Download Crustal structure of central and northern Iceland from analysis of

Geophys. J. Int. (2000) 143, 163±184
Crustal structure of central and northern Iceland from analysis of
teleseismic receiver functions
Fiona A. Darbyshire,1,* Keith F. Priestley,1 Robert S. White,1 Ragnar StefaÂnsson,2
Gunnar B. Gudmundsson2 and Steinunn S. JakobsdoÂttir2
1
2
Bullard Laboratories, University of Cambridge, Madingley Road, Cambridge, CB3 0EZ, UK. E-mail: [email protected]
Icelandic Meteorological Of®ce, Bustadavegur 9, 150 Reykjavõk, Iceland
Accepted 2000 May 12. Received 2000 May 5; in original form 1999 December 17
SUMMARY
We present results from a teleseismic receiver function study of central and northern
Iceland, carried out during the period 1995±1998. Data from eight broad-band seismometers installed in the SIL network operated by the Icelandic Meteorological Of®ce were
used for analysis. Receiver functions for each station were generated from events for a
wide range of backazimuths and a combination of inversion and forward modelling was
used to infer the crustal structure below each station.
The models generated show a considerable variation in the nature and thickness of the
crust across Iceland. The thinnest crust (20±21 km) is found in the northern half of
the Northern Volcanic Zone approximately 120 km north of the centre of the Iceland
mantle plume. Thicker crust (24±30 km) is found elsewhere in northern and central
Iceland and the thickest crust (37 km) is found close to the plume centre. Velocity±
depth pro®les show a distinct division of the crust into two main sections, an upper highvelocity-gradient section of thickness 2±8 km and a lower crustal section with small
or zero overall velocity gradient. The thickness of the upper crust correlates with the
tectonic structure of Iceland; the upper crust is thickest on the ¯anks of the northern and
central volcanic rift zones and thinnest close to active or extinct central volcanoes.
Below the Kra¯a central volcano in northeastern Iceland the receiver function models
show a prominent low-velocity zone at 10±15 km depth with minimum shear wave
velocities of 2.0±2.5 km sx1. We suggest that this feature results from the presence of
partially molten sills in the lower crust. Less prominent low-velocity zones found in
other regions of Iceland may arise from locally high temperatures in the crust or from
acidic intrusive bodies at depth.
A combination of the receiver function results and seismic refraction results constrains
the crustal thickness across a large part of Iceland. Melting by passive decompression of
the hot mantle below the rift zone in northern Iceland forms a crust of thickness
y20 km. In contrast, the larger crustal thickness below central Iceland probably arises
from enhanced melt production due to active upwelling in the plume core.
Key words: crustal structure, Iceland, mantle plume, receiver function.
1
INTRODUCTION
Iceland, the largest exposed landmass on the mid-ocean ridge
system, has been created by the interaction of the Mid-Atlantic
spreading centre with the Iceland mantle plume. Estimates
of the mantle temperature anomaly associated with the plume
vary from 150 uC (White & McKenzie 1995; White et al.
1995; White 1997) to 300 uC (Wolfe et al. 1997). The elevated
temperatures and active upwelling in the plume have elevated a
* Now at: Geological Survey of Canada, 1 Observatory Crescent,
Ottawa, ON, K1A OY3, Canada.
# 2000
RAS
large area of sea¯oor surrounding Iceland above normal depths
and have caused the Icelandic crust to be signi®cantly thicker
than normal oceanic crust (White 1997). The Mid-Atlantic
Ridge in the Iceland region is a slow-spreading ridge; the full
spreading rate is approximately 18 mm yrx1 (DeMets et al. 1994).
Across Iceland, the Mid-Atlantic Ridge is expressed as
three rift zones composed of central volcanoes transected by
rifts and ®ssure swarms (Fig. 1). In southern Iceland there are
two subparallel rift zones. The Western Volcanic Zone (WVZ)
extends from the Reykjanes Peninsula in southwest Iceland
to the LangjoÈkull icecap in central western Iceland. The WVZ
is separated by a transform fault system, the South Iceland
163
164
F. A. Darbyshire et al.
Figure 1. Tectonic map of Iceland showing the eight broad-band SIL stations used in receiver function analysis (triangles) and the four seismic
refraction pro®les discussed in the text. The SIL stations are identi®ed by a three-letter code referred to in the text. Stations ASB (SIL) and BORG
(IRIS) are the same and are referred to as station `ASB/BORG'. The volcanic rift zones are shown as pale shaded areas bounded by ®ne lines,
with the central volcanoes marked by circles. Glaciers are shaded white. Abbreviations used are as follows: NVZÐNorthern Volcanic Zone;
EVZÐEastern Volcanic Zone; WVZÐWestern Volcanic Zone; SISZÐSouth Iceland Seismic Zone; TFZÐTjoÈrnes Fracture Zone; RRÐReykjanes
Ridge; KRÐKolbeinsey Ridge; SNÐSnñfellsnes volcanic ¯ank zone; LjÐLangjoÈkull; VjÐVatnajoÈkull; krÐKra¯a central volcano.
Seismic Zone (SISZ), from the Eastern Volcanic Zone (EVZ),
which extends from south Iceland to below the VatnajoÈkull
icecap in central Iceland. The Northern Volcanic Zone (NVZ)
extends northwards from VatnajoÈkull to the northern coast,
and is bounded to the north by the TjoÈrnes Fracture Zone
(TFZ). The NVZ was the site of a signi®cant episode of crustal
spreading and volcanism from 1975±1984 (BjoÈrnsson et al. 1977).
The positions of the Icelandic rift zones have changed with
time (Sñmundsson 1979; Helgason 1984, 1985; Hardarson et al.
1997; Smallwood et al. 1999). Rifting in the south is presently
being transferred from the WVZ to the EVZ (JoÂhannesson
1980), and at least two major shifts of the northern rift zone
in the last 15 Myr have been identi®ed. The ridge jumps are
thought to arise from the interaction of the Mid-Atlantic Ridge
with the Iceland plume (e.g. Sñmundsson 1974, 1978). The
Mid-Atlantic Ridge drifts slowly westwards with respect to
the plume (Morgan 1981); therefore, eastward jumps of the rift
axis in Iceland are required to keep the ridge above the
plume centre. The centre of the plume is believed to lie below
the VatnajoÈkull icecap at present.
Although it has long been postulated that the plume causes
anomalously thick crust beneath Iceland compared to normal
oceanic crust, a variety of geophysical studies have led to
two contrasting models. The ®rst model has a hot, 10±15 km
thick crust overlying a partially molten upper mantle with
anomalously low seismic velocities (e.g. Gebrande et al. 1980;
Beblo & BjoÈrnsson 1978, 1980). The second model has a cooler,
20±40 km thick crust, with the thickest crust directly above the
centre of the plume (Zverev et al. 1976; Bjarnason et al. 1993;
Staples et al. 1997; Darbyshire et al. 1998; Menke et al. 1998).
In Iceland, the terms `upper crust' and `lower crust' are
used differently from elsewhere. The Icelandic crust is divided
vertically into two sections (FloÂvenz 1980). The `upper crust' is
characterized by seismic P-wave velocities of typically less than
y6.5 km sx1 (although this value may vary between different
regions of Iceland) and velocity gradients in excess of 0.2 sx1.
Typical upper crustal thicknesses determined from seismic
refraction pro®ling are y5 km for old Icelandic crust, 6±10 km
for crust on the ¯anks of the volcanic rift zones that have been
affected by ridge jumps, and 2±3 km for crust directly below
active and extinct central volcanoes (e.g. PaÂlmason 1963; FloÂvenz
1980; Bjarnason et al. 1993; Staples et al. 1997; Darbyshire
et al. 1998; Menke et al. 1998). The `lower crust' is characterized by seismic P-wave velocities of y6.5±7.4 km sx1 and
low (<0.02 sx1) velocity gradients.
In this study we divide the Icelandic crustal structure
resolved by receiver function modelling into an upper and
lower crust based purely on the changes in velocity gradient
in the velocity±depth pro®les, rather than using the absolute
P-wave velocity to classify the different parts of the crust. The
reason for this is twofold. First, receiver function analysis
is sensitive to the shear wave velocity structure. The P-wave
velocity shown in Figs 5±7 and 10 is an assumed value based
on estimates of Poisson's ratio from previous seismic studies of
the Icelandic crust (e.g. Menke et al. 1994). Second, a trade-off
exists in receiver function analysis between the depth to an
interface and the average seismic velocity above the interface. The absolute velocity modelled at any depth is of less
importance than the positions and magnitudes of velocity
contrasts in the models.
#
2000 RAS, GJI 143, 163±184
Receiver structure of central and northern Iceland
The composition and structure of the Icelandic crust provides important information about the processes taking place
in the mantle plume, and resolution of the crustal structure
of Iceland has important implications for the study of mantle
dynamics and plume±rift interaction. In this study, we use
receiver function analysis (Langston 1979; Ammon et al. 1990;
Ammon 1991) to obtain 1-D shear wave velocity models
of the crustal structure below eight three-component broadband seismographs in central and northern Iceland. We then
combine these results with published seismic refraction results
to map the crustal structure across a large part of Iceland.
2
DATA AND ANALYSIS
The broad-band seismographs used in this study were deployed
as part of the SIL network (StefaÂnsson et al. 1993) in central
and northern Iceland (Fig. 1). The stations have either a Guralp
CMG-3T (GIL, GRA, REN, SIG), a Guralp CMG-3ESP
(GRI, HVE, SKR) or a Streckheisen STS-02 (ASB/BORG)
seismometer and an RD3/OSD3 digitizer (RoÈgnvaldsson et al.
