Calibration of shear-wave splitting in the subcontinental upper mantle beneath active orogenic belts using ultramafic xenoliths from the canadian cordillera and alaska

TECTONOPHYSICS ELSEVIER

Tectonophysics 239 (1994) l-27

Calibration of shear-wave splitting in the subcontinental upper mantle beneath active erogenic belts using ultramafic xenoliths from the Canadian Cordillera and Alaska Shaocheng Ji a*, Xiaoou Zhao a, Don Francis b aD.6partement de gkologie, Uniuersite’de Mot&al, C.P. 6128, Succursale Centre-Ville, Montrkal, Qut H3C 3J7, Canada b Department of Geological Sciences, McGill University, 3450 Unicersity, Mont&al, Q&. H3A 2A 7, Canada

Received 31 January 1994; revised version accepted 5 July 1994

Abstract Despite the abundance of measurements of shear-wave splitting around the world, some fundamental questions about its geological interpretation have not yet been answered. In order to constrain (i) the orientation and magnitude of S-wave anisotropy, (ii) the thickness of the anisotropic layer and (iii) the possible variation of the anisotropy with depth in the subcontinental upper mantle beneath active erogenic belts, we have carried out a systematic investigation on shear-wave properties of mantle xenoliths from recent alkaline basalts of the Canadian Cordillera and Alaska. It is found that the polarization direction ($1 of the fast S-wave as well as the time delay (St) between the two arrivals are strongly dependent on the propagation direction with respect to the structural frame. At a single station, the variations of measured 4 and 6t values for SKS phases from different events can result from this dependence. Because the average grain size of olivine in upper mantle samples is commonly larger than 0.5 mm, dislocation creep prevails over a thickness of at least 250 km and results in olivine LPO and hence anisotropy and shear-wave splitting. A single thick anisotropic layer in mountain-parallel, subvertical shear zones may cause high values of St, while multiple, subhorizontal layers will produce multiple splitting, making the measurement and interpretation of S-wave splitting difficult. Our results also suggest that the S-wave splitting observed in cold, old, stable cratons such as the Canadian Shield is likely dominated by fossil anisotropy, whereas that observed in hot, active erogenic regions, such as the Northwest American Cordillera and the Tibetan Plateau, may be dominated by modern tectonic deformation in the lithosphere and flow in the asthenosphere.

1. Introduction

A propagating shear-wave (S-wave) splits into two S-waves which propagate at different velocities with orthogonal polarizations. This physical

* Corresponding author.

phenomenon is called acoustoelastic birefringence in materials sciences. Because most of the polycrystalline materials have a lattice preferred orientation (LPO) developed during their formation and particularly plastic deformation, they tend to be anisotropic. Thus, acoustoelastic birefringence has been used for a long time as a technique for non-destructive evaluation of tex-

0040.1951/94/$07.00 0 1994 Elsevier Science B.V. All rights reserved SSDI 0040-1951(94)00139-l

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S. Ji et al. / Tectonophysics 239 (1994) 1-27

ture in materials. Nur and Simons (1969) and Christensen (1971) were among the first to observe experimentally acoustoelastic birefringence in naturally deformed rocks. Crampin (1981) termed this phenomenon S-wave splitting. In-situ observations of LPO-induced S-wave splitting in the Earth did not begin until the pioneer work of Ando et al. (1983) using digital three-component seismograph stations. Due to the fact that steep incidence of S-waves is a pre-requisite for the observation of S-wave splitting, however, the Swaves emitted directly from earthquake focuses can be used only for studying crack-induced seismic anisotropy in the brittle upper crust (e.g., Crampin et al., 1984; Tanimoto and Anderson, 1984; Kaneshima et al., 1988; Savage et al., 1990a; Shih and Meyer, 1990) and in upper mantle wedges above Wadati-Benioff zones (Ando et al., 1983; Shih et al., 1991). In order to be able to study LPO-induced seismic anisotropy beneath the continents, Vinnik et al. (1984) measured S-wave splitting in SKS teleseismic phases. Since then, there has been considerable progress made in the measurement of S-wave splitting in all continents (Silver and Chan, 1988, 1991; Savage et al., 1990b; Vinnik et al., 1990, 1992, 1994; Milev and Vinnik, 1991; Makeyeva et al., 1992; Savage and Silver, 1993; Silver and Kaneshima, 1993). Depth resolution of SKS splitting observations is poor because of a trade-off between the thickness of the anisotropic layer and magnitude of its anisotropy. In principle, anisotropy can be anywhere along the final leg of the SKS path from the core-mantle boundary to the surface. The subcontinental upper mantle (about 350 km thick) is generally assumed to be the major source for S-wave anisotropy in SKS phases from the coremantle boundary and the surface (Silver and Chan, 1988, 1991; Ji and Salisbury, 1993; Mainprice and Silver, 1993). Seismic anisotropy in the subcontinental upper mantle has been known from measurements of Pn velocity (e.g., Babuska

et al., 1993) and of azimuthal and polarization anisotropy in long-period surface waves (e.g., Montagner and Tanimoto, 1991). The constituent minerals of the upper mantle such as olivine, orthopyroxene and clinopyroxene are very anisotropic. The maximum S-wave anisotropy is as high as 21.6% for olivine, 17.0% for orthopyroxene and 26.6% for clinopyroxene single crystals (Fig. 1). However, properties of S-wave are much more complex with respect to crystallographic axes than those of P-waves. In olivine, for example, there is a single VP maximum (9.6 km/s) which is parallel to the [loo] direction while there are two S-wave anisotropy [ A( maxima (> 17%) which correspond to the directions near [loll and [loll, respectively. The A(V,) for a given propagation direction is defined by the following equation (Mainprice and Silver, 1993):

In olivine, the angle between the V, maximum and minimum is 90” while that between the V, maximum and minimum is only about 45”. Similar situations occur for orthopyroxene; the single vp minimum (6.5 km/s) is parallel to the [OlO] direction while three A(V,) minima occur in the directions of the a-, b- and c-axes. This complex relationship in single crystals makes the tectonic interpretation of observed S-wave splitting more complicated than that of P-wave anisotropy. Mantle xenoliths from the Canadian Cordillera and Alaska were chosen for the present study in order to calibrate the composition, deformation mechanism, LPO of minerals and seismic properties of the upper mantle beneath modern active erogenic region. 2. Mantle xenoliths from the Canadian

Cordillera

and Alaska

A series of isolated late Tertiary to post-glacial alkaline volcanic centers are distributed along the

Fig. 1. P-wave velocity surfaces and S-wave anisotropies of olivine, orthopyroxene, clinopyroxene and spine1 single 1000°C and 1.4 GPa (corresponding to about 50 km depth). Equal-area projection of the lower hemisphere.

crystals

at

S. Ji et al. / Tcctonophysics 239 (1994) l-27

UIivine

CONTOURS (Kin/t)

Orthopyroxene

:oNTouw : 9

1:

riooi Clinopyroxene

Spifld

(X)

4

S. Ji et al. / Tectonophysics 239 (I 994) I-27

Northwest American Cordillera accreted terranes in British Columbia, the Yukon and Southern Alaska. These volcanic centers are thought to be related to local pull-apart basins associated with dextral transcurrent fault zones parallel to the boundary between the Pacific and North American plates (Souther, 1977). Samples of the mantle xenoliths studied in the present paper were collected from three localities: Castle Rock, Alligator Lake and Nunivak Island (Fig. 2). 2.1. Castle Rock The Castle Rock cinder cone (57”50’N, 130”35’W> is located northeast of the Mount Edziza volcanic complex, 52 km east of Telegraph Creek in northern British Columbia. Rounded ultramafic nodules 5-20 cm in diameter can be found in a Quaternary (Nixon, 1987) volcanic breccia made up of sub-round fragments of a fine-grained alkali basalt in a matrix of palago-

nite. In the Castle Rock suite, the mantle xenolith population is dominated by spine1 lherzolite (S&l%), with lesser quantitites of harzburgite (1.3%), orthopyroxenite (7.8%1, websterite (1.3%), plagioclase lherzolite (0.9%) and clinopyroxenite (0.6%). The average modal composition of 70 mantle xenolith samples from Castle Rock is: 64.2% olivine, 23.7% orthopyroxene, 10.6% clinopyroxene and 1.5% spine1 (Prescott, 1983). The xenoliths from Castle Rock display textures ranging from porphyroclastic (partially recrystallized; e.g., CR7, CR24, CR27 and CR651 to fine-grained granuloblastic (entirely recrystallized; e.g., CR34). The extensive recrystallization texture, indicative of large finite strains (e.g., Mercier and Nicolas, 1975; Karato, 19841, is believed to originate from ductile deformation in the subcontinental lithosphere, rather than from the sampling process by the host basalt (see discussion in 2.4). Typically, the olivine grain size in porphyroclastic xenoliths (e.g., CR651 ranges from

