Polyphase zircon in ultrahigh-temperature granulites (Rogaland,SW Norway): constraints for Pb diffusion in zircon

A. MOLLER1,* , P . J . O’BRIEN2, A. KENNEDY3 AND A. KRONER1

1Institut fur Geowissenschaften, Universitat Mainz, Postfach 3980, 55099 Mainz, Germany,2Institut fur Geowissenschaften, Universitat Potsdam, Postfach 601553, 14415 Potsdam, Germany,3Curtin University, Department of Applied Physics, Bentley 6102 WA, Australia

ABSTRACT SHRIMP U–Pb ages have been obtained for zircon in granitic gneisses from the aureole of the Rogalandanorthosite–norite intrusive complex, both from the ultrahigh temperature (UHT; >900 �C pigeo-nite-in) zone and from outside the hypersthene-in isograd. Magmatic and metamorphic segments ofcomposite zircon were characterised on the basis of electron backscattered electron and cathodolumi-nescence images plus trace element analysis. A sample from outside the UHT zone has magmatic coreswith an age of 1034±7 Ma (2r, n ¼ 8) and 1052±5 Ma (1r, n ¼ 1) overgrown by M1 metamorphicrims giving ages between 1020±7 and 1007±5 Ma. In contrast, samples from the UHT zone exhibitfour major age groups: (1) magmatic cores yielding ages over 1500 Ma (2) magmatic cores giving ages of1034±13 Ma (2r, n ¼ 4) and 1056±10 Ma (1r, n ¼ 1) (3) metamorphic overgrowths ranging in agebetween 1017±6 Ma and 992±7 Ma (1r) corresponding to the regional M1 Sveconorwegiangranulite facies metamorphism, and (4) overgrowths corresponding to M2 UHT contact metamorphismgiving values of 922±14 Ma (2r, n ¼ 6). Recrystallized areas in zircon from both areas define a furtherage group at 974±13 Ma (2r, n ¼ 4). This study presents the first evidence from Rogaland for newgrowth of zircon resulting from UHT contact metamorphism. More importantly, it shows the survivalof magmatic and regional metamorphic zircon relics in rocks that experienced a thermal overprint of c.950 �C for at least 1 Myr. Magmatic and different metamorphic zones in the same zircon are sharplybounded and preserve original crystallization age information, a result inconsistent with someexperimental data on Pb diffusion in zircon which predict measurable Pb diffusion under suchconditions. The implication is that resetting of zircon ages by diffusion during M2 was negligible in thesedry granulite facies rocks. Imaging and Th ⁄U–Y systematics indicate that the main processes affectingzircon were dissolution-reprecipitation in a closed system and solid-state recrystallization during andsoon after M1.

Key words: diffusion; Rogaland; SHRIMP ion microprobe; trace elements; UHT granulites; zirconchemistry.

INTRODUCTION

Open system behaviour, or the loss of the products ofradioactive decay by volume diffusion at high tem-peratures, is one of the most important limiting factorsin the reliable interpretation of ages from isotopegeochronology. It is well documented that Rb–Sr,K–Ar and Ar–Ar ages of mica and amphibole in high-grade rocks correspond to the temperature stage on thecooling path where closed system behaviour effectivelybegan, i.e. where diffusive loss of the daughter isotopesbecame negligible. Such so-called �closure� or �blocking�temperatures, an important aid to constraining tem-perature–time paths of natural processes, derive from acombination of diffusion rate, grain size and cool-ing history (e.g. Dodson, 1973, 1986) and have been

estimated for most geochronologically importantminerals. Experimental results on the diffusivity oftrace elements in zircon indicate that Pb diffusion isexceedingly slow even taking into consideration theorder of magnitude difference between the diffusivi-ties given by Lee et al. (1997) and those derived byCherniak et al. (1997a,b) and Cherniak & Watson(2000). The important consequence of exceedingly slowdiffusion rates in pristine ⁄ crystalline zircon is that�blocking� or �closure� temperatures lie outside therealm of crustal metamorphism (i.e. >950–1000 �C:Black et al., 1986; Claoue-Long et al., 1991; Williams,1992) and therefore measured U–Pb ages represent thetime of crystallization of the zircon. However, thecase for a lower �closure temperature� for zircon hasrepeatedly been made in the literature (e.g. Heaman &Parrish, 1991; Ashwal et al., 1999) even though theresistance of zircon to diffusive Pb-loss under extremethermal conditions, such as in xenoliths assimilated inmagmas or in high grade migmatites, has repeatedly

*Now at: Institut fur Geowissenschaften, Universitat Potsdam, Post-

fach 601553, 14415 Potsdam, Germany (amoeller@geo.uni-potsdam.de).

J. metamorphic Geol., 2002, 20, 727–740

� Blackwell Science Inc., 0263-4929/02/$15.00 727Journal of Metamorphic Geology, Volume 20, Number 8, 2002

been demonstrated (e.g. Gulson & Krogh, 1973;Compston et al., 1986; Harrison et al., 1987; Chen &Williams, 1990). Although Pb-loss, and associated agediscordance, is common in metamict zircon due tohydrothermal leaching (Geisler et al., 2001) and partialrecrystallization (for review see Mezger & Krogstad,1997), one of the main causes of discordant datapoints is the simple fact that zircon is often poly-phase. Electron backscatter imagery, cathodolumines-cence spectroscopy and electron-microprobe traceelement analysis (e.g. Sommerauer, 1976; Koppel &Sommerauer, 1974), has revealed that zircon com-monly displays complex growth and overgrowth fea-tures of both magmatic and metamorphic origin.Unless such growth stages are identified and isolated,ages derived by conventional, evaporation or vapourdigestion techniques will always be subject to uncer-tainty. In the present study we present SHRIMP U–Pbages for different zircon zones, identified on the basisof difference in cathodoluminescence and electronbackscatter images, from regionally metamorphosedmagmatic rocks that experienced a very high temper-ature (at least 950 �C) thermal overprint. The intentionis to demonstrate the preservation of age informationand trace-element zoning patterns in zircon with mul-tiple magmatic and metamorphic overgrowths, despiteultra-high temperature metamorphism, thus substan-tiating models of negligible diffusion in nonmetamictzircon, even at extremes of crustal conditions.

GEOLOGICAL SETTING

The investigated area is in the Proterozoic Rogaland-Vest Agder terrane of the Sveconorwegian Province,South Norway (Fig. 1). This part of the Baltic Shield iscomposed of folded units of banded and migmat-itic tonalitic and granitic augen-gneisses containingminor lenses of amphibolite, metapelite, quartzite,

calc-silicate rocks and marble (Hermans et al., 1975).This already folded and metamorphosed crust wasintruded by anorthosites and related rocks forming thec. 1200 km2Rogalandanorthosite–norite intrusive complex.This Rogaland complex is characterised by three

massif-type anorthosite plutons, the Egersund-Ogna(oldest), Ana-Sira and Haland-Helleren (youngest)bodies, intruded by the Bjerkreim-Søkndal layeredintrusion (anorthosite ⁄ leuconorite–norite–gabbronor-ite–mangerite cumulates, topped by quartz mangeritesand charnockites), cross-cut by hypersthene-monzodi-oritic (jotunitic) dykes (Duchesne et al., 1989; Wilsonet al., 1996; Duchesne & Wilmart, 1997). The mainstage of anorthosite plutonism is dated at 931± 2 Ma(zircon and baddeleyite) with ages of 920± 3 Ma fromthe ilmenite ore-body of the Tellnes jotunite dyke, and915± 4 Ma from the foliated margin of the Egersund-Ogna massif marking the end of the emplacementprocess (Scharer et al., 1996). Some zircon fractionsfrom the anorthosite–norite intrusive complex yieldedupper concordia intercept ages of c. 1240, 1450 and1690 Ma (Duchesne et al., 1987; Scharer et al., 1996),pointing to inheritance. The inherited zircon may havebeen derived from the (lower crustal?) sources ofthe magmas, or reflects crustal contamination duringintrusion (Scharer et al., 1996). The inherited ages arecompatible with the Gothian (1.75–1.5 Ga) andSveconorwegian (1.25–0.9 Ga) orogenic events estab-lished for the Baltic Shield (Gaal & Gorbatschev, 1987).A multistage metamorphic history for the gneisses,

best established in the volumetrically minor garneti-ferous metapelite layers, comprises an upper amphi-bolite to granulite facies stage (M1), overprinted athigh temperature granulite facies conditions (M2)and followed by an amphibolite facies overprint (M3)before final retrogression (M4) at greenschist orlower grade (Hermans et al., 1976; Maijer et al., 1981;Maijer, 1987). The M2 granulite facies stage is

