article in press - woods hole oceanographic institutionhirth_2006_15125.pdf · b woods hole...

etters xx (2006) xxx–xxx

+ MODEL

www.elsevier.com/locate/epsl

ARTICLE IN PRESS

Earth and Planetary Science L

Grain size sensitive deformation mechanisms in naturallydeformed peridotites

Jessica M. Warren a,⁎, Greg Hirth b

a MIT/WHOI Joint Program, Woods Hole, MA 02540, USAb Woods Hole Oceanographic Institution, Woods Hole, MA 02540, USA

Received 6 April 2005; received in revised form 5 June 2006; accepted 6 June 2006

Editor: R.D. van der Hilst

Abstract

Microstructural analyses of peridotite mylonites from the oceanic lithosphere indicate that shear localization results from thecombined effects of grain size reduction, grain boundary sliding and second phase pinning during deformation. The pinning effect,combined with experimental flow laws for olivine, suggests that a permanent transition from dislocation creep processes todiffusion creep occurs. This rheological transition provides a mechanism for long term weakening of the lithosphere for rocksdeforming in the brittle–ductile regime. In addition, our results support the hypothesis that a transition to diffusion creep promotesthe randomization of pre-existing lattice preferred orientations, which would reduce seismic anisotropy.© 2006 Elsevier B.V. All rights reserved.

Keywords: mylonites; olivine; grain boundary sliding; lattice preferred orientation; anisotropy

1. Introduction

The observation of fine grained mylonites in shearzones indicates that grain size sensitive creep processespromote strain localization (e.g., [1–7]). For the samplein Fig. 1, microstructural observations show that ductiledeformation resulted in three orders of magnitude grainsize reduction. Peridotite mylonites are commonlysampled from oceanic fracture zones [4]. The geologicprovenance of these mylonites suggests that they formedin the ductile region of the lithosphere along the down-dip extension of transform faults. We describe fabrics ina peridotite mylonite from the Shaka Fracture Zone that

⁎ Corresponding author. Tel.: +1 508 289 3749.E-mail addresses: [email protected] (J.M. Warren),

[email protected] (G. Hirth).

0012-821X/$ - see front matter © 2006 Elsevier B.V. All rights reserved.doi:10.1016/j.epsl.2006.06.006

support extrapolation of experimentally determinedrheological data [8] to geologic conditions appropriatefor the oceanic lithosphere. We conclude that grainboundary sliding and grain boundary pinning lead topermanent grain size reduction and strain localization.

Previous studies of peridotite mylonites in orogenicmassifs and ophiolites have interpreted grain sizereduction in mylonites as the result of (i) fluid addition(e.g., [9]), (ii) reaction with melt (e.g., [10]), and (iii) thebreakdown reaction of spinel to plagioclase withdecreasing pressure (e.g., [11,6]); see [3] for a reviewof earlier literature. Vissers et al. [9] interpreted themicrostructures in Erro Tobbio lherzolite mylonites asresulting from a wet olivine rheology with reaction-related grain size reduction. Dijkstra et al. [10]concluded that a melt-present reaction promoted grainsize sensitive deformation in shear zone mylonites from

EPSL-08233; No of Pages 13

mailto:jmwarren@whoi.�edu

mailto:[email protected]

http://dx.doi.org/10.1016/j.epsl.2006.06.006

Fig. 1. (A) Photomicrograph of ultra-mylonite peridotite sample AII-107-61-83 from the Shaka Fracture Zone on the Southwest Indian Ridge. Grainsize sensitive deformation has resulted in the formation of anastomosing bands of coarser (<100 μm) and finer (<10 μm) grained olivine, withdistributed grains of orthopyroxene, clinopyroxene and spinel. Large porphyroclasts of olivine, orthopyroxene, clinopyroxene, spinel and coarsergrained olivine aggregates remain and appear to have behaved as hard inclusions during deformation. (B) Photomicrograph of a coarser grained areaof the ultra-mylonite, with a grain size range of 1–100 μm. (C) Photomicrograph of a finer grained area, with an olivine grain size of 1–10 μm.

2 J.M. Warren, G. Hirth / Earth and Planetary Science Letters xx (2006) xxx–xxx

ARTICLE IN PRESS

the Othris peridotite massif. In the Turon de Técouéreperidotite shear zone, Newman et al. [6] suggested thatgrain size reduction resulted from the spinel to plagio-clase phase transition reaction.

Several lines of evidence indicate that the ShakaFracture Zone mylonites deformed in the absence ofwater and melt, and that breakdown reactions were notinvolved. We find no syn-deformational hydrous phasessuch as amphibole, talc or serpentine. The mantle atfracture zones is expected to be dehydrated and depleteddue to earlier on-axis melting and melt extraction [12].In addition to the lack of evidence for melt or waterinvolvement in the mylonite deformation, we find noindication of reaction enhancement via the breakdownof spinel to plagioclase. The rare earth elementconcentrations measured by ion microprobe in aclinopyroxene porphyroclast have the typical depletedcompositions found in other abyssal peridotites [13],with no Eu anomaly, which is a sensitive indicator ofplagioclase formation [14].

