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experiments and in nature
Elodie Amiguet, Bertrand van de Moortèle, Patrick Cordier, Nadège Hilairet,
Bruno Reynard
To cite this version:
Elodie Amiguet, Bertrand van de Moortèle, Patrick Cordier, Nadège Hilairet, Bruno Reynard.
Deformation mechanisms and rheology of serpentines in experiments and in nature.
Journal of
Geophysical Research : Solid Earth, American Geophysical Union, 2014, 119 (6), pp.4640-4655.
�10.1002/2013JB010791�. �hal-02107380�
Deformation mechanisms and rheology
of serpentines in experiments
and in nature
Elodie Amiguet1,2, Bertrand Van De Moortèle1, Patrick Cordier3, Nadège Hilairet3, and Bruno Reynard1
1
Laboratoire de Géologie de Lyon, Université de Lyon, CNRS, ENS de Lyon, Université Claude Bernard Lyon 1, Lyon, France,2Institute of Condensed Matter Physics, Ecole Polytechnique Fédérale de Lausanne, EPFL, Lausanne, Switzerland,3Unité Matériaux Et Transformations, UMR 8207 Université Lille 1-CNRS Villeneuve d’Ascq, France
Abstract
Microstructures in serpentine samples recovered from deformation experiments performed at high pressure (1–8 GPa), high temperature (150–500°C), and laboratory strain rates (4 10 4–10 6s 1) were studied using transmission electron microscopy on thin sections prepared by focused ion beam. Lizardite crystals deform by easy glide along the basal planes associated with microkink. This mechanism is associated with a plastic strength of ~100 MPa that defines the upper bound of lizardite strength in natural conditions. Antigorite crystals deform essentially by conjugated slip along (101) and 101planes observed in sections close to (010). This conjugate system results in an apparent global slip system akin to [100](001). In all samples, delamination and comminution producefine-grained interconnected regions at grain boundaries because intracrystalline deformation mechanisms are insufficient to satisfy the von Mises criterion. The deformation laws of lizardite (plasticflow) and antigorite (strain rate dependent) differ because of their differing intracrystalline deformation mechanisms. Subgrain boundary geometry in natural samples and [100](001) crystal preferred orientations indicate the activation of slip systems similar to those observed in experimentally deformed samples, suggesting that the strain rate-dependent rheology of antigorite derived from experiments applies to subduction zone conditions. Delamination of antigorite crystals of a few tens to hundreds of nanometers is not observed in natural samples from subduction zones, suggesting plastic deformation of serpentinites at natural low strain rates and stresses is likely accompanied by recrystallization through dissolution-precipitation processes that act as a low-temperature equivalent of dynamic recrystallization through diffusion at high temperature.1. Introduction
Serpentine minerals form by hydration of ultramafic rocks. The variety of serpentine minerals depends on pressure and temperature conditions [e.g., Chernosky et al., 1988; Evans, 2004; Schwartz et al., 2012]. Low-temperature serpentine minerals, lizardite and chrysotile, occur widely in serpentinites. While chrysotile is metastable, lizardite is believed to be stable from the surface up to as high as 400°C [Evans, 2004]. These low-temperature serpentines are particularly abundant at slow-spreading ridges [Mével, 2003], and along transform faults [Reinen et al., 1994] where they are associated with ductile deformation of the oceanic crust [Cannat et al., 2009]. Antigorite—the high-pressure and high-temperature serpentine—is a major water carrier in subduction zones due to its stability at temperatures and pressures up to 650°C and 6 GPa [Ulmer and Trommsdorff, 1995]. Antigorite serpentinites can form a viscous channel that accommodates plastic deformation at the top of the downgoing slab [Schwartz et al., 2001; Kawakatsu and Watada, 2007; Hilairet and Reynard, 2009]. Strong preferred orientation of antigorite in response to deformation could account for low-seismic velocities and high anisotropy in the mantle wedge above the subducting plate [Katayama et al., 2009; Mainprice and Ildefonse, 2009; Bezacier et al., 2010; Boudier et al., 2010; Jung, 2011; Mookherjee and Capitani, 2011; Nishii et al., 2011; Brownlee et al., 2013].
Deformation of serpentines has been extensively studied at low pressures (<1 GPa). Friction experiments have shown that lizardite and chrysotile possess low-friction coefficients (μ ~ 0.3–0.6) [e.g., Dengo and Logan, 1981; Reinen et al., 1994; Moore et al., 1997; Behnsen and Faulkner, 2012] and that antigorite has a friction coefficient comparable with other crustal rocks (μ ~ 0.6–0.8) [e.g., Reinen et al., 1991, 1994; Morrow et al., 2000; Okazaki et al., 2013]. The values of friction coefficients may vary with confining pressure, temperature,
Journal of Geophysical Research: Solid Earth
RESEARCH ARTICLE
10.1002/2013JB010791
Key Points:
• We characterize mechanisms involved in plastic deformation of serpentines • Deformation processes in serpentines
are comparable in experiments and in nature
• Strain rate-dependent rheology derived from experiments may be applicable
Correspondence to:
E. Amiguet,
elodie.amiguet@epfl.ch
Citation:
Amiguet, E., B. Van De Moortèle, P. Cordier, N. Hilairet, and B. Reynard (2014), Deformation mechanisms and rheology of serpentines in experiments and in nature, J. Geophys. Res. Solid Earth, 119, 4640–4655, doi:10.1002/ 2013JB010791.
