An investigation of closure temperature of the …...PII S0016-7037(00)00560-3 An investigation of...

PII S0016-7037(00)00560-3

An investigation of closure temperature of the biotite Rb-Sr system:The importance of cation exchange

GAWEN R. T. JENKIN,1,2,* ROBERT M. ELLAM ,1 GRAEME ROGERS,1, and FINLAY M. STUART1

1Isotope Geosciences Unit, Scottish Universities Environmental Research Centre (SUERC), Rankine Avenue,Scottish Enterprise Technology Park, East Kilbride, Glasgow G75 0QF, Scotland

2Department of Geology and Geophysics, University of Edinburgh, Grant Institute, West Mains Road, Edinburgh, EH9 3JW, Scotland

(Received February22, 2000;accepted in revised form September18, 2000)

Abstract—Factors controlling closure in the biotite Rb-Sr system were investigated in a detailed study of anamphibolite-facies metacarbonate from the central Swiss Alps. Oxygen isotope data suggest that the rockcooled as a closed system. Calcite-dolomite thermometry temperatures of;450°C to 500°C and feldsparthermometry temperatures of;300°C to 400°C provide evidence of extensive Ca-Mg and Na-K exchangeduring cooling. Biotite in the sample is 90 mol.% phlogopite and has high Rb (;900 ppm) compared to Sr(;0.3 ppm), giving precise Rb-Sr ages. Carefully separated and sized phlogopite shows a range of Rb-Sr agesthat do not simply decrease with grain size as predicted by current models of closure temperature. Rb-Sr agesdecrease from 18.1 Ma to 16.6 Ma with a decrease in mean grain diameter from 1.16 mm to 0.74 mm, butgrains with mean diameter of 0.54 mm show an increase again to 17.6 Ma. This contrasts with Ar-Ar data forsingle phlogopites, which do show a decrease in age with decreasing grain size. The Rb-Sr age pattern is dueto Rb-loss during cooling, which is most pronounced in the finest fraction. The phlogopites are restricted toa 2-cm-thick layer in calcite marble;87Sr/86Sr of the calcite decreases away from the phlogopite band over 4cm, indicating that the calcite was moving towards Sr-isotope equilibration with the phlogopites over thisdistance and that the phlogopite was not equilibrating with an “infinite reservoir.” Ion microprobe traversesacross grains of different minerals reveal systematic core-rim variations in major and trace element concen-trations. In particular, Sr decreases from calcite core to rim, but increases from core to rim in K-feldspar,whereas Rb decreases from core to rim in phlogopite but also increases from core to rim in K-feldspar. Thesegradients are interpreted as indicating the direction of transport of elements during cooling as a result of cationexchange reactions; calcite and phlogopite were sources for Sr and Rb, respectively, whereas K-feldspar actedas a sink for both elements. This chemical equilibration was taking place at the same time as isotopicequilibration during cooling, and was equally important in controlling the apparent ages recorded by the micagrains. In contrast, closure temperature calculations for geochronological systems based on classic Dodson-type models assume parent and daughterelementconcentrations are homogeneous across grains and do notchange with time, onlyisotopicexchange is modeled. Closure in mica Rb-Sr systems will depend both on thefactors that control isotopic exchange (grain size, mode,87Sr diffusion coefficients) and those that controlchemical exchange (grain size, mode, Rb and Sr diffusion coefficients, Rb and Sr contents of phases and theirpartition coefficients). Copyright © 2001 Elsevier Science Ltd

1. INTRODUCTION

Micas from slowly cooled igneous and metamorphic rocksgenerally give Rb-Sr ages younger than the time of peaktemperature. This is explained by micas having a “closuretemperature” for the Rb-Sr system lower than the peak tem-perature of the rock (Ja¨ger et al., 1967; Hunziker et al., 1992).If the closure temperatures for different geochronological sys-tems are known and combined with measured ages they givetemperature-time curves that are fundamental data for deter-mining the cooling and exhumation of orogenic belts (e.g.,Hunziker et al., 1992). Empirical field-based studies in the Alpsindicated closure temperatures of;300°C for Rb-Sr in biotiteand 500°C in muscovite (Armstrong et al., 1966; Ja¨ger et al.,1967), but other empirical studies indicated higher closuretemperatures of up to 400°C for biotite (Verschure et al., 1980;

Del Moro et al., 1982). The alternative to an empirical ap-proach, theoretical calculation of closure temperatures, dependson simple models of the closure process that have not generallybeen tested.

The low closure temperature of micas is generally attributedto retrograde equilibration of Sr isotopes by diffusion duringcooling, although it is acknowledged that recrystallisation, orreaction with fluids, may also play a part in certain cases(Hunziker et al., 1992; Freeman et al., 1997; Villa, 1998).However, although diffusion is proposed as the dominantpro-cessof retrograde equilibration of the Rb-Sr system, the precisemechanismby which closure ages are achieved remains indoubt. An early model of closure temperature (Dodson, 1973)focused on the effect of volume diffusion of Sr-isotopes withinthe mineral grain in question, assuming equilibration with ahypothetical “infinite reservoir.” A more recent model (Jenkinet al., 1995; Jenkin, 1997) suggests that interdiffusion of Sr-isotopes among different minerals in a closed-system rock maybe more realistic.

Petrologists have long attempted to estimate igneous andmetamorphic temperatures by using cation exchange geother-

* Author to whom correspondence should be addressed: ([email protected]).Present address:Department of Geology, University of Leicester,University Road, Leicester LE1 7RH, UK.

Pergamon

Geochimica et Cosmochimica Acta, Vol. 65, No. 7, pp. 1141–1160, 2001Copyright © 2001 Elsevier Science LtdPrinted in the USA. All rights reserved

0016-7037/01 $20.001 .00

1141

mometers (e.g., Ferry and Spear, 1978) and have frequentlyfound that, like many geochronological systems, these areprone to down-temperature re-equilibration during cooling, i.e.,that they record a closure temperature rather than a peaktemperature (e.g., Lasaga, 1983). Ganguly and Ruiz (1987)speculated that chemical exchange might also redistribute Rband Sr among minerals during cooling and that this process,rather than isotopic equilibration, could be controlling the clo-sure temperature.

Until the mechanism of Sr and/or Rb re-equilibration isproperly understood, our ability to make use of mica closureages is purely empirical, and tectonic reconstructions fromcooling curves may be erroneous. Here we investigate closurein the biotite Rb-Sr system and suggest that it may be theconsequence ofbothchemical exchangeand isotopic exchangeduring cooling.

We proceed as follows: first, we examine the sample usingtransmitted light and cathodoluminescence to identify possiblepathways for isotopic exchange. Then, we use stable isotopedata to confirm that this sample was not infiltrated by largeamounts of fluid during retrogression, i.e., that it cooled as aclosed system on at least a hand specimen scale. We then lookat carbonate and feldspar major element compositions to dem-onstrate that these have re-equilibrated during cooling. This isfollowed by the isotopic data where we determine how and whythe Rb-Sr ages of different biotite grains from the sample varyand what reservoir these grains equilibrated Sr-isotopes with.Ar-Ar ages are measured to show how closure in the Rb-Srsystem contrasts with the K-Ar system. Finally, we use ionprobe profiles across grains to demonstrate the mobility of bothmajor and trace elements, including Rb and Sr, during coolingand discuss how this affects closure in mica Rb-Sr systems.

2. SAMPLE INVESTIGATED

The sample investigated in this study (I.29) is an amphibo-lite-facies calcite marble containing a phlogopite-rich layerfrom Lago del Nare`t, Ticino, in the central Swiss Alps (Fig. 1).Subsidiary Rb-Sr isotopic data from a similar sample collected,15 m away (I.28) are also included. Although biotite Rb-Srages are typically determined on pelitic metasediments or gra-nitic gneisses, marble samples have a number of advantages inboth measurement and interpretation of data:

a) The calcite can be gently dissolved with ethanoic acidliberating whole, uncrushed, silicate grains (Section 4.2).

b) The d18O of carbonates is a sensitive indicator of fluidinfiltration, so can be measured to see if the rock behaved asa closed-system during cooling.

c) The planar geometry of the phlogopite-rich layer means thatit should be possible to measure the length scales of ex-change between phlogopite and calcite in the direction per-pendicular to the layer, assuming that concentration gradi-ents parallel to the layer are small.

d) The large proportion of calcite means that the rock couldhave behaved as an infinite reservoir for O and Sr isotopesduring cooling.

In addition, the Lepontine area of the central Alps has beenintensively studied and the regional geology is believed to berelatively well understood, although the metamorphism of the

area is still an area of active research (Todd and Engi, 1997).Considerable geochronological work has been carried out inorder to constrain the cooling history of the region (Hunziker etal., 1992).

3. GEOLOGICAL HISTORY

The sample comes from the core of a;40-m-thick infold ofTriassic marble on the northern margin of the Penninic Maggianappe (Hafner et al., 1975; Steiner, 1984; Gu¨nthert et al.,1996). The marble is part of the Mesozoic cover to the pre-Triassic metamorphic basement of the nappe; the two wereinterleaved during the Alpine orogeny. The relatively ductilemarbles, quartzites, and evaporites in the cover sequence weresynclinal zones of intense strain and are often highly attenuated(Grujic and Mancktelow, 1996). The Penninic nappes representthe distal (rifted) part of the northern pre-Triassic continentalmargin of Tethys, together with the Mesozoic oceanic sedi-ments and platforms (Tru¨mpy, 1980).

The marble outcrop sampled is massive, with intercalationsof more silicic material up to a few tens of cm thick. At its top,it is in thrust contact with alkali feldspar gneiss of the Maggianappe; at its base, it is conformably underlain by calc-micaschist “Bundnerschiefer” (Hafner et al., 1975). Alpine meta-morphism (meso-alpine phase) here reached peak temperaturesof ;575°C and pressures of;6.3 kb (Todd and Engi, 1997).

Fig. 1. Slab cut across layering of sample I.29. An arbitrary datumline was placed at the base of the phlogopite band. Inset shows samplelocation (circled) relative to nappe boundaries (after Gu¨nthert et al.,1996), and location in Switzerland (grid reference: 687.45, 148.20).Heavy dashed line is limit of amphibolite grade metamorphism (Toddand Engi, 1997).

1142 G. R. T. Jenkin et al.

Peak metamorphic conditions in this area were attained afterthe main deformation phase, and were maintained through tothe last pervasive deformation events (D2 and D3, respectively,of Grujic and Mancktelow, 1996; or F2–3 and F4, respectively,of Steck, 1998).

