The signature of devolatisation: Extraneous 40Ar systematics in … · 2013. 4. 25. · The signature of devolatisation: Extraneous 40Ar systematics in high-pressure metamorphic rocks

Available online at www.sciencedirect.com

www.else vier.com/locate /gca

Geochimic a et Cosmochim ica Acta 113 (2013) 94–112

The signature of devolatisation: Extraneous 40Ar systematics in high-pressure metamorphic rocks

Andrew J. Smye a,⇑, Clare J. Warren b, Mike J. Bickle c

a Jackson School of Geosciences, Univers ity of Texas, Austin, TX 78712, USA b Departme nt of Environme nt, Earth and Ecosystems , The Open University, Milton Keynes MK7 6AA, UK

c Departme nt of Earth Sciences, University of Cambridge, Cambridge CB2 3EQ, UK

Received 21 September 2012; accepted in revised form 11 March 2013; Available online 29 March 2013

Abstract

The validity of using the 40Ar/39Ar system for thermochronology relies on the assumption that the source mineral is sur- rounded by a grain boundary reservoir defined by an effective 40Ar concentration of zero. However, the presence of extraneous 40Ar (Are) in metamorphic rocks shows that this assumption is invalid for a significant number of cases. Are is common inmicas that have equilibrated under (ultra-)high pressure (ðUÞHP ) conditions: metasediments from six Phanerozoic ðUÞHPterranes yield apparent 40Ar/39Ar phengite ages K 50% in excess of the age of peak ðUÞHP conditions, whereas cogenetic mafic eclogites yield ages up to �700% older despite lower K2O concentrations. A model is developed that calculates Are

age fractions as a function of variable mica–fluid KD, bulk K2O and porosity under closed system conditions. Measured Are concentrations in mafic eclogites are reproduced only when porosities are K 10�4 volume fraction, showing that maficprotoliths operate as closed systems to advective solute transport during subduction. Porosities in eclogite-facies metapelites are K 10�2, reflecting loss of significant volumes of lattice-bound H2O relative to mafic rocks during subduction. Retention oflocally-generated 40Ar in mafic eclogites shows that the oceanic crust is an efficient vehicle for volatile transport to the mantle.� 2013 Elsevier Ltd. All rights reserved.

1. INTRODUCTION

The decay of 40K to 40Ar in mica has been routinely used for decades to constrain the rates of cooling and, therefore,exhumation of orogenic crust. Traditionally, the retention of radiogenic argon in a host mica grain is thought to becontrolled by thermally-activated diffusion (Harrison and McDougall, 1980a,b ). Using this concept, measured 40Ar/39Ar ratios correspond to the time since the mineral cooled below the temperature range over which 40Ar is lost from the grain by intracrystalline diffusion (Dodson, 1973 ).The concept of closure temperature (T c) is only valid given the following assumptions: (1) the distribution of 40Ar in amica grain is controlled by Fick’s law of diffusion in which

0016-7037 /$ - see front matter � 2013 Elsevier Ltd. All rights reserved.

http://dx.doi.org/10.10 16/j.gca.2013.03 .018

⇑ Corresp onding author.E-mail address: andrew.sm [email protected] xas.edu (A.J. Smye).

the diffusion of 40Ar produced by the decay of lattice-bound 40K is dependent on temperature and grain geometry; (2)the mica equilibrates with no argon present in the lattice (i.e. no inheritance), or contains argon of atmospheric composition (atmospheric 40Ar/36Ar = 295.5); (3) at temperatures above the closure temperature interval, the mica grain operates as an open system in that argon is efficientlyremoved from the grain and partitioned into a grain boundary reservoir with an effective zero concentration. Further- more, the analytical solution to the diffusion–productionequation requires that a target mica grain cools at a rate which is directly proportional to the inverse of the temperature. This is a reasonable assumption for slowly-cooled terranes, but a significant source of error for rocks that experience rapid cooling (Warren et al., 2012 ).

The closure temperature concept has been successfully applied to noble gas thermochronological studies across awide spectrum of metamorphic terranes. However, inhigh-dP/dT tectonic regimes such as subduction zones, the

http://dx.doi.org/10.1016/j.gca.2013.03.018

mailto:[email protected]

http://dx.doi.org/10.1016/j.gca.2013.03.018

A.J. Smye et al. / Geochi mica et Cosmochim ica Acta 113 (2013) 94–112 95

40Ar/39Ar system in mica commonly yields anomalously old ages, often in excess of the age of peak metamorphism estimated by U–Th–Pb, Sm–Nd, Lu–Hf and Rb–Sr geochronology—all systems in which the diffusion rates of the daughter nuclide are slower than that of argon (Ruffetet al., 1995; Scaillet, 1996; Sherlock and Kelley, 2002; DiVincenzo et al., 2006; Warren et al., 2010, 2011 ). Such ages show that (ultra-)high pressure (ðUÞHP ) mica grains have apropensity to include extraneous 40Ar (Are). Following the terminology developed by Damon (1968) and Dalrymple and Lanphere (1969), this extraneous 40Ar refers to the additional 40Ar responsible for measured 40Ar/36Ar ratios in excess of atmospheric values (>295.5). Excess 40Ar isthe component of 40Ar, apart from atmospheric 40Ar, that is incorporated into rocks and minerals by processes other than in-situ radioactive decay of 40K. Inherited 40Ar refers to any 40Ar that is produced within mineral grains by the decay of 40K before the event being dated. Both excess and inherited 40Ar are forms of extraneous 40Ar. It isimportant to note that once 40Ar becomes orphaned from parent 40K, it becomes excess 40Ar. For example, if 40Artransport distances are less than typical hand-sample length-scales, the sample can be described as containing inherited 40Ar despite the constituent minerals containing excess 40Ar. Inherited 40Ar is prevalent in granulite-facies metamorphic rocks (Foland, 1979 ) and mantle xenoliths (Kelley and Wartho, 2000 ), whereas excess 40Ar is common in fluid inclusions (Cumbest et al., 1994 ) and rocks adjacent to focused zones of high-strain (Vance et al., 1998b ) and granite intrusions (Kelley et al., 1986 ). Extraneous 40Ar isless common in greenschist and amphibolite grade metamorphic rocks and micas (Brewer, 1969 ).

The presence of Are shows that at least one of the govern- ing assumptions to the closure temperature concept is invalid. Given that Are concentrations often vary between, and within, individual ðUÞHP mica grains (e.g. Scaillet, 1996;Warren et al., 2010 ), uncertainty over the diffusion parameters Ea and D0 are unlikely to provide an explanation.Rather, the spuriously-old 40Ar/39Ar mica ages must becaused by either incomplete resetting of the 40K–40Ar system (inherited 40Ar), or the accumulation of radiogenic 40Ar (ex-cess 40Ar) in a grain-boundary reservoir to concentrations high enough to facilitate partitioning of argon back into the target mica grain. Incorporation of locally-derived excess 40Ar in the source mineral is controlled by mineral–fluidKD, porosity, and the external transmissive timescale of argon out of the local system (Baxter, 2003 ). Diffusive removal of inherited 40Ar is controlled by the interplay between time (t), effective diffusion radius (r) and temperature (the Fou- rier number: Dt=r2), such that it is possible for a 1 millimeter diameter mica grain to retain �95% inherited 40Ar for 0.5 Ma at 550 �C and 3 GPa (e.g. Warren et al., 2012 ). Fur- thermore, the strongly incompatible nature of argon (min-eral–fluid KD between 10�3 and 10�5; Kelley, 2002 ), means that the concentration of Are in a mica grain is critically dependent on the capacity of the grain-boundary reservoir to retain and accumulate argon, controlled by the availability and connectivity of the fluid phase.

Aqueous fluids promote chemical and physical evolution of the Earth’s lithosphere by accelerating reaction kinetics

during advective solute transport (Bickle and McKenzie,1987) and exerting a dominant control on rock-mechanical properties (Paterson and Kekulawala, 1979; Chopra and Paterson, 1984; Mackwell et al., 1985; Kohlstedt, 2006 ).In particular, fluids released by dehydration reactions insti- gate mantle melting and the production of arc magmas (Peacock, 1993; Schmidt and Poli, 1998 ) in addition tointermediate-depth earthquakes (Peacock, 1990 ). Despite their importance, values of permeability and porosity inmetamorphic rocks are poorly constrained. Measured grain-scale porosities in exhumed metamorphic rocks are between 10�3 and 10�6 volume fraction (Norton and Knapp, 1977 ). However, despite common usage, such measurements reflect time-integrated values, of which the rele- vance to metamorphic conditions is uncertain. The importance of understanding porosity is amplified by the fact that permeability of crustal media increases as the cu- bic or higher power of porosity (Norton and Knapp,1977), meaning that sub-percent-level porosities generated by metamorphic reactions have the potential to cause order-of-magnitude increases in permeability (Connolly,2010). Estimates of lithospheric permeabilities based onnear-surface (2–3 km) measurements and calculated metamorphic fluid-fluxes span 16 orders of magnitude (10�23–10�7 m2; Ingebritsen and Manning, 2010 ).

In this paper, we determine whether the concentrations of Are measured in ðUÞHP phengites can be explained byaccumulation of 40Ar produced by bulk-rock K2O in aclosed system, as a function of porosity and mica–fluidKD. This calculation places maximum limits on the porosity volume fraction present at the time of mica growth or equilibration under ðUÞHP conditions and provides a limiting case against which we consider whether removal of internally-derived 40Ar by fluid-loss during devolatisation is necessary in mafic eclogites and eclogite-facies metasediments.We show that porosities in mafic eclogites are consistent with hydration or limited fluid-loss during subduction,whereas Are concentrations in phengite from ðUÞHPmetasediments reflect loss of 40Ar related to devolatisation.Finally, we consider the combined effects of matrix phases,porosity, permeability and argon diffusivity on controlling the accumulation of Are in ðUÞHP rocks and the case inwhich valid “Dodsonian” (Dodson, 1973 ) cooling ages are predicted to occur.

2. OPEN VERSUS CLOSED SYSTEMS

In an open system, the diffusive flux of radiogenic argon produced by the decay of 40K within the mica grain isdependent on the temperature-sensitive diffusion coefficient,distance, and the chemical potential gradient of argon between the mineral and its surroundings (Fig. 1a. Once inthe grain-boundary reservoir, radiogenic argon is transported out of the mica–fluid system by advection and/or diffusion along grain boundary pathways at rates greater than the rate at which argon diffuses inside the host mica (Baxter, 2003 ). During devolatisation, elevated fluid pressures are expected to develop a connected porosity network through which advective solute transport will control the removal of 40Ar from the local system (e.g. Rubie, 1986;

40Ar

Open system

40Ar 40Ar 40Ar

40Ar

a b

Closed system

Fig. 1. Open and closed system scenarios for argon transport at T> T c. (a) Open system —the flux of argon through the grain boundary medium exceeds the diffusive flux of argon out of the source mineral. Radiogen ic argon diffuses out of the target mica, across the grain boundary, and is effectively transported by a combinat ion of advection and diffusion to an external sink. Either the volume or the transport velocity of the interconnected grain boundary reservoir ensures that it has an effectively infinite capacity for argon storage. (b) Closed system —the flux of argon from the source mineral exceeds the grain boundary flux of argon to an external sink. The grain boundary argon concentratio n increases as a function of time and K2O content of the protoli th, eventually reaching levels such that argon will partition back into the mica in detectable quantities.

96 A.J. Smye et al. / Geochi mica et Cosmochim ica Acta 113 (2013) 94–112

John et al., 2012 ). Following prograde fluid-loss, grain boundary transport of 40Ar will be governed by bulk diffu-sion which is expected to be effective over length-scales K 1 m in rock matrices that have experienced a 10-mil- lion-year-long regional metamorphic cycle (Rubie, 1986;Baxter et al., 2002 ). The end-member open system is definedby a grain boundary reservoir that has a zero 40Ar concentration, thus preventing any 40Ar from partitioning back into the host mica grain (Kelley, 2002 ). In reality no system is entirely open and all fluid–mineral systems contain trace quantities of Are (e.g. Kelley et al., 1986; Cumbest et al.,1994; Stuart et al., 1995 ). The ultimate macroscale sink for argon in any geological system is the atmosphere. How- ever, argon is highly incompatible and will strongly partition into grain boundary fluids in metamorphic rocks, orinto melts in magmatic systems (Kelley, 2002 ). More specif- ically, Etheridge et al. (1983) used Rayleigh-Darcy modeling to show that under lithospheric conditions where the metamorphic fluid pressure is greater than or equal to the minimum principal compressive stress, convective fluid flowwill be established through a network of connected veins,providing a mesoscale sink for crustal volatiles. Given that metamorphic porosities are expected to be extremely small (10�7–10�3 vol. fraction; Ganor et al., 1989; Skelton et al.,2000), the role of deformation in forming interconnected vein systems is likely an important factor in the transport of volatiles such as argon (Vance et al., 1998a ). Fluid inclusions also have the potential to act as a sink for smaller quantities of argon (Cumbest et al., 1994 ), relative to vein and melt systems. The general assumption of open system argon transport in a wide spectrum of crustal environments is supported by the fact that both the K–Ar and 40Ar/39Armethods commonly yield geologically meaningful age data

that are free from analytically-detectable Are (e.g. Wijbransand McDougall, 1987, 1988 ).

