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Seismic constraints on dynamics of the mantle of the Kaapvaal craton
PHYSICS OFTHE EARTH AND PLANETARY INTERIORS
ELSEVIER
Physics of the Earth and Planetary Interiors95 (1996) 139-151
Seismic constraints on dynamics of the mantle of the Kaapvaal craton L.P. Vinnik a,., R.W.E. Green b, L.O. Nicolaysen b a Institute of Physics of the Earth, B. Gruzinskaya 10, 123810 Moscow, Russia b Bernard Price Institute of Geophysical Research, University of the Witwatersrand, Johannesburg, PO WITS 2050, South Africa
Received 24 September 1994; accepted 31 August 1995
Abstract
Shear-wave splitting in the upper mantle of the Kaapvaal craton in southern Africa is measured by applying the SKS technique to the records of a portable array of digital broad-band seismographs. The data are interpreted in terms of lattice preferred orientation in olivine. Fast wave polarization directions at all stations of the array are close to the absolute plate motion (APM) direction of southern Africa since the end of the Jurassic. This alignment is similar to that previously reported for the North American craton and suggests that the large-scale component of mantle anisotropy on both sides of the Atlantic can be related to resistive drag at the base of the plates. A likely depth range of the corresponding deformation is 150-400km. Analysis of the mantle converted phases suggests that the high-velocity mantle root of the Kaapvaal craton, which translates coherently with the plate, resides in the same depth range: the bottom of the root is found at a depth close to 380 km. We conclude that the root is deformed by the recent plate motion, but the deformations are not strong enough to be seen in the available tomographic models.
1. Introduction Mobility of anisotropic minerals is sensitive to temperature (e.g. Nicolas and Christensen, 1987). Hence origin and age of mantle anisotropy may depend on depth. Anisotropy immediately below the Moho is probably frozen and related to processes of a distant past, when the rock within this layer was hot (temperature higher than about 1000°C). This is documented by numerous seismic refraction experiments in the oceans (e.g. Bibee and Shor, 1976; Shimamura, 1984). Deeper than about 100km, anisotropy beneath the oceans is related to the pre-
* Corresponding author,
sent-day plate tectonics; this is known from the surface-wave analyses (e.g. Montagner and Tanimoto, 1991). The frozen layer in the oceans is thin in comparison with the rest of the anisotropic mantle, and the cumulative effect of anisotropy in the teleseismic shear phases such as ScS reflects mainly processes not much older than about 15 Ma (Farra and Vinnik, 1994). Anisotropic properties of the continental mantle are less well understood, partly because the history and structure of the continental mantle are more complicated and partly because systematic seismic observations of continental anisotropy started much later, after the introduction of SKS techniques (Vinnik et al., 1984). Unfortunately, observations of splitting in SKS provide little information on the depth of
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140
L.P. Vinnik et al. / Physics of the Earth and Planetary Interiors 95 (1996) 139-151
the effect, and the data are interpreted mainly by comparing the directions of anisotropy with other observables, such as tectonic trends or absolute plate motion (APM) directions. These correlations lead to controversial interpretations. Silver and Chan (1991) claimed that the fast directions of splitting in SKS at several stations are aligned with the old crustal fabric and concluded that anisotropy in the stable continental regions has been frozen in the uppermost mantle since the last tectonic episode (the Precambrian orogenies in North America and Africa, the Hercynian in Europe). Babuska et al. (1993)argued that observations of splitting in Central Europe are related to frozen anisotropy in the subcrustal lithosphere. On the other hand, Vinnik et al. (1989, 1992) concluded that the major contribution to the observed shearwave splitting in stable continental regions, as in the oceans, comes from the asthenosphere, where anisotropy is related to the recent plate motions. This division of opinions is present in m a n y other papers, Correlation between fast direction of anisotropy and recent APM direction is especially strong in the east of North America, but it is less impressive in some other regions (Vinnik et al., 1992). In this paper we present and analyse observations of shearwave splitting at a number of seismograph stations within the Kaapvaal craton in southern Africa, where the only previously known estimate of the direction of anisotropy (Station SLR) is in a strong disagreement with the recent APM direction. Our seismic observations are carried out at seven digital broadband seismograph stations in a corridor 700km long crossing the craton in a SW-NE direction (Table 1 and Fig. 1). This portable array was fully operational
Table 1 List of portable stations with the effective values of the parameters of anisotropy: parameters of the events are given in Table 2
Station Latitude (deg) BPI -26.175 WAR -28.377 DOU -29.115 KAM -29.133
Events
a
(deg) 28.030 24.893 23.865 23.125
3,5,7, 10 2 5,6, l0 1,5
(deg) 20 20 40 20
st (s) 0.4 1.3 1.1 1.4
SAN PIL
-24.626 -25.221
27.618 27.101
5 , 6 , 7 , 9 , 11 4,5,7,8,10,11
30 50
0.7 0.6
KLI
-25.853 26.272
7, 10, II
40
1.0
Longitude
20°E
25°E
30°E
~ . ~ : , , ~ / \.~ KAAPVAAL . . . CRATON . . . "..... :l ~ ~:~s~ i ,/ ( ~\]-\ _ _ , '~/
'
~
k_
/
25°S
3oOs
~.,...,_..,.._.~-~ '
'
Fig. 1. Map of southern Africa. Boundaries of the Kaapvaal craton are adopted from De Wit et al. (1992). Direction of polarization of the fast split wave at every station is shown by a straight line; the length of the line is proportional to lapse time 8t of the slow wave. Direction of APM since the end of the Jurassic is shown by
an arrow.
from the end of 1989 until the end of 1990. The seismographs used in the experiments have been described by Green (1992).
2. Data and results
We invert the records of SKS and similar seismic phases for polarization direction of the fast split wave (ct) counted clockwise from north and time lag 8t of the slow wave, by using the algorithm by Vinnik et al. (1989). Polarization of the fast wave is roughly parallel to the [100] axis in olivine and, consequently, to the longest axis of the strain ellipsoid or to the direction of the mantle flow, if it is in the form of progressive simple shear (McKenzie, 1979; Nicolas and Christensen, 1987; Ribe, 1989, 1992). The parameters of anisotropy are found by minimizing the root mean square (r.m.s.) difference between the observed transverse component T of SKS and the transverse component Tc derived from the observed radial component (R). These estimates are related to the nearest vicinity of the station
L.P. Vinnik et a l . / Physics of the Earth and Planetary Interiors" 95 (1996) 139-151
within the first Fresnel zone. The optimal pair of the parameters is found by a grid search over the values of the parameters between 0° and 180° for t~ and between 0s and 2 s for ~t. The resulting penalty function is computed for every individual record and for the whole set of records of the seismograph station. E can be expressed as 1/2
E(ot,St)
E(a,~t)=(N_t~f[T(t)-T~(t,a,St)]2dt} i= 1
~-~ t'~'t
where N is the number of records. To distinguish between the individual estimates (N = 1) and those for the group of records, the latter will be referred to as the effective values, The accuracy of the estimates is easy to derive for an individual record, assuming that the only source of errors is additive seismic noise. For the estimates of a obtained from good-quality data (signal/noise ratio in the T component higher than three), the accuracy thus evaluated is usually within a range of a few degrees. However, the scatter of the individual estimates for events with strongly different back azimuths is usully much larger than can be explained by the additive seismic noise. This scatter is caused mainly by the Earth's complexity, and, using the effective values, one should somehow deal with the 'noise', which is not additive and the statistical properties of which are not known. The only possibility for judging this effect is to compare the estimates for events with strongly different epicentral parameters,
141
Seismic phases suitable for the analysis, SKS, SKKS and PKS, with a high signal/noise ratio are found in the records of 11 events (Table 2). At most stations, the estimates of the parameters of anisotropy are obtained by processing the records of several events with different back azimuths (Table 1). The records are individually filtered in the period range between 3 s and 15 s to attain the highest possible signal/noise ratio. Representative examples of the filtered records are shown in Fig. 2. The PKS records of Event 5 in Fig. 2(a) are interesting, because this phase is seldom used for the analysis of splitting. Large amplitudes of PKS are caused by the caustic at a distance near 128 °. The PKS record at SAN is of marginal quality, but a better quality SKKS is present on the same seismogram. The noise in the records in Fig. 2 is low, and there is a quarter-period phase shift between the R and T components, which presents a diagnostic property of azimuthal anisotropy. The effective penalty functions, obtained by combined processing of all available records, are shown in Fig. 3. To achieve a better understanding of the quality and implications of the results obtained, we will present and discuss some data in detail. Fig. 4 shows examples of the penalty functions for Station PIL. In the record of Event 4, with a back azimuth of 50 °, there is no signal in the T component. This indicates that the axis of symmetry beneath the station is either parallel or perpendicular to this direction. The minima of the respective penalty functions in Fig. 4 are observed at azimuths around
Table 2 List of events: back azimuths and distances are given with respect to BPI No.
Date
Latitude (deg)
Longitude (deg)
Depth (kin)
Back azimuth (deg)
Distance (deg)
Phase
1 2 3 4 5 6 7 8 9 10 11
34389:09 12 89 34989:15 12 89 10890:18 04 90 11690:26 04 90 13290:12 05 90 15090:30 05 90 16590:14 06 90 26090:17 09 90 28390:10 10 90 29090:17 10 90 36490:30 12 90
0.1N 8.3N 1.2N 36.2N 49.0N 6.0S 47.9N 53.2E 19.5S 11.0S 6.9S
123.3E 126.7E 122.9E 100.3E 141.9E 77.2W 85.1E 159.6E 66.6W 70.8W 151.0E
151 24 26 10 606 24 58 10 127 599 179
92 86 91 50 47 258 34 153 251 256 111
94.8 101.4 94.9 92.1 124.1 100.9 89.7 90.5 85.6 93.0 116.6
SKS SKS SKS SKS PKS, SKKS SKS SKS SKS SKS SKS SKS, SKKS
142
L.P. Vinnik et a l . / Physics (dthe Earth and Planetary hueriors 95 (1996) I39-151
(al
' j SAN
~
I
~ BP1
.
~
,
~ DOU
~
l l ' l i , ~ KaM
,
0.0
(
20.0
~
400
~
60 0
time(s)
| BPI
•,
.