1997).
P waves from a large number of teleseismic earthquakes have
been recorded by the SIL stations but, due to the high levels of
microseismic noise in Iceland, we were only able to analyse
waveforms from 19 earthquakes (Table 1). The events cover a
wide range of azimuths, and there are several events that fall into
clusters of azimuth and epicentral distance. This is extremely
important in the identi®cation of robust features in the waveforms, particularly where the data are signi®cantly contaminated
by noise.
2.1
Method
A receiver function is the response of a seismograph site to an
incident teleseismic P wave. The use of receiver function analysis
for determining crustal structure beneath a three-component
broad-band seismograph station is now a well-established technique. P waves dominate the vertical-component seismogram
from a steeply arriving P wave, and S waves in the P-wave
coda, which arise from P-to-SV conversions at interfaces below
the site, are preferentially recorded on the radial component.
The P-wave coda also contains information about the earthquake source, source site effects and deep mantle propagation.
Deconvolution of the vertical-component from the radialcomponent seismograms removes information common to
both, leaving a time-series showing the site response or `receiver
function'. In this study we use the deconvolution technique of
Langston (1979) as modi®ed by Ammon (1991) to give the trueamplitude receiver function, since preservation of the radialto-vertical amplitude ratio constrains the near-surface velocity
at the receiver site. Since deconvolution is inherently unstable,
the process is stabilized by ®lling spectral holes using the waterlevel method (Helmberger & Wiggins 1971; Clayton & Wiggins
1976). The receiver functions are smoothed by convolving
with a Gaussian function. We use a Gaussian parameter of 1.5,
which low passes the waveforms at y0.7 Hz. Water-level parameters in the range 0.0001±0.1 are used in this study, chosen
individually for each receiver function by visual inspection
of the trade-off between the form of the `averaging function'
(the deconvolution of the vertical component from itself;
this should ideally resemble a narrow Gaussian curve) and
the stability of the radial and tangential receiver functions. The
seismograms were not ®ltered before deconvolution since some
of the frequencies of interest to the study may also be lost when
microseismic noise is removed by high-pass ®ltering.
We determined the crustal velocity structure below each
site using a combination of inversion and forward modelling.
1-D velocity models were determined using the time-domain
linearized inversion procedure of Ammon et al. (1990). Starting
models, which were based on nearby refraction models, were
Table 1. Events used in receiver function analysis. Backazimuth and epicentral distance are expressed in degrees;
depth is expressed in kilometres. Event information comes from the IRIS hypocentre data ®les. The backazimuth and
epicentral distance values are representative values compiled from the northeast Iceland stations. Events from 1994 and
1995 (second section of table) were recorded by IRIS station BORG (site ASB/BORG) only.
Event
#
Region
97339112654
98152053403
96128232000
97195160935
96357145327
98123233021
97133141345
98150062228
97130075729
96283131052
97301061517
95231214331
97245121322
98158232013
Kamchatka
Kamchatka
Kuril Islands
Kuril Islands
Japan
Taiwan
Afghanistan
Afghanistan
Iran
Cyprus
Peru±Bolivia
Colombia
Colombia
Mexico
94230044257
94289051000
95118163007
94277132225
95147130352
Kuril Islands
Kuril Islands
Kuril Islands
Kuril Islands
Sakhalin Island
2000 RAS, GJI 143, 163±184
165
Backazimuth
Epicentral distance
Magnitude
Depth
0
2
11
11
18
34
74
75
86
112
240
242
242
265
61
61
70
72
70
88
57
56
55
43
82
73
74
70
7.6
6.5
6.2
5.9
6.5
7.5
6.4
6.9
7.3
6.8
6.6
6.5
6.5
6.3
33
40
54
33
227
33
196
33
33
33
125
126
232
87
6
7
8
9
11
71
70
71
71
62
6.6
6.7
6.9
8.3
7.5
15
117
29
14
11
166
F. A. Darbyshire et al.
parametrized as a stack of thin horizontal layers to a depth
of 60 km. The S-wave velocity was the free parameter in the
inversion, the P-wave velocity was set assuming a Poisson's
ratio of 0.25 and the layer thicknesses were ®xed. The starting
models were randomly perturbed into 40 new starting models
and the radial receiver function was inverted by minimizing
the difference between the observed receiver function and
synthetics computed from the models, while simultaneously
constraining the model smoothness. This pseudo-Monte Carlo
approach reduces the dependence of the inversion convergence
on the form of the initial starting model. The inversion produced a range of solutions that ®t the observed receiver function
to different degrees. Those geologically reasonable models that
gave a good match to the data, given the noise level present and
the waveform coherence, were selected for further study.
We used forward modelling to reduce the model complexity
and to assess how well individual model features were constrained by arrivals in the data. In the forward modelling we
required that the simpli®ed earth model should contain the
general features of the inversion models but with a smaller
number of parameters. Large tangential arrivals in the receiver
functions indicate laterally heterogeneous structures (Langston
1977; Cassidy 1992); where these occurred we attempted to
match only the largest-amplitude features of the radial receiver
function so as to minimize the effect of varying lateral structure
on the 1-D earth model.
An important consideration in receiver function analysis
is that of vertical and horizontal resolution. Our receiver
functions are low-pass ®ltered at y0.7 Hz, corresponding to
a shear wavelength of 5±6 km in the crust. Various studies
(see Sheriff & Geldart 1982 for a summary) have shown that
interfaces separated by more than a quarter-wavelength of the
seismic wave are resolvable. Thus, for the frequency content
of the receiver functions presented here we do not expect to
resolve features thinner than y2 km. Horizontal resolution is
determined by the horizontal averaging of the wave and this
is generally taken (Sheriff & Geldart 1982) as the radius, R, of
the ®rst Fresnel zone associated with the incoming P wave. This
is given by
q
R~ (zzj/2)2 {z2 ,
(1)
where z is the depth of the interface and l is the wavelength
of the incident P wave. The horizontal resolution is depthdependent; for our case the horizontal resolution at the depth
of the Moho is about 25±30 km.
All the crustal models discussed in this paper are expressed
in terms of shear wave velocity but we also include P-wave
velocities in the model plots in order to make comparisons of
the receiver function crustal models with existing models from
refraction pro®les. We use a Poisson's ratio of 0.27 to calculate
the P-wave velocity; this value is reported from previous
seismic studies of Iceland (e.g. Menke et al. 1994). The data sets
of Ludwig et al. (1970) and Carlson & Herrick (1990) are used
to calculate the rock densities used in the forward models.
The receiver function models for the Icelandic seismic stations
all have low surface seismic velocities. These are constrained
using the apparent shift of the ®rst arrival from zero time. The
direct P-wave arrival (referred to as Pp) arrives at zero time
but, for a crustal structure with low surface velocities, it is
closely followed by P-to-S conversions from interfaces such as
the base of a sediment layer or the base of a layer of recent lava
¯ows. If the time separation is small, the direct P-wave arrival
and the near-surface signals are smoothed by the Gaussian
®lter into a single peak that is delayed with respect to zero time.
We refer to this peak as the `apparent Pp arrival '.
For most of the stations studied, one event in a cluster of
azimuth and distance was superior to the others in terms of the
signal-to-noise (snr) ratio. We inverted and forward modelled
the receiver function of the superior event, and used the
receiver functions from the other events in the cluster to assess
the stable features of the receiver functions. Only at station
ASB/BORG did we have a number of low-noise waveforms
for a cluster, and we were able to stack a group of receiver
functions.
2.2
Analysis
As an example of our analysis procedure we describe here
the analysis of the data from REN. Details of the analysis of the
data from the other seven stations are given in the Appendix.
Most of the REN radial receiver functions have large amplitudes compared with the pre-signal noise and the tangential
amplitudes (Fig. 2), suggesting signi®cant impedance contrasts
below the station, and there is a strong coherence in arrivals
from similar azimuths. All receiver functions show a shift of
y0.2±0.4 s from zero time of the apparent Pp arrival. A strong
negative arrival 4±5 s after the direct arrival is seen in almost all
of the radial receiver functions.
Each receiver function discussed in this and subsequent
sections is identi®ed by the event origin time in the form
yyjjjhhmmss, where y=year, j=julian day, h=hour, m=minute
and s=second.
The crustal structure below REN is derived from 10 receiver
functions (Fig. 2), of which receiver functions from events
Figure 2. Receiver functions for station REN. Radial receiver functions
are plotted on the left and tangential receiver functions are plotted
on the right. All receiver functions are plotted to the same amplitude
scale. The receiver functions are identi®ed by the event origin timeÐ
year, julian day, hour, minute, second (left). The backazimuth and
epicentral distance of the event are plotted in the centre. The water-level
parameter used in the deconvolution is plotted on the right. The
Gaussian value used in all deconvolutions was 1.5.
#
2000 RAS, GJI 143, 163±184
Receiver structure of central and northern Iceland
97130075729 and 97301061517 were inverted. These receiver
functions are the quietest waveforms from the data set,
and represent two different event azimuths. Both inversions
give similar velocity models (Fig. 3). The uppermost 2±3 km
has a strong positive velocity gradient, giving a velocity of
y3.7 km sx1 at y3 km depth. Below this, there is a pronounced broad low-velocity zone (LVZ) extending to y15 km
depth. This is the cause of the strong negative arrival at 4±5 s.