2. Methods

To study deformation mechanisms in mylonites, weused electron backscatter diffraction (EBSD) to analyzeolivine and orthopyroxene crystal orientations in thinsections of mylonitized peridotites cut parallel tolineation and perpendicular to foliation. We madeorientation measurements down to a 1 μm grain size,

an order of magnitude smaller size than that accessiblewith a U-stage. In addition, EBSD permits identificationand quantification of the distribution of secondaryphases. Data were gathered using an electron back-scatter detector attached to a JEOL 840 SEM. Portionsof thin sections were mapped for lattice orientation in 1,2, or 4 μm steps (depending on the average grain size ofthe region being mapped), using a rasterized beamacross the sample surface. EBSD patterns wereprocessed and analyzed using HKL Technology'sChannel 5 software package.

Orientation maps in Figs. 2–4 contain from 52% to72% indexed data. Non-indexed (white) pixels representpoints with a mean angular deviation (MAD) number≥1°, which result from surface roughness and computermis-indexing. The MAD number quantifies the mis-match between lattice planes in the calculated orientationand those quantified from the digitized bands of thediffraction pattern. Raw data, consisting of all indexedpoints with a MAD number <1°, are shown as maps andpole figures for coarser and finer grained regions in Fig.2. Processed data, for which wild spikes have beenremoved and all pixels with zero solutions have beenextrapolated, are also shown in Fig. 2. Wild spikes aresingle pixels (i) which are misoriented by >10° from theaverage orientation of the surrounding eight pixels and(ii) for which the maximum misorientation between anytwo of the surrounding pixels is <10°. All pixels whichare wild spikes or have a MAD number >1° are replaced

Fig. 2.

3J.M. Warren, G. Hirth / Earth and Planetary Science Letters xx (2006) xxx–xxx

ARTICLE IN PRESS

Fig. 2. Maps of raw data collected in a coarser grained area (A) and afiner grained area (E) of the ultra-mylonite. The coarser grained areawas collected at a 2 μm step size with a total of 64% of the areaindexed; the finer grained area was collected at a 1 μm step size with52% indexed. Each grey point represents an olivine orientationcollected by EBSD and green points represent secondary phases(orthopyroxene, clinopyroxene, and spinel) for which orientationswere measured. White points have a MAD number ≥1°. Black linesindicate grain boundaries, defined by a 10° misorientation betweenadjacent points. Maps of the processed data, for which wild spikeshave been removed and data extrapolated, are shown in (B) and (F).The map in (I) of the fine-grained area shows an alternative method ofprocessing the data, in which the data is extrapolated prior to removingwild spikes. Pole figures of one orientation point per olivine grain areshown in (C), (D), (G), (H), and (J). All pole figures are equal arealower hemisphere projections, contoured using a 15° half width and asmultiples of a uniform distribution (MUD) to a maximum of 3 MUD.Contouring of 1 MUD indicates no significant difference from auniform distribution and corresponds to the absence of an LPO. Theshear direction is parallel to the horizontal axis in the pole figures.


ARTICLE IN PRESS

with zero, or null, solutions. In the processed data, pixelswith zero solutions were replaced with the most commonneighbor orientation. This method of processing the dataminimizes the overcounting of individual grains whenextracting one orientation point per grain for theconstruction of pole figures. However, it under-repre-sents the smallest grains, which are close to the samplingstep size and are filtered out with noise in the data.

Fig. 3. Orientation maps collected by EBSD for two coarser grained areas of(D) was collected at a 4 μm step size, with 72% of the area indexed. Olivindegree of misorientation of each pixel from the shear plane. Dark green correswhite areas have no data. Pole figures of one orientation per olivine for the arein (C). For the area in (D), data are only plotted for 10–100 μm grains in (E), dmaximum of 6 MUD. The shear direction is parallel to the horizontal axis in

Pole figures, shown in Figs. 2–4, were calculatedusing one point per grain, with all data sets containing≥180 grains, which has been demonstrated to providestatistically robust results [15]. Grain boundaries weredefined by misorientations ≥10° between adjacentpoints and subgrains by 2°–10° misorientations. Polefigures of raw and processed data in Fig. 2 demonstratethat in both the finer and coarser grained areas, theprocessing affects the strength of the lattice preferredorientation (LPO), but does not alter the overall LPOpattern. There are two main differences between the rawand processed data: (1) the removal of small grains,points with no solutions and wild spikes, whichdecreases the randomness of the data, leading to astronger LPO and (2) extrapolation of the data, whichfurther reduces the number of small grains but increasesthe randomness of the data, as groups of adjacent smallgrains with the same orientation are replaced byindividual large grains.

Using an alternative processing method, shown inFig. 2I and J, data are extrapolated prior to the removalof wild spikes. This provides a better representation ofthe distribution of the pyroxene and spinel grains in thefinest grained areas of the sample. However, as we wishto minimize potential bias in the olivine fabrics, weshow data processed by the first method only in Figs. 3and 4. These issues are inherent in analyzing such fine-grained material, which has a grain size close to theminimum step size.

3. Results

Microstructural observations of the mylonites indi-cate that grain size reduction leads to progressivelocalization of deformation into fine-grained bands.The influence of grain size on rheology can be seen in theoptical-scale microstructures in Fig. 1. The myloniteconsists of anastomosing bands of coarser (≤100 μm)and finer (≤10 μm) grained olivine, with distributedgrains of orthopyroxene, clinopyroxene and spinel.Large, round to elongate porphyroclasts of olivine,orthopyroxene, clinopyroxene and spinel, ranging in sizefrom 0.1 mm to 5 mm, are preserved in the fine-grainedmatrix. These microstructures indicate that the largeporphyroclast grains behaved as hard inclusions duringdeformation. In some cases, these inclusions are

the mylonite (A, D). The area in (A) is also shown in Fig. 2; the area ine is shaded from blue to yellow, with the color range representing theponds to orthopyroxene, light green to clinopyroxene, red to spinel, anda in (A) are shown for 1–10 μm grains in (B) and for 10–100 μm grainsue to the large step size of the map. Contouring in all pole figures is to athe pole figures.