At temperatures in excess of mineral closure temperatures, the accumulation of radiogenic 40Ar in a metamorphic system is controlled by the degree to which radiogenic 40Ar is removed to an external sink. This trans- missivity depends on the active mechanism of grain boundary solute transport and can be advective and/or diffusive.Baxter (2003) defined the transmissive timescale (sT ) asthe characteristic time for 40Ar to escape through the local intergranular medium to a sink for argon, located at a distance L from the source mineral. For diffusive transport,sT ¼ L2

D�, where D� is the total effective diffusivity of argon,whereas for advective transport, sT ¼ L

V �, where V � is the total effective velocity (Baxter, 2003 ). In order for a metamorphic system to retain locally-derived 40Ar, sT must becomparable or greater than the timescale of 40Arproduction.

On cooling through the closure temperature interval,attainment of argon thermodynamic equilibrium becomes limited by the rate of intracrystalline diffusion of argon through the source mineral, not solely by sT . Baxter(2003) showed that in systems where the ratio of the timescale of local diffusional exchange between source mineral and grain boundary, and sT (the diffusional exchange parameter, De, of Baxter, 2003 ) exceeds 15, the source mineral diffusively exchanges fast enough to maintain local equilibrium. In the case when De < 15, to be expected atsub-closure temperatures, the source mineral cannot exchange argon fast enough to maintain local equilibrium,leading to retention of radiogenic 40Ar.

Rates of intergranular diffusion along dry grain boundaries are up to 16 orders of magnitude slower than for


aqueous fluid-bearing grain boundaries (10�24 versus10�8 m2 s�1 Rubie, 1986; Scaillet, 1996 ), but still 103 ordersof magnitude faster than typical rates of volume diffusionthrough a crystalline lattice Manning (1974). Accordingly,fluid-poor systems will have longer values of sT comparedto fluid-rich systems. In a fully-closed system a strong correlation should exist between Are and bulk K2O as 40Ar islocally produced in the rock and not transported over length scales greater than grain dimensions. Foland (1979)and Sherlock and Kelley (2002) observed such a correlation between Are concentration and bulk K2O in fluid-poorrocks of the Arden pluton in the Appalachian mountains and the Tavsanli Zone, NW Turkey respectively. A wealth of evidence from in-situ UV laser and bulk step-heating studies demonstrate that ðUÞHP metamorphic rocks provide the most extreme examples of Are accumulation under fluid-poor, low sT conditions (e.g. Li et al., 1994; Ruffetet al., 1995; Scaillet, 1996; Sherlock and Kelley, 2002; DiVincenzo et al., 2006; Warren et al., 2011 ).

3. EXCESS ARGON IN ðUÞHP METAMORPHIC

TERRANES

Fig. 2a shows 40Ar/39Ar phengite ages from a global compilation of ðUÞHP terranes; the ages are normalised to the respective timing of ðUÞHP metamorphism, constrained by independent geochronology (U–Th–Pb; Sm–Nd; Lu–Hf; Rb–Sr). All data plotted correspond to phengite that grew during the peak of ðUÞHP metamorphism. The majority of these ages lie within 100% excess of the peak ðUÞHP age. However, phengite from the Dora Maira complex, Western Alps, records ages >700% older (Scaillet,1996) than the peak of Alpine ðUÞHP metamorphism

Age

of p

eak

(U)H

P ev

ent

1 10 100 1000

Marble

Lithologies Mafic

Orthogneiss Pelite

Excess 40Ar age equivalent (%)

Dora Maira

Dabie Sulu

Oman

Turkey

Eastern Alps

Sezia Zonea

Fig. 2. Excess argon in ðUÞHP rocks. (a) Global compilatio n of phengpercentage excess relative to independent ly-derived ages of peak ðUsquares = metapelite; diamonds = marble; triangles = orthogneiss. Data sDi Vincenzo et al. (2006), Sulu – Ames et al. (1993), Giorgis et al. (2000); Oet al. (1999), Sherlock and Kelley (2002); Eastern Alps – Smye et al. (201et al. (1999). (b) Contours of constant mica K2O plotted as a function o(atoms g�1). Represen tative single grain phengite 40Ar/39Ar data from an2011) are plotted over the contours. Atomic quantities calculated using aphengite K2O value of 10.34 (n = 15; 1.88 % 2r, Warren et al., 2011 ) and 1r.

(35.1 Ma, Rubatto and Hermann, 2001 ). Similarly, Giorgiset al. (2000) record phengite ages from the Sulu eclogite terrane, China, between 80% and 400% older than the accepted age of peak ðUÞHP metamorphism (217 Ma, Ameset al., 1993 ). It is also apparent from Fig. 2a that rock composition exerts a strong control on the concentration of Are

(Scaillet et al., 1992; Baxter et al., 2002 ). In general,40Ar/39Ar mica ages from mafic eclogites are older than those from white mica in cogenetic metapelites. For example, Di Vincenzo et al. (2006) showed that phengite from mafic eclogites of the UHP Dora Maira terrane of the Wes- tern Alps yielded ages >52 Ma older than phengite from associated orthogneisses and >34 Ma older than phengite derived from eclogite-facies marbles. Provided that Are islocally produced, these data show that despite higher concentrations of K2O, metapelites lose a larger fraction ofradiogenic 40Ar during devolatisation than maficlithologies.

Elevated 40Ar/39Ar white mica ages can be explained byincorporation of Are, or by the removal of K2O. The latter is typically associated with margarite formation in response to chemical weathering and is relatively uncommon (DeJong et al., 2001 ). It is important to note that the effect ofexcess 40Ar on apparent age will be maximized in low- K2O micas where the atomic ratio of excess 40Ar to supported, radiogenic, 40Ar exceeds that expected in micas with higher weight % K2O. To illustrate this, Fig. 2b shows contours of constant muscovite weight % K2O plotted as afunction of excess age and number of atoms of excess 40Ar. Single grain phengite 40Ar/39Ar data from a represen- tative mafic eclogite, collected from the Eastern Alps (CWT13 Warren et al., 2011 ), plot along a linear array,subparallel to �10 wt.% K2O, showing that intergrain age

30

25

20

15

10

5

0

)aM(

egAssecxE

3.02.52.01.51.00.50.0

Atoms of excess 40Ar (x1015 g-1)

8 wt.%

K 2O

9 10 11 12

b

ite single-grain 40Ar/39Ar ages from ðUÞHP terrane s expressed asÞHP metamorph ism. Symbol notation: circles = mafic eclogite;

ources: Dora Maira – Scaillet (1996), Rubatto and Hermann (2001),man – Warren et al. (2003), Warren et al. (2010); Turkey – Sherlock1), Warren et al. (2011); Sesia Zone – Ruffet et al. (1995), Rubatto f excess age in Ma, and the number of atoms of incorporate d Are

East Alpine mafic eclogite (orange circles; CWT13 Warren et al.,diffusive cooling age of 28.05 Ma (Warren et al., 2011 ), an average

a muscovite density of 2.81 g cc�1 (Deer et al., 1997 ). Error bars are


variations are explained by variable Are concentrationsalone and that the effect of intergrain variation in 40K concentration is small. Importantly, these calculations show that single phengite grain ages reflect local fluid–mica isotopic disequilibrium and that the scale of argon isotopic equilibrium under ðUÞHP conditions is smaller than the intergranular distance (10�3–10�6 m).

Stable isotope data from ðUÞHP rocks are also consistent with short length scales of fluid–rock interaction indi- cated by the ubiquity of Are. Several studies document observations that ðUÞHP rocks often retain recognizable vestiges of pre-subduction fluid–rock interaction. For example, elevated d18O and dD values from eclogitic par- ageneses of the Sulu UHP terrain, China, were acquired in an ancient hydrothermal system prior to subduction and were not reset beyond millimeter to centimeter length-scales during coesite-grade metamorphism (Rumbleand Yui, 1998 ). Similarly, pre-subduction oxygen isotope heterogeneities in eclogites of the Western Alps imply time-integrated fluid fluxes across lithological contacts aslow as 1.6 cm3 cm�2 (Philippot and Rumble, 2000 ). Fur- thermore, Getty and Selverstone (1994) showed that millimeter-scale d18O heterogeneities associated with tuffaceousbanding in an eclogite from the Eastern Alps were preserved throughout HP metamorphism and subsequent Bar- rovian overprinting.

4. FLUID–MICA ARGON EXCHANGE MODEL

At subduction zones hydrated lithosphere returns to the mantle, heating up during descent and releasing mineral- bound fluid by devolatisation reactions (Peacock, 1990 ).The fraction of protolith-derived argon that is retained bymica on equilibration under ðUÞHP conditions reflects the extent to which the rock has behaved as an open or closed system to volatile loss during subduction. In the following section we present a model for calculating the concentrations of Are in phengite as a function of variable KD, porosity, and mica modal-abundance for the case where no argon is lost between the timing of protolith formation and mica growth/equilibration. When applied to natural phengite 40Ar/39Ar datasets from the Eastern Alps and Oman, this allows the metamorphic porosity to be estimated and provides a limiting case to assess the extent of fluid-loss during subduction. Model results help constrain the mechanisms by which argon and other volatiles are transported into the mantle, in addition to better informing sampling strat- egy for 40Ar/39Ar geochronology in ðUÞHP terranes.

4.1. Solubility of Argon

The relative solubilities of argon between cogenetic ðUÞHP phases control the distribution of Are. The solubility of argon in hydrous fluids strongly depends on both temperature and salinity (Weiss, 1970; Crovetto et al., 1982;Smith and Kennedy, 1983 ), varying between 10 and 100 ppm bar �1 across the likely range of metamorphic temperatures and fluid salinities (Kelley, 2002 ). The solubility of argon in minerals is less well-constrained, largely due to a lack of experimental data and the propensity for argon

to adsorb onto mineral defects (e.g. Broadhurst et al.,1992). Existing data show that argon solubilities in silicate minerals are J 3 orders of magnitude smaller than fluidvalues: clinopyroxene – 0.09 ppb bar �1 (Brooker et al.,1998); olivine – 1 ppb bar �1 (Brooker et al., 1998 ); plagioclase – 1.8 ppb bar �1 (Kelley, 2002 ) and K-feldspar –0.7 ppb bar �1 (Wartho et al., 1999 ). Measurements of argon solubility in quartz vary over at least 3 orders of magnitude, between 3.75 ppb bar �1 (Roselieb et al., 1997 ) and �2000 ppb bar �1 (Watson and Cherniak, 2003 ). If the solubility data of Watson and Cherniak (2003) are applicable to natural systems, quartz has the potential to be the dominant sink of crustal argon (Baxter, 2003 ). Whereas fluidinclusions in vein quartz are abundant and likely have the potential to act as a sink for significant quantities of argon (Cumbest et al., 1994 ), further work is required to under- stand the role of fluid inclusions in rock-forming quartz as a sink for volatiles.

There is a dearth of experimental data constraining the solubility of argon in hydrous minerals used for 40Ar/39Argeochronology. Although Onstott et al. (1991) undertookpreliminary experiments to show that argon solubility inbiotite is between 3.6 and 36 ppm bar �1, the accuracy ofthe results was compromised by inclusion of atmospheric argon. Nevertheless, the observation that muscovite is commonly observed with lower Are concentrations than cogenetic biotite (e.g. Brewer, 1969; Roddick et al., 1980 )confirms a relative solubility order for the micas. Measure- ments on phlogopite indicate that argon solubilities in micas are �1 order of magnitude greater than K-feldspar (1.8 ppb bar �1: Roselieb et al., 1999 ). To date, there are no argon solubility data available for phengite. Accord- ingly, these relative solubility data suggest that KD betweenmuscovite and fluid lies between 10�5 and 10�3 under metamorphic conditions.