The timing of peak metamorphism is controversial. Only 4km to the south a concordant monazite U-Pb age of 22 Ma isbelieved to date the end of amphibolite-facies metamorphism inthat area (Ko¨ppel and Gru¨nenfelder, 1975; Ko¨ppel et al., 1980).Conversely,;10 km to the north, white mica below the stau-rolite isograd (T , 500°C; Todd and Engi, 1997) gives Rb-Srages of 35–38 Ma, which have been interpreted as dating thepeak of metamorphism there (Ja¨ger et al., 1967). Vance andO’Nions (1992) dated garnet growth at the climax of metamor-phism on the staurolite isograd 40 km along strike to the westat 32–29 Ma. For the Nare`t area, Steiner (1984) gives musco-vite Rb-Sr ages, and biotite Rb-Sr and K-Ar ages, of 22.76 2.0Ma and 16.56 1.2 Ma, respectively, giving an average coolingrate of ;32°C/Ma (assuming closure temperatures of 500°Cand 300°C, respectively). Extrapolating this cooling curve backto the peak temperature suggests that the onset of cooling atNaret commenced at;25 Ma, consistent with the southern-ward decrease in age of peak metamorphism from other studies.Hurford (1986) discusses how the trend in age from north tosouth could be explained by either, a) prolonged cooling froma common peak at 35 Ma; b) peak metamorphism being diach-ronous; or c) multiple peaks of metamorphism. In all of thesescenarios, rocks in the Nare`t area would have spent a consid-erable time at temperatures.500°C before the metamorphicpeak, perhaps as much as 10 Ma.

Fission track zircon and apatite ages of 12.56 0.4 Ma and4.5 6 1.4 Ma in the Nare`t area (Steiner, 1984), together withthe biotite ages, constrain the cooling rate in the interval 300°Cto 120°C to ;15°C/Ma, similar to the rates determined innorthern and central Valle Maggia by Hurford (1986).

Syn-metamorphic fluid infiltration in the central Alps isbelieved to have been limited (Baertschi, 1957; Hoernes andFriedrichsen, 1980; Todd and Engi, 1997). Retrograde meta-morphic (neo-alpine) fluids at temperatures of 400°C to 450°Care restricted to minor fissure veins formed at;22 to 18 Ma(Mullis et al., 1994), which may have been open as late as 13Ma in this area (Purdy and Stalder, 1973).

4. METHODS

4.1. Cathodoluminescence

Polished thin sections of marble were examined at Edinburgh byusing a CITL CCL-8200 Mk3A system (Cambridge, UK) mounted ona Nikon Optiphot microscope (Melville, NY, USA) using a;16-kVelectron beam and;700-mA gun current.

4.2. Mineral Separation and Sizing

Sample chips,1 cm were gently crushed by repeatedly passingthem through a roller mill and gradually decreasing the aperture. Anymaterial,0.75 mm was sieved out each time. This process produced alarge proportion of nearly monomineralic, unbroken grains close totheir original grain size, especially in the coarser fractions. Mineralseparation of various grain-size fractions from sieving was carried outby using a combination of magnetic separation, heavy liquids, andhand-picking.

To dissolve out unbroken silicate grains from the phlogopite-richband in I.29; 9 g of chips, 0.7 cm were repeatedly leached in 10-

to 30-mL fractions of 1 M ethanoic (acetic) acid (CH3COOH) for timesranging from 1.5 to 44 h (reaction was fast initially, but slower for laterleachings). In 10 steps the calcite was removed from the sample,leaving a silicate residue of single grains and loosely bonded aggre-gates of grains. To check that this prolonged leaching did not affect theRb-Sr isotopic composition of the phlogopite, a control sample ofphlogopite, previously separated by conventional methods from thesame sample, was immersed in CH3COOH for the duration of theexperiment (460 h) and was then analysed. Its isotopic composition andcalculated age was not substantially different from the unleachedportion (Table 6).

The undissolved silicate residue was sieved into four size fractions:.0.75 mm, 0.75 to 0.5 mm, 0.5 to 0.25 mm, and,0.25 mm, andphlogopite was separated from the three coarsest fractions. Because ofthe tabular shape of the phlogopite grains the largest dimension of thegrains tended to be rather larger than the nominal sieve size. For thisreason, the actual grain size of a representative sample of;22 grainsfrom each size fraction (including for the.0.75 mm and 0.75- to0.5-mm fractions all grains used for Ar-Ar analysis), was measuredunder a binocular microscope. Grains were typically elliptical prisms orrectangular plates with dimensionsa, b, g in ratios ' 8:5.3:1; theorientation ofg being parallel to thec-axis,a andb the dimensions ofthe plate face (aspect ratio; 1.5). For the phlogopite fractions analysedfor Rb-Sr isotopes the compositions measured were biased towardsrepresenting the larger grains, because these are volumetrically moresignificant. Therefore the meanb length, used to represent the grainsize of the fraction, was weighted according to the mass or volumetricproportions of each grain in it. Thus for the.0.75-mm fraction meanweightedb 5 1.16 mm, the 0.75 to 0.5-mm fraction5 0.74 mm, andthe 0.5 to 0.25 mm fraction5 0.54 mm.

4.3. Traverse Microsampling

Traverses of small whole-rock samples were taken with a diamond-tipped, hand-held dental drill on a surface cut perpendicular to theorientation of the phlogopite rich band, and running across the band.Sample trenches were elongated parallel to the datum line and pro-duced 2 to 4 mg of material for O and C isotope analysis, or 20 to 30mg for Sr isotope analysis (Fig. 1). The drill was cleaned between eachtrench with 1 M ethanoic acid and de-ionized water, the sample with aclean alcohol-soaked tissue.

4.4. Stable Isotope Analysis

At SUERC, oxygen was liberated from silicate minerals by usingBrF5 (Clayton and Mayeda, 1963), and was then reduced to CO2 ford18O analysis. CO2 for analysis ofd18O andd13C was liberated fromcalcite separates and calcite in whole-rock powders by reaction withphosphoric acid (McCrea, 1950). Whole-rock traverse samples weigh-ing 0.8 to 2.7 mg were analysed ford18O andd13C by using a VGIsogas Auto-Carbonate Preparation System on-line to a VG PRISMSeries II isotope ratio mass spectrometer (Manchester, UK). The reac-tion temperatures and times used for both techniques were such thatonly the composition of the calcite was analysed; any dolomite re-mained unreacted. Results are reported relative to V-SMOW and V-PDB. Repeat analyses of NBS-28 quartz at SUERC give an averaged18O of 9.6‰. Both carbonate techniques give averaged18O 5 26.66andd13C 5 21.06‰ for NBS-20 calcite, with a precision of,0.1‰.Replicate analyses were made on silicate samples until a precision(1sn21) of ,0.2‰ was obtained.

4.5. Rb-Sr Analysis

Before analysis at SUERC, separates obtained using heavy liquidswere washed with Analar acetone in an ultrasonic bath for;5 min. Allmaterials to be analysed were similarly washed with ion-exchangedwater. Micas were then leached in 1 M CH3COOH and 1 M HCl for 1 heach to remove adhering calcite. Whole-rock samples and calciteseparates were gently leached in an ultrasonic bath for;30 min with10 mL of 1 M NH4Cl, and then left overnight to remove any easilyexchangeable Sr that could be due to weathering (Ohde and Elderfield,1992). Examination of leachates from whole-rock samples showed that

1143Closure temperature of the biotite Rb-Sr system

these rigourous cleaning routines were unnecessary, and they were notcarried out for the drilled whole-rock traverse samples.

Silicate samples were spiked with87Rb and84Sr spikes, then dis-solved by HF-HNO3-HCl digestion. Rb and Sr were separated in 2.5NHCl by using Bio-Rad AG50W X8 200–400 mesh cation exchangeresin (Richmond, CA, USA).

Whole-rock samples were first treated with 1 M HCl to dissolve thecarbonate fraction, this solution was reserved and, before spiking, wasadded to that resulting from dissolution of the silicate fraction as above.This avoided the formation of large amounts of insoluble CaF2 whenHF was added.

Calcite samples and calcite in whole-rock traverse samples weredissolved in 5 mL of 1 M CH3COOH for 1 h. The supernatant wasquantitatively removed from any residual material by centrifuging andpipetting off the supernatant, and washing twice with water, centrifug-ing the solids down each time. Dolomite is not believed to have beendissolved by this treatment. The solution was dried, spiked, and sepa-rated as per the silicate samples.

For mass spectrometry, Sr samples were loaded onto single Tafilaments with 1N phosphoric acid, Rb samples onto triple Ta fila-ments. Sr samples were analysed on either a VG 54E single collectorthermal ionization mass spectrometer or a VG Sector 54-30 multiplecollector mass spectrometer. On both instruments, the87Sr/86Sr ratiowas corrected for mass fractionation by using86Sr/88Sr 5 0.1194 andan exponential law. Repeat analysis of the NBS-987 Sr standard gave87Sr/86Sr 5 0.71024 6 3 (1 SD, n 5 20) on the VG 54E and0.7102436 10 (1 SD,n 5 88) on the VG Sector 54-30. Rb sampleswere analysed on either a VG MM30 or a VG 54E single collector massspectrometer. Rb and Sr spike contribution, and concentrations werecalculated by using adaptations of the standard algorithms of Kroghand Hurley (1968).

Amounts of Rb analysed were 4 to 252mg for mica and whole-rocksamples and 0.08 to 9.8mg for calcite and feldspar. Amounts of Sranalysed were 2 to 23mg for calcite, feldspar and whole-rock samples,and 37 to 728 ng for mica. Total procedural blanks for Rb and Sr were,0.5 ng. Application of a pessimistic blank correction of 1 ng of Srwith 87Sr/86Sr 5 0.710 decreases Sr concentrations by,0.06 ppm,87Sr/86Sr of calcite, feldspar and whole-rock samples increases by,3 3 1026, whereas87Sr/86Sr of micas increases by as much as 0.1.However, applying this blank correction makes no difference to cal-culated ages, and all data here are uncorrected for blanks.

All errors are quoted as two standard errors (SE). Errors on Srabundances are typically,0.25% for all phases. A blanket error of 1%was put on all Rb abundances, apart from Rb measurements in calcite,in which the precision from the mass spectrometry runs was up to 9%,and this error was used. All errors were propagated through to theuncertainty on87Rb/86Sr and calculated ages. A realistic measure of theoverall precision on calculated ages is probably given by the repeatanalyses of the 1 to 1.5-mm phlogopite fraction of I.29, which agree towithin 0.3 Ma and are indistinguishable within error (Table 6).

4.6. Ar-Ar Analysis

The phlogopites were cleaned ultrasonically in acetone and sizefractions were wrapped in Al foil and irradiated along with splits ofTCR sanidine monitor (McDougall and Harrison, 1988) at 1 MW for18 h in the TRIGA research reactor at Oregon State University. ThecalculatedJ value was 0.004356 0.00002. After 14 months eachphlogopite grain, or collection of grains, was loaded into 2-mm-diam-eter holes in a Cu laser pan mounted in a Pyrex viewport at SUERC.Samples were fused for 2 min using a continuous wave 20 W Nd:YAGlaser with an external shutter focused through a modified petrographicmicroscope. Released gases were purified for 10 min on two SAESGP-50 getters at 300°C. Ar isotopes were analysed on a MAP 215 massspectrometer with a Nier-type source in static mode. Masses 36 to 40were measured on a Johnson electron multiplier in 10 cycles of peakjumping down mass. Absolute abundances of each isotope were cal-culated by extrapolation back to the inlet time. These raw data werecorrected for blanks,37Ar decay, neutron interferences and mass dis-crimination. Blanks were measured after every two or three samples.Because of the decrease in sample size, blank corrections become moresignificant for the smallest grains. The 1s uncertainty in the ages takes

into account all sources of error in the measurement and correction ofdata.