Below y15 km there is a positive velocity gradient, and mantle
velocities of y4.5 km sx1 ®rst occur at y20 km depth. Below
23 km depth, the 97301061517 model has a simple half-space
structure. Velocity variations below 25 km depth are required
to ®t the 97130075729 receiver function arrivals later than
y11 s. However, because of the large-amplitude tangential
arrivals at these times compared to the radial arrivals, we
do not consider that the sub-25 km structure shown in the
(a)
(b)
Figure 3. Modelling results for the 97130075729 (a, top and bottom
panels) and 97301061517 (b, top and bottom panels) REN receiver
functions. Top: radial and tangential (radial above tangential) receiver
function data (dotted lines), synthetic receiver functions generated from
the inversion of 40 starting models (grey solid lines) and a synthetic
receiver function from the best-®tting forward model (thick solid line).
Bottom: corresponding velocity±depth pro®les for the inversion results
(grey solid lines, expressed in terms of the shear wave velocity, VS)
and the velocity±depth pro®le of the best-®tting forward model (thick
solid line). The bounds of the model space explored by the inversion are
shown by broken lines. Event origin time, backazimuth (BAZ) and
epicentral distance (GCARC) are plotted in the top panels. Zero time
represents the arrival time of the direct P wave (Pp).
#
2000 RAS, GJI 143, 163±184
167
97130075729 inversion result is signi®cant. Instead, it is more
likely to be an artefact introduced by laterally heterogeneous
structure.
The forward modelling results are plotted as thick solid lines
in Fig. 3 and forward modelling tests for REN are shown in
Fig. 4. The delay in the apparent Pp arrival results from low
surface velocities (<1.6 km sx1). The upper crust is 2±3 km
thick, and has a strong positive velocity gradient. The lower
section of the crustal models contains a prominent LVZ with a
minimum shear wave velocity of 2 km sx1. The high velocity
contrast at 10±12 km depth is required to ®t the negative
arrival at 4±5 s in the receiver functions. The base of the LVZ is
not well constrained, and is shown as a gradual transition to
normal lower crustal velocities. The Moho is a sharp velocity
discontinuity at 20±21 km depth, at the base of a positive
velocity gradient in the lower crust, below which the halfspace velocity is y4.4 km sx1. Figs 4(c) and (f) show the ®ts
of the synthetics for these models to noisier receiver functions
from approximately the same azimuth; in general, the ®t is
satisfactory and this adds to our con®dence in the models.
Small crustal LVZs often arise from the mapping of lateral
structural variations into the 1-D model. Lateral variations are
certainly present in the vicinity of REN but, because the phase
in the receiver function controlled by the LVZ is seen strongly
at a wide range of azimuths, we consider the LVZ to be an
important feature of the model. The signi®cance of the features
modelled here is discussed in the following section.
The two forward models have sharper discontinuities
and lower minimum velocities for the mid-crustal LVZ than
the inversion models. The broad LVZ from the inversion is
characterized by velocity gradients rather than by ®rst-order
discontinuities due to the smoothing constraint. However, these
broad LVZs are insuf®cient to match the amplitudes of the
negative arrivals at 4±5 s in the two radial receiver functions,
and the remaining amplitude ®ts are taken up by additional
velocity changes in the models below 25 km. If the structure
below 25 km is taken as a half-space (as in Figs 3 and 4),
the amplitude ®t to the ®rst 10 s of the radial waveform
deteriorates signi®cantly unless the top of the LVZ is modelled
as a sharp discontinuity instead of a gradual decrease in
velocity. A sharp decrease in velocity down to y2 km sx1 in this
case is necessary to match the negative arrivals at 4±5 s in both
receiver function models, but an adequate ®t to the subsequent
positive arrival can only be achieved if the 2 km sx1 layer is
thin and the layers beneath simulate a gradational increase in
velocity. A comparison of the single crustal LVZ of the forward
model and the double LVZ of the inversion model shows some
of the trade-off inherent in receiver function analysis.
In the following discussion, we summarize the results of the
receiver function analysis and discuss their implications for our
understanding of the crustal structure of Iceland. All seismic
velocities quoted in the text are shear wave velocities, unless
otherwise stated.
3 RESULTS OF RECEIVER FUNCTION
ANALYSIS: CRUSTAL STRUCTURE OF
CENTRAL AND NORTHERN ICELAND
3.1
The Northern Volcanic Zone
Data from stations REN and GIL (Fig. 1) were used to determine the crustal structure across the northern part of the
168
F. A. Darbyshire et al.
(a)
(b)
(c)
(d)
(e)
(f)
Figure 4. Forward modelling tests for the REN 97130075729 (a±c) and 97301061517 (d±f) receiver functions. (a) Final velocity±depth model
(solid line) and test model (dashed line) with LVZ removed; (b) radial (top) and tangential (bottom) receiver function data (dotted lines), synthetic
receiver functions for the ®nal model (solid line) and for model with LVZ removed (dashed line); (c) ®t of the receiver function model
for 97130075729 (solid lines) to the 98150062228 and 97133141345 receiver functions (dotted lines); (d), (e) model and receiver function ®t for
97301061517; (f) ®ts of 97301061517 model to noisier data from 97245121322 and 98158232013.
NVZ. The crustal models for REN (Figs 5a and b) have a lowvelocity surface layer (1.1±1.5 km sx1), a thin upper crust
(2±3 km thick) and a total crustal thickness of 20±21 km. REN
lies 10 km southeast of the Kra¯a central volcano, and therefore the crust below the station is possibly affected by active
magmatic processes. Surface rocks in the Kra¯a region consist
mostly of lava ¯ows and hyaloclastites, all of which have low
seismic velocities. BrandsdoÂttir et al. (1997) noted a similar thin
upper crust in their refraction study across the Kra¯a region,
and also found a dome of high-velocity material below the
volcano. The crustal thickness is similar to the 19 km reported
for the FIRE refraction pro®le (Staples et al. 1997).
All REN radial receiver functions (Fig. 2) show a strong
negative arrival 4±5 s after the direct P arrival. This feature
results in a strong LVZ in the lower crust. Since the negative
arrival does not vary signi®cantly with azimuth in the radial
receiver functions, we believe that the crustal LVZ is a robust
feature of the crustal model, and not an artefact of scattering
from 3-D structures within the crust. The LVZ lies at 10±16 km
depth southwest of REN (Fig. 5a) and at 12±20 km depth east
of REN (Fig. 5b), and has a minimum shear wave velocity of
y2 km sx1 in both cases. Such a signi®cant velocity decrease
in the immediate vicinity of an active volcano suggests the
presence of near-solidus temperatures or melt within the lower
#
2000 RAS, GJI 143, 163±184
Receiver structure of central and northern Iceland
169
Figure 5. Final crustal models for REN and GIL. The models are expressed in terms of shear wave velocity, but the P-wave velocity (using a
Poisson's ratio of 0.27) is also labelled above each model. The labels on each panel denote station code, event approach direction and event origin
time. The approximate position of the Moho is marked with an arrow (solid stem for models shown as solid lines, broken stem for models shown as
broken lines). (a) Southwest of REN; (b) east of REN; (c) southeast of GIL; (d) east of GIL.
crust. The velocity decrease from partial melt is governed
strongly by the aspect ratio of the melt inclusions. For example,
if the melt occurs in disc-shaped cracks, less than 1 per cent
melt is required to explain the observed velocity decrease, but
if the melt inclusions are spherical, as much as 40 per cent
melt is required to explain the velocity decrease (Le Ravalec
& GueÂgen 1996). Hence, the information we have from the
velocity model is not suf®cient to quantify the amount of melt
present. The lateral extent of the LVZ cannot be resolved but,
from a consideration of the angles of approach of the rays used
in receiver function analysis, the LVZ is at least 15±20 km
across if it is a single feature. That melt exists in the crust below
Kra¯a over a large range of depths, as suggested by the crustal
models, has support from geochemistry. Pressure estimates
from variations in clinopyroxene compositions in rocks from
northern Iceland show that crystallization has taken place at a
range of depths in the crust and uppermost mantle (Maclennan,
personal communication, 1999), implying at least the past
existence of deep crustal magma chambers.
The crustal structure in the northeast part of the NVZ is
constrained by receiver functions from station GIL (Fig. A1).
Figs 5(c) and (d) show the crustal velocity models derived
from the GIL receiver functions. The upper crust is 5±8 km
thick, thicker than average for Iceland (FloÂvenz 1980), and
has a strong positive velocity gradient. The mid- and lower
crust exhibit almost uniform velocities. Some velocity steps are
modelled but the details are not well constrained by the data.
The Moho is a sharp positive step with a high velocity contrast
at 25±30 km depth.
The surface rocks in the vicinity of GIL are interglacial and
postglacial extrusives, subglacial pillows and hyaloclastites, but
#
2000 RAS, GJI 143, 163±184
older rocks (Plio-Pleistocene and Tertiary) crop out 15 km to
the east of GIL. The station lies near the northern limit of a
volcanic ridge, which is thought to be an aborted rift system
(Helgason 1989). The thickened upper crust east of GIL is
related to the tectonic history of the region. At 6±7 Ma, the
northern rift zone jumped to its present position in northeast
Iceland, and the old extrusive layers making up the upper crust
were buried by new lava ¯ows. Since the old upper crust is cool,
it can be buried deeply without the metamorphism that would
result in a signi®cant increase in seismic velocity; hence, the
upper crust increases in thickness on the ¯ank of the rift zone.
The Moho is sampled 10±30 km east of GIL, below Tertiary
surface geology. The models show the crust thickening from
north to south in the region, consistent with reported results for
eastern Iceland from Staples et al. (1997).
