Volcanic
Region
the Pacific plate and the Indian headed
arrows
represent
(C.V.R.)
with respect
plate (in mm/yr) back-arc
spreading
to the
are shown within
the
387
km
L -500
I
I
-400
-300
1
-200
,
-100
-50
0
+50
+100
+200
lwclneters
Fig. 2. Diagram structures
showing
of the North
some of the associations Island
that can be made
(top) and the generalised
island arc system as given by Karig (1974). Upright
triangles
structural represent
between
model
low-potash
have been active within
the past 10,000 years. The area1 extent of the Central
and the Taupo
Zone (T.V.Z.) are shown.
Volcanic
the various
(below)
geological
of a western
andesite Volcanic
Pacific
volcanoes
that
Region (C.V.R.)
388
is an area of crustal appears
extension
(the oceanic back-arc There
and high heat flow, and as Karig (1970) notes,
to be a direct continuation basin associated
is a fundamental
Volcanic
Region,
spreading
back-arc
into New Zealand
difference
however.
Whereas
basin (Karig,
of the young
with the Tonga-Kermadec between
the Havre
the Havre Trough
trench-arc
Trough
is thought
1974) within an oceanic
it
Havre Trough
and
system).
the Central
to be an actively
area, the Central
Volcanic
Region is classified by Karig and Jensky (1972) as a “ volcanic-tectonic rift zone” situated within a continental environment. Karig and Jensky postulate a general progression, with increasing amounts of extension, from a volcano-tectonic rift, which they also describe as the earliest phase of extension in a back-arc basin, to an actively spreading back-arc basin where new lithosphere is created. The purpose of this paper is to examine the response of continental lithosphere to back-arc processes by presenting interpretations data that bear upon the crust and upper mantle
of seismic, gravity and heat flow structure of the Central Volcanic
Region.
GEOLOGICAL
PROVINCES WITHIN THE CENTRAL
NORTH ISLAND
The area Karig (1970) referred to as the Taupo Volcanic Zone closely resembles what is called here the Central Volcanic Region. Figure 2 shows the area1 extent and location of what is commonly termed by New Zealand geologists (e.g., Healy, 1962; Cole, 1979) the Taupo Volcanic Zone (T. V.Z in Fig. 2). This distinction in terminology is important; past geological and geophysical studies have tended to concentrate on the Taupo Volcanic Zone, for here the economically important geothermal activity occurs and the impressive exposures of Quaternary rhyolitic volcanics
can
be seen. About
12,000 km’ of Quaternary
rhyolitic
volcanics
are
contained within the Taupo Volcanic Zone (Cole, 1979) the average heat flow of which is estimated to be about 700-800 mW/m’ (Studt and Thompson, 1969; Allis. 1979).
This
continental extension
value
is approximately
norm of 60 mW/m’. in historical
twelve
The Taupo
times (Walcott,
to fourteen Volcanic
1978; Sissons,
times
greater
than
the
Zone has also been an area of 1980). Analysis
of repeated
geodetic surveys spanning the past 70 years, shows significant shear strain rates that are interpreted by Sissons to be due to (back-arc) spreading in a NW-SE azimuth at a rate of 7 f 4 mm/yr. The Taupo Volcanic Zone forms the eastern half of a larger area of Quaternary volcanism in the North Island-the Central Volcanic Region-with which it shares a common eastern boundary. The western boundary of the Central Volcanic Region separates extrusions of Quaternary rhyolite to the east from Mesozoic greywacke and Tertiary rhyolites to the west (Fig. 3). So defined, the western boundary of the Central Volcanic Region is based on surface geology. However, the boundary also receives emphasis from gravity and seismic data.
Greywacke
RhyoMe domes (Q=Quaternary
T=Ter&y
Andesite volcanoes -active in past 18yrs 0,
,
5;Okm
BAY OF PLENTY
White Island A .:*
TAFiANAKl
1750
Fig.
3. Generalised
greywacke
ranges
the area within basalt about
200-300
cone-shaped
map
and immediately
and andesite
C.V.R. is a broad
geological
of the Central
that flank the C.V.R., outcrops
occur
North
and the Quatemary
adjacent
to the C.V.R.
in only minor
Island
an ignimbrite
slice of the axial greywacke
converging
the C.V.R.
Much of
is mantled
by extensive
ignimbrite
deposits;
volumes.
sheet. The plateau ranges
southwardly
domes within
The Kaingaroa
area where drilling and seismic work (Hochstein m beneath
showing
rhyolite
and Hunt,
has been
that has subsided
towards
Plateau
to the east of the
1970) show greywacke
interpreted the C.V.R.