Fig. 1. Simplified geological map of theRogaland anorthosite–norite intrusivecomplex with sample locations.

7 28 A . M O L L E R E T A L .

preserved in a 10–30 km-wide aureole around theanorthosite–norite intrusive complex with tempera-tures increasing towards the intrusive complex asreflected in mapped isograds for incoming ortho-pyroxene, osumilite and pigeonite (Fig. 1), with a peakof more than 1000 �C reached close to the contact.Low pressures of c. 4 kbar for the high temperatureM2 stage were deduced by Jansen et al. (1985), basedon osumilite stability and have subsequently beenreinforced by experiments in the osumilite-bearingmetapelite system (Carrington & Harley, 1995) as wellas from melting experiments on samples from thelayered complex (Van der Auwera & Longhi, 1994).Holland et al. (1996) estimated the pressure at5.5± 0.2 kbar at 800–850 �C from an osumilite, gar-net and orthopyroxene assemblage. Thermal modellingshows that the intrusion of later magmas into crustalready heated by earlier intrusions would yield thevery high temperatures recorded in the aureoles(Westphal et al. in press). Isobaric cooling after thisevent is reflected in M3 corona, symplectite and exso-lution textures formed at conditions of 550–700 �C,3–5 kbar (Jansen et al., 1985).Despite the obvious field relationships between

mineral isograds and the intrusive complex, the age ofthe high temperature M2 metamorphism is still indispute. An age of 1159± 5 Ma (U–Pb zircon, Zhouet al., 1995) for the deformed charnockitic gneiss fromHidderskog (60 km SE of Egersund) is a maximumlimit for regional metamorphism, whereas Rb–Srwhole rock isochrons of 980± 14 Ma on the unde-formed Holum granite in the Mandal area (Wilsonet al., 1977), and 998± 14 Ma for the D3 deformedHomme granite in the Flekkefjord area (Falkum &Pedersen, 1979) provide a lower age limit for regionaldeformation and metamorphism. Nijland et al. (1996)interpreted published zircon, monazite and titaniteages from migmatites and orthogneisses of 1.05–1.0 Ga as the age of M2 metamorphism, i.e. consid-erably older than intrusion of the anorthosite–noritecomplex. Their interpretation was heavily influencedby a K–Ar age for osumilite of 970± 30 Ma (Maijeret al., 1981). More recent conventional U–Pb datingof zircon, monazite and titanite in orthogneisses byBingen & van Breemen (1998a,b) allows a clearerpicture to emerge. Of these orthogneisses, a series ofcalc-alkaline augen-gneisses known as the Feda Suitelies predominantly east of the orthopyroxene-in iso-grads and contains magmatic hornblende, titanite,allanite and zircon as well as metamorphic monaziteand thorite. From these rocks, Bingen & van Breemen(1998a,b) deduced an intrusion age of c. 1050 Ma fromzircon of magmatic origin (based on morphology).Several fractions of metamorphic monazite in theserocks and other charnockitic gneisses, plotted con-cordantly or near concordantly at 1024–997 Ma. Thismonazite growth is interpreted as resulting from hightemperature breakdown of allanite, titanite and biotiteduring the regional M1 phase (Bingen & van Breemen,

1998b) with subsequent reactivation and coolingrecorded by a spread of monazite ages continuingdown to 970 Ma. Wielens et al. (1980) also derived anage of 1020–1000 Ma for metamorphic zircon.The thermal effect of anorthosite–norite intrusion

(M2) and subsequent cooling (M3) is represented by:(1) growth of clinopyroxene at the expense of horn-blende in the Feda Suite orthogneisses (isograd liesto the east of the orthopyroxene-in isograd, Fig. 1),accompanied by a distinct monazite growth phase at930–925 Ma; (2) titanite ages throughout the areaclustered around 918± 2 Ma; (3) hornblende Ar–Arages in the range 930–904 Ma; and (4) a further,distinct, low-U monazite group yielding an age of912–904 Ma (Bingen & van Breemen, 1998b; Bingenet al., 1998). Subsequent hydrothermal alteration ofmica and hornblende is reflected in Rb–Sr, K–Ar andAr–Ar ages in the range 895–853 Ma (Verschure et al.,1980; Bingen et al., 1998).Conventional thermal ionisation mass spectrometry

(TIMS) U–Pb zircon geochronology studies on gran-ulites in the regional aureole around the anorthosite–norite intrusive complex (Wielens et al., 1980; Bingen& van Breemen, 1998a) found no evidence for the ageof contact metamorphism. However, these studiesdid not include samples from within the pigeonite-inisograd. The application of imaging-controlled ion-microprobe analysis to samples from the zone closestto the intrusive contact and comparison with lowergrade augen-gneisses provides an excellent opportu-nity to study the behaviour of zircon during UHTmetamorphism.

ANALYTICAL METHODS

Selected samples were crushed and zircon was concentrated usingheavy liquids and magnetic separation. The 50–400 lm zirconfractions were handpicked under a binocular microscope to selectunbroken grains of different morphologies and sizes without inclu-sions. About 200 grains per sample were mounted on 1 incl diameterepoxy disks. After polishing to expose the grain interiors, theinternal structures of zircon were documented with back-scatteredelectron (BSE) and cathodoluminescence (CL) images on a JEOL8900 RL electron microprobe at the Institut fur Geowissenschaften,Mainz.Uranium, thorium and lead isotopic measurements were made

during one session on the Perth Consortium SHRIMP II ion-microprobe, employing operating, data-processing and error calcu-lation procedures described by Nelson (1997). The 1-sigma error inthe 206Pb ⁄ 238U of the standard analyses (n ¼ 14) for this study was0.64%. Pb ⁄U ratios were determined relative to that of the standardSri Lankan zircon CZ3, which has been assigned a 206Pb ⁄ 238U valueof 0.0914, corresponding to an age of 564 Ma (Pidgeon et al., 1994).Data reduction involved checking individual scans for outliers andaveraging the results of the remaining scans. An important factor inassigning a specific age to a group of data points is often the commonlead correction. The presence of trace amounts of the nonradiogenicisotope 204Pb indicates that either some Pb was incorporated into thecrystal structure during mineral growth or that surface lead waspresent on the grain mount. Almost all of the measurements onunknowns contained very little 204Pb (see Table 1). The uncertaintyon the amount and composition of common Pb for correction wouldin this case be larger than possible errors introduced by no correc-tion, and the data were therefore not corrected for common Pb.