4.2. The model

We calculate apparent mica ages in a model in which the exchange of 40Ar between mica and a fluid-filled porosity ina closed system is simulated according to mass balance.This is equally applicable to mica fractions containing either inherited or excess 40Ar. Controlled by a particular value for the partition coefficient, inherited 40Ar will diffuseout of the grain until a balance is attained between the grain-boundary network and the grain itself. Similarly, inthe case of excess 40Ar, argon will diffuse into the grain until the same balance is reached. Critical parameters are porosity (/), volume fraction of mica (mm), mica–fluid partition coefficient (KD) and relative densities (q/; qwr and qm). Mod- el mica is treated as muscovite, incorporating the range inKD discussed above. No argon is lost or gained in the rock volume under consideration between the time of protolith formation and the time of mica growth/equilibration. The parental link between 40K and 40Ar is preserved on the volume scale, but not necessarily on the grain scale.

Model assumptions are illustrated in Fig. 3. Following the approach of Nabelek (1987) and Banner and Hanson (1990), the distribution of 40Ar in both the mica and fluidare calculated before and after equilibration. Argon


exchange between mica and fluid is simulated as being instantaneous, which is acceptable for the model simula- tion; in reality argon exchange will occur over a finite time period (sss, Baxter, 2003 ), controlled by the transmissive timescale and total local sink capacity. Model mica contains 10.9 wt.% K2O and has a density (qm) of 2.81 g cc�1,analogous to typical ðUÞHP phengite (Deer et al., 1997 ).The pore-fluid has a density of (q/) 1 g cc�1, in the range appropriate for high P–T conditions (Norris and Henley,1976). The following section outlines the necessary equa- tions to simulate closed system 40Ar exchange.

Prior to fluid–mica interaction, the concentration of40Ar in the fluid–rock system (Ci) is defined by mass balance where the total argon is the sum of that in the fluidphase (C/;i) and the solid phase (Cwr;i):

Ci ¼ FC/;i þ ð1� F ÞCwr;i ð1Þ

F, the mass fraction of fluid present in the whole-rock, isrelated to the porosity volume fraction by:

F ¼/q/

/q/ þ ð1� /Þqwrð2Þ

Fig. 3. Schematic figure illustrating the differences between inherited an(metamorphic recrystallisatio n and isotopic equilibra tion), 40Ar builds minerals recrysta llise. In an open system (ai), all the protolith-deriv edprotolith-de rived 40Ar is retaine d in the local rock volume where it partsolubilities. In this case, new K-bear ing minerals incorporate excess Arduring metamorp hism (ci), will contain inherited 40Ar, which is removed athrough the mineral closure temperatu res, ”new” Ar (light gray text) starare contamin ated with excess (bii) and inherited (cii) 40Ar. Panel d links tundergoin g subduction- related metamorph ism.

The concentrations of 40Ar in the solid (Cwr) and fluid(C/) after equilibration are calculated using a bulk argon distribution coefficient (D ¼ KDmm, where mm is the mass fraction of mica present):

Ci ¼ FCwr

Dþ ð1� F ÞCwr ð3Þ

The equilibrated 40Ar concentration in mica is calculated by dividing Cwr by the mass fraction of mica. This yields a concentration which, together with the initial system 40K concentrations, is converted to atomic values and used to calculate the age of the mica fraction. It is important to recognize that the calculated ages are the maximum expected for a closed system given protolith and metamorphic ages.

Previous models of the accumulation of Are in mineral–fluid systems employ a reactive-transport approach, assess- ing the degree to which partitioning of 40Ar into the rock matrix retards advective–diffusive removal of Are out ofthe rock volume (Baxter et al., 2002; Baxter, 2003 ). The particular strength of this approach is that it allows predictions of the spatial and temporal distribution of Are in a

d excess 40Ar. Between time t1 (protolith formation) and time t2

up in K-bearing minerals in the protolith. At time t2, K-bearing 40Ar is removed from the local rock volume. In a closed system,itions between fluid and matrix phases according to relative argon into their structur es (bi). A mineral that escapes recrysta llization ccordin g to the efficiency of volume diffusion. As rocks start to cool ts to be retaine d. Either a ”true” cooling age is yielded (aii) or ages he events, t1, t2 and t3 with the temperature (T) evolution of a rock


rock volume to be made. Reactive-transport modeling ofAre removal during devolatisation is complicated by the time-variant nature of porosity (DePaolo and Getty,1996) and the need for detailed K–Ar isotopic profiles toconstrain the length scale of argon transport through the rock matrix (Baxter et al., 2002 ). In comparison, the mass balance model used here evaluates the time-integrated effects of fluid-loss on Are concentrations. By assuming that no internally-derived 40Ar is lost from the rock volume since the timing of formation, calculated porosities reflectthe time-integrated fluid fraction required to explain the concentrations of Are measured in ðUÞHP phengites. The porosities are then treated as maximum values against which thermodynamic predictions of fluid-loss can becompared.

4.3. Model results

Model results are shown as contours of constant fluid–mineral partition coefficients (Fig. 4a) and mica volume fraction (Fig. 4b) on a plot of excess age fraction against porosity. Excess age fractions are defined as the fraction of Are retained by the mica at T > T c. The maximum excess age fraction is the age of the protolith less the age of metamorphism; excess age fractions greater than 1 are caused bythe addition of excess 40Ar. For example, an excess age fraction of 0.5 for a mica grain which cools through T c at30 Ma after residing at T > T c for 10 Ma corresponds toa measured age of 35 Ma. Porosities correspond to the maximum volume fraction of fluid necessary to explain aspecific mass fraction of Are in muscovite at the time the rock cools through the closure temperature. At T < T c, argon exchange will be governed by reaction-rate-controlled dissolution–precipitation mechanisms (DePaolo and Getty,1996), which operate at faster rates than volume diffusion ofargon in mica (Baxter, 2003 ). Fig. 4 shows that for any given value of KD or mm, the concentration of Are within the mica fraction increases as a function of decreasing /. Con- tours of constant KD (Fig. 4)a and muscovite fraction asymptotically approach end-member closed (excess age fraction = 1) and open (excess age fraction = 0) system

1.0

0.8

0.6

0.4

0.2

0.0

noitcarFegA

ssecxE

a

10-5

10-4

10-3

10-6 10-5 10-4 10-3 10-2 10-1

Porosity (vol. frac.)

Fig. 4. Mica–fluid argon exchange model. (a) Excess age fraction plotteKDm�/

between 10�3–10�5. Calculatio ns employ a value of mm of 0.3 and as(b) Excess age fraction–/ curves for a range of mica volume fractions ra

conditions where all Are resides in the mica and fluid,respectively.

Typically, eclogite-facies metapelites contain �20–60 vol.% phengite, compared to <10 vol.% in mafic eclogites. Fig. 4b shows that in order for phengite in ðUÞHP pel-ites to have younger apparent 40Ar/39Ar ages than maficeclogites the effective porosity must be at least an order ofmagnitude larger in pelites than in mafic rocks. Given that the model maximizes phengite Are concentrations byassuming no loss of Are from the fluid–rock system along with minimum porosities, it acts as a limiting case against which the effects of devolatisation in ðUÞHP rocks can beevaluated. Violation of the model assumptions by either larger effective porosities, Are removal by fluid loss, or partitioning into local mineral sinks, reduces the magnitude ofthe calculated excess age fraction.

5. APPLICATION OF THE MODEL TO NATURAL 40AR/39AR DATASETS

In order to assess whether the Are in ðUÞHP rocks might be generated in-situ prior to the peak of ðUÞHP metamor-phism, we model argon systematics in phengites using data from two recently published 40Ar/39Ar single-grain phengite datasets: the Tauern Window HP nappes of the Eastern Alps (Warren et al., 2011 ) and the As Sifah HP terrane inOman (Warren et al., 2010 ). Due to poorly-constrained values of KD for matrix phases, we only consider argon exchange between the mica and fluid fractions, treating the remainder of the rock mass as inert.

5.1. The Tauern Window

The Tauern Window in south-central Austria exposes the metamorphosed remnants of the European continental margin and Tethyan oceanic basin that were subducted and underplated during the collision between Europe and Adria. The Eclogite Zone is a Triassic–Jurassic volcano- sedimentary succession (Kurz et al., 1998 ) which was subducted to �20–26 kbar and �550 �C between �30–37 Ma(Glodny et al., 2005; Smye et al., 2011 ). Following

1.0

0.8

0.6

0.4

0.2

0.010-6 10-5

0.1%1%

10%

100%

10-4 10-3 10-2 10-1

b

Porosity (vol. frac.)

d as a function of fluid volume fraction (porosity; /) for values ofsume the grain-boundary fluid has a zero 40Ar initial concent ration;nging from 0.001 to 1. Calculatio ns use a KD of 10�4.


eclogite-facies metamorphism, the Eclogite Zone was over- printed by a regional Barrovian event (ca. 7 kbar, 520–570 �C; Zimmermann et al., 1994 ) between 27 and 31 Ma(Inger and Cliff, 1997; Glodny et al., 2005; Gleiner et al.,2007). On the basis of numerical diffusion calculations (Wheeler, 1996; Warren et al., 2011, 2012 ), phengite (0.5 mm diameter) growing under HP conditions prior toBarrovian metamorphism, and subsequently cooling at30 �C Ma�1, is expected to yield an age of 28.05 Ma.

Sample CWT13 is a foliated eclogite collected from the central portion of the Eclogite Zone, containing: garnet + omphacite + quartz + glaucophane + muscovite + epidote + calcite + ankerite + rutile + paragonite.Paragonite is subordinate and smaller (<400 lm diameter)than muscovite; collectively the mica volume fraction is<0.05. Muscovite is celadonite-rich and contains �6.9–7.2 c.p.f.u. Si per 22 O; this, together with their textural alignment with omphacite and garnet grains, indicates that they grew during the HP event. Single-grain fusion 40Ar/39Ar laserprobe ages from 10 white-mica grains fall between 36 and 59 Ma. Apparent 40Ar/39Ar phengite ages cluster between 36 and 43 Ma whereas paragonite analyses yield ages in the range of 48–59 Ma (Warren et al., 2011 ).The protolith age of sample CWT13 is loosely constrained to �110 Ma (Bousquet et al., 2002 ).

Sample CWT12 is an Eclogite Zone metapelite comprised of garnet, quartz, phengite, chlorite, albite, rutile and kyanite. Phengite is abundant (�0.5 volume fraction)and has a high Tschermak content (6.73 c.p.f.u. Si),

Table 1Model parameter s used for the East Alpine and Oman 40Ar/39Arphengite datasets.

Eastern Alps Oman

CWT13 CWT12 CWO40 8 CWO410 tpr 110 300 298 298 tc 28.05 28.05 75 75tm 33 33 80 80mm 0.05 0.5 0.05 0.4 40Ar/39Ar phengite age

48–59 37–44 83–125 77–89

Excess age fraction 0.24–0.37

0.03–0.05

0.04–0.22

0.008–0.06

Fig. 5. Mica–fluid argon exchange model applied to single grain 40Ar/39Aterranes. Calculatio ns use a KD of 10�4. Model parameter s are present40Ar/39Ar ages measured.

indicative of growth under HP conditions. A total of 17phengite grains (>820 lm diameter) yielded single-grain apparent 40Ar/39Ar ages between 37 and 44 Ma (Warrenet al., 2011 ). A bulk-protolith age of 300 Ma is consistent with the Eclogite Zone metapelites being derived from subducted Variscan crust (Kurz et al., 1998 ).

Table 1 details the model parameters used to simulate closed system argon exchange in samples CWT13 and CWT12, the results of which are shown in Fig. 5a. Using a value of KD ¼ 10�4, porosity volume fractions of �3–6 �10�5 and 4–6 �10�3 are required to account for the observed excess age fractions in samples CWT13 and 12respectively.

5.2. The As Sifah eclogite terrane, Oman

Mafic eclogites, intercalated with pelitic and calcareous metasediments, are exposed toward the northeastern edge of a window through the Semail ophiolite in Saih Hatat,northeastern Oman. The eclogites and host metasediments experienced peak conditions of 16–22 kbar and 520–550 �C (El-Shazly, 2001; Warren and Waters, 2006 ) during subduction of the Permian shelf sequence. The protolith age of the succession is best constrained by a 298 Ma zircon ex- tracted from a tuffaceous horizon (Gray et al., 2004 ). U–Pbzircon ages of 79.06 �0.32 Ma (2r, ID-TIMS; Warrenet al., 2003 ) and 82.06 � 0.83 Ma (1r, SHRIMP; Grayet al., 2004 ) are interpreted as representing the timing ofeclogite-facies metamorphism. Rb–Sr phengite–whole-rockisochrons (El-Shazly et al., 2001 ) and U–Pb rutile ages (Warren et al., 2005 ) of 78 � 2 Ma and 79.6 � 1.1 Marespectively, suggest that the HP terrane cooled quickly during exhumation; zircon fission-track ages span 66–70 Ma (Saddiqi et al., 2006 ). Numerical diffusion modeling (Wheeler, 1996; Warren et al., 2012 ) yields ages between 72.5 and 78.4 Ma (Warren et al., 2010 ) assuming that mica grew concomitant with peak T (550 �C) at 79–82 Ma and cooled at rates between 10 and 50 �C Ma�1 in an open system.