(Healy,
to lie
1962) as a
390
CRUSTAL
STRUCTURE
Upper crustal structure Rock samples from drill holes (Grindley, 1965, 1982) and seismic refraction studies (Stern, 1982; Robinson et al., 1981) (Fig. 4) indicate that the upper 2 km of crust for the Central Volcanic Region consists of low-density, low-velocity volcanic rocks (the terms density and velocity used here refer to saturated density and compressional seismic velocity respectively). At a depth of 1.8 to 2.2 km a “basement” rock unit of velocity 4.8-5.5 km/s has been identified at three locations within the Central Volcanic Region (Fig. 4). The rock type represented by this basement
layer is debatable.
In drill holes at Wairakei
and Mokai (Fig. 3) andesite
and welded ignimbrite with densities of 2.55-2.65 Mg/m’ were encountered depth of about 2.1 km (Grindley, 1982; Stern, 1982). Rocks in this density would satisfactorily
account
for the observed
basement
At Broadlands (Fig. 3), however, Mesozoic greywacke depths of l-2 km (Hochstein and Hunt, 1970).
velocities
of 4.8-5.5
at a range km/s.
rocks were encountered
at
It has generally been assumed by New Zealand geologists and geophysicists (e.g., Grindley, 1965; Hochstein and Hunt, 1970, Rogan, 1982) that a continuous, albeit extensively faulted, greywacke basement extends beneath the Central Volcanic Region. An alternative view has, however, been put forward by Calhaem (1973) and Evison et al. (1976). They consider that the Central Volcanic Region is an area of active back-arc spreading and that a new lithosphere is in the process of being created. Under this proposal basement rocks will predominantly consist of intruded volcanic and igneous rocks that are in various stages of cooling. Furthermore, Evison et al. note that the drill holes at Broadlands are situated within 5 km of the eastern boundary of the Central Volcanic Region, and thus contend that the severely faulted greywacke surface found at depth here represents blocks that have been spalled from the eastern greywacke ranges during previous rifting episodes. This question of what rock type constitutes basement for the Central Volcanic Region is an important one in terms of unravelling the tectonic history of the region; it will be taken up again later in this paper when heat flow data are considered. Geophysical studies by Rogan (1982) and Stern (1982) point to the possible existence of molten rhyolite magma within the upper crust of the Central Volcanic Region. For example, consider the residual gravity anomaly map shown in Fig. 4b. These residual anomalies were obtained by removing a regional field from the observed gravity anomaliesthe regional field was numerically derived by fitting a polynomial surface of low-order to gravity values situated on greywacke outcrops found each side of the Central Volcanic Region (Fig. 3) (Stern, 1979). The residual field of Fig. 4b, reflects, therefore, the anomalous density distribution in the upper crust of the Central Volcanic Region, with respect to the density distribution in the greywacke crust elsewhere within the North Island. The residual anomalies are
391
0
3/TAUPO
2/MANGAKINO
I /ROTOEHU
Oi8
0.8
3.0 km/:
0.8
4 3.2
km/s
3.2
kmls
5.5
km/s
0.5 3.2 km/s
1.0
i.0 km/s
3.2 km/s
1.5
2.0
5.5 km/s 4.8 km/s
2.5
RESIDUAL
3
GRAVITY
ANOMALIES
contour
20
Fig
4. a. Interpretations
Volcanic
Region.
Rotoehu
(I ), Mangakino
4a for Rotoehu
of three,
b. Residual
gravity
reversed, anomaly
(2) and Taupo
and Mangakino
seismic
refraction
(3) refraction
refer to the structure
surveys
km
surveys
map of the C.V.R.
1@mgal)
intew
(after
carried Stem,
out within
the Central
1979). Positions
are shown. The velocity
at the south ends of these lines.
columns
of the in Fig.
392
almost wholly negative - 55 to -60 determined
mGal
2 km thickness
4) is estimated residual deeper
within the Central of volcanic
Region
with minimum
gravity
effect of the seismically
rocks at Taupo
and Mangakino
to be only - 35 to - 40 mGa1. The remaining
gravity than
Volcanic
(Fig. 4b). Yet, the anomalous
anomaly
2 km.