ARTICLE IN PRESS

Fig. 4. Orientation maps for three finer grained areas of the mylonite (A, C, E), using the same color scheme as in Fig. 3. The area in (A) is also shownin Fig. 2; the area in (C) was collected at a 1 μm step size, with 57% indexed; the area in (E) was also collected at a 1 μm step size, with 58% of thearea indexed. Pole figures of one orientation per olivine are shown (B, D, F) for the 3 areas, all of which have grain sizes of 1–10 μm. The contourscale is the same as in Fig. 3, to a maximum level of 6 MUD. In the first area, the actual maximum is 2 MUD; in the second and third areas it is 3MUD.


ARTICLE IN PRESS


ARTICLE IN PRESS

relatively coarse-grained aggregates of olivine (Fig. 1).Kinematic indicators, such as vergence of small folds andasymmetric recrystallization trails around porphyroclasts,indicate the relative sense of shear for the mylonite.

Olivine orientations in coarser and finer grainedregions of the mylonite are shown in maps and polefigures in Figs. 3 and 4. The data are plotted with respectto the banding and lineation orientations, which weassume are equivalent to the shear plane for this highstrain rock. Data are presented for two coarser grainedregions of the mylonite in Fig. 3, with the pole figuresdivided into 1–10 μm and 10–100 μm size grains foreach data set. In the first region (Fig. 3A–C), a [100]maximum is observed in the plane of the banding andsubparallel to the lineation, while a [001] maximum isobserved perpendicular to the banding. In some areas,the LPO is significantly stronger among 10–100 μmgrains than among 1–10 μm grains. In the other coarsegrained region, a pole figure is only shown for the 10–100 μm size grains (Fig. 3E), due to the large step size atwhich this data set was collected. The [100] maximum isnot oriented parallel to the banding. Avariable LPO wasalso found by limited universal stage work on coarsergrained regions in this sample [4].

In the finer grained regions, shown in Fig. 4, adecrease in LPO strength is observed in all three areasmeasured. Almost no LPO is observed in some regionsof the fine-grained bands (Fig. 4A–F). Locally, in thefine-grained region shown in Fig. 4E and F, we find aweak and dispersed “ghost” of the LPO observed in thecoarser-grained region in Fig. 3A–C.

The variation of LPO and grain size correlates withthe secondary phase distribution. Pyroxenes and spinelare inhomogeneously distributed in the mylonite, withsmall grains of these phases predominant in regions witha smaller olivine grain size. Based on the very fine grainsize, the raw data (Fig. 2A and E), and the alternativemethod of processing the data (Fig. 2I), the number ofsmall orthopyroxene grains is greater than that shown inFigs. 3 and 4. The EBSD orientation maps highlight thelarger number of secondary phases in finer grained areasrelative to coarser grained areas (compare Fig. 2B to F).The finer grained areas have a smaller range of grainsizes and a more random distribution of secondaryphases. In contrast, the coarser grained regions have alarger range of olivine grain sizes. Locally, small grainsof secondary phases are mixed with small grains ofolivine at the edges of larger olivine grains.

Fig. 2F and I provide lower and upper bounds on theamount of secondary phases pinning the grain bound-aries in the finest grained areas. We have also examinedback-scattered electron images and X-ray maps of

coarse- and fine-grained areas. These confirm ourobservation of a greater concentration of secondaryphases in the finest grained areas. Similarly, highresolution EBSD and X-ray maps of Michibayashi andMainprice [16,17] show a greater accumulation ofsecondary phases in finer grained regions of theperidotite mylonite in their study.

4. Discussion

4.1. Fabric variations with grain size

The variation in olivine fabric with grain size in themylonite provides an opportunity to constrain grain sizesensitive deformation mechanisms in the lithosphere.Three observations in particular can be exploited toconstrain the deformation mechanisms: (i) randomiza-tion of a pre-existing LPO in the finest grained regionsof the mylonite; (ii) observation of an LPO in coarsergrained regions; (iii) the preservation of porphyroclasts.

The randomization of olivine LPO in the finestgrained regions of the mylonite indicates a change indeformation mechanism to diffusion creep. Two obser-vations suggest that the randomization of LPO occurredas deformation localized into fine-grained regions. First,a weak LPO is observed in one fine-grained area (Fig.4F) which is similar to the LPO in a coarser grained area(Fig. 3B and C). Second, the LPO among 1–10 μmgrains in a coarser grained area is similar, but weaker,than that in the 10–100 μm size fraction (Fig. 3B and C).We conclude that the randomization of LPO is related tograin rotations that occur during the grain boundaryswitching process, which is required during diffusioncreep (e.g., [18]).

The LPO in the coarser grained regions indicates that adislocation process accommodated deformation duringthe early stages of mylonitization. At the same time, thepreservation of large, sometimes equant, porphyroclastsimplies that the rheology is grain size sensitive. If therheology was not grain size sensitive, then all grainsshould be deformed to the same degree and large,relatively undeformed porphyroclasts would not bepreserved. Based on comparison to experimental studies,we suggest that deformation in the coarser grained bandsoccurred by dislocation accommodated grain boundarysliding (DisGBS). DisGBS is a grain size sensitive creepmechanism thatwas identified experimentally on the basisof rheologic data on olivine aggregates deformed nearthe transition between dislocation creep and diffusioncreep [19,8]. Of particular significance for our analysisis that deformation byDisGBS has been shown to produceLPOs in calcite [20,21] and ice [22,23].

