crustal velocity structure of the rukwa rift in the western

Transcription

S. KIM, A.A. NYBLADE AND C-E. BAAG
251
CRUSTAL VELOCITY STRUCTURE OF THE RUKWA RIFT IN THE
WESTERN BRANCH OF THE EAST AFRICAN RIFT SYSTEM
S. KIM
School of Earth and Environmental Sciences, Seoul National University, Seoul, South Korea
e-mail: [email protected]
A.A. NYBLADE
Department of Geosciences, Pennsylvania State University, University Park, Pennsylvania, USA
University of the Witwatersrand, Private Bag 3, WITS, 2050 South Africa
C-E. BAAG
School of Earth and Environmental Sciences, Seoul National University, Seoul, South Korea
© 2009 December Geological Society of South Africa
ABSTRACT
We investigate the thickness and seismic velocity structure of the Rukwa Rift crust by modeling seismograms from the 1994 Mw
5.9 Rukwa earthquake and teleseismic receiver functions recorded on a broadband seismic station at the southeastern end of the
rift. Two methods have been used to model receiver functions, H- stacking and waveform inversion, yielding Moho depth
estimates of 34 ± 2 km (H- stacking) and 38 ± 2 km (waveform inversion), and a crustal Poisson’s ratio of 0.27 ± 0.01
(H- stacking). A 1D average velocity model for the rifted crust has been obtained by trial-and-error modeling of the Rukwa
earthquake seismograms. The best fitting velocity model obtained has a 4.5 km thick near-surface low-velocity section over a
33 km thick middle-to-lower crustal section. The velocities and thickness of the near-surface layers are consistent with the known
thickness and composition of the sedimentary basin fill. The middle-to-lower crustal section is characterised by a linear velocity
gradient with shear wave velocities of 3.9 km/s or higher at depths > 20 km. Crustal Poisson’s ratio in the model is 0.27. The Moho
depth in the model (37.5 km) is similar to the Moho depth obtained from analysing receiver functions, suggesting that there has
been little, if any, crustal thinning beneath the rift compared to the southeastern edge of the rift. The Poisson’s ratio of 0.27, in
combination with the high velocities in the middle and lower crust, indicate that a significant component of the rifted crust has a
mafic composition. This finding is important for understanding the occurrence of lower crustal earthquakes in east Africa, because
it suggests that at least some of the crust in east Africa could be sufficiently strong at mid- and lower crustal depths to support brittle
deformation.
Introduction
In this paper, we model broadband seismic data from
teleseismic and regional earthquakes recorded by a
seismic station at the southeastern end of the Lake
Rukwa Rift in Tanzania to investigate the thickness and
seismic velocity structure of the rifted crust (Figure 1).
Imaging crustal structure beneath the Rukwa Rift is not
only important for determining the amount of crustal
modification that may have occurred as a result of
lithospheric extension, but also for improving our
understanding of why lower crustal earthquakes occur
frequently in east Africa.
The only published estimates of crustal thickness
beneath the Rukwa rift come from modeling the travel
times of P reflections from the Moho (PmP) generated
by regional earthquakes and recorded on a local seismic
network around the rift (Camelbeek and Iranga, 1996).
Camelbeek and Iranga (1996) reported an average Moho
depth for the rift of 42 km, which is some 5 to 7 km
thicker than the crust at the southeastern end of the rift
determined by Last et al. (1997) from modeling receiver
functions for station PAND (Figure 1), and by Julià et al.
(2005) from the joint inversion of receiver functions and
Rayleigh wave phase and group velocities. The different
estimates of crustal thickness make it difficult to
determine the amount of crustal modification, if any,
associated with the rift.
Over the past few decades, a number of papers have
commented on the occurrence of earthquakes in the
East African Rift System at lower crustal depths of >20
km (e.g., Shudofsky, 1985; Shudofsky et al., 1987;
Jackson and Blenkinsop, 1993; Nyblade and Langston,
1995; Zhao et al., 1997; Foster and Jackson, 1998; Brazier
et al., 2005; Keir et al., 2009). Most of the earthquakes
have occurred in the southern half of the rift system,
either along the Western Branch in the Lake Tanganyika,
Rukwa and Malawi rifts, or in northeastern Tanzania
where the Eastern Branch impinges on the eastern
margin of the Archean Tanzania Craton. Many of the
earthquakes have magnitudes >5, representing a
significant portion of the seismic strain release across
east Africa. Nyblade and Langston (1995) investigated
what rheological and thermal properties of the lower
crust would be required within the standard crustal
SOUTH AFRICAN JOURNAL OF GEOLOGY, 2009, VOLUME 112 PAGE 251-260
doi:10.2113/gssajg.112.3-4.251
252
CRUSTAL VELOCITY STRUCTURE OF THE RUKWA RIFT, EAST AFRICAN RIFT SYSTEM
Figure 1. Map of the Rukwa Rift showing the location of the 18 August 1994 Lake Rukwa earthquake (star) and seismic stations of the 199495 Tanzania broadband seismic experiment (triangles). The gray colored area in the right-upper inset shows location of map with respect to
the main Precambrian terrains (Tanzania Craton, Ubendian Belt and Mozambique Belt), and political boundaries. The bold black lines and
the gray dashed lines denote border faults and seismic reflection lines, respectively.
