Antarctica II: Upper-mantle structure from velocities and anisotropy

,~ ~

ELSEVIER

Physics of the Earth and Planetary Interiors 84 (1994) 33—57

PHYSICS OFTHE EARTH ANDPLANETARY INTERIORS

Antarctica II: Upper-mantle structure from velocities and anisotropy G. Roult

*,a, D. Rouland b J.P. Montagner a Institut de Physique du Globe de Paris, 4 place Jussieu, 75252 Paris Cedex 05, France b Ecole et Observatoire de Physique du Globe, UniversitéLouis Pasteur, 5 rue Descartes, 67084 Strasbourg Cedex, France a Département de Sismologie,

(Received 5 May 1993; revision accepted 8 December 1993)

Abstract To improve the lateral resolution of three-dimensional seismic wave velocity models in Antarctica and the surrounding oceans, we have analysed direct earthquake-to-station Rayleigh-wave data observed on the vertical high-gain long-period and the very long period components of seven GEOSCOPE stations located in the southern hemisphere and three other stations at equatorial latitudes. The phase velocities of Rayleigh waves along 400 well-distributed paths are obtained in the period range 60—300 s, by fitting the data with synthetic seismograms computed with known source parameters in a reference earth model represented by the Preliminary Reference Earth Model (PREM). Corrections for shallow layers have been carefully applied to the observed phase velocities. The geographical distributions of phase velocities and azimuthal anisotropy are then computed with the tomographic method without any a priori regionalization developed by Montagner (Ann. Geophys., 4(B3): 283—294, 1986). The results show some new and important features of Antarctica and the southern hemisphere. The locations of velocity anomalies are well resolved. The eastern part of Antarctica corresponds to a craton-like structure down to depths of about 250 km, and the highest velocities are observed in Enderby Land, where some of the oldest rocks in the world have been sampled. The low velocities are located along the ridges encircling the Antarctic continent. The lowest velocities appear in some areas corresponding to hotspots (Crozet, Kerguelen, Macquarie and Balleny Islands). Also, an elongated low velocity is found on the western flank of the Transantarctic Mountains, which might be related to the existence of a rift zone similar to the African rift. The Australia—Antarctica Discordance (AAD) presents slow velocities near the surface but fast velocities below the lithosphere. These main features are discussed in the framework of the Gondwana hypothesis and the earlier supercontinent. The first azimuthal anisotropy results are also discussed. Anisotropy values are smaller within the Antarctic continent than in the surrounding oceans. They are also small in the AAD but particularly large in the areas around it, suggesting an active tectonic process characterized by a downward flow at depth, a good candidate for a cold spot or a new subduction zone.

*

Corresponding author.

Elsevier Science B.V. SSDI 0031-9201(94)05025-S

34

G. Roult et al. /Physics of the Earth and Planetary Interiors 84 (1994) 33—57

1. Introduction The Antarctic continent has been known since the beginning of the century (Du Toit, 1937) to have been part of a southern supercontinent, and its accurate positioning plays a key role in the reconstruction of this earlier continent. The International Geophysical Year (IGY, 1957—1958) contributed greatly to our knowledge of this continent. Early this century, the various orogenic phases in Antarctica were recognized (see review by Roult and Rouland, 1994). By developing scientific programmes in high latitudes, the IGY allowed the first geophysical investigations on the Antarctic crustal structure. The harsh conditions for maintenance of permanent seismic stations and the inaccessibility of this part of the world made it difficult to undertake geophysical studies until recent decades. The recent developments of world-wide seismological digital networks, such as the International Deployment of Accelerometers (IDA, Agnew et al., 1976), the Global Digital Seismographic Network (GDSN, Peterson and Orsini, 1976) and GEOSCOPE (Romanowicz et al., 1984), have provided us with high-quality data and have led to the construction of the first three-dimensional models of the upper mantle and to refinements of these Earth models (Nakanishi and Anderson, 1984; Natal et al., 1984, 1986; Woodhouse and Dziewonski, 1984; Tanimoto and Anderson, 1985). The increasing number of phase velocity observations on direct mdividual source—station paths from several stations

with locations throughout the world has allowed us to enhance the resolution of tomographic models obtained from global studies by greatcircle analysis or direct analysis along the paths (Romanowicz, 1990; Montagner and Tanimoto, 1990, 1991; Roult et al., 1990; Zhang and Tanimoto, 1992). Numerous regional studies are possible; the Indian Ocean is now well documented (Montagner, 1986a; Roult et al., 1987; Montagner and Jobert, 1988), but regions of high southern latitudes have not yet been investigated in detail (see Roult and Rouland, 1994). The favourable distribution of GEOSCOPE stations in the southem hemisphere has provided us with an important data set; the first tomographic models of Antarctica and the surrounding regions have been presented in a preliminary paper with a limited data set of 213 paths (Rouland and Roult, 1992). A complementary and more precise investigation is presented in this paper, including the first results on anisotropy. The results ar~discussed in the framework of some important geodynamical problems.