Fig. 5. (A) An olivine deformation mechanism map, on axes ofdifferential stress versus grain size, contoured for strain rate, plottedusing the flow laws for dry, melt free olivine at 700 °C and 12 kmdepth [25,26,8]. The black recrystallization piezometer field delimitsthe grain size predicted by the empirical piezometric relationship forolivine deformed under dry conditions [31,32]. The thick whiteboundaries between deformation fields indicate field boundaries in theabsence of DisGBS. The thin white boundaries delineate the DisGBSfield and the dashed boundaries on either side represent uncertainty inthe location of the DisGBS boundaries due to uncertainty in itsactivation energy. The change in shading within the DisGBS fieldindicates the transition from the regime where creep on the easy slipsystem limits the strain rate to the regime where GBS limits the strainrate. (B) The variation of total and individual strain rates with grainsize, at a differential stress of 300MPa, and at 700 °C and 12 km depth.The dashed lines represent the individual components of DisGBS.


ARTICLE IN PRESS

During DisGBS, strain is dominantly accommodatedby the relative movement of grains (i.e. grain boundarysliding) and is limited by either this process or themovement of dislocations within grains on the easy slipsystem [24,8]. These two processes act in series, asdeformation is not accommodated independently byeither grain boundary sliding or easy slip. Thus thestrain rate for DisGBS (ϵ̇DisGBS) is limited by the slowestof the two mechanisms through the relationship:

�ϵDisGBS ¼�

1�ϵGBS þ1�ϵEasy

��1

ð1Þ

where ϵ̇GBS is the strain rate for grain boundary sliding(GBS) and ϵ̇Easy is the strain rate for easy slip. The grainsize sensitivity of DisGBS results from the influence ofgrain size on the GBS component of the flow law.

There is considerable confusion regarding theterminology used for grain size dependent deformationmechanisms. For example, deformation in both thediffusion creep and DisGBS regimes involves grainboundary sliding, with relative grain movement accom-modated either by diffusion or dislocation processes,respectively. Furthermore, the term superplasticity isoften used when referring to deformation accommo-dated by GBS (see discussion in Goldsby and Kohlstedt[24]). To clarify our terminology, we refer to the serialdeformation process (Eq. (1)) as DisGBS and the grainboundary sliding component of the process as GBS.

4.2. Deformation mechanism maps

We use olivine deformation mechanism maps toconstrain the deformation conditions of the mylonites(Fig. 5A). The total strain rate (ϵ̇Total) during deforma-tion is contoured in stress versus grain size space, usingthe constitutive law [8]:

�ϵTotal ¼ �ϵDif þ �ϵDis þ �ϵLTP þ �ϵDisGBS ð2Þwhere ϵ̇Dif is the strain rate for diffusion creep, ϵ̇Dis fordislocation creep, and ϵ̇LTP for low-temperature plasti-city (LTP). The flow law formulation and experimentaldata for LTP are from [25,26]. These four deformationmechanisms are independent, so the mechanism withthe fastest strain rate controls the rheology.

The boundaries for the fields of diffusion creep,dislocation creep, and DisGBS are calculated using acompilation of olivine experimental data from Hirth andKohlstedt [8]. The flow laws for melt- and water-freeaggregates in these regimes have the form:

�ϵ ¼ Arnd�mexp �E þ PVRT

� �ð3Þ

where A is a constant, σ is the differential stress, n is thestress exponent, d is the grain size, m is the grain sizeexponent, E is the activation energy, P is the pressure, V


ARTICLE IN PRESS

is the activation volume, R is the gas constant, and T is theabsolute temperature. For the mylonite, d is in the range of1–100 μm. The deformation temperature is constrainedfrom thermometry to lie in the range 600–750 °C [4]. Weuse a nominal pressure of 400 MPa for these calculations.The pressure of deformation is not well constrained,however the PV term in Eq. (3) is negligible fordeformation in the shallow oceanic lithosphere.

Application of the experimental olivine data to theconditions of mylonite formation involves a significantextrapolation of temperature, but not of grain size orstress. Therefore, the greatest uncertainty in the map inFig. 5A is in the activation energies for the deformationmechanisms. We calculated the error range for the fieldboundaries due to error in these activation energies. Theactivation energies for DisGBS are the most uncertain.For the grain boundary sliding component, the activa-tion energy is assumed to be the same as that for the easyslip system, which is constrained experimentally at hightemperature [27]. However, results from Carter andAve'Lallemant [28] on the experimental deformation ofdunites indicate that the easy slip system in olivinechanges at low temperature. Despite uncertainty in theextrapolation to low temperature, the range of estimatedmylonite deformation conditions lies within the DisGBSfield, as shown in Fig. 5A.

To emphasize the conditions where DisGBS is asignificant deformation mechanism, and its significancefor LPO development, we plot strain rate versus grainsize for a stress of 300 MPa in Fig. 5B. This plotrepresents a constant stress section through Fig. 5A. Thetotal strain rate is dominated by DisGBS at grain sizes inthe range ∼15–700 μm. Furthermore, the strain rate forDisGBS is at least an order of magnitude greater than thatfor dislocation creep at grain sizes <∼300 μm. The strainrate for diffusion creep becomes greater than that forDisGBS at grain sizes <∼15 μm. For comparison, at agrain size of∼15 μm, the strain rate for diffusion creep isapproximately three orders of magnitude greater thanthat for dislocation creep. Assuming that LPOs formwhen a dislocation process accommodates a significantcomponent of the total strain–and recalling that ourmicrostructural observations indicate a randomization ofLPO when grain size decreases below ∼10 μm–weconclude that the LPOs observed in coarser grainedregions of the mylonite resulted from DisGBS.

