Download Rheology and Tectonic Significance of Serpentinite

Rheology and Tectonic
Significance of Serpentinite
Sheared serpentinite,
Oman
Greg Hirth1 and Stéphane Guillot2
1811-5209/13/0009-107$2.50
DOI: 10.2113/gselements.9.2.107
S
erpentinites occur in many active geologic settings and control the
rheology of the lithosphere where aqueous fluids interact with ultramafic rocks. The crystal structure of serpentine-group minerals results in
diagnostic physical properties that are important for interpreting a wide range
of geophysical data and impart unique rheological behaviors. Serpentinites
play an important role during continental rifting and oceanic spreading, in
strain localization along lithospheric strike-slip faults, and in subduction zone
processes. The rheology of serpentine is key for understanding the nucleation
and propagation of earthquakes, and the relative weakness of serpentinite
can significantly affect geodynamic processes at tectonic plate boundaries.
KEYWORDS : seismicity, ocean–continent transition, seafloor,
subduction zone, serpentinization
INTRODUCTION
Plate tectonics is controlled by a combination of ductile
flow and frictional resistance in fault zones. At plate boundaries, where deformation is localized along centimeter- to
kilometer-scale shear zones, the influence of serpentinite
on tectonic processes is linked to its unique rheological
properties. While the rheological weakness of serpentinite
relative to other rocks has been appreciated through field
observations in various geological settings, such as along
slow-spreading oceanic ridges, along strike-slip faults, and
in subduction zones, the quantification of its influence
on tectonic processes requires an understanding of the
factors that control ductile flow and frictional resistance.
Experimental observations constrain these properties and
highlight the rheological behaviors that are important for
understanding the role of serpentinization in strain localization in the lithosphere, its role in both fault creep and
dynamic fault rupture, and the influence of the dehydration of serpentinite on the seismicity of subducting slabs
and the exhumation of high- to ultrahigh-pressure rocks.
RHEOLOGICAL PROPERTIES
OF SERPENTINITE
Since the pioneering work of Raleigh and Paterson (1965),
who documented how dehydration reactions in serpentinite result in weakening and strain localization, numerous
experimental studies have been conducted to investigate
the physical properties of the serpentine minerals. These
1 Department of Geological Sciences, Brown University
Providence, RI 02912, USA
E-mail: [email protected]
2 CNRS, Institut des Sciences de la Terre, Université de Grenoble
1381 rue de la Piscine, 38041 Grenoble cedex, France
E-mail: [email protected]
E LEMENTS , V OL . 9,
PP.
107–113
studies provide fundamental data
for understanding fracture and
plastic flow, friction, and the
role of dehydration reactions on
fault rheology.
Fracture and Plastic Flow
Both brittle and crystal plastic
de for mat ion pro cesses a re
strongly controlled by the anisotropic properties of serpentine.
Dislocation glide (a deformation
process where strain is accommodated by stress-driven motion of
strike-slip fault,
crystal defects) is easy along the
basal plane of serpentine minerals,
leading to ductility of serpentinite
at low temperatures. Nonetheless,
experimental studies have produced apparently contradictory results regarding high-temperature flow behavior.
Several studies indicate that serpentinites deform by macroscopically ductile processes at high pressure (Hilairet et
al. 2007 and references therein). In contrast, Chernak
and Hirth (2010) observed localized faulting under hightemperature/high-pressure conditions, although these
authors document macroscopically ductile flow at strains
less than ~0.1 to 0.2. These results can be reconciled by
considering the von Mises strain compatibility criterion,
which states that five independent slip systems are required
to accommodate homogeneous flow of a polycrystalline
material. Dislocation glide on the basal plane and kinking
can accommodate a limited amount of deformation, but
to achieve higher strains additional processes are required.
Experimental studies on serpentinite (Chernak and Hirth
2010) indicate that microcracking initiates at moderate
strains, leading to a semibrittle rheology for which strength
remains pressure sensitive. Thus, while anisotropy in
plasticity facilitates dislocation glide at low temperatures,
this process is insufficient to promote fully crystal plastic
deformation at high strain.