must, therefore,
Interpretations
be explained
of these
(Figs. 3 and
- 15 to - 25 mGal of
by negative
“secondary”
values of
mass anomalies
residual
anomalies
at
Mangakino and Taupo (Stern, 1982) have been made in terms of discrete bodies of molten magma at depths of 3 to 10 km. Although it is acknowledged that other interpretations are possible, the molten rhyolite hypothesis receives support from the geochemical studies of Ewart et al. (1975). They show from phenocryst equilibration studies that the rhyolitic rocks erupted from the Central Volcanic Region had, prior to eruption, equilibrated in magma chambers at depths of 7--8 km. Moreover, their data suggest temperature and density ranges for the rhyolite eruption, of 725”-915°C and 2.0-2.38 Mg/m” respectively.
magma,
prior
to
Deep crustal structure
Studies made on dispersive earthquake waves (Thompson and Evison, 1962) and the average value of Bouguer gravity anomalies in New Zealand (Reilly, 1962) point to a crustal thickness of 30-40 km for New Zealand as a whole. There may be considerable variation from this average, however, particularly in the North Island where both Bouguer and isostatic gravity anomalies vary from - 150 to 60 mGa1. Moreover, from a study of earthquake travel-times, Haines (1979) finds that the velocity of upper mantle, compressional seismic waves (P,) varies from 8.5 to 7.4 km/s
within
between
depth
the North
Island.
to the upper mantle
Woollard
(1970)
and Pn velocity
observes
a linear
for North
America,
relationship thus the P,
velocity of 7.4 + 0.1 km/s that Haines associates with an area that covers much of the Central Volcanic Region, suggests that the crust may be thinner here than elsewhere
in the North
Island.
Figure 5 shows the travel-time data and plane-layer interpretation for a two-way deep seismic sounding survey carried out within the Central Volcanic Region. One 400 kg shot was fired in the western bay of Lake Taupo and three 400 kg shots were fired within the vicinity of White Island. Seismograph stations were distributed between these shot points. The interpretation given in Fig. 5 is a horizontally layered model because the apparent velocity and time intercepts of the straight-line segments for both branches of the time-distance graph are about equal. The velocity of 7.4 rt 0.2 km/s at a depth of 15 km beneath the Central Volcanic Region (Fig. 5) is the same as the velocity observed by Haines (1979) for the region, and also is similar to P,, velocities detected in other parts of the world where magmatic activity high heat flow, extension and/or spreading are occurring. For example, P, velocities of 7.1-7.5 km/s, coupled with a thinner than normal crust,
393
CENTRAL
VOLCANIC
REGION TRAVEL-TIME
DATA
0 first arrival
cl-
+
second arrival
o earthquake
travel time(WTZ-
KRP)
3-
o-
L
0
White Island
~STAN~E(km)
Lake TAUPO
PLANE LAYER SOLUTION
O3.0 km/s
2-
6.3 km/s
42
6-
$
8-
: t
lo-
6.0 km&
1214-
..- 14.8
km
7.4 km/s Fig. 5. Data and plane layered Central
Volcanic
Karapiro
(KRP)
beneath
Whakatane
Region.
solution
for a long-range
The open circle
(Fig. 6) for an eathquake (Dr M. Reyners,
represents
seismic refraction
on the 28th February
N.Z. Seismological
survey carried
the travel time between Observatory,
out within
the
(WTZ)
and
Whalcatane
1983 located
apprbximately
pers. commun.,
1983).
1 km
394
are observed in Iceland Atlantic Ridge (Talwani western
(Bott, 1965) the Africa Rifts (Fairhead, et al., 1965) and the Basin and Range
U.S.A. (Thompson
1976). the Midprovince of the
and Burke, 1974; Scholz et al., 1971). Accordingly,
results of Fig. 5 are interpreted in terms of a 15 km thick crust underlain upper mantle with an anomalously low velocity of 7.4 km/s.