Our results suggest that an LPO is maintained duringDisGBS when the easy slip component is dominantduring deformation, though rate limited by the GBScomponent. Grain boundary sliding, whether diffusionor dislocation accommodated, tends to randomize LPOs.Prior and co-workers have demonstrated that a grain

boundary sliding process leads to an increased dispersionin the misorientation between grains, which weakens aninitial single crystal LPO [7,29,30]. Our observation of arandomized LPO during diffusion creep supports thesuggestion by Bestmann and Prior [29] that a grainboundary sliding process weakens LPOs.

4.3. Recrystallization and grain boundary pinning

To understand how grain size sensitive creeppromotes strain localization, we consider the processesthat control the evolution of grain size during mylonitedeformation. As deformation continues to lower tem-perature, and therefore higher stress, grain size decreasesas a result of dynamic recrystallization. Empiricalrelationships between stress and grain size have beendetermined for olivine and this recrystallization piezo-meter [31,32] is shown in Fig. 5A. Dynamic recrystalli-zation is driven by gradients in dislocation density,which promote subgrain rotation and grain boundarymigration. In the diffusion creep field, the applicabilityof the empirical piezometer is problematic owing to apotential lack of driving force for recrystallization. Thisinsight motivated the field boundary hypothesis [33,34],in which the piezometer is defined by the boundarybetween the dislocation creep and diffusion creep fields.Thus, the piezometric relationship may be temperaturedependent due to differences in the activation energy forthe creep processes. However, Drury [35] emphasizedthe lack of evidence for temperature dependence in theempirical relationships for olivine, based on theobservation of statistically equivalent piezometric rela-tionships for samples deformed at 1100–1300 °C [32]and at 1500 °C [31]. Thus, while the field boundaryhypothesis has merit, the higher temperature experi-mental data indicate that a recrystallization piezometerapplies within the dislocation creep regime.

The magnitude of grain size reduction in the mylonitecan be explained by either the empirical piezometricrelationship or a modified field boundary hypothesis. Asillustrated in Fig. 5A, the grain size at which we seeevidence for LPO randomization (i.e., ∼10 μm) fallsnear the boundary between diffusion creep and DisGBSat geologic strain rates, consistent with a modified fieldboundary hypothesis in which the piezometer is definedby the diffusion creep/DisGBS field boundary, ratherthan the diffusion creep/dislocation creep field boundary.At the same time, the empirical piezometric relationshipfalls in the DisGBS field (within uncertainty), indicatingthat the driving force for recrystallization arises from theeasy slip component of the DisGBSmechanism. In eithercase, we emphasize that the extreme grain size reduction

Fig. 6. The variation of effective viscosity with temperature, fordeformation within a shear zone compared to that outside a shear zone.Within the shear zone, viscosity is defined by deformation at theDisGBS–diffusion creep boundary at a total strain rate of 10−12 s−1.Deformation outside the shear zone is by dislocation creep, calculatedat the same stress as within the shear zone.


ARTICLE IN PRESS

observed in the mylonites is difficult to explain withoutthe DisGBS mechanism.

Grain boundary pinning can lead to permanentgrain size reduction and a transition to diffusion creep.Without a mechanism to limit grain growth, grain sizereduction may not promote a transition to diffusioncreep in single phase materials, due to a dynamicbalance between grain growth and recrystallization[33,34]. However, we suggest that the translation ofgrains during DisGBS by grain boundary sliding (andphase boundary sliding) results in the mixing ofdifferent mineral phases, which eventually inhibitsgrain growth. In the mylonite, the randomization ofLPO in the fine-grained regions indicates that thepresence of pyroxenes and spinel results in the pinningof olivine grain boundaries, leading to a transition todiffusion creep. Previous studies have emphasized therole of grain size pinning, resulting from fluidaddition, melt crystallization or metamorphic reac-tions, in promoting grain size sensitive creep processes[9,11,6,10].

4.4. Strain localization

To explore the implications of grain size sensitivecreep mechanisms for strain localization in the oceaniclithosphere, in Fig. 6 we compare the viscosities of afine-grained shear zone and the surrounding coarse-grained mantle. Based on our analysis of microstructuraldata and experimental flow laws, we assume thatdeformation in the shear zone occurs at a stress andgrain size on the boundary between the DisGBS anddiffusion creep fields on the deformation mechanismmap, which leads to the relationships:

�ϵDisGBS ¼ �ϵDif ð4Þ�ϵSZ ¼ �ϵDisGBS þ �ϵDif ¼ 2 �ϵDisGBS ¼ 2 �ϵDif ð5Þ

where ϵ̇SZ is the shear zone strain rate.The effective viscosity (η) of the shear zone in Fig. 6

is calculated using Eqs. (3) and (5), the relationshipη=σ/ϵ̇SZ and assuming ϵ̇SZ=10

−12 s−1. The calcula-tions are made for both the GBS and the easy slipcomponents of DisGBS, as shown in Fig. 6. Attemperatures greater than ∼700 °C, viscosity in theshear zone is controlled by the GBS component of theDisGBS flow law and this component may dominate,within error, at lower temperatures as well.