Flow laws for antigorite (a high-pressure type of serpentine)
are reported by Hilairet et al. (2007). Based on analysis
of the stress exponent (n) and the ratio of strength to
confi ning pressure [(σ1 – σ3 )/σ3 ], they concluded that
samples at 4 GPa deformed by dislocation creep (where
the motion of crystal defects is thermally controlled by
diffusive processes) while samples at 1 GPa deformed by
dislocation glide. These data represent results of pioneering
high-pressure deformation experiments, though the strains
in the experiments of Hilairet et al. are below those for
which semibrittle processes were observed in the study by
Chernak and Hirth (2010). The observation of a power-law
relationship between stress and strain rate (i.e. ė ∝ σn with
n ≈ 3) suggests that deformation is limited by dislocation
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climb (the diffusive process that generally controls dislocation creep), which requires rapid diffusion of all components of the crystal structure (i.e. Si, Mg, O, and H for
serpentine). However, this seems unlikely at the temperature (down to 200 oC) of these experiments. Nonetheless,
the high strain-rate sensitivity of serpentine deformation
has important implications for geologic processes.
A
The tendency for strain localization during semibrittle flow
(where deformation is accommodated by a combination of
brittle and ductile processes) of serpentine has important
implications for understanding the stability of fault slip
when an earthquake rupture propagates into a serpentinized region. Constraining the conditions where strain
localization occurs is also important for extrapolating lab
data to natural conditions, owing to the trade-offs between
effective viscosity and shear zone width (e.g. Wada et al.
2008). Escartín et al. (1997) documented a transition from
localized to distributed deformation with an increase in
confi ning pressure from 200 to ~400 MPa at room temperature. They inferred that the transition occurred when
the bulk sample strength became less than the frictional
strength. However, with increasing temperature at constant
pressure, a transition back to localized semibrittle deformation is observed (Chernak and Hirth 2010). The transition
back to localized semibrittle behavior at high temperature
suggests that the temperature dependence of the friction
coefficient is greater than the temperature dependence of
the bulk sample strength.
While the details of semibrittle flow remain unresolved,
the observation of a deformation-induced alignment of
minerals [i.e. lattice-preferred orientation (LPO)] in both
experimental and natural samples of serpentinite (e.g.
Katayama et al. 2009; Bezacier et al. 2010) has important
implications for interpretating seismic anisotropy (FIG. 1).
Serpentine has a high compressibility normal to the basal
plane (parallel to the c-axis), which leads to a low isotropic
P-wave velocity (Vp) (e.g. Bezacier et al. 2010). Similarly,
the shear modulus (ratio of shear stress to shear strain)
parallel to the basal plane is very low. Thus, where a strong
LPO arises owing to deformation, the elastic anisotropy
leads to strong seismic anisotropy (Katayama et al. 2009).
Nonetheless, the limited knowledge of elastic properties makes interpretation of seismic data nonunique. For
example, minor amounts of lizardite/chrysotile and/or LPO
can strongly impact the Vp /Vs ratio and the magnitude of
shear wave splitting (e.g. Bezacier et al. 2010). In addition,
the magnitude of shear wave splitting depends on both the
strength of LPO and the width of the anisotropic region
that waves pass through. Thus, if strain is highly localized under geologic conditions, shear wave splitting will
be negligible.
Friction
Examining the frictional behavior of fault materials
over a range of sliding velocities is essential for understanding the dynamics of stress evolution and slip during
earthquakes. Phyllosilicate minerals are often inferred to
control the behavior of creeping sections of faults (e.g.
Moore and Rymer 2007). Friction experiments on serpentinite document velocity-strengthening behavior (which
promotes aseismic fault slip) at tectonic displacement rates
(e.g. Reinen et al. 1994), consistent with this hypothesis.
Experiments also indicate a decrease in the friction coefficient (µ) of antigorite with increasing temperature. Reinen
et al. (1994) measured µ = 0.5–0.85 for antigorite at room
temperature; Moore et al. (1996) determined µ in the
range of 0.4–0.6 for antigorite gouge at temperatures of
25–194 °C. Chernak and Hirth (2010) determined values
of µ from 0.1 to 0.35 at temperatures of 400–550 °C. In
E LEMENTS
B
C
(A) Photomicrograph (crossed polarizers) of an antigorite sample deformed in the lab at 300 oC and 1 GPa.