2. Data collection

2.1. Seismological stations Seven GEOSCOPE stations with updated instrumentation installed in the southern hemisphere (Romanowicz et al., 1984), on sites operated by the IFRTP (Institut français de la

Table 1 Geographical coordinates of all GEOSCOPE stations used in our study Station

Location

Latitude

Longitude

Operational since

BNG CAN CAY CRZ DRV MBO NOC PAF PPT RER

Bangui, Central African Republic Canberra, Australia Cayenne, Guyana Crozet Islands Dumont d’Urville, Antarctica M’bour, Senegal Noumea, New Caledonia Kerguelen Island Papeete, Tahiti Reunion

4.43°N 35.32°S 4.95°N 46.43°S 66.67°S 14.39°N 22.28°S 49.35°S 17.57°S 21.16°S

18.55°E 149.00°E 52.32°W 51.86°E 140.01°E 16.96°W 166.43°E 70.21°E 149.58°W 55.75°E

1987 1987 1985 1986 1986 1985 1985 1983 1986 1982

G. Roult et aL /Physics of the Earth and Planetary Interiors 84 (1994) 33—57

Recherche et Technologie polaires), ORSTOM (Office de la Recherche Scientifique et Technique d’Outre Mer) and the IPGP (Institut de Physique du Globe de Paris), have provided us with long-period records of high signal-to-noise ratio. These stations and their geographical coordinates are listed in Table 1, and their locations are shown in Fig. 1, where we have represented the whole southern hemisphere and a part of the northern hemisphere (up to latitude 30°N).Some stations have been operating since 1982, such as RER (Reunion), or 1983, such as PAF (Kerguelen), and have already been used in regional studies on the southeast Indian Ocean (Rouland et al., 1985; Montagner, 1986b; Roult et al., 1987). In this study, we used three complementary GEOSCOPE stations located in the northern

35

hemisphere but near the equatorial meridian, MBO (M’bour) and BNG (Bangui) in Africa, and CAY (Cayenne) in South America, to add a few different paths to ensure a good spatial and azimuthal coverage. 2.2. Selection of data All events of magnitude larger than 5.5 with known centroid solution and located on the boundaries of the Antarctic plate, in the Chilean subduction zone, the Mid-Atlantic Ridge and the boundaries of the South Pacific Ocean plate were systematically used in this study when the corresponding minor arc crossed the Antarctic plate. More than 500 seismograms recorded at 10 GEOSCOPE stations from 1987 to 1989 have 0°

MB

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~?1Oo 1800 Fig. 1. Southern hemisphere and a part of the northern hemisphere (up to latitude 30°). Map of GEOSCOPE stations and events used in this study. The projection is a general stereographic projection centred on the South Pole. In all figures, we assume Greenwich meridian to represent the north direction at the South Pole.

36

G. Roult et al. /Physics of the Earth and Planetary Interiors 84 (1994) 33—57

been selected with these criteria, and after more restrictive selection (no redundancy and high signal-to-noise ratio), 399 paths were kept. The geographical distribution of these events is shown in Fig. 1. The map of the corresponding 399 propagation paths considered in this study is given in Fig. 2, but only latitudes south of 30°S are represented. We notice the good spatial and azimuthal coverage with numerous crossing paths. The southern hemisphere is fairly uniformly sampled, except in the Atlantic Ocean at low latitudes, where the path coverage density is lower; however, this region was not of major interest for this study. The Antarctic continent itself is very well sampled. Details of data acquisition and instrumental response have been described by Romanowicz et al. (1984). For the stations installed according to

the Strasbourg acquisition system (DRy, CRZ, NOC and PAF; Pillet et a!., 1990), we have used the high-gain long-period channel recorded with a 1 s sampling rate. Fig. 3(a) gives an example of signals recorded on the vertical component at PAF with three different channels, i.e. the broad-band channel (BRB), the high-gain longperiod channel (HGLP) and the very long period channel (VLP), for the same Kermadec event of 14 May 1989. The corresponding response curves in acceleration are plotted in Fig. 3(b). In the period range used in this study (40—300 s), the HGLP and VLP channels are very well suited. Fig. 4 gives examples of seismograms recorded on the vertical-component HGLP at the Antarctic DRV station (Dumont d’Urville, Terre Adélie) for various intermediate size events, at epicentral distances ranging from 2000 to 9500 km. The visible difference in Rayleigh train waveforms is

00

N

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180° Fig. 2. Distribution of geographical paths used in this study(399 paths).