The shear zone viscosity at 700 °C is approximatelyfour orders of magnitude lower than outside the shearzone (Fig. 6). Outside the shear zone, we assume (i) thatdislocation creep accommodates deformation of the

coarse-grained mantle and (ii) that stress is limited bythe strength of the shear zone. The effective viscosity forthe shear zone remains lower than the coarse-grainedmantle until a temperature of ∼950 °C, at which pointdeformation is no longer localized. Where grain sizepinning results in a transition to diffusion creep, theviscosity contrast would be even larger than thatillustrated in Fig. 6. A convenient flow law fordeformation in mantle shear zones – which accountsfor grain size dependence – can be derived by solvingEq. (4) for grain size and substituting into Eq. (5):

�ϵSZ ¼ 2A3DisGBS

A2Dif

r8:5exp2HDif � 3HDisGBS

RT

� �ð6Þ

where H=E+PV. Using the values for these parametersfrom [8], Eq. (6) can then be written as:

�ϵSZ ¼ 2:4� 10�7r8:5exp�4:7� 105

RT

� �ð7Þ

for σ in MPa. This relationship is applicable for shearzone deformation at temperatures <950 °C and


ARTICLE IN PRESS

lithospheric depths, where the PV term is negligible. Wereiterate that the largest uncertainty is related to theactivation energy, as shown in Fig. 5A. At highertemperature, and therefore lower stress, errors asso-ciated with extrapolation in stress and grain size shouldalso be considered.

4.5. Relationship of mylonites to earthquake processes

A combination of geophysical, experimental andgeochemical observations indicate that the transitionfrom brittle to ductile processes occurs at a temperatureof ∼600 °C in the oceanic lithosphere. Seismologicalstudies of the depth of earthquakes along transformfaults indicate that seismicity is confined to depths in theoceanic lithosphere with temperatures <600 °C (e.g.,[36,37]). Extrapolation of experimental data for olivineaggregates indicates that the 600 °C isotherm also marksthe transition from stable to unstable frictional sliding[38]. Finally, Jaroslow et al. [4] estimated that theminimum temperature for mylonite deformation is∼600 °C, based on olivine–spinel geothermometry.

In Fig. 7, we compare the estimated strain rate formylonite deformation to the strain rates for differentgeologic processes. The stress and strain rates ofmylonite deformation are estimated using the micro-structural observations combined with the olivine flowlaws. From the recrystallization piezometer [31,32], theestimated differential stress for a recrystallized grain

Fig. 7. Duration versus estimated strain rate for the mylonite and forearthquakes, post-seismic slip and plate spreading. Earthquakes last 1–10 s, post-seismic slip lasts 1–100 days, and spreading events from1 yr to 1 My.

size of 10 μm is 240 MPa and of 3 μm is 580 MPa. At600–700 °C and these stress and grain size conditions,the estimated strain rate for mylonite deformation is10−14–10−10 s−1. The strain rate for post-seismic slip inFig. 7 is calculated as:

�ϵ ¼ slipwidth

� �� 1

Dt

� �ð8Þ

where Δt is the duration of post-seismic slip. Theestimated strain rate for post-seismic slip is 10−10–10−5 s−1, for slip<1 m, width∼1 m to <1 km, andΔt≈1–100 days. Therefore, the estimated strain rate formylonites overlaps the low end of post-seismic slip andthe high end of tectonic strain rates (i.e., 10−15–10−12 s−1). Our results place constraints on theconditions of deformation at the base of the seismo-genic zone in the oceanic lithosphere, which may beused to interpret geodetic data for post-seismic slip(e.g., [39]).

4.6. Mantle anisotropy

The microstructural observations of the mylonitedemonstrate that the transition to diffusion creep resultsin a randomization of pre-existing LPOs. Evidence thatpre-existing anisotropy may be destroyed by thetransition to diffusion creep is important for under-standing mantle anisotropy. In the upper mantle, seismicanisotropy reflects high temperature flow by dislocationcreep [40–43]. However, with increasing depth andtemperature in the mantle, the dominant mechanism ofolivine deformation may change from dislocation creepto diffusion creep [44]. Our results indicate that such achange can lead to a decrease in mantle anisotropy, ifenough strain is accumulated to randomize the fabric.We emphasize that in the case of shear zones, we do notexpect fabric randomization to have a detectable effecton regional mantle anisotropy. However, following thesuggestion of Karato and Wu [44] that a transition todiffusion creep may occur at depth in the upper mantle,our results imply that this would lead to an isotropiclayer in the lower portion of the upper mantle.

5. Conclusions

Peridotite mylonites demonstrate that three orders ofmagnitude grain size reduction occurs during localizeddeformation in the oceanic lithosphere. Microstructuralanalyses of the mylonite fabric suggest that deformationof aggregates with a grain size of 10–100 μm occurredby DisGBS and produces an LPO. Grain boundary


ARTICLE IN PRESS

pinning, due to the mixing of pyroxenes and spinelamong olivine grains during DisGBS, resulted inpermanent grain size reduction. Reduction in grainsize down to 1–10 μm, combined with second phasepinning, led to a transition to diffusion creep andresulted in a randomization of the pre-existing LPO.