Fig. 2. Simplified tectonic map of Northwest America and locations of the studied mantle samples. Shear-wave splitting parameters (4 and at) have been measured at stations COL, LON and RSNT by Silver and Chan (1991) and at other stations by Bostock and Cassidy (in press).

S. Ji et al. / Tectonophysics 239 (1994) l-27

0.1 to 2.4 mm with an average value of about 0.9 mm (Fig. 3b). However, in the completely recrystallized samples (e.g., CR341, olivine neoblasts are more restricted in grain size (0.1-1.5 mm) with an average value of about 0.6 mm. There-

(a)

fore, the inferred which proceeded in their alkaline according to the (1980). -

(W

CR34

t.

5

5o

CR65

45

d&013

d.O.585 mm

40 -

0

0.3

0.6

0.9

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1.5

1.5

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2.4

2.7

0

3

0.3

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0.3

0.5

0.9

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1.5

1.5

2.1

2.4

2.7

3

0

0.3

0.5

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1.5

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2.4

2.7

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2.4

2.7

3

Grab &a (mm)

Gnh mka (mm)

0

1.2

AL58 dAOO7mm

45 -

0

0.9

mm

Groh &a (mm)

Grab &a (mm)

w

flow stress in the upper mantle, the entrainment of the xenoliths host magmas, is about 8 MPa, geopiezometer of Karat0 et al.

1.5 2.1 2.4 2.7

3

3.3

Odn(In(nml)

Fig. 3. Histograms displaying the grain-size CR65. (c) AL58. (.d) N164. (e) NI56.

distribution

of olivine in some selected

samples

for fabric measurements.

(a) CR34. (b)

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S. Ji et al. / Tectonophysics 239 (1994) l-27

Equilibrium temperatures for the mantle xenoliths from Castle Rock, calculated using the geothermometer of Wells (1977), vary from 950 to 1100°C (Prescott, 1983). This range of temperatures corresponds to a range of pressures from 1.33 to 1.69 GPa (46-59 km) using Ranalli’s (1980) estimated geotherm for the Canadian Cordillera (see also Fig. 8a). This geotherm is constructed by combining estimates of the surface heat-flow (about 84 mW/m*, see also Lewis, 1991) with an assumed lithological profile (see also Clowes et al., 1992), and assuming that the temperature at a depth of 35 km beneath the Canadian Cordillera is 750°C (Caner, 1970). The estimated temperatures and pressures, as well as the mineralogical assemblage, indicate that the xenoliths from Castle Rock had equilibrated in the spine1 lherzolite stability field. 2.2. Alligator Lake The Alligator Lake (60”25’N, 135”25’W) volcanic complex is located about 30 km southwest of Whitehorse, Yukon Territory, lying within the Coast Plutonic Belt of the Canadian Cordillera, near its eastern contact with the Intermontane Belt (Tipper et al., 1981). The Alligator Lake volcanic complex is of Quaternary age (EichC et al., 1987) and is dominated by alkali olivine basalts and basanites. All lavas containing more than 9 wt.% MgO are rich in fragments of spine1 lherzolite xenoliths and xenocrystals derived from their disaggregation (Francis, 1987). The spectrum of Alligator Lake xenoliths ranges from relatively fertile pyroxene-rich spine1 lherzolite to relatively depleted olivine-rich harzburgite. The average modal composition is: 65.6% olivine, 26.2% orthopyroxene, 6.3% clinopyroxene and 1.9% spine1 (Francis, 1987). Porphyroclastic textures with remarkable foliations are common in the lherzolites from the Alligator Lake locality (e.g., samples AL40 and AL58), but tabular textures with elongated olivine grains are observed in some lherzolite samples (e.g., AL88). Some harzburgites, such as AL41 may exhibit little or relatively poor foliations and have granuloblastic equant texture (Mercier and Nicolas, 1975). These microstructural features in-

dicate that the peridotites in the upper mantle, similar to the quartzofeldspathic rocks in the deep crust, are deformed by cyclic dislocation glide, dynamic recrystallization and grain growth processes. Although there is a tendency for the grain size of olivine to decrease with increasing pyroxene content, the suite as a whole is more fine grained (0.3-2.7 mm) than the Nunivak Island suite (Fig. 3). Orthopyroxene-clinopyroxenespine1 clusters are found to be completely dispersed in the olivine matrix, indicating that the rocks have been subjected a very large plastic strain (Nicolas, 1986). Hydrous minerals such as amphibole and phlogopite are generally absent, indicating that the upper mantle beneath the Alligator Lake locality from which the xenoliths were carried is undersaturated with water. Fe/ Mg partitioning between olivine and spine1 and the Cr content of spine1 (Francis, 1987) indicate that most of the xenoliths from Alligator Lake have equilibrated at similar sub-solidus temperatures (940 _t 8000, calculated after the method of Fabries (1979). The miscibility gap between coexisting orthopyroxene and clinopyroxene yields slightly higher temperatures (1000 f 100°C Francis, 1987) using the geothermometer of Wells (1977). These temperatures and corresponding pressures (1.25-1.69 GPa), calculated using Ranalli’s (1980) Cordilleran geotherm, indicate that all samples from the Alligator Lake fall in the spine1 lherzolite stability field. However, the Si content and rare earth element profiles of the primary alkaline magmas at Alligator Lake suggest that they were derived from a deeper garnet-bearing source (EichC et al., 1987; Francis, 1987), and not produced from the mantle source represented by the xenolith population. 2.3. Nunivak Island Nunivak Island (60”N, 166”W), located approximately 50 km from the west coast of Alaska, is one of a number of late Tertiary to Recent (< 6.1 Ma), basaltic eruption centers, located behind the Aleutian Arc on the Bering Sea shelf (Fig. 2). Olivine tholeiite basalts dominate, but numerous small cones of alkaline basalts contain abundant xenoliths (Hoare et al., 1968). A count of more

S. Ji et al. / Tectonophysics 239 (1994) l-27

than 3000 ultramafic xenoliths from 30 vents revealed that the xenolith population is dominated by spine1 lherzolite (92.3%), with lesser quantitites of harzburgites (7.4%) and websterite (10.3%) (Francis 1976a,b, 1978). Average modal composition of the mantle xenoliths is: 68.2% olivine, 20.6% orthopyroxene, 9.7% clinopyroxene and 1.5% spine1 (Prescott, 19831. The mantle xenoliths from Nunivak Island, Castle Rock and Alligator Lake in the Northwest American Cordillera have lower olivine contents and higher clinopyroxene contents than the average modal compositions of oceanic uppermost mantle given by Dick (1987). His modal analyses of 266 peridotite samples dredged at six oceanic ridges gave an average composition of 74.8% olivine, 20.6% opx, 3.6% cpx, 0.5% sp and 0.5% pl. The xenoliths from Nunivak Island are characterized by transitional textures (e.g., N16 and NI76) between protogranular (e.g., NI56) and tabular (NI64.1 types. The tabular texture is considered to be formed by fluid-assisted grain boundary migration (e.g., Drury and Van Roermund, 1989). The olivine grains are relatively coarse with the average size generally larger than 1.0 mm. The protogranular and tabular xenoliths display similar olivine grain size distributions and mean grain sizes (Fig. 3d and e>. Subgrain boundaries (tilt walls) are developed in most olivine grains larger than 1.0 mm, but are usually absent in the grains smaller than 0.5 mm. Xenoliths having fine-grained mosaic texture are relatively rare at this locality. The equilibrium temperatures calculated using the geothermometer of Wells (1977) vary from 930 to 1040°C for the xenoliths from Nunivak Island (Prescott, 1983). This suggests that a rather restricted depth range of the upper mantle (45-53 km) was sampled by the alkaline basaltic magmas of Nunivak Island. Corresponding pressures are estimated to range from 1.29 to 1.51 GPa using Ranalli’s (1980) Cordilleran geotherm, with all samples fall in the spine1 lherzolite stability field. 2.4. Are the xenoliths representative of the upper mantle? Our lithologic and microstructural observations rely on a statistical study carried out on