G. Roult et al. /Physics of the Earth and Planetary Interiors 84 (1994) 33—57

37

PAF

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Fig. 3. (a) Examples of vertical records obtained for the same event (Kermadec event of 14 May 1989), at the same station (PAF), and for three channels (from top to bottom: BRB, broad-band; FIGLP, high-gain long-period; VLP, very long period). (b) Corresponding transfer functions (in digits ~m s—i s —

clear evidence for the existence of large lateral heterogeneities in this area. The earthquakes used are events of magnitude ranging from 5.5 to 6.5, and all the selected records have a high signalto-noise ratio.

3. Data processing For each record, both the fundamental mode group and phase velocities are computed. However, only phase velocities will be presented in

38

G. Roult et al. /Physics of the Earth and Planetary Interiors 84 (1994) 33—57

this paper. The processing consists of computing the synthetic seismogram by normal mode summation, as described by Woodhouse and Girnius (1982), taking account of the instrumental response, and using the focal parameters given by the Harvard centroid moment tensor (Dziewonski and Woodhouse, 1983). We then determine the difference between the observed phase and the computed phase, according to a technique similar to that used by Suetsugu and Nakanishi (1985). At this stage, some seismograms are rejected because of an evident discrepancy between the observed and the synthetic seismograms; this is often caused by the presence of higher modes. An example of the processing is given in Fig. 5. In Fig. 5(a) are plotted two traces—at the top the observed seismogram filtered in the period range 70—400 s, and at the bottom the corresponding synthetic one computed in the case of a spherically symmetrical transversely isotropic Earth de-

40—

20

i

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I

4. Geographical distribution of phase velocities

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The phase velocities along each path have been computed by the process described above, in the period range 60—300 s. This data set can be

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scribed by the Preliminary Reference Earth Model (PREM) (Dziewonski and Anderson, 1981). A phase shift is clearly visible between the traces: the observed signal is in advance compared with the synthetic one, which corresponds to higher velocities than those of PREM, as can be seen in Fig. 5(b). This result seems normal because the corresponding path between the Bouvet Islands (52°S,13°E)and DRY is essentially continental. These phase velocities, which are particularly high, are similar to those for cratons, as found by Knopoff (1972) in other parts of the world.

I

FEB 05 ~36), 19881

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2000

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TIME

,

sec

Fig. 4. Examples of vertical records at DRy, from HGLP channel, corresponding to intermediate-size earthquakes of epicentral distances ranging from 2000 to 9500 km. (Note that the second record (Sandwich islands—DRV) corresponds to a typical continental path and that the third (Indian Ocean—DRV) corresponds to a composite path.)

G. Roult et al. /Physics of the Earth and Planetary Interiors 84 (1994) 33—57

regionalized, to locate the lateral heterogeneities of phase velocities. Two approaches are possible —either a discrete parametrization in terms of a spherical harmonics expansion, as adopted by Woodhouse and Dziewonski (1984) and Natal et al. (1986), or a continuous parametrization, as developed by Montagner (1986a) and Montagner and Nataf (1988). For regional investigations, the second approach is appropriate; it has already

A ~

-

39

been used by Montagner (1986b) and Roult et al. (1987) for a study on the Indian Ocean region. The general procedure, without a priori constraints and based on the Tarantola and Valette algorithm (1982), makes it possible to retrieve simultaneously the spatial distribution of local phase velocities and azimuthal anisotropy. Instead of parametrization using a set of basis functions such as spherical harmonics, this method

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period (s) -

Fig. 5. Bouvet Islands event of 30 July 1989, recorded at DRy on the vertical VLP component: (a) Top: comparison of the observed seismogram (surface-wave train Ri), at the top, and the corresponding synthetic seismogram (computed for a spherical Earth represented by PREM), at the bottom. The latter is seen to be late. (b). Bottom: corresponding phase velocities. The continuous line corresponds to the synthetic signal and the dashed line to the observed one (higher velocities).