rheological model (Kirby, 1980; Brace and Kohlstedt,
1980, Jackson, 2002) for brittle failure to occur at depths
of > 20 km in east Africa, where average heat flow is
about 60 to 65 mWm-2 (Nyblade, 1997). They concluded
that the lower crust must have a mafic composition,
otherwise the lower crust would not be sufficiently
strong to support brittle failure.
Further to that conclusion, Julia et al. (2005)
examined details of crustal structure in Tanzania by
jointly inverting receiver functions and Rayleigh wave
phase and group velocities using data recorded by the
Tanzania broadband seismic experiment stations
(Nyblade et al., 1996). Julia et al. (2005) found shear
wave velocities of >3.9 km/s in the bottom ~5 to 8 km
of the crust beneath many stations away from the rift
valleys, indicating the presence of mafic lithologies.
Given this result, earthquakes in east Africa nucleating
within the lowermost part of the crust outside of the rifts
can be explained by brittle deformation within mafic
rock, provided that temperatures in the lowermost crust
have not been greatly elevated by the extensional
tectonism. But this finding leaves open the question of
why earthquakes beneath the rift valleys nucleate at
depths between about 20 km and the bottom 5 to 8 km
of the crust, where the crustal composition is less likely
to be mafic. One possibility is that beneath some of the
rifts a larger portion of the lower crust is mafic than
beneath the unrifted terrains. And given that possibility,
the second purpose of this paper is to obtain an estimate
of the average seismic velocity structure of the Rukwa
rift crust to determine whether or not velocities below
about 20 km depth are consistent with the presence of
mafic lithologies.
To estimate the thickness and velocity structure of
the Rukwa Rift crust, we first examine crustal structure
at the southeastern end of the rift through a re-analysis
of receiver functions for station PAND (Figure 1) using
two methods, H- stacking (Zhu and Kanamori, 2000)
and waveform inversion (Ammon et al., 1990). We then
model regional waveforms from the Mw 5.9 August 18,
1994 Rukwa earthquake recorded on station PAND
(Figure 1) to obtain a 1D velocity model for the rifted
crust.
Background
The Rukwa Rift is located in the Proterozoic Ubendian
mobile belt to the south of the Archean Tanzania Craton
(Figure 1). The rift is surrounded by exposed basement
SOUTH AFRICAN JOURNAL OF GEOLOGY
253
Table 1. List of teleseismic events used for computing receiver functions
Date
(yyyymmdd)
19940615
19940630
19940725
19940808
19940915
19940916
19940928
19941025
19941031
19941114
19941115
19950103
19950506
Time
(hhmm)
0923
0923
2200
2108
2347
0620
1639
0054
1148
1916
2018
1612
0159
Latitude
-10.34
36.33
-56.36
24.72
-57.8
22.53
-5.79
36.36
3.02
13.52
-5.59
-57.7
24.99
Longitude
113.66
71.13
-27.36
95.2
-8.77
118.71
110.35
70.96
96.19
121.07
110.19
-65.88
95.29
(deg)
79.1
57.4
66.5
69.0
58.4
89.3
76.4
57.4
63.9
90.0
76.3
87.3
69.2
BAZ
(deg)
98.8
36.0
211.9
59.3
204.9
67.2
93.8
35.9
82.2
76.3
93.6
212.1
59.0
Magnitude
(Mb)
6.0
6.5
6.3
6.0
5.2
6.5
5.9
6.2
5.7
6.1
6.2
6.2
6.4
Depth
(km)
19.9
226.6
81.3
121.7
10
13.1
637.5
238.7
29.4
31.5
560.6
13.9
117.5
Fit
(%)
97.2
93.2
90.1
91.5
79.8
98.7
96.6
82.9
83.7
97.4
96.8
78.0
95.9
rock and is bounded to the northeast and southwest by
two major border faults (Lupa and Ufipa, Figure 1).