7

several thousand samples examined in the field and more than a hundred representative samples studied in the laboratory from each locality with reference to other sites in the North American Cordillera. We assume that the observed modal composition, texture and fabrics of the xenoliths are representative of those of the in-situ upper mantle in the depth range 40-60 km at the time when they were picked up by the basaltic melts. This assumption is mainly based on the following arguments: (1) Xenolith-bearing magmas are expelled through conduits opened rapidly by hydraulic fracturing. The mantle rocks at the upper tips of the vertical hydrofractures where melt pressure exceeds the rupture strength of rocks are rapidly dismembered and carried to the surface by the ascending melts. Because the basaltic melts (p = 2.8 g/cm31 are less dense than the peridotite xenoliths (p = 3.3 g/cm31 the xenolith-bearing magma should ascend so rapidly that the xenoliths have no time to sink. Its ascent velocities were estimated to be about 5-10 km/h (Kushiro et al., 1976; Fujii and Scarfe, 1982). This means that it takes only several hours for the alkali basaltic magma to carry the spine1 peridotite xenoliths from the upper mantle to the surface. The time is too short for the xenoliths to have any significant changes in petrological composition and in microstructure. (2) Experiments of Brearley and Scarfe (1986) shows that a peridotite xenolith with an initial diameter of 10 cm will be dissolved completely after 1000 hours in a superheated (due to adiabatic decompression) alkali basaltic melt. This gives an upper bound for the life of xenoliths in the magma. (3) The volume fraction of xenoliths in the ascending magma is less than 2%. In such a dilute solid-fluid system, strain is strongly localized in the melt, and shear stress is too low to deform the solid xenoliths suspending in the melt. In addition, previous experimental studies (e.g., Nicolas and Poirier, 1976) indicated that a large strain (several tens of percents) is necessary to attain significant dynamic recrystallization in olivine. Therefore, it is not acceptable that the porphyroclastic texture of xenoliths was consid-

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S. Ji et al. / Tectonophysics 239 (1994) 1-27

ered to be formed by a short-term deformation during the volcanic eruption process (e.g., Goetze, 1975). (4) In the Archean cratons, the ultramafic xenoliths were extracted by Mesozoic kimberlite magmas long after the cessation of plastic deformation. However, the ultramafic xenoliths in the North American Cordillera were carried to the surface from the deforming upper mantle by recent (late Tertiary and Quaternary) basaltic volcanic eruptions and thus reflect the present state of underlying upper mantle. (5) In addition, the rapidly quenched ultramafic xenoliths in the recent volcanic eruptions are believed to be more representative, in terms of both composition and microstructure, of the subcontinental upper mantle than the ultramafic massifs of oceanic lithosphere which were formed at mid-ocean ridges under high-temperature and low-pressure conditions and transported to the surface by slower (lo’-lo* years) tectonic obduction (Suhr, 1993).

3. Petrofabrics Four to five representative xenolith samples from each locality were selected for LPO studies in order to estimate the overall seismic properties of the uppermost mantle under the Cordillera. Petrofabric measurements were carried out using a petrographical microscopy equipped with a five-axes U-stage. Within a given area of each thin section, all mineral grains except spine1 were measured, with each grain being located on a magnified photograph, so that no grains were repeated or omitted for measurement. LPO diagrams were plotted for each sample and contoured by computer using an unpublished program written by D. Mainprice.

3.1. Olivine Olivine LPO diagrams for each of thirteen representative xenolith samples from Castle Rock, Alligator Lake and Nunivak Island are shown in Fig. 4. All samples except AL41 (granuloblastic equant texture) and AL88 (tabular texture) display well-developed olivine fabrics with [lo01 crystallographic directions strongly aligned parallel or subparallel to the extension lineation (X>. In the xenoliths from Castle Rock and Nunivak Island, the LPOs are remarkably consistent from sample to sample. However, the xenoliths from these two localities show slightly different orientation patterns of olivine [OlO] and [OOll. The [OlO] crystallographic directions in the xenoliths from Nunivak Island form a single point maximum approximately perpendicular to the foliation (XY-plane), while in the xenoliths from Castle Rock the [OlO] directions form a complete (CR7, CR24, CR27, CR651 or almost complete (CR341 girdle approximately normal to the foliation and lineation. For both localities, the maximum concentration of [OlO] is always located near the normal to the foliation (Z-direction). In the xenoliths from Nunivak Island, the [OOl] directions are preferentially concentrated in an elliptic region whose long and short axes parallel to the X- and Z-directions, respectively. However, in the xenoliths from Castle Rock, these ellipses are elongated in the YZ-plane. In the both cases, the center of this elliptic region is located near the Y-axis. Compared with other samples from Castle Rock, olivine in CR34 has a smaller average grain size cd= 0.585 mm, Fig. 3a) and a stronger LPO. In the xenoliths from Alligator lake, the olivine LPOs (Fig. 4b) are significantly weaker than those of the Castle Rock and Nunivak Island xenoliths, but the average fabrics show a similar pattern to those of the xenoliths from Nunivak Island.

Fig. 4. Preferred orientations of olivine [loo], [OlOl and [OOl] directions in selected samples from Castle Rock (a), Alligator Lake (b) and Nunivak Island cc). The contours are 1, 2, 3 and 4% of the number of measurements per 1% of the hemispherical area. The highest contour interval is shaded. The number of measurements for each sample is given by the value of n. Equal-area, lower hemisphere projection. The XY-plane (foliation) is the N-S solid line and is perpendicular to the page; the X-direction (lineation) is N-S. W, q and oindicate the maximum density, the calculated best fabric axis and the pole to the best fabric plane, respectively (Bouchez et al., 1971).

S. Ji et al. / Tectonophysics 239 (1994) 1-27

CR65

n-139

All

n-591

S. Ji et al. / Tectonophysics 239 (1994) l-27

10

n-1 03

Fig. 4 (continued).