Sample CWO408 is a mafic eclogite comprised ofgarnet + omphacite + phengite + crossitic glaucophane +epidote + quartz + rutile. Phengites are characterised byhigh-Si core regions (>7.2 Si c.p.f.u. per 22 O) and

r phengite datasets from the Tauern Window (a) and Oman (b) HP

ed in Table 1. Shaded rectangles represen t the range in apparen t


marginally lower rim concentrations (6.9 Si c.p.f.u.); texturally, phengite belongs to the HP assemblage. Whole-rock XRF analysis yields a K2O content of 2.06 wt.% (Warren,unpublished data ). Single-grain fusion 40Ar/39Ar analysis of15 phengite grains (420–840 lm size fraction), yielded ages between 83 and 125 Ma (Warren et al., 2010 ) .

Sample CWO410 is an eclogite-facies phengite schist collected from a horizon enclosed within a mafic eclogite bou- din. Bulk-rock K2O content is 4.11 wt.% (Warren,unpublished data ). The sample is comprised of major phengite and quartz with subordinate rutile, epidote, calcite, garnet and hematite. Celadonite–muscovite exchange is only weakly present in phengite (typical core: �7 Si c.p.f.u.; typical rim: 6.9 Si c.p.f.u.). Greenschist-facies alteration is restricted to regions surrounding quartz-filled veins,commonly measuring between 1 mm and 20 cm. Single- grain fusion analysis of 19 phengite grains (840–1000 lmsize fraction) yielded apparent 40Ar/39Ar ages between 76.8 and 89.2 Ma (Warren et al., 2010 ).

Model results for CWO408 and CWO410 are shown inFig. 5b; further parameters are detailed in Table 1. Assum- ing that no argon is lost from the rock since the time of protolith formation, phengite in sample CWO408 equilibrated with a grain-boundary reservoir � 6� 10�5–4� 10�4 vol-ume fraction. Fluid fractions up to two orders of magnitude larger, � 3� 10�3–2� 10�2, are calculated for the pelitic sample, CWO410.

6. DISCUSSION

Extraneous argon can originate from both the radiogenic argon produced in-situ within a ðUÞHP mica grain at temperatures above the closure temperature and also from diffusive equilibration, or crystallization, of a mica grain in a grain boundary fluid that contains 40Ar. Incorpo- ration of in-situ derived Are would lead to mica ages between the time of equilibration under ðUÞHP conditionsand cooling through the closure temperature range. Pheng- ite 40Ar/39Ar ages greater than the residence time of the ðUÞHP terrane (<30 Ma; Kylander-Clark et al., 2012 ) can only be explained by retention of inherited argon. By defi-nition, inheritance does not account for apparent 40Ar/39Arages older than the age of the protolith, which must be explained by inclusion of excess 40Ar. The model developed above shows that extraneous argon concentrations inðUÞHP phengites can be explained by partitioning of inherited or locally-generated 40Ar in a closed system with porosity K 10�4 volume fraction for mafic eclogites and K 10�2

for metapelites. These porosities are upper-bound estimates of the time-integrated fluid volume fraction present at temperatures above phengite T c. It is important to note that innatural systems, rock porosity is transient and fluctuatesduring metamorphism (Connolly, 2010 ). Accordingly, the accuracy of the porosity estimates is contingent on estab- lishment of mica–fluid equilibrium for argon, dependent on the mica fraction being diffusively open for argon exchange. Should thermodynamic equilibrium not be attained, for example, due to short residence times atT > T c, the porosity estimates represent upper-bounds asthe mica fraction is under-filled with respect to Are.

Furthermore, if 40Ar is removed from the system by grain boundary transport mechanisms or partitioning into other mineral phases, apparent K–Ar (40Ar/39Ar) ages will be re- duced, increasing the porosity estimates. Therefore, it isimportant to consider the factors that will control the loss of 40Ar during prograde metamorphism: (1) storage and release of argon from local matrix phases; (2) the diffusivemobility of argon; (3) the availability of a fluid phase; (4)the connectivity of the fluid phase.

6.1. Storage and release of argon from local matrix phases

Storage of locally-derived 40Ar in matrix minerals will lower the amount of Are available for partitioning between the fluid-filled porosity and mica fractions, leading to lower apparent 40Ar/39Ar mica ages. However, the propensity for argon to partition strongly in the fluid phase (KD between103 and 105) means that fluids will dominate the argon storage capacity of the rock if they are present in sufficient mass fractions. The total argon capacity of the local mineral sink reservoir is controlled by mineral–fluid KD and mineral modes, in addition to the mass fraction of fluid-filled porosity present (the Total Local Sink Capacity of Baxter, 2003 ).Given that argon solubilities in rock-forming phases are poorly-constrained (see Kelley, 2002 for a review) and that metamorphic porosities are vanishingly small (10�6–10�3;Norton and Knapp, 1977; Skelton, 2011 ), it is important to evaluate the potential for local matrix phases to act assignificant sinks for Are.

Fig. 6a shows contours of constant mineral sink mode (1, 10 and 70 volume %) calculated for the cases in which the sink phase strongly (KD ¼ 10�2) and weakly (KD ¼ 10�5) partitions argon. Values of KD reflect the range in experimentally-determined argon solubility for silicate minerals. As expected, when argon solubility in the local mineral sink phase is lower than in the mica fraction, the fluid-filled porosity dominates the total local sink capacity and the mica excess age fractions are lowered by < 20%at extremely small values of porosity (< 10�6), even when the entire remainder of the rock volume (0.7 volume fraction) is open to sequester argon. However, when argon partitions strongly (KD ¼ 10�2) into local mineral phases, assuggested by the solubility data of argon in quartz derived by Roselieb et al. (1997), the effect on decreasing mica Are

concentrations is dramatic (dashed lines Fig. 6a). Mineral modes as small as 1% lower mica excess age fractions by� 80% at realistic metamorphic porosities (10�4), showing that the total local sink capacity will be dominated by the mineral phase. The presence of such a matrix phase would play a critical role in preventing the accumulation of Are incoexisting K-bearing phases, permitting successful thermochronology.

In addition to acting as sinks for argon, local K-bearing mineral phases may also break down at temperatures greater than T c, liberating excess radiogenic argon into the grain boundary reservoir, which may subsequently partition between the fluid-filled porosity and mica fraction. For example, in mafic protoliths, chlorite and actinolite both contain trace amounts of K2O (�1 weight %) and are predicted tobecome unstable at temperatures greater than 500 �C at

Fig. 6. The role of matrix phases. (a) Effect of local mineral sink phases on excess age fraction –/ curves. Relations calculated for a system with mm ¼ 0:3;KDm�/

¼ 10�4 and assume that the grain bound ary has a zero 40Ar initial concentratio n. The solid line is the case where mica isthe only argon-bea ring phase. Dashed lines are calculat ed for the case when 0.01, 0.1 and 0.7 volume fraction of a sink phase with a KD of 10�2

coexist with the mica fraction . Dash-dot lines use a KD of 10�5 for the sink phase; (b) Effect of local mineral source phases. System parameter sare as used in (c) solid line represents the case in which no source phases are present. Dotted line (scenario 1) is the case for a source phase of0.01 volume fraction, containin g 1 wt.% K2O and a tsc : tres ratio of 1:1. Dashed line (scenario 2) is the case for a source phase of 0.1 volume fraction, 11 wt.% K2O and a tsc : tres ratio of 5:1. Dash-dot line (scenario 3) is the case for a source phase of 0.7 volume fraction , 11 wt.% K2Oand a tsc : tres ratio of 10:1.


�21 kbar (Fig. 8a). Prior to the initiation of garnet growth,plagioclase is likely to be an important carrier of inherited Ar in both mafic and pelitic rocks during the early stages ofsubduction. In alkali-rich protoliths, primary K-feldspar and low- P metamorphic muscovite also have the potential to act as ‘vehicles’ for argon transport. The importance ofthis effect is evaluated by calculating the increase in excess age fraction observed after introducing quantities of excess 40Ar produced by (1) a small modal fraction (1 volume %)of a mineral containing trace quantities of K (1 weight %)that is stable for a period of time (denoted tsc), equivalent to the duration of time the mica fraction resides above T c

(tres; tsc : tres ¼ 1 : 1); (2) a large volume fraction (10 volume %) of a high-K mineral (11 weight %), stable for an interme- diate duration prior to breakdown (tsc : tres ¼ 5 : 1); (3) alarge volume fraction (10 volume %) of a high-K mineral (11 weight %), stable for a long duration prior to breakdown tsc : tres ¼ 10 : 1. Calculations use a mica volume fraction of 0.3 and a KD of 10�4; results are shown in Fig. 6b ascurves of excess age fraction plotted against porosity for each scenario. As expected, scenario 1 shows that the effectof small modal proportions of minerals containing trace amounts of K2O is small, increasing mica excess age fractions by < 1%. Scenarios 2 and 3 illustrate the importance of the age of the source phase, chosen here to reflect metastable biotite, expected to break down in place of muscovite at T > 550�C and P > 20 mbar in mafic rocks (Fig. 8a). For the extreme case when the duration of radiogenic in-growth for the source fraction exceeds that of the muscovite fraction by a factor of ten (scenario 3), excess age fractions are elevated by � 30%. These calculations highlight that mica ages in excess of the protolith age require the addition of excess 40Ar, plausibly sourced from metastable biotite and other high-K minerals that escaped recyrstallisation during prograde metamorphism due to sluggish kinetics.However, the absence of phengite 40Ar/39Ar ages in excess of protolith ages from ðUÞHP rocks (Fig. 2b) suggests that

this phenomenon is unlikely in ðUÞHP systems and that Are

is locally-derived.

6.2. The diffusive mobility of argon during subduction

Diffusive transport of argon can proceed via solid-state volume diffusion within mica grains, along dry or wetted grain boundaries, and within a fluid-filled porosity. Intra- crystalline volume diffusion of argon operates between � 10�15 and 10�22 cm2 s�1 (Rubie, 1986; Scaillet, 1998 ) significantly slower than � 10�4–10�13 cm2 s�1 (Walther and Wood, 1984; Rubie, 1986; Scaillet, 1998 ) typical of diffusionthrough a fluid phase. Therefore, it is important to address whether the length scales of diffusion characteristic of asubducting rock will hinder the effective removal of protolith-derived argon from the grain boundary network and also the removal of lattice-bound 40Ar from phengite grains residing at ðUÞHP conditions.

Using the experimental parameters of Harrison et al.(2009), Fig. 7a shows that argon within muscovite following typical subduction P–T trajectories (Peacock, 1993 ) will diffuse over sub-micrometer length-scales (< 10�7 m; Fig. 7a),significantly less than typical dimensions of metamorphic mica. Fig. 7b shows the effect of convergence velocity oncumulative diffusion length-scales; P–T path parameters are otherwise the same as path 2 in Fig. 7a. These plots show that, independent of convergence velocity, diffusive removal of argon from mica equilibrating during prograde metamorphism along typical subduction-related temperature gradients will be inefficient. On reaching peak ðUÞHPconditions mica that has equilibrated continuously during subduction will retain a memory of its protolith, or prograde history. This suggests that release of protolith-derived argon from mica, and other K-bearing phases, is controlled bynon-diffusive processes driven by gradients in chemical potential, such as dissolution–precipitation mechanisms (Wijb-rans and McDougall, 1986; Putnis, 2009 ).