. -.,
', ,
PIL
~
,
,~,. KLI
~
~
50 ° and 140°, and there are no estimates of ~t. A similar solution is obtained for SKKS of Event 5, with a back azimuth of 47 ° (not shown in Fig. 4). For the record of Event 7, with a back azimuth of 34°, there are also two minima at the same azimuths, with the minimum at 140° being slightly deeper than the other one (Fig. 4). The solutions for Events 4, 5 and 7, whose back azimuths are close to each other, are, clearly, very similar; practically the only source of scatter of the estimates within this subset of data is the additive seismic noise in the T component, and its disturbing contribution to the estimates of fast azimuth is less than 10°. A comparable solution is also obtained for Event 8, the back azimuth direction of which (153 °) is roughly perpendicular to those of Events 4 and 5. In the SKS record of Event 10, the signal in the T component is clearly seen (Fig. 2(b)), and a record of comparable quality is obtained for SKKS of Event 11. Back azimuths of these events are 256 ° and 111 °. The minima of the respective penalty functions are observed at 30° and 0.5 s, and at 60 ° and 0.7 s (Fig. 4). The difference of 30° between the two solutions is much larger than the effect of the additive seismic noise. Discrepancies of such magnitude are reported for many stations in other regions (e.g. Vinnik et al., 1994), and should not be considered as something unusual. The effective estimates of the parameters for the whole set of records of PIL are 50° and 0.6 s (Fig. 3), which differ by up to 20 ° and 0.1 s from the individual estimates. A formal evaluation of the accuracy of the effective estimates is difficult because the main source of 'noise' in this case is complexity of the Earth's medium, and the properties of this noise are hard to quantify. We may assume, however, that the likely error is within the range of scatter of the individual estimates, or less than about 20 ° for the azimuth and 0.2 s for ~t. The data for BPI are comparable in signal/noise ratio with those for PIL. For the PKS record of Event 5 (Fig. 2(a)), there is a region of small values of the
DOU
O0
~
200
~
40.0
~ 60 0 time(s)
Fig. 2. Examples of the R (continuous lines) and T component records (dashed lines): (a) PKS phase of Event 5; (b) SKS phase of Event 10. Vertical lines indicate time windows used to invert the data for the parameters of splitting.
L.P. Vinnik et al./ Physics of the Earth and Planetary Interiors 95 (1996) 139-151
penalty function between 150° and 40°, with the minimum at 170° and 0.4s (Fig. 5). In the record of Event 7, the signal in the T component is practically missing; in the resulting penalty function there are two minima, at about 400 and 130° (Fig. 5). The other record with a missing signal in the T component is Event 3, with a back azimuth of 91 °. The minima for this record are around 10° and 100°. For Event 10 (Fig. 2(b)), there is a minimum of the penalty function at around I0 ° and 0.7s (Fig. 5). It is clear that, in spite of scatter of the individual estimates, the regions of the smallest values of the penalty functions overlap in the vicinity of 20 ° and 0.4 s, and the effective solution (Fig. 3) is close to these values. The accuracy of the effective values for BPI is comparable with that for PIL, and results of comparable quality are obtained for KLI and KAM. The results for DOU (Fig. 6) deserve special discussion because of a peculiar effect in St. In all three records of this station (Events 5, 6 and 10), the (a)
BPI
DOU
signal in the T component is well seen; two of them (Events 5 and 10) are shown in Fig. 2. The respectire penalty functions (Fig. 6) have clear minima at 40 ° and 1.8 s, 60 ° and 2 s, and 60 ° and 1.6 s. Back azimuths of Events 6 and 10 differ by only 2 °, and their solutions demonstrate once more that the influence of the additive seismic noise upon the estimate of fast azimuth is within a range of a few degrees. The influence upon the estimate of 8t in this case is around 0.4s. The effective estimate of ~t (Fig. 3) is 1. l s, which is lower than each of the individual estimates, and this bias should not be ignored when the effective estimates are interpreted. The effective estimate of the azimuth is 40°, which coincides with the solution for Event 5. The data for SAN present a very singular character because of the unusually large scatter of the individual estimates. For the PKS record of Event 5, shown in Fig. 2(a), the solution (Fig. 7) is 140° and 2 s. This record is of marginal quality, but a similar (b)
oI ......... oi o I
180. O.
DELRY (S)
2.
180. I O.
KAN
~ 2.
I8
, O.
18 DELAY {S)
2.
O.
DELRY ISi
2.
NRR
{ ~("~'~(, '~'~'
_
5AN
. . . . . . . . . .
KL I
_
o12:
PIL
o
, , DELR" iS)
143
-
%. 180.
18o.
O.
DELR'i' IS)
2.
, ,
O.
I8O.
~LR'i" IS)
2.
O.
DELAY IS}
2.
Fig. 3. Plots of the penalty function E(oL,St) obtained by combined processing of all available records and used to find effective values of the parameters.
144
L.P. Vinnik et al. / Physics of the Earth and Planetary Interiors 95 (1996) 139-151
solution is obtained for the record of Event 7, the back azimuth for which differs little from that of Event 5 (Fig. 7). A comparable solution is obtained also for the SKKS record of Event 5. For the record of Event 6, with a back azimuth near 258° and a high signal/noise ratio in both components, the solution is around 40° and 0-5 s (Fig- 7); a similar solution is obtained for the record of Event 9, with almost the same back azimuth. The solution for the well-recorded SKS of Event 11, with a back azimuth of 111° (80° and 1.6 s; Fig. 7), is very different from the others, and a close solution (100 ° and 2 s) is obtained for SKKS of the same event. Remarkably, in spite of the enormous scatter of the individual estimates, the effective estimates for SAN (30° and 0.7 s; Fig. 3) are close to those for the neighbouring Stations PIL and BPI. To summarize, the effective estimates of the fast
i
1
k
i
~
I
I
J
i
i
i
i
i
i
BP; #5
BPI #7 ' ........ ~ ~~--___~
0.
0.
~
~ l ( / ~
~
~ l ' ~ ~ ?-
,
, ,', , ~(~/~,f~,/f' f , ~ 180.
0.
O.
DELRT (51
BP; #10 ' ........ / ~
2.
180./
O.
DELflT [Sl
2.
-i
i
t ff-o.~o~,
180 . . . . . . . . . . O. DELl:IT IS)
--
1 2.
Fig. 5. The same as Fig. 4 but for BPI.
~: ( 180.
~
~ 180.
,
BELl:ITIS1 PIL #10
O. ~
' ...... ~
2. _
. DELAftS~ PIL #1t 0. J ' ~ ~ 1 ~ ~ O.
2.
~ Jt / . " ° f ~~ -~
~
~
1
. ~ -
~
]80.
3. S-velocity structure o f the m a n t l e beneath the Kaapvaal craton
18o. I . . . . . . . . . O.
~LR'f IS1
2.
O.
OELfff ISI
2.
Fig. 4. Examples of plots of the penalty function E(t~,St) for the
records of PIL.
direction for the stations of the network are in the interval between 20° and 50°' with a mean value near 30°. The differences between the effective estimates of a for the various stations are generally within the range of the accuracy of measurement. The estimates of Bt, considering both the individual and effective estimates, are about two times lower in the north-east (BPI, SAN and PIL) than in the south-west (DOU, KAM and WAR). SAN, the station with anomalously strong variability of the individual estimates, is very close to the normal Station PIL; hence, the source of the anomaly is very shallow, perhaps crustal.
To understand the splitting observations, additional data on the deep structure of the study region
L.P. Vinnik et al. / Physics of the Earth and Planetary Interiors' 95 (1996) 139-151
are required. Generally, the upper mantle of Precambrian platforms is characterized by higher than average P and S velocities (e.g. Suet al., 1994). The high velocities can be explained by the presence of old mantle roots, which may differ from the normal mantle in chemical composition; the roots, a few hundred kilometres deep, translate coherently with the plates (Jordan, 1988). The maximum depths of the roots are not known, owing to the limited resolution of the available seismic data. For the study region, high S velocities in the upper mantle are confirmed by the multimode surface-wave dispersion data (Bloch et al., 1969). More recently, the records of the array shown in Fig. 1 have been used to infer a detailed S-velocity profile of the upper mantle beneath the Kaapvaal craton. Here we present our preliminary results, and a comprehensive report will be published elsewhere (Vinnik et al., 1996); our preliminary results are not much different from the final ones.
145
SnN#s
SAN#6
i ..... :' '~L o., ~ ~ l ~ ~......... vq~ i
~
~
1
~
~
~ ~~,~.~ q \ ~'-..~o~"--~_
~
180. . . . . . . . . . . 180. 0. t~LAY~SJ 2. 0. t~LnY tS~ 2. SRN#~ SAN#11 0. ~ ' ' ~ ' \,.,~., 0. e'0 " ~ .~'/"---~--~-~ ,
~
m
Fig. 7. The same as Fig. 4 but for SAN.
°I ]80 . . . . . . . . . . . 0. Itl.flY (SI o0u #10 0. ' . . . . . . . .
.......
100. 2.
0.
OELIq~"~SJ
"~ _
o o . 1 ~ ~ , , 180 ! 0.
0ELAYtS~
2.
Fig. 6. The same as Fig. 4 but for DOU.
2.
S-Velocity gradients in the mantle beneath the Kaapvaal craton have been inferred from the observations of the teleseismic P-to-SV converted phases. These phases are detected by using a technique (Vinnik, 1977) that consists of source equalization (deconvolution), rotation of the seismogram into the L and H coordinate frame and stacking the H component records of many events with appropriate moveout corrections. The L and H directions coincide with the directions of the principal motion components in the P wave and the converted phases, respectively. Standardization is performed by using a filtering technique which makes the P waveform look like a 8 function. Details of the techniques and the results obtained elsewhere, have been described by, for example, Kind and Vinnik (1988), Stammler et al. (1992) and Petersen et al. (1993). We processed the records of 6, 8, 6, 8 and 10 events in the distance range 30-90 ° for Stations BPI, DOU, SAN, PIL and KLI, respectively. To improve
146
L.P, Vinnik et al. / Physics of the Earth and Planetary Interiors 95 (1996) 139-151 PIL+KLI+DOU
800- ~ 750- ~--~-~ ~ 640--~'~-/"
I ~
I
~
560480- ! a-400-
350240-
t6080-
0' T' ' ' I , , , I ' ' ' I ' ' ' I ' ' ' t . . . .