PO L Y PH A S E Z I R C O N I N U H T G R A N U L I T E S 72 9

Major elements and selected trace elements (Hf, Y, P in some casesalso U, Th, Pb, Yb) in zircon were measured with the JEOL 8900 RLelectron microprobe, using a 20 kV and 100 nA beam focused to5 lm. The Ka lines for Al, P, K, Ca and Fe, the La line for Yb, LaIIlines were used for Zr and Si, and Ma lines for Hf, Th, and U.Natural standards were used for Ca and P (apatite), K (orthoclase),interference-free silicate glasses doped with REE were used for Yb,silicate glass was used for Th, pure metals were used for Hf and U,Fe2O3 for Fe, synthetic Y3Al5O12 for Y and Al, and synthetic ZrSiO4for Zr and Si. Some of the elements served as monitors for possibleinclusions or secondary alteration (Fe, K, Al, Ca). Counting time of30 s was chosen for Si, 40 s for Zr and Hf, 50 s for P, 60 s for Yb, and100 s for Th, U and Y. Detection limits for trace elements wereestimated at about 80–100 ppm.

SAMPLE DESCRIPTIONS

Samples of the country rock to the anorthosite complex were col-lected along a traverse perpendicular to the margin of the intrusivecomplex and to the metamorphic isograds (Fig. 1), along state roadNo. R42 running ENE from Egersund to Tonstad. Here, only datafor orthogneiss samples 17 and 19, from within the �pigeonite in�zone, and for orthogneiss sample 2, from outside the �hypersthene-in�zone, are presented.Sample NR2B: (roadcut at Osen, UTM: 36205 ⁄ 650480). This

sample is a granodioritic augen-gneiss from outside the hypersth-ene-in isograd showing well developed, late cleavage. Large perthitic

K-feldspar augen (cm-size) and quartz and minor plagioclase(myrmekitised) constitute the bulk of the leucocratic layers. Grainrims are recrystallized and have serrated edges. Quartz shows adistinct chessboard subgrain pattern. Mafic layers contain abundantamphibole and biotite, often chloritised. Apatite, zircon and ilmeniteare often found intergrown or in clusters and are more abundantin the mafic layers. Accessory titanite is found intergrown withamphibole.Sample NR17C was collected along Gyavatnet lake

(34385 ⁄ 649708), within the pigeonite-in isograd, from the same out-crop as sample B649 of Bingen & van Breemen 1998a). The sampleconsists of layered leucocratic and melanocratic charnockite. Macro-scopically, pyroxene + magnetite layers are visible whereas layerswith garnet are rare. In thin section the leucocratic parts consist, inorder of abundance, of K-feldspar, quartz and plagioclase (myr-mekite). The feldspar has lobate grain boundaries, often recrystal-lized to small subgrains, and shows evidence of strong deformation.Quartz is deformed into parallel subgrains (gliding on the prismbase). Anhedral inverted pigeonite and clinopyroxene is found inbands where it is commonly intergrown with accessory apatite(abundant), ilmenite and zircon. Zircon is often found included in thepyroxene or grown on ilmenite.NR19A was collected at the western end of Gyavatnet lake dam

(34245 ⁄ 649605) at around 2.5 km from the contact of the intrusivecomplex. The sample is charnockitic, showing a strong late, more orless vertical, foliation. The sample is petrographically similar toNR17C except for the presence of minor biotite.

Table 1. Summary of ion-microprobe U-Th-Pb results for zircon from mineral separates.

Grain

spot

U

p.p.m.

Th

p.p.m. Th ⁄UPbrad

ppm †% f206

206Pb ⁄

238U ±

207Pb ⁄

235U ±

207Pb ⁄

206Pb ±

Ages (in Ma)*

206Pb ⁄ 238U ± 207Pb ⁄ 235U ± 207Pb ⁄ 206Pb ± % conc

Sample NR2B

4.61.1re 259 64 0.246 42 0.00 0.16411 0.0012 1.64218 0.0189 0.07257 0.0006 980 7 987 7 1002 16 98

4.67.2r 283 70 0.246 47 0.00 0.16822 0.0012 1.70358 0.0191 0.07345 0.0006 1002 7 1010 7 1026 15 98

4.64.2r 314 85 0.270 53 0.00 0.16985 0.0012 1.71535 0.0186 0.07325 0.0005 1011 7 1014 7 1021 15 99

4.69.1r 272 75 0.276 46 0.00 0.17144 0.0013 1.77658 0.0199 0.07516 0.0006 1020 7 1037 7 1073 15 95

4.75.2c 409 140 0.343 72 0.01 0.17255 0.0011 1.77653 0.0169 0.07467 0.0005 1026 6 1037 6 1060 12 97

4.75.3re 858 98 0.114 142 0.00 0.17355 0.0010 1.75866 0.0131 0.07350 0.0003 1032 5 1030 5 1028 9 100

4.67.1c 860 237 0.276 149 0.00 0.17417 0.0010 1.78245 0.0132 0.07422 0.0003 1035 5 1039 5 1048 8 99

4.64.1c 714 237 0.332 126 0.01 0.17439 0.0010 1.81021 0.0141 0.07529 0.0003 1036 6 1049 5 1076 9 96

4.63.1c 761 197 0.258 132 0.00 0.17472 0.0010 1.78863 0.0137 0.07425 0.0003 1038 5 1041 5 1048 9 99

4.63.3c 640 166 0.259 111 0.00 0.17477 0.0010 1.79373 0.0147 0.07444 0.0004 1038 6 1043 5 1053 10 99

4.63.2c 864 191 0.221 148 0.00 0.17513 0.0010 1.78219 0.0132 0.07380 0.0003 1040 5 1039 5 1036 8 100

4.75.1c 927 192 0.207 160 0.00 0.17721 0.0010 1.81156 0.0131 0.07414 0.0003 1052 5 1050 5 1045 8 101

Sample NR17C

2.42.1r 101 176 1.744 21 0.00 0.15236 0.0016 1.49016 0.0271 0.07094 0.0010 914 9 926 11 956 28 96

2.34.2r 127 261 2.052 28 0.01 0.15408 0.0015 1.48581 0.0248 0.06994 0.0009 924 8 925 10 927 25 100

2.34.1r 64 194 3.042 17 0.03 0.15657 0.0020 1.50638 0.0344 0.06978 0.0012 938 11 933 14 922 35 102

2.37.2r 240 210 0.876 43 0.00 0.15746 0.0012 1.52020 0.0190 0.07002 0.0006 943 7 939 8 929 18 101

2.42.2 198 194 0.980 39 0.01 0.17011 0.0014 1.69223 0.0220 0.07215 0.0007 1013 8 1006 8 990 18 102

2.36.1c 247 506 2.051 62 0.00 0.17324 0.0013 1.77522 0.0207 0.07432 0.0006 1030 7 1036 8 1050 16 98

2.36.2c 169 399 2.354 45 0.01 0.17372 0.0015 1.77869 0.0241 0.07426 0.0007 1033 8 1038 9 1049 19 98

2.37.3c 251 547 2.181 65 0.01 0.17453 0.0013 1.79742 0.0210 0.07469 0.0006 1037 7 1045 8 1060 16 98

2.37.1c 113 222 1.961 29 0.01 0.17794 0.0018 1.82179 0.0300 0.07425 0.0009 1056 10 1053 11 1048 24 101

2.44.1c 249 130 0.524 75 0.00 0.27939 0.0021 3.77279 0.0370 0.09794 0.0005 1588 10 1587 8 1585 10 100

Sample NR19A

1.67.2r 415 306 0.737 70 0.00 0.15171 0.0010 1.49414 0.0148 0.07143 0.0005 911 6 928 6 970 13 94