West of GIL, the structure is uncertain. The radial receiver
functions have low amplitudes and the tangential receiver
functions have high amplitudes. These observations suggest
that the eastern edge of the NVZ might have a low velocity
contrast at the Moho and a strongly 3-D structure.
3.2
Northern Iceland
The crustal structure of northern Iceland is constrained by data
from stations GRA, GRI and SIG (Fig. 1). Crustal structure
close to the western ¯ank of the NVZ is constrained by receiver
functions from GRA, and the results of the analysis (Fig. A2)
for these data are shown in Figs 6(a)±(c). To the east and
southwest of GRA, the upper crust is y5 km thick, which
is average for Iceland (FloÂvenz 1980), but to the north of
GRA the upper crust is somewhat thinner (y3.5 km). To the
170
F. A. Darbyshire et al.
southwest, there is a velocity discontinuity at the base of the
upper crust (Fig. 6a). The crustal models to the north and
southwest of GRA both have LVZs at approximately 15 km
depth (Figs 6a and b). To the east of GRA no LVZ is required
by the data, and here the velocity increases in a set of positive
steps (Fig. 6c). The Moho occurs at 20±22 km depth, at the
base of a positive velocity gradient.
The azimuthal differences in the crustal structure around
GRA can be interpreted in terms of the local tectonic structure.
To the west of GRA there is a faulted region, west of
which Plio-Pleistocene and Tertiary rocks (age <7.5 Myr) lie
unconformably on older Tertiary rocks (age >9.5 Myr). The
unconformity crops out y10 km west of GRA, and is thought
to be angular in nature (Jancin et al. 1985). Two extinct volcanic
complexes lie in the region y5±20 km to the northwest of
the station (Jancin et al. 1985). Thinning of the upper crust
below active and extinct central volcanoes in Iceland has been
reported by several authors (e.g. PaÂlmason 1963; FloÂvenz
1980), and `dome' structures of high-velocity material in the
crust of this region have been reported by Menke et al. (1998).
The velocity discontinuity at the base of the upper crust southwest of GRA may represent the unconformable transition
from Plio-Pleistocene to Tertiary age rocks at depth, although
the details of this transition are not known from geological
information.
The best ®t to the north and southwest GRA receiver
functions (Fig. A2) is achieved by the inclusion of an LVZ at
y15 km depth. However, the LVZ is small compared to that
at REN discussed above and its removal from the GRA model
results in minor degradation of the receiver function match.
Since GRA lies approximately 30 km from the rift axis, and
the crust is at least 5 Myr old, it is unlikely that partial melt
due to rift axis volcanism causes the LVZ. A more plausible
explanation for an LVZ in this region is a change in crustal
composition at depth. Acidic extrusives have been mapped on
the peninsula west of GRA, and acidic intrusive bodies may
occur at depth in the crust. These are likely to have a lower
seismic velocity than the surrounding basaltic rocks (Christensen
& Mooney 1995).
Crustal structure in the TjoÈrnes Fracture Zone is constrained by receiver functions from GRI (Fig. A3). The radial
receiver functions for GRI show large-amplitude arrivals, and
both the degree of variation with event azimuth and the large
amplitudes of the tangential waveforms suggest signi®cant 3-D
structure. Two possible crustal models for GRI were derived
from the receiver function analysis. The ®rst model has a
5 km thick upper crust with a strong positive velocity gradient
(Fig. 6d, solid line). Lower crustal seismic velocities are high
(y4.1 km sx1) and the Moho is a sharp positive discontinuity
at 16 km depth. There is an LVZ at 22±28 km depth. The
second model (Fig. 6d, broken line) has a 4 km thick upper
crust and two LVZs in the mid-crust. The Moho lies at the base
of a positive velocity gradient at 25 km depth.
No a priori information about crustal thickness around GRI
exists. We favour the model with a 16 km thick crust for two
reasons. First, the size of the gravity anomaly in the region
suggests thinner crust than on the mainland, if we assume local
isostatic compensation beneath Iceland. Second, evidence from
oceanic fracture zones at slow-spreading ridges (e.g. White et al.
1984; Detrick et al. 1993) suggests that the crust within fracture
zones is signi®cantly thinner than that along the main ridge
segments.
The crustal structure of north-central Iceland is constrained
using receiver functions from SIG (Fig. A4). Two crustal
Figure 6. Final crustal models for GRA, GRI and SIG. Refer to Fig. 5 for plotting conventions. (a) Southwest of GRA; (b) north of GRA;
(c) east of GRA; (d) east of GRI; (e) east of SIG.
#
2000 RAS, GJI 143, 163±184
Receiver structure of central and northern Iceland
models are obtained for the structure east of SIG (Fig. 6e).
No information could be obtained for other azimuths, since
the amplitudes of the tangential receiver functions are larger
than those of the radial receiver functions. Both models
have a 2±3 km thick upper crust with higher surface velocities
(2.2±2.3 km sx1) than those found in the neovolcanic zones.
The surface layers around SIG are made up of well-compacted
rocks of Tertiary age with high seismic velocities. The ®rst
model has an LVZ at 11±16 km depth and a positive velocity
gradient between 22 and 30 km depth. The Moho is a sharp
positive discontinuity at 30 km depth. The second model has
lower average seismic velocities (y3.2 km sx1 at 2±20 km
depth and y4.1 km sx1 at 26±50 km depth), a positive velocity
gradient between 20 and 26 km depth and a strong positive
velocity step at 50 km depth. The Moho lies at 26 km depth.
The two forward models for SIG give a crustal thickness
of 25±30 km. The shear wave velocity below the Moho in the
thinner crust model is low compared to mantle S-wave velocities
inferred from P-wave velocities measured by recent refraction
pro®les. The thicker crust model corresponds better to existing
refraction pro®le models of the Icelandic crust, if the assumptions made about the average value of Poisson's ratio in Iceland
are correct. The LVZ in the thicker crust model is not a robust
feature, since it is modelled by only a single event receiver
function. It may be an artefact of scattering from dipping
structures in the crust below SIG.
There is strong evidence for 3-D structure below both
GRI and SIG; this may be attributed to the proximity of the
TjoÈrnes Fracture Zone. The TFZ is an active transform zone
composed of three major active fault systems, striking approximately northwest±southeast (RoÈgnvaldsson et al. 1998). SIG lies
between two of the fault lineaments in the southwest of the TFZ,
171
and GRI lies within the northeastern lineament. Southwest of
GRI there is a large graben (FloÂvenz & Gunnarsson 1991) that
gives rise to a prominent gravity anomaly. This extreme 3-D
structure may explain why it was not possible to obtain a 1-D
model for the structure southwest of GRI.
3.3
Central Iceland
Data from stations ASB/BORG, HVE and SKR (Fig. 1) were
used to determine the crustal structure of central Iceland. The
crustal structure in west-central Iceland is constrained using
receiver functions from station ASB/BORG (Fig. A5), and the
crustal models from the receiver function analysis are shown
in Figs 7(a) and (b). North of ASB/BORG, the upper crust is
4±5 km thick, with a strong positive velocity gradient (Fig. 7a).
Seismic velocity increases with a smaller positive velocity gradient
from 4±8 km depth. A small decrease in velocity is modelled
between 12 and 17 km depth. The Moho is a sharp velocity
increase (to y4.5 km sx1) at y25 km depth. Southwest of
ASB/BORG, the upper crust is 4±5 km thick (Fig. 7b). An
LVZ occurs at 10±12 km depth and the Moho lies within a
strong positive velocity gradient between 24 and 28 km depth.
The difference in the crustal structure to the north and
southwest of ASB/BORG probably results from the station
location on the eastern edge of the Snñfellsnes volcanic ¯ank
zone (Fig. 1). The surface layers in this area consist predominantly of Tertiary lava ¯ows, but postglacial basaltic and
acidic rocks are found to the west and southwest of ASB/
BORG. The LVZ in the crust southwest of ASB/BORG is not a
well-resolved feature but it may indicate the presence of acidic
intrusive material at these depths.
Figure 7. Final crustal models for ASB/BORG, HVE and SKR. Refer to Fig. 5 for plotting conventions. (a) North of ASB/BORG; (b) southwest
of ASB/BORG; (c) southwest of HVE; (d) north of HVE; (e) east of SKR.
#
2000 RAS, GJI 143, 163±184
172
F. A. Darbyshire et al.
The crustal structure at the northern end of the WVZ in
central Iceland (Fig. 1) is determined from receiver functions
from station HVE (Fig. A7). The crustal models from the
receiver function analysis are shown in Figs 7(c) and (d).
Surface shear wave velocities are less than 1.5 km sx1, and the
upper crust is 5 km thick to the southwest of HVE (Fig. 7c) and
3.5 km thick to the north of HVE (Fig. 7d). Southwest of HVE
there are two LVZs between 7 and 18 km depth. The Moho is
a sharp velocity discontinuity at 32 km depth, below a 14 km
thick section with a positive velocity gradient. North of HVE,
the velocity increases in a set of positive velocity discontinuities.
The Moho lies at 29 km depth.
The surface layers around HVE are made up of Upper
Pleistocene hyaloclastites and postglacial lava ¯ows, all of
which have low seismic velocities, consistent with the receiver
function models. In addition, HVE lies close to a major geothermal ®eld in which the unconsolidated sediments exhibit low
velocities. The LVZs in the crust to the southwest of HVE are
not well constrained since they ®t waveforms in a narrow range
of azimuths, but they may arise from small quantities of melt or
acidic intrusive bodies associated with the nearby LangjoÈkull
central volcano.