(B) Contoured, lower-hemisphere stereographic projections
showing the lattice-preferred orientation (LPO) of antigorite with a
[001] maximum oriented perpendicular to the shear plane (the
shear plane pole is horizontal in the pole figure) and a maximum of
poles to (100) oriented parallel to the slip direction. (C) Anisotropic
velocity imparted by this LPO: (LEFT) contours of P-wave velocity;
( MIDDLE) contours of % anisotropy of S-wave velocity; ( RIGHT) polarization direction of shear waves. After Katayama et al. (2009)
FIGURE 1
general, theories for rock friction suggest that µ should
be relatively insensitive to temperature (e.g. Scholz 1990).
In adhesion theory, µ is related to the ratio of the shear
strength of asperity contacts to the indentation hardness
of the asperities (FIG. 2). A lack of temperature dependence
for µ is inferred based on the notion that the processes that
control the shear strength of asperities are the same as
those that control indentation hardness (e.g. Scholz 1990).
However, when considering a strongly anisotropic material,
another possibility arises. Assume that the basal planes of
phyllosilicate grains at asperity contacts are dominantly
aligned parallel to the fault surface. In this scenario, the
shear strength (e.g. controlled by dislocation glide) may
be significantly lower than the indentation hardness (e.g.
controlled by semibrittle flow). The temperature dependence of basal slip would be expected to be greater than
that for indentation hardness, resulting in temperature
dependence and a lower value for µ.
High-velocity friction experiments on serpentinite
document dramatic dynamic weakening and dehydration
at slip velocities above ~0.1 m/s (see Kohli et al. 2011 and
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A
B
C
Schematics of a fault surface (shown perpendicular
to the fault plane) in serpentinite at different scales.
(A) The fault surface is rough, making contact only at asperities.
(B) The basal planes of serpentine grains (dashed lines) are aligned
parallel with the fault surface at asperities. (C) The shear deformation of asperities is accommodated by dislocation glide (as represented by symbols parallel to dashed lines). Grains are much
stronger for deformation normal to the fault zone, resulting in a
low, temperature-dependent coefficient of friction.
FIGURE 2
references therein). These results indicate that a material
exhibiting velocity-strengthening behavior at tectonic
displacement velocities (~10 -9 m/s) may become dynamically unstable at near-seismic slip velocities after only a
few millimeters of displacement (FIG. 3). Such behavior
may be particularly important for understanding processes
that control the dynamics of great subduction earthquakes,
which likely propagate across patches that creep during
interseismic periods. When the asperity-contact lifetime
during frictional slip is short compared to the time it
takes to diffuse heat away from asperities, the resulting
“flash heating” can produce dramatic weakening. Kohli
et al.’s experiments indicate that weakening occurs at a
critical weakening velocity (Vw), at which dehydration
of serpentine to talc occurs at asperities (500–700 o C).
Based on extrapolation of Kohli et al.’s data, Vw is about
25 mm/s at an ambient T = 300 °C. This analysis suggests
a possible explanation for the spectrum of fault-slip behaviors observed along fracture zones and subduction zones,
where the alteration of mantle peridotite produces serpentinites. The effectiveness of dynamic weakening depends
strongly on the width of the actively deforming zone. Thus,
an understanding of how deformation localizes in serpentinite during interseismic deformation remains critical for
constraining the significance of dynamic weakening.