30

25

20

15

10

5

0

)rabk(erusser P

6005004003002001000

Temperature (°C)

LD=10-7m

1

2

3

10-8

10-9

10-10

10-11

10-12

10-13

10-14

10-15

10-16

Temperature (°C)0 001 002 003 004 005 006 700

10-24

10-22

10-20

10-18

10-16

10-14

10-12

10-10

10-8

10-6

10-4

10-2

100

10 cm.yr-1

1 cm.yr-1

)m(

htgnelnoisu ffid

e vit alumu

C

a b

)m(

htgnelnois uffi

D

Residence time (Ma)0 5 10 15 20 25 30

10-14

10-12

10-10

10-8

10-6

10-4

10-2

100

100°C

200°C

300°C

400°C

500°C

600°C

700°C800°C900°C

1000°C

Grain-scale

c

Grain-scale

Fig. 7. Argon diffusion calculations. (a) Calculated steady-state P–T paths followed by the top of a subduc ting slab using the analytical solutions of Molnar and England (1990) and Peacock (1993). The following parameters were used for each path: convergen ce velocity (V) = 8 cm yr�1, subduction angle (hsub) = 15�, basal heat flow (qm) = 0.05 W m�2, crustal density (qc) = 3000 kg m�3, thermal conductivity (K) and diffusivity (j) of 2.5 W m�1 C�1 and 1 �10�6 m2 s�1, respective ly. Temperatur e gradients are controlle d by the amount of shear heating along the slab–wedge interface (Molnar and England , 1990; Peacock, 1993 ). Path 1 is calculated using a shear stress (s) of 30 MPa,path 2 s = 45 MPa and path 3 s = 60 MPa. P–T paths are contoured for the cumulativ e diffusion distance (LD ¼ 2

ffiffiffiffiffi

Dtp

, where D = D0:e�ðEaþV 0ðP�P 0Þ=R:T Þ) of argon in muscovite, using the experimenta l diffusion param eters of Harrison et al. (2009):Ea = 263 � 29 kJ mol �1; D0 = 2:3þ70

�2:2 cm2 s�1; V 0 = 14 cm3 mol�1 (P 0 = 1 GPa). (b) Effect of convergence rate on cumulativ e diffusion length during subduct ion. Shaded rectangle highlights typical phengite grain dimensions . (c) Effect of ðUÞHP residenc e time on diffusion length for the T range: 100–1000 �C. Shaded rectangle highlights dimensions of typical phengite grains.


At temperatures greater than �400 �C length scales ofargon diffusion become spatially significant. Fig. 7c is a plot of diffusion length as a function of residence time at peak T,between 0 and 30 Ma (a plausible range for ðUÞHP terranes;Kylander-Clark et al., 2012; Agard et al., 2009 ) for a range of geologically realistic temperatures. Millimeter-scale argon diffusion is shown to occur only at temperatures J 500 �C, independent of ðUÞHP residence time. These results are consistent with Warren et al. (2012) who calculated that 0.1 and 1 mm diameter muscovite grains residing at25 kbar and 550 �C for 10 Ma will retain <5% and �65%radiogenic argon, respectively. They also showed that a0.5 mm muscovite grain residing at the same P–T condi-tions for 0.5 and 20 Ma, respectively, will retain �95%and �25% radiogenic argon. Provided that the macroscopic grain radius represents the effective diffusion length scale,this modeling shows that excess age fractions will be consid- erable in grains J 1 mm in diameter from ðUÞHP terranescharacterized by residence times <10 Ma, even with aneffective open system grain boundary network. However,previous work suggests that diffusion of argon in muscovite is non-Fickian and most accurately simulated by the presence of multiple diffusion domains of differing length scales but identical activation energies (Lovera et al., 1991; Harri- son et al., 2009 ). Structural deformation is also expected toenhance bulk grain argon diffusivity as a result of concomitant reductions in grain size and effective diffusion length scale (see Baxter, 2010 for a review). Both multi-domain diffusion and deformation of muscovite will lead to enhanced loss of 40Ar and younger cooling ages, increasing the concentrations of Are required to explain measured 40Ar/39Ar ages.

In the absence of a free-fluid phase, inter-crystalline diffusion (� 103 times faster than intra-crystalline diffusion;Manning, 1974 ) will control the rate of inter-granular argon exchange. The presence of an inter-granular fluid-phase will increase the efficiency of argon transport by a combination of advective solute transport, operating between 10�3 and10�1 m a�1, and aqueous diffusion, up to 16 orders of magnitude faster than solid-state diffusion (Rubie, 1986 ). There- fore, effective removal of inherited or locally-generated Are

from a mica-bearing rock will critically depend on the availability of fluids.

6.3. The availability of a grain-boundary fluid phase

Subducting sediments and oceanic crust contain free H2O residing in pores and lattice-bound H2O in hydrous phases. Pore fluids, occupying K 50 vol.% in subducting sediment and K 7 vol.% (Becker and Sakai, 1989 ) in the upper oceanic crust, are expelled by porosity collapse atdepths K 20 km and T K 150–300 �C, depending on the thermal structure of the subduction zone. Despite the large fluid fluxes caused by porosity collapse in subducting sediments (0.1–1.2 �10�3 m3 m�2 a�1 Hyndman and Peacock,2003), the lack of pervasive mineralogical change, with the exception of the smectite ) illite transition, means that release of inherited 40Ar by dissolution is likely to benegligible.

Release of lattice-bound H2O from hydrous phases iscontrolled by dehydration reactions. In order to illustrate the different dehydration histories for subducting maficand pelitic lithologies, P–T pseudosections (THERMOCALC

version 3.33; Holland and Powell, 1998 ) were calculated

+5%

-65%

-65%

-15%

-10%

-5% +10%

+35%+30%+25%+20%

-55%

-50%

-45%

-40%

-35%

-30%

-25%

400 450 500 550 600400 450 500 550 600

hb ep bi pl sphact chl epbi ab sph

gl act chl ep bi ab sph

gl act chl ep bi spho gl chl ep mu sph

law o gl act chl mu ru

law o gl ta act chl mu ru

a NCKFMASHTO(+SiO2)

o g ta law mu ru

o gl g ta law mu ru

law o act chl mu ru

o gl chl ep mu ru

o gl act chl ep mu ru

gl act chl ep mu ru

+act

gl act chlep bi ru hb ch

l ep bi ru

hb ep bi ru

hb o ep bi ru

hb ep bi ru sph

hb ep bi ab sph

g glact chlep mu ru

hb a

ct c

hlbi

ab

sph

ep

SiO 2 absent

coeqtz

b

5

10

15

20

25

30

Pres

sure

(kba

r)

-20%

-15% 0%

400 450 500 550 600Temperature (°C)

+5%

+10%

+15%+20%

+25%+30% +35% +40%

+60%

-10% -5

%

g ep car jd

MnNCKFMASHO(+mu+SiO2)

c g ky jdg kycar jd

g ctd ep car jd

g ep c

ar ctd

g ky

gctdky

g ctd ep

g ctd chl epctd chl ep

ctd chl ep pa

g ctd ep pa chl car

g chl ma ep

g ctd chl ma epg chl ma

g ky chl ma

g ky chl

g chl m

a bi

g ma bi st

g ky b

i ma

kysill

g ky car

400 450 500 550 600Temperature (°C)

5

10

15

20

25

30

Pres

sure

(kba

r)

dcoeqtz

Fig. 8. P–T pseudosecti ons calculated for mafic and pelitic composit ions between 400–650 �C and 5–30 kbar. Panel (a) shows stable phase equilibria for an oxidised MORB bulk compo sition: 2.66 Na2O; 12.43 CaO; 0.23 K2O; 12.93 MgO; 9.26 Al2O3; 53.4 SiO 2; 1.07 TiO 2, 1.5 O,(Rebay et al., 2010 ). The system is H2O saturated. Panel (b) is contoured for amounts of mineral- bound H2O released/ consumed relative to4 kbar and 400 �C. Contours are labelled at 5% intervals and calculated for assemblag es stable on the low- T side of the amphibole solvus.Panel (c) shows a P–T pseudosec tion constructed for a peraluminous pelite (TH-680; Smye et al., 2011 ), using the fully-exp anded MnNCKFM ASHO chemica l system. Bulk composit ion in molar %: 0.09 MnO; 0.24 Na2O; 1.29 CaO; 3.33 K2O; 5.55 FeO; 3.64 MgO; 13.76 Al2O3; 70.31 SiO 2; 0.185 O. Both quartz and H2O are set to excess. Panel (d) shows contour s of constant H2O flux relative to 4 kbar and 400 �C. Contours are labelled at 5% intervals .


between 400 and 600 �C and 5–30 kbar for an oxidised MORB (Fig. 8a and b; NCKFMASHTO) and a peraluminous pelite (Fig. 8c and d; MnNCKFMASHO) composition. The pseudosections are contoured for modal proportions of mineral-bound H2O liberated (þDH2O) orconsumed (�DH2O), expressed as a percentage relative tothe amount of free H2O present in respective greenschist-facies assemblages stable at 4 kbar and 400 �C. Fig. 8a and bare curves corresponding to the flux in lattice-bound H2Oproduced by each of the P–T paths calculated in Fig. 7a,for MORB and pelitic compositions respectively. Phase relations considered here are restricted to temperatures >400 �C due to uncertainty over initial water contents ofthe respective protoliths.

In a rock of MORB composition, dehydration occurs dominantly by discontinuous, univariant reactions. Specif- ically, the stability of lawsonite exerts a strong control onthe amount of free H2O present during subduction. The decomposition of lawsonite (red band, Fig. 8b) produces �50% change in relative free H2O content. A rock following P–T path 1 (Fig. 7a) will encounter SiO 2 undersatu-rated conditions at pressures >20 kbar. P–T path 2 has a sub-parallel trajectory to the lawsonite-out band in P–T space (Fig. 9a); lawsonite decreases in modal abundance during the low-temperature portion (<450 �C) of the path,whereas after the path crosses the sphene ) rutile transition, lawsonite modes increase, requiring the addition ofH2O. Oceanic crust following the high- s P–T path (path


3; Fig. 7a) will dehydrate continuously between 400 and 600 �C, liberating �30% H2O relative to greenschist satu- ration values though the decomposition of actinolite and chlorite. Importantly, path 3 does not facilitate lawsonite growth.

During subduction, pelitic rocks lose up to 65%(Figs. 8a, b and 9b) of the mineral-bound H2O stable atgreenschist facies conditions, 25% more than the maximum DH2O yielded by oceanic crust. Devolatisation in the pelitic system is characterised by divariant, continuous reactions that release H2O throughout the P–T space between 400 and 600 �C and 5–30 kbar. Path 1 is dominated by the modal abundance of the hydrous phase carpholite, which, like lawsonite, requires hydration during subduction. Paths 2and 3 liberate similar quantities of H2O (DH2O ¼ 56 and 52% respectively) through the successive breakdown ofchlorite, epidote and chloritoid. Phengite is stable throughout the P–T space considered.

-60

-40

-20

0

20

40

600550500450400

60

50

40

30

20

10

0

-10-20

600550500450400

Temperature (˚C)

Temperature (˚C)

Δ H

2)

%( O

a

b

Δ H

2)

%( O

Path 1

Path 3

Pelite

[ctd]

[chl,ctd,ep]

Path 2[ctd]

[chl,ep]

Path 1

Path 3

Path 2

[sph↔ru]

[act]

[act,chl]law-out

law-in

MORB

Fig. 9. Calculate d flux in mineral-bou nd H2O for a MORB (panela) and pelite (panel b) undergoin g subduction -related devolati za- tion between 400 �C and 600 �C. Bulk composit ions are the identical to those used to calculate pseudosec tions shown inFig.8. Curves correspon d to the amount of mineral-bou nd H2Oliberated (þDH2O), or consumed (�DH2O), expressed as a percenta gerelative to the modal proportio n of free H2O present in respective greenschist- facies assemblag es stable at 4 kbar and 400 �C. Fluxes are plotted as a function of T for the subduction P–T paths 1(brown), 2 (green) and 3 (blue), described in Fig. 7. Square brackets represent the position of discontinuou s reactions; the labelled phase correspon ds to the phase that is consumed . Dashed line in panel adenotes silica-undersa turated conditions.

The order-of-magnitude larger porosity estimates calculated from mica in pelitic protoliths (Figs. 4b and 5) relative to those calculated from mafic eclogites correlates with the greater extent of devolatisation that is predicted to occur during subduction to ðUÞHP conditions. The continuous supply of lattice-bound H2O to the grain-boundary network in pelitic rocks will cause elevated hydrostatic pressures, driving crack propagation and advective fluidtransport of protolith-derived argon. In contrast, oceanic crust has the potential to undergo hydration during subduction due to lawsonite growth and its ability to incorporate an entire H2O molecule, effectively buffering H2O activity in the mafic system. Under such conditions, mafic eclogites may remain dry until the blueschist ) eclogite transition where glaucophane, garnet, and talc are stabilized at the ex- pense of actinolite and chlorite. The buffering of free H2Oby lawsonite has the potential to limit available grain boundary fluid to a layer only nanometers in thickness, thus preventing effective dilution and removal of internally- cycled argon from the system. Under dry conditions, as predicted to occur in mafic rocks, local matrix phases could potentially dominate the total local sink capacity for argon (Fig. 6a), whereas the elevated fluid fluxes expected inmetasediments suggest that pore fluids will be the dominant repository for Are.