Twenty-four grains withb ranging from 1.7 to 0.4 mm, with weightsranging from 2.52 to 0.13 mg, were analysed individually. All thesegrains were examined under a binocular microscope and, apart fromsome containing small inclusions of calcite or apatite, they were trans-parent pale brown, with no obvious intragrain boundaries, flaws orfractures. For the finest grain fraction (0.5 to 0.25-mm sieve);20grains were analysed together in 8 groups to yield more reliable results.

4.7. Electron Microprobe (EMP)

Analysis was carried out by Wavelength Dispersive Spectrometry(WDS) on a Cameca Camebax Microbeam electron microprobe (Paris,France) at the Department of Geology and Geophysics, University ofEdinburgh. All samples were analysed at 20 kV gun potential and aprobe current of 20 nA, except carbonates at 10 nA. The beam wasrastered over an area 203 20 mm for carbonates and feldspars. Themodal analyses given in Section 5.1 were determined by point countingwith the electron microprobe.

4.8. Ion Microprobe (IMP)

At the Ion Microprobe Facility at the Department of Geology andGeophysics, University of Edinburgh, samples were analysed using an16O2 primary beam of 15 keV impact energy. Positive secondary ionswere measured and molecular peaks substantially reduced by usingenergy filtering (Zinner and Crozaz, 1986). Essentially only high-energy positive secondary ions with energies of 55 to 95 eV weremeasured. Secondary ion count rates were normalised to30Si or 44Ca.Isotopes26Mg, 30Si, 44Ca, 54Fe, 55Mn, 88Sr, and138Ba were recordedfor calcites. Isotopes23Na, 30Si, 40Ca,41K, 85Rb, 88Sr, and138Ba wererecorded for feldspars, with the addition of26Mg for phlogopite. Mass130.5 was measured to monitor background. Ionisation efficiencies foreach element relative to these species were determined by using acombination of glass and mineral standards (OKA, NCC for calcite andSRM 610 and SHF1 feldspars and Corning borosilicate glass for thesilicates). In these cases, precision in the absolute concentration shouldbe better than 5%. No appropriate standards were available for thephlogopite but although the absolute concentrations are probably onlyaccurate to 20% to 30%, variations within the crystal will largely belimited by the counting statistics (6 a few %).

5. PETROGRAPHY

5.1. Mineralogy

The 2-cm-wide phlogopite-bearing layer is made up of phl-ogopite (;19% by volume), calcite (;51%), and dolomite(;29%). Elsewhere in the sample phlogopite and dolomite areabsent, apart from;4% of fine dolomite exsolved from thecalcite (;78%) during cooling. Quartz (;13%), K-feldspar(;2%), and albite (;3%) also occur outside the phlogopitelayer, and trace apatite and muscovite occur throughout. Thealbite contains many inclusions, especially in grain cores. K-feldspar is free of inclusions, untwinned, and is easily mistakenfor quartz in thin section (c.f., Wenk et al., 1986).

Dolomite grains in the phlogopite band are;0.25 mmacross, calcite;0.5 mm, phlogopite averages;0.6 mm per-pendicular toc. Outside the phlogopite layer, calcite is rathercoarser (1.5–2.0 mm) and the quartz, K-feldspar, and plagio-clase average;0.5 mm. Apatite prisms are 0.1 to 0.2 mmacross.

The phlogopite is believed to have formed during progrademetamorphism in a dolomite- and K-feldspar-rich layer by thereaction:


KAlSi 3O8 1 3CaMg(CO3)2 1 H2O 5 KMg3AlSi3O10(OH)2

1 3CaCO3 1 3CO2

until all K-feldspar was consumed (Flowers and Helgeson,1983). Up to 33% of the OH2 in the phlogopite is replaced byF2 (Section 7.3), reducing the amount of water required for itsformation and enhancing its stability field (Abercrombie et al.,1987), so that it probably developed in the interval 400°C topeak temperature.

Calcite exhibits four textural features that could have influ-enced transport through it: grain boundaries, twin planes, rarefractures, and fine exsolution (Fig. 2). Grain boundaries can bevery narrow, only distinguishable by the change in orientationof twin planes across them, but they can also appear as darklines in transmitted light, up to 20mm wide. Twin lamellae aretypically a few microns to 20mm thick and are always straight.These are Type II twins, believed to be formed at 150°C to300°C (Burkhard, 1993). Fractures are often discontinuous, upto 60mm wide, and are a green-brown color, or dark and dustyin transmitted light. Fractures sometimes show offsets at twin

planes (Fig. 2a) suggesting that the twin planes were alreadydeveloped when the fractures formed. Fine exsolution of dolo-mite (as identified by IMP/EMP, Section 7.1) forms dark dustypatches in variable amounts through the sample, being mostdeveloped in and around the phlogopite layer.

5.2. Cathodoluminescence

Calcite shows three types of luminescence (Fig. 2): i) areasthat are clear in transmitted light luminesce orange; ii) irregularpatches, 10 to 40mm across, within the orange luminescingmaterial are dull red and believed to be exsolved dolomite; iii)bright yellow luminescing calcite usually speckled with tinygrains of dull-red luminescent material. The latter type occurs:a) along grain boundaries where they appear to be wide intransmitted light; b) sometimes along twin lamellae; c) alongfractures; and d) in cloudy areas of fine exsolution. Graham etal. (1998) interpret the development of yellow luminescentcalcite in similar sites in a high-grade marble as mapping outthe pathways of infiltration of an externally derived fluid into

Fig. 2. Calcite textures in transmitted light (TL) and cathodoluminescence images (CL). Abbreviations: Ab5 Albite, F 5Fracture, GB5 Grain Boundary, Kf5 K-feldspar, Ph5 Phlogopite, Q5 Quartz, TP5 Twin Planes (orientations marked),X 5 Area of fine dolomite exsolution. (a) Large calcite grain with fractures at;159 mm relative to datum. Width of fieldof view 1.6 cm. TL: Note arrowed fracture being offset at twin plane. CL: Grain luminesces mostly orange, but fracturesbright yellow. Twin lamellae in small exsolution patch also yellow luminescing. Note grain boundary at W contains noluminescent calcite. IMP traverse W is shown in Fig. 7c. (b) Large calcite grain at26 mm relative to datum. Most calcitein the sample is less altered than this. Width of field of view 1.6 cm. TL: Note thick dark rim to grain, twin planes and cloudyareas to bottom left. CL: Grain boundaries, cloudy areas, and some twin planes are all yellow luminescing. (c) Calcite at130 mm relative to datum with only minor exsolution. Width of field of view 1.3 cm. TL: Note twin planes and cloudyexsolution at top of grain. Regular black spots are EMP traverse points. IMP traverse Y is shown in Fig. 7a, traverse Z inFig. 7b. CL: Grain is clear orange luminescent calcite. Twin lamellae and grain boundaries crossed by traverse Y do notshow yellow luminescent calcite; this is restricted to the area of fine exsolution at the top of the grain.


the sample. In contrast, we show below that there is no evi-dence for fluid infiltration into our sample, and that the fineexsolution that the yellow luminescence is related to appears tobe an isochemical replacement. We relate the yellow lumines-cence along some twin boundaries to incipient fine exsolution,concentrated there due to the higher surface energy.

Coarse dolomite grains exhibit dull red luminescence that isslightly brighter at the rims. Plagioclase luminesces deep azure-blue; K-feldspar pale blue, but sometimes pale yellow-green inthe cores. Quartz luminesces very dark brown, speckled withorange—typical of high-grade metamorphic rocks that havebeen slowly cooled (Zinkernagel, 1978). Rare fractures in thequartz have a dull red luminescence. Apatite luminesces green-ish-gray. Phlogopite is nonluminescent.

6. OXYGEN AND CARBON ISOTOPE DATA—EVIDENCEFOR CLOSED SYSTEM EXCHANGE

Whole-rock carbonate and calcite mineral separates wereanalysed from slices of the sample from the phlogopite layer (0to 114 mm relative to datum; Fig. 1) and adjacent to it (215to 230 mm relative to datum). The ranges in bothd18O(120.79 to 121.03‰) andd13C (12.65 to 12.80‰) arewithin the precision of the technique (Table 1). Whole-rocktraverse powders from sampling points221 mm to119 mmalso showed no significant trend in composition across thephlogopite band and gave mean values, comparable with thebulk samples. Almost identicald13C andd18O values, differingby ,0.5‰, were obtained on I.28, 15 m away in the sameoutcrop, suggesting that these values are typical for the area.

These values can be compared with values for unmetamor-phosed Triassic carbonates to assess the possibility of fluidinfiltration during metamorphism. Thed13C values are withinthe normal range for Triassic limestones (0.926 2.46‰;Veizer et al., 1980), but thed18O values are slightly lower than

the normal range (123 to 129‰; Veizer and Hoefs, 1976).Within the Pennine nappes Burkhard and Kerrich (1988) founda range ofd18O of 120.6 to125.7‰ for Triassic limestonematrix calcites in the Rhone Valley and suggested that thelower values could relate to oxygen isotope exchange withdetrital silicates. Baertschi (1957) found a range of125 to130‰ in pure marbles within 10 km of Nare`t, but impuremarbles hadd18O values ranging from129 to 115‰. Calcu-lations show that equilibration of the original calcite withdetrital silicate grains with a lowd18O of ' 110‰ could havecaused as much as an 8‰ downward shift in the calcite inthis sample, in addition, decarbonation during the progradeformation of phlogopite could cause downwards shifts of,2‰. Thus infiltration of this rock by an exotic fluid is notrequired to explain the carbonate stable isotope composi-tions. Any fluid infiltration associated with the formation ofphlogopite would have taken place before the peak of meta-morphism and not affected the retrograde effects being ex-amined in this paper.

The minerald18O values and apparent temperatures are alsoinconsistent with the rock being infiltrated with a low-d18Ofluid during cooling. Because of the relatively slow rate ofoxygen isotope exchange of quartz with fluid compared tocalcite and phlogopite (Farver and Yund, 1991; Farver, 1994;Fortier and Giletti, 1991), fluid infiltration would produce largeand nonequilibriumDquartz-calciteand Dquartz-phlogopiteand un-reasonably low apparent equilibration temperatures. Calculatedtemperatures (Table 1) are not unreasonably low and are sim-ilar to calcite-dolomite solvus equilibration temperatures, ex-cept for Tquartz-calcite in the phlogopite-rich layer, which isunreasonably high. These temperatures are qualitatively similarto what would be expected as a result of closed-system equil-ibration (Jenkin et al., 1991, 1994), with the smallDquartz-calcite

in the phlogopite-rich layer being caused by low-temperatureexchange between calcite and phlogopite raising calcited18Otowards the quartz value.