(Jung et al. 2004) suggest that antigorite dehydration can
at least generate microseismicity. However, Chernak and
Hirth (2010) observed macroscopically ductile flow and
velocity-strengthening behavior in dehydrated antigorite at
1.5 GPa and 700 °C, challenging the hypothesis that intermediate-depth earthquakes result directly from dehydration embrittlement. Motivated by these results, Chernak
and Hirth (2011) conducted tests in which temperature
was ramped across the dehydration boundary, reasoning
that earthquake-like instabilities would be most likely if
samples dehydrated at high load under conditions where
strain was already localized. No frictional instabilities were
observed; rather, stress relaxed stably over several minutes,
providing experimental documentation of stable fault slip
during dehydration reactions. Similar interpretations were
made based on the lack of detectable acoustic emissions
during dehydration experiments on antigorite in a multianvil device (Gasc et al. 2011). These observations suggest
that a mechanism other than dehydration may be responsible for intermediate-depth earthquakes (see discussion
and references in Chernak and Hirth 2011). One possibility
is that the migration of fluids produced via dehydration
promotes seismogenic failure in the surrounding unaltered
rocks. Alternatively, earthquakes may nucleate via
thermally induced viscous flow instabilities. In this case,
the fi ne-grained and polycrystalline products of dehydration reactions may promote the initial strain localization
required for the thermally induced viscous instability.
GEOLOGICAL OBSERVATIONS
Since their discovery on the seafloor of the central Atlantic
in the 1970s, serpentinites have been recognized worldwide
in different active tectonic settings. Due to their unique
physical properties, described above, serpentinites play a
major role in strain localization.
The Ocean–Continent Transition
and Slow-Spreading Ridges
Continental rifting and subsequent ocean-ridge spreading
involve extensional faulting, exhumation of subcontinental mantle, and magmatism, processes that reflect the
tendency for localized deformation in the lithosphere.
Dehydration Embrittlement
Dehydration embrittlement (where fluids resulting from
dehydration reactions induce brittle deformation) has
become the accepted mechanism invoked to explain
intermediate-depth seismicity (e.g. Hacker et al. 2003).
Several issues are important for the application of these
results to higher pressures. First, the canonical experiments were conducted at confi ning pressures where the
Clapeyron slope of the dehydration reaction is positive
(where dehydration results in a volume increase, promoting
microcracking) and the thermal stability of serpentine
is relatively modest (inhibiting plastic flow processes).
Second, the experiments did not show evidence for stick–
slip behavior (or velocity weakening), which is required to
nucleate earthquakes. Recent studies have extended these
results to higher pressure. Acoustic emissions during deformation experiments at P = 1–6 GPa and T = 550–820 °C
E LEMENTS
Influence of sliding velocity on the frictional behavior
of antigorite (after Kohli et al. 2011). At tectonic slip
rates, the behavior is velocity strengthening, resulting in creep.
However at seismic slip rates, samples exhibit dramatic dynamic
weakening, which could facilitate earthquake rupture through aseismically creeping patches.
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FIGURE 3
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The denudation of serpentinized mantle along the ocean–
continent transition (OCT) and by seafloor spreading elicits
the following questions: when and how is lithosphere
serpentinized and what role does serpentinization play
during rifting processes and ongoing slow-spreading-ridge
activity?
Along the western Iberia margin, the OCT forms a 100
km wide basin dominated by completely (95–100%) to
moderately (50%) serpentinized peridotite (e.g. Afilhado et
al. 2008). Gravimetric and seismic profiles indicate that the
serpentinized mantle is about 5 km thick. Locally, tectonic
breccia and semibrittle gouge composed of gabbro, amphibolite, and serpentinite clasts in a matrix of calcite- and
chlorite-rich cataclasite are recovered during drilling. The
geochemical signatures of primary minerals indicate that
the serpentinized mantle is of depleted, subcontinental
type. The oxygen isotope signature and the absence of
antigorite in these rocks suggest temperatures less than
300 °C during fluid infi ltration (Skelton and Valley 2000).
Later infi ltration of cooler seawater (<100 °C) along brittle
normal faults accompanied the seafloor exhumation of
the mantle. Mohn et al. (2012) proposed the following
scenario of mantle exhumation: Crustal thinning is fi rst
accommodated by pure shear along conjugate, crustal-scale
shear zones and by ductile flow of the middle crust (FIG. 4A).
The upper mantle is progressively thinned and replaced
by deeper, hotter mantle impregnated by mafic magmas
(FIG. 4B). Faults cut from the surface into the mantle, and
seawater infi ltrates down to the upper mantle, initiating
moderate-temperature serpentinization
in the lizardite stability field (FIG. 4 C ).