Received 1 NOV 2013 Accepted 23 APR 2014
Accepted article online 29 APR 2014 Published online 9 JUN 2014
proportion of accessory phases, and porefluid pressure. A transition in the deformation regime from brittle to ductile was observed at 0.2 GPa for lizardite and 0.4 GPa for antigorite [Raleigh and Paterson, 1965; Murrel and Ismail, 1976; Verlag and Byerlee, 1978; Escartín et al., 1997a; Moore et al., 1997]. The behavior of serpentines at higher pressure (P) and temperature (T) is still debated. Semibrittle deformation and strain localization have been observed in antigorite serpentinite at conditions up to 1.5 GPa and 600°C [Chernak and Hirth, 2010]. In situ deformation experiments on antigorite [Hilairet et al., 2007] and lizardite [Amiguet et al., 2012] up to 8 GPa and 600°C suggest plastic deformation at confining pressures above 1 GPa. These experiments constrain serpentineflow strengths to be in the range of 0.1–2 GPa, lower than that of olivine—representative of dry lithospheric rocks—which is in the range of 2–3 GPa at comparable conditions [Raterron et al., 2004; Mei et al., 2010].
Discrepancies among experimental results at high strain rates and their applicability to natural conditions are debated [Reynard, 2012; Hirth and Guillot, 2013] and must be assessed by thorough comparison of active deformation mechanisms in experimental and naturally deformed samples. Crystallographic preferred orientations (CPO) in antigorite serpentinites have been studied in different geological settings. Some have been interpreted as evidence of dislocation creep with different dominant slip systems, while others have been attributed to topotactic transformation with inherited textures from the peridotites protoliths [Bezacier et al., 2010; Boudier et al., 2010; Hirauchi et al., 2010b; Soda and Takagi, 2010; Van de Moortèle et al., 2010; Jung, 2011; Nishii et al., 2011; Padrón-Navarta et al., 2012; Brownlee et al., 2013]. Other processes, such as dissolution-precipitation creep may occur at low stress [Hoogerduijn Strating and Vissers, 1991; Andréani et al., 2005; Rutter et al., 2009; Wassmann et al., 2011]. The relationship between CPO and grain-scale microstructures could be the key to understanding serpentine deformation behavior, but only a few
transmission electron microscopy (TEM) studies have been performed on these minerals [Auzende et al., 2006;
Table 1. Set of Conditions Experienced by Recovered Serpentine Samples on Which the TEM Microstructural Study Was Performed
# Sample Grain Size (μm) P (GPa) T (K) Strain Rate (10 5s 1) Number of Deformation Cycle
D1064a(Liz) < 10 8 523 0.66 4 9.65 623 8.41 0.41 D1065a(Liz) 50–100 7 423 0.59 4 11.94 523 0.94 15.70 D1130a(Liz) < 10 3.5 523 1.06 4 4.34 673 0.83 3.80 D1131a(Liz) < 10 1.5 523 1.02 3 3.43 673 1.0
D0539s 5–30 length < 5 thick 4 473 no deformation D0539b(Atg) 5–30 length < 5 thick 4 473 8.26 14
1.77 573 1.35 5.51 673 1.34 5.94 773 1.27 4.45 1 473 1.02 3.73 573 0.86 4.27 673 1.27 3.36
a(Liz = lizardite and Atg = antigorite). Mechanical results detailed in Amiguet et al. [2012]. b
Figure 1. (a) Optical microscope image of a lizardite serpentine recovered from deformation experiments (sample D1064) and the surrounding cell assembly with 1: crushable alumina piston; 2: densified alumina piston; 3: gold foil used as strain markers at the top and the bottom of the sample; 4: h-BN sleeve; 5: sample; 6: graphite furnace; and 7: boron epoxy cubic pressure medium and (b) FIB thin section. The differential stress compression axis is reported as red arrows.
Figure 2. Environmental SEM images of the deformed sample. (a) Lizardite D1064 sample, (b) Lizardite D1065 sample, (c) Lizardite D1131 sample, and (d) Antigorite D0539 sample. Figures 2a–2c are secondary electrons (SE) images and Figure 2d is a backscattered electron image. All samples display homogeneous deformation. The horizontal cracks usually form during cold decompression. The scale bars are all 50μm and the compression axis is vertical.
Hirauchi et al., 2010a]. As a result, the relationship between rheological properties and microstructures induced by deformation is still poorly constrained for serpentines.
Here we present a detailed TEM study on antigorite and lizardite serpentinites experimentally deformed at high pressure and high temperature using a Deformation-DIA (D-DIA) apparatus [Hilairet et al., 2007; Amiguet et al., 2012]. Our observations show intracrystalline deformation by glide and delamination at grain boundaries in both serpentine minerals. The large difference observed in rheological properties of antigorite and lizardite is rationalized in terms of crystallographic structures and deformation geometries. Implications for the interpretation of natural textures, rheology at natural deformation rates, and seismic anisotropy are discussed.