1.67.1r 494 246 0.497 79 0.00 0.15334 0.0010 1.48719 0.0139 0.07034 0.0004 920 5 925 6 938 12 98

1.38.3re 282 8 0.028 42 0.01 0.16132 0.0012 1.62993 0.0184 0.07328 0.0006 964 7 982 7 1022 16 94

1.111.3r 349 25 0.072 53 0.00 0.16245 0.0011 1.60021 0.0169 0.07144 0.0005 970 6 970 7 970 15 100

1.67.3re 567 47 0.084 88 0.00 0.16409 0.0010 1.64275 0.0141 0.07261 0.0004 979 6 987 5 1003 11 98

1.38.1r 237 40 0.171 38 0.01 0.16629 0.0013 1.68751 0.0203 0.07360 0.0006 992 7 1004 8 1031 17 96

1.70.1r 431 40 0.093 68 0.01 0.16763 0.0011 1.68633 0.0160 0.07296 0.0004 999 6 1003 6 1013 12 99

1.111.2 392 32 0.083 63 0.00 0.17080 0.0011 1.73738 0.0171 0.07377 0.0005 1017 6 1022 6 1035 13 98

1.38.2c 358 80 0.222 61 0.01 0.17409 0.0012 1.78277 0.0179 0.07427 0.0005 1035 6 1039 7 1049 13 99

1.93.1re 1095 42 0.038 277 0.00 0.26591 0.0014 3.44640 0.0217 0.09400 0.0003 1520 7 1515 5 1508 5 101

† f 206% ¼ 100 · (common 206Pb ⁄ total 206Pb).% conc ¼ concordance as 100 · (206Pb ⁄ 238U age) ⁄ (207Pb ⁄ 206Pb age).r: rim; c: core; re: recrystallized. *: with 1r error.Results of each sample in order of increasing 206Pb ⁄ 238U age.

7 30 A . M O L L E R E T A L .

RESULTS

Ion-microprobe data on zircon are listed in Table 1.Cathodoluminescence (CL) and backscatter-electronimages for studied zircon usually show anticorrelatedbrightness with CL-dark areas usually appearingBSE-bright. Details in zoning patterns are usually onlyseen in one or the other rather than both types ofimage. Figure 2 therefore shows both CL and BSEimages for representative grains from each sample.Those parts of grains which show discrete fine-scaleoscillatory zoning have been interpreted to reflectmagmatic growth. Metamorphic zircon is representedby recrystallized parts of grains as well as overgrowthson existing grains. We interpret those areas as over-growths that show little or no internal zoning andshapes unrelated to existing features such as magmaticcores. These areas may be the product of a dissolution-precipitation process. Areas which have lobateboundaries into zoned, magmatic parts of grains andshow some internal zoning or relicts of earlier featuresare interpreted to results from complete or partialsolid-state recrystallization; a process faster than solid-state diffusion. Note that early postmagmatic solid-state recrystallization can produce similar features torecrystallization during metamorphic events (probablybecause the mechanism is essentially the same).The results of microprobe analyses from spots

directly adjacent to the ion-microprobe analysis pitsare listed in Table 2. Totals are generally lower than100%, but stoichiometry is perfect, and duplicateanalyses made on different days usually gave very goodreproducibility of results. The Th content is often closeto or below the estimated detection limit, but Th ⁄Uratios estimated from the electron microprobe analysesare usually in very good agreement with the resultsobtained by SHRIMP (except for Th-poor areas).

U–Pb and geochemical data on zircon

The zircon population in augen-gneiss sample NR2Bcomprises mainly long-prismatic, pink, translucent tobrownish grains, and very few pinkish, anhedralgrains. The prismatic grains all preserve oscillatoryzoning, visible in CL and BSE images and typicalfor magmatic crystallization environments (Fig. 2a).Unzoned or weakly zoned rims of medium brightnessare present on most grains but vary in thickness from<5 lm (grain 4.75) to >50 lm (grain 4.67) and arehere interpreted as metamorphic overgrowths. It isclear from textural evidence that these rims areyounger than the oscillatory-zoned cores and alsoyounger than some weakly zoned CL-dark areas (seeovergrowth covering broad CL-dark area in grain 4.75,patches in grains 4.64 and 4.67). Oscillatory zoningis truncated by these CL-dark areas as well as byCL-dark patches with lobate boundaries, and theseare interpreted to represent recrystallized areas. Therecrystallization is possibly associated with mechanical

weaknesses in the grains (crack in grain 4.75, inclusionsin grains 4.64 and 4.67). Corrosion of oscillatory-zoned parts is visible on the upper left side of the corein grain 4.67, where a CL-dark area with little zoningtruncates a prismatic oscillatory zoned core. Notableis a �reaction front� of medium brightness, which islocated at the interface between a CL-dark zone andthe core. A similar feature can be seen in grain 4.61where the lower part preserves some remnants ofoscillatory zoning, whereas the upper part consistsentirely of a medium brightness, broadly zoned areawhich impinges on the older structured part, alsoforming a reaction front – in this case CL-darker. Suchdiscordant features have been described by Hoskin &Black (2000) and were also interpreted to be the resultof solid-state recrystallization.Magmatic crystallization in this sample is dated at

1034± 7 Ma by eight analyses, the majority of analysesfrom this sample (Fig. 3). Recrystallization resulting inCL-dark areas occurred shortly after crystallization(and did not affect the U–Pb system, i.e. the age) asshown by the 206Pb ⁄ 238U date of 1032± 5 Ma obtainedon grain 4.75, which is in the range obtained on theparts of the grain which preserve magmatic growthfeatures. Three analyses of unzoned, almost featurelessrims have 206Pb ⁄ 238U dates between 1020± 7 and1002± 7 Ma (Fig. 2, grain 4.64 & 4.67) and are inter-preted to result from zircon growth during M1 regionalmetamorphism. The youngest date obtained from thissample is near-concordant at 980± 7 Ma, analysed onthe unzoned area withmedium brightness interpreted asa solid-state recrystallization front (Fig. 2a, grain 4.61).All analyses from this sample yield moderate Th and Ucontents resulting in a narrow range of Th ⁄U between0.2 and 0.34, regardless of the age or zoning featureobserved (Fig 4a,b). One exception is the recrystallizedzone in grain 4.75 giving a lower Th ⁄U of 0.11. Theobservation that analyses from the sample form ahigher and a lower U group in a Th vs. U plot may be anartefact produced by the limited number of analyses(Fig. 4c). Electron microprobe data show that the traceelement content of metamorphic and most recrystal-lized zircon in this sample is not significantly differentfrom the original igneous zircon (Table 2). For example,all analyses from this sample define a cluster in the Y(as a proxy for heavy rare earth element behaviour)vs. Th ⁄U diagram (Fig. 4a).Zircon from pigeonite-bearing granulite NR17C

(T ¼ 900 �C) is also mainly prismatic, albeit usuallywith low aspect ratios. Some grains have roundedterminations or are present as fragments; anhedralgrains are rare. Many grains preserve magmatic oscil-latory and sector zoning (Fig. 2b, fir-tree type, grain2.36). Magmatic growth ages of 1033± 20 Ma aremost abundant (n ¼ 3), one grain centre also yielded a206Pb ⁄ 238U date of 1056± 10 Ma (Figs 2b & 3).Concordant magmatic ages have been obtaineddirectly at the rim of grains and agree within error withdates obtained on the more internal parts of the same

PO L Y PH A S E Z I R C O N I N U H T G R A N U L I T E S 73 1

grain (Fig. 2b, grain 2.36). Another core preserves aconcordant 206Pb ⁄ 238U date of 1588± 10 Ma (Fig. 2b,grain 2.44) interpreted to represent magmatic crystal-lization of this inherited grain. Recrystallization

resulting in BSE-bright reaction fronts truncating theearlier zoning pattern can be observed in this grain.Such recrystallized areas are not as common as in theaugen-gneiss sample (NR2B).