-20
0
20 40 Time (s)
60
BO
Fig. 8. Result of stacking the H components of Stations DOU, KLI and PIE Moveout corrections for stacking were calculated for the trial depths of conversion, indicated on the left. Time is shown relative to the P wave at 0s. Times of the detected mantle phases are shown by marks at the top.
the signal/noise ratio, the results of the stacking were additionally low-pass filtered with a cutoff frequency between 0.1 and 0.2 Hz. Fig. 8 shows the result of stacking for DOU, PIL and KLI; comparable results are obtained by stacking the records of all stations. The phases between 0 and 30s are formed by conversions and multiple reflections from the Moho and free surface. The phase with a time (lapse time relative to P) of 66.0s is related to the 6 5 0 k m discontinuity. A remarkable result of data processing is the phase with a downward motion at 38 s. Such motion is possible if the conversion occurs at the upper boundary of a low-velocity layer. This phase is detected reliably, because its amplitude (amplitude ratio between the converted phase and the P wave) is 0.06, which is much higher than the level of the preceding noise and the standard error of the stack (0.007); moreover, its amplitude is largest at trial depths around 400km, thus indicating that its pattern of amplitude variation vs. trial depth is in agreement with the depth of conversion that can be inferred from its time. The last criterion is especially important because it provides evidence that the detected phase is a true converted phase rather than an artefact of data processing. Additionally, to test the
stability of the observed wavefieid, the records were divided into subgroups, with roughly the same amount of records in each subgroup, and these subgroups were processed separately. A large downward motion at 38 s has been detected in every subgroup. The time of the 6 5 0 k m phase (66.0s) is 2 s less than for the standard Earth model IASP91 (Kennett and Engdahl, 1991). On a global scale, the travel time of the 6 5 0 k m phase in the mantle transition zone is stable, and the variations observed at the Earth's surface are controlled mainly by the lateral velocity variations within the upper, most heterogeneous zone of the Earth, of 400 km thickness (Stammler et al., 1992). The observed travel time residual of the converted phase is practically the difference between the teleseismic S and P residuals that are accumulated in that zone. Assuming, in agreement published data (e.g. Romanowicz and Cara, 1980), that the ratio between the S and P residuals is close to three, the teleseismic S residual in the study region can be estimated as - 3 s. This estimate is in good quantitative agreement with the data for the other Precambrian shields (e.g. Su et al., 1994). The S-velocity model SVKC explaining the waveform of the 38 s conversion (Fig. 9) is found by a trial-and-error method. The theoretical seismograms for the trial models are computed in the 3 0 - 9 0 ° with
5.5 5.4
53 ~, 5.2 "~ 51
IASP91,-'""
~ 50 ~ 4.9 ~ 48
'-4 ~
~
4..7
46
,
4.5 '
360
i
4.60 . .440. . . 480 o~prH ( ~ l
5~,o '
Fig. 9. SVKC S-velocity model inferred from the data shown in Fig. 8 (continuous line) and standard model 1ASP91 (Kennett and
Engdahl,1991).
L.P. Vinnik et al./ Physics of the Earth and Planetary Interiors 95 (1996) 139-151
distance range by using the reflectivity technique (Fuchs and Mueller, 1971) and processed like the observed seismograms. The P waveform found by processing the observed seismograms, with a correction for mantle attenuation, was used as the input for computing theoretical seismograms. The Q values were taken from EURSQ model (Der et al., 1986), which is based on seismic data for the Precambrian platforms. To account for the presence of the highvelocity mantle root of the craton, the S and P velocities in the mantle at depths less than 380 km are increased relative to IASP91 by 0.1 km s-~ and 0.15 km s -~, respectively. The depth and S-velocity contrast of the 380km discontinuity in our model are constrained by the time and amplitude of the 38 s phase, respectively. The discontinuity is made sharp for reasons of simplicity. The adopted thickness of the layer where S velocity is lower than in IASP91 is, practically, the smallest acceptable thickness. If it is made substantially smaller, a strong converted phase at about 43 s which is missing in the observations appears in the theoretical seismograms. SVKC provides a very good fit between the observed and synthetic waveforms of the 38 s phase, Any alternative explanation of the 38 s signal is difficult. It could be assumed, for example, that the waveform of the Ps phase is distorted by a multiple reflection of the type Ps(d)s (P converted at the surface to S and then reflected once from an interface at depth d). By appropriate selection of d, the times of the multiples and the 400 km Ps phase at a certain epicentral distance can be equalized, but the moveout corrections which are needed for detecting the multiples and the Ps phases are very different. The delay of Ps relative to P grows with increasing ray parameter or decreasing epicentral distance; the dependence of the delay of the multiples is opposite (Vinnik and Mikhailova, 1988). The dependence of the amplitude on trial depth in Fig. 8 precludes the explanation of the 38 s phase in terms of the multipies: if this phase was produced by the multiples, the largest amplitudes were obtained at zero trial depth, as for the crustal multiples in Fig. 8. It could also be proposed that the anomalous phase is converted at the normal 400 km discontinuity, but the waveform is distorted by topography of the discontinuity. There are a few objections to this interpretation. The normal time delay of the 650 km
147
phase relative to the 400km phase is very stable and close to 24s (Stammler et al., 1992). This means that, in the study region, the 400km phase is expected to arrive at around 42 s. A lowered S velocity beneath the 400km discontinuity may change this estimate by only a fraction of a second. The anomalous phase arrives a few seconds earlier, which means that the related discontinuity is displaced upward by a few tens of kilometres relative to the standard depth of the 400km discontinuity. If the 400 km discontinuity is interpreted conventionally as the olivine-spinel phase transformation with a positive Clapeyron slope, to provide this uplift the temperature should be lowered by a few hundred degrees, which is out of the question for the study region. The topography would lower the amplitude of the detected signal, because the signal coherency, which is a prerequisite for our signal detection techniques, would deteriorate. In reality, the amplitude of the downward motion at 38 s is twice as large as the normal amplitude of the standard 400 km Ps phase. Finally, we do not think that any topography can transform the upward motion of the 400 km Ps wave into the downward motion in Fig. 8; there is a phase shift in the Sp waves for some angles of incidence (Aki and Richards, 1980), but this is not applicable to the Ps waves. Similar analyses are known for a number ofcontinental stations (e.g. Stammler et al., 1992), but nothing comparable with that shown in Fig. 8 has ever been seen. Numerous S-velocity models derived from various kinds of seismic data for many regions usually show that in the interval between 300 and 650km, the P and S velocities rise with depth. Some complications may exist in the vicinity of the currently active subduction zones, but far outside subduction zones the low S velocity at depths between 350 and 500km should be viewed as a strong anomaly. The mantle beneath Southern Africa is anomalously heated (Nyblade and Robinson, 1994), but within the Kaapvaal craton the heat flow data give no evidence of a thermal perturbation (Jones, 1988); this can be explained by a large thickness of the lithosphere of the craton. With the indications of heating taken into account, the low-velocity layer in the SVKC model can be explained by elevated ternperature that comes close to the solidus temperature.
148
L.P. Vinnik et al. / Physics of the Earth and Planetary Interiors 95 (1996) 139-151
The upper part of the low-velocity layer may correspond to 'anticrust', the layer enriched by, among other components, volatiles (Gasparik, 1992). The temperature of the mantle solidus is lowered by the volatiles (Thompson, 1992). The 380kin discontinuity may separate the root of the craton or 'tectosphere' from the underlying mantle; the S velocity above the discontinuity is higher, because the root may have a different composition and a higher solidus temperature (Jordan, 1988). To conclude this section, the seismic data give reasons to believe that the mantle root of the Kaapvaal craton comes close to a 400kin depth. Temperatures in the lower part of the root are well in excess of 1300°C, and the root should not be mistaken for the lithosphere,
4. Discussion The fast directions in Fig. 1 are fairly uniform, Has this uniform direction been frozen in the uppermost mantle since the Archaean, or does it reflect a process which operated much later at greater depths? Observations of splitting within the currently active tectonic zones (collisional belts, zones of rifling) indicate that the fast directions of mantle anisotropy in the upper mantle are often either parallel or perpendicular to the strikes of these zones (collisional belts and riffs, respectively), but there are notable exceptions to these regularities (e.g. Vinnik et al., 1992). One may speculate that the tectonic processes of a distant geologic past are very similar to those active at present, and that the anisotropy thus created can be preserved in the mantle for a billion years or so, but to become credible, these ideas must be confirmed by seismic observations sensitive to depth of anisotropy. There are a few currently stable regions where such observations are available, namely, the east of North America, southem Germany and the Iberian peninsula, In North America, there is a clear trend in the fast directions of SKS, which coincides with the recent absolute plate motion direction (Vinnik et al., 1992). One can see this trend also in more recent observations (Silver and Kaneshima, 1993). Moreover, seismic long-range refraction observations (Iyer et al., 1969) do not show any anisotropy in the upper
100 km of the mantle where frozen anisotropy should reside; a detailed discussion of the implications of the long-range profiling data in North America has been given elsewhere (Vinnik et al., 1992). In southern Germany, seismic refraction observations (Bamford, 1977)reveal frozen anisotropy whose direction has nothing in common with either the Hercynian crustal fabric or the fast directions in SKS; additional data confirming this direction have been given by Farra et al. (1991). The observations of SKS at some stations in southern Germany can be inverted for two-layer anisotropic structures. In the resulting models (Vinnik et al., 1984), the anisotropic layer in the subcrustal lithosphere is relatively thin, and ~t in the asthenosphere is twice as large as in the subcrustal lithosphere. In the Iberian peninsula, as shown by seismic refraction experiments, the subcrustal lithosphere at depths less than 100km includes three thin anisotropic layers with the fast direction around NNE (Diaz et al., 1993). This direction has nothing to do with either the fast direction of anisotropy in SKS or the Hercynian tectonics. Moreover, longperiod surface-wave data which are sensitive to depth of anisotropy indicate that anisotropy beneath the Iberian plate resides at depths exceeding 100km (Maupin and Cara, 1992). The layers observed by Diaz et al. (1993) are too thin, and anisotropy within them is too weak to affect the observations of SKS and long-period surface waves. All observations (SKS, surface waves, P-wave refraction) are fully compatible, if the effect in SKS is accumulated mainly in the asthenosphere, where it is related to the relatively recent (most likely, Cretaceous) motion of the Iberian plate. To conclude, when seismic data sensitive to depth are available, they strongly suggest that (1)the dominant effect in SKS comes from the asthenosphere, and (2) the fast direction of frozen anisotropy in the subcrustal lithosphere differs from the fast direction in SKS. The early observations of splitting in SKS in southern Africa (Station SLR, a around 80 °) were interpreted by Silver and Chan (1991) in terms of frozen anisotropy in the subcrustai lithosphere. The reason for this interpretation is a similarity between the fast direction at SLR and the strike of the Limpopo belt, a feature several hundred kilometres to the north of the station. This interpretation is incompatible with the data, available at present (Fig.
