u–pb geochronology from tonagh island, east antarctica ... · island, enderby land, east...

Precambrian Research 116 (2002) 237–263

U–Pb geochronology from Tonagh Island, East Antarctica:implications for the timing of ultra-high temperature

metamorphism of the Napier Complex

C.J. Carson a,*, J.J. Ague a, C.D. Coath b

a Department of Geology and Geophysics, Yale Uni�ersity, P.O. Box 208109, New Ha�en, CT 06520-8109, USAb Department of Earth and Space Science, Uni�ersity of California, Los Angeles, CA 90095-1567, USA

Accepted 13 February 2002

Abstract

Ion microprobe U–Pb zircon geochronology of an orthopyroxene-bearing felsic orthogneiss from central TonaghIsland, Enderby Land, East Antarctica provides insight into the chronological-metamorphic evolution of theArchaean Napier Complex, the details of which have been the source of debate for over two decades. The orthogneisscrystallised at 2626�28 Ma, predating peak, ultra-high temperature (UHT) metamorphism and development of anintense regional S1 gneissosity. Two subsequent episodes of zircon growth/resetting can be identified. A minor periodof zircon growth occurred at 2546�13 Ma, the regional significance and geological nature of which is unclear. Thiswas followed by an episode of abundant zircon growth, as mantles on �2626 Ma cores and as anhedral grains,partly characterised by high Th/U (�1.2), at �2450–2480 Ma. This age coincides with both lower and upperconcordia intercept ages from other U–Pb zircon studies, and several Rb–Sr and Sm–Nd whole-rock isochron agesfrom the Napier Complex. We conclude that UHT metamorphism occurred at �2450–2480 Ma, and find nocompelling evidence that UHT occurred much earlier as has been postulated. The zircon U–Pb data from this studyalso indicates a lower intercept age of �500 Ma, which coincides with the emplacement of Early Palaeozoicpegmatite swarms and synchronous infiltration of aqueous fluids into the southwestern regions of the NapierComplex. © 2002 Elsevier Science B.V. All rights reserved.

Keywords: Napier Complex; East Antarctica; U–Pb geochronology; Zircon; Ion probe

www.elsevier.com/locate/precamres

1. Introduction

The granulite-facies Archaean Napier Complexof Enderby Land, East Antarctica has received

considerable attention over the past four decades(e.g. Sheraton et al., 1987; Harley and Black,1987, 1997), initially by Soviet and Australiangeological expeditions, and, in recent years, byJapanese expeditions. The identification of un-usual metamorphic mineral assemblages on a re-gional scale, such as sapphirine+quartz(Dallwitz, 1968), osumilite�garnet, and orthopy-

* Corresponding author. Present address: Geological Surveyof Canada, 601 Booth Street, Ottawa, Ont., Canada K1A 0E8.

E-mail address: [email protected] (C.J. Carson).

0301-9268/02/$ - see front matter © 2002 Elsevier Science B.V. All rights reserved.

PII: S0301 -9268 (02 )00023 -2

C.J. Carson et al. / Precambrian Research 116 (2002) 237–263238

roxene+sillimanite+quartz, indicative of ultra-high temperature (UHT) metamorphism, gener-ated great interest in the geological community.Subsequent detailed petrological studies confi-rmed that the terrane experienced extreme tem-peratures approaching, even surpassing, 1000 °Cat pressures of 8–11 kbars, temperatures quiteunusual for mid-crustal metamorphism (Ellis,1980; Ellis et al., 1980; Grew, 1980, 1982; Harley,1985a, 1987, 1998; Sandiford and Powell, 1986a;Harley and Hensen, 1990; Harley and Motoyoshi,2000). These stimulating discoveries in the petro-logical arena were matched by chronological re-ports that suggested southwestern regions of theterrane were composed of extremely ancientcrustal material (Pb–Pb ages �4000 Ma;Sobotovich et al., 1976; Lovering, 1979). Al-though, subsequent isotopic work questioned theveracity of these very old protolith ages, an earlyto mid Archaean origin for many of the or-thogneiss units and paragneiss sequences acrossthe Napier Complex is now nevertheless well es-tablished (Black and James, 1983; Black et al.,1983b, 1986a,b; DePaolo et al., 1982; Owada etal., 1994; Harley and Black, 1997). However, thetemporal framework of the structural-metamor-phic evolution of the Napier Complex, in particu-lar the timing of UHT metamorphism, remains asource of some debate amongst researchers (De-Paolo et al., 1982; Harley and Black, 1997; Grew,1998). In this paper, we present new zircon U–Pbion probe data from Tonagh Island, AmundsenBay, in which we offer additional evidence andconstraints concerning the timing of UHT tec-tonothermal events within the Napier Complex.