40

G. Roult et al. /Physics of the Earth and Planetary Interiors 84 (1994) 33—57

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G. Roult et aL/Physics of the Earth and Planetary Interiors 84 (1994) 33—57

only requires the definition of a covariance function which acts as a spatial filter. This covariance function between P and Q depends on a correlation length L~and on the a priori uncertainty of the model at each point, ff(P) and o(Q). It is defined as C1 ~ =

o-(P)

x o-(Q) x exp[(cos

A

—

41

layers (related to bathymetry, topography, sediments and crustal thickness); we know that this may introduce some bias, especially at short penods (60 s), for which the errors on phase velocities can reach 0.07 km s~. We have increased the number of paths across Antarctica to achieve apaths. map with coverage, with data set both of 399a In a better first experiment we acompare

1)/Lfl

posteriori error maps obtained at the same period, T 76s, from regionalization processing with the same correlation length of 1730 km, for both our old and the new data sets. The distribution of the corresponding a posteriori error maps is shown in Fig. 6. The error distribution maps are uniform, and of course the result is better when the path coverage is the greatest, especially in the region of DRY, where many paths cross. The computed errors do not exceed 0.08 km s~ in the area considered, which corresponds to a maximum error of about 2% of the average value (at short periods). The robustness of our inversion is tested and proved by the fact that the same general trends are obtained in- both cases on the regionalized velocity maps. =

where P and Q are two points on the Earth’s surface, A is the distance between P and Q, L,~ is the correlation length, and o~(N)is the a priori error of the model at the point N. The solution depends on the chosen correlation length, and, as in all inversion methods, there is a trade-off between the errors and the obtained resolution. Therefore the analysis of the sensitivity of the final model with respect to the choice of the correlation length is very important. Various correlation lengths, from 500 km to 2000 km, have been tested. The optimum correlation length depends on the surface of the area under investigation, on the number of data and on the number of azimuthal terms to be inverted (one, if we do not resolve azimuthal anisotropy, and three or five, if we want to resolve azimuthal anisotropy). In the isotropic case, a correlation length of 500 km is still resolvable and will allow us to refine detailed structures. In this study we have used, in many cases, correlation lengths of 1000 km or 1730 km, for which the location of velocity anomalies is stable. These correlation lengths correspond to a large degree of redundancy, display robust long-wavelength heterogeneities and allow comparison of our results with global tomographies (Montagner and Tanimoto, 1990, 1991; Zhang and Tanimoto, 1992). Phase velocities are calculated on a 5°x 5°grid and the results are represented in a general stereographic projection with the South Pole at the centre on all figures. In previous papers (Rouland and Roult, 1992; Roult and Rouland, 1994), we have presented results with a preliminary set of data (213 paths, with regionalization processing with a correlation length of 2000 km). In this first regionalization, we did not perform any correction for shallow -

-

4.1. Corrections for shallow layers It is well established that surface-wave velocities are sensitive to the uppermost layers of the Earth (Dziewonski, 1971; Souriau, 1976; Montagner and Jobert, 1981). Therefore corrections for shallow structure have to be computed carefully before any interpretation of the results. The lateral variations of parameters for shallow layers are very large. Four parameters have to be taken into account—topography, ocean bathymetry, sediment thickness and Moho depth. The dominant one is crustal thickness, and numerous workers have taken only this simple correction into account (Woodhouse and Dziewonski, 1984). The dependence on shallow structure of partial derivatives of phase velocity with respect to elastic parameters is very complex and non-linear (Anderson and Dziewonski, 1982). The corrections made by using a linear perturbation process (Woodhouse and Dziewonski, 1984; Nataf et al., 1986) are insufficient even at long periods; it is necessary to take into account other parameters