Fig. 2. Cathodoluminescence (CL, upper rows) and backscatter-electron (BSE, lower rows) images taken after SHRIMP analysis (noteion-probe pits, best visible on BSE images). Ion-microprobe U–Pb results are indicated as 206Pb ⁄ 238U dates with 1r errors. (a) zirconfrom amphibolite facies sample NR2 with oscillatory zoned magmatic cores and variably wide featureless rims, yielding fromM1 dates.Grain 4.61 shows a recrystallization front truncating earlier fine-scale zoning, an analysis in the recrystallized area yields an ageintermediate between M1 and M2, indicative of incomplete recrystallization. (b) zircon from UHT (pigeonite-bearing) granulite faciessample NR17, with preservation of magmatic oscillatory and sector zoning. Featureless rim gives M2 UHT metamorphic age.Note concordant result for inherited Mid-Proterozoic core. (c) zircon from UHT (pigeonite-bearing) granulite-facies sample NR19showing oscillatory zoned magmatic, inherited cores, truncated by wide featureless M1 as well as M2 metamorphic rims.

7 32 A . M O L L E R E T A L .

Two different ages were obtained for featurelessareas, a CL-bright grain interior with a 206Pb ⁄ 238U dateof 1013± 8 Ma is overgrown by a c. 50 lmwide lower-brightness rim of 914± 11 Ma (grain 2.42, not shownin Fig. 2). Distinct, featureless rims dating theM2UHT

event are also found on other grains that preserveoscillatory zoned cores and comprise overgrowths(grain 2.34, not shown in Fig. 2) as well as recrystallizedareas (e.g. Fig. 2b, grain 2.37). These four analyses yielda combined 206Pb ⁄ 238U date of 931± 22 Ma (2r).

Table 2. Representative electron-microprobe analyses of zircon.

Sample

Analysis No.

SHRIMP No.

Description

NR2B

56

4.75.1c

p. recr.

59

4.75.2c

osc. z.

62

4.69.1r

no z.

61

4.67.1c

psc. z., CL-br.

64

4.64.1c

osc. z.

66

4.64.2r

no z.

68

4.63.2c ⁄ 3cosc. z.

78

4.67.2r

no z.

79

4.75.3re

CL-br.

83

4.63.1c

osc. z.

85

4.61.1.re

p. no z.

SiO2 32.18 32.17 32.63 32.35 32.36 32.42 32.76 32.23 32.25 32.51 32.47

P2O5 0.05 0.09 0.06 0.09 0.07 0.06 0.07 0.06 0.03 0.07 0.05

ZrO2 65.44 65.70 65.54 65.40 65.28 65.40 65.80 65.43 65.58 65.65 64.93

HfO2 1.36 1.02 1.16 1.18 1.15 1.16 1.27 0.99 0.98 1.08 1.03

Y2O3 0.03 0.09 0.05 0.10 0.09 0.04 0.07 0.03 0.01 0.10 0.03

UO2 0.10 0.03 0.03 0.12 0.03 0.04 0.08 0.03 0.06 0.09 0.04

ThO2 0.01 0.02 0.01 0.04 0.00 0.01 0.03 0.00 0.02 0.02 0.00

Yb2O3 0.02 0.03 0.05 0.04 0.03 0.04 0.04 0.02 0.02 0.05 0.04

Total 99.19 99.14 99.54 99.32 99.00 99.16 100.11 98.80 98.95 99.57 98.59

Formula

Si 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.01

P 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Zr 0.99 0.99 0.98 0.98 0.98 0.98 0.98 0.99 0.99 0.98 0.98

Hf 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01

Total 2.00 2.00 2.00 2.00 2.00 2.00 2.00 2.00 2.00 2.00 2.00

Trace elements in p.p.m.

Hf 11533 8615 9803 9964 9718 9837 10769 8387 8319 9150 8692

Y 130 335 213 402 339 138 264 134 47 374 134

P 98 354 248 354 283 236 283 131 59 157 118

Yb 70 123 228 180 119 162 167 92 83 224 162

Th 79 132 70 378 0 79 228 0 132 193 0

U 908 300 300 1031 256 317 696 300 538 829 326

Hf ⁄Y 89 26 46 25 29 71 41 63 176 24 65

Th ⁄U EPMA 0.09 0.44 0.23 0.37 0.00 0.25 0.33 0.00 0.25 0.23 0

no K, Ca, Al, Fe above 0.015 wt%.

Abbreviations: br. bright; fir-tr. fir-tree; osc. oscillatory; ov. overlapping; p. partly; recr. recrystallized; v. very; z. zoning.

Table 2. (Cont’d).

Sample

Analysis No.

SHRIMP No.

description

NR17C

87

2.44.1c

osc. z.

89

2.42.1r

no z.

91

2.42.2c

CL-br.

92

2.36.2c

sector z.

93

2.36.1c

fir-tr. z.

95

2.37.1c

z.

96

2.37.2r

weak z.

97

2.37.3c

z.

98

2.34.1r

no z.

99

2.34.2r

no z.

SiO2 31.92 32.16 32.40 32.30 32.65 32.28 32.38 32.08 32.35 32.34

P2O5 0.21 0.13 0.17 0.17 0.18 0.20 0.12 0.18 0.15 0.13

ZrO2 64.79 64.94 65.45 64.65 65.28 64.79 65.17 64.77 65.31 65.58

HfO2 0.78 1.04 1.09 1.14 1.17 1.16 1.09 1.15 1.07 1.03

Y2O3 0.69 0.10 0.14 0.21 0.22 0.21 0.11 0.20 0.12 0.14

UO2 0.06 0.01 0.02 0.03 0.03 0.03 0.02 0.03 0.01 0.01

ThO2 0.04 0.01 0.03 0.05 0.06 0.04 0.02 0.06 0.03 0.02

Yb2O3 0.18 0.05 0.11 0.08 0.08 0.09 0.06 0.07 0.04 0.05

Total 98.68 98.45 99.41 98.64 99.67 98.80 98.97 98.53 99.07 99.30

Formula

Si 0.99 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00

P 0.01 0.00 0.00 0.00 0.00 0.01 0.00 0.00 0.00 0.00

Zr 0.98 0.98 0.98 0.98 0.98 0.98 0.98 0.98 0.98 0.99

Hf 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01

Total 1.99 2.00 2.00 1.99 1.99 1.99 2.00 1.99 2.00 2.00

Trace elements in p.p.m.

Hf 6623 8853 9234 9692 9896 9837 9243 9769 9056 8751

Y 2725 406 535 831 874 843 441 768 465 532

P 463 284 364 380 393 426 262 402 321 292

Yb 795 224 479 369 356 408 250 294 189 228

Th 387 114 264 439 501 378 167 492 237 141

U 494 106 212 229 273 220 212 247 79 79

Hf ⁄Y 2.4 22 17 12 11 12 21 13 19 16

Th ⁄U EPMA 0.78 1.08 1.25 1.92 1.83 1.71 0.79 1.99 2.99 1.77

no K, Ca, Al, Fe above 0.015 wt%.

Abbreviations: br. bright; fir-tr. fir-tree; osc. oscillatory; ov. overlapping; p. partly; recr. recrystallized; v. very; z. zoning.