L.P. Vinnik et a l . / Physics of the Earth and Planetary Interiors 95 (1996) 139-151
1). There are Proterozoic mobile belts with a suitable orientation beyond the north-western border of the Kalahari craton (Hartnady et al., 1985), but on tectonic maps at such distances (about 1000 km from the observation site) one can find not only this but also many different directions. On the other hand, the direction of anisotropy at our stations practically coincides with the APM direction of southern Africa in the hotspot reference frame from the end of the Jurassic to the present (Dietz and Holden, 1970). As shown by the trails of the hotspots in the Southern Atlantic, during this time the plate moved by approximately 3000kin. Approximately 25Ma ago Africa halted (Burke and Wilson, 1972), but apparently the accumulated deformation is still preserved. Within the North American craton, the fast direction of anisotropy is close to the present-day APM direction. This APM direction is related to the opening of the North Atlantic and has been preserved without a significant change since the end of the Cretaceous, 65Ma ago (Dietz and Holden, 1970). The relationship between the fast direction of anisotropy and the relatively recent APM direction in southern Africa is, clearly, very similar to that in North America, and suggests the same origin of the large-scale component of mantle anisotropy in both regions: resistive drag at the base of the moving plate. In that case, the deformation near the base of the plate is in the form of progressive simple shear which yields orientation of the [100] axis of olivine in the direction of plate motion (Ribe, 1989). Lattice preferred orientation (LPO) of olivine is sensitive to deformation only at temperatures higher than about 1000°C (Nicolas and Christensen, 1987). A temperature of 1000°C underneath the Kaapvaai craton, as deduced from the mineral equilibrium studies of kimberlite inclusions, is reached at depths around 150kin (Boyd and Gurney, 1986). This estimate is remarkably consistent with the geotherm calculated from heat flow data; the present-day geotherm differs little from that in the Cretaceous (Jones, 1988). The properties of minerals at depths exceeding 400km seem to be unfavourable for developing significant anisotropy (Mainprice and Silver, 1993). Hence, anisotropy created during the last 150Ma in the mantle of the Kaapvaal craton is probably located between 150 and 400km, in the depth range occupied by the deep continental root of
149
the craton. To obtain ~t of 1 s in a layer of 250km thickness, the required difference between the split wave velocities is 2%. This difference is about onethird of the maximum values found in samples of the upper-mantle rocks. The indications of relatively young deformations of the deep continental root are not necessarily at odds with seismic evidence of its coherent motion with the plate. Significant LPO develops by intracrystalline slip when ln(cl/c 2) or ln(c2/c 3) (where c I, c 2 and c 3 are respectively the largest, intermediate and smallest axes of the strain ellipsoid), is around 0.3 (Ribe, 1992). The equivalent shear strain y is close to unity. These theoretical predictions can be complemented by observational data: experiments on simple shear in ice (Bouchez and Duval, 1982) demonstrated that a good fabric was induced for a y of only 0.7. This range of strains is reached when the root at a depth of 400km is displaced laterally by less than 200kin with respect to a 150km depth. A shift of such magnitude at a 400 km depth is difficult to detect in the currently available tomographic models (e.g. S u e t al., 1994). The estimates of ~t are relatively low for the north-eastern stations and relatively high in the south-west. The number of Jurassic-Cretaceous kimberlite intrusives within our corridor is much larger in the south-west than in the north-east (Skinner, 1989). This correlation suggests a relationship between kimberlites and mantle anisotropy. We speculate that both phenomena could be affected by the influx of volatiles: volatiles are abundant in the kimberlite magma, and they could intensify recrystallization of olivine. At temperatures higher than about 1200°C, recrystallization takes place on grain boundaries and, in the presence of strain, increases anisotropy (D. Mainprice, personal communication, 1993). To conclude, the large-scale component of mantle anisotropy (scale on the order of 1000km) within the Kaapvaal craton and similar regions of North America can most easily be explained by the relatively recent (ages around 50-100Ma) deformations in the lower (deeper than about 150km) part of the continental root. The temperatures at 200-400km depths beneath the Kaapvaal craton are in the range which is commonly ascribed to the asthenosphere, and it should not be regarded as a great surprise that the
150
L.P. Vinnik et al.// Physics of the Earth and Planetary Interiors 95 (1996) 139-151
mantle at these depths is d e f o r m e d by shearing, W h a t is surprising is the w e a k n e s s o f the d e f o r m a -
crust and upper mantle of southern Africa from multi-mode surface wave dispersion. Bull. Seismol. Soc. Am., 59: 1599-
tions. A contribution to the o b s e r v e d s h e a r - w a v e
1630. Bouchez, J.L. and Dural, P., 1982. The fabric of polycrystalline ice deformed in simple shear: experiments in torsion, natural deformation and geometrical interpretation. Text. Microstruct., 5: 1-17. Boyd, F.R. and Gumey, J.J., 1986. Diamonds and the African lithosphere. Science, 232: 472-477. Burke, K. and Wilson, J.T., 1972. Is the African plate stationary? Nature, 239: 387-390. Der, Z.A., Lees, A. and Cormier, V.F., 1986. Frequency dependence of Q in the mantle underlying the shield areas of Eurasia, Part 3: The Q model. Geophys. J. R. Astron. Soc., 87:
splitting f r o m e v e n greater depths cannot be excluded, although with the data available at present this does not look very likely. W e do not want to say that older deformations are absent in the subcratonic upper mantle. S a m p l e s o f rocks f r o m the South African subcrustal lithosphere are w e a k l y anisotropic (Mainprice and Silver, 1993), and this anisotropy could contribute to the small-scale lateral variations o f the parameters of splitting. Unfortunately, interpretation o f small-scale and azimuthal variations o f the inferred parameters o f anisotropy is difficult, because there are too m a n y possible reasons for them. W a v e scattering at small-scale lateral i n h o m o geneities within the crust c o u l d be one o f these reasons. S o m e effects o f layered anisotropic structure in seismic w a v e f i e l d s can be used to infer the eters o f anisotropy in the layers (e.g. Farra 1991; Vinnik et al., 1994), but either these are w e a k at our stations or the a m o u n t o f
paramet al., effects data is
insufficient to identify them.
Acknowledgements The authors are m u c h indebted for h e l p to E.O. Kostlin, L.I. M a k e y e v a , R. K i n d and G. Asch. W e thank an a n o n y m o u s r e v i e w e r and M. G r a n e t for helpful reviews. Part o f this research was carried out in the S e i s m o l o g i s c h e s Z e n t r a l o b s e r v a t o r i u m o f G e r m a n y (Erlangen), where one o f the authors (L.V.) was supported by the A l e x a n d e r - y o n - H u m b o l d t -
Stiftung.
References Aid, K. and Richards, P.G., 1980. Quantitative Seismology, Vol. 1. W.H. Freeman, San Francisco, CA, 557 pp. Babuska, V., Plomerova, J. and Sileny, J., 1993. Models of seismic anisotropy in the deep continental lithosphere. Phys. Earth Planet. Inter., 78: 167-192. Bamford, D., 1977. Pn velocity anisotropy in a continental upper mantle. Geophys. J. R. Astron. Soc., 49: 29-48. Bibee, L. and Shor, Jr., G., 1976. Compressional wave anisotropy in the crust and upper mantle. Geophys. Res. Lett., 3: 639-642. Bloch, S., Hales, A.L. and Landisman, M., 1969. Velocities in the
1103-1112. De Wit, M.J., Roering, Ch., Hart, R.J., Armstrong, R.A., de Ronde, C.E.J., Green, R.W.E., Tredoux, M., Peberdy, E. and Hart, R.A., 1992. Formation of an African continent. Nature, 357: 553-562. Diaz, J., Him, A., Gallart, J. and Senos, L., 1993. Evidence for azimuthal anisotropy in southwest Iberia from deep seismic sounding data. Phys. Earth Planet. Inter., 78: 193-206. Dietz, R.S. and Holden, J.C., 1970. Reconstruction of Pangaea: breakup and dispersion of continents, Permian to present. J. Geophys. Res., 75: 4939-4956. Farra, V. and Vinnik, L., 1994. Shear-wave splitting in the mantle of the Pacific. Geophys. J. Int., 119: 195-218. Farra, V., Vinnik, L.P., Romanowicz, B., Kosarev, G.L. and Kind, R., 1991, Inversion of teleseismic S particle motion for azimuthal anisotropy in the upper mantle: a feasibility study. Geophys. J. Int., 106: 421-431. Fuchs, K. and Mueller, G., 1971. Computation of synthetic seismograms with the reflectivity method and comparison with observations. Geophys. J., 23: 417-433. Gasparik, T., 1992. Enstatite-jadeite join and its role in the Earth's mantle. Contrib. Mineral. Petrol., 111: 283-298. Green, R.W.E., 1992. PC-based continuous recording field seismograph. Bull. Seismol. Soc. Am., 82: 497-504. Hartnady, C., Joubert, P. and Stowe, C., 1985. Proterozoic crustal evolution in south-western Africa. Episodes, 8: 236-244. Iyer, H.M., Pakiser, L.C., Stuart, D.J. and Warren, D.H., 1969. Project Early Rise: seismic probing of the upper mantle. J. Geophys. Res., 74: 4409-4441. Jones, M.Q.W., 1988. Heat flow in the Witwatersrand Basin and environs and its significance for the South African geotherm and lithosphere thickness. J. Geophys. Res., 93: 3243-3260. Jordan, T.H,, 1988. Structure and formation of the continental tectosphere. J. Petrol., special lithosphere issue, 11-37. Kennett, B.L.N. and Engdahl, E.R., 1991. Traveltimes for global earthquake location and phase identification. Geophys. J. Int., 105: 429-465. Kind, R. and Vinnik, L.P., 1988. The upper mantle discontinuities underneath the GRF array from P-to-S converted phases. J. Geophys., 62: 138-147. Mainprice, D. and Silver, P.G., 1993. Interpretation of SKS waves using samples from the subcontinentai lithosphere. Phys. Earth Planet. Inter., 78: 257-280.