The upper crustal structure of the Rukwa Rift has been
investigated in a number of studies (e.g. Peirce and
Lipkov, 1988; Kilembe and Rosendahl, 1992; Morley
et al., 1992; Mbebe, 1993; Wheeler and Karson, 1994).
Interpretations of seismic reflection profiles show rift
faults extending to a depth of at least 10 km (Morley
et al., 1992; Mbebe, 1993) while seismicity studies of
lower crustal earthquakes suggest that faulting may
extend to lower crustal depths (Camelbeeck and Iranga,
1996; Zhao et al., 1997).
The region has been affected tectonically several
times since the Late Neoproterozoic, and the
sedimentary record reflects periods of tectonically
controlled sedimentation during the Karoo, Cretaceous,
Paleogene, and Neogene-Quaternary (Dypvik et al.
1990; Roberts et al., 2004). The Karoo Supergroup
consists predominantly of continental deposits
(conglomerates, sandstones, shales, and coal) sitting on
Precambrian basement. The Karoo sediments are
overlain by the Cretaceous-Paleogene Red Sandstone
group, and at the top of the stratigraphic sequence are
unconsolidated Neogene-Quaternary lake deposits.
The total thickness of the sedimentary layers is 5 to
10 km (Morley et al. 1992; Mbebe, 1993). Volcanism
during the Neoproterozic, Cretaceous and Miocene
occurred within the vicinity of the Rukwa Rift
(Rasskazov et al. 2003; Ebinger, 1989), and periods of
tectonic inversion and lateral shearing occurred during
the Late Neoproterozoic and Permo-Triassic (Delvaux,
2001; Ebinger, 1989; Klerkx et al., 1998; Lenoir et al.,
1994; Theunissen et al., 1996).
Figure 2. Computed receiver functions and event location map.
(a) Distribution of teleseismic earthquakes used for computing
receiver functions. The earthquakes are located within an
epicentral distance range of about 57° to 90° from station PAND.
(b) Plot of the receiver functions aligned on the direct P-wave
arrival time and sorted by backazimuth. The recognisable Moho
converted phases (Ps, PpPs, PsPs+PpSs) are denoted by arrows.
Data
The data used in this study come from the Tanzania
broadband seismic experiment (Nyblade et al. 1996).
In this experiment, twenty seismic stations were
deployed for one year (1994-1995) across Tanzania, with
station PAND located at the southeastern end of the rift
(Figure 1). Other nearby stations included TUND, AMBA
254
Table 2. Crustal models for the Rukwa Rift and Tanzania
Source
Average Vp
(km/s)
6.4 ± 0.25
6.2 ± 0.2
Poisson’s
Ratio
0.24 ± 0.02
Langston et al., 2002
Moho depth
(km)
Receiver function
35 ± 4
PmP traveltime analysis
42 +4 and -5
Joint inversion of receiver function 38 ± 2
and Rayleigh wave dispersion
Phase time inversion
40.4 ± 0.6
6.47 ± 0.05
0.25 ± 0.005
This study
H- stacking
6.3
6.4
6.5
6.6
6.7
6.8
6.92 ± 0.06
6.51
0.27
0.27
0.27
0.27
0.27
0.27
Last et al., 1997
Camelbeeck and Iranga, 1996
Julià et al., 2005
Method
Receiver function inversion
Waveform modeling
32.6
33.0
33.6
34.2
34.8
35.4
38.0
37.5
±
±
±
±
±
±
±
0.10
0.06
0.10
0.07
0.11
0.06
2.0
0.27
±
±
±
±
±
±
0.001
0.001
0.001
0.001
0.001
0.001
Region
station PAND
Rukwa Rift
station PAND
Tanzania Craton
and vicinity
station PAND
station PAND
station PAND
station PAND
station PAND
station PAND
station PAND
Rukwa rift
Figure 3. Plot of the H- stacking results. Contours map out normalised values of the H- stacking objective function. The P-wave velocity
varies from 6.3 km/s to 6.8 km/s.
and MTAN, but they were situated either just outside the
rift (TUND) or well away from the rift (AMBA, TUND)
(Figure 1). Details of the station equipment and recording
parameters can be found in Nyblade et al. (1996).
During the experiment, a Mw 5.9 earthquake
occurred on 18 August 1994 at a depth of 25 km at the
northwestern end of the Rukwa Rift (Figure 1).