Olivine, whose volume fraction is higher than 60%, constitutes a stress-supporting continuous matrix in the Cordilleran subcontinental upper mantle. The strong olivine LPOs observed in the xenoliths suggest that dislocation creep prevailed

throughout the upper mantle from which the xenoliths were derived. Unlike face-centeredcubic metals, low-symmetry rock-forming minerals such as olivine and pyroxene do not have five independent slip systems to accommodate an ar-

S. Ji et al. / Tectonophysics 239 ( 1994) l-27

11

Nl64

Nl76

n-122

All

Fig. 4 (continued).

bitrarily imposed deformation by slip alone, and the critical resolved shear stress for slip varies greatly between different glide systems (e.g., Bai et al., 1991). These minerals deform by slip on only one or two easiest slip planes and directions

(Carter and AvC Lallement, 1970; Nicolas and Christensen, 1987), with dynamic recrystallization accommodation serving as a supplementary mechanism (e.g., Drury and Urai, 1990). Both theoretical (Wenk and Christie, 1991; Ribe and

S. Ji et al. / Tectonophysics 239 (I 994) I-27

12

Yu, 1991) and experimental studies (Nicolas et al., 1973; S. Zhang and Karato, 1993) clearly show that mineral LPOs form in polycrystalline aggregates as the result of the alignment of slip planes and directions, respectively, with the shear plane and direction during progressive rotational deformation. Obviously, the LPO pattern ([OlOl concentration parallel to the Z-axis) observed in the xenoliths from Nunivak Island and Alligator Lake may be explained as a result of [lOOl(OlO) slip, whereas the LPO pattern ([OlO] shows complete or almost complete girdle in the YZ-plane) in the xenoliths from Castle Rock as a result of

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olivine fabric has been reported in mantle nodules in kimberlites at Mothae, Thaba, Putsoa and Kimberley (South Africa) (Boullier and Nicolas, 1975) and in ultramafic massifs such as Lanzo (Peselnick et al., 1974), Wadi Bani Kharus (Oman), Table Mountain, Lewis Hills and North Arm Mountain (Newfoundland), Twin Sisters of Washington (USA), Dun Mountain and Red Hills (New Zealand) (Christensen, 1984). The latter olivine fabric has been observed in mantle xenoliths in France (Mercier and Nicolas, 1975), as well as in the Samail ophiolite (Oman), Troodos

10101

[lOOI

(a>

[ lOO](OlO)-dominant [ lOO](Okl) slips. The former

.--

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Lake

n-172

Nunivak

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Fig. 5. Preferred orientations of orthopyroxene Lake (b) and Nunivak Island (c). The number

[loo], [OlO] and [OOl] directions in selected samples from Castle Rock (a), Alligator of measurements for each sample is given by the value of n. Symbols as in Fig. 4.

S. Ji et al. / Tectonophysics 239 (1994) l-27

(Cyprus), Blow Me Down Mountain (Newfoundland) and Red Mountain (New Zealand) ultramafic massifs (Christensen, 1984). 3.2. Pyroxenes Composite fabric diagrams of orthopyroxene for the three Cordilleran xenolith localities are shown in Fig. 5. At Alligator Lake and Castle Rock, the orthopyroxene LPOs are very weak. At Nunivak Island, the orthopyroxene crystals show a tendency for the [OOl] crystallographic direction to be preferentially aligned in the foliation plane, with a maximum parallel to the lineation, and for the [loo] and [OlO] directions to be preferentially aligned in the YZ-plane. This fabric is similar to those for enstatite in other alkali basalt xenoliths (Mercier and Nicolas, 1975) and in kimberlite nodules such as samples M57, LMA2 and PHN1925 (Boullier, 1977). Clinopyroxene in the Cordilleran xenoliths (not shown) has an almost random fabric. Therefore, pyroxene, whether orthopyroxene or clinopyroxene, exhibits much weaker LPOs than olivine in the Cordilleran xenoliths. This contrast may be due to the fact that pyroxene is stronger, but volumically much less important, than olivine in the upper mantle. The pyroxene crystals act as rigid inclusions rotated in the ductile flowing olivine matrix (Ji and Zhao, 1993, 1994). However, the fabrics of orthopyroxene in the peridotite massifs reported by Christensen and Lundquist (1982), appear to be distinctly stronger than those in the Cordilleran ultramafic xenoliths. This difference may reflect the difference in the magnitude of flow stress between the ultramafic massifs and in-situ subcontinental upper mantle. The peridotite massifs were last deformed during tectonic emplacement under a much higher flow stress (100-300 MPa) than that found in the subcontinental upper mantle (< 50 MPa) beneath regions of high thermal flow, such as the Canadian Cordillera and Alaska. This contrast in orthopyroxene fabric is evidence that the Cordillera ultramafic xenoliths have not been affected by their emplacement mechanism, and thus retain the original microstructure of the subcontinental upper mantle.

13

4. Seismic wave properties The seismic velocities (V,, Vs,, vs.), shear-wave anisotropy and polarization direction of V,i of the Cordilleran xenoliths were calculated on the basis of the LPO, density, volume fraction and elastic stiffness coefficients of each of the constituent minerals after the method of Crosson and Lin (1971). The computations were performed using an interactive FORTRAN program (Mainprice, 1990; Mainprice and Humbert, in press). The Voigt average was used for velocity calculations because it gives the closest agreement between the calculated values of seismic velocities and those of laboratory measurements (e.g., Ji and Salisbury, 1993; Ji et al., 1993). For olivine, the experimental coefficients, including elastic constants [C,j(T,,P,)] at ambient temperature (25°C) and pressure (0.1 MPa), the first-order pressure (aCjj/aP> and temperature (aC,,/U) derivatives, the density [p(T,,P,)] at ambient conditions, the isotropic adiabatic bulk modulus (K > and the average thermal expansion coefficient (a) were taken from Kumazawa and Anderson (1969). Orthopyroxene’s coefficients were taken from Frisillo and Barsch (1972). Cij(TO,P,,) and K for clinopyroxene was taken from Levien et al. (19791, but aCij/aZ’ and aC,,/U are not presently available. Bass et al. (1981) estimated that the bulk modulus pressure derivative of clinopyroxene is about 7.9, smaller than the value of 9.6 found by Frisillo and Barsch (1972) for orthopyroxene. We assume, for want of accurate experimental data, that the clinopyroxene pressure derivatives of C, , , C,,, C,,, C,,, C,, and C,, (those C,, making up the bulk modulus pressure derivative) are equal to 82% (7.9/9.6) of the corresponding pressure derivatives of orthopyroxene. In contrast, C,,, C,, and C,, are assumed to have the same pressure derivatives in both clinopyroxene and orthopyroxene (Estey and Douglas, 1986). We also assume all the temperature derivatives of clinopyroxene to be equal to those of orthopyroxene. Spinel, being cubic, required three elastic moduli to completely describe its single-crystal elastic character. Spinel’s, C,j(To,Po>, p(T,,P,,) and K are taken from Weidner et al. (1984), XZ’,j/aP from Yoneda

S. Ji et ul. / Tectonophysics 239 (1994) I-27

14

symmetry axis is within 20”). In such a cone, the upper mantle may not be homogeneous in terms of composition and LPO. The ultramafic xenoliths of different petrographic and textural types occurring at each locality represent accidental pieces of the subcontinental upper mantle randomly scavenged by the basalt magmas. Unfortunately, there is no way to know the spatial distri-

(1990>, K,j/G from Askarpour et al. (19931 and N from Suzuki et al. (1979). As SKS teleseismic waves sample a large volume of the upper mantle beneath stations in subvertical propagation directions, the recorded S-wave splitting should reflect the average anisotropy in a vertical, circular mantle cone (the angle between the generatrix and the vertical

Castle Rock X

Allig+m X

Lake

Nunivak Island X

VP

Fig. 6. Average seismic properties of the upper mantle beneath Castle Rock, Alligator Lake and Nunivak Island at 1000°C and 1.4 GPa (corresponding to about 50 km depth), calculated from the average fabrics and modal compositions. d is the fast S-wave polarization direction. The delay time between the two S-wave arrivals (61) was calculated for a path length of 100 km.