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Maupin, V. and Cam, M., 1992. Love-Rayleigh wave incompatibility and possible deep upper mantle anisotropy in the Iberian peninsula. Pageophys., 138: 429-444. McKenzie, D., 1979. Finite deformations during fluid flow. Geophys. J. R. Astron. Soc., 58: 689-715. Montagner, J.P. and Tanimoto, T., 1991. Global upper mantle tomography of seismic velocities and anisotropies. J. Geophys. Res., 96: 20337-20351. Nicolas, A. and Christensen, N.I., 1987. Formation of anisotropy in upper mantle peridotites--a review. Am. Geophys. Union Geodyn. Ser., 16:111-123. Nyblade, A.A. and Robinson, S.W., 1994. The African Superswell. Geophys. Res. Lett., 21: 765-768. Petersen, N., Vinnik, L., Kosarev, G., Kind, R., Oreshin, S. and Stammler, K., 1993. Sharpness of the mantle discontinuities, Geophys. Res. Lett., 20: 859-862. Ribe, N.M., 1989. Seismic anisotropy and mantle flow. J. Geophys. Res., 94: 4213-4223. Ribe, N.M., 1992. On the relation between seismic anisotropy and finite strain. J. Geophys. Res., 97: 8737-8747. Romanowicz, B.A. and Cara, M., 1980. Reconsideration of the relations between S and P station anomalies in North America. Geophys. Res. Lett., 7: 417-420. Shimamura, H., 1984. Anisotropy in the oceanic lithosphere of the northwestem Pacific basin. Geophys. J. R. Astron. Soc., 76: 253-260. Silver, P.G. and Chan, W.W., 1991. Shear wave splitting and subcontinental mantle deformation. J. Geophys. Res., 96: 16429-16454. Silver, P.G. and Kaneshima, S., 1993. Constraints on mantle anisotropy beneath Precambrian North America from a transportable teleseismic experiment. Geophys. Res. Lett., 20: 1127-1130. Skinner, E.M.W., 1989. Contrasting Group 1 and Group 2 kimber-
151
lite petrology: towards a genetic model for kimberlites. In: J. Ross et al. (Editors), Kimberlites and Related Rocks 1. Blackwell, Melbourne, Vic., pp. 528-544. Stammler, K., Kind, R., Petersen, N., Kosarev, G., Vinnik, L. and Liu, Q., 1992. The upper mantle discontinuities: correlated or anticorrelated? Geophys. Res. Lea., 19: 1563-1566. Su, W.-J., Woodward, R.L. and Dziewonski, A.M., 1994. Degree 12 model of shear velocity heterogeneity in the mantle. J. Geophys. Res., 99: 6945-6980. Thompson, A.B., 1992. Water in the Earth's mantle. Nature, 358: 295-302. Vinnik, L.P., 1977. Detection of waves converted from P to SV in the mantle. Phys. Earth Planet. Inter., 15: 39-45. Vinnik, L.P. and Mikhailova, N.G., 1988. Structure of the mantle transition zone by the observations of the Ps converted phases. In: I.L. Nersesov (Editor), Structure and Dynamics of the Lithosphere from Seismic Data. Nauka, Moscow, pp. 19-61. Vinnik, L.P., Kosarev, G.L. and Makeyeva, L.I., 1984. Anisotropy in the lithosphere from the observations of SKS and SKKS (in Russian). Dokl. Akad. Nauk SSSR, 278: 1335-1339. Vinnik, L.P., Farra, V. and Romanowicz, B., 1989. Azimuthal anisotropy in the Earth from observations of SKS at Geoscope and NARS broadband stations. Bull. Seismol. Soc. Am., 79: 1542-1558. Vinnik, L.P., Makeyeva, L.I., Milev, A. and Usenko, A.Yu., 1992. Global patterns of azimuthal anisotropy and deformations in the continental mantle. Geophys. J. Int., 111: 433-447. Vinnik, L.P., Krishna, V.G., Kind, R., Bormann, P. and Stammler, K., 1994. Shear-wave splitting in the records of the German Regional Seismic Network. Geophys. Res. Lett., 2t: 457-460. Virmik, L.P., Green, L.O., Nicolaysen, Kosarev, G.L. and Petersen, N.V., 1996. Deep seismic structure of the Kaapvaal craton. Tectonophysics, in press.
ELSEVIER
Physics of the Earth and Planetary Interiors95 (1996) 139-151
Seismic constraints on dynamics of the mantle of the Kaapvaal craton L.P. Vinnik a,., R.W.E. Green b, L.O. Nicolaysen b a Institute of Physics of the Earth, B. Gruzinskaya 10, 123810 Moscow, Russia b Bernard Price Institute of Geophysical Research, University of the Witwatersrand, Johannesburg, PO WITS 2050, South Africa
Received 24 September 1994; accepted 31 August 1995
Abstract
Shear-wave splitting in the upper mantle of the Kaapvaal craton in southern Africa is measured by applying the SKS technique to the records of a portable array of digital broad-band seismographs. The data are interpreted in terms of lattice preferred orientation in olivine. Fast wave polarization directions at all stations of the array are close to the absolute plate motion (APM) direction of southern Africa since the end of the Jurassic. This alignment is similar to that previously reported for the North American craton and suggests that the large-scale component of mantle anisotropy on both sides of the Atlantic can be related to resistive drag at the base of the plates. A likely depth range of the corresponding deformation is 150-400km. Analysis of the mantle converted phases suggests that the high-velocity mantle root of the Kaapvaal craton, which translates coherently with the plate, resides in the same depth range: the bottom of the root is found at a depth close to 380 km. We conclude that the root is deformed by the recent plate motion, but the deformations are not strong enough to be seen in the available tomographic models.
1. Introduction Mobility of anisotropic minerals is sensitive to temperature (e.g. Nicolas and Christensen, 1987). Hence origin and age of mantle anisotropy may depend on depth. Anisotropy immediately below the Moho is probably frozen and related to processes of a distant past, when the rock within this layer was hot (temperature higher than about 1000°C). This is documented by numerous seismic refraction experiments in the oceans (e.g. Bibee and Shor, 1976; Shimamura, 1984). Deeper than about 100km, anisotropy beneath the oceans is related to the pre-
* Corresponding author,
sent-day plate tectonics; this is known from the surface-wave analyses (e.g. Montagner and Tanimoto, 1991). The frozen layer in the oceans is thin in comparison with the rest of the anisotropic mantle, and the cumulative effect of anisotropy in the teleseismic shear phases such as ScS reflects mainly processes not much older than about 15 Ma (Farra and Vinnik, 1994). Anisotropic properties of the continental mantle are less well understood, partly because the history and structure of the continental mantle are more complicated and partly because systematic seismic observations of continental anisotropy started much later, after the introduction of SKS techniques (Vinnik et al., 1984). Unfortunately, observations of splitting in SKS provide little information on the depth of
0031-9201/96/$15.00 © 1996 Elsevier Science B.V. All rights reserved SSDI 0031-9201(95)03123-5
140
L.P. Vinnik et al. / Physics of the Earth and Planetary Interiors 95 (1996) 139-151
the effect, and the data are interpreted mainly by comparing the directions of anisotropy with other observables, such as tectonic trends or absolute plate motion (APM) directions. These correlations lead to controversial interpretations. Silver and Chan (1991) claimed that the fast directions of splitting in SKS at several stations are aligned with the old crustal fabric and concluded that anisotropy in the stable continental regions has been frozen in the uppermost mantle since the last tectonic episode (the Precambrian orogenies in North America and Africa, the Hercynian in Europe). Babuska et al. (1993)argued that observations of splitting in Central Europe are related to frozen anisotropy in the subcrustal lithosphere. On the other hand, Vinnik et al. (1989, 1992) concluded that the major contribution to the observed shearwave splitting in stable continental regions, as in the oceans, comes from the asthenosphere, where anisotropy is related to the recent plate motions. This division of opinions is present in m a n y other papers, Correlation between fast direction of anisotropy and recent APM direction is especially strong in the east of North America, but it is less impressive in some other regions (Vinnik et al., 1992). In this paper we present and analyse observations of shearwave splitting at a number of seismograph stations within the Kaapvaal craton in southern Africa, where the only previously known estimate of the direction of anisotropy (Station SLR) is in a strong disagreement with the recent APM direction. Our seismic observations are carried out at seven digital broadband seismograph stations in a corridor 700km long crossing the craton in a SW-NE direction (Table 1 and Fig. 1). This portable array was fully operational
Table 1 List of portable stations with the effective values of the parameters of anisotropy: parameters of the events are given in Table 2
Station Latitude (deg) BPI -26.175 WAR -28.377 DOU -29.115 KAM -29.133
Events
a
(deg) 28.030 24.893 23.865 23.125
3,5,7, 10 2 5,6, l0 1,5
(deg) 20 20 40 20
st (s) 0.4 1.3 1.1 1.4
SAN PIL
-24.626 -25.221
27.618 27.101
5 , 6 , 7 , 9 , 11 4,5,7,8,10,11
30 50
0.7 0.6
KLI
-25.853 26.272
7, 10, II
40
1.0
Longitude
20°E
25°E
30°E
~ . ~ : , , ~ / \.~ KAAPVAAL . . . CRATON . . . "..... :l ~ ~:~s~ i ,/ ( ~\]-\ _ _ , '~/
'
~
k_
/
25°S
3oOs
~.,...,_..,.._.~-~ '
'
Fig. 1. Map of southern Africa. Boundaries of the Kaapvaal craton are adopted from De Wit et al. (1992). Direction of polarization of the fast split wave at every station is shown by a straight line; the length of the line is proportional to lapse time 8t of the slow wave. Direction of APM since the end of the Jurassic is shown by
an arrow.
from the end of 1989 until the end of 1990. The seismographs used in the experiments have been described by Green (1992).