7. MINERAL MAJOR ELEMENTCHEMISTRY—EVIDENCE FOR RETROGRESSION

7.1. Calcite and Dolomite

The average composition of all orange-luminescing calcite isgiven in Table 2. Assuming calcite and dolomite,2.5 mmapart are in equilibrium gives a calcite-dolomite solvus equil-ibration temperature (Anovitz and Essene, 1987) of 490623°C (sn21, n 5 19). This indicates that Ca and Mg havere-equilibrated between these minerals during cooling frompeak temperatures, because calcite Mg contents at peak tem-perature would have been nearly twice as great (;7 mol.%MgCO3). Dolomite compositions are also consistent withequilibration with calcite at;490°C, i.e., they fall close to thesteep dolomite limb of the solvus at this temperature. Allorange-luminescing calcite throughout the sample has almostidentical Mg and Fe concentrations and gives the same solvusequilibration temperature (4826 17°C, 1sn21, n 5 179),suggesting that Ca and Mg could have re-equilibrated betweencalcite and dolomite over distances of up to 5 cm.

Chemical re-equilibration between calcite and dolomite dur-ing cooling is rate limited by Ca-Mg interdiffusion in thecoarser calcite. For these nearly pure calcites, the Ca-Mg in-

Table 1. Stable isotope data (in ‰) and apparent oxygen isotopeequilibration temperatures, sample I.29.

Sample: 0 to114 mm 215 to 230 mm 221 to 119 mm

d13CCca 2.656 0.05 2.806 0.05

d13CWRCc 2.696 0.05 2.776 0.05 2.816 0.14b

d18OCca 21.036 0.05 20.866 0.05

d18OWRCc 20.846 0.05 20.796 0.05 21.226 0.22b

d18OWRCc 20.846 0.05 20.796 0.05 21.226 0.22b

d18OOc 21.246 0.01 21.636 0.18

d18OPhd 17.716 0.19 —

DQ-Ph 3.536 0.19 —TQ-Ph (°C) 4776 20 —DCC-Ph 3.326 0.20 —TCC-Ph (°C) 4236 21 —DQ-Cc 0.216 0.05 0.776 0.19TQ-Cc (°C) 10726 160 4296 87

Cc 5 calcite, Q5 quartz, Ph5 phlogopite, WR5 whole-rock.a 120–160-mm fraction separate for 0 to114 mm, 160–500-mm

fraction for 215 to 230 mm.b Average for 15 traverse points across this interval.c 160–500-mm fractions.d 1–1.5-mm grains.Fractionation factors used to calculate apparent temperatures: Cc-Ph;

Chacko et al. (1996), Q-Cc; Clayton et al. (1989). Phlogopite assumedto contain 33 mol.% fluorphlogopite.


terdiffusion coefficient will be approximated by the Mg tracerdiffusion coefficient (e.g., Chakraborty, 1995). Using the Mgtracer diffusion coefficient of Fisler and Cygan (1999) the(Dodson) closure temperature for Ca-Mg exchange in the 0.5-mm-diameter calcite grains in the phlogopite band cooling at32°C/Ma should be;510°C, slightly higher than the averagecalcite-dolomite solvus equilibration temperature. Lower clo-sure temperatures that better fit the measured temperatureswould result from: a) a higher diffusion coefficient; b) a slowercooling rate; or c) smaller effective grain size for diffusion. Theanalytical uncertainty on the diffusion coefficient is approxi-mately a factor of 2.5 (Fisler and Cygan, 1999); using a value2.5 times higher gives closure temperatures for 0.5-mm grainsof 490°C, similar to the measured values. To reduce the closuretemperatures to equal the measured temperatures would requirereducing the cooling rate to;10°C/Ma for a 0.5-mm grain.This would not be unreasonable for cooling down to 500°C ifthe metamorphic peak was as old as 38 Ma. A smaller effectivegrain size for diffusion could be accomplished by exsolution ofdolomite within the calcite grains, such as the 10- to 40-mmdolomite grains observed within some of the orange lumines-cent areas. However this is not always observed and cannottherefore explain all the low temperatures. More detailed mod-elling of the shapes of the Mg profiles across grains shouldallow discrimination between options (a) and (b).

Both EMP and IMP traverses in dark cloudy calcite withbright yellow speckled luminescence (e.g., Fig. 7b) reveal thatMg, Fe, and Mn show minor variations over short wavelengths(;20 mm) to values above and below the values in adjacentorange luminescent calcite (Fig. 7a). Variations in Mg, Fe, andMn are all well-correlated. Bulk compositions of the traversesin clear and cloudy calcite in the same grains are indistinguish-able (e.g., Fig. 7a,b). This yellow speckled luminescence is,therefore, interpreted as fine-grained (smaller than the;20-mmanalysis spot) isochemical exsolution of small quantities of Mg,Fe, and Mn-rich dull-luminescent dolomite from the orange

luminescent calcite, leaving a yellow-luminescent Mg-poorcalcite (composition in the “erratic” luminescence field ofSavard et al., 1995). The isochemical nature of the exsolutionsuggests that, if it were initiated by fluid, the fluid was probablypresent in only small amounts. The variable Mg and Fe con-tents give a much larger range of calcite-dolomite solvus tem-peratures with some Mg- and Fe-rich (i.e., dolomite-rich) sam-ples giving unrealistic temperatures over 600°C. The lowestMg and Fe contents measured in an exsolved area probablycorrespond to the calcite composition after exsolution and giveequilibration temperatures of 300°C to 400°C, although oneMg-poor spot gives a temperature of 100°C. These low tem-peratures, along with the fine-grained nature of the exsolution,and the fact that it is developed along low-temperature twinplanes, indicate that it is a late, low-temperature, effect.

7.2. Feldspars

Plagioclase is dominantly albite (;Ab98An1Or1) showing nosignificant core-rim variation (Table 3). A single;30mm grainof oligoclase (;Ab87.5An11.5Or1) was found at the margin ofone albite. K-feldspar shows smooth zonations in composition,with K- and Ba-rich rims (;Ab5.9An0.4Or93.3Cn0.4) comparedto the cores (;Ab9.3An0.0Or90.5Cn0.2). These feldspar compo-sitions are typical for K-feldspar bearing marbles from this partof the Alps (Wenk et al., 1986). Intergrowths of more than oneplagioclase mineral have previously been identified in metacar-bonates from the central Alps, and are believed to relate tolow-temperature miscibility gaps in the plagioclase series(Wenk et al., 1986, 1991).

The compositions of feldspar pairs can be used to estimatethe temperature on the solvus at which they last equilibrated, ifpressure is known. Application of two-feldspar thermometry toslowly cooled metamorphic rocks is fraught with difficultybecause,inter alia: a) the feldspars are likely to have becomeSi-Al ordered to some extent, whereas all available geother-

Table 2. Average compositions of carbonatesa analysed by electron microprobe (EMP) and ion microprobe (IMP).

Calcite(EMP) (sn21)

Calcite(IMP) (sn21)

Dolomite(EMP) (sn21)

Dolomite(IMP) (sn21)

MgO 1.61 0.17 1.54 0.11 20.48 0.34 20.63 0.13CaO 53.96 0.56 53.96 0.16 30.31 0.25 30.39 0.11FeO 0.27 0.06 0.24 0.05 1.34 0.16 1.49 0.12MnO 0.09 0.02 0.04 0.01 0.13 0.04 0.12 0.02CO2 44.30 0.38 44.21 0.02 47.05 0.18 47.37 0.03Total 100.28 2.03 100.00b 0.00 99.54 0.79 100.00b 0.00Fe (ppm) 2117 470 1829 356 10386 1260 11552 959Mn (ppm) 688 167 281 76 978 275 931 170Sr (ppm) bdl 138 13 bdl 63 4Ba (ppm) bdl 2.64 2.19 bdl 1.10 0.30% end-member:MgCO3 4.0 0.4 3.8 0.3 47.5 0.7 47.6 0.3CaCO3 95.5 0.5 95.8 0.3 50.6 0.5 50.3 0.2FeCO3 0.4 0.1 0.3 0.1 1.7 0.2 1.9 0.2MnCO3 0.1 0.1 0.1 0.0 0.2 0.1 0.2 0.0T (°C)c 486 19 479 14n 76 103 12 13

a Calcite compositions are for orange-luminescing calcite free from late exsolution. Dolomite compositions are for coarse dolomite grains in thephlogopite band, not from the late exsolution.

b Ion probe analyses are recalculated to total 100%.c Calcite-dolomite solvus temperature by using equation 31 of Anovitz and Essene (1987).


mometers are for disordered feldspars; b) during cooling therate of K-Na exchange will be greater than Ca-(Na,K) ex-change, which is coupled to slow Si-Al exchange; thus equil-ibration of different components may be last achieved at dif-ferent temperatures (Kroll et al., 1993). Nevertheless, a two-feldspar thermometer is applied here to test if re-equilibrationhas taken place since peak temperature (Table 4). Three tem-peratures,TAb, TOr, andTAn, are calculated which relate to theequilibrium partitioning of each component respectively be-tween the two phases. A fourth temperature,Tmod, is calculatedby adjusting feldspar compositions within their uncertainty inorder to minimise the difference between the first three tem-peratures, and then finding the average. Because of the verylow An contentsTAn is believed to be unreliable, as smallmeasurement errors propagate to large uncertainties inT. It ispossible that all An values are overestimates due to contami-

nation by the calcite host. The generally lower An values foundby IMP, in which a surface layer is sputtered away beforeanalysis, would support this.

Excluding TAn, all results are,400°C atP 5 3 kb (theeffect of pressure is small,;10°C/kb). These results are con-firmed by comparing the data with the solvus isotherms on theternary diagram. Plagioclase lies to the low temperature, Ab-rich side of the 400°C, 3 kb isotherm; at peak conditions(575°C and 6.3 kb) it would be expected to have had a com-position ;Ab92Or8. K-feldspar also generally lies to the lowtemperature, An-poor side of the 400°C, 3 kb isotherm, butbecause of the steepness of the solvus here, analyses with evenminor amounts of An plot at higher temperatures. At peakconditions the K-feldspar would have had a composition of;Ab22Or78. Thus it is clear that significant Na-K exchange has

Table 3. Average composition of feldspars analysed by electron microprobe (EMP) and ion microprobe (IMP).

