# P-wave anisotropy in eclogites and relationship to the omphacite crystallographic fabric

Alexandra Mauler, Luigi Burlini, Karsten Kunze, Pascal Philippot, Jean

Pierre Burg

► To cite this version:

Alexandra Mauler, Luigi Burlini, Karsten Kunze, Pascal Philippot, Jean Pierre Burg. P-wave anisotropy in eclogites and relationship to the omphacite crystallographic fabric. Physics and chemistry of the earth. Part A: Solid earth and geodesy, 2000, 25 (2), pp.119-126. 10.1016/S1464-1895(00)00020-X. hal-03867204

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# P-Wave Anisotropy in Eclogites and Relationship to the Omphacite Crystallographic Fabric

A. Mauler¹, L. Burlini¹, K. Kunze¹, P. Philippot², J.-P. Burg¹

¹Geologisches Institut, ETH Zentrum, CH-8092 Zürich, Switzerland

²Laboratoire de Pétrologie Métamorphique, Université Paris 7, F-75252 Paris Cédex 05, France

**Abstract.** Measurements of P-wave velocity at room temperature and confining pressures up to 500 MPa were carried out on three eclogite mylonites collected from a shear zone in the Monviso area (Western Alps). P-wave velocities at a pressure of 400 MPa range from 7.7 km/s to 7.9 km/s, yielding to a maximum anisotropy of 6%. From the CPO of omphacite we estimated a maximum contribution of omphacite to the P-wave anisotropy of only 1.3%. These results suggest that primarily the compositional layering and secondary the fabric of minor constituent minerals significantly contribute to the seismic anisotropy. Because of the anisotropy, the seismic reflectivity of subduction zones may vary with the direction of observation.

## 1 Introduction

Laboratory measurements of seismic properties of eclogites have demonstrated that eclogites have a week seismic anisotropy (Christensen, 1974; Kumazawa et al., 1971; Maghanani et al., 1974). As an example, Fountain et al. (1994) observed a mean P-wave anisotropy of 3% in highly sheared eclogites converted from Grenvillian granulite during the Caledonian orogeny. This anisotropy seems to be weak with respect to the strong compositional layering, the shape and crystallographic preferred orientations (SPO and CPO respectively). If such weak anisotropy is also confirmed in eclogites originating from subduction zones, it would make any seismic distinction between eclogites and surrounding mantle peridotites impossible, as pointed out by several authors (Fountain & Christensen, 1989; Kern, 1993; Kumazawa et al., 1971).

In this paper, we present laboratory measurements of P-wave velocities and their anisotropies performed at pressures up to 500 MPa and room temperature on mylonitic eclogites from the Monviso ophiolitic complex (Western Alps). P-wave

velocities were also calculated using a volume averaging scheme based on the CPOs, elastic constants of major minerals and the modal composition of the rock. Results are discussed with respect to seismic properties of eclogites and mantle peridotites from other localities.

<figure><img src="image_2.png" /><figcaption>Fig.1. Geological map of the Lago Superior unit of the Monviso complex and location of the sampling sites (modified after Philippot, 1988).</figcaption></figure>

*Correspondence to:* A. Mauler
## 2 Geological Setting

The Monviso ophiolitic complex is located on the French-Italian border about 80 km SW of Torino. It is interpreted as a remnant slice of oceanic crust originating from the Piemonte ocean basin that separated the Apulian plate from the European plate during Jurassic times (Dal Piaz, 1974). The closure of this ocean is recorded in the Monviso area by Alpine metamorphism, Mid-Eocene in age (Monié & Philippot, 1989, Duchêne et al., 1998), and of decreasing PT conditions from eclogite facies (570°C<T<670°C, P>2400 MPa) to blueschist and greenschist facies (Messiga et al., 1999, Lombardo et al., 1978).