6.4. The connectivity of the grain-boundary fluid phase

Mafic rocks release smaller quantities of structurally- bound H2O than metapelitic rocks during subduction.However, removal of locally-generated 40Ar in a subducting rock will not only depend on the volume of available fluid,but critically, the connectivity of the fluid phase. Permeabil- ity is dependent on rock rheology and exists as both fracture networks, observable over micro- to meter lengthscales, and texturally-controlled inter-granular pathways in metamorphic rocks (Ingebritsen and Manning,2010). In order to maintain porosities similar to those calculated above (10�3–10�5) over the duration of mica equilibration/growth under ðUÞHP conditions, a fraction of the metamorphic fluid liberated during prograde metamorphism must remain in the rock. Thermal-petrologic models predict that integrated fluid fluxes from subducting maficoceanic crust and associated sediments are between � 10�4 and 10�6 m3 m�2 a�1 at depths between 20 and 80 km (Hyndman and Peacock, 2003 ). Using this range ofvalues, which account for the differences in thermal structure observed in subduction zones, a 5 km-thick slab undergoing dehydration will produce enough fluid to fill or form the calculated porosities over timescales of 500 to 5,000,000 years, respectively. Adopting Manning and Ingebritsen (1999)’s formulation for the conditions under which fluidpressures become hydrostatic, slab permeabilities K 10�20

and K 10�22 m2 are required to retain metamorphic fluidsreleased at these production rates. These values are compa- rable to estimates of background crustal permeability:� 10�20 m2 (Connolly, 2010 ).

The hydraulic connectivity of a rock undergoing metamorphism is controlled by a combination of pervasive flowthrough interconnected pore spaces and channelized flow


through fractures. An interconnected porosity network iscontrolled by the dihedral angle (h) between fluid–min-eral–mineral junctions (Graham et al., 1997 ). A fundamen- tal change in fluid mobility occurs at h = 60 � (Smith, 1948 ).At values of h <60�, the surface energy of a wetted grain- edge region is smaller than a dry grain-edge region and the system favors open grain-edge channels, independent of the value of /. For h >60�, dry grain-edge surface energies are smaller than the wetted grain-edge configurationand grain-edge channels pinch-off along their length, forming a closed grain-boundary reservoir. The quartz–fluid hwill dominate texturally-defined permeability in metapelites. Holness (1993) and Holness and Graham (1995)showed that porosity in quartz-rich rocks will be connected within two permeability windows to C–O–H fluids atP <11 kbar: one at temperatures <400 �C and the other at temperatures >1100 �C. The topology of P–T–h con-tours constructed by Holness (1993) suggests a further permeability window at pressures above �11 kbar.Connectivity in the mafic system is best represented bythe forsterite–fluid h. Watson et al. (1991) and Mibe et al.(1999) showed that h decreases with increasing pressure and temperature between 0.5–5 GPa and �400–1300 �C;the critical h of 60 �C is attained at �2 GPa at 1000 �C.

Whereas pervasive flow through an interconnected porosity network reflects thermodynamic equilibrium, fractures are nonequilibirum phenomena, capable of connect- ing pores that would otherwise be isolated. Rocks fracture in response to externally-applied or internally-generated (i.e. devolatization) stresses exceeding the tensile strength limit (Etheridge, 1983 ). The low tensile strength of rocks, K 5 MPa, means that hydrofracturing provides a self-regulating mechanism to control pore fluid pressures (Morris and Henley, 1976 ). This is supported by recogni- tion of petrographic evidence for multiple cracking and sealing events in the history of a single vein (Ramsay,1980). Fracturing also occurs in response to deviatoric stresses under conditions of P fluid < P rock (Yardley, 1983 ),forming boudinage veins and shear fractures, commonly larger than tensile fractures e.g. Philippot and Selverstone (1991). Both tensile and shear fractures are common inðUÞHP terranes and often contain ðUÞHP phases, highlight- ing the importance of channelized fluid flow as an effectivemechanism for solute transfer during subduction (e.g.Franz et al., 2001; Widmer and Thompson, 2001; Zhang et al., 2008; John et al., 2012 ). Given that garnet (Ji and Martignole, 1994 ) and clinopyroxene (Kolle ´ and Blacic,1983) are two of the strongest minerals in the crustal lithosphere, eclogites have higher ultimate tensile and shear strengths (Hacker, 1996; Jin et al., 2001 ) compared to cogenetic pelites in which quartz controls effective rock strength (Hoek and Brown, 1997 ). Coupled with the larger volumes of fluid released from metasediments, this shows that eclogites are less susceptible to fracture formation during subduction. Furthermore, metapelites typically comprise 10–70 volume % of sheet silicates, compared to K 2% in eclogites, promoting the formation of planar foliation fabrics that act as fluid percolation pathways.

Collectively, these data suggest that metapelites are expected to be more permeable to fluids released during pro-

grade metamorphism than cogenetic mafic rocks.Accordingly, pelites will lose protolith-derived argon byadvective solute transport and have true porosities < 10�4

volume fraction at the time of mica equilibration.

6.5. A conceptual model for the accumulation of Are in

metamorphic rocks

Fig. 10 shows the competing effects of available fluid,permeability and argon diffusion efficiency on rocks ofpelitic and mafic compositions during increasing temperature (and implied pressure). Note that the dependence ofhydraulic connectivity on dihedral angle assumes that fracture permeability is negligible. The model is equally applicable to the scenario where deformation drives hydraulic connectivity. During the early stages of burial pores are connected in both mafic and pelitic lithologies (stage Oi

Fig. 10a–c) and pore fluids are expelled by compaction.As temperatures are < T c, volume diffusion is ineffective(LD < 10�10 m; Fig. 7) and argon exchange occurs bygrain-boundary dissolution–precipitation, concomitant with the growth and consumption of new and relict metamorphic phases respectively. As H2O is released into a connected porosity network, inherited 40Ar is effectively flushedout of the system to an external sink, ultimately residing inthe atmosphere. Pelites retain a connected porosity and lose large amounts of H2O over the temperature interval atwhich diffusion in white mica becomes efficient over grain-scale dimensions (stage Oii Fig. 10a and c; see also Fig. 7), thereby facilitating open system 40Ar transport toan external sink before the rock becomes impermeable,either when h < 60� or the dehydration fluid flux decreases.Conversely, mafic rocks hydrate or liberate small amounts of H2O during prograde metamorphism, becoming closed to advective solute transport at values of T < T c (stagesOi and Ci Fig. 10b and c). Accordingly, the volume fraction of free fluid in mafic rocks will be significantly less than cogenetic pelites at values of T > T c (stage Cii Fig. 10band c). On exhumation and associated cooling through T c, diffusion will become inefficient and, in the absence ofa high- T fluid phase (Camacho et al., 2005 ), the 40Ar concentration acquired by the mica under ðUÞHP conditionsis effectively ‘locked’ into the grain.

6.6. Implications for the transport of volatiles to the mantle

Recent work has shown that the isotopic composition ofheavy noble gases and halogens in the mantle wedge and the convecting mantle is remarkably similar to that of seawater (Holland and Ballentine, 2006; Sumino et al., 2010 ).This challenges the popular concept of a noble gas ‘subduction barrier’ (Staudacher and Alle `gre, 1988 ) in which volatiles, with the exception of S and CO2, are returned to the surface as pore-fluids are expelled from oceanic crust atdepths <20 km (Hyndman and Peacock, 2003; Jarrard,2003). The retention of locally-derived 40Ar to mantle depths in mafic eclogites provides further compelling evidence that the oceanic crust behaves as a closed system tovolatiles during subduction. Marine pore fluids could plausibly become encapsulated during the early stages of burial

H2O

rele

ased

(

)

( )

Tc

θ>60°

Temperature

H2O

rele

ased

(

)

θ

Tc

θ>60°

θ

a: pelite b: mafic

40Ar

40Ar

40Ar

40Ar

40Ar40Ar

40Ar

40Ar

40Ar

40Ar

Temperature

c

( )

40Ar

40Ar

40Ar40Ar

40Ar

Oi Oii Ci Cii

CiiOii

OiOi

Ci Cii

Release of 40Ar by dissolutioninto an open system

Release of 40Ar by diffusioninto an open system

Release of 40Ar by dissolutioninto a closed system

Diffusive equilibration of 40Ar in a closed system

Fig. 10. Conceptual model for the accumulatio n of Are under ðUÞHP conditions. Panels a and b show the relation ships betwee n the amount of H2O released during prograde metam orphism (black solid line), argon diffusivity (the boundary between efficient and inefficient diffusion,T c is marked schematically as a region bounded by grey dashed lines) and dihedral angle (h, black dashed line) with increasin g T during subduction for pelitic and mafic protoliths respectively. The topology of the schematic h and H2O curves are discussed in the text. The shaded region represent s values of h > 60�, the value at which the system becomes impermeab le to grain boundary flow at low porosities. Panel cshows the four possible configurations of argon transport that a mica-bear ing rock will experience during subduction to ðUÞHP conditions;these four stages are marked on panels a and b. Solid arrows represent dissolution exchang e wherea s dashed arrows represent diffusionalexchange. (Oi) Early-stag e flushing of protolith-deriv ed argon controlle d by mineral dissolution and advective solute transport in a connected grain-bou ndary pore network. (Oii) “Dodson” open system: as T exceeds the range of T c for argon in mica, diffusion becomes efficient at the grain, leading to removal of lattice-bound 40Ar into a connected (open, infinite) grain boundary reservoir. (Ci) Release of 40Ar by dissolution reactions into a closed system at T < T c. (Cii) Closed system: whilst the T conditions still allow efficient diffusion, the rock is no-longer permeable and each individua l mica grain equilibrates with its local immobile fluid argon concentratio n.


(h > 60� and T < T c; Fig. 10). During subduction and concomitant heating, pore fluids will participate in dissolution–precipitation reactions leading to mineral hydration or formation of fluid inclusions. The fact that phengites crystalliz- ing at ðUÞHP conditions incorporate protolith-derived 40Arshows that the grain boundary reservoir behaves as a closed ‘sink’ for incompatible elements, such as noble gases,throughout subduction. Release of volatiles into the mantle reservoir could plausibly be facilitated by increased dehydration fluid fluxes driven by lawsonite breakdown, orwhen h <60� (80–150 km depth; Mibe et al., 1999 ).

7. CONCLUSIONS

Extraneous argon is ubiquitous in phengite that has equilibrated under ðUÞHP conditions. Apparent phengite 40Ar/39Ar ages K 700% older than the age of ðUÞHP meta-morphism can be explained by incorporation of locally-generated 40Ar. Measured Are concentrations in mafic eclogites are reproduced only when porosities are K 10�4 volumefraction, showing that mafic protoliths operate as closed systems to advective solute transport during subduction.Porosities in eclogite-facies metapelites are K 10�2, reflect-ing loss of significant volumes of lattice-bound H2O relative

to mafic rocks during subduction. The concentration of Are

in ðUÞHP phengite strongly reflects the devolatisation history of the protolith. Phase equilibria calculations show that pelitic lithologies dewater continuously throughout subduction to ðUÞHP conditions, flushing accumulated 40Ar from the system. Mafic lithologies either require hydration, or produce less H2O during subduction, therefore retaining protolith-derived radiogenic argon under closed system conditions. Such small porosities are main- tained by permeabilities K 10�20 m2 during subduction.Retention of protolith-derived 40Ar to ðUÞHP conditionsimplies that the oceanic crust is an efficient vehicle for the transport of volatiles to the mantle.

ACKNO WLEDGME NTS

A.J.S. acknowledg es support of a Jackson Postdoct oral Fellow- ship. C.J.W. acknowledg es funding from the Natural Environ ment Research Council (NE/E0114038/1 and NE/H016279 /1). We are grateful to David Shuster for his editoria l handling of the manuscript. Formal reviews by Jan Wijbrans , Ethan Baxter and ananonymous reviewer served to greatly improve the clarity of the manuscript. This work has benefitted from fruitful discussion with Simon Kelley, Chris Hawkesw orth and Marian Holness.


REFERENCE S

Agard P., Yamato P., Jolivet L. and Burov E. (2009) Exhumation of oceanic blueschists and eclogites in subduction zones: timing and mechan isms. Earth Sci. Rev. 92, 53–79.

Ames L., Tilton G. and Zhou G. (1993) Timing of collision of the Sino-Ko rean and Yangtse cratons: U–Pb zircon dating ofcoesite-be aring eclogites. Geology 21, 339.

Banner J. and Hanson G. (1990) Calculatio n of simultan eous isotopic and trace element variation s during water–rock interaction with applicatio ns to carbonate diagenesis. Geochim.

Cosmochi m. Acta 54, 3123–3137.Baxter E. (2003) Natural constraints on metamorph ic reaction

rates. Geol. Soc., Lond. Spec. Publ. 220, 183–202.Baxter E. (2003) Quantification of the factors controllin g the

presence of excess 40Ar or 4He. Earth Planet. Sci. Lett. 216,619–634.