K-fsp(rims)(EMP) (sn21)

K-fsp(rims)(IMP) (sn21)

K-fsp(core)(EMP) (sn21)

K-fsp(core)(IMP) (sn21)

Plag(rims)(EMP) (sn21)

Plag(rims)(IMP) (sn21)

Plag(core)(EMP) (sn21)

Plag(core)(IMP) (sn21)

SiO2 64.62 0.93 64.61a 64.60 0.30 64.61a 68.37 0.69 68.35a 68.32 0.42 68.35a

Al2O3 18.20 0.20 18.20a 18.19 0.07 18.20a 19.43 0.29 19.40a 19.34 0.11 19.40a

CaO 0.12 0.06 0.02 0.00 bdl 0.03 0.00 0.27 0.18 0.24 0.18 0.22 0.25 0.12 0.06SrO — 0.02 0.00 — 0.02 0.00 — 0.00 0.00 — 0.00 0.00BaO 0.38 0.12 0.22 0.10 0.23 0.12 0.16 0.01 bdl 0.00 0.00 bdl 0.00 0.00Na2O 0.60 0.09 0.75 0.11 0.97 0.05 1.18 0.26 11.83 0.21 11.80b 0.32 11.80 0.16 11.91b 0.38K2O 15.52 0.17 15.68b 0.29 15.11 0.19 15.19b 0.57 0.16 0.06 0.16 0.03 0.19 0.04 0.17 0.04Rb2O — 0.04 0.01 — 0.04 0.01 — 0.00 0.00 — 0.00 0.00Total 99.32 1.15 99.54 0.23 99.01 0.44 99.42 0.30 100.05 0.87 99.95 0.43 99.86 0.37 99.96 0.35Sr (ppm) — 186 37 — 150 14 — 16 11 — 8 4Ba (ppm) 3360 1058 1988 864 2089 1072 1405 50 bdl 2.4 2.0 bdl 1.4 1.2Rb (ppm) — 357 123 — 341 130 — 0.29 0.08 — 0.20 0.04Formula per 8(O):Si 3.003 0.005 3.001 0.002 3.004 0.003 2.999 0.002 2.989 0.010 2.991 0.007 2.993 0.008 2.991 0.006Al 0.997 0.006 0.996 0.001 0.997 0.005 0.996 0.001 1.001 0.011 1.001 0.002 0.998 0.006 1.011 0.002Ca 0.005 0.004 0.001 0.000 — 0.002 0.000 0.013 0.009 0.011 0.008 0.010 0.012 0.006 0.003Sr — 0.001 0.000 — 0.000 0.000 — 0.000 0.000 — 0.000 0.000Ba 0.005 0.004 0.004 0.002 0.002 0.003 0.003 0.000 — 0.000 0.000 — 0.000 0.000Na 0.054 0.008 0.067 0.010 0.088 0.005 0.106 0.024 1.003 0.021 1.001 0.025 1.003 0.013 1.011 0.030K 0.920 0.014 0.929 0.016 0.896 0.009 0.900 0.033 0.009 0.003 0.009 0.002 0.010 0.002 0.010 0.002Rb — 0.001 0.001 — 0.001 0.000 — 0.000 0.000 — 0.000 0.000SXc 0.984 0.012 1.003 0.012 0.986 0.006 1.012 0.009 1.025 0.015 1.021 0.030 1.023 0.011 1.026 0.029S Al,Si 4.000 0.004 3.997 0.003 4.001 0.002 3.995 0.002 3.990 0.004 3.992 0.009 3.992 0.005 3.992 0.008% end-member:An 0.5 0.4 0.2 0.0 — 0.2 0.0 1.2 0.8 1.1 0.8 0.9 1.1 0.6 0.3Ab 5.5 0.8 6.7 1.0 8.9 0.5 10.5 2.4 97.9 1.0 98.1 0.8 98.1 1.1 98.5 0.2Or 93.5 0.8 92.7 0.9 90.9 0.5 89.0 2.4 0.9 0.3 0.8 0.1 1.0 0.2 0.9 0.2Cn 0.5 0.4 0.4 0.2 0.2 0.3 0.3 0.0 — 0.0 0.0 — 0.0 0.0n 9 4 8 2 17 5 14 4

a Not measured, average value by EMP used to calculate total.b Average feldspar value adjusted to be comparable to average EMP value.c X 5 Ca, Sr, Ba, Na, K, Rb.

Table 4. Feldspar thermometry temperatures for P5 3 kb.

K-feldspar Plagioclase TAb TOr TAn Tmod

Average IMP rims Average IMP rims 318 295 1960 3586 9Average IMP cores Average IMP cores 377 305 4042 3766 148Average IMP rims Oligoclase; Ab87An12Or1 341 370 675 3966 0

Temperatures calculated using the program SOLVCALC 2 (Wen and Nekvasil, 1994) and the solution model of Elkins and Grove (1990). Thecelsian component in K-feldspar was added to Or for these calculations.


taken place between the feldspars during cooling from peaktemperature.

Ordering will raise the solvus temperatures; Brown andParsons (1989) propose an increase of;100°C for fully or-dered feldspar pairs. However, the low 2V (#50°) measured onI.29 K-feldspar suggests that it is still disordered, as has beenfound for K-feldspar in other Alpine marbles that remainedundeformed during cooling (Wenk et al., 1986).

Given the homogeneity in plagioclase compositions it isunlikely that plagioclase cores retain equilibrium with K-felds-par cores. Nevertheless, calculated temperatures are higher forcores compared to rims (Table 4), as would be expected if thezoning in K-feldspar reflects exchange in the cores laggingbehind the rims due to the slow rate of transport within thegrain (almost certainly by diffusion) compared with the rate ofexchange between grains (e.g., Lasaga, 1983; Jenkin et al.,1995). This core-rim zonation in K-feldspar occurs whether ornot the K-feldspar is touching a plagioclase grain, indicatingthat K and Na were efficiently transported between feldsparsalong the grain boundaries. The lack of zonation in the plagio-clase could be due to a number of reasons, including: a)relatively fast exchange of Na and K within plagioclase ironingout gradients, either because diffusion was intrinsically fast, orbecause it was enhanced by the short-circuiting along the grainboundaries of its many inclusions; or b) peristerite exsolutionironing out gradients.

7.3. Phlogopite

The biotite mica is near end-member phlogopite (;Ph90)with appreciableF substitution (Table 5). No significant vari-ation in major element composition between cores and rims, orin grains of different size was detected.

8. ISOTOPIC DATA

8.1. Mineral and Whole-rock Rb-Sr Data

Mineral separate and whole-rock data for I.29 and nearbysample I.28 are given in Table 6 and plotted on an isochrondiagram in Fig. 3. Where measured, calcites were found to havenegligible Rb contents (,0.07 ppm) and a value of zero isassumed where Rb was not measured. All calcite87Sr/86Srratios are significantly higher than Triassic seawater values(;0.7071–0.7078; Faure, 1986) indicating that they have ex-changed with high87Rb/86Sr phases during metamorphism. Ifthe whole-rock samples originally contained no high87Sr/86Srdetrital input and remained closed systems through metamor-phism, using their Rb contents to correct their87Sr/86Sr toTriassic ages (205–245 Ma) should result in Triassic seawatervalues. Although Triassic whole-rock87Sr/86Sr values clusteraround Triassic seawater values (Table 6), suggesting littledetrital input, relatively low values in the mica-bearing bandsand high values in the host marbles imply significant loss of87Sr from the former to the latter over distances of a few cm.

Phlogopites have high Rb/Sr ratios, resulting in precise ageswith errors, 0.2 Ma. Phlogopite Sr contents are very low, butvariable. The lowest Sr values (;0.3 ppm in I.29,;1.8 ppm inI.28) probably reflect the true Sr content; higher values mostlikely relate to small inclusions of a Sr-bearing phase such ascalcite or apatite. The presence of these inclusions does not

affect calculated ages as the phlogopite data lie on a mixing linebetween the pure phlogopite and a low Rb/Sr compositionsimilar to that used to calculate ages, e.g., compare the coarsephlogopite fractions for I.29, which all have essentially thesame ages but very different Rb/Sr ratios.

Mineral ages are calculated according to convention usingwhole-rock or calcite data to estimate initial87Sr/86Sr ratios.Mineral-calcite ages are considered more meaningful, as thecalcite represents the largest reservoir of Sr in these rocks. Formicas and K-feldspar, the high Rb/Sr means that the choice ofinitial ratio hardly affects the calculated age, so that agescalculated using either whole-rock or calcite data are identicalwithin error (Table 6). Indeed, for the mica data the Rb/Sr is sohigh that using any of the whole-rock or calcite data wouldmake little difference in the calculated age. The K-feldspar age

Table 5. Average composition of phlogopite analysed by electronmicroprobe (EMP) and ion microprobe (IMP).

Phlogopite(EMP) (sn21)

Phlogopite(IMP) (sn21)

SiO2 41.01 0.85 41.08a

TiO2 0.56 0.14 0.56a

Al2O3 14.21 0.99 14.21a 0.00FeO 2.10 0.17 2.10a

MgO 25.05 0.67 24.60 1.04MnO 0.02 0.01CaO 0.07 0.07 0.02 0.00BaO 0.03 0.04 0.01 0.00Na2O 0.11 0.03 0.09 0.01K2O 10.13 0.18 10.11b 0.29Rb2O — 0.08c 0.02Total 93.30 0.29 92.98 0.80Sr (ppm) —d 0.44 0.11Ba (ppm) 243 397 110 31Rb (ppm) —d 767 142Formula per 22 (O):Si 5.867 0.097 5.899 0.060Ti 0.060 0.015 0.060 0.001Al 2.397 0.175 2.406 0.025Fe 0.251 0.021 0.252 0.003Mg 5.343 0.131 5.269 0.172Mn 0.002 0.001 —Ca 0.011 0.011 0.004 0.001Ba 0.002 0.002 0.001 0.000Na 0.031 0.009 0.025 0.003K 1.849 0.039 1.852 0.055Rb — 0.008 0.001S Si, Ti, Al 8.325 0.097 8.365 0.086S Ca, Ba, Na, K, Rb 1.892 0.042 1.889 0.056S cations 15.813 0.046 15.779 0.071Al IV 2.072 0.091 2.041 0.061AlVI 0.325 0.097 0.365 0.086S (VI) 5.919 0.028 5.886 0.088% end-member:Annite 4.2 0.4 4.3 0.1Phlogopite 90.3 1.9 89.5 1.7Al biotite 5.5 1.7 6.2 1.6n 14 49

a Not measured, average value by EMP used to calculate total.b Value adjusted to be comparable to EMP value.c Value adjusted to be comparable to value measured by isotope

dilution.d Also measured by isotope dilution on bulk separates (see Table 6).

Similar phlogopites from the Nare`t area were found to contain up to 33mol.% fluorphlogopite (H. Townley, private communication).


is less robust, and is only reasonable if the host calcite orwhole-rock data are used to calculate its age (Fig. 3).

Mineral ages decrease in the order K-feldspar. musco-vite . phlogopite, consistent with the closure temperatures for87Sr calculated or assumed for these systems. For a cooling rateof 32°C/Ma, 1 to 0.5-mm-diameter K-feldspar in I.28 wouldhave a Dodson closure temperature for Sr of 532°C to 492°Cusing the diffusion data of Giletti (1991). Muscovite Rb-Srclosure temperature is generally assumed to be;500°C. Theage of;20 Ma for I.28 muscovite is just below the range of21.1 to 24.9 Ma given by Steiner (1984) for muscovite from thebasement gneisses within 2 km.

Phlogopite ages show a rather larger range (13.9–18.1 Ma)compared with biotite Rb-Sr and K-Ar ages from the adjacentgneisses (15.8–17.3 Ma; Steiner, 1984). Crystal chemical con-siderations (Dahl, 1996) suggest that the Mg- and F-rich phl-ogopite in the metacarbonates would tend to have higher Arand Sr closure temperatures (and therefore ages) than biotite inthe basement gneisses, which is likely to be more Fe-rich andF-poor. Experimental studies of Ar diffusion do indicate higherclosure temperature in phlogopite compared to biotite (curveson Fig. 4a). In contrast, the Sr closure temperature in biotiteappears to be slightly higher than that in phlogopite (Fig. 4a),but both estimates are semi-empirical, and no significanceshould be attached to this difference. However, the Rb-Sr dataon Figure 4a lie close to these curves, suggesting that they maybe broadly correct.