Our microstructural observations are consistent withextrapolation of experimental olivine flow laws totemperatures of 600–800 °C during mylonite deforma-tion. The empirical recrystallization piezometer plotswithin the field of DisGBS, providing a driving force forgrain size reduction. Alternatively, grain size reductionmay have followed a modified version of the fieldboundary hypothesis [33,34]. At temperatures below∼950 °C, deformation in the DisGBS and diffusioncreep fields occurs at lower stress and viscosity than inthe dislocation creep field, providing a mechanism forweakening and strain localization in the oceanic litho-sphere at transform faults.

Acknowledgements

We thank Henry Dick for bringing our attention tothe peridotite mylonites. Louie Kerr provided invaluableassistance with the SEM at the Marine BiologicalLaboratory. In addition we would like to thank MargaretBoettcher, Mike Braun, Peter Kelemen, Jeff McGuire,and Laurent Montési for helpful discussions. Reviewsby David Prior and an anonymous reviewer helpedimprove our manuscript. This work was supported byNSF Grant EAR-0230267 and a Hollister Fellowshipfrom the WHOI Academic Programs Office.

References

[1] S.H. Kirby, Rock mechanics observations pertinent to therheology of the continental lithosphere and the localization ofstrain along shear zones, Tectonophysics 119 (1985) 1–27.

[2] J. Tullis, L. Dell'Angelo, R.A. Yund, Ductile shear zones frombrittle precursors in feldspathic rocks: the role of dynamicrecrystallization, in: A.G. Duba, W.B. Durham, J.W. Handin,H.F. Wang (Eds.), The brittle–ductile transition in rocks, Geo-physical Monograph, vol. 53, American Geophysical Union,1990, pp. 67–82.

[3] M.R. Drury, R.L.M. Vissers, D. Van der Wal, E.H. HoogerduijnStrating, Shear localisation in upper mantle peridotites, PureAppl. Geophys. 137 (4) (1991) 439–460.

[4] G.E. Jaroslow, G. Hirth, H.J.B. Dick, Abyssal peridotitemylonites: implications for grain-size sensitive flow and strainlocalization in the oceanic lithosphere, Tectonophysics 256(1996) 17–37.

[5] D. Jin, S.-I. Karato, M. Obata, Mechanisms of shear localizationin the continental lithosphere: inference from the deformationmicrostructures of peridotites from the Ivrea zone, northwesternItaly, J. Struct. Geol. 20 (2/3) (1998) 195–209.

[6] J. Newman, W.M. Lamb, M.R. Drury, R.L.M. Vissers,Deformation processes in a peridotite shear zone: reaction-softening by an H2O-deficient, continuous net transfer reaction,Tectonophysics 303 (1999) 193–222.

[7] Z. Jiang, D.J. Prior, J. Wheeler, Albite crystallographic preferredorientation and grain misorientation distribution in a low-grademylonite: implications for granular flow, J. Struct. Geol. 22(2000) 1663–1674.

[8] G. Hirth, D. Kohlstedt, Rheology of the upper mantle and themantle wedge: a view from the experimentalists, in: J. Eiler (Ed.),The Subduction Factory, Geophysical Monograph, vol. 138,American Geophysical Union, 2003, pp. 83–105.

[9] R.L.M. Vissers, M.R. Drury, E.H. Hoogerduijn Strating, C.J.Spiers, D. van der Wal, Mantle shear zones and their effect onlithosphere strength during continental breakup, Tectonophysics249 (1995) 155–171.

[10] A.H. Dijkstra,M.R. Drury, R.L.M.Vissers, J. Newman, On the roleof melt-rock reaction in mantle shear zone formation in the Othrisperidotite massif (Greece), J. Struct. Geol. 24 (2002) 1431–1450.

[11] M. Furusho, K. Kyuichi, Transformation-induced strain localiza-tion in a lherzolite mylonite from the Hidaka metamorphic belt ofcentral Hokkaido, Japan, Tectonophysics 313 (1999) 411–432.

[12] G. Hirth, D.L. Kohlstedt, Water in the oceanic upper mantle:implications for rheology, melt extraction and the evolution of thelithosphere, Earth Planet. Sci. Lett. 144 (1996) 93–108.

[13] K.T.M. Johnson, H.J.B. Dick, N. Shimizu, Melting in the oceanicupper mantle: an ion microprobe study of diopsides in abyssalperidotites, J. Geophys. Res. 95 (1990) 2661–2678.

[14] J.M. Warren, N. Shimizu, H.J.B. Dick, Melt impregnationrevealed by clinopyroxene geochemistry in abyssal peridotites,Geochim. Cosmochim. Acta 67 (2003) A526.

[15] W. Ben Ismaïl, D. Mainprice, An olivine fabric database: anoverview of upper mantle fabrics and seismic anisotropy,Tectonophysics 296 (1998) 145–157.

[16] K. Michibayashi, D. Mainprice, Understanding the deformationof rocks from the earth's mantle using simultaneous EBSD andEDS, www.hkltechnology.com/data/0-mantle.pdf, 2003.

[17] K. Michibayashi, D. Mainprice, The role of pre-existingmechanical anisotropy on shear zone development withinoceanic mantle lithosphere: an example from the Oman ophiolite,J. Petrol. 45 (2) (2004) 405–414.

[18] R. Raj, M.F. Ashby, On grain boundary sliding and diffusionalcreep, Metall. Trans. 2 (1971) 1113–1127.

[19] G. Hirth, D.L. Kohlstedt, Experimental constraints on thedynamics of the partially molten upper mantle: 2. Deformationin the dislocation creep regime, J. Geophys. Res. 100 (1995)15441–15449.