S. Ji et al. / Tectonophysics Table 1 Elastic constants in GPa of the upper Canadian Cordillera and Alaska derived ultramafic xenoliths C!,

Castle Rock

Alligator

Cl, C22

228.05 195.91 201.65 65.00 71.06 69.61 66.26 68.07 -0.12 -0.12 - 0.94 68.25 ~ 0.08 0.11 -0.18 PO.37 0.02 0.36 0.07 0.12 -0.41

214.57 200.38 211.04 67.79 11.24 69.5 1 67.34 68.54 -0.09 0.38 0.40 68.01 0.50 0.26 0.94 - 0.43 0.49 0.02 0.56 0.20 0.01

c 33 c 44 CSS C66

Cl2 Cl.3

Cl4 C,S CM C23 c24

C ?S C 2h C 14 C 35 C Sh C 4s C 46 G6

Lake

mantle beneath the from average data of Nunuvak

Island

224.44 193.13 210.52 65.52 73.03 68.65 67.27 70.71 - 0.28 - 0.33 - 0.94 68.69 0.29 0.19 -0.88 0.57 -0.25 -0.25 -0.65 0.00 -0.05

bution of the xenolith types or the orientation of their foliation and lineation in the mantle prior to their entrainment by the basaltic magma which carried them to the surface. Therefore, the seismic properties calculated from the average modal composition and average LPO of the xenoliths are taken to be representative of the upper mantle beneath each locality. The average results for each locality are presented in contour diagrams in Fig. 6 and the resultant elastic constants (C,j) are given in Table 1. The three localities in the Canadian Cordillera and Alaska have remarkably similar seismic properties. V, has a unique maximum (8.3 km/s for Castle Rock and Nunivak Island and 8.1 km/s for Alligator Lake), which corresponds to the olivine u-axis maximum concentrations, parallel to the extension lineation (XI. Contours of Swave velocity anisotropy (Fig. 6) indicate a quasiorthorhombic geometry in which the largest Swave anisotropy, which coincides with the olivine c-axis maxima, occurs for propagating paths parallel to the Y-direction. Hence it may be inappropriate to consider the subcontinental upper man-

15

239 (1994) l-27

tle to have hexagonal anisotropies with a horizontal or vertical axis of symmetry. The calculated seismic geometry mainly reflects the olivine LPOs and the presence of pyroxene serves simply to reduce the bulk anisotropy of rocks by decreasing the volume fraction of olivine (also see Mainprice and Silver, 1993), because of the weak LPOs of pyroxene. The change from an (010) point maximum at Z for Nunivak Island to a YZ-girdle for Castle Rock results in a change of orientation of the maximum A(V,) from the XY-plane to somewhere within the YZplane, which may affect the interpretation of SKS observations. The magnitude of the delay time between the fast and slow arrivals depends not only on the magnitude of S-wave anisotropy, but also on the length of ray path. For the upper mantle beneath Castle Rock, the maximum value of A(t() is 5.0%, corresponding to 1.1 s delay time for a 100 km thick anisotropic layer. For each locality, the largest 6t occurs for the propagating path parallel to the Y-direction.

5. Discussion 5.1. Thickness mantle

of anisotropic

layer in the upper

There is a trade off between intrinsic S-wave anisotropy [A( and thickness of anisotropic layer in the upper mantle in interpretation of the time delay in SKS splitting. Studies of the seismic properties of mantle xenoliths which directly sample the subcontinental mantle provide direct evaluation of A(V). The ultramafic xenoliths from Northwest America in this study, and those from South Africa by Mainprice and Silver (1993), may be representative of the subcontinental mantle underneath active erogenic regions and stable cratonic regions, respectively. If so our findings indicate that the maximum S-wave anisotropy of the upper mantle is slightly higher in active orogenie regions (4.5%) than in stable cratonic regions (3.7%). It is necessary to study more samples from other continental areas, however, to verify whether this interpretation has general validity.

S. Ji et al. / Tectonophysics

16

We still need to know to what depth the anisotropic layer extends in the upper mantle. Previous theoretical analyses (e.g., Karat0 and Wu, 1993) and experiments (Karat0 et al., 1986; S. Zhang and Karato, 1993) show that LPO development in polycrystalline aggregates is controlled by the deformation mechanism. LPO develops where the deformation mechanism is dominated by recovery-accommodated dislocation creep. But no LPO develops when grain-size-sensitive diffusion creep prevails, resulting in an isotropic seismic structure. Therefore, to a first approximation, we assume that the lower boundary of the anisotropic layer coincides with the transition from dislocation creep to diffusion creep in the upper mantle. The transition depth from dislocation to diffusion creep can be qualitatively estimated using olivine flow laws determined experimentally by Karat0 et al. (1986). The flow law for each mechanism has the following form:

Ig(Strain-rate)=-13 (is)

1::: 10'

b

Diffusion

of olivine

as functions

of grain

creep

-I 400

600

600

1000

1200

1400

1600

1800

-f (“‘7 Fig. 7. Critical grain size for the transition between dislocation creep and grain-size-sensitive diffusion creep in olivine aggregates as a function of temperature. The data about olivine grain size and equilibrium temperature in naturally deformed mantle samples (Table 2) were plotted in boxes Z-13. The open and black boxes represent dislocation creep and diffusion creep, respectively.

is the activation energy, P is the pressure, V is the activation volume, T is the temperature in K, and R is the gas constant. For olivine in the dislocation creep regime, A = 3.5 X 1O22 SC’, n = 3.5, m = 0, E = 540 kJ/ mol and I/ = 15 cm3/ mol

where i is the strain-rate, A is a pre-exponential coefficient, u is the differential stress, p is the shear modulus (81 GPa), II is the stress exponent, m is the grain-size exponent, b is the length of the Burgers vector (0.5 nm), d is the grain size, E

Table 2 Deformation mechanism different continents

239 (1994) I-27

size, temperature

and pressure

conditions

in the mantle

xenoliths

from

Locality

Lithology

T (“0

P (GPa)

d (mm)

Mechanism

Source

2 3 4 5 6

Matsoku, South Africa Bulfontein, South Africa Monastery, South African Hanoba, Hebei, China Silong, Zhejiang, China Luihe, Jiangsu, China Jacques Lake, Canada Big Timothy Mt., Canada Nunivak Island, Alaska Alligator Lake, Canada Castle Rock, Canada Lesotho, South Africa Voltri massif, Italy

1050 900 900 980-1120 1200 1000-1200 1000-1130 900-1180 930-1040 900-1100 950-1100 1400 800-925

3.5 3.5 3.5 1.25-2.4 2.80 1.73-3.01

7 8 9 10 11 12 13

gt-peridotite gt-peridotite gt-peridotite sp-peridotite gt-peridotite harzburgite sp-peridotite sp-peridotite sp-peridotite sp-peridotite sp-peridotite gt-peridotite sp-peridotite

0.5-3.0 5.0-20 2.5-5.0 1 .O-5.0 0.5-5.0 0.2-4.0 0.5-5.0 2.0-10 0.1-3.6 0.3-3.0 0.1-2.4 0.07 0.01-0.03

Dis Dis Dis Dis Dis Dis Dis Dis Dis Dis Dis Dif Dis/Dif

Mainprice and Silver, 1993 Mainprice and Silver, 1993 Mainprice and Silver, 1993 Cong and Zhang, 1983 R. Zhang and Cong, 1984 Yang and Wang, 1985 Littlejohn and Greenwood, 1974 Ross, 1983

No.

I

Dis = dislocation

creep;

Dif = grain-size

sensitive

diffusion

1.25-2.60 1.29-1.51 1.25-1.69 1.33-1.69 5.8 0.6-0.8 creep.