The earthquake was well recorded on station PAND, and
the source-receiver path extends along the axis of
the Rukwa Rift (Figure 1). Source parameters for the
earthquake were determined by Zhao et al. (1997) from
modeling teleseismic waveforms.
The arrival time of the initial P wave (Pn) at station
PAND was found to be about 10 seconds late compared
to the expected time arrival time using the average
crustal model for Tanzania from Langston et al. (2002).
The time delay is too large to be caused by realistic
event-station path effects related to crustal and upper
mantle structure, and therefore there must have been a
timing problem with station PAND. The arrival time at
nearby TUND station is consistent with the average
crustal model of Langston et al. (2002). Given this, and
the similar epicentral distance for TUND (231 km)
compared to PAND (237 km), we shifted forward the
Pn arrival time for the PAND waveforms by 10.6 seconds
to make them consistent with the Pn arrival time at
TUND, taking into consideration the slight difference in
epicentral distance.
Receiver Function Analysis
For re-examining crustal structure beneath station PAND,
we have modeled teleseismic receiver functions from
13 events using two methods, H- stacking and
waveform inversion (Table 1). These events cover
backazimuths of 36 to 212 degrees and epicentral
distances of 57 to 90 degrees. Receiver functions were
calculated using the iterative time domain deconvolution
method of Ligorria and Ammon (1999), and are
illustrated in Figure 2.
In the H- stacking technique (Zhu and Kanamori,
2000), Moho depth (H) and Vp /Vs () are determined by
a grid search procedure to find a combination of H and that yield receiver functions which best match the
arrival times of the Ps, Pp Ps and Pp Ss+Ps Ps phases.
The objective function in the stacking technique is
defined by
s (H,) = w 1r (t 1 ) + w2r (t 2 ) + w 3r (t 3 )
(1)
255
where, r (t ) are the radial receiver functions, w i is the
weighting factors and t i is the calculated arrival time of
Ps, Pp Ps and Pp Ss+PsPs corresponding to a given H and value. The results of the H- stacking are presented in
Figure 3 and Table 2 using weights of 0.6, 0.3 and
0.1 for Ps, Pp Ps and Pp Ss+PsPs, respectively. Estimated
Moho depth varies from 33 to 35 km/s depending on the
assumed average crustal P-wave velocity (6.3 to
6.8 km/s) used for the stacking. A Poisson’s ratio of
0.27 (Vp/Vs value of 1.77) is also obtained. The
uncertainty values shown in Table 2 are 1 errors
calculated using the bootstrap method of Efron and
Tibshirani (1991).
In addition to the H- stacking, the receiver function
inversion method described in Ammon et al. (1990) was
used to obtain a one-dimensional crustal velocity model
for station PAND. For this method, only the receiver
functions with a >90% fitting rate in the iterative
deconvolution process were used (Table 1). The model
consisted of 2 km thick layers, and for the starting
model, the Vs in all layers was set to a mean crustal
value. The mean crustal values ranged from 3.5 km/s to
3.8 km/s (Figure 4). The receiver functions were
inverted with the Poisson’s ratio fixed to the value
obtained from the H- stacking. The percent of receiver
Figure 4. S-wave velocity profiles obtained from receiver function inversion. Dashed lines show the starting models used for the inversion.
Right panel shows examples of receiver function comparisons between receiver functions from data and synthetics obtained using a one-layer
crustal model with specified mean P- and S-wave velocities.
256
Figure 5. Three-component displacement seismograms of the 18 August 1994 Lake Rukwa earthquake recorded at station PAND. Vertical,
radial and tangential seismograms are shown. Major phases are labeled and denoted by arrows.
function fit by the final models in each case was >81%,
and the resulting velocity models are very similar to each
other (Figure 4). The models are characterised by lowvelocity surface layers with a thickness of about 4 km, a
high velocity layer with a thickness of about 10 km just
below the surface layers, and a Moho at about 38 km
depth. The Moho is placed where there is a strong
velocity discontinuity marked by a Vs increase to
between 4.4 and 4.7 km/sec (Figure 4). The uncertainty
in the Moho depth estimate is ± 2 km, which arises
primarily from the use of 2 km thick layers in the model
and the influence of the starting model on the inversion
result.