42

G. Roult et al. /Physics of the Earth and Planetary Interiors 84 (1994) 33—57 .0

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such as ocean depth, topography and sediment thickness, and the non-linearity introduced by the structural difference between oceans and continents (Montagner and Jobert, 1988). In our study, we chose to perform the shallow layer corrections on the direct path phase velocity data, rather than on the regionalized phase velocity distributions, in accordance with the results of Montagner and Jobert (1988), who proved that this is the only correct way to make these corrections. The processing for shallow layer corrections was extensively described in the appendix of Montagner and Jobert (1988) and by Montagner and Tanimoto (1991), and we only describe briefly here the principles of the technique. Three reference models (ocean, continent and average) are considered, which differ only in their crustal discontinuities (bathymetry, topography, thickness of sediments, two intermediate discontinuities, and depth of Moho discontinuity). A different dispersion curve is associated with each reference model; the correction is applied to each set of data. This procedure allows us to take account of the non-linear correction for the difference between oceans and continents. It is assumed that it is possible to choose the right reference model everywhere. When points are clearly continental (or oceanic), a linear correction is calculated with respect to the continental (or oceanic) reference model and added to the non-linear correction between the continental and the intermediate reference model. The influence of shallow layers on long-period dispersion is evident, and we have calculated the amount of contamination of deep structure by shallow layers. This may reach 0.03 km s’ at a period of 300 s and 0.07 km s~ at 100 s, which is not negligible. The corrections for shallow layers tend to increase the contrast of lateral heterogeneities, as will be seen below. We have tested the effects of shallow layers at a period of 125 s, which corresponds to a depth of penetration of about 200 km, with the data set of 399 paths. Fig. 7 illustrates the distribution of phase velocities obtained before and after correction for surface layers. In both cases, the correlation length is 1730 km. We can see that the distribution of anomalies is not modified by correction for shal-

43

low layers, but the contrast between low and high velocities increases.

5. Long-wavelength regionalization 5.1. Phase velocity maps In Fig. 8 we have plotted four maps corresponding to regionalization results at four Selected periods, 76, 100, 125 and 166 s, after correction for surface layers and for a correlation length of 1730 km. We observe that the maps at low periods fit the general tectonic structure well. First, the velocity contrast between East and West Antarctica is clearly established. The highest yelocities correspond to East Antarctica, commonly recognized by geologists as a craton. The maximum observed value at a period of 76 s (Fig. 8(a)) is 4.16 km s’ and corresponds to a typical shield velocity on curves established by Knopoff (1972); it occurs in Enderby Land, where the oldest sampled rocks have been identified (Napier complex rocks; see Roult and Rouland (1994)). West Antarctica displays lower, but still continental, velocities. This region shows -an uniform low velocity distribution not limited to the region of the Antarctic mountains. Numerous geological and geophysical studies point to a mixed structure where tensional stresses might have been highly active for a long time, probably owing to the fact that West Antarctica has undergone several orogens (from 650 m.a. to 100 m.a.). The lowest velocities underline the succession of active tectonic zones as mid-oceanic ridges around the Antarctic continent which delimit the boundaries of the Antarctic plate, with increasing values from ridge axis to old oceans; the lowest phase velocity anomaly is found at T = lOOs and is located southwest of Macquarie Island. This anomaly is related to a broad and deep thermal anomaly of the asthenosphere, as suggested by some previous studies (Xu, 1984; Rouland et al., 1985). The same general trends appear on maps at periods of 100 s, 125 s and 166 s (Figs. 8(b), 8(c) and 8(d), respectively), but the amplitude of anomalies decreases as the period increases.

44

G. Roult et aL/Physics of the Earth and Planetary Interiors 84 (1994) 33—57

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Contrasts between the high velocities of the eastem part of Antarctica and the low velocities of the central parts of ridges are still observed. The correlation of velocities with the shallow tectonic structure remains important but begins to vanish, At long periods, the resolution of our regionalization deteriorates slightly (as a result of small events and low signal-to-noise ratio) and the a posteriori errors increase, but robust features still hold; for instance, East Antarctica remains a fast velocity region, and the contrast between the eastern and western regions is still visible.

47

The regionalization technique designed by Montagner (1986a) makes it possible to invert not only for phase velocity but also for geographical distribution of azimuthal anisotropy. Smith and Dahien (1973) demonstrated that for a slightly anisotropic medium, the local phase or group

of azimuthal anisotropy for the rest of the world, but in the present study, devoted to the southern polar areas, it was necessary to design a new technique which allows us to invert correctly for the azimuthal anisotropy. The solution of this problem is to rotate the coordinate system to suppress the particular role of the poles. For regional studies this is particularly simple, because it is always possible to find new poles outside the area under investigation. For instance, they can be located on the equator for polar studies. Each point of the study area will be characterized by new coordinates, a new north, N’, and a new east, E’, such that E’, N’ form a new direct trihedron. In that case, it is necessary to redefine the azimuth ‘I’ at each point along the paths. The new azimuth ‘I” is the angle between N’ and the direction of the path. Eq. (1) is still valid in the new coordinate system and does not change the regionalization technique. After obtaming the distributions a1 and a2, an inverse

velocity can be expanded to first order, as a Fourier series of the azimuth ‘I’ according to the relationship

rotation can be applied to recover the geographical coordinate system. Obviously, it will not be possible to calculate the azimuth at the southern