An eclogites facies shear zone is part of the Lago Superiore unit (Fig. 1) and is a km-scale zone that formed at a minimum depth of 70 km in a subduction zone. It consists of highly mylonitized Mg-metagabbros that contain boudins of mylonitic Fe-metagabbros and locally lenses of low-strained rocks preserving igneous textures (Philippot, 1988; Philippot & Van Roermund, 1992).

## 3 Sample Description

Modal composition and density of the three studied samples are summarized in Table 1, and optical micrographs are shown in Fig. 2.

Samples V16 and V15 (Fig. 2a, 2b) are annealed eclogites presenting a plano-linear fabric. Alignment of quartz and broken rutile grains along thin layers marks the foliation. The stretching lineation is defined by elongated omphacite grains. Some retrogression attending blueschist facies conditions is indicated by rims of chloromelanite around omphacite and flakes of glaucophane and phengite oriented at high angle to the foliation plane. Late stage greenschist facies overprint is negligible and occurs as randomly dispersed grains of chlorite and plagioclase and overgrowth zones of actinolite after glaucophane.

Sample V10 (Fig. 2c) is a partially retrogressed eclogite. Apart from the eclogite facies-minerals, omphacite, garnet and rutile, it contains glaucophane, phengite and zoisite, associated with blueschist facies conditions and plagioclase,

<figure><img src="image_3.png" /></figure>

**Fig.2.** Micrographs illustrating the microstructures of a) sample V16, b) sample V15 and c) sample V10. Left: Sections parallel to lineation and perpendicular to foliation. Right: Sections normal to lineation and perpendicular to foliation. For samples V16 and V15, the measured omphacite shape ratio in each section is represented as segment of an ellipse
actinote and chlorite, representative of a late greenschist facies stage. Omphacite grain boundaries are commonly corroded by a late clinopyroxene and plagioclase symplectite. All these minerals form moderately to highly strained thin layers (mm scale) that alternate parallel to the foliation plane. Grain sizes range from about 100 to 300 mm in the moderately strained layers to around 10 mm in the highly strained domains.

<table><caption>**Table 1.** Modal composition (in volume %) and density (in g/cm³) of the studied samples. On the right the modal proportions used for the calculation of P-wave velocities are presented.</caption><thead><tr><th>Sample</th><th>V16</th><th>V15</th><th>V10</th><th>V16</th><th>V15</th><th>V10</th></tr></thead><tbody><tr><td>Omphacite</td><td>61</td><td>59</td><td>43</td><td>63</td><td>64</td><td>46</td></tr><tr><td>Garnet</td><td>30</td><td>26</td><td>26</td><td>31</td><td>28</td><td>28</td></tr><tr><td>Quartz</td><td>5</td><td>1</td><td>1</td><td>6</td><td>-</td><td>-</td></tr><tr><td>Rutile</td><td>1</td><td>1</td><td>-</td><td>-</td><td>8</td><td>18</td></tr><tr><td>Epidote</td><td>-</td><td>7</td><td>17</td><td>-</td><td>-</td><td>8</td></tr><tr><td>Amphibole</td><td>2</td><td>4</td><td>3</td><td>-</td><td>-</td><td>8</td></tr><tr><td>Plagioclase</td><td>-</td><td>1</td><td>7</td><td>-</td><td>-</td><td>8</td></tr><tr><td>Other</td><td>1</td><td>1</td><td>3</td><td>-</td><td>-</td><td>8</td></tr><tr><td>Bulk Density</td><td>3.49</td><td>3.49</td><td>3.23</td><td>-</td><td>-</td><td>-</td></tr><tr><td>Grain Density</td><td>3.506</td><td>3.516</td><td>3.244</td><td>-</td><td>-</td><td>-</td></tr><tr><td>Calc. Density</td><td>3.527</td><td>3.519</td><td>3.467</td><td>-</td><td>-</td><td>-</td></tr></tbody></table>

# 4 Experimental determination of P-wave velocities

## 4.1 Experimental Technique

The measurements of P-waves velocities were performed at room temperature and pressures up to 500 MPa in a Heard type pressure vessel at ETH Zürich using the pulse transmission technique (Birch, 1960) on cores (26 mm diameter, 30 to 55 mm long) of oven-dried specimens.