Regional Waveform modeling
To image crustal structure beneath the Rukwa Rift, we
have modeled regional waveforms from the 18 August
1994 Rukwa earthquake recorded at station PAND
(Figure 5). For computing synthetic seismograms, the
reflectivity method of Kennett (1983) was used, which
gives full waveform responses for a one-dimensional
layered medium. In applying this method, we used
the focal mechanism and source-time function for the
Rukwa earthquake reported by Zhao et al. (1997) and
modified by Langston et al. (2002).
The 1D average crustal model for Tanzania from
Langston et al. (2002) produces synthetic seismograms
that match well the seismograms recorded at most of the
stations in the Tanzania broadband seismic experiment
(Langston et al., 2002). However, the synthetic
seismograms from this 1D model do not match the
waveforms at station PAND (Figure 6, model ‘L’;
Figure 7). The data show relatively high frequency peaks
near the Pn and Sn arrivals, unlike the waveforms at
other stations in Tanzania (Langston et al. 2002). The
high frequency phases could be multiples partially
trapped or scattered in near-surface low-velocity layers
(Dreger and Helmberger, 1990), and large impedance
changes between those layers could explain the large
amplitude peaks of these phases.
In the initial modeling step, we used the 1D model
from Langston et al. (2002), and added several lowvelocity near-surface layers to simulate the effects of the
sedimentary basin fill (Figure 6). One of the models is
shown in Figure 6 (model a). From doing this, it was
noticed that the phases arriving after Pn and Sn in the
synthetics using the Langston et al. (2002) velocity
model (“L”) split into several small peaks (Figure 7),
better matching the data. Models having two or more
low-velocity near-surface layers were found to produced
better matches to the waveform shape than models with
only one low-velocity near-surface layer.
A trial-and-error procedure was followed to obtain a
best fitting velocity model by imposing perturbations of
Figure 6. P- and S-wave velocity models. Model L is from Langston
et al. (2002). Models (a) and (b) are examples of models with nearsurface low-velocity layers and a perturbed lowermost crustal
structure. The model c is the final velocity model that gives
synthetics that best match the data.
velocities and layer thickness on the initial velocity
model (model ‘L’), with emphasis placed on model
parameter variation of the near-surface layers. In this
257
process, the following model parameters were adjusted;
crustal thickness, thickness and velocities of the lowvelocity near-surface layers, Poisson’s ratio, Pn and
Sn velocities, lowermost crustal velocities, and the
velocity gradient of mid-crustal layers. Figure 8 shows a
total of 133 velocity models used to compute synthetic
seismograms obtained by varying these model
parameters.
A comparison between the data and a few
representative synthetic seismograms, along with the
synthetics for the best-fitting ‘final’ model are shown in
Figure 7. The corresponding velocity models are given
in Figure 6. The synthetic seismograms in Figure 7
model (c) provide the best match to the data, and the
corresponding velocity model is given in Figure 6
(model c) and Table 3. Synthetic seismograms from the
final model are also shown with the data in Figure 9.
Key features in the final model required to match the
data include low-velocity near-surface layers (Vp range
from 2.5 to 4.0 km/s) and a large contrast in velocities at
the bottom of the low velocity near-surface layers. If the
low-velocity near-surface layers are not included in
the model, then the high-frequency phases following
Pn and Sn are not generated. To obtain the high
frequency peaks in synthetic seismograms following Pn
and Sn, a large velocity contrast at the base of the
sedimentary layers is needed.
Although the broadband seismograms are
complicated, the synthetic waveforms from the final
model show a reasonable match to the data, as can be
seen by the alignment of the phases in Figure 7, models
Figure 7. Comparison of synthetic seismograms with the broadband displacement waveforms of the 18 August 1994 Lake Rukwa event
recorded at station PAND. The top traces of each component (vertical, radial, tangential) are the data and the other traces are synthetic
seismograms corresponding to the models (L), (a), (b) and (c) in Figure 6. Characteristic phases are denoted by arrows on the data and
synthetic seismograms for model (c). The number given at the beginning of each trace gives the maximum amplitude of the seismogram in
centimeters.
258
Figure 8. Plot of all models used to compute synthetic seismograms. Bold and thin lines show P- and S-wave velocity models, respectively.