An alternative model to the Dodson infinite reservoir as-sumption for closure temperature is the closed-system Rb-Srmodel of Jenkin et al. (1995). This predicts that the closuretemperature for a mineral pair will be intermediate between theDodson closure temperatures of the two minerals involved,depending on the proportions and Sr contents of the two phases.For a mica-calcite pair, the closure temperature would onlyequal the mica Dodson closure temperature at a small mica/calcite ratio and low [Sr]mica/[Sr]calcite. In this marble samplethese conditions are probably well-satisfied, but if not, theeffect of assuming a closed-system model would be to slightly

raise phlogopite closure temperature values towards a calciteclosure temperature corresponding to its grain size (Fig. 4a).

Although the mineral ages are broadly consistent with thosedetermined by earlier work, it is clear that the phlogopite showsa significant range of ages, both between samples only 15 mapart, and within individual samples. Such a range of ages isinconsistent with the assumption that the closure temperature ofthe Rb-Sr system in biotite is 300°C, or indeed any singletemperature. Part of the variation may be explained by varia-tions in grain size, because Dodson (1973) predicts that closuretemperature and mineral ages should decrease with mica grainsize (Fig. 4a). The 2.4 Ma decrease in age for the convention-ally separated I.28 phlogopitescould be consistent with thedecrease in grain size, if the minimum grain sizes of theseparates apply (Table 6). In contrast, the carefully separatedand sized phlogopites from I.29 do not show a simple decreasein age with grain size parallel to the curves on Figure 4a formodel systems. Coarsest grains give an age of;18 Ma, and theage decreases to 16.6 Ma in the 0.75 to 0.5-mm fraction, similarto the decrease predicted by the Dodson equation, but then,however, the ageincreasesagain to 17.6 Ma in the 0.5 to0.25-mm fraction (Fig. 4a). Differences between ages are sig-nificantly larger than analytical uncertainty. Another explana-tion must be sought for this variation in age.

The anomalously old age of the finest phlogopite fractioncorrelates with a 15% lower Rb content compared with thecoarsest fraction (Fig. 4b), confirmed also by the Rb content ofthe 160 to 120mm conventionally prepared fraction (Table 6).The effect of this has been to decrease the amount of87Rb morethan the decrease in87Sr resulting from the decrease in grainsize, so that some of the87Sr has become “orphaned” and theapparent age actually increases. This can only have happenedduring cooling, because loss of Rb prior to cooling would resultin reduced87Sr production during cooling and no effect on theage. Late loss of Rb due to weathering can be excluded as thesample is very fresh. Moreover, we show in Section 9.2 thatRb-loss in the phlogopite is one of the many retrograde ex-change processes operating in this rock.

Fig. 3. Isochron plots of Rb-Sr data given in Table 6. Representative calcite-mineral isochrons are shown for eachsample.


Tab

le6.

Min

eral

and

who

le-r

ock

Rb-

Sr

isot

ope

data

(no

blan

kco

rrec

tion)

.

Sam

plea

Rb

(ppm

)S

r(p

pm)

Rb/

Sr

87 R

b/8

6 Sr

87 S

r/86 S

r

Age

(Ma)

WR

87 S

r/86 S

r i

Age

(Ma)

Cc

87 S

r/86 S

r i8

7 Sr/8

6 Sr 2

05

87 S

r/86 S

r 24

5

I.2

82

28

to2

10

mm

Ph

0.5–

2m

mb12

051.

824

660.

720

011.

1875

266

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8.2. Ar-Ar Ages

The Ar-Ar ages show a range from 15.5 to 17.3 Ma (Table7; Fig. 4a). This may partly relate to the presence of inclusions,to embayed shapes, to the presence of minor muscovite, or tosubgrains not identified when selecting grains for analysis. Forexample, grains 16, 20, and 21 have the three lowest ages of the.0.75 mm fraction, and they were also the only grains ob-served to contain inclusions (Table 7). This might be explainedby the inclusions reducing the effective grain size for40Ar

diffusion. Strongly embayed grains have relatively young agesfor theirb value, which might be explained in the same way. Inaddition, blank corrections are most significant for small grainsanalysed individually. However, simply comparing the ages ofthe largest 3 grains (16.956 0.21 Ma) with the 8 aggregates of20 grains of 0.5- to 0.25-mm-fraction phlogopites (16.0760.12 Ma), there appears to be a decrease in age with grain sizesimilar in magnitude to that expected for a Dodson infinitereservoir model (Fig. 4a), as has been observed in other Ar-Arstudies (Wright et al., 1991; Hess et al., 1993). The data plotbelow the intermediate biotite curve of Grove and Harrison(1996) and far below the F-rich, near end-member phlogopitecurves of Giletti (1974) and Hammouda and Arnaud (2000).This suggests that the40Ar diffusion rate in the phlogopiteshere is closer to that measured by Grove and Harrison, in linewith their suggestion that the diffusion rate is not a linearfunction of Mg/Fe, but only decreases very close to end-member fluorphlogopite. The fact that the data plot below theGrove and Harrison curve could be attributed to a number offactors:

1) The Grove and Harrison (1996) diffusion parameters under-estimate the diffusion rate.

2) All grains in this study having an effective grain radius forAr diffusion of ;50 to 150mm, similar to the effectivediffusion radius of 150mm inferred for biotite by Harrisonet al. (1985).

3) The presence of high diffusivity pathways (Lee, 1995), ormultiple diffusion domains (Lovera et al., 1989) in oursamples.

4) Inaccuracy in the position of the closure temperature curvedue to the assumptions made about cooling rate and thermalhistory of the sample.

The last is perhaps most likely; it is reiterated that the estimatedlocal cooling history is itself dependent on assumptions aboutmica closure temperature, and what is most relevant here is theform of these curves, rather their absolute position.

The important observation here is that the Ar-Ar ages do notshow the peculiar increase in age of the finest grain-size frac-tion shown by the Rb-Sr ages, and could conform to a Dodsoninfinite reservoir model. Compared to the Rb-Sr system, aninfinite reservoir model for closure temperature is likely to bemore realistic for40Ar in micas, (except in highP terranes;Kelley et al., 1994), because40Ar is not bonded in mineralstructures or strongly partitioned into them. Once40Ar hasdiffused to the grain boundary there is little to stop it diffusingout of the rock altogether. The phlogopite Ar-Ar ages do notappear to have been affected by chemical exchange effects;both EMP and IMP analyses demonstrate that the K contentvaries by,2% across the grains.

8.3. Calcite Sr-Isotope Traverse

The phlogopites have Rb-Sr ages of;18 Ma, yet the peak ofmetamorphism was at;25 Ma, and the phlogopite formedbefore this on the prograde path. Therefore it is implied that, formore than 7 Ma, the87Sr produced by decay of87Rb (87Sr*)was lost from the phlogopites. The model of closure tempera-ture appropriate for these phlogopites depends on the transportmechanism of the87Sr* and the reservoir that assimilated it.

Fig. 4. Rb-Sr and Ar-Ar data for sized I.29 phlogopites from disso-lution experiment. (a) Grain radius versus age. (b) Grain radius versusRb concentration. Rb-Sr data are for grain size fractions, Ar-Ar data arefor individual grains (single points) or splits of the 0.5–0.25-mm sizefraction (with error bars). Grain radius refers to measuredb/2 forindividual grains analysed by Ar-Ar and the weighted meanb/2 for sizefractions measured for Rb-Sr and Ar-Ar (Section 4.2). Error bars showthe total range of grain size in size fractions. Curves on 4a show biotite(Bi), phlogopite (Ph) and calcite (Cc) closure temperatures and biotiteand phlogopite closure ages as a function of grain size for a Dodsoninfinite reservoir model for various estimates of Ar and Sr diffusion.Dotted curve for calcite assumes a spherical model. Solid curves forbiotite and phlogopite are for an infinite cylinder model; fine dashedcurves are for rectangular grains with shapea 5 2b, for which the Aterm in the Dodson equation has a value of 14.8 (R. Holme, privatecommunication). Comparing these two curves shows that the effect ofassumed diffusion geometry is minor. Data sources: Sr Ph; Hammoudaand Cherniak (2000), Ar Ph1; Giletti (1974), Ar Ph2; Hammouda andArnaud (2000), Sr Cc; Cherniak (1997), Ar Bi; Grove and Harrison(1996), Sr Bi; Jenkin (1997). All data are based on experimentalmeasurements, although the Sr diffusion data for biotite and phlogopitealso involve major assumptions and should only be viewed as semi-empirical. Closure age (tc) is related to calculated closure temperature(Tc) via the temperature and time of peak metamorphism (Ti 5 575°C,t i 5 25 Ma) and the inferred local cooling rate (dT/dt 5 328C/Ma) bytc 5 t i 2 [(Ti 2 Tc)/(dT/dt)]. Below ;300°C the cooling rate isinferred to drop to about 15°C/Ma, the effect of which would be todecrease closure temperatures#300°C by 10–20°C. Because the in-ferred local cooling history is itself partly dependent on assumptionsabout the closure temperature of micas, the position of these curves canonly be considered approximations; the generalform of the curves ismore relevant here.


For the Dodson (1973) infinite reservoir model the87Sr* isassumed to be transported either out of the rock completely, orinto a reservoir in the rock that is large enough to assimilate itwithout its composition being significantly affected. In theclosed-system exchange model (Jenkin et al., 1995) the87Sr* istransported into other minerals in the rock, but the compositionof these reservoirscan be significantly affected. The calcite isthe obvious candidate as the major sink for the87Sr* from thephlogopites, as it is by far the largest reservoir of Sr in the rock.Therefore, to test this hypothesis, calcite Sr isotope composi-tions were measured in a traverse running across and perpen-dicular to the phlogopite band.

Calcite fractions were separated from the whole rock bydissolution in ethanoic acid (Sections 4.3, 4.5). Calcite87Sr/86Sr is found to be highest in the phlogopite-rich band(;0.7154) and decreases away from it over a distance of 4 cminto the marble down to;0.7122 (Table 8; Fig. 5). At thisdistance the decrease in87Sr/86Sr appears to flatten out. Wheremeasured (Table 8), the Rb concentrations, are significantlyhigher than in the conventional calcite separates from the samearea of rock (Table 6), although still low. This effect occurs in

both the phlogopite band and outside it, and must be due eitherto fine Rb-rich silicate being dissolved in the calcite fractions,or to Rb being leached from the silicates by the ethanoic acid.Provided that Sr is not preferentially leached relative to Rbfrom the silicate material, the effect of incorporating minor Rband Sr from the silicates into the calcite fraction is small; usingthe Rb contents to correct to;18 Ma (zero age for the phl-ogopite) makes a negligible difference to the calcite87Sr/86Sr(Table 8; Fig. 5). The high87Sr/86Sr in the phlogopite bandcalcite is therefore a real effect. Taken at face value, these datasuggest that87Sr* lost from the phlogopites has been trans-ported into the surrounding calcite.