A
Serpentinization weakens the upper
mantle layer, promoting strain localization and allowing the normal faults in
the distal margin to root at low angle. The
ongoing extension leads to buoyant flow
of partially serpentinized mantle along
low-angle detachment faults, resulting in
the horizontal stretching of the subcontinental mantle at the seafloor (FIG. 4D).
B
At slow- to ultraslow-spreading ridges
complexes (OCCs) (FIG. 5). The domes exhibit corrugated
surfaces that are interpreted to be caused by exposed
detachment faults. OCCs are usually younger than 2
Ma and, when active, show hydrothermal activity and
moderate seismicity down to 3–4 km depth along the main
detachment fault. OCCs cover areas of ~200 to 500 km 2,
extending ~25 km along the ridge axis and ~15 km perpendicular to the ridge axis. The corrugated surfaces include
serpentinized peridotite and altered gabbro; some basalt
is observed close to the axial ridge (FIG. 5). Seismic studies
show relatively high velocities (Vp > 7 km/s) within a 2 to
3 km thick zone parallel to the surface, suggesting that
this zone is dominantly composed of a mixture of gabbro
and partially serpentinized peridotite (Karson et al. 2006).
Deformation is localized in a ~100 to 200 m thick shear
zone at the footwall of the detachment zone, and the rest
of the massif is mostly undeformed (Karson et al. 2006). Far
from the shear zone, serpentinites are massive and sparsely
crosscut by joints. Partially serpentinized peridotites retain
porphyroclastic textures acquired during high-temperature mantle deformation. Mesh-textured serpentinites
show variable densities of crosscutting serpentine veins.
Within the detachment faults, the mineralogy is more
complex; the presence of talc, amphibole, and chlorite
indicates the addition of silica, calcium, and aluminum
from hydrothermal fluids derived through seawater–
gabbro interactions. Serpentinization along the detachment faults involves at least several stages of alteration
and veining (e.g. Andreani et al. 2007). Early serpentine
(<40 mm/y), 20 to 25% of the seafloor is
composed of serpentinite (Cannat et al.
2009). These areas of outcropping serpentinite are associated with domes, up to
500 meters high, known as oceanic core
C
FIGURE 4
Dynamic evolution of a rifted continental
margin up to the formation of the ocean–
continent transition (modified after Mohn
et al. 2012). UCC = upper continental crust;
MCC = middle continental crust; LCC = lower
continental crust; OCT = ocean–continent
transition; LT = low-temperature; HT = hightemperature; dark green = lower oceanic
crust; blue to green = partially serpentinized
upper mantle; yellow = marine sediments
D
E LEMENTS
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mation in a closed chemical system. The
last generation of veins show a granular
texture formed by micron-scale lamellae
of Al-rich lizardite. These compositions
and microtextures require a change in
ambient conditions, with significant water
circulation at shallow depths (<2 km) and
temperatures of <200 °C.
Strike-Slip Faults
(B OTTOM) Schematic 3-D block diagram of an oceanic
core complex (OCC) (after Cannat et al. 2009).
(UPPER LEFT) Corrugated surface of an OCC (area shown is 10 × 5 km).
(U PPER RIGHT) Photomicrograph (crossed polarizers) of a weakly
serpentinized sample of troctolite from IODP Hole U1309D
(Bar scale: 200 µm). PHOTO IODP, EXPEDITION 305
FIGURE 5
veins are irregular and thin (<0.5 mm) and exploit preexisting cracks in olivine. The local presence of antigorite
and Fe-rich brucite suggests that early serpentinization
occurred at about 300–400 °C at a low water/rock ratio.
Second-generation veins, composed dominantly of lizardite
or chrysotile and magnetite, are usually interconnected and
a few centimeters in length, and they form the primary
network of serpentine veins. These veins are often oriented
parallel to the detachment fault, suggesting that they
opened during tectonic unroofi ng. The third generation
of veins contain oblique synkinematic fibers that record
a shear component; the fibers are typically chrysotile and
polygonal serpentine and document incremental defor-
A
Large-scale strike-slip faults can either
slip by aseismic creep or move abruptly
during earthquakes. Weak fault behavior
can be caused by high fluid pressures,
locally high geothermal gradients, or
the presence of weak materials such as
serpentinite. The San Andreas Fault (SAF),
one of the most active strike-slip faults
on Earth, extends at least 15 km into the
crust and is marked at the surface by a
complex network of crushed and fractured
rocks a few hundred meters to several
kilometers wide (FIG. 6). Just north of the
Parkfield segment, the SAF creeps at a rate of 28 mm/y.