The crustal structure close to the centre of the Iceland plume
is constrained by receiver functions from SKR (Fig. A8). East
of SKR, two possible 1-D crustal models are obtained from the
receiver function analysis, but at other azimuths the amplitudes
of the tangential receiver functions are higher than those of
the radial receiver functions, indicating pronounced lateral
variation in the crustal structure. The crustal models to the east
of SKR are shown in Fig. 7(e). For model 1 (Fig. 7e, solid line),
the upper crust is 5 km thick, which is average for Iceland
(FloÂvenz 1980), and has a strong positive velocity gradient.
There is a positive velocity discontinuity at 19 km depth. It is
not clear what this feature of the model represents. The Moho
lies at 37 km depth, at the base of a 9 km thick section with a
positive velocity gradient. Model 2 (Fig. 7e, broken line) has a
2 km thick upper crust and a transition to mantle velocities at
y22 km depth. Velocities below 7 km depth are greater than
4 km sx1. We consider that Model 1 is the more likely solution,
since the crustal velocities and the depth to the Moho are more
in agreement with the results from refraction pro®les.
4
DISCUSSION
Results from our receiver function analysis indicate that the
crust of central and northern Iceland varies from y20 km
thick in the northern rift zones to y40 km thick in central
Iceland above the plume centre. The upper crustal thickness
varies from 2±3 km close to central volcanoes to 7±8 km on the
edges of the rift zones. Several of the velocity±depth pro®les
show LVZs within the crust. Most of these LVZs are not large
and the shear wave velocity within them decreases by less
than 0.5 km sx1. However, below the Kra¯a central volcano, a
prominent LVZ exists in which shear wave velocities decrease
to 2±2.5 km sx1 (a velocity anomaly of >1.5 km sx1).
The seismic velocity in the uppermost mantle is not well
constrained. In most cases, the mantle has been modelled as
a half-space with a shear wave velocity of 4.4±4.6 km sx1.
The receiver function method is not sensitive to the absolute
velocity of the half-space, and therefore these velocities should
not be considered a reliable measurement of the upper mantle
velocity below Iceland.
4.1
Low-velocity zones in the Icelandic crust
Crustal LVZs may arise from compositional changes, high
temperatures, ¯uids under high pore pressure, or small amounts
of partial melt. The LVZs modelled from Icelandic receiver
functions from a narrow range of event azimuths (e.g. SIG-E,
Fig. 6e) may result from scattered energy from 3-D structures
within the crust. However, we believe the LVZs at HVE and
GRA to be robust features since their effects appear in the
receiver functions from a wider range of azimuths. At REN,
the negative arrival at 4±5 s to which we attribute a prominent
mid-crustal LVZ is seen in every radial receiver function.
The crustal LVZs beneath HVE and GRA may result from
compositional changes in the lower crust such as acidic
intrusive bodies. The acidic intrusive rocks mapped in Iceland
(JoÂhannesson & Sñmundsson 1989) consist mainly of granites
and granophyres (Soesoo 1998). The difference in seismic
velocities between granites and gabbros (Christensen & Mooney
1995) is consistent with the magnitudes of the LVZs modelled
at GRA and HVE. Alternatively, localized increases in the
concentration of plagioclase in lower crustal rocks may also
reduce seismic velocities.
The LVZ below REN is likely to be associated with the nearby
Kra¯a central volcano. Shallow magma chambers associated
with Kra¯a and other Icelandic central volcanoes have been
detected previously using results from geodetic measurements
 rnadoÂttir
(e.g. Tryggvasson 1986; Rymer & Tryggvasson 1993; A
et al. 1998) and seismic undershooting (e.g. Gudmundsson
et al. 1994; BrandsdoÂttir et al. 1997). Geodetic measurements
(e.g. Tryggvasson 1986) also suggest that deeper magma
reservoirs exist below Kra¯a but do not constrain their depths.
From geochemical analysis of clinopyroxene crystals found in
the Kra¯a basalts, Maclennan (personal communication, 1999)
suggested that melt crystallizes at two or more different depths
below Kra¯a. This appears to be similar to the generation of
the oceanic crust observed in the Oman ophiolite. Kelemen
et al. (1997) studied the geochemistry of the layered gabbros
in the Oman ophiolite and suggested that these gabbros were
formed by the emplacement of sills in the lower crust, which
then cooled and crystallized in situ. The LVZ below Kra¯a may
therefore be a section of crust that contains partially molten
sills (Fig. 8). Alternatively, the LVZ could represent a region
of distributed partial melt or material close to its solidus
temperature.
4.1.1 Tests on the Kra¯a LVZ: is the new result compatible
with seismic refraction models?
Results from the FIRE refraction pro®le (BrandsdoÂttir et al.
1997; Staples et al. 1997) give a crustal model for the Kra¯a
region that has a magma chamber at 3 km depth but no LVZ in
the mid-lower crust, in contrast to the prominent LVZ inferred
from the REN receiver function data in the same area. While
an LVZ cannot be resolved directly by seismic refraction
pro®les, its presence may be inferred from shadow zones in the
data. Shadow zones are reported by BrandsdoÂttir et al. (1997)
at offsets corresponding to the shallow magma chamber but
no shadow zones are obvious in the data at longer offsets. The
data at large offsets are sparse, however, so shadow zones
might be dif®cult to identify. In addition, the FIRE traveltime
models of Staples et al. (1997) show no crustal diving rays
#
2000 RAS, GJI 143, 163±184
Receiver structure of central and northern Iceland
173
Figure 8. Left: an interpretation of crustal accretion processes for the Oman ophiolite proposed by Kelemen et al. (1997). The crust is formed in a
shallow magma chamber and also at depth by the emplacement and crystallization of basaltic material in sills. Right: a similar interpretation
proposed to explain the LVZ modelled at Kra¯a by receiver function analysis. The LVZ is caused by a high concentration of sills (shown as dark
grey ellipses). Note that the two models are not to the same scale. The numbers shown in the Kra¯a model are seismic P-wave velocity contours
taken from Staples et al. (1997).
turning in the mid-lower crust below Kra¯a. It is important to
assess the effect of an LVZ in the mid-lower crust (i.e. at
10±15 km depth) below Kra¯a on the FIRE refraction data
in order to ensure that no direct contradictions exist between
the FIRE and receiver function crustal models. We used two
different synthetic seismogram modelling techniques to examine
the effects of introducing an LVZ into the FIRE model, based
®rst on full-waveform re¯ectivity modelling and second on ray
theory.
Fig. 9(a) shows a section-normalized P-wave common-shot
gather from an explosive shot ®red in the Kra¯a caldera.
Arrivals from crustal diving rays and Moho re¯ections are seen
in the section. The re¯ectivity method (Sandmeier & Wenzel
1986; Fuchs & MuÈller 1971) was used to create synthetic shot
gathers for a 1-D crustal model of the Kra¯a region (Figs 9b
and c). Ray tracing through the 2-D FIRE crustal model of
Staples et al. (1997) using the method of Zelt & Ellis (1988) was
used to provide a more realistic test of the effect of the LVZ on
the FIRE refraction data. The FIRE model, with a crustal LVZ
at 10 km depth added, is shown in Fig. 9(d). Fig 9(e) shows a
synthetic shot gather for the Kra¯a shot (using the ray tracing
method) with no LVZ, and Fig. 9(f) shows the synthetic data
for the model with an LVZ. The presence of the LVZ has little
effect on the crustal diving ray arrivals or the Moho re¯ections.
Strong re¯ections from the top and bottom of the LVZ are
observed in the synthetic section at offsets of <50 km (these
are also seen in the re¯ectivity synthetics). The Kra¯a shot data
(Fig. 9a) show considerable ringing in the seismograms at these
offsets, and it is clear that any re¯ections from an LVZ would
be obscured. Similar tests were carried out for the other FIRE
shot gathers (Darbyshire 1999). We conclude that comparison
of the synthetic seismogram tests with the FIRE data shows
that the ®ts of the synthetic shot gathers to the amplitudes of
the observed P-wave data are equally good whether or not an
LVZ at 10±15 km depth is present.
BrandsdoÂttir et al. (1997) reported clear S phases and some
indication of SmS phases in the FIRE refraction data for all the
explosive shots, thus ruling out the presence of large quantities
of interconnected melt in the mid- and lower crust below
Kra¯a. The receiver function results also show that S waves are
propagating throughout the crust, although their velocity is
#
2000 RAS, GJI 143, 163±184
reduced by up to 1.5 km sx1 in the mid-lower crust. If there
is melt at 10±15 km depth below Kra¯a, the propagation of
S waves shows that the amounts must be small, and that the
melt cannot be interconnected.
The ratio of P- to S-wave velocity in northeast Iceland is
also reported by BrandsdoÂttir et al. (1997). For most of the
FIRE shots, VP /VS lies in the range 1.76±1.79, suggesting
crustal temperatures well below the basalt and gabbro solidus.
However, SmS traveltimes for a path along the NVZ north of
Kra¯a give a higher value of VP /VS of 1.88, which is consistent
with high temperatures in the lower crust.
The tests performed on the FIRE crustal model suggest
that the inclusion of an LVZ in the mid-lower crust below
Kra¯a is permitted by the resolution of the FIRE refraction
data and the result of the receiver function analysis presents no
contradiction to the refraction results.