2. Data and results
We invert the records of SKS and similar seismic phases for polarization direction of the fast split wave (ct) counted clockwise from north and time lag 8t of the slow wave, by using the algorithm by Vinnik et al. (1989). Polarization of the fast wave is roughly parallel to the [100] axis in olivine and, consequently, to the longest axis of the strain ellipsoid or to the direction of the mantle flow, if it is in the form of progressive simple shear (McKenzie, 1979; Nicolas and Christensen, 1987; Ribe, 1989, 1992). The parameters of anisotropy are found by minimizing the root mean square (r.m.s.) difference between the observed transverse component T of SKS and the transverse component Tc derived from the observed radial component (R). These estimates are related to the nearest vicinity of the station
L.P. Vinnik et a l . / Physics of the Earth and Planetary Interiors" 95 (1996) 139-151
within the first Fresnel zone. The optimal pair of the parameters is found by a grid search over the values of the parameters between 0° and 180° for t~ and between 0s and 2 s for ~t. The resulting penalty function is computed for every individual record and for the whole set of records of the seismograph station. E can be expressed as 1/2
E(ot,St)
E(a,~t)=(N_t~f[T(t)-T~(t,a,St)]2dt} i= 1
~-~ t'~'t
where N is the number of records. To distinguish between the individual estimates (N = 1) and those for the group of records, the latter will be referred to as the effective values, The accuracy of the estimates is easy to derive for an individual record, assuming that the only source of errors is additive seismic noise. For the estimates of a obtained from good-quality data (signal/noise ratio in the T component higher than three), the accuracy thus evaluated is usually within a range of a few degrees. However, the scatter of the individual estimates for events with strongly different back azimuths is usully much larger than can be explained by the additive seismic noise. This scatter is caused mainly by the Earth's complexity, and, using the effective values, one should somehow deal with the 'noise', which is not additive and the statistical properties of which are not known. The only possibility for judging this effect is to compare the estimates for events with strongly different epicentral parameters,
141
Seismic phases suitable for the analysis, SKS, SKKS and PKS, with a high signal/noise ratio are found in the records of 11 events (Table 2). At most stations, the estimates of the parameters of anisotropy are obtained by processing the records of several events with different back azimuths (Table 1). The records are individually filtered in the period range between 3 s and 15 s to attain the highest possible signal/noise ratio. Representative examples of the filtered records are shown in Fig. 2. The PKS records of Event 5 in Fig. 2(a) are interesting, because this phase is seldom used for the analysis of splitting. Large amplitudes of PKS are caused by the caustic at a distance near 128 °. The PKS record at SAN is of marginal quality, but a better quality SKKS is present on the same seismogram. The noise in the records in Fig. 2 is low, and there is a quarter-period phase shift between the R and T components, which presents a diagnostic property of azimuthal anisotropy. The effective penalty functions, obtained by combined processing of all available records, are shown in Fig. 3. To achieve a better understanding of the quality and implications of the results obtained, we will present and discuss some data in detail. Fig. 4 shows examples of the penalty functions for Station PIL. In the record of Event 4, with a back azimuth of 50 °, there is no signal in the T component. This indicates that the axis of symmetry beneath the station is either parallel or perpendicular to this direction. The minima of the respective penalty functions in Fig. 4 are observed at azimuths around
Table 2 List of events: back azimuths and distances are given with respect to BPI No.
Date
Latitude (deg)
Longitude (deg)
Depth (kin)
Back azimuth (deg)
Distance (deg)
Phase
1 2 3 4 5 6 7 8 9 10 11
34389:09 12 89 34989:15 12 89 10890:18 04 90 11690:26 04 90 13290:12 05 90 15090:30 05 90 16590:14 06 90 26090:17 09 90 28390:10 10 90 29090:17 10 90 36490:30 12 90
0.1N 8.3N 1.2N 36.2N 49.0N 6.0S 47.9N 53.2E 19.5S 11.0S 6.9S
123.3E 126.7E 122.9E 100.3E 141.9E 77.2W 85.1E 159.6E 66.6W 70.8W 151.0E
151 24 26 10 606 24 58 10 127 599 179
92 86 91 50 47 258 34 153 251 256 111
94.8 101.4 94.9 92.1 124.1 100.9 89.7 90.5 85.6 93.0 116.6
SKS SKS SKS SKS PKS, SKKS SKS SKS SKS SKS SKS SKS, SKKS
142
L.P. Vinnik et a l . / Physics (dthe Earth and Planetary hueriors 95 (1996) I39-151
(al
' j SAN
~
I
~ BP1
.
~
,
~ DOU
~
l l ' l i , ~ KaM
,
0.0
(
20.0
~
400
~
60 0
time(s)
| BPI
•,
.
. -.,
', ,
PIL
~
,
,~,. KLI
~
~
50 ° and 140°, and there are no estimates of ~t. A similar solution is obtained for SKKS of Event 5, with a back azimuth of 47 ° (not shown in Fig. 4). For the record of Event 7, with a back azimuth of 34°, there are also two minima at the same azimuths, with the minimum at 140° being slightly deeper than the other one (Fig. 4). The solutions for Events 4, 5 and 7, whose back azimuths are close to each other, are, clearly, very similar; practically the only source of scatter of the estimates within this subset of data is the additive seismic noise in the T component, and its disturbing contribution to the estimates of fast azimuth is less than 10°. A comparable solution is also obtained for Event 8, the back azimuth direction of which (153 °) is roughly perpendicular to those of Events 4 and 5. In the SKS record of Event 10, the signal in the T component is clearly seen (Fig. 2(b)), and a record of comparable quality is obtained for SKKS of Event 11. Back azimuths of these events are 256 ° and 111 °. The minima of the respective penalty functions are observed at 30° and 0.5 s, and at 60 ° and 0.7 s (Fig. 4). The difference of 30° between the two solutions is much larger than the effect of the additive seismic noise. Discrepancies of such magnitude are reported for many stations in other regions (e.g. Vinnik et al., 1994), and should not be considered as something unusual. The effective estimates of the parameters for the whole set of records of PIL are 50° and 0.6 s (Fig. 3), which differ by up to 20 ° and 0.1 s from the individual estimates. A formal evaluation of the accuracy of the effective estimates is difficult because the main source of 'noise' in this case is complexity of the Earth's medium, and the properties of this noise are hard to quantify. We may assume, however, that the likely error is within the range of scatter of the individual estimates, or less than about 20 ° for the azimuth and 0.2 s for ~t. The data for BPI are comparable in signal/noise ratio with those for PIL. For the PKS record of Event 5 (Fig. 2(a)), there is a region of small values of the
DOU
O0
~
200
~
40.0
~ 60 0 time(s)
Fig. 2. Examples of the R (continuous lines) and T component records (dashed lines): (a) PKS phase of Event 5; (b) SKS phase of Event 10. Vertical lines indicate time windows used to invert the data for the parameters of splitting.
L.P. Vinnik et al./ Physics of the Earth and Planetary Interiors 95 (1996) 139-151
penalty function between 150° and 40°, with the minimum at 170° and 0.4s (Fig. 5). In the record of Event 7, the signal in the T component is practically missing; in the resulting penalty function there are two minima, at about 400 and 130° (Fig. 5). The other record with a missing signal in the T component is Event 3, with a back azimuth of 91 °. The minima for this record are around 10° and 100°. For Event 10 (Fig. 2(b)), there is a minimum of the penalty function at around I0 ° and 0.7s (Fig. 5). It is clear that, in spite of scatter of the individual estimates, the regions of the smallest values of the penalty functions overlap in the vicinity of 20 ° and 0.4 s, and the effective solution (Fig. 3) is close to these values. The accuracy of the effective values for BPI is comparable with that for PIL, and results of comparable quality are obtained for KLI and KAM. The results for DOU (Fig. 6) deserve special discussion because of a peculiar effect in St. In all three records of this station (Events 5, 6 and 10), the (a)
BPI
DOU
signal in the T component is well seen; two of them (Events 5 and 10) are shown in Fig. 2. The respectire penalty functions (Fig. 6) have clear minima at 40 ° and 1.8 s, 60 ° and 2 s, and 60 ° and 1.6 s. Back azimuths of Events 6 and 10 differ by only 2 °, and their solutions demonstrate once more that the influence of the additive seismic noise upon the estimate of fast azimuth is within a range of a few degrees. The influence upon the estimate of 8t in this case is around 0.4s. The effective estimate of ~t (Fig. 3) is 1. l s, which is lower than each of the individual estimates, and this bias should not be ignored when the effective estimates are interpreted. The effective estimate of the azimuth is 40°, which coincides with the solution for Event 5. The data for SAN present a very singular character because of the unusually large scatter of the individual estimates. For the PKS record of Event 5, shown in Fig. 2(a), the solution (Fig. 7) is 140° and 2 s. This record is of marginal quality, but a similar (b)
oI ......... oi o I
180. O.
DELRY (S)
2.
180. I O.
KAN
~ 2.
I8
, O.
18 DELAY {S)
2.
O.
DELRY ISi
2.
NRR
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_
5AN
. . . . . . . . . .
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_
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PIL
o
, , DELR" iS)
143
-
%. 180.
18o.
O.
DELR'i' IS)
2.
, ,
O.
I8O.
~LR'i" IS)
2.
O.
DELAY IS}
2.
Fig. 3. Plots of the penalty function E(oL,St) obtained by combined processing of all available records and used to find effective values of the parameters.
144
L.P. Vinnik et al. / Physics of the Earth and Planetary Interiors 95 (1996) 139-151
solution is obtained for the record of Event 7, the back azimuth for which differs little from that of Event 5 (Fig. 7). A comparable solution is obtained also for the SKKS record of Event 5. For the record of Event 6, with a back azimuth near 258° and a high signal/noise ratio in both components, the solution is around 40° and 0-5 s (Fig- 7); a similar solution is obtained for the record of Event 9, with almost the same back azimuth. The solution for the well-recorded SKS of Event 11, with a back azimuth of 111° (80° and 1.6 s; Fig. 7), is very different from the others, and a close solution (100 ° and 2 s) is obtained for SKKS of the same event. Remarkably, in spite of the enormous scatter of the individual estimates, the effective estimates for SAN (30° and 0.7 s; Fig. 3) are close to those for the neighbouring Stations PIL and BPI. To summarize, the effective estimates of the fast
i
1
k
i
~
I
I
J
i
i
i
i
i
i
BP; #5
BPI #7 ' ........ ~ ~~--___~
0.
0.
~
~ l ( / ~
~
~ l ' ~ ~ ?-
,
, ,', , ~(~/~,f~,/f' f , ~ 180.
0.
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O.
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-i
i
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Fig. 5. The same as Fig. 4 but for BPI.
~: ( 180.
~
~ 180.
,
BELl:ITIS1 PIL #10
O. ~
' ...... ~
2. _
. DELAftS~ PIL #1t 0. J ' ~ ~ 1 ~ ~ O.
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. ~ -
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3. S-velocity structure o f the m a n t l e beneath the Kaapvaal craton
18o. I . . . . . . . . . O.
~LR'f IS1
2.
O.
OELfff ISI
2.