This study This study This study Boullier and Gueguen, Drury et al., 1990

1975

S. Ji et al. / Tectonophysics

whereas at higher T and P, lower stress and smaller grain size, deformation occurs by diffusion. Table 2 summarized the data on olivine grain size and deformation temperature in the subcontinental mantle xenoliths from different continents. Most spine1 and garnet peridotite xenoliths from the continents, wherever active erogenic or stable cratonic regions, have olivine grain sizes ranging from 0.1 to 20 mm (with an average value between 1 and 2 mm), and have been equilibrated at temperatures between 850 and 1250°C. The commonly observed occurrences of strong

(Karat0 et al., 1986). For olivine in the diffusion creep regime, A = 8.7 X 1015 s-l, n = 1.0, m = 2.5, E =300 kJ/mol and I/= 6 cm3/mol (see Karat0 and Wu, 1993). Based on the assumption that at a given T, P, grain size and strain-rate, the mechanism which gives the lowest flow stress is the dominant deformation mechanism, we calculated the critical grain size for the transition between dislocation creep and diffusion creep as a function of temperature in the upper mantle (Fig. 7). As shown in Fig. 7, line ab separates two domains; at lower T and P, higher stress and larger grain size, dislocation creep dominates,

(4

1m

17

239 (1994) 1-27

tb)

. Canadlan Cordlllera

‘aoo: Canadlan Shield

15cQ-

1500-

12UO-

1200G e

G L900

I-

lSOO-

300-

0

50

100

150

200

250

300

350

0

50

100

150

200

250

300

350

Death

DepthWm)

lo3

(d)

102

-z

10'

E. 0

; (D

10

c -6

10-l

6 10-2

10-3 35

85

135

185 ~pth(W

235

285

335

10-3 35

a5

135

1.35

235

285

335

Depth(h)

Fig. 8. Variations of temperature and critical grain size for the transition between dislocation creep and grain-size-sensitive diffusion creep in olivine aggregates with depth in the Canadian Cordillera (a and c) and Canadian Shield (b and d). The strain rate is assumed to be 10W’3/s for the Canadian Cordillera and 1Om’5/s for the Canadian Shield.

1X

S. Ji et al. / Tectonophysics

LPOs (Fig. 4) and dislocation recovery structures, such as subgrain boundaries in olivine grains, within mantle xenoliths suggest that dislocation creep is important in the mantle from which they were sampled by magma. As shown in Fig. 7, these peridotite xenoliths are located well within the dislocation creep field (above line ab) predicted by extrapolating the laboratory flow laws of Karat0 et al. (1986) to natural strain-rates (e.g., IO-‘-‘/s for active erogenic zones). Boullier and Gueguen (1975) reported that grain-size-sensitive diffusion creep mechanism occurs in fine-grained (- 0.07 mm) garnet peridotite xenoliths deformed at very high temperature (- 1400°C). Drury et al. (1990) also showed that microstructures in very fine grained spine1 peridotite mylonites from the Voltri massif (Italy) are consistent with a transition of deformation mechanism from dislocation creep to some type of grainsize-sensitive diffusion creep. There is thus a good correspondence between theoretical models and the features observed in mantle peridotites. In Fig. SC and d the relationship is shown between critical grain size and deformation mechanism in the upper mantle beneath the active erogenic belts (Canadian Cordillera) and stable cratons (Canadian Shield). Seismic reflection profiling (Clowes et al., 1992) suggests that the Moho occurs at 35 km in both regions. The temperature profile for the Canadian Shield (Fig. 8b) was estimated using a conductive thermal model based on the observed geotherm and the relatively low heat flow (40 mW/m*) observed in the Abitibi region of Canadian Shield. The Canadian shield has a much lower geothermal gradient than the Cordillera. This difference is clearly reflected in seismic velocities and attenuation (Grand, 1987). Obviously, the geothermal profile and particularly the olivine grain size are critical factors in determining the thickness of the mantle anisotropic layer. In active erogenic zones, where the geothermal gradient is high and the strain-rate is rapid (e.g., lo- “/s), as long as the average olivine grain size exceeds 0.5 mm, all the upper mantle above the olivine/spinel transition zone at about 350-400 km will be dominated by dislocation creep, and consequently be seismically anisotropic

239 (1994) l-27

(a)

-350

1 ! ! 0

1

2

Ig(Straiwate)=~l3(/s)

3

!

!

]:

:

f

-I

4

5

6

7

6

6

7

6

Ig(Stress)

(MPa)

(W

0

1

2

3

4

Ig(Stress)

5

(MPS)

Fig. 9. Variation of the upper mantle flow strength with depth in the dislocation creep field beneath the Canadian Cordillera (a) and Shield (b). The strain-rate is assumed to be 10-“/s for the Canadian Cordillera and to be 1O-15/s for the Canadian Shield.

(Fig. 8~). But the olivine fabric intensity and hence bulk seismic anisotropy may decrease with depth as the differential flow stress attenuates with increasing temperature (Fig. 9). This agrees with the results of surface-wave tomography (Montagner and Tanimoto, 1991) and S-wave splitting (Russo and Silver, 1993). However, in stable cratons, where the geothermal gradient is low and the strain-rate is slow (e.g., 10-“/s), the average olivine grain size must exceed 1.0 mm in order to keep the upper mantle in the dislocation creep field (Fig. 8d). To summarize, our calculations indicate that as long as the average olivine grain size is larger than 1.0 mm, the mantle

S. Ji et al. / Tectonophysics

anisotropic layer may extend from the Moho to the olivine/spinel transition zone and its thickness is independent of the geothermal gradient. However, if the average grain size of the upper mantle is less than 0.2 mm, the thickness of anisotropic layer will depend on heat flow and thus the mantle temperature profile. The seismically anisotropic layer will be thinner where the heat flow is higher and thus should be thicker beneath the cold cratons than beneath warm orogenie belts. Upper mantle rocks are commonly coarse grained (> 0.5 mm) and deforms in the dislocation creep regime over a thickness of at least 250 km. Dislocation creep results in olivine LPO and, therefore, seismic anisotropy and S-wave splitting. Grain-size-sensitive diffusion creep may locally occur in extremely fine grained ( < 0.1 mm> polyphase recrystallized zones, where olivine grain growth is effectively inhibited by the presence of other phases such as pyroxenes and spine1 (Boullier and Gueguen, 1975; Karato, 1989). The thickness of the anisotropic layer in the upper mantle estimated in this study coincides with the recent S-wave splitting measurement results of Alsina and Snieder (1994) and Gledhill and Gubbins (1994). Gledhill and Gubbins (1994) investigated the anisotropic structure of the upper mantle beneath the Hikurangi subduction zone underlying the North Island of New Zealand by analyzing both local subduction zone earthquakes and the SKS phases from teleseisms at each station. They found that the delays between the two split shear waves increase with source depth, reaching the SKS values at between 400 and 600 km. The study of Alsina and Snieder (1994) suggests that the anisotropy recorded by SKS and SKKS splitting is partly located in the sublithospheric mantle at about 400 km depth. Other evidence for the presence of anisotropy in all the upper mantle is found in a comparison of the results of Vinnik et al. (1992) with those of Shih et al. (1991). At the same station (BOCO: 4.6”N, 74”W), Vinnik et al. and Shih et al., respectively, measured S-wave splitting in SKS phases and shear waves from intermediate-depth earthquakes in the neighboring Bucaramanga nest. The fast polarization directions determined using these two methods is

239 (1994) I-27

1Y

very similar (4 = 16 + 4”). The largest delays reported by Shih et al. for the events at depths near 160 km, however, are only about 0.4 s, which is three times lower than the value obtained by Vinnik et al. This discrepancy implies that a larger thickness (300-400 km) of anisotropic mantle was sampled by the SKS waves. Partial melting may take place when the local geothermal gradient crossed the solidus of mantle rocks. This occurs particularly in the uppermost part of the asthenosphere. Effects of melt on the transition of deformation mechanism and on the seismic anisotropy have been controversial. Cooper and Kohlstedt (1986), Ji and Mainprice (1986) and Dell’Angelo and Tullis (1988) proposed that the presence of melt may promote a transition from dislocation creep to diffusion-assisted grain boundary glide. Van der Molen and Paterson (1979) and Davidson et al. (1994) suggested that partial melting can lead to melt-enhance embrittlement. Neither diffusion-assisted grain boundary glide nor melt-enhance embrittlement can result in mineral LPO and consequently seismic anisotropy. In contrast, Nicolas (1992) and McNamara et al. (1994) proposed that the presence of melt can increase efficiency of mineral preferential alignment. The presence of melt or water on grain boundaries may affect olivine grain size and consequently the transition between dislocation and diffusion creeps. A quasicontinuous basaltic melt film along the olivine grain boundaries may increase the grain size by overgrowth of selected neoblasts (e.g., Drury and Van Roermund, 1989). However, if the melt is superheated, the grain size of olivine in peridotites will decrease by corrosion (Donaldson, 1985; Boudier, 1991) and dissolution (Brearley and Scarfe, 1986). Clearly, there is ample scope for further theoretical and experimental studies on the role of fluids (melt or water) on the transition of deformation mechanisms in the upper mantle. 5.2. Fast polarization direction and magnitude of the delay time Geological interpretations ting involves relationships