5.2. Anisotropy maps

~

(

‘

=

1T ‘

)

a0( 0, F, T)

+

a1 cos 2~I’

+a2sin2-)I~+a3cos4-)I’+a4sin4~ (1) where a0, a1, a2, a3 and a4 are the azimuthal anisotropy distributions, which also depend on 0, F, T. The term a0 is the azimuthally averaged phase velocity, 0 and 4 are the polar coordinates and T is the period~This equation, derived for a plane Earth, is also asymptotically valid for a spherical medium (Mochizuki, 1986), and has been applied by Montagner and Tanimoto (1990) in a global investigation of azimuthal anisotropy. Montagner and Anderson (1989) showed that, from petrological considerations, it is sufficient to invert for the first two distributions, a1 and a2, for Rayleigh waves, and for the last two distributions for Love waves. However, Montagner and Tanimoto (1990) noted that Eq. (1) is not well suited for polar areas, where the azimuth varies very rapidly, and is not valid at the pole itself. They showed that this limitation does not affect the determination

~‘,

pole. For a global study, the situation is slightly more complex, and the procedure described with one .rotation is not possible, because a single rotation will always define two new poles which are again singular points and the same problem arises with these poles. In that case, a solution can be found by defining a local coordinate systern at each point of the Earth, such that the poles are always at 90° from this point. The degree of freedom left is suppressed by applying the second rotation along the meridian of the point. The results of this technique will be presented later for a global study (J.P. Montagner, personal communication, 1993). According to Fig. 2, the azimuthal coverage is correct only for latitudes further south than 30°S, and therefore we present maps of azimuthal anisotropy for only these latitudes at various penods obtained for Rayleigh waves (Fig. 9). The first point is that the azimuthal anisotropy can be as large as several per cent in the whole period range 70—300 s, and therefore it is not a secondorder effect, in contrast to what is usually as.

48

G. Roult et aL /Physics of the Earth and Planetary Interiors 84 (1994) 33—57

sumed in tomographic studies. The second point is that the introduction of anisotropy is significant; if the same number of parameters is inverted for, the variance reduction, if we take account of azimuthal anisotropy, is more important (by almost 8%) at 62%. Azimuthal anisotropy is present not only below oceanic areas but also below continents, although it is smaller in the latter case. At the shortest periods, in oceanic areas, the direction of maximum velocity is orthogonal to most of the ridges (the Southern East Pacific Ridge, Mid-Atlantic Ridge and Southwest Indian Ridge) and in agreement with plate velocities. The only exception is the central part of the Southeast Indian Ridge, but this zone, including the Australia—Antarctica Discordance (AAD), is known to have a very anomalous structure, with high velocities below the lithosphere, which are indicative either of a cold spot or of the initiation of a subduction process (Montagner, 1986; Forsyth et al., 1987; Roult et aL, 1987; Zhang and Tanimoto, 1992). Within the Antarctic continent, the azimuthal anisotropy is small but still significant, and it is large around the Transantarctic Mountains. However, it is difficult to know at this stage whether this anisotropy is related to a fossil strain field emplaced during the tectonically active periods or is related to the present flow pattern below Antarctica. In a second stage, we should obtain fundamental information on the processes involved at depth, which will allow us to accommodate the continuing increase of the surface of the Antarctic plate. The inversion at depth of the azimuthal anisotropy distributions, according to the technique described by Montagner and Nataf (1986), is currently in progress and will be presented elsewhere.

6. Short-wavelength regionalization of phase yelocity The good path coverage allows us to obtain velocity distribution maps with shorter correlation lengths (down to 500 km). For such correlation lengths, the variance reduction is still increased; its absolute value is up to 60%, which