2. Geological setting

The Archaean Napier Complex is locatedwithin the East Antarctic Shield between latitudes66°–68°S and longitudes 48°–57°E and occupiesmuch of Enderby Land (Fig. 1). The south-west-ern regions of the Complex, well exposed as near-shore islands and coastal mountains in the vicinityof Amundsen and Casey Bays, also include sev-eral broadly east–west trending ranges of nuna-taks and massifs of the Tula, Scott and Raggatt

Mountains. Inland, to the north and east ofAmundsen Bay, lie the scattered and sparselyexposed nunataks of the Napier Mountains. Al-though the precise location of the boundary isuncertain (Fig. 1), to the south and southeast theNapier Complex is juxtaposed against the gran-ulite-facies early Neoproterozoic Rayner Complex(�1000 Ma; e.g. Black et al., 1987; Harley andHensen, 1990; Kelly et al., 2000). The south-west-ern Napier Complex is now thought to be boundby Rayner Complex material, reworked by theeastward extension of Early Palaeozoic Lutzow–Holm Complex tectonism (520–550 Ma; Shirashiet al., 1997).

The geological evolution of the Napier Com-plex presented here is condensed from Sheraton etal. (1987) and Harley and Black (1987) and repre-sents the generally accepted framework for theterrane. The Complex is dominated by tonalitic togranitic orthogneisses, with locally extensive lay-ered paragneiss sequences, which include quartz-ites, ironstones, aluminous meta-pelites and-psammites. Mafic and ultramafic units arepresent, usually as narrow layers and boudinagedpods within felsic gneisses and layered paragneisssequences. Calc-silicates and anorthosite bodiesare rare. The first recognised deformation, D1,resulted in the development of a regional intense,flat-lying, strongly-lineated S1 gneissosity and thetransposition and interleaving of lithologicalunits. The second deformation, D2, resulted inrefolding of S1 about F2 tight assymetrical foldswith recumbent to reclined axial planes but onlylocalised S2 development. The third deformation,D3, resulted in upright broad (1–10 km) foldingand produced regional dome and basin structures.In the Field Islands (Fig. 1), D3 was sufficientlyintense to have developed a penetrative fabric, butover much of the Napier Complex S3 developmentwas not pervasive and only locally developed inF3 hinges. UHT metamorphic conditions arethought to have evolved during D1–D2 with up-per-amphibolite conditions prevailing during D3,after which the Complex was essentially cra-tonised. Petrological studies have demonstratedthat the P–T evolution of the Napier Complexfollowed a near-isobaric cooling trajectory follow-

C.J. Carson et al. / Precambrian Research 116 (2002) 237–263 239

ing peak UHT conditions (Harley and Hensen,1990). Tholeiitic dyke swarms were emplaced at�2350 and �1190 Ma (Sheraton and Black,1981), preceding the development of various lo-calised amphibolite-facies shear zones and pseu-dotachylites (D4–D5; Harley, 1985b; Sandiford,1985), movement along which is thought to haveresulted in partial exhumation of the terrane.Features associated with D4−D5 have been gen-erally attributed to the local manifestation ofgranulite-facies tectonism which developed in theadjacent Neoproterozoic Rayner Complex (�1000 Ma), though this is largely based on indirectisotopic evidence, such as imprecise U–Pb lowerconcordia intercepts. The precise timing of these

shear zones and pseudotachylites remain some-what speculative. The final rock-forming eventsoccurred with the emplacement of EarlyPalaeozoic planar pegmatites (500–522 Ma; Grewand Manton, 1979; Black et al., 1983a; Carson etal., 2002) and rare mafic alkaline dykes (466–482Ma; Black and James, 1983; Miyamato et al.,2000).