2. Sample and
Experimental Techniques
2.1. Experimental Details
Two natural serpentine minerals were used as the starting material. Well-characterized antigorite serpentinite from central Cuba is a one-layer polytype with an average
superperiodicity of m = 14. Its structural formulas is (Mg2.62Fe0.16Al0.15)∑ = 2.93 (Si1.96Al0.04)∑ = 2O5(OH)3.57[Auzende et al., 2004; Hilairet et al., 2007]. The lizardite serpentinite used is a euhedral 1 T polytype from Elba Island, Italy, with a structural formula
(Mg2.74Fe0.16Al0.09)∑ = 2.99
(Si1.93Al0.07)∑ = 2O5(OH)4[Mellini and Viti, 1994].
The serpentinites were ground and sieved to obtain three powders with grain size of 5–30 μm in length and less than 5μm in thickness for antigorite, and two powders for lizardite, one with grain size below 10 μm and one with grain sizes ranging from 50 to 100μm. Serpentinite powders were cold pressed manually into 1.2 mm long cylinders, with diameters 1.2 mm for the starting specimens used in high-pressure and high-temperature deformation experiments. Sample cylinder axis is set parallel to the differential compression axis in the press. Serpentine samples were experimentally deformed using a Deformation-DIA apparatus in a 250 T large volume press coupled with a monochromatic synchrotron light source at the GeoSoilEnviroCARS beamline of the Advanced Photon Source (Argonne National Laboratory, IL, USA), allowing in situ stress and strain measurements [Wang et al., 2003, 2009]. Deformation cycles were performed on a single antigorite specimen and on four lizardite specimens. A control sample was hot pressed in the D-DIA at high pressure and temperature and recovered without deformation only for antigorite. Temperature in the range 150–500°C
Figure 3. Low-magnification TEM bright-field images of recovered lizar-dite. (a) Sample D1064 deformed at 8 GPa with starting grain size <10 μm, (b) sample D1065 deformed at 7 GPa with a starting grain size between 50 and 100μm, and (c) sample D1131 deformed at 1.5 GPa with a starting grain size< 10 μm. Areas of intense grain size reduction, when visible on the image, are outlined by white dotted contours and are labeled by white asterisks. The compression direction is represented as red arrows. The scale bars are all 1μm.
and strain rates in the range 10 4–10 6s 1were applied to both serpentine minerals. The pressure range was 1–8 GPa for lizardite and 1–4 GPa for antigorite experiments (Table 1). Additional details on sample preparation and experimental protocol can be found in Hilairet et al. [2007] and Amiguet et al. [2012], for antigorite and lizardite samples, respectively.
2.2. Preparation of Recovered Samples
High-pressure cells recovered from deformation experiments were embedded in epoxy resin, cut in two parts parallel to the uniaxial
compression direction, and polished with alumina down to 1/20μm (Figure 1a). Half-cells were observed using an environmental scanning electron microscope XL30 from FEI with partial water pressure, in order to confirm distributed deformation and for choosing the location for extraction of TEM thin sections. Samples were then coated with a 20 nm thick carbon layer for scanning electron microscope (SEM)/ focused ion beam (FIB) observations. TEM thin sections were extracted parallel to the compression axis a using a ZEISS NVision40 FIB. A 10–15 nm thick W or C layer was deposited by induced electron beam [Utke et al., 2008] to avoid amorphization of the subsurface on an area about 2.5 × 25μm2. A thicker layer (0.2μm to 1–2 μm for W and C, respectively) was then added by ion beam-induced deposition [Utke et al., 2008] to protect TEM lamellae during the full process of preparation. Excavations were made on both sides of the TEM lamella location using ion beam current from 27 nA down to 3 nA at 30 kV accelerating voltage. TEM lamellae were separated from the bulk of the sample and lifted out when a thickness of about 1–2 μm was reached. They were fixed by C ion beam deposition on a half copper grid on V position as illustrated in Figure 1b. The TEM lamellae thickness was progressively decreased to 100– 150 nm for TEM observations with ion
Figure 4. TEM bright-field images of lizardite samples. The relationship between kink band, delamination of grains into slabs, and intense grain size reduction is illustrated by (a and b) Kink-band and lizardite slab bending in samples D1064 and D1131, respectively, and corresponding selected area electron diffraction (SAED) pattern (inset) and (c) delami-nation of grains into slabs leading to intense grain size reduction in sample D1065; transitional zone is delimited by dotted white lines. Note the presence of polygonal serpentine (in white dotted square) as minor phase in sample D1131. The compression axis is reported as red arrows. The scale bars are all 0.4μm.
beam current of 700 pA down to 40 pA at 30 kV from the initial to thefinal steps of the thinning process. To accessfine microstructures in the antigorite sample, an X2 sample holder provided by ZEISS was used to decrease the thickness of the sample below 75 nm without any bending of the thin area. The TEM lamellae werefinally cleaned for traces of Ga ion implantation by a last milling at 2 kV and 50 pA during 3–5 min on each side.
Twelve TEM thin sections (Figure 1b) parallel to the compression axis were extracted using FIB from six serpentine specimens recovered from high-P and high-T experiments:five on the single recovered antigorite sample [Hilairet et al., 2007], one on the control antigorite sample that was hot pressed in the D-DIA but not deformed, and six on the four recovered lizardite samples [Amiguet et al., 2012]. Details on the deformation conditions experienced by each of the samples (e.g., P, T, and strain rate) are summarized in Table 1. The TEM observations have been performed with three TEM microscopes: a Philips CM30 TEM operating at 300 kV and a Tecnai G2-20 TEM operating at G2-200 kV both at the centre commun de microscopie (CCM) of the University of Lille, France, and a Jeol 200CX TEM operating at 200 kV at MATEIS laboratory in University of Lyon, France.