4.2 Comparison of results from receiver function and
refraction seismic models
Fig. 10 shows 1-D P-wave velocity±depth pro®les taken from
refraction pro®les and receiver function models for similar
regions of Iceland. The P-wave velocities labelled for the
receiver function models are based on an assumed Poisson's
ratio of 0.27. Although the comparison of these models is
limited due to the different nature of the techniques used, the
models show a high degree of similarity in the large-scale
features of crustal structure. Comparisons of crustal thickness at SIL stations GRA and REN with the results from
the corresponding B96 and FIRE refraction pro®les are of the
greatest interest due to the close proximity of the SIL stations
to the locations of the refraction pro®les. In both cases the
receiver function and refraction results are in excellent agreement. LVZs in the middle and lower crust are modelled using
receiver function analysis but are not observed in the refraction
pro®le models since the subtle effects of LVZs in refraction
data are easily masked by noise. In addition, changes in seismic
wave speeds due to high temperatures or partial melt may result
in a strong shear wave velocity anomaly that will affect the
receiver function results, but a less prominent P-wave velocity
change that may not be clear in refraction data.
174
F. A. Darbyshire et al.
Figure 9. Kra¯a LVZ tests on FIRE (Staples et al. 1997) refraction data. (a) FIRE refraction pro®le record section for a shot at Kra¯a. The
®rst arrivals to y100 km are crustal diving rays and the Moho re¯ection, PmP, is seen clearly at offsets >70 km and y3 s reduced traveltime.
(b) Re¯ectivity synthetic shot gather for 1-D structure at Kra¯a, with no LVZ. (c) As (b) but with an LVZ at 10±15 km depth in the crust. (d) Crustal
model of the Northern Volcanic Zone from Staples et al. (1997). An LVZ is inserted below Kra¯a (distance 0 km along model). (e) Synthetic shot gather
from 2-D ray tracing through the FIRE model with no LVZ. (f) Synthetic shot gather from 2-D ray tracing through the FIRE model with LVZ inserted.
The comparisons of crustal thickness at stations HVE and
SKR with pro®les derived from the ICEMELT crustal model
are not expected to be as good as those in northeast Iceland,
since the stations are not as close to the refraction pro®le, and
the crustal structure is known to be strongly 3-D. However,
the Moho depths modelled by receiver function analysis are
within 5 km of those modelled from the ICEMELT refraction
pro®le.
#
2000 RAS, GJI 143, 163±184
Receiver structure of central and northern Iceland
175
Figure 10. Tectonic map of Iceland showing crustal thickness estimates from this study (larger text) and previous seismic studies (smaller text).
Triangles: receiver function analysis (this study); diamonds: ICEMELT refraction pro®le (Darbyshire et al. 1998); circles: FIRE refraction pro®le
(Staples et al. 1997); stars: B96 refraction pro®le (Menke et al. 1998); square: reinterpretation of RRISP refraction pro®le (Menke et al. 1996, 1998);
inverted triangles: SIST refraction pro®le (Bjarnason et al. 1993). Also shown are representative 1-D velocity±depth models for the receiver
function results and for the refraction pro®les in the corresponding regions. The models are expressed in terms of P-wave velocity; for the receiver
function models, this value was calculated using a Poisson's ratio of 0.27.
4.3 Moho depth variation across central and northern
Iceland
We now combine the results from published wide-angle seismic
pro®les with the receiver function results presented in this
paper to map out the crustal thickness variations across much
of central and northern Iceland (Fig. 10). On the Icelandic
mainland a large variation in crustal thickness is reported,
ranging from a minimum of 19 km below the Kra¯a central
volcano in northeast Iceland to a maximum of 40 km below
northwest VatnajoÈkull (the location of the current plume centre,
according to geophysical observations). The crustal thickness
variations show several trends. An increase in Moho depth
along the NVZ and its western ¯ank from north to south (i.e.
towards the plume centre) is observed. Along the ICEMELT
refraction pro®le, the crust thickens towards the plume centre,
and a similar trend is observed in western central Iceland. A
decrease in crustal thickness along the FIRE refraction pro®le
#
2000 RAS, GJI 143, 163±184
from the eastern fjords to the Kra¯a central volcano and an
increase in crustal thickness from Kra¯a to the western ¯ank of
the NVZ is also observed.
A clear increase in Moho depth towards the plume centre
is shown. Similar observations have been noted in crustal
structure studies above other major mantle plumes such as the
Hawaiian plume (Zucca et al. 1982; Watts & ten Brink 1989),
the Azores hotspot (Detrick et al. 1995) and the GalaÂpagos
hotspot (Ito & Lin 1995). On the small scales considered here
(up to 250 km from the plume centre), some temperature increase
towards the plume core may be a contributing factor to the
increase in crustal thickness but the dominant mechanism for
enhanced melt production is most likely to be active upwelling
associated with the central rising core of the plume.
At Kra¯a, y120 km north of the plume centre, crustal
thickness estimates from rare earth element inversions of the
Kra¯a lavas (Nicholson & Latin 1992) are consistent with estimates from geophysical studies. Since the rare earth element
176
F. A. Darbyshire et al.
inversions assume passive upwelling of the mantle, the close
agreement with the seismic crustal thickness suggests that
the dominant mechanism of crustal generation at Kra¯a is
passive decompression melting of the mantle in response to
the separation of the plates at the spreading axis (White et al.
1995). The 20±22 km crustal thickness suggests a mantle
potential temperature of 1480±1500 uC (Bown & White 1994;
White 1997).
Close to the plume centre, the crustal thickness is 40 km.
If this crust was created from melting due to passive mantle
upwelling alone, the required mantle potential temperature
would be y1650 uC. This temperature is greater than those
estimated from other geophysical and geochemical observations.
It is much more likely that the mantle temperature below
northwest VatnajoÈkull is not signi®cantly different from that
below Kra¯a, and that the change in crustal thickness can
be explained by taking active upwelling into account. Active
upwelling in the plume core increases the volume of mantle
material passing through the melt zone, hence melt production
is enhanced. Recent plume models proposed by Ito et al. (1999)
suggest that active upwelling is restricted to a radius of less than
100 km around the plume centre; the geophysical observations
from central and northern Iceland appear to support this
hypothesis.
5
CONCLUSIONS
The azimuthal dependence of receiver function waveforms for
the eight SIL seismic stations studied indicates that the structure
of the Icelandic crust is highly 3-D in nature. However, 1-D
approximations to the crustal structure below each station may
be modelled successfully using the receiver function technique
provided that simple velocity models are sought for the main
crustal features.
The crustal thicknesses from receiver function analysis
are broadly consistent with the crustal structure from recent
refraction pro®les, showing a crust of at least 20 km thickness across the Icelandic mainland. The thickest crust (up to
y37 km) is found close to the plume centre in central Iceland,
and in regions of old Tertiary surface geology. Thinner crust
is found within the northern and northwestern parts of the
Northern Volcanic Zone, and off the northern coast in the
TjoÈrnes Fracture Zone.
Low-velocity zones are found within the Icelandic crust.
In some cases these features may result from compositional
changes at depth such as acidic intrusions. However, the
prominent shear wave LVZ below the Kra¯a central volcano
may be explained by the presence of partially molten material
in the form of a high concentration of basaltic melt sills.
ACKNOWLEDGMENTS
We would like to acknowledge numerous fruitful discussions
with Rob Staples, John Maclennan, Bryndõs BrandsdoÂttir, Bill
Menke, Paul Saunders and Steve Mangino. The manuscript was
 lafur Gudmundsson and
improved by valuable reviews from O
Takuo Shibutani. This research was supported by NERC grant
GR3/9410. FD was supported by NERC and Schlumberger
Cambridge Research. The broad-band seismometers at stations
GRI, HVE and SKR were provided by the PASSCAL geophysical equipment pool. Figs 1 and 10 were created using the
GMT software of Wessel & Smith (1991). Department of Earth
Sciences, Cambridge, contribution number 5976.
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F. A. Darbyshire et al.
APPENDIX A: RECEIVER FUNCTION
ANALYSIS IN CENTRAL AND
NORTHERN ICELAND
A1
Station GIL
The receiver functions recorded by GIL show low-amplitude
radial signals for events approaching from the west, but stronger
signals for events approaching from the east. The tangential
receiver functions have signi®cant amplitude arrivals suggesting
that the structure below the station is laterally heterogeneous.
Two receiver functions were chosen for inversion from 10
events recorded by GIL: 96283131052 and 98150062228. These
are the quietest of the receiver functions for which we believe a
1-D analysis is valid. The results of the inversions are shown in
Fig. A1 (grey lines). Each inversion yielded two distinct model
families, both of which ®t the data adequately. The inversion
models for 96283131052 show a similar crustal structure, but
family 2 has consistently lower seismic velocities and a shallower
Moho than family 1. For 98150062228 the ®rst family of models
shows higher velocities and a deeper Moho than family 2.
The models have an upper crustal section of thickness 5±8 km
with strong positive velocity gradients, and the Moho lies at
25±30 km depth.
Forward modelling results for GIL are shown in Fig. A1 as
thick solid lines. 96283131052 model 1 gives a crustal thickness
(a)
(b)
of 30 km with a sharp Moho. Above 7 km depth the velocity
gradients are high but the lower section of the model shows
no overall velocity increase. The surface velocity required to ®t
the time of the apparent Pp arrival is 1.4 km sx1. Variations
in the velocity structure between depths of 12 and 21 km is
required to ®t the features of the waveform between 4 and 10 s.
Model 2 shows consistently lower velocities than model 1. The
surface velocity is 1.1 km sx1 and the upper crustal section is
6 km thick. This model also requires a velocity variation in the
middle section of the crust to ®t the receiver function between 4
and 10 s. The Moho is shallower than for model 1 (28 km).