gives us confidence in our results. The improvement is particularly evident at short periods. The results with a correlation length of 1000 km are presented in Fig. 10, for two periods (76 s and 200 s); the regionalized maps show more detailed patterns than before. Results obtained with a correlation length of 500 km (the lower limit for the period range used) are presented for the same two periods in Fig. 11. The corresponding a posteriori errors are, of course, higher; they reach values of 4% or 5%, instead of 1.5% in the case of a correlation length of 1730 kin, but their distribution maps are uniform and the anomalies are still well resolved. They are presented in Fig. 12. Both Figs. 10 and 11 allow detailed interpretation. It is important to note the large phase velocity gradient from the western part of Antarctica to the eastern part, with continental values for both regions and typical shield values in the eastern part. Instead of one velocity maximum obtained with a 1730 km correlation length, two zones of maximum velocities appear, one in Enderby Land and the other in the region of Terre Adélie. This result agrees with the age of 3.8 Ga obtained for rocks found in Enderby Land, and also with dating in the Terre Adélie region, where rocks of 1.7—2.4 Ga have been collected (0. Monnier, personal communication, 1993). Global results of Zhang and Tanimoto (1992) show the same two zones of high velocities in the eastern part of Antarctica at depth h = 90km, as will be seen below in Fig. 14. The Transantarctic Mountains are also characterized by high velocities at T = 76s, but this is less pronounced at 200 s. The craton areas and the mountain range are separated by a zone of average velocity which has no surface tectonic signature. Low velocity anomalies again correlate very well with the active tectonic regions, including mid-ocean ridges, which underline the boundaries between plates all around Antarctica. On the Antarctic continent itself, between the westem part and the eastern part, a zone of lower velocity could correspond to the beginning of a rift zone along the western flank of the Transantarctic Mountains. In the Ross Sea, there is a region of active volcanism (e.g. Mount Ere-

G. Roult et aL /Physics of the Earth and Planetary Interiors 84 (1994) 33—57

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bus), which is seen as a hotspot. Recently, an-

might be expected for a spreading centre above

other volcano, situated not very far from the South Pole, has become active (Blankenship et al., 1993). This volcanic activity along the western part of the Tnansantarctic Mountains could be seen as a manifestation of rifling, and our results may help in finding the answer to this important question. Such a rifting process could imply a maximum horizontal extensional axis in the SW— NE direction, which corresponds to the direction of ancient orogens. Following recent geophysical and geological investigations in the Ross Sea basins, several workers have documented the evident current rifting of the West Antarctic Rift system (Stern and Ten Brink, 1989; LeMasurier, 1990; Behrendt et al., 1991). Regions with particularly slow velocities are observed in the South Atlantic Ridge, the East Pacific Rise and the Southeast Indian Ocean Ridge. In Fig. 11, the lowest velocities appear in some regions corresponding to hotspots (Crozet, Kerguelen, Macquarie and Balleny Islands). However, one low velocity area (50°S,105°E)has an unknown origin. The low velocities under the Indian Ocean Ridge are a little shifted from the current position of ridges, as hotspots often are in this region. A region of high velocities is found in the southern part of the Pacific Ocean; it could be related to the increase of velocity with the age of the sea-floor or could correspond to an ancient subduction zone. The South Sandwich subduction zone is not clearly marked, but there is, however, an important gradient connecting low velocities to high velocities from west to east, at the rear of the subducted plate. Another feature can be also noted: the anomabus region of the AAD (between 120°and 130°E) is clearly visible on our maps. Indeed, we observe slow velocities near the surface (T = 76s) but fast velocities at greater depth (T = 200s), corresponding to the deep cold anomaly of the AAD. This region is well documented (Weissel and Hayes, 1974; Vogt et al., 1983; Klein et al., 1988). From Rayleigh surface-wave analysis, the shear velocity throughout the upper 150—200 km of the mantle is known to be faster beneath the discordant zone than beneath a normal mid-ocean ridge (Montagner, 1986b; Forsyth et al., 1987). This

an unusually cool asthenosphere (Palmer et al., 1993). The Rayleigh phase velocities of our maps (Figs. 10 and 11) increase with increase in period, and the signature of the AAD then appears very cleanly. The high velocities are indicative either of a cold spot or of the initiation of a subduction process (Montagner-, 1986b; Forsyth et al., 1987; Roult et al., 1987; Zhang et al., 1991). On our maps, the subduction zone of the South Sandwich Islands is visible, but the signature of the South Shetland subduction zone (Grad et al., 1993) is less marked. The distribution of the corresponding a postenon errors for both correlation lengths (1000 km and 500 km) is shown in Fig. 12, at the same period (76 s). The error distribution maps are fairly uniform, and the results are better when the path coverage is greatest, especially in the region of DRy, where many paths cross. The computed errors do not exceed 0.16 km s~for a correlation length of 1000 km and 0.20 km s’ for a correlation length of 500 km, which correspond-to maximum errors of about 4% and 5% of the average value (at short periods), respectively. Fig. 13 shows a comparison of our results with those of Montagner and Tanimoto (1991) for the southern hemisphere (latitudes south of 60°S);in their study the correlation length is approximately 500 km, as in our results in Fig. 11. Their global tomographic models are in good agreement with our regional results, but the contrast that they obtained between high and low velocities is lower by a factor of two, as is often observed on smooth global models. The active tectonic zones (with low velocities) are less visible, and velocities in East Antarctica are lower. Nevertheless, the main features are observed; their number of paths across Antarctica is about the same as ours, but they used the first and second wave train of Rayleigh waves, Ri and R2. In our study we only used the first Rayleigh wave train, Ri, which may provide a better resolution. In the velocity maps of Zhang and Tanimoto (1992), the spherical harmonics have been ~xpanded up to degree 1 = 36. -Comparison of our results with their global Earth model shows no large discrepancy between the general features; -