3. Results
Lizardite and antigorite samples recovered from deformation experiments were observed at both sample and thin section scale. The cylinders present distributed deformation at the millimeter scale (Figure 2). Deformation is homogeneous down to dimensions commensurate with grain size. Heterogeneities are noted between large grains with intracrystalline deformation and comminuted grains at grain boundaries, except in the hot pressed nondeformed antigorite control sample. No traces of dehydration were observed. TEM observations are detailed below for each serpentine mineral.
Figure 5. TEM bright-field images of kink bands in lizardite grains. (a) Kink band in sample D1064 where brightfissures (white arrows) corre-spond to (001) cleavage planes and correcorre-sponding SAED pattern (inset). (b and c) Triangular lizardite sectors resulting from intense kinking with nonparallel kink-band boundaries (noted KBB and represented as orange dotted lines) in samples D1130 and D1131, respectively, and corre-sponding SAED pattern (inset). The compression axis is reported as red arrows. The scale bars are all 0.2μm.
3.1. Microstructures in Lizardite
TEM observations on the four lizardite samples show large elongate grains surrounded by interconnected regions with extremelyfine grain size formed by grain size reduction during deformation (Figure 3). These regions represent approximately 30% of the TEM specimens regardless of the confining pressure (from 1.5 to 8 GPa; Figures 3a and 3c). The average grain size of large grains is comparable to the initial grain size in the starting material. In samples with initial average grain size<10 μm, the average grain size in the recovered samples is around 5μm (Figures 3a and 3c). After deformation of the sample with coarse-grained starting material (initial grain size between 50 and 100μm), the grains are larger and regions of fine-grained material are proportionally less (Figure 3b).
Within crystals, deformation of lizardite was accommodated by the activation of two major mechanisms over the entire range of experimental conditions: (1) bending and slip along (001) planes and (2) intense kinking (Figures 3–5). These mechanisms are strongly controlled by deformation on the weak basal planes where only weak OH―O bonds are broken. Almost all observed lizardite grains present at least one of these microstructural features.
In lizardite crystals deformed by kinking, the kink-band boundaries (KBB) between the rotated and unrotated portion of the crystals (also called kink plane, see orange dotted lines on Figures 4b and 4c) show no clear evidence of preferred orientation related to the principal compression direction. In some regions KBB are not distributed parallel to each other, leading to the formation of triangular lizardite sectors (Figures 5b and 5c). Kink formation is accompanied by cracks at the edges of the kinked grains (see white arrows on Figures 4b and 5a) probably formed by elastic relaxation. Some microcracks are present within regions where intense grain size reduction took place (Figures 3c and 4b) in the lizardite specimen deformed at the lowest confining pressure (~1.5 GPa). These microcracks represent a marginal microstructural feature and likely formed during the last deformation cycle of this sample when deformation of the cell assembly led to a decrease in confining pressure down to 0.8 GPa or as a result of cold decompression. The margins of lizardite crystals affected by bending and cleavage along (001) planes are delaminated into thin slabs (0.2 to 0.5μm wide; Figures 4a and 4b) in the case of small starting grain size; similar slabs are slightly larger (0.5 to 0.8μm wide; Figure 3b) in the case of the coarser starting grain size (Figure 4c). The slabs are also affected by slip along basal planes. The division of lizardite crystals into slabs forms transitional zones—an intermediate step before the further delamination of crystals. Comminution reaches a stage where partial amorphization of the material is observed (Figure 4c). Whether amorphization is due to deformation or to the TEM thin section preparation and electron irradiation damage cannot be discerned.
With dominant basal slip resulting from two independent intracrystalline slip systems, lizardite is far from satisfying the von Mises criterion, which requiresfive independent slip systems for ensuring isochoric deformation compatibility in a crystal or an aggregate [Mises, 1928]. Delamination and comminution at grain boundaries are therefore necessary for accommodating geometrical incompatibilities in the deforming lizardite aggregate.
Figure 6. Low-magnification TEM bright-field images of recovered antigorite samples. (a) Sample D0539s hot pressed 8 h at 4 GPa and 350°C, and (b) sample D0539 deformed at 1 and 4 GPa and several temperatures with a starting grain size between 5 and 30μm. Areas of intense grain size reduction, when visible on the image, are bounded by white dotted lines and are labeled by white asterisks. The compression axis is reported as red arrows. The scale bars are all 1μm.
3.2. Microstructures in Antigorite
TEM observations were carried out on two antigorite samples: a hot pressed comparison sample (cold compressed to 4 GPa and then heated at 350°C for 8 h without applying deformation other than compression) and the sample recovered after deformation cycles (cold compressed to the target pressure, heated, and then deformed at high P and T during several cycles, see Table 1 for details on deformation conditions). These observations allow estimating the effect of cold compression on the polycrystalline sample and characterizing the microstructures actually induced by high-pressure and high-temperature
Figure 7. TEM bright-field images of recovered hot pressed and deformed antigorite samples D0539s and D0539. (a) Large antigorite grain with little diffraction contrast in undeformed sample D0539s. Mechanisms involved in plastic deformation include (b) glide along weak basal and conjugate planes. (c and d) Thin kink-band in a large grain oriented with [010] in the observation plane, and (e) SAED pattern corresponding to the large grain in Figures 7b and 7c. The differential stress compression axis is reported as red arrows. The scale bars are all 0.2μm.