Fig. 4. Examples of plots of the penalty function E(t~,St) for the
records of PIL.
direction for the stations of the network are in the interval between 20° and 50°' with a mean value near 30°. The differences between the effective estimates of a for the various stations are generally within the range of the accuracy of measurement. The estimates of Bt, considering both the individual and effective estimates, are about two times lower in the north-east (BPI, SAN and PIL) than in the south-west (DOU, KAM and WAR). SAN, the station with anomalously strong variability of the individual estimates, is very close to the normal Station PIL; hence, the source of the anomaly is very shallow, perhaps crustal.
To understand the splitting observations, additional data on the deep structure of the study region
L.P. Vinnik et al. / Physics of the Earth and Planetary Interiors' 95 (1996) 139-151
are required. Generally, the upper mantle of Precambrian platforms is characterized by higher than average P and S velocities (e.g. Suet al., 1994). The high velocities can be explained by the presence of old mantle roots, which may differ from the normal mantle in chemical composition; the roots, a few hundred kilometres deep, translate coherently with the plates (Jordan, 1988). The maximum depths of the roots are not known, owing to the limited resolution of the available seismic data. For the study region, high S velocities in the upper mantle are confirmed by the multimode surface-wave dispersion data (Bloch et al., 1969). More recently, the records of the array shown in Fig. 1 have been used to infer a detailed S-velocity profile of the upper mantle beneath the Kaapvaal craton. Here we present our preliminary results, and a comprehensive report will be published elsewhere (Vinnik et al., 1996); our preliminary results are not much different from the final ones.
145
SnN#s
SAN#6
i ..... :' '~L o., ~ ~ l ~ ~......... vq~ i
~
~
1
~
~
~ ~~,~.~ q \ ~'-..~o~"--~_
~
180. . . . . . . . . . . 180. 0. t~LAY~SJ 2. 0. t~LnY tS~ 2. SRN#~ SAN#11 0. ~ ' ' ~ ' \,.,~., 0. e'0 " ~ .~'/"---~--~-~ ,
~
m
Fig. 7. The same as Fig. 4 but for SAN.
°I ]80 . . . . . . . . . . . 0. Itl.flY (SI o0u #10 0. ' . . . . . . . .
.......
100. 2.
0.
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0ELAYtS~
2.
Fig. 6. The same as Fig. 4 but for DOU.
2.
S-Velocity gradients in the mantle beneath the Kaapvaal craton have been inferred from the observations of the teleseismic P-to-SV converted phases. These phases are detected by using a technique (Vinnik, 1977) that consists of source equalization (deconvolution), rotation of the seismogram into the L and H coordinate frame and stacking the H component records of many events with appropriate moveout corrections. The L and H directions coincide with the directions of the principal motion components in the P wave and the converted phases, respectively. Standardization is performed by using a filtering technique which makes the P waveform look like a 8 function. Details of the techniques and the results obtained elsewhere, have been described by, for example, Kind and Vinnik (1988), Stammler et al. (1992) and Petersen et al. (1993). We processed the records of 6, 8, 6, 8 and 10 events in the distance range 30-90 ° for Stations BPI, DOU, SAN, PIL and KLI, respectively. To improve
146
L.P, Vinnik et al. / Physics of the Earth and Planetary Interiors 95 (1996) 139-151 PIL+KLI+DOU
800- ~ 750- ~--~-~ ~ 640--~'~-/"
I ~
I
~
560480- ! a-400-
350240-
t6080-
0' T' ' ' I , , , I ' ' ' I ' ' ' I ' ' ' t . . . .
-20
0
20 40 Time (s)
60
BO
Fig. 8. Result of stacking the H components of Stations DOU, KLI and PIE Moveout corrections for stacking were calculated for the trial depths of conversion, indicated on the left. Time is shown relative to the P wave at 0s. Times of the detected mantle phases are shown by marks at the top.
the signal/noise ratio, the results of the stacking were additionally low-pass filtered with a cutoff frequency between 0.1 and 0.2 Hz. Fig. 8 shows the result of stacking for DOU, PIL and KLI; comparable results are obtained by stacking the records of all stations. The phases between 0 and 30s are formed by conversions and multiple reflections from the Moho and free surface. The phase with a time (lapse time relative to P) of 66.0s is related to the 6 5 0 k m discontinuity. A remarkable result of data processing is the phase with a downward motion at 38 s. Such motion is possible if the conversion occurs at the upper boundary of a low-velocity layer. This phase is detected reliably, because its amplitude (amplitude ratio between the converted phase and the P wave) is 0.06, which is much higher than the level of the preceding noise and the standard error of the stack (0.007); moreover, its amplitude is largest at trial depths around 400km, thus indicating that its pattern of amplitude variation vs. trial depth is in agreement with the depth of conversion that can be inferred from its time. The last criterion is especially important because it provides evidence that the detected phase is a true converted phase rather than an artefact of data processing. Additionally, to test the
stability of the observed wavefieid, the records were divided into subgroups, with roughly the same amount of records in each subgroup, and these subgroups were processed separately. A large downward motion at 38 s has been detected in every subgroup. The time of the 6 5 0 k m phase (66.0s) is 2 s less than for the standard Earth model IASP91 (Kennett and Engdahl, 1991). On a global scale, the travel time of the 6 5 0 k m phase in the mantle transition zone is stable, and the variations observed at the Earth's surface are controlled mainly by the lateral velocity variations within the upper, most heterogeneous zone of the Earth, of 400 km thickness (Stammler et al., 1992). The observed travel time residual of the converted phase is practically the difference between the teleseismic S and P residuals that are accumulated in that zone. Assuming, in agreement published data (e.g. Romanowicz and Cara, 1980), that the ratio between the S and P residuals is close to three, the teleseismic S residual in the study region can be estimated as - 3 s. This estimate is in good quantitative agreement with the data for the other Precambrian shields (e.g. Su et al., 1994). The S-velocity model SVKC explaining the waveform of the 38 s conversion (Fig. 9) is found by a trial-and-error method. The theoretical seismograms for the trial models are computed in the 3 0 - 9 0 ° with
5.5 5.4
53 ~, 5.2 "~ 51
IASP91,-'""
~ 50 ~ 4.9 ~ 48
'-4 ~
~
4..7
46
,
4.5 '
360
i
4.60 . .440. . . 480 o~prH ( ~ l
5~,o '
Fig. 9. SVKC S-velocity model inferred from the data shown in Fig. 8 (continuous line) and standard model 1ASP91 (Kennett and
Engdahl,1991).
L.P. Vinnik et al./ Physics of the Earth and Planetary Interiors 95 (1996) 139-151
distance range by using the reflectivity technique (Fuchs and Mueller, 1971) and processed like the observed seismograms. The P waveform found by processing the observed seismograms, with a correction for mantle attenuation, was used as the input for computing theoretical seismograms. The Q values were taken from EURSQ model (Der et al., 1986), which is based on seismic data for the Precambrian platforms. To account for the presence of the highvelocity mantle root of the craton, the S and P velocities in the mantle at depths less than 380 km are increased relative to IASP91 by 0.1 km s-~ and 0.15 km s -~, respectively. The depth and S-velocity contrast of the 380km discontinuity in our model are constrained by the time and amplitude of the 38 s phase, respectively. The discontinuity is made sharp for reasons of simplicity. The adopted thickness of the layer where S velocity is lower than in IASP91 is, practically, the smallest acceptable thickness. If it is made substantially smaller, a strong converted phase at about 43 s which is missing in the observations appears in the theoretical seismograms. SVKC provides a very good fit between the observed and synthetic waveforms of the 38 s phase, Any alternative explanation of the 38 s signal is difficult. It could be assumed, for example, that the waveform of the Ps phase is distorted by a multiple reflection of the type Ps(d)s (P converted at the surface to S and then reflected once from an interface at depth d). By appropriate selection of d, the times of the multiples and the 400 km Ps phase at a certain epicentral distance can be equalized, but the moveout corrections which are needed for detecting the multiples and the Ps phases are very different. The delay of Ps relative to P grows with increasing ray parameter or decreasing epicentral distance; the dependence of the delay of the multiples is opposite (Vinnik and Mikhailova, 1988). The dependence of the amplitude on trial depth in Fig. 8 precludes the explanation of the 38 s phase in terms of the multipies: if this phase was produced by the multiples, the largest amplitudes were obtained at zero trial depth, as for the crustal multiples in Fig. 8. It could also be proposed that the anomalous phase is converted at the normal 400 km discontinuity, but the waveform is distorted by topography of the discontinuity. There are a few objections to this interpretation. The normal time delay of the 650 km
147
phase relative to the 400km phase is very stable and close to 24s (Stammler et al., 1992). This means that, in the study region, the 400km phase is expected to arrive at around 42 s. A lowered S velocity beneath the 400km discontinuity may change this estimate by only a fraction of a second. The anomalous phase arrives a few seconds earlier, which means that the related discontinuity is displaced upward by a few tens of kilometres relative to the standard depth of the 400km discontinuity. If the 400 km discontinuity is interpreted conventionally as the olivine-spinel phase transformation with a positive Clapeyron slope, to provide this uplift the temperature should be lowered by a few hundred degrees, which is out of the question for the study region. The topography would lower the amplitude of the detected signal, because the signal coherency, which is a prerequisite for our signal detection techniques, would deteriorate. In reality, the amplitude of the downward motion at 38 s is twice as large as the normal amplitude of the standard 400 km Ps phase. Finally, we do not think that any topography can transform the upward motion of the 400 km Ps wave into the downward motion in Fig. 8; there is a phase shift in the Sp waves for some angles of incidence (Aki and Richards, 1980), but this is not applicable to the Ps waves. Similar analyses are known for a number ofcontinental stations (e.g. Stammler et al., 1992), but nothing comparable with that shown in Fig. 8 has ever been seen. Numerous S-velocity models derived from various kinds of seismic data for many regions usually show that in the interval between 300 and 650km, the P and S velocities rise with depth. Some complications may exist in the vicinity of the currently active subduction zones, but far outside subduction zones the low S velocity at depths between 350 and 500km should be viewed as a strong anomaly. The mantle beneath Southern Africa is anomalously heated (Nyblade and Robinson, 1994), but within the Kaapvaal craton the heat flow data give no evidence of a thermal perturbation (Jones, 1988); this can be explained by a large thickness of the lithosphere of the craton. With the indications of heating taken into account, the low-velocity layer in the SVKC model can be explained by elevated ternperature that comes close to the solidus temperature.