of shear-wave splitbetween seismic

20

S. Ji et al. / Tectonophysics 239 (1994) 1-27

anisotropy and strain (foliation and lineation), and finally between strain and tectonic processes. The relationship between strain and tectonic processes is generally clear while that between seismic anisotropy and strain is usually complicated. Previous authors (e.g., Silver and Chan, 1988; Vennik et al., 1992; Savage and Silver, 1993) have commonly considered that the polarization direction (4) of v,, coincides with the stretching lineation in the upper mantle. There are strong doubts about the uniqueness of this interpretation, however, because the polarization direction is strongly dependent on the propagation direction with respect to the structural frame (foliation and lineation; Ji et al., 1991; Ji and Salisbury, 1993; Kern, 1993). In the peridotite xenoliths studied here, the 4 corresponds to the stretching lineation (X) only for rays parallel to Y and 2 (Fig. 6). For all other ray paths, the 4 is oblique to these structural directions. Although it is known that the propagating rays of SKS phases are subvertical and the vectors of polarization of Sl and S2 lie in a nearly horizontal plane, one still needs to know the orientation of the structural frame of the upper mantle of interest in order to constrain the tectonic implication of 4. If the 4 direction does not coincide with the stretching lineation in the upper mantle, there is no reason to relate this direction to surface geologic/ tectonic features, nor to absolute plate motion. As shown in Fig. 6, the biggest delay time between two arrivals is observed when the ray path is parallel to the foliation and perpendicular to the lineation. In the other propagation paths the delay time is smaller than the maximum value. For example, in the upper mantle beneath Nunivak Island, the delay times for the path in the Xand Z-directions are only about 60 and 40% of the value in the Y-direction (Fig. 6). The angular distance between the contours of the highest and lowest S-wave anisotropies is generally less than 45”. Therefore, the S-wave splitting (AK or at) also depends on the propagation direction with respect to the structural frame in the anisotropic upper mantle. Because the exact attitude of the structural frame in the subcontinental upper mantle is usually unknown, we may conclude that

(a)

g r

P 5

,151 llO-

n

105n

5 100._ t 95!! i

nn

n

90n

2c

wmm

n

#m

nn

85 Gi .‘p 2 5 0

n

aQ-

n

75-8

I

I

80

a5

90

.

I

I

. 105

100

95

A Weg)

W

2.5l 2.03 k

l l 1.5-

le

E ';

0

1.0-

8 l l

s 6

2

l

he

le

0.5-

0.04 . 80

I

85

,

,

90

95

,

100

4 105

A (dw) Fig. 10. Variations of the observed fast polarization direction (a) and delay time (b) for SKS with epicentral angles (A) at station TUNL (36.199”N, 94.815”E). This station is located just off the northern edge of the Tibetan plateau. The fast polarization directions range from 79 to 109” with an average value of 92”, and the delay times vary from 0.75 to 2.20 s with an average value of 1.11 s. The source regions for 24 SKS arrivals were the subduction zones at Solomon, New Hebrides and Tonga, and the active regions of western North America. Data from McNamara et al. (1994).

while small delay times (at> between S-wave arrivals are not necessarily indicative of low anisotropy, large delay times are diagnostic of strong anisotropy even if it is unclear whether the observed St or AI/, is the maximum possible value. It is often noted that at a single station, both fast polarization direction ($1 and delay time (St)

S. Ji et al. / Tecfonophysics 239 (1994) l-27

in SKS phases vary considerably from event to event (Fig. 10). The variations demonstrate that the SKS ray paths from the core/mantle boundary to the surface is not entirely vertical. The SKS waves from different events have different incident angle with respect to the foliation and lineation in the upper mantle. Thus, the variation of measured splitting parameters (4 and at> at the same station can be a result of their strong dependence on the propagation direction. Rather than simply the incident angle varying from event to event, 4 and 6t may also be affected by heterogeneities in lithology, strain and LPOs of olivine and pyroxenes. Therefore, numerous SKS observations at each station are extremely important for correct interpretations of shear-wave splitting in terms of tectonic deformation. Increasingly, geologists and geophysicists are coming to believe that the crust and upper mantle are involved differently in erogenic deformation. During a continent-continent collision orogeny, the crust may be dominated by thrustrelated thickening (mountain-perpendicular movement) while the underlying upper mantle is dominated by mountain-parallel movement (Vauchez and Nicolas, 1991; Nicolas, 1993). The mountain-parallel movement along a series of large-scale strike-slip fault zones allows lateral extrusion of the crustal and particularly the upper mantle materials from the hinterland ahead of a relatively rigid indenter (Tapponnier et al., 1982, 1990). Within these mountain-parallel movement zones, intensive ductile deformation (displacement of hundreds to thousands of kilometers) can produce strong olivine LPO, steeply dipping foliations and horizontal extension lineations in the upper mantle. Under these circumstances, the SKS waves have a propagating path nearly parallel to the Y-direction of the structural frame with the 4 direction nearly parallel to the regional extension lineation and prominent S-wave splitting will occur. The resultant 6t may be as high as 2.0 s for a S-wave propagating vertically through a 200 km thick upper mantle (Fig. 6) and the 4 direction is nearly parallel to the strike of the movement zones. This was demonstrated by recent observations of SKS and SKKS splitting in regions of present-day convergence. For example,

21

in Eurasia, the coincidence of the d, direction with the structural trend of mountain ranges has been found in the Tibetan Plateau (TUNL, BUDO, ERDO, WNDO, AMDO, XIGA and GANZ), the Tien Shan (TAS, FRU), Balcans (VTS), the Caucasus (BKR, GRS), the Pamirs (GAR) and the Hindu-Kush (KAAO) (Silver and Chan, 1991; Vinnik et al., 1990, 1992; Milev and Vinnik, 1991; McNamara et al., 1994). At stations USHU, SHIO and CHTO, the average d directions are apparently parallel to the Jinsha-Red River and Sangall strike-slip fault zones and hence parallel to the general southeast extrusion direction of Indochina due to the penetration of India into Eurasia (Tapponnier et al., 1982). The SKS splitting measurements in the Appalachians and the Pyrenees (Barruol et al., 1994) are also characterized by a seismic anisotropy with the 4 direction parallel to the structural trend of the mountain belts. Moreover, the average 4 directions in the Andean (Shih et al., 1991; Russo and Silver, 1993), central Japanese, Izu-Bonin, Taiwan and Indonesian subduction zones are also parallel to the trench, indicating that the mantle flow above subduction slabs are primarily trench parallel. McNamara et al. (1994) observed very large delay times (2.0-2.4 s> in northern Tibet. This high St region correlates with a zone of poor Sn propagation (Ni and Barazangi, 1983; Beghoul et al., 19931, large teleseismic S-P travel time residuals (Molnar, 1990) and slow Rayleigh phase velocities (Brandon and Romanowicz, 19861, where heat flow is anomalously high and widespread Cenozoic volcanism with both basaltic and granitic compositions has been found. The basaltic volcanism was due to melt extraction from partially melted mantle peridotites by hydrofracturing and porous flow. The above geological and geophysical facts may suggest that melt-filled, preferentially oriented cracks in the upper mantle may contribute, besides olivine LPO, to the seismic anisotropy inferred from shear-wave splitting. Geological, volcanological and geochemical data suggest that many parts of continents (e.g., Central Europe, western North America, the Balkans and southern part of central Sweden) originated in oceanic settings through plate tec-