A GRAVITY
INTERPRETATION
ADJOINING
MODEL
FOR THE CENTRAL
VOLCANIC
the by an
REGION
AND
AREAS OF THE NORTH ISLAND
A summary
of heat flow (Pandey,
and this paper),
crustal
thickness
1981)
estimates
seismic sounding
data (Garrick,
from S to P conversion
studies
1968; (Smith,
1970 and W.D. Smith, pers. commun., 1983) and seismic attenuating properties of the upper mantle (Mooney, 1970) is given in Fig. 6. East and southeast of the Central Volcanic Region, heat flow is lower than normal, the crustal thickness is about 35 km, and no attenuating region within the upper mantle is observed. West of the Central Volcanic Region heat flow is higher than normal. the upper mantle attenuates seismic waves, and the crust is thinner than normal continental crust (apart from the S to P conversion estimate of 28 km at TUA (Fig. 6), Eiby (1958) reports a crustal thickness estimate of 23 km just to the north of Auckland (AUC in Fig. 6) based on unpublished seismic reflection data). Thus on the basis of the data given in Fig. 6 three distinct tectonic provinces within the North Island can be identified: the Northwest North Island, the Central Volcanic Region, and the East Coast-Wanganui Basin area. This partition of the North Island, as first noted by Hatherton (1970), is also reflected in both the Bouguer and isostatic gravity anomalies of the North Island (Fig. 7). The area of the North
Island
associated
marked
by attenuation
within
with positive gravity anomalies,
the upper
the Central
mantle
Volcanic
and
thin
crust
is
Region is an area of
relatively low gravity flanked by gravity highs, and the East Coast-Wanganui Basin area is dominated by an intense gravity low. The cross-sections AA’ and BB’ of Fig. 7 are shown in Fig. 8. Of note in these cross-sections
is that
markedly
different
gravity
anomaly
profiles
are generated
above different portions of the subducted Pacific Plate. Profile AA’ shows the characteristic dipolar gravity anomaly pattern commonly observed across convergent margins (Hatherton, 1969); a gravity low over the point of major downturn of the subducted plate coupled with a broad positive gravity anomaly over the back-arc area. The gravity anomaly of profile AA’ does, however, differ from the dipolar anomalies observed over other convergent margins like, say, Japan. In the Japanese situation the gravity high peaks within the vicinity of the volcanic arc (e.g. Yoshii, 1979), whereas for the North Island the maximum in the gravity high is reached some 50-100 km westward, or landward, of the volcanic arc. Hatherton (1970) interpreted a profile across the North Island slightly to the
395
f%km)-S
to P conversion
ECZ 0 =
depth
SEISMOGRAPH STATION
= UPPER MANTLE ATTENUATING REGION ~~11
l-y---l 6. 2
- 15 km
7.0 km/s
NORTHWEST NORTH I Mea,fbv,=80”22
mWi&53)
NI
heat
flow(convective)
(49km)
fo
Ir SOUTH WANGANUl i TARANAK$S”JMNG~~~~,,,>
1OOkm
/;F -10
-20
35 km 118’)
I Fig. 6. Geophysical
data bearing
flow data from Pandey depths Garrick
(Smith,
upon the crust and upper
(1981), upper mantle
attenuating
1970 and W.D. Smith, pers. commun.,
(1968) and this paper.
mantle
regions
structure
of the North
from Mooney
(1970), S-P
1983), and seismic refraction
I Island.
Heat
conversion
interpretations
from
TATIC
GRAVfTY
ANOM
area
beneath
subducted
which
Pacific
is present
<
-50
the
plate il.l
mgal
NORTH ISLAND
Fig. 7. isastatic
{Airy-Heiskanen,
1965). Area beneath direction
which
of dip for the subducted
refer to profites in Fig. 8.
T-
30 km) gravity
the subducted
Pacific
plate isnorthwest,
anomaly
plate
map of the
is inferred
i.e. in the direction
North Island (after Reilly,
to be present
is shown.
Note
the
of A’-,4 and B’- 8. AA’ and BB’
AA’
in gravity
anomalies
and BE’ of Fig. 7 showing
Note the gross dissimilarity
Fig. 8. Profiles
between
the position
of the subducted the two profiles.
-d)
studies, and associated
PROFILEtB
from seismicity
WANGANUI
Pacific plate, as inferred
250
250
50
0
200
Bay
2co
Hawke
PROFILECA-h
150
,
BAY
150
C.V.R.
HAWKE
100
I
-
too
Waikato
WAIKATO
gravity
anomalies.
5
north of AA’ (Fig. 7). He ascribed a mass deficiency
within
component
to a mass excess within
BB’
S), however,
(Fig.
seismological present North
evidence
the negative
the upper exhibits
(Adams
here. This observation
component
of the dipolar
80 km of the subducted
to
the plate for the depth range 8OG250 km. Profile no such
dipolar
gravity
and Ware, 1977) indicates suggests
Island are not so much controlled
anomaly
even
the subducted
that the principal
gravity
by mass anomalies
the subducted Pacific plate, but are more influenced mantle inhomogeneities within the overriding Indian Figure 9 shows an interpretation
anomaly
plate. and the positive
directly
by crustal plate.
though
plate to be
anomalies associated
thickness
of the with
and upper
of the 280 km long profile AA’ of Fig. 7. In this
interpretation a model for the suducted plate is included that used by Sager (1980) for the Mariana Arc structure.