<figure><img src="image_4.png" /><figcaption>**Fig.3.** Sketch of the Heard Rig used for the experimental determination of P-wave velocities. On both sides of the specimen, a brass foil ensures electric continuity between the transducers, which are connected to the external part of the pressure vessel and the electrically common pole.</figcaption></figure>

Three perpendicular cores were cut from each sample and related to the inherent fabric elements with the following convention: X core parallel to lineation, Y core normal to lineation and parallel to foliation, Z core normal to foliation. In sample V15, a fourth core W oriented within the XY plane as shown in Fig. 4 was drilled to better constrain the seismic properties. The specimen assembly is represented in Fig. 3. The gas pressure is provided by an air driven booster up to 300 MPa, and then by an intensifier. The electronic system consists of a HP Function generator and an oscilloscope (Le Croy, 125 MHz).

An electric pulse is transmitted to the upper piezoelectric transducer (lead zirconate, 1MHz) and converted into a mechanical pulse that propagates through the rock specimen. The pulse is converted again into an electric signal by the lower piezoelectric transducer, and then recorded by the oscilloscope. The pulse delay time is measured and converted into velocity after appropriate calibration of the system.

## 4.2 Results

Graphs of P-wave velocities vs. pressure are shown in Fig. 4. Each specimen shows a marked increase in P-wave velocities with pressure. Velocities were recorded during both pressurisation and depressurisation. The non-linear part of the curves corresponds to closure of pores and cracks. Higher velocities during depressurisation are attributed to the fact that some pores and cracks do not reopen immediately (hysteresis). The quasi-linear part of the curves represents the intrinsic properties of the rocks and is approached at pressure of 250 MPa for samples V16 and V15, and already of 150 MPa for sample V10. At 450 MPa, average P-wave velocities range from 7.71 km/s in sample V10 to about 7.87 km/s in sample V15 and V16.

The P-wave anisotropy $A = 100\% * (V_{max} - V_{min}) / V_{mean}$ (Birch, 1961) ranges at maximum pressures from 3% to more than 6%. The fastest P-waves are within the foliation plane in all samples, and the slowest are normal to foliation (Z). Sample V16 is transversally isotropic for P-waves as indicated by similar P-wave velocities in the X and Y directions. In sample V15, the fastest P-waves are parallel to lineation (X), while in V10 the fastest propagation direction is perpendicular to lineation (Y).

# 5 Calculation of P- and S-wave velocities

The P-wave velocities were measured along 3 resp. 4 different directions per sample. This is not enough to completely constrain the anisotropy and to determine the maximum anisotropy between any direction. In order to establish the distribution of P-waves in all directions, the complete shape of anisotropy and to identify the minerals responsible for the anisotropy, calculations of seismic properties were performed. These included calculations of P- and S-wave velocities from modal composition, single crystal
Sample V16

<figure><img src="image_5.png" /></figure>

<figure><img src="image_6.png" /></figure>

Max.Vp = 8.31

Min.Vp = 8.22

Anisotropy = 1.1 %

Vp Contours (km/s): 8.24 8.25 8.26

8.27 8.28 8.29

Sample V15

<figure><img src="image_7.png" /></figure>

<figure><img src="image_8.png" /></figure>

Max.Vp = 8.45

Min.Vp = 8.35

Anisotropy = 1.1 %

Vp Contours (km/s): 8.37 8.38 8.39

8.40 8.41 8.42 8.43

Sample V10

<figure><img src="image_9.png" /></figure>

<figure><img src="image_10.png" /></figure>

Max.Vp = 8.06

Min.Vp = 7.96

Anisotropy = 1.3 %

Vp Contours (km/s): 7.98 7.99 8.00

8.01 8.02 8.03 8.04

<caption>**Fig.4.** P-wave velocity vs. confining pressure curves. Open symbols represent measurements recorded during pressurization and filled symbols represent measurements recorded during depressurization. Shaded rectangles define P-wave velocity intervals over all sampled directions as calculated using the Voigt, VRH and Reuss average schemes at room temperature and pressure (see text for further explanations).
Directional distributions of P-wave velocities calculated using the VRH scheme at room temperature and pressure, are shown as upper hemisphere equal area projections. Trace of the foliation plane is indicated by the horizontal line and lineation is E-W.</caption>
elastic constants, density and crystallographic preferred orientation of the main constituent minerals.