Baxter E., DePaolo D. and Renne P. (2002) Spatially correlated anomalous 40Ar/39Ar age variations in biotites about a litho- logic contact near Simplon Pass, Switzerland : a mechanist icexplanatio n for excess Ar. Geochim. Cosmochim. Acta 66, 1067–1083.

Baxter E. F. (2010) Diffusion of noble gases in minerals. Rev.

Mineral. Geochem. 72, 509–557.Becker K. and Sakai H. (1989) Drilling deep into young oceanic

crust, Hole 504B, Costa Rica Rift. Rev. Geophys. 27, 79–102.Bickle M. and McKenzi e D. (1987) The transport of heat and

matter by fluids during metam orphism. Contrib. Mineral.

Petrol. 95, 384–392.Bousquet R., Goffe B., Vidal O., Oberha ¨nsli R. and Patriat M.

(2002) The tectono-m etamorphic history of the Valaisan domain from the Western to the Central Alps: new constraints on the evolution of the Alps. Geol. Soc. Am. Bull. 114, 207.

Brewer M. (1969) Excess radiogen ic argon in metamorph ic micas from the eastern Alps, Austria. Earth Planet. Sci. Lett. 6, 321–331.

Broadhurst C., Drake M., Hagee B. and Bernatowicz T. (1992)Solubility and partitioning of Ne, Ar, Kr and Xe in minerals and synthetic basaltic melts. Geochim . Cosmochi m. Acta 56,709–723.

Brooker R., Wartho J., Carroll M., Kelley S. and Draper D. (1998)Preliminary UVLAMP determin ations of argon partition coefficients for olivine and clinopyroxen e grown from silicate melts. Chem. Geol. 147, 185–200.

Camacho A., Lee J., Hensen B. and Braun J. (2005) Short-lived orogenic cycles and the eclogitizatio n of cold crust byspasmodic hot fluids. Nature 435, 1191–1196.

Chopra P. and Paterson M. (1984) The role of water in the deformatio n of dunite. J. Geophys. Res. 89, 7861–7876.

Connolly J. (2010) The mechanics of metamorp hic fluid expulsion.Elements 6, 165–172.

Crovetto R., Ferna ńdez-Prini R. and Japas M. (1982) Solubilities of inert gases and methane in HO and in DO in the temperatu rerange of 300 to 600 K. J. Chem. Phys. 76, 1077.

Cumbest R., Johnson E. and Onstott T. (1994) Argon composition of metamorph ic fluids: implications for 40Ar/39Ar geochronology. Geol. Soc. Am. Bull. 106.

Dalrymple G. and Lanphere M. (1969) Potassium –argon Dating:

Principles, Techniques and Application s to Geochronolo gy.Freeman, San Francisco, California, p. 258 .

Damon P. (1968) Applicatio n of the potassium–argon method tothe dating of igneous and metamorph ic rocks within the Basin Ranges of the southwest. Arizona Geol. Soc., Southern Arizona

Guideboo k 3, 7–20.De Jong K., Feraud G., Ruffet G., Amouric M. and Wijbrans J.

(2001) Excess argon incorporation in phengite of the Mulhace ń

complex: submicrosc opic illitization and fluid ingress during late Miocene extension in the Betic Zone, south-easte rn Spain.Chem. Geol. 178, 159–195.

Deer W., Howie R. and Zussman J. (1997) Rock-form ing minerals.Geol. Soc. 4.

DePaolo D. and Getty S. (1996) Models of isotopic exchange inreactive fluid–rock systems: implications for geochronolo gy inmetamorph ic rocks. Geochim. Cosmochim. Acta 60, 3933–3947.

Di Vincenzo G., Tonarin i S., Lombardo B., Castelli D. and Ottolini L. (2006) Comparison of 40Ar/39Ar and Rb–Sr data onphengites from the UHP Brossaco -Isasca unit (Dora Maira Massif, Italy): implications for dating white mica. J. Petrol. 47,1439–1465.

Dodson M. (1973) Closure temperat ure in cooling geochronolo g-ical and petrologic al systems. Contrib. Mineral. Petrol. 40, 259–274.

El-Shazly A. (2001) Are pressures for bluesch ists and eclogites overestimat ed? Lithos 56, 231–264.

El-Shazly A., Bro ¨cker M., Hacker B. and Calvert A. (2001)Formation and exhumation of blueschists and eclogites from NE Oman: new perspecti ves from Rb–Sr and 40Ar/39Ar dating.J. Metamorph . Geol. 19, 233–248.

Etheridge M. (1983) Differential stress magnitud es during regional deformation and metam orphism: upper bound imposed bytensile fracturing. Geology 11, 231–234.

Etheridge M., Wall V. and Vernon R. (1983) The role of the fluidphase during regional metamorph ism and deformatio n. J.

Metamorph . Geol. 1, 205–226.Foland K. (1979) Limited mobility of argon in a Metamorphic

Terrain. Geochim. Cosmochi m. Acta 43, 793–801.Franz L., Romer R., Klemd R., Schmid R., Oberhan sli R., Wagner

T. and Shuwen D. (2001) Eclogite-fac ies quartz veins within metabasites of the Dabie Shan (eastern China): pressure–temperature–time-deform ation path, composit ion of the fluidphase and fluid flow during exhumation of high-press ure rocks.Contrib. Mineral. Petrol. 141, 322–346.

Ganor J., Matthews A. and Paldor N. (1989) Constraint s oneffective diffusivity during oxygen isotope exchange at a marble- schist contact , Sifnos (Cyclades), Greece. Earth Planet. Sci.

Lett. 94, 208–216.Getty S. and Selverston e J. (1994) Stable isotopi c and trace element

evidence for restricted fluid migratio n in 2 GPa eclogites. J.

Metamorph . Geol. 12, 747–760.Giorgis D., Cosca M. and Li S. (2000) Distribu tion and significance

of extraneous argon in UHP eclogite (Sulu terrain, China):insight from in situ 40Ar/39Ar UV-laser ablation analysis. Earth

Planet. Sci. Lett. 181, 605–615.Gleiner P., Glodny J. and Franz G. (2007) Rb–Sr isotopic dating of

pseudomorp hs after lawsonite in metabasalts from the Glock- ner nappe, Tauern Window , Eastern Alps. Eur. J. Mineral. 19,723–734.

Glodny J., Ring U., Kuhn A., Gleissner P. and Franz G. (2005)Crystallizatio n and very rapid exhuma tion of the younges tAlpine eclogites (Tauern Windo w, Eastern Alps) from Rb/Sr mineral assemblag e analysis. Contrib. Mineral. Petrol. 149, 699–712.

Graham C., Skelton A., Bickle M. and Cole C. (1997) Lithological,structural and deformation controls on fluid flow during regional metamorph ism. In Deformatio n-enhanced Fluid Trans-

port in the Earth’s Crust and Mantle (ed. M. Holness). Chapman and Hall, pp. 196–226.

Gray D., Hand M., Mawby J., Armstrong R., Miller J. and Gregory R. (2004) Sm–Nd and zircon U–Pb ages from garnet- bearing eclogites, NE Oman; constrai nts on high-P metamorphism. Earth Planet. Sci. Lett. 222, 407–422.

http://refhub.elsevier.com/S0016-7037(13)00160-9/h0005

















































































































































Hacker B. (1996) Eclogite formation and the rheology , buoyancy,seismicity, and H2O content of oceanic crust. Geophys. Monogr.

Ser. 96, 337–346.Harrison T., Celerier J., Aikman A., Hermann J. and Heizler J.

(2009) Diffusion of 40Ar in muscovite. Geochim . Cosmochim.

Acta 73, 1039–1051.Harrison T. and McDougall I. (1980a) Investig ations of an

intrusive contact, northwest Nelson, New Zealand –I. Thermal,chronolog ical and isotopic constraints. Geochim. Cosmochim.

Acta 44, 1985–2003.Harrison T. and McDouga ll I. (1980b) Investig ations of an

intrusive contact, northwest Nelson, New Zealand –II. Diffusionof radiogen ic and excess 40Ar in hornblende revealed by40Ar39Ar age spectrum analysis. Geochim. Cosmochi m. Acta

44, 2005–2020.Hoek E. and Brown E. (1997) Practical estimates of rock mass

strength. Int. J. Rock Mech. Min. Sci. 34, 1165–1186.Holland G. and Ballentine C. (2006) Seawate r subduction controls

the heavy gas composition of the mantle. Nature 441, 186–190.Holland T. and Powell R. (1998) An internally consistent thermo-

dynamic data set for phases of petrological interest. J.

Metamorph . Geol. 16, 309–344.Holness M. (1993) Temperatur e and pressure dependenc e of

quartz–aqueous fluid dihedral angles: the control of adsorbe dH2O on the permeability of quartzit es. Earth Planet. Sci. Lett.

117, 363–377.Holness M. and Graham C. (1995) P–T–X effects on equilibrium

carbonate–H2O–CO2–NaCl dihedral angles: constraints oncarbonate permeability and the role of deformatio n during fluid infiltration. Contrib . Mineral. Petrol. 119, 301–313.

Hyndman R. and Peacock S. (2003) Serpentiniz ation of the forearc mantle. Earth Planet. Sci. Lett. 212, 417–432.

Ingebritsen S. and Mannin g C. (2010) Permeability of the continental crust: dynamic variations inferred from seismicity and metam orphism. Geofluids 10, 193–205.

Inger S. and Cliff R. (1997) Timing of metamorph ism in the Tauern Window , Eastern Alps: Rb–Sr ages and fabric formatio n. J.

Metamorph . Geol. 12, 695–707.Jarrard R. (2003) Subductio n fluxes of water, carbon dioxide,

chlorine, and potassium. Geochem. Geophys. Geosyst. 4.Ji S. and Martignole J. (1994) Ductility of garnet as an indicator of

extremely high temperature deformatio n. J. Struct. Geol. 16,985–996.

Jin Z., Zhang J., Green H. and Jin S. (2001) Eclogite rheology:implication s for subducted lithosp here. Geology 29, 667–670.

John T., Gussone N., Padladchi kov Y., Bebou t G., Dohmen R.,Halama R., Klemd R., Magna T. and Seitz H. (2012) Volcanic arcs fed by rapid pulsed fluid flow through subduct ing slabs.Nat. Geosci. 5, 489–492.

Kelley S. (2002) Excess argon in K–Ar and Ar–Ar geochronology.Chem. Geol. 188, 1–22.

Kelley S., Turner G., Butterfield A. and Shepherd T. (1986) The source and significance of argon isotopes in fluid inclusions from areas of mineraliz ation. Earth Planet. Sci. Lett. 79, 303–318.

Kelley S. and Wartho J. (2000) Rapid kimberlite ascent and the significance of Ar–Ar ages in xenolith phlogopit es. Science 289,609–611.

Kohlstedt D. (2006) The role of water in high-temp erature rock deformatio n. Rev. Mineral. Geochem. 62, 377–396.

Kolle ´ J. and Blacic J. (1983) Deformatio n of single-cryst alclinopyroxe nes: 2. Dislocation -controlled flow processes inhedenber gite. J. Geophys. Res. 88, 2381–2393.

Kurz W., Neubauer F., Genser J. and Dachs E. (1998) Alpine geodynamic evoluti on of passive and active continent al margin

sequences in the Tauern Window (eastern Alps, Austria, Italy):a review. Geol. Rundsch. 87, 225–242.

Kylande r-Clark A., Hacker B. and Mattinson C. (2012) Size and exhumation rate of ultrahigh-p ressure terranes linked toorgenic stage. Earth Planet. Sci. Lett. 321–322, 112–120.

Li S., Wang S., Chen Y., Liu D., Qiu J., Zhou H. and Zhang Z.(1994) Excess argon in phengite from eclogite: evidence from dating of eclogite minerals by Sm–Nd, Rb–Sr and 40Ar/39Armethods. Chem. Geol. 112, 343–350.

Lovera O. M., Richter F. M. and Harrison T. M. (1991) Diffusiondomains determin ed by 39Ar released during step heating. J.

Geophys. Res. 96, 2057–2069.Mackwe ll S., Kohlsted t D. and Paterson M. (1985) The role of

water in the deformatio n of olivine single crystals. J. Geophys.

Res. 90, 11319–11333.Manning C. and Ingebritse n S. (1999) Permea bility of the

continental crust: the implications of geothermal data and metamorph ic systems. Rev. Geophys. 37, 127–150.

Manning J. (1974) Diffusion kinetics and mechanism s in simple crystals. Geochem. Transport Kinet., 3–13.