22

S. Ji et al. / Tectonophysics 239 (1994) I-27

tonic processes such as subduction and accretion at convergent boundaries or island arcs. In the subcontinental lithosphere, remnants of the paleosubduction slabs of the oceanic lithosphere may retain original olivine fabrics and result in dipping anisotropic structures (Babuska and Cara, 19911, where the fast-velocity directions and intermediate-velocity directions are within the dipping planes and the low-velocity direction is perpendicular to the plane. Both the 6t and 4 direction of SKS phases vertically through dipping paleosubduction slabs may vary largely and their geological interpretation may be very complicated. In the asthenosphere, the attitudes of flow foliation and lineation at a given locality will depend on its position in the convection cell. Within a common convection cell, the foliation and lineation are expected to be subvertical in the center of upwelling and downwelling plumes, subhorizontal in the advection layers and oblique in the transition zone between the plume and advection layer. As the ray path of the nearly vertical propagating SKS waves is closely parallel to the Z-direction in the advection layer and to the X-direction in the upwelling or downwelling plume, the delay time between two arrivals cannot have the maximum value (Fig. 6). Therefore, the contribution of the asthenosphere to the total time delay in SKS phases may not be larger than 1.0-1.5 s. 5.3. Variation of seismic anisotropy with depth It is important to determine how uniform seismic anisotropy is in the subcontinental upper mantle. If the seismic anisotropy is concentrated in distinct layers rather than distributed continuously in the upper mantle, one may expect a multiple splitting in the SKS phases. Each split shear wave out of a deeper anisotropic layer could split again when it enters shallower anisotropic layer. Multiple splitting will complicate the waveform (Savage and Silver, 1993) and make the measurement and interpretation of Swave splitting difficult. Hence shear-wave splitting in a single, thick anisotropic layer may be much easier to detect than in multiple, thin

anisotropic layers. For this reason, we propose that in mountain-parallel movement zones the anisotropic layer has such a continuity in the vertical direction that large 6t values can be measured in the SKS and SKKS phases. Discrete bands of reflections below the Moho, observed in deep seismic profiling (e.g., McGeary and Warner, 1985; Clowes et al., 1992), have been interpreted as shear zones in the lithospheric upper mantle. Mantle xenoliths with mylonitic textures and those with coarse-granular textures may be derived, respectively, from the shear zones and their surrounding rocks. Variation in the strength and style of olivine LPO with microstructure or texture appears to be complex in the mantle xenoliths studied here. In the xenoliths from Castle Rock and Nunivak Island, the style and strength of olivine LPO are consistent from sample to sample, irrespective of microstructure and equilibrium temperature. This may indicate that the upper mantle beneath Castle Rock and Nunivak Island is relatively homogenous in seismic anisotropy in the depth range (40-60 km), where the xenoliths were sampled by basaltic magmas. At Alligator Lake, however, the peridotite xenoliths with a granuloblastic equant texture (e.g., sample AL411 or a tabular texture (e.g., AL88) show much weaker LPO than those with porphyroclastic texture. This indicates that the upper mantle beneath Alligator Lake is likely heterogeneous in seismic anisotropy in the depth range between 35 and 60 km. It is impossible to determine the variation of seismic anisotropy in the deeper portion of the upper mantle due to the absence of garnet-bearing xenoliths in these suites. 5.4. Source of seismic anisotropy The source of the anisotropy indicated by shear-wave splitting has been debatable. Silver and Chan (1988, 19911, Babuska et al. (1993) and Silver and Kaneshima (19931 have proposed that anisotropy is due to ancient tectonic deformation that has been frozen in the continental lithosphere. Vinnik et al. (1992), however, suggested that anisotropy is due to present-day flow in the subcontinental asthenosphere and to deformation

S. Ji et al. / Tectonophysics

in the lithosphere. If the contribution of the asthenosphere predominates over that of the lithosphere, the fast directions of anisotropy will coincide with the directions of absolute plate velocity (Vinnik et al., 1992). Seismic anisotropy is almost certainly formed by dislocation-creep-induced LPO of upper mantle minerals (mainly olivine). Since dislocation creep is temperature sensitive, the preservation of LPO in olivine formed by an ancient tectonic deformation in the modern subcontinental lithosphere will depend largely on the thermal regime. The olivine flow law in the dislocation creep regime (Karat0 et al., 1986) can be used to address this problem (Fig. 9). At natural strain-rates (e.g., lo-‘“l’s for the hot active regions and 10-“/s for cold stable regions), olivine is able to undergo a significant amount of plastic deformation and to form new LPO when flow stresses are lower than 50 MPa (Nicolas and Poirier, 1976). This threshold roughly corresponds to a depth of 210 km in the Canadian Shield, and therefore anisotropy in the upper mantle above this depth is most probably “frozen” from the past. The extremely high values of 6t (1.5-1.7 s, Silver and Kaneshima, 1993) found in a 250 km wide, NEEtrending belt in the western Superior Province of the Canadian Shield may be dominated by anisotropy formed by mountain-parallel movement during the Archean. However, in the Cordillera of North America, where the geotherm is steeper, the threshold corresponds to a depth of only about 45 km. As a consequence, the anisotropy of the upper mantle in the SKS path beneath the active erogenic regions should be dominated by present-day tectonic deformation. Furthermore, there may be a correlation between thickness of fossil anisotropic layers or apparent elastic thickness of continental lithosphere, and present-day surface heat flow. The elastic thickness decreases with increasing surface heat flow (e.g., Karato, 1984). Old, cold (presentday heat flow 30-50 mW/m*, Meissner, 1986) continental lithosphere is strong and elastic throughout its entire thickness so that fossil anisotropy can be preserved, whereas in areas of higher heat flow (So-100 mW/m2), the elastic thickness has decreased and the ancient fabric

239 (I 994) l-27

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has been destroyed. New fabric can form as a result of a recent thermal events associated with rifting or continental collision. The contribution of fossil anisotropy to the S-wave splitting in the SKS phases becomes important when the effective elastic thickness of continental lithosphere exceeds 100-200 km in cratons and shields as well as in old, eroded mountain belts.

Acknowledgments

This study was supported by grants from the NSERC of Canada (OGP0121552) and from the FCAR of Qutbec (NC-10781 to S. Ji. The first author thanks V. Babuska, K. Gledhill, H.M. Kern, D. Mainprice, M. Mareschal, A. Nicolas, P.G. Silver, L.P. Vinnik and P. Zhao for helpful discussions and J. Martignole for reading an early version of the manuscript. Careful reading and constructive comments of D. Mainprice and S. Karat0 are grateful acknowledged. This is LITHOPROBE contribution No. 587. References Alsina, D. and Snieder, R., 1994. Small-scale sublithospheric continental mantle deformation: constraints from SKS splitting observations. 2nd Int. Workshop on Dynamics of the Subcontinental Mantle from Seismic Anisotropy to Mountain Building, La Grande Motte. France, p. 21 (abstract). Ando, M., Ishikawa, Y. and Yamazaki, F., 1983. Shear wave polarization anisotropy in the upper mantle beneath Honshu, Japan. J. Geophys. Res., 88: 5850-5864. Askarpour, V., Manghnani, M.H., Fassbender, S. and Yoneda, A., 1993. Elasticity of single-crystal MgAl,O, spine1 up to 1273 K by Brillouin spectroscopy. Phys. Chem. Miner.. 19: 511-519. Babuska, V. and Cara, M., 1991. Seismic Anisotropy in the Earth. Kluwer, London, 217 pp. Babuska, V., Plomerova, J. and Sileny, J., 1993. Models of seismic anisotropy in the deep continental lithosphere. Phys. Earth Planet. Inter., 78: 167-191. Bai, Q., Ma&well, S.J. and Kohlstedt, D.L., 1991. High-temperature creep of olivine single crystals, 1. mechanical results for buffered samples. J. Geophys. Res., 96: 24412463. Barruol, G., Silver, P., Vauchez, A. and Souriau, A., 1994. Seismic anisotropy in erogenic areas: Results from the Pyrenees and the Appalachians. Eos, Am. Geophys. Union Trans., 75(16): 349 (abstract).

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