(a) and (c) marked by arrows. Even though some
of high-frequency peaks in the synthetic vertical
component seismogram differ from the data, the overall
resemblance of the synthetic waveforms to the data is
good, especially for the timing and shape of major
phases. A difference between the synthetic seismograms
and data for the radial component can also be seen in
Table 3. Final velocity model
Depth
(km)
0.0 – 1.5
1.5 – 3.0
3.0 – 4.5
4.5 – 36.3
36.3 – 37.5
Moho
Vp
(km/s)
2.50
3.20
4.00
6.50 – 7.30
(Linear gradient)
7.70
8.29
Vs
(km/s)
1.40
1.80
2.25
3.65 – 4.10
4.32
4.75
the timing of Sn. Because the Sn arrival time for both the
vertical and tangential component synthetics match
the data fairly well, the mismatch in the timing of Sn on
the radial component is probably caused by a local
structural effect. Similarly, waveform mismatches for the
S wavetrain in later parts of the tangential seismogram
are likely due to three-dimensional wave propagation
effects. To model three-dimensional structure, different
methods to compute synthetic seismograms are
required, such as the finite-difference method using
discontinuous grids with locally variable time steps
(Kang and Baag, 2004).
Discussion and conclusion
The estimates of Moho depth obtained from analysing
receiver functions (34 km for H- stacking; 38 km from
waveform inversion) for station PAND are consistent
with the Moho depth of 35 ± 4 km obtained by Last
et al. (1997) and 38 ± 2 km by Julià et al. (2005), within
Figure 9. Comparison of the broadband displacement waveforms
of the 18 August 1994 Lake Rukwa earthquake recorded at station
PAND with synthetic seismograms for the best fitting velocity model
(Table 3 and Figure 7, model c).
the reported uncertainties. The Poisson’s ratio of 0.27
that we obtained is larger than the Poisson’s ratio of 0.24
reported by Last et al. (1997) for station PAND.
The receiver functions show variation with back-azimuth
(Figure 2b), and a second arrival can be seen on some of
the receiver functions following the Moho Ps (Figure 2b).
These observations indicate that crustal structure
beneath station PAND is laterally heterogeneous.
In the 1D velocity model for the Rukwa Rift
(Figure 6, model c), the 4.5 km-thick near-surface lowvelocity layers in the model overlay a 33 km-thick
middle-to-lower crustal section. The velocities and
thicknesses of the near-surface layers are consistent
with the known thickness and composition of the
sedimentary basin fill. The middle-to-lower crustal
259
section is characterised by a linear velocity gradient with
higher-than-average velocities for Precambrian crust
in Tanzania, which are represented by model ‘L’ in
Figure 6. The crust is 37.5 km thick, with an average
P-wave velocity of 6.51 km/s and a Poisson’s ratio of
0.27. The Moho depth is similar to the Moho depth
beneath station PAND, but somewhat shallower than the
Moho depth of 42 km reported by Camelbeeck and
Iranga (1996).
Given this result, there appears to be little, if any,
variation in crustal thickness beneath the Rukwa
Rift compared to the southeastern edge of the rift.
Thus, if crustal thinning has occurred beneath the
Rukwa Rift, it is too small (i.e., not more than a few kms)
to be detected by the data and modeling methods used
in this study.
The Poisson’s ratio of 0.27 determined for the Rukwa
rift crust is comparable to that for the crust beneath
station PAND and indicates that a significant component
of the crust is mafic in composition (Zandt and Ammon,
1995). This conclusion is supported by the high midand lower crustal velocities found beneath the rift and
station PAND (Figures 4 and 6). For example, shear
wave velocities within the middle and lower parts of the
crust approach or exceed 3.9 km/s. Shear velocities in
the crust that are >3.9 km/s are indicative of mafic
lithologies (Holbrook et al. 1992; Christensen and
Mooney, 1995; Rudnick and Fountain, 1995; Rudnick
and Gao, 2003). For the Rukwa Rift, shear wave
velocities of >3.9 km/s are found at depths of >20 to
25 km (Figure 6, model c). For station PAND, shear wave
velocities that high are found in both the upper and
lower crust (Figure 4).
The high velocity (i.e, Vs >3.9 km/s) rock beneath the
rift at depths >20 to 25 km has important implications
for understanding the occurrence of lower crustal
earthquakes in east Africa, because, as was discussed in
the introduction, it suggests that some of the crust in east
Africa could have a predominately mafic composition
throughout its lower half and thus be sufficiently strong
to support brittle deformation.
Acknowledgements
We would like to thank Raymond Durrheim, Damien
Delvaux and Artur Cichowicz for their constructive
reviews that improved the manuscript. This research was
funded by the Korea Meteorological Administration
Research and Development Program (Grant CATER
2006-5205) while CEB was on sabbatical leave at Penn
State. Funding for this study has also been provided
by the National Science Foundation (grant OISE
0530062).