The Voigt and Reuss averaging schemes, representing upper and lower bounds of the elastic properties of a polycrystalline aggregate, were calculated. In addition, the Voigt-Reuss-Hill (VRH) average was derived as the arithmetic mean of the elastic tensors obtained from the Voigt and Reuss schemes.

## 5.1 Elastic constants

The elastic constants used for the calculation are those of Bhagat & Bass (1992) for omphacite, Babuska (1978) for garnet, Ryzkova (1966) for epidote and Haermon (1979) for albite and quartz. The computations were performed using the interactive program reported by Mainprice (1990). The calculation is based on the CPOs of omphacite for samples V15 and V10 and of omphacite and quartz for sample V16. Garnet, epidote and plagioclase have been assumed to be randomly oriented. Minor minerals have been ignored in the calculations (Table 1).

## 5.2. CPO determination

Crystallographic preferred orientations were measured using electron backscatter diffraction (EBSD) (Venables & Harland, 1973; Dingley et al., 1987; Kunze et al., 1993). EBSD patterns were collected in a SEM CamScan CS44LB at the

ETH Zürich and automatically indexed using the software OIM 2.0 (TSL Inc.) based on a defined reference file of 40 reflections for omphacite (Mauler et al., 1998). The technique allows the determination of the full crystallographic orientation of grains with size down to 1µm and with an absolute angular resolution of 1-2°. In samples V15 and V16, orientation data were collected on a 6000µm x 2460µm grid with step size of 60µm (4242 data points). In sample V10, patterns were analyzed in an omphacite-rich area on a 250µm x 400µm grid with step size of 5µm (4131 measurements). Additionally, the quartz CPO was obtained for sample V16 by collecting and manually indexing EBSD patterns from 215 quartz grains.

In Fig. 5, pole figures of the b-axis [010] and -b-axis [0-10], of the c-axis [001] and a-plane (100) are displayed. The distribution of the axes [207] and [-101] are also shown, as they correspond respectively to the fastest and slowest propagation direction for P-waves in an omphacite single crystal of the composition described by Bhagat & Bass (1992) (Fig. 6). Omphacite fabrics in samples V16 and V15 have the same characteristics. The [010] and [0-10] pole figures have similar distributions with a strong single maximum normal to foliation. The [001] and (102) pole figures have a point maximum close to lineation. The (100) and (-502) pole figures do not show any strong maximum. The fabric of sample V10 differs from the fabrics of the other samples as the [001] and (102) pole figures display a girdle lying in the foliation with a maximum parallel to lineation. The fabric of quartz in sample V16 does not show any clear CPO.

<figure><img src="image_11.png" /><figcaption>**Fig.5.** CPO of omphacite. a) sample V16, b) sample V15, c) sample V10 and d) quartz (sample V16). Logarithmic scale on upper hemisphere equal area projection. Trace of the foliation plane and lineation are horizontal</figcaption></figure>
## 5.3 Results