98150062228 model 1 is simple with surface velocities of
1.2 km sx1, an upper crust of thickness 5 km and a Moho
depth of 25 km. A LVZ beneath a `lid' structure at the base of
the upper crust is required to produce a peak at 6±8 s in the
waveform. Model 2 shows similar general features to model 1
but the structure is a set of transitional velocity changes rather
than sharp discontinuities, with a velocity inversion below
9 km depth. More complex crustal models than those shown
are required to ®t the amplitudes of the arrivals in the
98150062228 radial receiver function. However, the waveform
shows signi®cant side-lobes around the apparent Pp arrival.
The averaging function (the deconvolution of the verticalcomponent seismogram from itself) shows that these side-lobes
result from the band-limited nature of the signal (Darbyshire
1999). Since the side-lobes distort the amplitudes of the converted
(c)
(d)
Figure A1. Modelling results for the 96283131052 (a,b; top and bottom panels) and 98150062228 (c,d; top and bottom panels) GIL receiver
functions. Top: radial and tangential (radial above tangential) receiver function data (dotted lines), synthetic receiver functions generated from
the inversion of 40 starting models (grey solid lines) and a synthetic receiver function from the best-®tting forward model (thick solid line).
Bottom: corresponding velocity±depth pro®les for the inversion results (grey solid lines, expressed in terms of the shear wave velocity, VS) and the
velocity±depth pro®le of the best-®tting forward model (thick solid line). The bounds of the model space explored by the inversion are shown as
broken lines. Event origin time, backazimuth (BAZ) and epicentral distance (GCARC) are plotted in the top panels. Zero time represents the
arrival time of the direct P wave (Pp).
#
2000 RAS, GJI 143, 163±184
Receiver structure of central and northern Iceland
and reverberative phases close to the apparent Pp arrival, it
is inappropriate to attempt to ®t them with a more complex
model than those shown in Fig. A1.
The crust beneath GIL is 25±30 km thick with a 5±8 km
thick upper crust.
A2
Station GRA
The radial receiver functions at GRA have large amplitudes
compared to the tangential amplitudes and pre-signal noise.
Waveforms clustered in event azimuth and distance show strong
similarities and some arrivals are common to all azimuths. The
differences in amplitude of these arrivals indicate a difference in
the crustal structure to the east and to the west of GRA.
Three events, 97301061517, 98150062228 and 98152053403,
were chosen for inversion. The receiver functions cluster into
groups according to event azimuth and distance and the chosen
events were the quietest in each cluster. The inversion results
are shown in Fig. A2 (grey lines). The model for 97301061517
has a strong positive velocity gradient in the uppermost 5 km,
an LVZ at y15 km depth and a positive velocity gradient
between 17 and 24 km depth. Mantle velocities occur below
y24 km depth. The model for 98152053403 has a crustal
thickness of y25 km and a strong positive velocity gradient
in the uppermost 5 km. The shear wave velocity reaches high
velocities (4.5 km sx1) at 10±14 km depth but subsequent
forward modelling allows us to reduce this velocity without
degrading the waveform ®t signi®cantly. The model for
98150062228 has a strong positive velocity gradient in the
near-surface layer and positive steps in velocity between 4 and
(a)
(b)
179
22 km depth. The high-velocity `lid' structure at y25 km depth
is required to ®t the arrivals between 12 and 18 s after the
apparent Pp arrival.
The forward modelling results for GRA are shown in
Fig. A2 (thick solid lines). The 97301061517 model has a 5 km
thick upper crust with a surface velocity of 1.8 km sx1 and a
strong positive velocity gradient. A step in velocity at 4.5 km
depth is required to ®t the negative arrival in the waveform
at 4 s and an LVZ at 13±16 km depth is required to ®t the
negative arrival at 7 s. The Moho lies at 22 km depth at the base
of a section with a positive velocity gradient. The model also
provides a good ®t to the main features of the noisier receiver
functions from the southwest azimuth cluster (95231214331,
97245121322 and 98158232013) shown in Fig. A2(d) (top). The
receiver function for 98152053403 shows high levels of noise
and a long-wavelength signal. A simple structure was therefore
used to model the main features of the waveform without
striving for a close amplitude ®t. The uppermost 3.5 km has a
strong positive velocity gradient and the waveform is ®tted by
an LVZ at 14±18 km depth and a sharp Moho at 20 km depth.
Although the other receiver functions from the northern event
azimuth (96128232000 and 98123233021) are noisy, their main
features are reasonably matched by the 98152053403 model
(Fig. A2d, middle). The 98150062228 waveform is oscillatory
and clear identi®cation of the phases was dif®cult. We used a
simple model composed of positive steps in velocity to ®t the
times of the arrivals in the ®rst 12 s. The crustal thickness
is 20 km. The model provides a reasonable ®t to the noisier
receiver functions from a similar event azimuth (97130075729
and 97133141345) shown in Fig. A2(d) (bottom).
(c)
(d)
Figure A2. Modelling results for the 97301061517 (a), 98152053403 (b) and 98150062228 (c) GRA receiver functions. Refer to Fig. A1 for
plotting conventions. (d) Shows the ®ts of the models to noisier receiver functions from similar backazimuths (top panel: SW; middle panel: N;
bottom panel: E).
#
2000 RAS, GJI 143, 163±184
180
F. A. Darbyshire et al.
In summary, the crust beneath GRA is 20±22 km thick with
a 3.5±5 km thick upper crust. LVZs exist in the lower crust to
the north and southwest of GRA but the crust to the east is
characterized by a set of positive steps in velocity.
A3
Station GRI
The receiver functions at GRI have strong amplitude
arrivals on both the radial and tangential waveforms and
there is signi®cant variation in the waveforms with event
azimuth. Receiver function clusters from similar event azimuths
have strong similarities in the waveforms. We interpret the
prominent arrival at 2 s in the eastern azimuth receiver
functions as a P-to-S conversion at the Moho, suggesting a
relatively thin crust.
Inversion results for events 98150062228 and 98158232013
are shown in Fig. A3 (grey lines). These receiver functions were
chosen for analysis because they are the quietest waveforms
from two of the three event azimuth clusters. The northern
azimuth receiver functions were not analysed since the tangential receiver functions have a higher amplitude than the
radial receiver functions. Two model families were obtained for
98150062228. The ®rst has a strong positive velocity gradient to
y6 km depth, where the velocity reaches almost 4.5 km sx1.
Between 6 and 17 km the velocity is nearly constant. At 17 km
the velocity increases abruptly to almost y5 km sx1. The
second model family has lower average velocities than model 1.
There is a strong positive velocity gradient to y4 km depth,
beneath which the velocity is nearly constant to y20 km
depth. Between 20 and 25 km depth there is another positive
(a)
(b)
velocity gradient, with the Moho at y25 km depth. Inversion
of 98158232013 gives very low surface velocities (<1 km sx1)
and a strong positive velocity gradient to y4 km depth. Below
this the velocity gradient is weaker and a velocity of 4.1 km sx1
is reached by y15 km depth. There is no obvious structure in
the model that could represent the base of the crust and we do
not consider this model further.
Fig. A3 (thick solid lines) shows the forward modelling
results for 98150062228. The simpli®cation of inversion model
1 has a strong positive velocity gradient in the uppermost 5 km
and a sharp Moho at 16 km depth. This model ®ts the arrivals
at 2.5 and 7 s, which are modelled as the Moho conversion
and reverberation, respectively. Later arrivals are ®tted by
an LVZ at 22±28 km depth. The average crustal velocities are
high (>4 km sx1 below 6 km depth). The second model has
a thinner (4 km) uppermost section, lower average seismic
velocities and a thicker crust. The Moho lies at 25 km depth,
at the base of a section of positive velocity gradients. The two
LVZs in the model are required to improve the ®t to the
waveform at 5±10 s. Fig. A3(d) shows the ®ts of the models to
two noisier receiver functions (97133141345 and 97130075729)
from similar azimuths to 98150062228. The two models are
able to provide some ®t to the early part of these waveforms.
Neither model can be said to ®t the 98150062228 receiver
function better than the other, therefore we have two different
crustal structures possible for the region east of GRI. The ®rst
has high seismic velocities, an upper crustal thickness of 5 km
and a Moho depth of 16 km. The second has consistently lower
seismic velocities, an upper crustal thickness of 4 km and a
Moho depth of 25 km.
(c)
(d)
Figure A3. Modelling results for the 98150062228 (a,b) and 98158232013 (c) GRI receiver functions. Refer to Fig. A1 for plotting conventions.
(d) Shows the ®ts of the 98150062228 models to noisier receiver functions from similar backazimuths.
#
2000 RAS, GJI 143, 163±184
Receiver structure of central and northern Iceland
A4
Station SIG
Radial receiver functions at SIG are mostly of low amplitude
compared to the tangential waveforms and the pre-signal noise.
Similarities between receiver functions in azimuthal clusters are
observed, but only event 98150062228 has a suf®ciently large
radial receiver function for 1-D modelling.
Two different model families were obtained from the inversion
of 98150062228 (Fig. A4, grey lines). Both have a thin (y2 km)
upper crust with a strong positive velocity gradient overlying
a thick section of near-constant velocity, below which there
is a gradational transition to a half-space. Model family 1 has
consistently higher seismic velocities than model family 2 and,
correspondingly, a deeper transition to a half-space (25±30 km
for family 1, 20±25 km for family 2). Family 2 has a strong
positive velocity gradient at y50 km depth. Whether this is a
signi®cant feature or an artefact from 3-D structure is not clear.
The forward models for SIG are shown in Fig. A4 (thick
solid lines). Both have surface velocities of 2.3 km sx1 and a
3 km thick upper crust with a strong positive velocity gradient.