References
Ammon, C.J., Randall, G.E. and Zandt, G. (1990). On the nonuniqueness of
receiver function inversions. Journal of Geophysical Research,
95, 15303–15318.
Brace, W.F. and Kohlstedt, D.L. (1980). Limits on lithospheric stress imposed
by laboratory experiments. Journal of Geophysical Research,
85, 6248–6252.
260
Brazier, R.A., Nyblade, A.A. and Florentin, J. (2005). Focal mechanisms and
the stress regime in NE and SW Tanzania, East Africa, Geophysical
Research Letters, 32, L14315, doi:10.1029/2005GL023156.
Camelbeeck, T. and Iranga, M.D. (1996). Deep crustal earthquakes and active
faults along the Rukwa trough, eastern Africa. Geophysical Journal
International, 124, 612–630.
Christensen, N. and Mooney, W. (1995). Seismic velocity structure and
composition of the continental crust: A global view. Journal of
Geophysical Research, 100, 9761–9788.
Delvaux, D. (2001). Tectonic and palaeostress evolution of the TanganyikaRukwa-Malawi rift segment, East African Rift System. In: P.A. Ziegler, W.
Cavazza and A.H.F. Robertson and S. Crasquin-Soleau (Editors), PeriTethys Memoir 6: PeriTethyan Rift /Wrench Basins and Passive Margins.
Mém. Mus. Natn. Hist. Nat., 186, 545–567. Paris.
Dreger, D.S. and Helmberger, D.V. (1990). Broadband modeling of local
earthquakes. Bulletin of the Seismological Society of America,
80, 1162–1179.
Dypvik, H., Nesteby, H., Rudent, F., Aagaard, P., Johansson, T., Msindai, J.
and Massay, C. (1990). Upper Paleozoic and Mesozoic sedimentation in
the Rukwa-Tukuyu Region, Tanzania. Journal of African Earth Sciences,
11, 437–456.
Ebinger, C.J. (1989), Tectonic development of the western branch of the East
African rift system. Geological Society of America Bulletin, 101, 885– 903.
Efron, B. and Tibshirani, R. (1991). Statistical data analysis in the computer
age. Science, 253, 390–395.
Foster, A.J. and Jackson, J.A. (1998). Source parameters of large African
earthquakes; implications for crustal rheology and regional kinematics.
Geophysical Journal International, 134, 422–448.
Holbrook, W.S., Mooney, W.D. and Christensen, N.I. (1992). The seismic
velocity structure of the deep continental crust, Chapter 1, In: D.M.
Fountain, R. Arculus and R.W. Kay (Editors), Continental Lower Crust,
Elsevier, Amsterdam, 1–43.
Jackson, J.A. and Blenkinsop, T. (1993). The Malawi earthquake of 10 March,
1989: Deep faulting within the East African rift system. Tectonics,
12, 1131–1139.
Jackson, J.A. (2002). Strength of the continental lithosphere: Time to
abandon the jelly sandwich? GSA Today, 12, 4–10.
Julià, J., Ammon, C.J. and Nyblade, A.A. (2005). Evidence for mafic lower
crust in Tanzania, east Africa, from joint inversion of receiver functions
and Rayleigh wave dispersion velocities. Geophysical Journal
Kang, T.-S. and C.-E. Baag (2004). An efficient finite-difference method for
simulating 3d seismic response of localized basin structures. Bulletin of
the Seismological Society of America, 94, 1690–1705.
Keir, D., Bastow, I.D., Whaler, K.A., Daly, E., Cornwell, D.G and Hautot, S.
(2009). Lower crustal earthquakes near the Ethiopian rift induced by
magmatic processes. Geochemistry, Geophysics, Geosystems, 10, Q0AB02,
doi:10.1029/2009GC002382.
Kennett, B.L.N. (1983). Seismic wave propagation in stratified media.
Cambridge University Press, Cambridge, 342pp.
Kilembe, E.A. and B.R. Rosendahl (1992). Structure and stratigraphy of the
Rukwa rift. Tectonophysics, 209, 143–158.
Kirby, S. (1980). Tectonic stress in the lithosphere: constraints provided by
the experimental deformation of rocks. Journal of Geophysical Research,
85, 6353–6363.
Klerkx, J., Theunissen, K. and Delvaux, D. (1998). Persistent fault controlled
basin formation since the Proterozoic along the western branch of the
East African Rift. Journal of African Earth Sciences, 26, 347–361.