deformation. In the hot pressed sample, antigorite grains are large and show angular shape (Figure 6a). The diffraction contrast shows that few defects have accumulated in grains, and the amount of small grain size material, probably partly inherited from crushing and partly formed during the cold compression, is estimated on the TEM thin section at ~10% (Figures 6a and 7a). In comparison, the deformed antigorite sample shows smaller grains with smoother shapes and stronger orientation contrast inside grains, indicating intracrystalline deformation (Figure 6b). The proportion of small grain size material is higher in the deformed sample (40–50%) than in the hot pressed sample (Figure 6), suggesting that the small grain size material created during cold compression may have played a role during thefirst stage of the experiment, but is not sufficient to accommodate the deformation. These observations indicate that microstructures in the recovered sample were activated during the deformation of the specimen and are not a reactivation of defects created during cold compression.
Within antigorite crystals, deformation occurred mainly by slip along weak planes, and to a smaller extent by bending (Figures 7 and 8). In most of deformed grains, conjugate glide planes are observed in crystals with a and c axis close to the thin section plane. The conjugate planes are inclined with respect to the basal planes with orientations that match those of the (101) and 101planes (Figure 7a). Glide along these conjugate planes leads to delamination into thin slabs (~ 0.2μm wide) and formation of extremely fine-grained regions (Figure 7a). Large grains oriented with the b and c axes close to the TEM section plane (Figures 7b and 7c) appear different from those with a and c axis close to the thin section plane (Figure 7a). Thin planar contrast similar to stacking faults perpendicular to the c* direction are interpreted as traces of the conjugate glide systems along (101) and 101that both appear parallel to (001) in this orientation (see diffraction pattern in Figure 7d). The delamination by glide into thin slabs is not observed in this orientation, suggesting that [010] is a“hard” direction in the crystal where stresses were relaxed only by the formation of thin kink bands (Figures 7b and 7c). Those kinks are a marginal microstructure accounting for a negligible amount of the total deformation.
Figure 8. TEM bright-field images of recovered antigorite sample D0539. Major mechanisms involved in intracrystalline plastic deformation are displayed. (a) Antigorite grain showing conjugate glide planes on the left and bending on the right. (b) Close-up on the (101) and 101glide planes showing contrast akin to those of elastic strain around dislocations. (c) High-resolution image with (001) lattice fringes of dislocation-like contrast in the glide plane and (d) high-resolution image on the bent region with (001) lattice fringes. Fourier transform of areas on both sides of a kink zone are shown.
The dominant conjugate (101) and 101glide planes are lined with localized elastic contrasts that are reminiscent of those of dislocations (Figure 8). In spite of attempts to image potential dislocations by both conventional and high-resolution transmission electron microscopy (HRTEM), we could not constrain the dislocation Burgers vector or line orientation. In the conjugate (101) and 101planes, the smallest lattice translation is ~0.9 nm in the [010] direction, but little glide is observed in that direction. Glide occurs mainly in directions close to101or [101] in (101) and 101, respectively, which would correspond to a Burgers vector of ~4 nm (Figure 9). Such high values imply that dislocation glide, if active, likely occurs by dissociation into partial dislocations and stacking faults, the structure of which remains to be determined. Numerous elastic contrasts along glide planes would suggest piling of dislocations.
Similar to lizardite, the two dominant slip systems in antigorite are not sufficient to satisfy the von Mises criterion, and comminution at grain boundaries occurs to accommodate geometrical incompatibilities during deformation of the aggregate. Partial amorphization may have resulted from deformation, or TEM
preparation and irradiation damage. Comminution may have been eased in some areas by the presence of minor chrysotile nanotubes that were present in the starting material due to retrogressive alteration at grain boundaries [Auzende et al., 2002].
4. Discussion
4.1. Deformation Mechanisms and Rheology of Lizardite and Antigorite in Experiments
The two serpentine minerals studied, lizardite and antigorite, differ not only in their rheological properties but also in their microstructures. Mechanical results reveal a strain rate-dependent rheology for antigorite, whereas lizardite has a plastic behavior with strain rate-independent yield stress [Hilairet et al., 2007; Amiguet et al., 2012]. Differences in intracrystalline deformation mechanisms provide a likely explanation for this observation. Comminution and delamination at grain boundaries are observed in both materials and cannot alone account for their different rheology.
Lizardite rheology at high pressure is highly anisotropic and is not (or very little) influenced by temperature, pressure, or strain rate. Lizardite deforms plastically, and its low strength (20–200 MPa with most likely value around 100 MPa) can be assumed to be constant from laboratory to tectonic strain rates in the P-T conditions of subduction zones [Amiguet et al., 2012]. At low pressures (<300 MPa) brittle deformation occurs, and strength shows a positive pressure dependence. These frictional processes control the rheological behavior in the pressure conditions at which deformation is likely to occur in an oceanic context [Raleigh and Paterson, 1965; Murrel and Ismail, 1976; Escartín et al., 1997b; Moore, 2004]. In our study, lizardite’s low strength results from distributed plastic deformation by grain size reduction at grain boundaries. This grain size reduction
Figure 9. Structures of serpentine minerals and their weak glide planes for (a) antigorite, where thin lines show the trace of the (101) in blue, 101in red, and (001) planes in black. Broken lines show the likely atomistic structure following weak OH bond along dashed lines, cutting strong ionic Mg―O (lines) of octahedral layers and strongest ionocovalent Si―O bonds (thick lines) at tetrahedral layer reversals. This minimizes the number of strong bonds that have to be broken, and (b) Lizardite where thin black line show the trace of (001) planes.
occurs by the activation of two major mechanisms, equivalent in proportion: glide along the basal plane, and kink-band formation (Figure 10).