Model 1 shows no overall velocity increase between 3 and
22 km depth but an LVZ at 11±16 km depth is required to ®t
the waveform between 4 and 9 s. The Moho is a positive step in
velocity at 30 km depth below an 8 km thick section of positive
velocity gradient. Model 2 has a similar structure but there is
(a)
(b)
Figure A4. Modelling results for the 98150062228 SIG receiver
function. Refer to Fig. A1 for plotting conventions.
#
2000 RAS, GJI 143, 163±184
181
no LVZ. A section of positive velocity gradient lies between
20 and 26 km depth and a velocity step is modelled at 50 km
depth. The seismic velocities are lower than those in model 1
and the velocity of 4.1 km sx1 between 26 and 40 km depth
is more often associated with the lower crust than the upper
mantle in previous seismic studies of Iceland.
Since the arrivals close to the apparent Pp are distorted by
negative side-lobes, which are also present in the averaging
function, we do not attempt to ®t their amplitudes closely. The
models shown are the simplest that we consider to be an
adequate ®t to the data.
In summary, the crust beneath SIG is y25±30 km thick. The
upper crust is 3 km thick with higher surface velocities than
those found elsewhere in central and northern Iceland.
A5
Station ASB/BORG
The receiver functions for ASB/BORG have good snrs and
similar features within azimuthal clusters of waveforms. A
strong arrival at y4 s after the direct P wave seen on most of
the receiver functions is likely to be a P-to-S conversion at the
Moho. Four receiver functions of similar snr from the Japan
and Kuril Islands region were suf®ciently close in azimuth and
distance to be stacked (Cassidy 1992), and the stacked receiver
functions (Fig. A5) were used in further analysis. In addition,
the quietest receiver function from the southwest azimuth
(97301061517) was analysed separately.
We inverted the Japan/Kurils stacked receiver function and
the receiver function from event 97301061517 (Fig. A6, grey
lines). The inversion of the Japan/Kurils waveform gave three
model families. All have a similar structure; the difference
between the models is the average seismic velocity and, correspondingly, the depth to the Moho. The uppermost 3±5 km of
the models has a strong positive velocity gradient and there is a
smaller positive velocity gradient between 4 and 8 km depth.
Between 10 and 25 km depth, there is a very small overall velocity
Figure A5. Receiver function stack for ASB/BORG from events
from the Japan/Kurils region. Top: stacked waveform (solid line) and
one standard deviation stacking bounds (dotted lines). The receiver
functions making up the stack are plotted below, and are identi®ed by
the event origin times.
182
F. A. Darbyshire et al.
(a)
(b)
(c)
(d)
(e)
Figure A6. Modelling results for the Japan/Kurils (a±c) and 97301061517 (d) ASB/BORG receiver functions. Refer to Fig. A1 for plotting
conventions. (e) Shows the ®ts of the models to noisier receiver functions from similar backazimuths (top three panels: N; bottom panel: SW) .
increase. The Moho is a prominent velocity increase with a high
gradient at 25±30 km depth. Inversion of 97301061517 yields a
similar model except that the transition to mantle velocities is
much more gradual than that modelled from the Japan/Kurils
receiver function stack.
Fig. A6 (thick solid lines) shows the forward modelling
results for ASB/BORG. The three models from the Japan/
Kurils receiver function stack have an uppermost 3±4 km thick
section with a strong positive velocity gradient, below which is
a y5 km section of lower velocity gradient. The lower crustal
section has little or no velocity increase, and LVZs with small
velocity contrasts are modelled between 12 and 18 km depth.
The LVZs are required to bring the synthetic waveforms within
the bounds of the receiver function stack between 6 and 8 s.
The Moho is a sharp increase in velocity at 24 km (model 3),
25 km (model 1) or 28 km (model 2) depth. The forward model
for 97301061517 has a similar upper crust to the Japan/Kurils
models. An LVZ at 10±12 km depth is required to ®t the
waveform between 4 and 8 s. The section of the model between
18 and 28 km depth has a strong positive velocity gradient
and the Moho is interpreted to lie within this transitional zone
between 24 and 28 km depth. Fig. A6(e) shows the ®ts of the
forward models to noisy receiver functions from similar event
azimuths to the receiver functions analysed. In each case the
main features of these waveforms are matched by the crustal
models shown.
The crust beneath ASB/BORG is 24±28 km thick. To the
north, the structure is simple, with a sharp Moho. Southwest
of ASB/BORG, the Moho is a more gradational feature. The
upper crust is 3±4 km thick.
A6
Station HVE
The arrivals in the radial receiver functions for HVE have
strong amplitudes, and receiver functions in clusters of similar
event azimuth show similar features. A peak at 4±5 s after the
direct P-wave arrival is visible on most of the radial waveforms.
Receiver functions from events 97301061517 and 98152053403
were inverted; these waveforms are the quietest from the two
dominant arrival azimuths (north and southwest) in the data
set. The results are shown in Fig. A7 (grey lines). Model family
1 for 97301061517 has a 4 km thick upper crust with a strong
positive velocity gradient, below which the velocity is nearly
uniform to 18 km depth. From 18 to y40 km depth, there
is another positive velocity gradient and mantle velocities exist
below y35 km depth. Model 2 shows a similar velocity gradient
in the upper crust, a weak positive velocity gradient to y12 km
depth and small oscillations in the velocity below this depth.
The 98152053403 model has a 3 km thick upper crust below
which the velocity increases in a series of steps to y35 km
depth.
Forward models for HVE are shown in Fig. A7 (thick solid
lines). Model 1 for 97301061517 has a 5 km thick upper crust
with a surface velocity of 1.1 km sx1 and a strong positive
velocity gradient. Between 5 and 18 km depth there is no velocity
increase and two LVZs are required to ®t the waveform
between 3 and 10 s. Between 18 and 32 km depth the velocity
gradient is positive and a velocity discontinuity at 32 km depth
is interpreted to be the Moho. Model 2 has a similar uppermost
structure to model 1. There is a strong positive velocity gradient
at y5 km depth, a near-constant velocity to y12 km depth
#
2000 RAS, GJI 143, 163±184
Receiver structure of central and northern Iceland
(a)
(b)
(c)
183
(d)
Figure A7. Modelling results for the 97301061517 (a,b) and 98152053403 (c) HVE receiver functions. Refer to Fig. A1 for plotting conventions.
(d) Shows the ®ts of the models to noisier receiver functions from similar backazimuths (top four panels: SW; bottom two panels: N).
(a)
(b)
(c)
Figure A8. Modelling results for the 97130075729 SKR receiver function. Refer to Fig. A1 for plotting conventions. (c) Shows the ®ts of the
models to a noisier receiver function from a similar backazimuth.
#
2000 RAS, GJI 143, 163±184
184
F. A. Darbyshire et al.
and two LVZs between 15 and 30 km depth. Seismic velocities
of >4.1 km sx1 occur below 14 km depth and mantle velocities
of 4.5 km sx1 and above appear from y30 km depth. The
forward model for 98152053403 has a 3 km thick upper crust
with a strong positive velocity gradient. Positive steps in
velocity occur at 10, 19 and 29 km depth, the deepest of which
is interpreted to be the Moho. The small negative velocity
steps at 15 and 24 km depth increase the amplitude of the
synthetic receiver function to near to that of the observed
receiver function, but are not likely to be signi®cant features.
Fig. A7(d) shows the ®ts of the crustal models to noisy receiver
functions from the southwest and north event clusters. The
main features of these receiver functions are ®tted by the
97301061517 (top four panels) and 98152053403 (bottom two
panels) models.
The crust beneath HVE is 29±32 km thick with a 3±5 km
thick upper crust. North of HVE the crust contains a set of
positive velocity steps. To the southwest the velocity increase is
more gradual and two LVZs occur in the lower crust.
A7
Station SKR
We observe similar amplitudes on the radial and tangential
receiver functions at SKR for all events except 97130075729.
Two strong peaks at 2.5 and 5 s are seen in the 97130075729
radial waveform and these are also observed as weaker arrivals
in receiver functions from other azimuths.
The receiver function from 97130075729 was inverted, giving
the models shown in Fig. A8 (grey lines). The ®rst model family
has a strong positive velocity gradient in the uppermost 5 km,
a nearly constant velocity between 8 and 18 km depth and a
gradual increase in velocity between 18 and 35 km depth. The
second model family has an oscillatory structure with high
(>4.5 km sx1) seismic velocities below 13 km depth.
Forward modelling results for 97130075729 are shown in
Fig. A8 (thick solid lines). Model 1 has a surface velocity of
1.3 km sx1 and a 5 km thick upper crust with a strong positive
velocity gradient. A `lid' structure occurs at 5±8 km depth
and there is a positive step in velocity at 19 km depth. The
Moho lies at 37 km depth below a 9 km thick section with
a positive velocity gradient. Model 2 has a strong positive
velocity gradient to y2 km depth, below which there is a
weaker positive velocity gradient to y22 km depth, where the
velocities reach the Moho value of 4.6 km sx1. The seismic
velocities in the model are high (i4 km sx1 below 7 km depth).
Both models show acceptable ®ts to the 97130075729 receiver
function, taking the pre-signal noise levels into account, and
similar degrees of ®t are also made to a noisier receiver function
(97133141345) from a similar azimuth (Fig. A8c).
In summary, two different crustal structures are possible
for SKR. The ®rst has a Moho depth of 37 km with a 5 km
thick upper crust and a velocity step at 19 km depth. The
second has high average seismic velocities and a Moho depth
of 22 km.
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2000 RAS, GJI 143, 163±184