Results are displayed in Figures 4 and 6 and in Table 2. The calculated P-wave velocities vary from 7.72 (Reuss average in sample V10) to 8.55 km/s (Voigt average in sample V16) with anisotropy values of about 1.3% in all the samples. In samples V16 and V10, Voigt and Reuss bounds are widely separated because they include quartz and plagioclase, components with much weaker elastic moduli than garnet, omphacite or epidote. The directions of highest P-wave velocities are subparallel to lineation and within the foliation plane in all the samples and closely related to the maxima of omphacite [207] and [001], which are both fast directions in an omphacite single crystal. Directions of slow P-wave velocities are perpendicular to foliation and correspond to maximum concentrations of the b-axis, although that is not the slowest direction for P-waves in omphacite single crystals. This is due to the fact that the slowest direction [-101] is more randomly distributed, while the b-axis is a slow direction (Vp along [010] is 8.1 km/s) and is clearly concentrated in the direction normal to foliation. Shear-wave velocities (Table 2) display more complicated patterns, with the maximum and minimum velocities oriented away from foliation and lineation. Fast and slow S-wave anisotropies are lower than P-wave anisotropies.

For sample V16, the calculations were performed with and without the contribution of quartz. Results show that the quartz reduces the calculated Vp by about 3% (Table 2). Quartz contributes in a destructive manner to the P-wave anisotropies and in a constructive manner to the S-wave anisotropies.

## 6 Comparison between measured and calculated velocities

Figure 4 shows that the measured P-wave velocities are significantly lower than the calculated velocities. Possible explanations are an incomplete closure of the microcracks,

even at 500 MPa confining pressure, or symplectization along grain boundaries and the presence of blueschist facies minerals resulting from retrogression.

Despite the differences in absolute values, the distributions of calculated P-wave velocities are consistent with the measured velocities: in samples V16 and V15, velocities are high parallel to X and low parallel to Z. In sample V10, velocities are high within the foliation and low parallel to Z. This shows that the distribution of P-waves in the rock is correlated to the omphacite CPO in all the samples, even in sample V10, where the proportion of omphacite is less than 45%. Calculated anisotropies (about 1.3%) using either the Voigt or the VRH or the Reuss schemes are significantly lower than the measured anisotropies (up to 6% in sample V10). One explanation for the discrepancy is that the Voigt, Reuss and VRH averages are exclusively based on the elastic constants of the constituents, their volume fractions and their CPO. No information about the microstructure, such as composition layering, microfracturing, grain size and shape or grain distribution (foliated or lineated microstructures, vs. random orientation of isolated planar or elongated or equidimensional grains) is included. In fact, the clear anisometric omphacite grain shape (Fig. 2), as well as microfractures statistically oriented parallel to foliation could produce a small anisotropy, with lower velocity in the direction normal to the foliation. Moreover, in samples V15 and V10

<table><caption>Table 2. P-, S1-, S2-wave velocities and anisotropies, shear wave splitting (ΔVs) calculated for room pressure and temperature with the VRH average scheme. Velocities of sample V16 were calculated using the CPO of quartz, while in sample V16*, the contribution of quartz was ignored.</caption><thead><tr><th rowspan="2"></th><th colspan="3">Vp</th><th colspan="3">Vs1</th><th colspan="3">Vs2</th><th>ΔVs</th></tr><tr><th>Max (km/s)</th><th>Min (km/s)</th><th>A (%)</th><th>Max (km/s)</th><th>Min (km/s)</th><th>A (%)</th><th>Max (km/s)</th><th>Min (km/s)</th><th>A (%)</th><th>Max (km/s)</th></tr></thead><tbody><tr><td>V16</td><td>8.31</td><td>8.22</td><td>1.1</td><td>4.84</td><td>4.79</td><td>0.9</td><td>4.82</td><td>4.78</td><td>0.7</td><td>0.05</td></tr><tr><td>V16*</td><td>8.55</td><td>8.45</td><td>1.2</td><td>4.90</td><td>4.86</td><td>0.7</td><td>4.88</td><td>4.85</td><td>0.6</td><td>0.04</td></tr><tr><td>V15</td><td>8.45</td><td>8.35</td><td>1.1</td><td>4.85</td><td>4.80</td><td>0.9</td><td>4.83</td><td>4.80</td><td>0.6</td><td>0.02</td></tr><tr><td>V10</td><td>8.06</td><td>7.96</td><td>1.3</td><td>4.60</td><td>4.58</td><td>0.4</td><td>4.59</td><td>4.57</td><td>0.4</td><td>0.05</td></tr></tbody></table>