The fact that the calcite87Sr/86Sr has been changed byexchange with the phlogopites indicates that a Dodson infinitereservoir assumption is not valid here for the phlogopite. Themechanism by which87Sr* might have been transported overthis distance into the calcite is not determined by the data, butcurrent models (e.g., Jenkin et al., 1995) would invoke rapiddiffusion along the grain boundary network and slower volumediffusion into the calcite. Some evidence for this may be givenby the calcite mineral separate data. The87Sr/86Sr of the

Table 7. Argon isotope data.

Grain#

aa

(mm)b

(mm)g

(mm)wt

(mg)age

(Ma)1s

(Ma) Commentsb

.0.75-mm sieve1 2.5 1.5 0.25 2.52 16.93 0.04 Composite?2 1.5 1.2 0.25 1.23 16.86 0.063 1.8 1.6 0.20 1.38 16.60 0.05 Embayed4 1.7 1.7 0.15 0.66 17.32 0.105 1.6 0.8 0.15 0.83 17.12 0.116 1.0 0.8 0.10 0.38 16.33 0.05 Embayed7 2.1 1.2 0.35 2.03 16.75 0.08 Inclusion

11 1.5 0.9 0.20 0.92 16.23 0.17 Embayed13 1.1 1.1 0.20 0.99 16.46 0.0614 1.2 1.1 0.15 0.52 16.95 0.1615 1.1 1.0 0.15 0.65 16.86 0.11 Some musc?16 1.6 1.2 0.15 0.62 16.07 0.11 Inclusions20 1.1 0.5 0.10 0.27 15.85 0.06 Inclusion21 1.2 1.2 0.20 0.62 15.84 0.10 Inclusion0.5–0.75-mm sieve

24 1.0 0.6 0.10 0.27 16.60 0.0725 1.1 0.5 0.20 0.41 16.70 0.06 Some musc?26 1.5 0.9 0.10 0.58 16.65 0.0629 0.8 0.6 0.03 0.13 16.93 0.0931 1.2 0.7 0.10 0.51 16.71 0.0432 1.0 0.4 0.15 0.31 16.30 0.0734 0.8 0.8 0.10 0.40 16.91 0.0835 1.2 1.0 0.10 0.40 17.33 0.0540 0.8 0.7 0.07 0.12 16.02 0.0944 1.0 0.6 0.07 0.17 17.18 0.200.25–0.75-mm sievec

A 0.67 0.54 0.16 0.03 15.89 0.03B 0.67 0.54 0.16 0.03 16.04 0.04C 0.67 0.54 0.16 0.03 16.17 0.04D 0.67 0.54 0.16 0.03 16.24 0.07E 0.67 0.54 0.16 0.03 16.62 0.04F 0.67 0.54 0.16 0.03 16.31 0.33G 0.67 0.54 0.16 0.03 15.54 0.15H 0.67 0.54 0.16 0.03 15.76 0.06

a See Section 4.2.b Grains may be composite, contain inclusions of calcite or apatite, show pronounced embayments, or contain minor muscovite.c Fractions of;20 grains analysed together. Grain sizes given are weighted mean of 22 randomly sampled grains. Note that these are biased by

one very large grain; excluding this graina, b, andg are 0.55, 0.41 and 0.08 mm respectively. Grain weight is the mean value.


separate215 to 230 mm away from the phlogopite bandagrees well with the traverse samples, but the separate in thephlogopite band has a distinctly lower87Sr/86Sr (Fig. 5). Anexplanation of this, other than sample inhomogeneity within theband, is that mineral separates tend to be biased towards graincore compositions—fragments of grain rims, often compositewith other grains, are discarded during separation—so that thegrain cores could have lower87Sr/86Sr. If so, in comparisonwith the bulk drilled calcite, this could indicate that the calcite

grains are zoned in87Sr/86Sr from relatively low 87Sr/86Srcores to higher87Sr/86Sr rims, with rim values higher than thebulk values. Such an isotopic zonation is what would be ex-pected if there was fast equilibration along the grain boundariesand slower equilibration by volume diffusion into the calcitegrains.

Calcite Sr concentration was measured by isotope dilutionduring Sr isotope analysis (Fig. 5) and large variations werefound along the traverse, from low values of 130 ppm in thephlogopite band, increasing away from it up to;185 ppm inthe marble. The variation in Sr concentration is inversely pro-portional to87Sr/86Sr and occurs over the same length scale,implying a common transport mechanism produced these gra-dients. Again, there is a discrepancy between the mineralseparate and drilled whole-rock samples in the phlogopite band.Using the same argument as that used to explain the discrep-ancy in87Sr/86Sr data, this might be interpreted to mean that thecalcite is also zoned in Sr content, from high-Sr cores to low-Srrims. If such a gradient existed it could suggest movement of Srout of the calcite at the same time as inward gradients in87Sr/86Sr were being developed. Although models of Rb-Srclosure assume87Sr* mobility, none incorporate the effects ofsimultaneous mobility of other Sr isotopes, such that changes inSr concentration occur. We investigate further evidence for thisbelow.

9. ION MICROPROBE TRAVERSES—FURTHEREVIDENCE FOR RETROGRADE EXCHANGE INVOLVING

Rb AND Sr

9.1. Calcite and Dolomite

Dolomite is enriched in Mg, Fe, and Mn and depleted in Srand perhaps Ba compared to calcite (Table 2; Fig. 6). Average

Table 8. I.29 Calcite Sr isotope traverse data.

SampleaSr

(ppm)Rb

(ppm) Rb/Sr 87Rb/86Sr 87Sr/86Sr 87Sr/86Sr18Ma

231a 175.7 2.053 0.0117 0.0338 0.712796 3 0.712786 3225a 164.0 15.43 0.0941 0.2724 0.713196 6 0.713126 6219a 167.5 0.713646 329a 159.9 0.714096 325a 165.5 0.714296 321a 156.9 1.403 0.00894 0.0259 0.714466 3 0.714456 311a 144.5 22.10 0.1529 0.4428 0.714966 4 0.714856 415a 137.2 20.16 0.1469 0.4254 0.715306 7 0.715196 719a 137.9 12.12 0.0878 0.2543 0.715016 3 0.714956 3115a 132.7 0.715216 4119a 131.3 0.715766 2125a 147.2 0.714976 2131a 151.1 0.714786 3244b 185.3 0.712306 3239b 183.3 0.712296 2235b 182.4 0.712276 2231b 187.6 0.712316 2227b 178.0 0.712416 2131c 156.1 0.713906 2141c 157.4 0.713506 2150c 163.5 0.713166 2159c 155.0 0.712926 2

Calcite was selectively extracted from the whole-rock samples with ethanoic acid.a The sample number refers to mm from the datum, the letter a, b, or c, refers to offsets in the traverse parallel

to the phlogopite band—see Fig. 1.

Fig. 5. Sr isotope data for calcite from drilled small samples in atraverse across the phlogopite band. For sampling locations see Fig. 1.Gaps in the traverse mark offsets parallel to the phlogopite band.Calcite was selectively extracted from the whole-rock samples withethanoic acid. Mineral separate data are for calcite separated by con-ventional methods from the distances shown in the;2-cm-thick slabopposite to the surface shown in Figure 1.


ratios of concentration [ ]dolomite/[ ] calcite measured by IMPare approximately Mg5 12, Fe5 5, Mn 5 3, Sr5 0.5, Ba50.4. Dolomite rims are slightly more Mg-rich than cores (47.8mol.% MgCO3, compared to 47.0%; Fig. 6), as would beexpected if the rims record Ca-Mg exchange with calcite tolower temperature. Fe and Mn increase systematically from rimto core, but this may not necessarily imply that the dolomitewas losing Fe and Mn when these gradients were frozen in by

decreasing temperature. Both Fe and Mn are expected to bepreferentially concentrated in dolomite relative to calcite withfalling temperature (Anovitz and Essene, 1987; Hickmott et al.,1997; G. Roselle, pers. commun., 1999) so that these gradientscould represent “uphill diffusion” gradients resulting from mul-ticomponent exchange on cooling. Alternatively, Mn and Fecould be being preferentially partitioned into another phase,such as phlogopite.

Fig. 7a,c show IMP traverses across orange-luminescingcalcite grains, mostly free from dolomite exsolution. The fol-lowing observations can be made:

1) Sr concentration shows a smooth zonation, varying from;150 ppm in the grain cores, down to,100 ppm at grainboundaries, confirming indications of zoning from the tra-verse samples (Section 8.3). Cusp-shaped profiles are ob-served across calcite-calcite grain boundaries suggestingthat the boundaries represented sinks for Sr during cooling.Sr also decreases towards calcite-quartz and calcite-phlogo-pite grain boundaries, indicating that these were similarlyefficient Sr transport routes. In contrast, no deviations in theprofiles are apparent in relation to twin boundaries andfractures, suggesting that these structures were not impor-tant transport pathways for Sr. The smoothness of the pro-files in these well-preserved areas could be interpreted asthe result of, but does not demand, a volume diffusionprocess.

2) Other trace and major elements also vary in concentrationacross the traverses in relation to distance from the grainboundaries. Ba varies sympathetically with Sr. Mg and Fealso tend to decrease towards the grain boundaries, but inFigure 7a highest values are not in the grain core butbetween the core and the rim. In Figure 7c Mn also de-creases towards the grain boundaries, but in 7a this trend

Fig. 6. Ion microprobe traverse across a coarse dolomite grain incalcite from the phlogopite-rich band. Almost identical patterns werefound in a traverse into another dolomite grain.

Fig. 7. Ion microprobe traverses across calcite grains: (a) Clear (orange-luminescent) calcite shown in Figure 2c. (b)Adjacent cloudy (bright yellow-luminescent) calcite. Note that orientation is reversed so traverse crosses grain boundary (A)in the same orientation as in Figure 7a. (c) Orange-luminescent calcite with fractures shown in Figure 2a.


appears to be reversed. Decreases in Mg and Fe towards thegrain boundaries would be expected if the rims recordexchange with dolomite to lower temperatures (Anovitz andEssene, 1987); highest Mg contents in Figure 7a,c translateto calcite-dolomite equilibration temperatures of 500°C and486°C, respectively, lowest Mg contents at the grain bound-aries translate to temperatures of 463°C and 455°C, respec-tively. Mn would be expected to decrease towards grainboundaries by exchange with dolomite to low temperatures,as in Figure 7c. The opposite behaviour of Mn in Figure 7a,could be a multicomponent effect.

3) The profiles are asymmetric in shape. This is expected sincethe traverses run through grains of different sizes and areunlikely to run perpendicular to the grain boundaries. Fur-thermore, it is not known whether all grain boundaries are ofthe same width, or how close the traverse is to grain bound-aries both above and below the plane of the section. Theinflections in Mg and Fe (and possibly Mn) profiles withinthe central grain in Figure 7a could be interpreted as growthzoning that was modified at the rims by retrograde exchangeduring cooling, or just reversals in the direction of Mg andFe transport during cooling. However, we do not believethat growth zoning will be preserved in the calcite (Section10) and a more likely explanation of these inflections, whichare only observed in one grain, is that they are the result ofinterference from another grain boundary out of the plane ofthe section. The different shapes of the profiles for differentelements will be due to the combination of the differentdiffusion and partition coefficients for each element and thereservoirs involved. Nevertheless, the length scales of pos-sible diffusive exchange are indicated.