Serpentinized ultramafic rocks derived from the Franciscan
Complex mélange have been associated with fault creep,
as well as low fault strength, in this region (e.g. Moore
and Rymer 2007). The SAFOD project (San Andreas Fault
Observatory at Depth) sampled and instrumented the SAF
9 km northwest of Parkfield. At about 3 km depth, the drill
hole intersects the Great Valley Formation, a sedimentary
unit rich in serpentinites associated with the Coast Range
ophiolite (e.g. Moore and Rymer 2007). Microstructural
observations from core samples show evidence of deformation across the damaged zone. The deforming zones
contain serpentinite clasts and highly sheared siltstones.
Moore and Rymer (2007) also reported talc in the cuttings.
This discovery was noteworthy, as the frictional strength
of talc at elevated temperatures is low enough to meet the
constraints on the shear strength of the fault, and its stable
sliding behavior is consistent with fault creep.
(A) Distribution
of serpentinite
outcrops in central California
and localization of major faults
with creeping (aseismic)
segments. SF, San Francisco;
SB, Santa Barbara; SJ, San Jose;
M, Monterey; P, Parkfield; C,
Cholame; O, SAFOD; ID, Idria;
SAF, San Andreas Fault; HF,
Hayward Fault; CF, Calaveras
Fault; SYF, Santa Ynez Fault;
LPF, Little Pine Fault.
(B) Geology of the SAFOD
Borehole with location of
cuttings (yellow circle)
containing serpentinite-derived
clay minerals. Red lines indicate
active fault strands.
(C) Scanning electron microscope image of gouge
showing microshear zones
in serpentinite-derived clay
minerals (SAFOD Photography
Atlas). Width of field of view =
2 mm
B
FIGURE 6
C
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Recent investigations on the main 200 m thick damage
zone at ca. 3 km depth at SAFOD show that the creeping
area consists of clasts of serpentinite and sedimentary rock
dispersed in a matrix of saponite (Mg-rich smectite clay),
which is one of the weakest phyllosilicates. Saponite is a
low-temperature (<150 °C) product of metasomatic reactions
between the quartzofeldspathic wall rocks and serpentinite blocks in the fault (e.g. Lockner et al. 2011). Thus,
serpentinite-derived clay minerals alone appear sufficient
to explain the low fault strength and creeping behavior of
the San Andreas Fault. However, considering the temperature stability field of saponite and the geothermal gradient
within the drill hole, it is unlikely that this very weak
mineral is present deeper than 4 km along the fault zone.
Thus, stable creep and low strength in the SAF down to
a minimum depth of 10 km must reflect the presence of
other low-strength minerals (such as talc), elevated fluid
pressure, or other deformation mechanisms (e.g. pressure
solution) in serpentinites or quartzofeldspathic rocks.
Subduction Zones
Decoupling between the mantle wedge and the subducting
slab is inferred from modeling of heat flow data down to
75 ± 15 km (Wada et al. 2008). The decoupled zone corresponds to the stability of phyllosilicates such as serpentine,
whose weakness and/or ductility can explain a rheological
weakening of the interface (Hilairet et al. 2007; Wada et
al. 2008). Hydrous silicates (e.g. serpentine, amphibole,
talc, chlorite, lawsonite) destabilize at increasing pressure
and temperature in subduction zones and release water
at different depths (Hacker et al. 2003). With ~13 wt%
water in its structure, serpentine is an efficient carrier and
source of water at depth during subduction (see Evans et al.
2013 this issue). Dehydration of serpentine has also been
hypothesized to play an important role in the origin of
double seismic zones, observed in some subduction zones,
and in the origin of intermediate and deep seismicity
along the slab (e.g. Yamasaki and Seno 2003). However, as
noted above, the exact mechanism by which dehydration
promotes seismicity remains a matter of debate. In addition,
high-resolution seismic tomography indicates that seismic
velocities in the lower seismicity plane are better explained
by seismic anisotropy of anhydrous peridotite than by the
presence of serpentinite (Reynard et al. 2010).