[20] S.M. Schmid, R. Panozzo, S. Bauer, Simple shear experiments oncalcite rocks: rheology and microfabric, J. Struct. Geol. 9 (5/6)(1987) 747–778.

[21] E.H. Rutter, M. Casey, L. Burlini, Preferred crystallographicorientation development during the plastic and superplastic flowof calcite rocks, J. Struct. Geol. 16 (10) (1994) 1431–1446.

[22] W.B. Durham, L.A. Stern, S.H. Kirby, Rheology of ice I at lowstress and elevated confining pressure, J. Geophys. Res. 106 (6)(2001) 11031–11042.

[23] D.L. Goldsby, D.L. Kohlstedt, Reply to comment by P. Duval andM. Montagnat on “Superplastic deformation of ice: experimentalobservations”, J. Geophys. Res. 107 (B11) (2002) 2313.

[24] D.L. Goldsby, D.L. Kohlstedt, Superplastic deformation of ice:experimental observations, J. Geophys. Res. 106 (B6) (2001)11017–11030.

http://www.hkltechnology.com/data/0antle.pdf


ARTICLE IN PRESS

[25] C. Goetze, The mechanisms of creep in olivine, Philos. Trans. R.Soc. Lond., A 288 (1978) 99–119.

[26] B. Evans, C. Goetze, The temperature variation of hardnessof olivine and its implication for polycrystalline yield stress,J. Geophys. Res. 84 (1979) 5505–5524.

[27] Q. Bai, S.J. Mackwell, D.L. Kohlstedt, High-temperature creepof olivine single crystals: 1. Mechanical results for bufferedsamples, J. Geophys. Res. 96 (B2) (1991) 2411–2463.

[28] N.L. Carter, H.G. Ave'Lallemant, High temperature flow of duniteand peridotite, Geol. Soc. Amer. Bull. 81 (1970) 2181–2202.

[29] M. Bestmann, D.J. Prior, Intragranular dynamic recrystallizationin naturally deformed calcite marble: diffusion accommodatedgrain boundary sliding as a result of subgrain rotationrecrystallization, J. Struct. Geol. 25 (2003) 1597–1613.

[30] C.D. Storey, D.J. Prior, Plastic deformation and recrystallizationof garnet: a mechanism to facilitate diffusion creep, J. Petrol. 46(12) (2005) 2593–2613.

[31] S.-I. Karato, M. Toriumi, T. Fujii, Dynamic recrystallization ofolivine single crystals during high-temperature creep, Geophys.Res. Lett. 7 (9) (1980) 649–652.

[32] D. Van der Wal, P. Chopra, M. Drury, J. Fitz Gerald, Relation-ships between dynamically recrystallized grain size and defor-mation conditions in experimentally deformed olivine rocks,Geophys. Res. Lett. 20 (14) (1993) 1479–1482.

[33] J.H.P. De Bresser, C.J. Peach, J.P.J. Reijs, C.J. Spiers, Ondynamic recrystallization during solid state flow: effects of stressand temperature, Geophys. Res. Lett. 25 (18) (1998) 3457–3460.

[34] J.H.P. De Bresser, J.H. Ter Heege, C.J. Spiers, Grain sizereduction by dynamic recrystallization: can it result in majorrheological weakening? Int. J. Earth Sci. 90 (2001) 28–45.

[35] M.R. Drury, Dynamic recrystallization and strain softening ofolivine aggregates in the laboratory and the lithosphere, in: D.Gapais, J.P. Brun, P.R. Cobbold (Eds.), Deformation Mechan-isms, Rheology and Tectonics: FromMinerals to the Lithosphere,Special Publication of the Geological Society of London, vol.243, Geological Society of London, 2005, pp. 143–158.

[36] J.F. Engeln, D.A. Wiens, S. Stein, Mechanisms and depths ofAtlantic transform earthquakes, J. Geophys. Res. 91 (B1) (1986)548–577.

[37] R.E. Abercrombie, G. Ekström, Earthquake slip on oceanictransform faults, Nature 410 (2001) 74–77.

[38] M.S. Boettcher, G. Hirth, B. Evans, Olivine friction at the base ofoceanic seismogenic zones, J. Geophys. Res. (in press).

[39] L.G.J. Montési, G. Hirth, Grain size evolution and the rheologyof ductile shear zones: from laboratory experiments to post-seismic creep, Earth Planet. Sci. Lett. 211 (2003) 97–110.

[40] A. Nicolas, N.I. Christensen, Formation of anisotropy in uppermantle peridotites—a review, in: K. Fuchs, C. Froidevaux (Eds.),Composition, Structure and Dynamics of the Lithosphere–Asthenosphere System, Geodynamics Series, vol. 16, AmericanGeophysical Union, 1987, pp. 111–123.

[41] C.E. Nishimura, D.W. Forsyth, The anisotropic structure of theupper mantle in the Pacific, Geophys. J. 96 (1989) 203–229.

[42] D. Mainprice, P.G. Silver, Interpretation of SKS-waves usingsamples from the subcontinental lithosphere, Phys. Earth Planet.Inter. 78 (1993) 257–280.

[43] H. Jung, S.-I. Karato, Water-induced fabric transitions in olivine,Science 293 (2001) 1460–1463.

[44] S.-I. Karato, P. Wu, Rheology of the upper mantle: a synthesis,Science 260 (1993) 771–778.

article in press - woods hole oceanographic institutionhirth_2006_15125.pdf · b woods hole...

Documents