53

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Fig. 14 represents the S-wave velocity deviations at two depths, 90 km and 190 km, which correspond approximately to our regionalization maps of Fig. 10(a) or Fig. 1 i(a) at a period of 76 s, and to our results at i25 s. Their contrast is lower than ours by a factor of four. At greater depths, the contrast is insignificant, around 2% at 290 km depth and 1% at 390 km depth. Compared with the results of Montagner and Tanimoto (1991) and Zhang and Tanimoto (1992), all our maps present a higher contrast between slow and fast phase velocities, as do all regional

are located precisely along the active tectonic zones, the ridges encircling the Antarctic conti-

studies; we know that there is a bias between the a priori errors in the data, and the contrast between high and low velocities is dependent on the a priori errors introduced. Nevertheless, the positions of anomalies remain constant and are not affected by this problem, proving the robustness of our regionalization. Our results are in good agreement with the existence of the supercontinent Gondwanaland (Du Toit, 1937; see Roult and Rouland, 1994); the high velocities found for the East Antarctic craton are also observed in the western part of Australia (Montagner, 1986b;

great depths. The gradient of velocities from the western part to the eastern part of Antarctica (with continental values in both parts) is a clear signature of the various orogens of West Antarctica. The elongated low velocity on the western flank of the Transantarctic Mountains might be related to the existence of a rift zone similar to the African rift. A very important question is to understand the geodynamic behaviour of the Antarctic plate. It is a growing plate in extension, encircled by ridges all around its perimeter, except near the South

Roult et al., 1987), the Brazilian shield (Fouda, 1973), and the South African shield (Hofmann and Weber, 1983; Groenewald et al., 1991). These

Sandwich Islands and South Shetland Islands (Grad et al., 1993), where we see the only subduetion zone of the Antarctic plate boundaries. We

regions correspond to the oldest part of Gondwanaland. By using the high velocity areas, a precise reconstruction of Gondwanaland might

find evidence of zones of rifting in the Antarctic continent itself. A similar study using Love wave trains with determination of the corresponding

be possible, following the same approach as Tanimoto (1987).

attenuation maps, and with simultaneous inversion at depth of both Rayleigh and Love wave phase velocities, will be the subject of a forthcoming paper, and should provide new insight into the three-dimensional distribution of velocity and

7. Conclusion

anisotropy within the Antarctic continent.

nent. The lowest velocity anomalies correspond to the location of hotspots (Marion, Prince Edward, Crozet, Kerguelen, Macquanie and Balleny Islands). Another, less pronounced, low-velocity structure zone is observed, located in the vicinity

of the Drake Passage. Because of the great number of paths, the detailed maps obtained allow us to distinguish new details. The AAD is perfectly located and clearly seen as a cold region at large period or

The availability of high-quality data obtained

with the GEOSCOPE network in high southern latitudes has allowed us to construct high-resolution maps of the lateral heterogeneities of Antarctica (500 km) and of anisotropy in this region. The location of velocity anomalies is well resolved in our study. The eastern part of Antarctica shows a craton-like structure down to depths of about 250 km, and the highest velocities are observed in Enderby Land, where samples of the oldest rocks have been found. The slow velocities

Acknowledgements We thank the GEOSCOPE team for providing us with the data, and the operators in TAAF— ORSTOM stations of the southern hemisphere and Africa. We thank Olivier Monnier for fruitful discussions, and Jeannot Trampert for reading this manuscript. This is IPGP Contribution 1303.

56

G. Roult et aL /Physics of the Earth and Planetary Interiors 84 (1994) 33—57

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