that is similar in concept to A negative gravity effect is
ascribed to a 6 km thick oceanic crust layer, that caps the subducted plate, being taken down to a depth of 80 km. At this depth the oceanic crust is assumed to transform to eclogite of density 3.56 Mg/m3. The eclogite layer continues as the upper 6 km of the subducted plate to a depth of 150 km, thus constituting part of the positive mass anomaly directly associated with the subducted plate. The remainder of the positive mass anomaly is due to that part of a 80 km thick, relatively cool subducted plate, less the eclogite layer, in the depth range of 100 to 250 km. A density contrast of to.02 MS/m3 was found to give the best fit for this part of the subducted plate. The Central Volcanic Region is represented in Fig. 9 by three isodensity layers. The top layer represents the 2 km thick section of low-density, low-velocity volcanic rocks as detected
in drill holes and by seismic refraction
data (Fig. 4). Between 2 and
13 km is a 2.60 Mg/m3 layer. The choice of this lower than normal crustal density, with respect to the 2.75 Mg/m’ density used to represent the crust elsewhere in the North Island (Fig. 9), was found necessary to ensure a satisfactory fit between observed and calculated gravity anomalies. However, some support for this choice of density
comes from the observed
seismic
velocities
in this layer of 5.3-6.0
km/s
(Fig. 5), which are also somewhat lower than normal for continental crust. Profile AA’ passes through the Central Volcanic Region on a line just north of Lake Taupo where a deep gravity low of short wavelength is observed (Fig. 4b). In order to accommodate this gravity low a negative mass anomaly of 2.35 Mg/m’ is included within the 2.60 Mg/m’ layer, and is interpreted to represent a volume of molten or partially molten rhyolite (see discussion in previous section). The 3.18 Mg/m3 upper mantle density for the Central Volcanic Region is chosen on the basis of the low (7.4 km/s) upper mantle seismic velocity (Fig. 5), although there is no data that controls the depth to which the 7.4 km/s layer extends. The reduction of the upper mantle density for the area to the west of the Central Volcanic Region from 3.40 to 3.30 Mg/m3 is chosen on the basis of the low (7.887.9 km/s) P, velocities for this area, relative to the more “normal” P, velocities of 8.1-8.5 km/s observed for the area to the east of the Central Volcanic Region
COAST
’
I
side.
left-hand
/
are calculated
II
/-
COAST
Coast
with respect to the standard
density
_
_
Mg/m’
-
basin
is model shown on the
of +0.02
_
sedimentary
plate with a density
1
EAST
of the subducted
~~______~________~~~~_________~__~~_~~________,___
2p0
AA’ of Fig. 7. The portion
*c’
c:
;(_d ’0
to a depth of 250 km. Crust. and upper mantle mass anomalies
field for the profile
f
3.18
of the Bougucr
anomaly
LK”5
as extending
’ //
-’
modelled
150 km
L.IP
(2.35)
C.V.R.
Fig. 9. Interpretation
MODEL
DENSITY
ll STANDARD
3.3
1
3.4
Y/A V//‘/.
WEST
%
(Haines, 1979). The crustal thickness and mantle densities of the western and eastern North Island, and the Central Volcanic Region have all been adjusted in such a fashion
that
pensation
between
achieved Region
local isostatic
compensation
the western
at a depth
North
is maintained.
Island
and the Central
of 30 km and compensation
and the eastern
North
Island
On the East Coast of the North and Lewis (1981) consider
is achieved
Island
More
between
specifically, Volcanic
the Central
com-
Region
is
Volcanic
at a depth of 100 km.
is the East Coast Sedimentary
this basin to be part of an accretionary
Basin. Cole
prism containing
sediments whose ages span from early Tertiary to Quaternary. The exact depth of the basin is unknown but is modelled in Fig. 9 as a 5 km deep basin with a density contrast of -0.25 to - 0.35 Mg/m3. Compensation for this basin is presumably achieved in a regional sense by the excess mass of the subducted plate. That is, the basin is maintained by “Trench flexure” as defined by Davies (1981). The calculated gravity anomaly curve in Fig. 9 is comprised of two principal components: (1) The effect of mass anomalies directly associated with the subducted Pacific plate. (2) The gravity effect of anomalous upper mantle and crustal structures within the overlying Indian Plate. These two components are shown plotted separately in Fig. 10. Of note in Fig. 10 is that the calculated effect of the upper mantle and crustal structure of the Indian plate accounts for much of the character of the observed gravity anomaly curve (Fig. 9). In particular, it explains about 40 of the 50 mGa1 back-arc gravity highs observed west of the Central Volcanic Region. This finding, therefore, concurs with that of Watts
and
Talwani
(1974,
1975) who
argue,
on the
basis
of an elastic
plate
approximation to a subducted plate, that some back-arc gravity highs are due not so much to the excess mass of the subducted plate, but to gravity anomaly effects associated
with back-arc
extension.