Langston, C.A., Nyblade, A.A. and Owens, T.J. (2002). Regional wave
propagation in Tanzania, east Africa. Journal of Geophysical Research,
107, doi:10.1029/2001JB000167.
Last, R.J., Nyblade, A.A., Langston, C.A. and Owens, T.J. (1997). Crustal
structure of the east African plateau from receiver functions and Rayleigh
wave
phase
velocities.
Journal of Geophysical Research,
102, 24469–24484.
Lenoir, J.-L., Liégeois, J.-P., Theunissen, K. and Klerkx, J. (1994). The
Paleoproterozoic Ubendian shear belt in Tanzania: geochronology and
structure. Journal of African Earth Sciences, 19, 169–184.
Ligorria, J.P. and Ammon, C.J. (1999). Iterative deconvolution and receiverfunction estimation. Bulletin of the Seismological Society of America,
89, 1395–1400.
Mbede, E. I. (1993). Tectonic development of the Rukwa rift basin in SW
Tanzania. FU Berlin, 92 pp.
Morley, C.K., Cunningham, S.M., Harper, R.M. and Wescott, W.A. (1992).
Geology and geophysics of the Rukwa rift, east Africa. Tectonics,
11, 69–81.
Nyblade, A.A. (1997). Heat flow across the East African Plateau. Geophysical
Research Letters, 24, 2083–2086.
Nyblade, A.A., Birt, C., Langston, C.A., Owens, T.J. and Last, R.J. (1996).
Seismic experiment reveals rifting of craton in Tanzania, EOS,
Transactions American Geophysical Union, 77, 520–521.
Nyblade, A.A. and C.A. Langston (1995). East African earthquakes below 20
km depth and their implications for crustal structure. Geophysical Journal
Peirce, J.W. and Lipkov, L. (1988). Structural interpretation of the Rukwa rift,
Tanzania. Geophysics, 53, 824–836.
Prodehl, C., Keller, G.R. and Khan, M.A. (1994). Crustal and upper-mantle
structure of the Kenya rift. Tectonophysics, 236, 483.
Rasskazov S.V., Logachev N.A., Ivanov A.V., Boven A.A., Maslovskaya M.N.,
Saranina E.V., Brandt I.S. and Brandt S.B. (2003). Magmatic episode in the
western rift of East Africa (19-17 Ma). Russian Geology and Geophysics.
44, 317–324 (Translated from Geologiya i Geofizika, 2003, 44(4)).
Roberts, E.M., O’Connor, P.M., Gottfried, M.D., Stevens, N.J., Kapilima, S. and
Ngasala, S. (2004). Revised stratigraphy and age of the Red Sandstone
Group in the Rukwa Rift Basin, Tanzania. Cretaceous Research,
25, 749–759.
Rudnick, R.L. and Gao, S. (2003). Composition of the continental crust.
The Crust, Treatise in Geochemistry, 3, 1–64.
Rudnick, R.L. and Fountain, D.M. (1995). Nature and Composition of the
Continental Crust: A Lower Crustal Perspective. Reviews of Geophysics,
33, 267–309.
Shudofsky, G. (1985), Source mechanisms and focal depths of East African
earthquakes using Rayleigh-wave inversion and body wave modeling.
Geophysical Journal of the Royal Astronomical Society, 83, 563–614.
Shudofsky, G. N., Cloetingh, S., Stein, S. and Wortel, R. (1987). Unusually
deep earthquakes in East Africa: Constraints on the thermo-mechanical
structure of a continental rift system. Geophysical Research Letters,
14, 741–744.
Theunissen, K., Klerkx, J., Melnikov, A. and Mruma, A. (1996). Mechanisms
of inheritance of rift faulting in the western branch of the East African Rift,
Tanzania. Tectonics, 15, 776–790.
Wheeler, W. H. and J. A. Karson (1994). Extension and subsidence adjacent
to a “weak” continental transform: An example from the Rukwa rift, east
Africa. Geology, 22, 625.
Zandt, G. and Ammon, C.J. (1995). Continental crust composition
constrained by measurements of crustal Poisson’s ratio. Nature,
374, 152–154.
Zhao, M., Langston, C.A., Nyblade, A.A. and Owens, T.J. (1997). Lowercrustal rifting in the Rukwa graben, east Africa. Geophysical Journal
Zhu, L.P. and Kanamori, H. (2000). Moho depth variation in southern
California from teleseismic receiver functions. Journal of Geophysical
Research, 105, 2969–2980.
Editorial handling: R.J. Durrheim

crustal velocity structure of the rukwa rift in the western

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