Glide along the basal plane was documented as the dominant deformation mechanism in plastically deformed lizardite during shear at 1 GPa [Hirauchi et al., 2010a]. At lower pressure, near the semibrittle/ductile transition, when deformation of lizardite is mainly accommodated by cracks opening, glide along the basal plane remains a minor mechanism [Escartín et al., 1997b]. Kink-band formation is the second major microscopic feature we characterized in our samples deformed in compression. Few kinks have been previously observed in lizardite deformed in shear [Hirauchi et al., 2010a] and in partially dehydrated lizardite samples at high strain [Hirose et al., 2006; Viti and Hirose, 2009].
The present microstructural observations confirm that glide along the basal plane is the easiest lizardite deformation mechanism in the plastic regime above 1 GPa, particularly in shear geometry. Kink-band formation in phyllosilicates is strongly geometry and strain dependent in this regime (Figure 10). The strength of lizardite in aggregates is around 100 MPa [Amiguet et al., 2012], close to the values of sheared foliated mica aggregates [Misra and Burg, 2012]. This is consistent withfirst-principle calculations that give low
(80–120 MPa) critical resolved shear stress (CRSS) for translation in the basal plane [Amiguet et al., 2012]. Deformation experiments on micas, phyllosilicates with comparable structure to lizardite, yielded similar rheology and microstructures [Kronenberg et al., 1990; Mares and Kronenberg, 1993; Misra and Burg, 2012]. The strength of micas is low, typically 20–70 MPa, when recorded in samples oriented to promote glide on the
Figure 10. Microstructures involved in plastic deformation of lizardite and antigorite serpentines depending on deforma-tion geometry and their associated strengths. In the strain rate versus stress diagram lizardite strength is extrapolated to natural strain rate using the plasticflow law determined in Amiguet et al. [2012], and antigorite strength is extrapolated from the strain rate-dependent law from Hilairet et al. [2007].
basal plane (basal plane at 45° to compression axis) and higher in samples oriented to promote kink-band formation (basal plane parallel to compression axis) [Kronenberg et al., 1990; Mares and Kronenberg, 1993]. At laboratory strain rates (10 6–10 4s 1), antigorite is about 3 to 10 times stronger than lizardite under comparable conditions at confining pressures below 500 MPa [Escartín et al., 1997b] and above 1 GPa [Hilairet et al., 2007; Chernak and Hirth, 2010; Amiguet et al., 2012]. A similar rheology contrast was observed in shear experiments at 1 GPa, 250–300°C, and high strain rates in the range 2.5–60 10 5s 1[Hirauchi and Katayama, 2013]. Antigorite rheology in the ductile regime above 1 GPa has beenfitted with a positive dependence of strength on strain rate using either a Peierls law or a power law [Hilairet et al., 2007]. A semibrittle deformation regime was suggested for antigorite from experiments below a confining pressure of 1.5 GPa where antigorite strength increases with increasing confining pressure and decreasing temperature [Chernak and Hirth, 2010].
Above 1 GPa, antigorite crystals display intense intracrystalline deformation and grain size reduction at grain boundaries. The major intracrystalline deformation mechanism responsible for grain size reduction is glide on conjugate (101) and 101planes, close to the101and [101] directions contained in the (100) plane where major deformation is observed. In grains with observation plane oriented near (010), negligible deformation is observed, [010] glide in the basal plane is little activated, and thin kink bands have formed to accommodate plastic deformation (Figure 10). The dominant conjugate slip systems result in an average glide direction along [100]. Easy deformation along [100] was previously suggested from electron backscatter diffraction (EBSD) observations on experimentally deformed antigorite serpentinite from the strong alignment of [100] axis close to the shear direction [Katayama et al., 2009].
Easy glide along101and [101] directions in conjugate (101) and 101planes is consistent with the structure of antigorite (Figure 9), corresponding to planes where a minimum number of strong ionic Mg―O and ionocovalent Si―O bonds have to be broken during glide. In the 101 and [101] directions, the strongest Si―O―Si bonds can be torn one by one at tetrahedral layer reversals of the stacking sequence by screw-type dislocations, further minimizing the number of new strong bonds to break during slip. In the opposite, glide along [010] requires the simultaneous breaking of several Si―O―Si bridges, making this slip direction unfavorable. Glide along the (100) plane is disfavored since it requires either simultaneously breaking a large number of Mg―O bonds in the fully occupied trioctahedral layer, or tearing of Si―O―Si at each tetrahedral layer reversal. Modeling of the associated CRSS for the different possible slip systems from first principles would require intense computational effort due the large unit cell of antigorite, and further TEM determination of potential dislocation substructure.
HRTEM imaging suggests minute amorphization at the nanometer scale (Figure 8d). Glide might also be controlled by the formation of thin amorphous layers in the complex antigorite structure, which may explain the low-temperature dependence of antigorite deformation. Dominant dislocation glide and piling in conjugate (101) and 101planes suggests that Peierls law may be more appropriate than power law [Katayama and Karato, 2008] for describing the mechanical results on antigorite [Hilairet et al., 2007]. Extrapolation to natural strain rates, however, depends little on the chosen law [Amiguet et al., 2012].