There are, however, many speculative aspects of the model in Fig. 9. In particular, the depth at which oceanic crust transforms to eclogite is debatable; some authors (e.g., Grow, 1973) place the transformation as shallow as 30 km. Moreover, the model for the crust and upper mantle of the North Island is essentially a static, locally compensated one. In a dynamically active area such as that found at a subduction zone, one cannot rule out the possibility that some lithospheric mass anomalies may be supported by dynamic flow in the asthenosphere rather than by the finite strength of the lithosphere (McKenzie, 1967; Lambeck, 1972). Nevertheless, the import of the gravity anomaly analysis presented in Figs. 9 and 10 is that the gravity anomaly field of the North Island appears to be more dominated by those effects associated with extension within the overlying Indian plate, rather the gravity effect of mass anomalies within the subducted Pacific plate.
than
\d
calculated
(i) the calculated
50
gravity effect of mass anomalies
Fig. 10. A graph showing:
0
u
coast
---
V Vest
-40
-20
o-
20
40
100
effect
the
subducted
150
directly
assockted
plate
200
with the subducted
plate.
uppermantle structure
Jc= C.V.R. =$I
Y__;.
of
gravity effect of the crust and
gravity
1
gravity
Centrasl
1. \ the
I Y
for
combined
crust
250
North
3001km
Island,
Island
structure
North
mantle
western
upper
the
coast II
and
and
and northwestern
East
Region
of
of the western
Volcanic
effect
and (ii) the g
402
HEAT
FLOW AND LITHOSPHERIC
The principal
EVOLUTION
feature of the structural
model developed
Region
in the previous
section,
mantle.
Such a structure
could be reconciled
(1) A greywacke
is the thin crust
for the Central
underlain
Volcanic
by low-velocity
upper
with either:
crust that has been stretched
and thinned
by back-arc
sional forces; or, (2) a newly emplaced lithosphere arising from active spreading. One may possibly distinguish between these two models by examining
exten-
heat flow
data from the Central Volcanic Region. One of the most impressive geophysical features of the Central Volcanic Region is its high heat output. The total natural output, almost all of which arrives at the surface via hydrothermal systems on the eastern side of the Central Volcanic Region 17PE
177O
0 L
Fig. 11. Map showing potash
andesites
area within the Central considered Volcanic
the asymmetrical
within
I I
distribution
and immediately Volcanic
Region.
adjacent
Region
coincides.
at its northern volcanism.
,
for both the K-Ar to the Central
ages (after CaIhaem,
Volcanic
Region,
The down flow region, which includes
in the text as the area of the present
B.P. axis of active andesite
50km
day heat source.
end, with Calhaem’s
The western
(1973) proposed
The vector arrow of 3 cm/year
and rate (at the Bay of Plenty coast) of migration
represents
for the axis of active andesitic
1973) of low
and the geothermal
the geothermal boundary position
for the 4 m.y.
the apparent volcanism.
region. is
of the Central direction
403
(Fig. 11) has been estimated
to be (3.3-3.5).
Allis, 1979). The hydrothermal about
systems
20 km2 in area, where
300°C is found.
upflowing
In the areas separating
tions in bore holes show the geothermal to be due to downflowing fields (Studt and Thompson,
meteoric
10’ W (Studt
consist
hot water
as high as
fields, temperature
observa-
to be close to zero. This is thought
water providing
1969; Thompson,
1969;
fields, each of
with a temperature
the geothermal gradient
and Thompson,
of small geothermal
the recharge
for the geothermal
1977).