4) In areas of fine exsolution (Fig. 7b) it is notable that thezoning of Sr and Ba is still observed. If it is assumed that thecation zonations across calcite grains were dominantly de-veloped at 450°C to 500°C (calcite-dolomite thermometry),and the exsolution was developed at 300°C to 400°C (ther-mometry and relationship to twinning), then zoning devel-oped before exsolution. The preservation of Sr and Bazoning, despite the fine exsolution, seems therefore to bedue to the fact that they are not strongly partitioned betweencalcite and dolomite.

9.2. Phlogopite

Isotope dilution measurements of phlogopite separates indi-cated lower Rb contents in the finer-grained phlogopite thatappeared to be due to Rb-loss during cooling. IMP profilesacross phlogopite grains confirm this, showing that althoughthey are homogeneous in major element composition, Rb con-tents are highest in the cores of large grains, but decreasetowards the rim, and are low throughout in small grains (Fig.8). Variations in Ba are complimentary to Rb, suggestingBa-gain during cooling. The relatively smooth shapes of theprofiles again are consistent with a volume diffusion process.

The pronounced anisotropy in the mica lattice is thought tomake diffusion parallel to the cleavage;3 orders of magnitudefaster than diffusion parallel to thec-axis (Hofmann et al.,1974; Fortier and Giletti, 1991). Thus it would be expected thatRb exchange would dominantly occur in the cleavage direction

and be negligible parallel toc. In apparent contradiction to this,Fig. 8a,b show similar decreases in Rb towards the edge of thegrain in traverses parallel toc and parallel to the cleavage.However, it is stressed once again that we do not know theshape of the grains out of the plane of the section, and thelow-Rb rims measured parallel toc in this instance could beexplained by the closeness of the grain boundary along thecleavage direction above or below the plane of the section.

9.3. K-feldspar

IMP and EMP profiles across K-feldspar grains show K-richrims, developed by low temperature exchange with plagioclase.IMP profiles also show gradients of increasing Ba, Rb, and Srtowards K-feldspar rims suggesting that it was a sink for theseelements down to low temperatures (Fig. 9). In contrast, Ca andNa decrease towards the rims indicating, that these are beinglost, as would be expected due to exchange with plagioclase.Albite generally contains low Ba, Rb, and Sr concentrations.

10. DISCUSSION

It has been shown that calcite, dolomite, phlogopite, andK-feldspar show zonations in trace and major elements. Thesecannot be premetamorphic or prograde (growth) gradients be-cause:

a) They would have been destroyed by the intense deformationconcentrated in these rocks; the current texture is due tocontinued mineral growth at peak temperatures after perva-sive deformation.

b) Peak temperature would have been high enough, and per-sisted long enough, to have annealed out pre-existing gra-dients by diffusion according to available diffusion data(e.g., see calcite curve in Fig. 4), except perhaps Rbgradients (and by inference Ba) in K-feldspar grains,which are inferred to have closure temperatures. 600°Cfor grains. 0.5-mm diameter (Giletti, 1991).

c) Rb gradients in phlogopite affect the Rb-Sr age and there-fore must have been developed as Sr-exchange was haltingduring cooling.

It is therefore concluded that these gradients were developed

Fig. 8. Ion microprobe traverses across phlogopite grains: (a) Parallelto cleavage along the centre of a large phlogopite grain. (b) At rightangles to cleavage across the same grain; dashed line marks the pointof intersection of the two profiles. (c) At right angles to cleavage acrossa small grain only;150-mm-wide parallel to the cleavage.


during cooling of the rock. This is indicated by some cationsthat show these gradients also giving subpeak temperatures bycalcite-dolomite (450–500°C) or feldspar (300–400°C) solvusthermometry.

The cation zonations observed indicate a complex pattern ofchemical re-equilibration during cooling. The multicomponentnature of some exchange reactions means that it is possible thatsome “uphill” diffusion gradients could be developed. How-ever, assuming that in most cases, the direction of transport ofa cation is down its concentration gradient, then the observedgradients indicate the following movements: Mg, Fe, and prob-ably Mn, are lost from calcite and gained by dolomite, beingbalanced by Ca moving in the opposite direction; Ca and Na arelost from K-feldspar and gained by plagioclase, being balancedby K moving in the opposite direction; Ba is lost from calciteand gained by K-feldspar and phlogopite; Sr is lost from calciteand gained by K-feldspar; and Rb is lost from phlogopite and

also gained by K-feldspar. There may be exchanges of othercations and anions which were not analysed, as well as ex-changes involving other phases not analysed, such as apatite.Such chemical partitioning as a function of temperature is notunexpected; indeed many geothermometers depend on it. Norshould it be surprising that exchange takes place to low tem-peratures during cooling—the relatively fast diffusion rate ofthe alkali metals and alkali earth cations, especially the smallerions, is well known. Qualitatively the exchanges observed areconsistent with closed-system exchange. To prove this wouldrequire quantitative balancing of fluxes in and out of mineralsin terms of the inferred changes in composition of each mineraland their modal proportions, followed by a demonstration thatthese fluxes are consistent with the known partitioning data asa function of temperature, and with the diffusion data, whichcontrol how much of a mineral is actually available for ex-change at a given temperature. This is beyond the scope of thispaper, but forms the focus of ongoing study.

Whereas cation exchange during cooling may not be remark-able in itself, what is of great importance here is that elementsinvolved in dating schemes (K, Rb, and Sr) are taking part inthese exchange processes, and this has influenced mineral ages.These elements show zonations across grains which developedwith falling temperature and evolving time, yet all but onemodel of closure temperature for geochronological systemsassume that parent and daughterelementconcentrations arehomogeneous across grains, and do not change with time;usually onlyisotopicexchange is modeled. The exception is thetheoretical study of Ganguly and Ruiz (1987), who suggestedthat the cessation of cation exchange is what actually controlsclosure temperature in the Rb-Sr system. Although our resultssupport that suggestion, we show that in addition isotopicexchange also clearly occurred in this rock, and appears to bedecoupled from chemical exchange.

An example of the decoupling of chemical and isotopicexchange is given by the Sr data for the calcite. IMP profilesshow Sr decreasing from core to rim, suggesting the flux of Sris out of the grains. In contrast, it is inferred that87Sr/86Srdecreases from rim to core, suggesting a flux of87Sr into thegrains. Similar opposed gradients are seen on a larger scale inthe calcite traverse across the phlogopite band. The decrease inSr concentration towards the phlogopite band by as much as30%, compared with the87Sr/86Sr increase of only 0.4% in thesame direction (Fig. 5), means that the actual gradient in87Srconcentration along the traverse is downwards towards thephlogopite band, implying that87Sr is actually flowingtowardsthe phlogopite band! The decrease in87Sr/86Sr away from thephlogopite band may therefore not be due to a flow of87Sraway from the phlogopite, but to the decrease in87Sr towardsthe phlogopite being less than the decrease in the other Sr-isotopes. The calcite is therefore moving towards isotopicequilibration with the phlogopite, yet on a macroscopic scale itis losing87Sr. For diffusion of a trace element such as Sr, eachindividual atom will diffuse independently of any others by arandom walk. Thus some atoms of87Sr* released from thephlogopite will have moved into the calcite, but on average themacroscopic flux of87Sr will be out of it. On a grain scale,especially in those calcite grains actually touching the phlogo-pite, the87Sr/86Sr gradient may have been high enough thatthere was bulk flow of87Sr into the calcite, but it is possible

Fig. 9. Ion microprobe traverses in K-feldspar grains: (a) AcrossK-feldspar with small albite grain at edge at152 mm relative to datum.The first point in albite is probably a mixed analysis of the twominerals. Background shading shows the variation in major elementcomposition in terms of the feldspar end-members Ab, An, and Or (lefty-axis). (b) Micro-traverse taken at 1.8-mm steps towards the edge of aK-feldspar nearby to that in (a). Typical errors on these curves are64%. For comparison, changes in concentrations are shown relative tothe first point. The Ca profile stops before the others due to interferencefrom the calcite host.


that another sink is required for the87Sr*. This may have beenthe K-feldspar, which appears to have assimilated Sr withcooling.

11. CONCLUSIONS

Closure temperature in the mica Rb-Sr system is shown to bemuch more complicated than implied by simple infinite reser-voir (Dodson) models, being the result of both isotopic andchemical exchange with finite reservoirs in the surrounding fewcm of rock. This exchange was probably achieved by volumediffusion within grains and fast grain boundary diffusion be-tween grains. Isotopic exchange will therefore depend on: a)mineral grain sizes, both of sources and sinks; b) tracer87Srdiffusion coefficients in each mineral containing it; c) themodal proportions of each mineral and their Sr contents. Thesefactors are already incorporated in the model of Jenkin et al.(1995). Chemical exchange will also depend on mineral grainsizes, modal proportions, and the Sr and Rb contents of eachphase involved. Diffusion coefficients of Sr and Rb in eachparticipating phase will control the rate of chemical equilibra-tion. Finally, Rb and Sr partitioning data will control thedirection of chemical exchange among minerals during cooling.There is no reason why the factors described above should notalso apply in the granitic gneisses and pelitic metasedimentsmore typically used for biotite Rb-Sr dating, and investigationof these rocks for similar effects is currently underway.

Taking all of the above factors into account to calculate aclosure temperature will be exceedingly complicated, and allthe data required may not be currently available. It is clear,however, that the actual Rb-Sr age recorded by a mineral willdepend on the relative rates of gain or loss of Rb, Sr, and87Sr,to or from the mineral during cooling. For example, the K-feldspar examined in this study appears to be gaining both Rband Sr, as well as probably87Sr* from the phlogopite, but themeaning of the age recorded is currently obscure withoutknowing the relative importance of these effects.

The complexity revealed above for the biotite Rb-Sr systemmay explain why, to the best of our knowledge, no convincingage-grain size relationships have been reported for this system,in contrast to the simpler K-Ar mica system where such trendshave been identified (Wright et al., 1991; Hess et al., 1993).

Acknowledgments—This work was carried out while the senior authorwas a Royal Society of Edinburgh/Scottish Office Education Depart-ment Research Fellow. The ion microprobe work was funded by NERCScientific Services (IMP/86/0196). Thanks go to the following: CalumMacaulay and Zachary Sharp for assistance in the field; Terry Donnellyfor the stable isotope analysis of the small carbonate samples; VincentGallagher and Anne Kelly for excellent support for the Rb-Sr work;Richard Holme for calculating the Dodson A coefficient for rectangularplates; Tony Fallick for allowing access to all facilities at SUERC;Peter Hill, Paula McDade, and John Craven for excellent technicalsupport on the electron and ion microprobes. Sidney Hemming, DerekVance and Jibamitra Ganguly are thanked for their positive reviews.The ion microprobe is supported by NERC Science Programmes Di-rectorate. SUERC is funded by NERC and a consortium of ScottishUniversities.

Associate editor:S. L. Goldstein

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