4.2. Microstructures in Naturally Deformed Serpentinites
Deformation mechanisms in naturally deformed antigorite serpentinites have been previously inferred from microstructural features observed at thin section scale by optical microscopy or EBSD. Undulatory
extinctions, subgrains and strong shape preferred orientation and CPO are usually attributed to
intracrystalline deformation by dislocation glide in antigorite basal planes. Few previous TEM observations were motivated mainly to relate modulation variability in antigorite crystals with metamorphic grade and have little explored intracrystalline deformation microstructures [Auzende et al., 2002, 2006]. CPOs in naturally deformed antigorite serpentinites have been interpreted as dislocation glide within the (001) plane with contradictory evidences on the dominant slip direction. CPOs in natural samples imply the activation of two possible slip systems, [100](001) [Bezacier et al., 2010; Van de Moortèle et al., 2010; Brownlee et al., 2013] and [010](001) [Hirauchi et al., 2010a; Soda and Takagi, 2010; Nishii et al., 2011; Brownlee et al., 2013], or the contribution of both [Jung, 2011; Padrón-Navarta et al., 2012; Brownlee et al., 2013]. To our knowledge, there is no report of lizardite in association with ductile deformation in nature, but only with brittle failure
Understanding antigorite serpentinite deformation through CPO relies not only on considering the orientation of [100] and [010] axis with respect to the lineation but also on the distribution of (001) planes. Numerical simulations suggest no dependence of antigorite CPO on the slip systems activity [Padrón-Navarta et al., 2012], and the variability in (001) distribution could indicate alternative processes to deformation for the formation of CPO [Brownlee et al., 2013]. CPO in natural antigorite serpentinites do not result exclusively from deformation and could ensue from the topotactic growth of antigorite from olivine, thus inheriting the CPO of olivine in the peridotite protolith [Boudier et al., 2010]. Two topotactic relationships were identified as (001)Atg//(100)Oland (001)Atg//(010)Ol, suggesting that a growth of antigorite from an oriented olivine will result in a bidirectional or girdle distribution of (001) antigorite planes. The (001) girdle distributions in CPO of natural antigorite serpentinites associated with strong antigorite b axis parallel to the lineation have been interpreted as [010](001) dislocation glide [Soda and Takagi, 2010; Nishii et al., 2011] but are more likely to result from topotactic replacement than dislocation creep [Boudier et al., 2010]. Transition between CPOs with (001) girdles and those with marked concentrations of (001) perpendicular to the foliation with increasing deformation have been observed in naturally deformed serpentinites [Hirauchi et al., 2010a; Padrón-Navarta et al., 2012; Brownlee et al., 2013] and were demonstrated by shear deformation experiments [Katayama et al., 2009].
Textural observations indicate that CPOs with a strong (001) pole concentration perpendicular to the foliation are clearly related to intense ductile deformation [Hirauchi et al., 2010a] and generally show strong orientation of [100] direction along the lineation [Bezacier et al., 2010; Van de Moortèle et al., 2010; Padrón-Navarta et al., 2012; Brownlee et al., 2013]. The easy activation of [101]101and101(101) conjugate systems observed in our experiments results in an apparent slip system akin to [100](001) and accounts for the CPO observed in foliated antigorite mylonites. It is consistent with the observation of subgrain boundaries at low angle to the (100) plane, with rotation around the [010] direction, described as deformation features associated with dislocation creep [Padrón-Navarta et al., 2012]. The good agreement between
microstructures in naturally and experimentally deformed antigorite validates the extrapolation of rheology determined at high pressure and laboratory strain rates to natural conditions [Hilairet et al., 2007], for which Peierls law may be preferred to power law [Amiguet et al., 2012], because it corresponds better to high-stress low-temperature deformation with no diffusion-assisted dislocation climb, and stacking of dislocations at high densities in the glide planes.
5. Geophysical Implications
5.1. Serpentine Strength in Oceanic Lithosphere and Subduction
Antigorite and lizardite specimens experimentally deformed using the Deformation-DIA display differences in their microstructures. Lizardite deforms in the plastic regime with the activation of two major mechanisms over the entire range of experimental conditions: glide along the weak basal planes and microkink leading to the delamination of crystals and the formation of interconnected regions withfine-grained material. These microstructures are associated with a plastic strength of ~100 MPa (20–200 MPa) showing little to no dependence on temperature and strain rate, likely representing the upper bound of lizardite strength in natural conditions. Glide is active on different planes of antigorite, owing to its structural complexity, and is consistent with rheologies implying strong stress dependence on strain rate, like a Peierls law proposed by Amiguet et al. [2012]. Considering that microstructures in naturally and experimentally deformed serpentines display strong similarities, extrapolation of rheological laws determined from deformation experiments to geological conditions can be attempted.
At natural strain rates, serpentine minerals exhibit brittle behavior at low pressures (Figure 11). The mechanical strength of lizardite is described by its unique frictional properties instead of the general Byerlee’s law that overestimates the shear strength [Reinen et al., 1994; Escartín et al., 1997b]. The presence of lizardite in oceanic transform faults may explain the relative weakness of these faults. In the deep part of crustal faults and in the upper part of subduction zones where low-T serpentine could be present, lizardite shows a semibrittle behavior moving toward a ductile behavior associated with shear stress near 100 MPa (Figure 11a) [Amiguet et al., 2012]. The brittle-ductile transition occurs at low pressure and accounts for shallow ductile deformation in serpentinized detachment faults at the top of megamullions [Tucholke et al., 1998; Cannat et al., 2009].