PO L Y PH A S E Z I R C O N I N U H T G R A N U L I T E S 73 3

Zircon in this sample is characterised by high, butvery variable Th ⁄U between c. 0.8 and 3.0 both inmagmatic and metamorphic Sveconorwegian grains(Fig. 4a,b). The inherited Mesoproterozoic magmaticcore is distinct with its somewhat lower Th ⁄U, butmost distinctly by its Y content, about five times higherthan in the younger grains. Note that it is not clearlydistinguishable in the Th vs. U plot (Fig. 4c). The threehigh Th analyses which set themselves apart in this plotare those of the magmatic 1033± 20 Ma zircon,

whereas older magmatic and younger metamorphicanalyses fall into the same range and cannot be dis-tinguished chemically.The granitic pigeonite-bearing granulite NR19A, is

the closest to the contact with the intrusive anortho-site–norite complex. Zircon is less euhedral than thatfound in the other samples, but still preserves mag-matic oscillatory zoning visible in CL and BSE images(e.g. Fig. 2c, grain 1.38). A zoned grain interior datedat 1035± 6 Ma is consistent with the results obtained

Table 2. (Cont’d).

Sample

Analysis No.

SHRIMP No.

description

NR19A

103

1.111.3r

ov. weak z.

104

1.111.2

v. weak z.

107

1.67.1r

no z.

108

1.67.2r

no z. CL-br.

109

1.67.3re

no z.,

112

1.70.1r

ov. weak z.

115

1.38.2c

weak osc. z.

117

1.38.3re

weak z.

118

1.38.1r

weak z. CL-br.

120

1.93.1re

weak z.,

SiO2 32.10 31.86 31.98 31.92 31.93 32.32 31.96 32.21 32.17 32.36

P2O5 0.35 0.37 0.41 0.41 0.41 0.41 0.39 0.32 0.35 0.06

ZrO2 64.60 64.94 64.91 65.11 64.66 65.41 64.22 64.59 64.66 65.17

HfO2 1.06 1.09 1.18 1.12 1.09 1.09 1.05 1.09 1.05 1.073

Y2O3 0.31 0.34 0.35 0.36 0.37 0.37 0.36 0.27 0.32 0.045

UO2 0.05 0.03 0.06 0.05 0.06 0.05 0.04 0.02 0.03 0.117

ThO2 0.01 0.00 0.03 0.03 0.01 0.00 0.00 0.00 0.01 0.006

Yb2O3 0.16 0.14 0.12 0.11 0.17 0.18 0.16 0.15 0.15 0.055

Total 98.64 98.78 99.02 99.10 98.70 99.82 98.19 98.65 98.74 98.88

Formula

Si 1.00 0.99 0.99 0.99 0.99 0.99 1.00 1.00 1.00 1.00

P 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.00

Zr 0.98 0.98 0.98 0.98 0.98 0.98 0.98 0.98 0.98 0.98

Hf 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01

Total 1.99 1.99 1.99 1.99 1.99 1.99 1.99 1.99 1.99 2.00

Trace elements in p.p.m.

Hf 9022 9277 9972 9455 9226 9260 8921 9234 8938 9099

Y 1213 1354 1362 1421 1461 1441 1421 1079 1256 177

P 768 805 899 888 903 888 855 687 768 120

Yb 681 632 518 483 725 773 703 650 650 242

Th 44 0 220 246 70 0 0 0 105 53

U 458 282 485 423 555 432 353 176 229 1031

Hf ⁄Y 7.4 6.8 7.3 6.7 6.3 6.4 6.3 8.6 7.1 51

Th ⁄U EPMA 0.10 0 0.45 0.58 0.13 0.00 0.00 0.00 0.46 0.05

no K, Ca, Al, Fe above 0.015 wt%.

Abbreviations: br. bright; fir-tr. fir-tree; osc. oscillatory; ov. overlapping; p. partly; recr. recrystallized; v. very; z. zoning.

Fig. 3. Concordia diagram for SHRIMPU–Pb analyses of zircon from the Rogalandmetamorphic aureole. Pooled results are206Pb ⁄ 238U ages with 2r errors. Mesoprote-rozoic inherited cores in granulite samplesNR17C and NR19A are not shown.

7 34 A . M O L L E R E T A L .

in the other two samples farther away from the hightemperature contact and is interpreted to date mag-matic crystallization. Some grains are strongly alteredby recrystallization, which produced CL-dark zoneswith some internal detail visible in BSE imaging.However, this recrystallization appears to haveoccurred shortly after magmatic crystallization. Ananalysis on this recrystallized zone very close to theedge of the grain yielded a concordant 206Pb ⁄ 238Uage of 1520± 7 Ma (Fig. 2c, grain 1.93). UHT

metamorphism did not affect the U–Pb system, andthere is no evidence for Pb-loss by diffusion.Other grains in this sample are anhedral and tex-

turally show at least two phases of overgrowth and ⁄ orrecrystallization with weak internal zoning (Fig. 2c,grains 1.67, 1.38 & 1.70). The earlier phase commonlytruncates older strongly zoned grain interiors (e.g.grain 1.67) or forms thick rims on pre-existing grains(Fig. 2c, grains 1.38 & 1.70). It yielded two age groups,the older one corresponding to M1 metamorphismwith one concordant analysis at 999± 6 Ma and twonear-concordant dates of 1017± 6 Ma and 992±7 Ma (all 1r errors) up to the very edge of the grains.Three analyses on tips of grains showing textural evi-dence of further growth (e.g. Fig. 2c, grain 1.38) orrecrystallization (e.g. Fig. 2c, grain 1.67; incompleterecrystallization is consistent with the truncatingnature of the boundary with the texturally older part ofthe grain) returned a combined date of 972± 20 Ma(2r error). A late overgrowth is best discernible ingrain 1.67, where two weakly zoned overgrowths areseparated by a thin highly luminescent zone. The areaoutwards yielded two 206Pb ⁄ 238U dates of 920±5 and911± 6 Ma, respectively, identical within error andcorresponding to the UHT M2 phase.Most analysed grains in this sample have uniformly

low Th ⁄U in the narrow range 0.03–0.2, with someclearly in the range below 0.1 that was used byWilliams et al. (1996) as an indicator of metamorphiczircon growth. However, these analyses cover theentire magmatic and M1 age range in this sample,whereas the two analyses with distinctly higher Thcontents grew during M2 metamorphism (Fig. 4b,c).In the Th ⁄U vs. Y plot there is some distinction pos-sible between the M2 medium Th ⁄U zones andmost other analyses but, more distinctly, the stronglyrecrystallized Mesoproterozoic grain is discernible byits very low Y content.Microprobe trace element analyses of zircon from

all three samples dated by SHRIMP show thesame correlation trends commonly observed in zircon(e.g. Benisek & Finger, 1993). There are positivecorrelations between Y and P as expected for theincorporation of a xenotime solid solution in zircon,between U and Hf (albeit less well constrained), andbetween Hf ⁄Y and 1 ⁄Y suggesting a coupled exchangerelationship. An exception are some recrystallizedareas in grains, uniformly characterised by lower Ycontents close to the detection limit, whereas tetrava-lent ion contents appear unchanged, and Y and P lieon the same positive correlation trend observed inareas of preserved magmatic zoning.

DISCUSSION

Diffusional Pb-loss in zircon?