Indirect petrological evidence for serpentinization at
subduction zones is provided by the intimate association of
serpentinites with eclogites (high- and ultrahigh-pressure
rocks). Guillot et al. (2009) estimated that almost 30%
of Phanerozoic eclogitic massifs worldwide are associated
with serpentinites. Evidence for prograde metamorphism
of serpentinite includes the transformation of lizardite to
antigorite (FIG. 7) and the presence of metamorphic olivine
and enstatite associated with Ti clinohumite, which reflects
partial dehydration above 450 °C. The rheological properties of serpentine may help explain the mechanical evolution of subduction zones. Single-sided subduction appears
to require a weak hydrated slab interface and high slab
strength (Gerya et al. 2008). Tectonic settings that favor
serpentinization of strong mantle rocks near the surface
(such as the OCT and fracture zones) are thus good candidates for the location of subduction initiation. Serpentinites
can lubricate the nascent interplate fault zone, facilitating
asymmetric plate movement. Similarly, during exhumation
of high-pressure or ultrahigh-pressure rocks, the opposite
trajectories of exhumation and subduction require a decoupling zone within the subducting slab. A serpentinized
layer such as the OCT prior to subduction may become a
decoupling zone between the oceanic crust and underlying
lithosphere. Moreover, the buoyancy of serpentinite likely
contributes to eclogite exhumation because the average
density of eclogite and serpentinite mélange is lower than
that of anhydrous mantle peridotite.
CONCLUSION
Our understanding of the rheological properties of
serpentinite has advanced significantly during the last
ten years, and improved geophysical observations allow
detection of serpentinite bodies in active geological
A low-velocity zone above the high-velocity anomalies of
a cold subducting slab has been imaged by seismic tomography below the Cascadia forearc and is interpreted as
partially serpentinized mantle wedge (Bostock et al. 2002).
The hydrated subduction interface, called the serpentinite
channel, is more difficult to detect as it is probably thinner
than a few kilometers. Receiver functions show a slow layer
forming on top of the subducting plate below Japan at
depths between 80 and 140 km, a structure interpreted as
evidence for a serpentinite layer (Kawakatsu and Watada
2007). Bezacier et al. (2010) argued that serpentinite layers
with kilometer-size folds could result in an orientationindependent anisotropy, with bulk anisotropy close to the
mean value of 0.05–0.06 s km−1. In that case, the observed
S-wave delay times (up to 1 s or more) would require a
20 km thick serpentinite layer.
Serpentinite mud volcanoes containing clasts of serpentinized mantle peridotite in the Mariana forearc provide
direct evidence of serpentinites in subduction zones (e.g.
Fryer et al. 1999). The faults in the upper plate link mud
volcanism to the main decollement fault and provide a
route for buoyant serpentine-rich mud. Both the unconsolidated serpentine and the serpentinized peridotite clasts
are dominated by chrysotile ± lizardite and antigorite that
formed at 350–450 °C and 600–700 MPa.
E LEMENTS
Schematic diagram illustrating the geological context
during subduction-related serpentinization (COURTESY
OF F. D ESCHAMPS) and associated photomicrographs (COURTESY OF S.
SCHWARTZ). With increasing depth, low-grade chryzotile-lizardite (Lz)
transforms into high-grade antigorite (Atg). Below 50 km depth,
antigorite breakdown releases fluids and metamorphic olivine
(Ol) crystallizes.
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FIGURE 7
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settings. In spite of emerging consistency among laboratory,
geophysical, and geological observations, uncertainties
remain concerning the rheological laws used to fit the
data and the extrapolation of the laws to natural strain
rates. Future work will focus on comparing experimental
deformation mechanisms with natural examples and on
investigating the processes that control the brittle–ductile
transition and strain localization. These properties are
key for understanding the nucleation and propagation of
earthquakes and the significance of localized deformation
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ACKNOWLEDGMENTS
We acknowledge funding from the National Science
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