Figure 11 shows the area of the Central Volcanic mal fields and the area of downflow as interpreted
Region containing the geotherfrom the distribution of bore
holes where a zero geothermal gradient was found. The assumed downflow region shown in Fig. 11 is roughly the same as that suggested by Allis (1979) and has an area of about 5000 km’. Taking the heat output of the geothermal systems to be 3.5.10’ W, the average heat flow for the eastern half (the downflow area) of the Central Volcanic Region is estimated at 700 mW/m’. Insufficient measurements do not allow an estimate of the average heat flow for the whole of the Central Volcanic Region. A heat flow of 700 mW/m’ 80-110 mW/m’ for continental
is considerably rifts (Morgan,
higher than the average value of 1982). It is also higher than values
reported for many oceanic back-arc basins (e.g. Anderson, 1975). Comparisons with oceanic back-arc basins may not be valid, however, as obtaining an estimate of the average heat flow from spot measurements in these basins effects of hydrothermal convection within the oceanic
is difficult because of the crust (Sclater, 1972). A
volcanic spreading centre that is above sea level, and thus has a readily available heat output estimate, is Iceland. Although Iceland occupies a different plate tectonic setting to that of the Central Volcanic Region, comparison of the pertinent thermal and kinematic properties of both regions is straightforward and is therefore useful. In Iceland the active volcanism and associated geothermal activity (Palmason and Saemundsson, 1974) as for the Central Volcanic Region, are concentrated in an elongated volcanic zone several tens of kilometres wide. Therefore, comparison of volcanic productivity and heat flow output values between the Central Volcanic Region
and Iceland
are normalised
to volcanic
productivity
and heat output
per km
(of strike) of volcanic zone; these data are given in Table 1. From Table 1 it can be seen that by summing the heat output of both the geothermal fields and that represented by volcanic eruptions, the total heat output for the Central Volcanic Region (30 MW/km) and Iceland (35 MW/km) are similar. This suggests that the mode of extension within the Central Volcanic Region is more akin to that of an active spreading centre rather than that of a continental rift where average heat flow values are typically 80-110 mW/m2 (Morgan, 1982). One qualitative proposal to explain the heat output of the Central Volcanic Region is the lithospheric spreading model of Calhaem (1973). He noted a decrease in age of low-potash andesites from west to east across the Central Volcanic Region (Fig. 11). Adopting an observation of Dickinson and Hatherton (1967) that the
404
potash content Zone,
of andesite
Calhaem
andesites
volcanoes
postulated
was accompanied
the Benioff Central
Zone,
Volcanic
is related
that the apparent
by a corresponding
and the eastern Region
to the depth of the subjacent migration
portion
accordingly
the last 4 m.y. by back-arc
southeastwards
of the North
is envisaged
spreading.
Benioff
with time of the low-potash migration
Island.
The wedge-shaped
to have developed
From the distribution
of both
entirely
within
and ages of the various
geothermal fields, in addition to the age distribution of the low-potash andesites. Calhaem suggests the spreading to be asymmetric with intrusion presently occurring along the eastern boundary of the Central Volcanic Region, and the newly emplaced lithosphere being in various stages of cooling to the west. A brief order of magnitude attempt can be made to quantify details and assumptions
inherent
this model. Specific
in the model are:
(1) The geothermal systems are driven by a heat source that consists of molten rock that has been intruded at the eastern margin of the Central Volcanic Region and is in progressively greater stages of cooling to the west. Thus the heat source is envisaged to be migrating southeastwards with time, as an accompaniment to the migration
TABLE
of the low-potash
andesites,
so that
as new geothermal
fields
on the
1
A comparison
of spreading
rates and thermal Central
Spreading
rate
7+4
Volcanic
mm/yr.
20-30 0.9’10
Rate of production of eruptives
data from the C.V.R. and Iceland Region
Iceland * Up to 20 mm/yr.
(Sissons. 1980) ***
mm/yr.
(Calhaem.
1973)
(double drift rate)
4 km3/yr. (Ewart
1.4.10
4 km’/yr.
et al., 1975) **
per
km of strike of volcanic zone Heat output
22 MW/km
from
geothermal
15 MW/km
**
activity
per km strike of volcanic
zone
Equivalent output
8 MW/km
heat
20 MW/km
(Stem, 1982) **
of
eruptives Sum of heat outputs
30 MW/km
35 MW/km
**
from geothermal activity and volcanic
extrusives
* All Icelandic
data from Palmason
** These calculations
*** S&son’s rate is based Calhaem’s Volcanic
and Saemundsson
(1974).
assume a strike length for the C.V.R. of 160 km.
rate is based
on an analysis
of repeated
on the apparent
migration
Region in the last 4 m.y.
geodetic rate
observations
of low-potash
that span andesites
the past 70 yrs.
across
the Central
405
eastern
boundary
become
active older fields to the west will shut down.
(2) The area of the present
day heat
source
is taken
downflow
area in Fig. 11, i.e. 5000 km2. The thickness
parameter
to be determined
(3) The considers
typical
lifetime
that the Wairakei
from the following
calculation.
of a geothermal
system
geothermal
to be the same
as the
of the heat source
is the
is 1 m.y.