Comparisons were made across the metamorphicisograds in the study area between rocks which expe-

Fig. 4. (a) Y vs. Th ⁄U diagram of zircon from the studiedsamples. Analyses from each sample cover a distinct range ofvalues. Clearly distinguishable exceptions are recrystallized areasand inherited cores. (b) Th ⁄U vs. age diagram for the analysedzircon. (c) Th vs. U diagram for zircon from Rogaland meta-morphic rocks. Magmatic and metamorphic zircon do notdisplay consistent differences in Th- and U-contents andTh ⁄U between magmatic and metamorphic grains.

PO L Y PH A S E Z I R C O N I N U H T G R A N U L I T E S 73 5

rienced a thermal overprint of about 700 �C, and thosewhich were subjected to temperatures above 900 �C(pigeonite-in). Zircon predating the UHT overprint(igneous and metamorphic M1) was found at eachlocation. Comparison of samples from outside theorthopyroxene-in isograd with samples in the UHTdomain shows that there is no difference in the agesdetermined for the igneous and M1 events. In otherwords, each age population that can be foundthroughout the sampled temperature profile has,within analytical uncertainty, constant U-Th-Pb iso-topic systematics. There is no correlation betweendiscordance of results and distance from grainboundaries or distance from the boundaries of dif-ferent growth zones within a zircon. Previous, con-ventional U–Pb zircon studies (Wielens et al., 1981;Bingen & van Breemen, 1998b) did not find M2 zir-con in the high-grade metamorphic rocks from Ro-galand, possibly because no samples were taken frominside the pigeonite-in isograd. In our study, M2metamorphic zircon growth was found only in thehigher grade samples NR17 and NR19, and we sug-gest that the influence of M2 metamorphism on theaugen-gneisses, at least outside the orthopyroxene-inisograd, may not have been sufficient to trigger newzircon growth (i.e. metamorphic reactions releasingZr, formation of partial melts). Mesoproterozoicmagmatic cores were found to be inherited only in thegranitic, higher grade samples interpreted to haveintruded together with the protoliths of the augen-gneisses during the Sveconorwegian orogeny at1034± 5 Ma (see also Bingen & van Breemen,1998b). The only analyses which cannot be attributedto a magmatic or metamorphic event in the area arethose which yielded ages of 974± 13 Ma (n ¼ 4;Fig. 3). These either straddle two different growthzones (analysis 1.111.3, not in Fig. 2) and can there-fore be explained as mixed ages, or occur in zoneswhich have the characteristics of solid state recrys-tallization (e.g. Fig. 2a, grain 4.61; Fig. 2c, grain1.67). Such solid state recrystallization may be in-complete and probably resulted in Pb and other traceelements being inherited from the pre-existing grain(Hoskin & Black, 2000; Pidgeon, 1992; Pidgeon et al.,1998). Alternatively, in the case of complete recrys-tallization, the obtained U–Pb results may date aphase of post-M1 recrystallization, possibly associat-ed with fluids infiltrating during cooling following M1peak metamorphic conditions. These possibilitiescannot be resolved with data from this study.

Comparison with experimental data

Laboratory experiments on solid state diffusion inminerals are limited by the short duration for whichthe charge can be subjected to high temperatures. Theexperiments are therefore commonly conducted attemperatures far in excess of natural geological con-ditions and results are extrapolated into the relevant

temperature range. Natural zircon from UHT granu-lites is material that has been subjected to high tem-peratures for several million years and may thereforebe able to provide a test for the validity of the exper-imental results.Experimental studies which evaluated Pb diffusion

parameters in zircon have yielded different results, butleft open the possibility for Pb-loss by diffusion undercrustal conditions. Lee et al. (1997) for examplederived a closure temperature of c. 940 �C for 200 lmzircon cooled at 10 �C Myr)1. A summary of theavailable recent experimental estimates on trace ele-ment diffusion distances in zircon is listed in Table 3.For 1000 �C, Lee et al. (1997) deduced 6.52 ·10)19 m2 s)1 for the diffusion coefficient (D), whereasD ¼ 3 · 10)24 m2 s)1 can be calculated for Pb2+ fromthe data of Cherniak & Watson (2000) for the sametemperature.Using the data of Lee et al. (1997), zircon from

rocks within the pigeonite-in isograd of the Rogalandmetamorphic aureole should exhibit evidence for dif-fusion distances (x ¼ �(D · t)) for Pb of about 3 lm,c. 12 lm and 45 lm calculated for zircon exposed totemperatures of 900, 950 and 1000 �C, respectively, fora period of 1 Myr. The diffusion parameters for Pb inzircon deduced by Cherniak & Watson, 2000) predict amuch lower migration distance of c. 10 lm for 1 Myrat 1000 �C. Note that these estimates are much moresensitive to temperature than to time because of thenature of the diffusion equation. Reducing the dura-tion of the high temperature event by a factor of 10, forexample, reduces migration distance of Pb by a factorof c. 3.2, the square root of 10; whereas reducingtemperature by only 50 �C (from 1000 to 950 �C) hasalmost the same effect. Lowering the temperature to900 �C reduces migration length by a factor of almost10. Uncertainty in the duration of the metamorphicevent therefore has far less importance for migrationdistances than the uncertainty on peak temperatureconditions. These temperatures are comparatively wellconstrained from petrological arguments, thermo-barometry and thermal modelling (Westphal et al. inpress). A simple model for diffusion involving a single

Table 3. Diffusion distances for trace elements in zircon(�(D · t)). Calculated from experimental diffusion data for950 �C maximum T and a temperature history modelled for2.5 km distance from intrusion (Westphal et al. in press; Fig. 8).

Element

Migration

distance

log D0

(m2 s)1)

Ea

(kJ ⁄mol) Reference

Pb 12–13 lm 5.59 )677 Lee et al. (1997)

Pb c. 4 lm )0.96 )550 Cherniak & Watson (2000)

Sm (MREE) c. 0.1 lm 8.46 )841 Cherniak et al. (1997a)

Dy (MREE) c. 0.5 lm 5.36 )734 Cherniak et al. (1997a)

Yb (HREE) c. 1 lm 7.40 )769 Cherniak et al. (1997a)

U c. 2 nm 0.21 )726 Cherniak et al. (1997b)

Hf c. 1 nm 3.21 )812 Cherniak et al. (1997b)

Th c. 0.5 nm 1.94 )792 Cherniak et al. (1997b)

7 36 A . M O L L E R E T A L .

stage at 950 �C over a period of 1 Myr yields nearlyidentical diffusion distances compared to a complexmodel obtained from a calculated T-t path for rocks at2.5 km distance from the intrusive complex (Westphalet al. in press; Fig. 8, with a maximum temperatureof 940 �C and cooling from 900 to 800 �C atc. 24 �C Myr)1).Using the experimental parameters on diffusion of