148
L.P. Vinnik et al. / Physics of the Earth and Planetary Interiors 95 (1996) 139-151
The upper part of the low-velocity layer may correspond to 'anticrust', the layer enriched by, among other components, volatiles (Gasparik, 1992). The temperature of the mantle solidus is lowered by the volatiles (Thompson, 1992). The 380kin discontinuity may separate the root of the craton or 'tectosphere' from the underlying mantle; the S velocity above the discontinuity is higher, because the root may have a different composition and a higher solidus temperature (Jordan, 1988). To conclude this section, the seismic data give reasons to believe that the mantle root of the Kaapvaal craton comes close to a 400kin depth. Temperatures in the lower part of the root are well in excess of 1300°C, and the root should not be mistaken for the lithosphere,
4. Discussion The fast directions in Fig. 1 are fairly uniform, Has this uniform direction been frozen in the uppermost mantle since the Archaean, or does it reflect a process which operated much later at greater depths? Observations of splitting within the currently active tectonic zones (collisional belts, zones of rifling) indicate that the fast directions of mantle anisotropy in the upper mantle are often either parallel or perpendicular to the strikes of these zones (collisional belts and riffs, respectively), but there are notable exceptions to these regularities (e.g. Vinnik et al., 1992). One may speculate that the tectonic processes of a distant geologic past are very similar to those active at present, and that the anisotropy thus created can be preserved in the mantle for a billion years or so, but to become credible, these ideas must be confirmed by seismic observations sensitive to depth of anisotropy. There are a few currently stable regions where such observations are available, namely, the east of North America, southem Germany and the Iberian peninsula, In North America, there is a clear trend in the fast directions of SKS, which coincides with the recent absolute plate motion direction (Vinnik et al., 1992). One can see this trend also in more recent observations (Silver and Kaneshima, 1993). Moreover, seismic long-range refraction observations (Iyer et al., 1969) do not show any anisotropy in the upper
100 km of the mantle where frozen anisotropy should reside; a detailed discussion of the implications of the long-range profiling data in North America has been given elsewhere (Vinnik et al., 1992). In southern Germany, seismic refraction observations (Bamford, 1977)reveal frozen anisotropy whose direction has nothing in common with either the Hercynian crustal fabric or the fast directions in SKS; additional data confirming this direction have been given by Farra et al. (1991). The observations of SKS at some stations in southern Germany can be inverted for two-layer anisotropic structures. In the resulting models (Vinnik et al., 1984), the anisotropic layer in the subcrustal lithosphere is relatively thin, and ~t in the asthenosphere is twice as large as in the subcrustal lithosphere. In the Iberian peninsula, as shown by seismic refraction experiments, the subcrustal lithosphere at depths less than 100km includes three thin anisotropic layers with the fast direction around NNE (Diaz et al., 1993). This direction has nothing to do with either the fast direction of anisotropy in SKS or the Hercynian tectonics. Moreover, longperiod surface-wave data which are sensitive to depth of anisotropy indicate that anisotropy beneath the Iberian plate resides at depths exceeding 100km (Maupin and Cara, 1992). The layers observed by Diaz et al. (1993) are too thin, and anisotropy within them is too weak to affect the observations of SKS and long-period surface waves. All observations (SKS, surface waves, P-wave refraction) are fully compatible, if the effect in SKS is accumulated mainly in the asthenosphere, where it is related to the relatively recent (most likely, Cretaceous) motion of the Iberian plate. To conclude, when seismic data sensitive to depth are available, they strongly suggest that (1)the dominant effect in SKS comes from the asthenosphere, and (2) the fast direction of frozen anisotropy in the subcrustal lithosphere differs from the fast direction in SKS. The early observations of splitting in SKS in southern Africa (Station SLR, a around 80 °) were interpreted by Silver and Chan (1991) in terms of frozen anisotropy in the subcrustai lithosphere. The reason for this interpretation is a similarity between the fast direction at SLR and the strike of the Limpopo belt, a feature several hundred kilometres to the north of the station. This interpretation is incompatible with the data, available at present (Fig.
L.P. Vinnik et a l . / Physics of the Earth and Planetary Interiors 95 (1996) 139-151
1). There are Proterozoic mobile belts with a suitable orientation beyond the north-western border of the Kalahari craton (Hartnady et al., 1985), but on tectonic maps at such distances (about 1000 km from the observation site) one can find not only this but also many different directions. On the other hand, the direction of anisotropy at our stations practically coincides with the APM direction of southern Africa in the hotspot reference frame from the end of the Jurassic to the present (Dietz and Holden, 1970). As shown by the trails of the hotspots in the Southern Atlantic, during this time the plate moved by approximately 3000kin. Approximately 25Ma ago Africa halted (Burke and Wilson, 1972), but apparently the accumulated deformation is still preserved. Within the North American craton, the fast direction of anisotropy is close to the present-day APM direction. This APM direction is related to the opening of the North Atlantic and has been preserved without a significant change since the end of the Cretaceous, 65Ma ago (Dietz and Holden, 1970). The relationship between the fast direction of anisotropy and the relatively recent APM direction in southern Africa is, clearly, very similar to that in North America, and suggests the same origin of the large-scale component of mantle anisotropy in both regions: resistive drag at the base of the moving plate. In that case, the deformation near the base of the plate is in the form of progressive simple shear which yields orientation of the [100] axis of olivine in the direction of plate motion (Ribe, 1989). Lattice preferred orientation (LPO) of olivine is sensitive to deformation only at temperatures higher than about 1000°C (Nicolas and Christensen, 1987). A temperature of 1000°C underneath the Kaapvaai craton, as deduced from the mineral equilibrium studies of kimberlite inclusions, is reached at depths around 150kin (Boyd and Gurney, 1986). This estimate is remarkably consistent with the geotherm calculated from heat flow data; the present-day geotherm differs little from that in the Cretaceous (Jones, 1988). The properties of minerals at depths exceeding 400km seem to be unfavourable for developing significant anisotropy (Mainprice and Silver, 1993). Hence, anisotropy created during the last 150Ma in the mantle of the Kaapvaal craton is probably located between 150 and 400km, in the depth range occupied by the deep continental root of
149
the craton. To obtain ~t of 1 s in a layer of 250km thickness, the required difference between the split wave velocities is 2%. This difference is about onethird of the maximum values found in samples of the upper-mantle rocks. The indications of relatively young deformations of the deep continental root are not necessarily at odds with seismic evidence of its coherent motion with the plate. Significant LPO develops by intracrystalline slip when ln(cl/c 2) or ln(c2/c 3) (where c I, c 2 and c 3 are respectively the largest, intermediate and smallest axes of the strain ellipsoid), is around 0.3 (Ribe, 1992). The equivalent shear strain y is close to unity. These theoretical predictions can be complemented by observational data: experiments on simple shear in ice (Bouchez and Duval, 1982) demonstrated that a good fabric was induced for a y of only 0.7. This range of strains is reached when the root at a depth of 400km is displaced laterally by less than 200kin with respect to a 150km depth. A shift of such magnitude at a 400 km depth is difficult to detect in the currently available tomographic models (e.g. S u e t al., 1994). The estimates of ~t are relatively low for the north-eastern stations and relatively high in the south-west. The number of Jurassic-Cretaceous kimberlite intrusives within our corridor is much larger in the south-west than in the north-east (Skinner, 1989). This correlation suggests a relationship between kimberlites and mantle anisotropy. We speculate that both phenomena could be affected by the influx of volatiles: volatiles are abundant in the kimberlite magma, and they could intensify recrystallization of olivine. At temperatures higher than about 1200°C, recrystallization takes place on grain boundaries and, in the presence of strain, increases anisotropy (D. Mainprice, personal communication, 1993). To conclude, the large-scale component of mantle anisotropy (scale on the order of 1000km) within the Kaapvaal craton and similar regions of North America can most easily be explained by the relatively recent (ages around 50-100Ma) deformations in the lower (deeper than about 150km) part of the continental root. The temperatures at 200-400km depths beneath the Kaapvaal craton are in the range which is commonly ascribed to the asthenosphere, and it should not be regarded as a great surprise that the
150
L.P. Vinnik et al.// Physics of the Earth and Planetary Interiors 95 (1996) 139-151
mantle at these depths is d e f o r m e d by shearing, W h a t is surprising is the w e a k n e s s o f the d e f o r m a -
crust and upper mantle of southern Africa from multi-mode surface wave dispersion. Bull. Seismol. Soc. Am., 59: 1599-
tions. A contribution to the o b s e r v e d s h e a r - w a v e
1630. Bouchez, J.L. and Dural, P., 1982. The fabric of polycrystalline ice deformed in simple shear: experiments in torsion, natural deformation and geometrical interpretation. Text. Microstruct., 5: 1-17. Boyd, F.R. and Gumey, J.J., 1986. Diamonds and the African lithosphere. Science, 232: 472-477. Burke, K. and Wilson, J.T., 1972. Is the African plate stationary? Nature, 239: 387-390. Der, Z.A., Lees, A. and Cormier, V.F., 1986. Frequency dependence of Q in the mantle underlying the shield areas of Eurasia, Part 3: The Q model. Geophys. J. R. Astron. Soc., 87:
splitting f r o m e v e n greater depths cannot be excluded, although with the data available at present this does not look very likely. W e do not want to say that older deformations are absent in the subcratonic upper mantle. S a m p l e s o f rocks f r o m the South African subcrustal lithosphere are w e a k l y anisotropic (Mainprice and Silver, 1993), and this anisotropy could contribute to the small-scale lateral variations o f the parameters of splitting. Unfortunately, interpretation o f small-scale and azimuthal variations o f the inferred parameters o f anisotropy is difficult, because there are too m a n y possible reasons for them. W a v e scattering at small-scale lateral i n h o m o geneities within the crust c o u l d be one o f these reasons. S o m e effects o f layered anisotropic structure in seismic w a v e f i e l d s can be used to infer the eters o f anisotropy in the layers (e.g. Farra 1991; Vinnik et al., 1994), but either these are w e a k at our stations or the a m o u n t o f
paramet al., effects data is
insufficient to identify them.
Acknowledgements The authors are m u c h indebted for h e l p to E.O. Kostlin, L.I. M a k e y e v a , R. K i n d and G. Asch. W e thank an a n o n y m o u s r e v i e w e r and M. G r a n e t for helpful reviews. Part o f this research was carried out in the S e i s m o l o g i s c h e s Z e n t r a l o b s e r v a t o r i u m o f G e r m a n y (Erlangen), where one o f the authors (L.V.) was supported by the A l e x a n d e r - y o n - H u m b o l d t -
Stiftung.
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