Grindley
(1965)
field has been active for at least the past 0.5
m.y. on the basis of hydrothermally altered ignimbrite fragments found in a 0.5 m.y. old conglomerate. Given that Wairakei is now situated approximately mid-way between the eastern and western boundaries of the geothermal lifetime estimate for Wairakei of 1 m.y. appears reasonable.
area (Fig. 11) a total
(4) The average heat flow from the source over the past 1 m.y. is the same as the present day average of 700 mW/m2. With these assumptions the total heat required per unit area of source (Eo) to maintain the geothermal systems for 1 m.y. is: E, = 2.2. 1013 J/m2
The heat available heat = Vp(K+
(1)
from molten
rock is given by Jaegar (1968) as:
L)
(2) where V = volume; p = density; B = temperature drop; C = specific heat; L = latent heat. Taking a density for the intruded, molten rock of 2.6 . lo3 kg/m3, its temperature drop 1000°C and its specific and latent heat to be lo3 J kg-’ ‘C-’ and 4. lo5 J/kg respectively, it can be calculated from eqn. (1) and (2) that the volume of molten rock required to maintain the geothermal systems for 1 m.y is about 3.0. lo4 km3. Given that the area of the heat source is assumed fixed at 5000 km2 its thickness must be about 6 km. In the above account no consideration is given to the heat required to produce the extensive rhyolite eruptives. The prevailing view, with regard to the genesis of the rhyolites
erupted
from the Central
Volcanic
Region,
is that they arise from a partial
melt of greywacke rocks at a shallow level (Ewart and Stipp, 1968; Reid, 1983). If so, the heat source would have to be about 9 km thick, as the heat associated with the erupted rocks from the Central Volcanic Region amounts to about one third of the heat discharged by the geothermal systems (Table 1). With the above model, the observed heat output can be accounted for if about two-thirds of the 13 km (15 km less the 2 km of superficial pyroclastic cover) thick crust consists of cooling intruded rock. But the result rests on the initial assumptions, and other more complex models to explain the large heat output may be possible. CONCLUSIONS
A major contribution Region displays many
of this study has been to show that the Central Volcanic of the characteristics commonly associated with actively
spreading oceanic back-arc basins-i.e. a thin crust, low seismic velocities crust and upper mantle, possible evidence for upper crustal magma tensional
tectonics
to extension
and a very high heat flow. The mechanical
within
the Central
extension
manifested
spreading
accompanied
Volcanic
by the stretching
Region
by the intrusion
process that gives rise
is, nevertheless,
of an old continental of a new lithosphere
within the chambers,
still contentious.
lithosphere, the principal
Is
or is active means of
extension? The case for the active spreading alternative, as argued in this paper, rests on both the high heat flow and the absence of Mesozoic greywacke of greywacke derived
sediments
in any of the deep bore holes that have been drilled
median area of the Central Volcanic Region. The critical likely come from further deep drilling associated with exploration programme. It is unlikely that extension
or spreading
within
taking place in a simple, symmetrical manner Indeed, one of the consequences of the Central
within
the
test will, however, most the current geothermal
the Central
Volcanic
about a central Volcanic Region
Region
is
spreading ridge. being a back-arc
basin that is above sea level, is that we can observe a strong asymmetry in the pattern of heat discharge. The grouping of the active geothermal fields within the eastern half of the Central Volcanic Region, and on a more detailed scale, the tendency for the individual geothermal fields to become younger to the east (Calhaem, 1973), points towards intrusion now taking place close to the line of presently active andesite volcanoes on the eastern margin of the Central Volcanic Region. Although asymmetric spreading has been claimed for some oceanic back-arc basins on the basis of magnetic anomaly lineations (Barker and Hill, 1980), it is generally regarded that back-arc spreading, like its mid-oceanic counterpart, is symmetric
(Weissel,
1981).
Asymmetric
spreading
within
the
Central
Volcanic
Region should be amenable to testing by microearthquake studies (e.g., Evison et al., 1976). If intrusion is occurring along the eastern margin of the Central Volcanic Region, and the newly emplaced lithosphere is at progressively more advanced stages of cooling to the west, then the thickness of the new lithosphere should also increase
to the west.
ACKNOWLEDGEMENTS
This paper is partly derived from a PhD thesis undertaken at the Institute of Geophysics, Victoria University of Wellington. I thus readily acknowledge the assistance of my supervisors Prof. F.F. Evison, Dr R.R. Dibble and Dr T. Hatherton. In particular, I am indebted to Prof. Evison for many useful discussions with regard to the heat flow problem for the Central Volcanic Region. 1 also thank Drs R.I. Walcott, and F.J. Davey of Geophysics Division, D.S.I.R. for their critical reviewing
of this paper.
407
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