Mibe K., Fujii T. and Yasuda A. (1999) Control of the location ofthe volcanic front in island arcs by aqueous fluid connectivit y inthe mantle wedge. Nature 401, 259–262.

Molnar P. and England P. (1990) Tempe ratures, heat flux and frictional stress near major thrust faults. J. Geophys. Res. 95,4833–4856.

Morris R. and Henley R. (1976) Dewatering of a metamorph ic pile.Geology 4, 333–336.

Nabelek P. (1987) General equation s for modeling fluid–rockinteraction using trace elements and isotopes. Geochim. Cos-

mochim. Acta 51, 1765–1769.Norris R. and Henley R. (1976) Dewatering of a metamorph ic pile.

Geology 4, 333–336.Norton D. and Knapp R. (1977) Transport phenomena in

hydrotherm al systems: the nature of porosity. Am. J. Sci.

(United States) 277.Onstott T., Phillips D. and Pringle-Goo dell L. (1991) Laser

microprobe measureme nt of chlorine and argon zonation inbiotite. Chem. Geol. 90, 145–168.

Paterson M. and Kekulawa la K. (1979) The role of water in quartz deformation. Bull. Miner. 102, 92–98.

Peacock S. (1990) Fluid processe s in subduction zones. Science 248,329–337.

Peacock S. (1993) The importance of bluschist ! eclogitedehydration reactions in subducting oceanic crust. Geol. Soc.

Am. Bull. 105, 684–694.Philippot P. and Rumble D. (2000) Fluid–rock interactions during

high-pressure and ultrahig h-pressure metamorph ism. Int. Geol.

Rev. 42, 312–327.Philippot P. and Selverst one J. (1991) Trace-ele ment-rich brines in

eclogitic veins: implications for fluid composition and transport during subduction. Contrib . Mineral. Petrol. 106, 417–430.

Putnis A. (2009) Mineral replacemen t reactions. In Thermodyn am-

ics and Kinetics of Water–Rock Interactio n, vol. 70 (eds. E.Oelkers and J. Schott). Reviews in Mineralogy & Geochem istry,pp. 87–124.

Ramsay J. (1980) The crack-seal mechanism of rock deformation.Nature 284, 135–139.

Rebay G., Powell R. and Diener J. (2010) Calculate d phase equilibria for a MORB composition in a P–T range, 450–650 �C and 18–28 kbar: the stability of eclogite. J. Metamorph .

Geol. 28, 635–645.Roddick J., Cliff R. and Rex D. (1980) The evolution of excess

argon in alpine biotites–40Ar–39Ar analysis. Earth Planet. Sci.

Lett. 48, 185–208.




















































































































































Roselieb K., Blanc P., Buttner H., Jambon A., Rammen see W.,Rosenhaue r M., Vielzeuf D. and Walter H. (1997) Experim en- tal study of argon sorption in quartz: evidence for argon incompati bility. Geochim. Cosmochi m. Acta 61, 533–542.

Roselieb K., Wartho J., Buttner H., Jambon A. and Kelley S.(1999) Solubility and diffusivity of noble gases in syntheti cphlogopite: a UV LAMP investigation. Terra Abstr. 4.

Rubatto D., Gebauer D. and Compa gnoni R. (1999) Dating ofeclogite-f acies zircons: the age of Alpine metam orphism in the Sesia-Lan zo Zone (Western Alps). Earth Planet. Sci. Lett. 167,141–158.

Rubatto D. and Hermann J. (2001) Exhumation as fast assubducti on? Geology 29, 3.

Rubie D. (1986) The catalysis of mineral reactions by water and restrictio ns on the presence of aqueous fluid during metam orphism. Mineral. Mag. 50, 399–415.

Ruffet G., Feraud G., Bale ´vre M. and Kie ńast J. (1995) Plateau ages and excess argon in phengite s: an 40Ar/39Ar laser probe study of Alpine micas (Sesia Zone, Western Alps, northern Italy). Chem. Geol. 121, 327–343.

Rumble D. and Yui T. (1998) The Qinglongs han oxygen and hydrogen isotope anomaly near Donghai in Jiangsu Province,China. Geochim. Cosmochi m. Acta 62, 3307–3321.

Saddiqi O., Michard A., Goffe B., Poupeau G. and Oberhansli R.(2006) Fission-trac k thermochro nology of the Oman mountains continental windows, and current problems of tectonic inter- pretation. Bull. Soc. Geol. Fr. 177.

Scaillet S. (1996) Excess 40Ar transport scale and mechanism inhigh-press ure phengite s: a case study from an eclogitize dmetabasite of the Dora-Ma ira nappe, western Alps. Geochim.

Cosmochi m. Acta 60, 1075–1090.Scaillet S. (1998) K–Ar (40Ar/39Ar) geochronology of ultrahig h

pressure rocks. In When Continents Collide: Geodynamics and

Geochemistry of Ultrahigh-Pre ssure Rocks (eds. B. Hacker and J. Liou). Kluwer, pp. 161–201.

Scaillet S., Feraud G., Ballevre M. and Amouric M. (1992) MgFe and [(Mg, Fe) Si–Al2=] composition al control on argon behav- iour in high-press ure white micas: a 40Ar39Ar continuou s laserprobe study from the Dora-Mai ra nappe of the internal western Alps, Italy. Geochim . Cosmochim. Acta 56, 2851–2872.

Schmidt M. and Poli S. (1998) Experim entally based water budgets for dehydrating slabs and consequences for arc magma gener- ation. Earth Planet. Sci. Lett. 163, 361–379.

Sherlock S. and Kelley S. (2002) Excess argon evolution in HP-LT rocks: a UVLAMP study of phengite and K-free minerals, NWTurkey. Chem. Geol. 182, 619–636.

Sherlock S., Kelley S., Inger S., Harris N. and Okay A. (1999)40Ar–39Ar and Rb–Sr geochrono logy of high-press ure metamorphism and exhumation history of the Tavsanli Zone, NWTurkey. Contrib. Mineral. Petrol. 137, 46–58.

Skelton A. (2011) Metamorphic carbon fluxes: how much and how fast? Geology 39, 95–96.

Skelton A., Valley J., Graham C., Bickle M. and Fallick A. (2000)The correlation of reaction and isotope fronts and the mechanism of metamorph ic fluid flow. Contrib . Mineral. Petrol. 138,364–375.

Smith C. (1948) Grains, phases and interfaces: an interpr etation ofmicrostruc ture. Am. Inst. Min., Metall. Petrol. Eng. Trans. 175,15–51.

Smith S. and Kennedy B. (1983) The solubility of noble gases inwater and in NaCl brine. Geochim . Cosmochim. Acta 47, 503–515.

Smye A., Bickle M., Holland T., Parrish R. and Condon D. (2011)Rapid formation and exhuma tion of the youngest Alpine eclogites: a thermal conundrum to Barrovian metamorph ism.Earth Planet. Sci. Lett. 306, 193–204.

Staudacher T. and Alle `gre C. (1988) Recycling of oceanic crust and sediments: the noble gas subducti on barrier. Earth Planet. Sci.

Lett. 89, 173–183.Stuart F., Burnard P., Taylor R. and Turner G. (1995) Resolving

mantle and crustal contributions to ancient hydrothe rmal fluids: He Ar isotopes in fluid inclusions from Dae Hwa WMo mineralisatio n, South Korea. Geochim. Cosmochim. Acta

59, 4663–4673.Sumino H., Burgess R., Mizukami T., Wallis S., Holland G. and

Ballentine C. (2010) Seawate r-derived noble gases and halogens preserved in exhumed mantle wedge peridotite. Earth Planet.

Sci. Lett. 294.Vance D., Ayers M., Kelley S. and Harris N. (1998a) The thermal

response of a metamorph ic belt to extension: constraints from laser Ar data on metamorp hic micas. Earth Planet. Sci. Lett.

162, 152–164.Vance D., Ayres M., Kelley S. and Harris N. (1998b) The thermal

response of a metamorph ic belt to extension: constraints from laser Ar data on metamorp hic micas. Earth Planet. Sci. Lett.

162, 153–164.Walther J. and Wood B. (1984) Rate and mechani sm in prograde

metamorph ism. Contrib. Mineral. Petrol. 88, 246–259.Warren C., Hanke F. and Kelley S. (2012) When can muscovite

40Ar/39Ar dating constrain the timing of metamorph ic exhumation? Chem. Geol. 291, 79–86.

Warren C., Parrish R., Searle M. and Waters D. (2003) Dating the subduction of the Arabian continent al margin beneath the Semail ophiolite, Oman. Geology 31, 889–892.

Warren C., Parrish R., Waters D. and Searle M. (2005) Dating the geologic history of Oman’s Semail Ophiolite; insights from U/Pb geochronolo gy. Contrib. Mineral. Petrol. 150, 403–422.

Warren C., Sherlock S. and Kelley S. (2010) Interpr eting high- pressure phengite 40Ar/39Ar laserprobe ages: an example from Saih Hatat, NE Oman. Contrib. Mineral. Petrol. 161, 991–1009.

Warren C., Smye A., Kelley S. and Sherlock S. (2011) Using white mica 40Ar/39Ar data as a tracer for fluid flow and permeability under high-P conditions: Tauern Window, Eastern Alps. J.

Metamorph . Geol . http://dx.d oi.org/ 10.1111/j.152 5-1314.2011.00 956.x .

Warren C. and Waters D. (2006) Oxidised eclogites and garnet blueschists from Oman; P–T path modellin g in the NCFMA- SHO system. J. Metamorph . Geol. 24, 783–802.

Wartho J., Kelley S., Brooker R., Carroll M., Villa I. and Lee M.(1999) Direct measurement of Ar diffusion profiles in a gem- quality Madagas car K-feldsp ar using the ultra-violet laser ablation microprob e (UVLAMP). Earth Planet. Sci. Lett. 170,141–153.

Watson E., Brenan J. and Baker D. (1991) Distribut ion of fluids inthe continental mantle . In Continental Mantle (ed. M. Menzies).Oxford University Press, pp. 111–125.

Watson E. and Cherniak D. (2003) Lattice diffusion of Ar inquartz, with constrai nts on Ar solubility and evidence ofnanopores. Geochim . Cosmochi m. Acta 67, 2043–2062.

Weiss R. (1970) The solubility of nitrogen, oxygen and argon inwater and seawater. In Deep Sea Research and Oceanographic

Abstracts. Elsevier, pp. 721–735.Wheeler J. (1996) DIFFARG: a program for simulating argon

diffusion profiles in minerals. Comput . Geosci. 22, 919–929.Widmer T. and Thompso n A. (2001) Local origin of high pressure

vein material in eclogite facies rocks of the Zermatt-Saa s zone,Switzerland. Am. J. Sci. 301, 627–656.

Wijbrans J. and McDouga ll I. (1986) 40Ar/39Ar dating of white micas from an Alpine high-press ure metam orphic belt onNaxos (Greece): the resetting of the argon isotopic system.Contrib. Mineral. Petrol. 93, 187–194.



















































































































http://dx.doi.org/10.1111/j.1525-1314.2011.00956.x

http://dx.doi.org/10.1111/j.1525-1314.2011.00956.x






























Wijbrans J. and McDouga ll I. (1987) On the metamorph ic history of an Archaean granitoid greenstone terrane, East Pilbara,Western Australia, using the 40Ar/39Ar age spectrum technique.Earth Planet. Sci. Lett. 84, 226–242.

Wijbrans J. and McDouga ll I. (1988) Metamorphic evolution ofthe Attic Cycladic Metamorph ic Belt on Naxos (Cyclades,Greece) utilizing 40Ar/39Ar age spectrum measurem ents. J.

Metamorph . Geol. 6, 571–594.Yardley B. (1983) Quartz veins and devolatiliza tion during

metamorph ism. J. Geol. Soc. 140, 657–663.Zhang Z., Shen K., Sun W., Liu Y., Liou J., Shi C. and Wang J.

(2008) Fluids in deeply subducted continental crust: petrology,

mineral chemistry and fluid inclusion of UHP metamorp hic veins from the Sulu orogen, eastern China. Geochim. Cosmo-

chim. Acta 72, 3200–3228.Zimmerman n R., Hammer schmidt K. and Franz G. (1994) Eocene

high pressure metamorph ism in the Penninic units of the Tauern Windo w (Eastern Alps): evidence from 40Ar–39Ardating and petrological investigat ions. Contrib. Mineral. Petrol.

117, 175–186.

Associate editor: David L. Shuster



























Documents

The signature of devolatisation: Extraneous 40Ar systematics in … · 2013. 4. 25. · The signature of devolatisation: Extraneous 40Ar systematics in high-pressure metamorphic rocks