The transition between the brittle and ductile regimes for antigorite occurs around 30 km depth in subduction zones (Figure 11b). At lower depths, ductile deformation of antigorite occurs by dislocation creep, potentially assisted by dissolution recrystallization processes. The rheology of antigorite is bound in the deep part of subduction zones by the power law or Peierls creep, and the Newtonian behavior suggested by dissolution-precipitation creep.
5.2. Dislocation Glide and Creep Versus Frictional Behavior and Solution Creep
The deformation experiments performed here were designed to favor plastic behavior. Antigorite mainly deforms by glide along conjugate planes resulting in an apparent global slip system akin to [100](001), and lizardite by basal glide. Delamination of serpentine crystals by glide occurs because of the limited number of slip systems and results in the formationfine-grained interconnected regions at grain boundaries. In lizardite, as well as in micas, high stresses are observed in locked geometries and textures when the glide plane is perpendicular to the compression axis, leading to brittle failure. Plastic deformation at low stresses (typically 20–200 MPa) is observed when the glide plane is at a lower angle to the compression axis or shear direction [Kronenberg et al., 1990; Mares and Kronenberg, 1993; Amiguet et al., 2012; Misra and Burg, 2012]. Grain rotation in the shear direction should result in rapid weakening of lizardite serpentinites and to a transition from brittle to plastic behavior.
Semibrittle deformation associated with microcracking along basal planes was observed at slightly higher stresses [Chernak and Hirth, 2010] than ductile deformation of antigorite [Hilairet et al., 2007] at the same pressure and temperature conditions. Similar to lizardite, we suggest that the transition between plastic and frictional behavior is likely controlled by textural parameters such as grain size and orientation distribution, since frictional behavior is more readily observed during experimental deformation of rock cores [Rutter et al., 2009; Chernak and Hirth, 2010], and plastic behavior in the sintered aggregates studied here. Brittle-ductile transition occurs over a pressure range of 0.5–1.5 GPa (or depth range of 15–45 km) at experimental strain rates. This corresponds to the depth interval of the transition between the frictional (seismogenic) and ductile (aseismic) behavior along the subduction interface.
Atomic force microscopy (AFM) observations of the antigorite (001) surface along the basal plane suggest that [010] is an easy direction for frictional glide, and [100] an uneasy direction [Campione and Capitani, 2013]. The present observations suggest the opposite scheme for plastic deformation, with easy glide along [100], and uneasy glide along [010]. This may account for the complex mechanical behavior of the serpentinized plate interface near the seismogenic-aseismic transition [Hilairet et al., 2007; Campione and
Figure 11. Yield strength envelope for (a) lizardite and (b) antigorite as a function of depth following a model of 20°C/km for lizardite and the P-T profile i50 from Conder [2005] for antigorite, for strain rate of 10 10s 1[Behnsen and Faulkner, 2012]; [Reinen et al., 1994]; [Amiguet et al., 2012].
Capitani, 2013]. Further determination of active deformation mechanisms and associated CRSS in antigorite is needed to quantitatively resolve these issues, but this remains a challenge for both computational simulations and microstructural observations because of the structural complexity of antigorite. Intense delamination of serpentine crystals intofine-grained interconnected regions is a major
microstructural feature in experimentally deformed specimens and is not present in natural samples from subduction zones. This variation in microstructures implies that other deformation mechanisms than dislocation glide have taken place at low strain rates and stresses in geological conditions. Some
microstructural features in naturally deformed serpentinites have been attributed to dissolution-precipitation processes [Hoogerduijn Strating and Vissers, 1994; Andréani et al., 2005; Wassmann et al., 2011; Padrón-Navarta et al., 2012] that may be acting as a low-temperature equivalent of dynamic recrystallization for serpentine minerals. Therefore, laws that take into account both dislocation and dissolution-precipitation creep should be investigated to describe serpentinite mechanical behavior in deep and relatively hot natural contexts.
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Acknowledgments
We thank Sarah Brownlee and Ikuo Katayama for their constructive com-ments which improved the manuscript. This study was supported by ANR pro-ject SUBDEF grant ANR-08-BLAN-0192 to BR. Deformation experiments were performed at GeoSoilEnviroCARS (sector 13), Advanced Photon Source (APS), Argonne National Laboratory. GeoSoilEnviroCARS is supported by the National Science Foundation-Earth Sciences (EAR-1128799) and Department of Energy-Geosciences (DE-FG02-94ER14466). Use of the Advanced Photon Source was sup-ported by the U.S. Department of Energy, Office of Science, Office of Basic Energy Sciences, under contract DE-AC02-06CH11357. FIB thin sections were prepared at CLYM (CNRS, FED 4092) in Lyon (France). TEM was performed both at the CCM of the University of Lille (France) and at MATEIS laboratory in University of Lyon, (France). The TEM national facility in Lille (France) is supported by the Conseil Regional du Nord-Pas de Calais, the European Regional Development Fund (ERDF), and the Institut National des Sciences de l’Univers (INSU, CNRS).
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