<figure><img src="image_12.png" /><figcaption>Fig. 6. P-wave distribution in an omphacite single crystal of the composition described by Bhagat & Bass (1992) and relation with the main crystallographic directions and poles. Upper hemisphere equal area projections.</figcaption></figure>
the calculations are based only on the CPO of omphacite (omphacite and quartz in sample V16) while secondary phases displaying a strong CPO may have some effect on anisotropy. The difference between the measured maximum and minimum velocity is similar to the difference between the Voigt and Reuss bounds (Fig.4). That suggests that the narrow compositional layering has a strong effect on anisotropy.

## 7 Discussion

The measurements performed in this study show that average P-wave velocities reach 7.7 km/s to 7.9 km/s at 450 MPa in the sheared eclogites from the Lago Superiore unit of the Monviso. P-wave anisotropy ranges from 3% to more than 6%. These values are similar to those obtained on other eclogites. Typical P-wave velocities lie between 7.5 km/s and 8.7 km/s and anisotropies can reach 9% but are generally below 4% (Christensen, 1974; Fountain et al., 1994; Fountain & Christensen, 1989; Kumazawa et al., 1971; Maghanani et al., 1974; Rao et al., 1974).

Seismic studies of upper mantle rocks show that peridotites are characterised by P-wave velocities ranging from 8.0 km/s to 8.7 km/s and by a high anisotropy (typically from 4% up to 16%) (Fountain & Christensen, 1989; Kern, 1993).

In eclogites, the CPO of omphacite creates a concentration of c-axes parallel to lineation and therefore fast velocities in a direction parallel to lineation. In peridotites, seismic anisotropy is strongly controlled by the CPO of olivine with the olivine a-axes (which corresponds to the maximum P-wave velocity in olivine) producing fast velocities parallel to the lineation. Fountain & Christensen (1989) pointed out in a compilation of eclogite and peridotite seismic properties, that the differences between the mean velocity of these two rock types are too small to be discriminated by seismic investigations. However, in subduction zones, where rocks are strongly anisotropic due to the large amount of strain, seismic differentiation between eclogites and the surrounding peridotitic mantle might depend on the direction of observations because of their large difference in anisotropy. The reflection coefficient (Rc) and the difference in P-wave velocity ΔVp for deep seismic sounding (DSS) between peridotites and eclogites were calculated from the data of Ismael & Mainprice (1998) for subduction zone peridotites and from our data for eclogites. Provided that the layers are sufficiently thick to be imaged, no distinction is possible between the two types of rock in a steep subduction zone, as Rc is less than 0.01 and ΔVp less than 0.16 km/s. If the subduction is shallow, distinction is possible in reflection mode (Rc about 0.03) and in DSS, if the propagation direction is parallel to the dip of the subduction (ΔVp around 0.5km/s). If the propagation direction is parallel to the strike of the subduction, distinction will be very difficult (ΔVp below 0.1km/s).

## 8 Conclusion

The comparison of measured and calculated P-wave anisotropies indicates that the seismic anisotropy can qualitatively be interpreted by the omphacite CPO. However, no quantitative correlation was noted. This suggests that compositional layering, shape fabric and contributions of minor constituents along grain boundaries may have a strong effect on the anisotropy. More advanced models (self-consistent) should be used to better constrain seismic anisotropies in eclogites. A consequence of the low P-wave anisotropy observed in eclogites is that the recognition of the contact eclogites-peridotites in a subducting slab will depend on the direction of observation.

**Acknowledgements:** We thank David Mainprice for providing computer programs and for helpful discussions about crystallographic conventions. Comments by two anonymous reviewers helped to improve the manuscript. This work was supported through ETH project no. 17094/41-2704.5 and by Schweizer Nationalfonds project R'EQUIP no. 2160-53289.98.

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