Lee et al. (1997), Pb-diffusion within grains or Pb-lossfrom the grains in those rocks which experienced950 �C and more should be measurable by ion-microprobe, producing diffuse age zoning in zirconand discordance of U–Pb results in grains where sub-micron-scale zoning in BSE images is perfectly pre-served (see Discussion below). This is not consistentwith the results obtained in this study which clearlyshows that grains which preserve igneous oscillatoryzoning in the images always preserve igneous ageinformation. Our preferred explanation for thisobservation is that the diffusion coefficients deter-mined by Lee et al. (1997) are overestimated. Evalua-tion of the Pb diffusion data by Cherniak & Watson(2000) is beyond the scope of analyses presented in thisstudy and cannot be realised with these samples. Onlyrocks which have been subjected to temperatureshigher than 1000 �C or 1050 �C for about 1 Myrwould exhibit diffusion profiles in REE and Pbmeasurable by ion- or electron- microprobe.BSE brightness is positively correlated with the

mean atomic number of a mineral (e.g. Benisek &Finger, 1993; Hanchar & Miller, 1993). In zircon itincreases with the presence of the heavy elements Hfand U, but probably also Y, heavy rare earth elements(HREE), Th and Pb (in approximate order ofdecreasing abundance). Migration distances of theheavy tetravalent ions (Hf, Th, U) are several orders ofmagnitude lower than those of Pb; all yield migrationdistances below 0.01 lm for a duration of 1 Myr at1000 �C (Cherniak et al., 1997b). The exact cause forcathodoluminescence in zircon is still debated, but it iswell established that part of it is a property of ZrSiO4itself (Hanchar & Miller, 1993), that there is an acti-vating influence of REE (in particular by Dy3+ andTb3+, e.g. Marfunin, 1979; Remond et al., 1992), andprobably a suppression of luminescence by metam-ictisation (Nasdala et al., In Press). It can be expectedthat the high-grade M1 event completely annealed allmetamictisation accumulated prior to M1 in theMesoproterozoic zircon and that little structuraldamage could accumulate between the Sveconorwe-gian igneous crystallization and M1, and between M1and M2, because the metamorphosed rocks remainedat depth and above the annealing temperature.Although the resolution of CL and BSE images are

not strictly equivalent, a qualitative comparison for theeffects of diffusion can be attempted. With the exper-imental diffusion data of Cherniak et al. (1997b) forDy3+, a diffusion distance of 2.4 lm can be calculatedfor a timescale of 1 Myr at 1000 �C (Table 3). REE

diffusion is 4–5 times faster than Hf, U and Th(Cherniak et al., 1997b), consequently a diffusiveeffect, i.e. blurring of micron-scale zoning patterns,should be seen in CL images of zircon heated to1000 �C for about 1 Myr, whereas BSE zoning shouldretain sharp contrasts. At 950 �C maximum tempera-ture, maintained for 1 Myr, REE as well as tetravalentelement diffusion are negligible at 1 lm or below. Thecomparison of BSE and CL images shows that wheredetailed zoning is present in the samples studied here, itis visible in both types of images. This implies thatthe REE, which are largely responsible for the CLresponse, have remained immobile as have the tetra-valent elements (U, Th, Hf).

CONCLUSIONS

Ion-microprobe U–Pb data were collected on zirconfrom three samples from a traverse perpendicular toisograds in the metamorphic aureole surrounding theRogaland anorthosite–norite intrusive complex. Thedata define five groups of 206Pb ⁄ 238U ages (Fig. 2):(i) 922± 14 Ma (n ¼ 6) for the M2 UHT metamorphicevent (ii) 974± 13 Ma (n ¼ 4) is interpreted as partialrecrystallization (iii) 1006± 11 Ma (n ¼ 6) definesthe timing of the M1 metamorphic event, (iv) 1034±5 Ma (n ¼ 12) and 1053± 5 Ma (n ¼ 2) are theSveconorwegian magmatic intrusion events (v) twomore 207Pb ⁄ 206Pb ages from the granulites are older at1585 and 1507 Ma. One date is from an oscillatoryzoned, inherited core (NR17C, grain 2.42), the other arecrystallized zone with weak internal features cross-cutting an oscillatory zoned core (NR19A, grain 1.93).These Mesoproterozoic grains define the pre-Sveco-norwegian source of the Rogaland metamorphosedgranitoids.The youngest group of results from this study with a

mean age of 922± 14 Ma is the first direct evidencefor M2 zircon growth during regional contact meta-morphism related to intrusion of the Rogaland anor-thosite–norite complex (M2). The analysed areasconsist of near-featureless rims that have been foundonly in the higher grade samples.We conclude that (possibly fluid-induced) recrystal-

lization of zircon, which is visible in CL and BSEimages (Fig. 2a, grain 4.75; Fig. 2c, grain 1.93), did notreset the U–Pb system and therefore preserved thesame age information as nonrecrystallized zircon fromthe same grains. Alternatively, this type of recrystalli-zation occurred during or shortly after magmaticcrystallization (within analytical precision) and aninfluence on the U–Pb systematics is therefore notdiscernible.

Trace element chemistry of zircon as a guideto geochronological interpretation

This study accentuates the limitations of the sole use ofTh and U as an aid in the interpretation of zircon ages

PO L Y PH A S E Z I R C O N I N U H T G R A N U L I T E S 73 7

and specifically the distinction of magmatic vs. meta-morphosed or metamorphically grown zircon. Meta-morphic zircon is not always recognisable by lowTh ⁄U as the study of Williams et al. (1996) may haveimplied. It may, in fact, have relatively high Th ⁄U orreveal no change in comparison to magmatic zircon(Fig. 4). We suggest that geochemical characterisationof zircon by trace elements such as Hf and Y and, tosome extent, by REE has the potential to become auseful interpretative tool and may be able to providemore powerful arguments for the interpretation of thegrowth environment of zircon. However, the mech-anisms controlling trace element behaviour have to bebetter understood for a more than empirical interpre-tation of zircon geochemistry.

Diffusion in zircon and implications for U–Pbgeochronology of high grade metamorphic rocks

Our ion-microprobe investigations on zircon fromUHT granulites from Rogaland do not show a mea-surable effect of diffusional Pb-loss or Pb mobilityaffecting geochronological results. Pre-UHT growthzones preserve their age information and remainconcordant despite the high temperatures reachedduring the UHT (T ¼ 950 �C) metamorphic overprint.Magmatic zircon preserves micrometre-scale traceelement zoning. We conclude therefore, that thermallydriven diffusion does not appear to be a viable andimportant mechanism for the mobility of Pb in natural,nonmetamict zircon under even the most extremecrustal conditions.This conclusion is based on the following obser-

vations. Imaging of zircon from rocks containinginverted pigeonite, sampled near the intrusive contactreveals sharp contacts between pre-UHT growthphases (Fig. 2). Each pre-UHT age population can befound throughout the sampled temperature profile,and there is no correlation between discordance ofresults and distance from grain rims or distance fromthe boundaries of growth phases. Therefore, our datado not support the experimental evidence of Lee et al.(1997) which suggested that diffusion in zircon atT ¼ 950 �C for t ¼ 1 Myr will produce visible ⁄mea-surable effects and lead to Pb-loss at zircon margins.We suggest instead that zircon discordance, or the

spread of apparently concordant results over a timelonger than geologically viable for continued growth(e.g. Ashwal et al., 1999), is the result of other pro-cesses such as metamictisation, leaching, and incom-plete solid-state recrystallization (e.g. Vavra et al.,1996; Mezger & Krogstad, 1997; Hoskin & Black,2000).

ACKNOWLEDGEMENTS

We thank M. Westphal, J.C. Schumacher, H. Degelingand especially B. Mocek for useful discussions, andN. Groschopf, M. Muller and B. Schulz-Dobrick for

assistance with the electron microprobe analyses. Ourthanks also to B. Bingen and R. Pidgeon for reviewsand to D. Robinson for editorial comments whichhelped to improve the manuscript. Zircon analyseswere carried out on the Sensitive High Resolution IonMicroprobe mass spectrometer (SHRIMP II) operatedby a consortium consisting of Curtin University ofTechnology, the Geological Survey of WesternAustralia and the University of Western Australia withthe support of the Australian Research Council. Thisresearch was financially supported by the DeutscheForschungsgemeinschaft (DFG) through grantsKr 590 ⁄ 62–1 and Kr 590 ⁄ 62–2 to A.K.

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Received 8 October 2001; revision accepted 15 May 2002.

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