2.1. Pre�ious geochronology

A considerable amount of geochronological in-formation has been obtained from the NapierComplex. Initial work focused of the identifica-tion of very old crustal fragments (e.g. Sobotovich

Fig. 1. Location of Tonagh Island (indicated in box) within Archaean Napier Complex. Various localities discussed in the text arehighlighted and labelled. Boundary between Napier and Rayner Complexes (dashed with question marks) after Harley and Black(1987) and Harley and Hensen (1990). Regional extent of sapphirine+quartz assemblages delineated by dashed line with barbs(after Harley and Motoyoshi, 2000); the area enclosed within this line also approximately delineates the areal extent of UHTmetamorphism. Upper inset: location of Napier Complex within Antarctic continent. Lower inset: enlargement of Casey Bay region,with localities discussed in text (from Black et al., 1983a).


et al., 1976; Lovering, 1979) and was refined anddeveloped by subsequent work (e.g. DePaolo etal., 1982; Black et al., 1983a,b, 1986a,b; Harleyand Black, 1997). Much of the research conductedsince the early 1980’s has focused on the construc-tion of a structural-metamorphic timeframe forthe evolution of the Napier Complex.

The generally adopted chronological schemewas largely developed by L.P. Black and cowork-ers during the early to mid 1980’s, with subse-quent refinement by Harley and Black (1997), theresults of which are summarised here. Initial con-straints on the timing of D1 were based on aRb–Sr whole-rock isochron age of 3072�34 Ma(Sheraton and Black, 1983) from a syn-S1 or-thogneiss from Proclaimation Island (Fig. 1), alocation somewhat distal to the region of extrememetamorphism recorded in the Tula Mountainsand Amundsen Bay regions to the SW. Based onthese data, a number of workers maintained thatUHT metamorphism occurred around this time(Black and James, 1983; Black et al., 1983a,b,1986a,b). This estimate was later reassessed by theSHRIMP study of Harley and Black (1997) whoconcluded that UHT metamorphism and UHT-related D1 could be no older than �2840 Ma,based largely on the emplacement age of a syn-D1

orthogneiss at Dallwitz nunatuk, located withinthe region of UHT metamorphism. They pro-posed a best estimate of 2837�15 Ma for thetiming of ‘highest grade metamorphism’, evalu-ated on zircon populations from three widelyspaced samples.

The timing of D2 has been subject to someuncertainty. Initially, the timing of D2 was largelybased on a Rb–Sr whole rock isochron age of2840 +220/−280 Ma (Black et al., 1986a) froma granite body at Mt Bride (Fig. 1) for which asyn-D2 emplacement age was inferred. An U–Pbzircon age of 2948 +38/−17 Ma from Mt Sonesis also considered to represent an event at thistime (Black et al., 1986b; Black, 1988). Other, lesswell-constrained, U–Pb ages have also been usedas evidence to constrain the timing of D2. Theseinclude a concordant population of zircon ionprobe U–Pb ages from a paragneiss from MtRiiser-Larsen and a population of conventionalU–Pb analyses from Mt Tod and Dallwitz Nuna-

tak (Black and James, 1983), all of which indicateevidence of an event at �2900 Ma. Harley andBlack (1997) revised these earlier estimates on thetiming of D2 based on their additional zircon dataand stated that the entire D1–D2 evolution mayhave been achieved during the time interval of2850–2820 Ma, followed by an extended periodof near-isobaric cooling.

Black et al. (1983a) presented conventional U–Pb analyses and centimetre-scale whole-rock Rb–Sr isochrons from felsic gneisses sampled from theCasey Bay region (Fig. 1) which record evidencefor an event at �2450–2480 Ma. Black et al.(1983a,b)) concluded that these data indicate re-setting of Rb–Sr systems, zircon growth and ex-tensive resetting of the U–Pb system inpre-existing zircon during D3, the effects of whichare pervasive in Casey Bay. A conventional zirconU–Pb age from a syn-D3 pegmatite (AyatollahIsland; Fig. 1), yielded an upper intercept of 2456+8/−5 Ma (Black et al., 1983a); a date that hasbecome widely accepted as the best estimate forthe timing of D3 (Black and James, 1983; Black etal., 1983a,b, 1986a,b; Sheraton et al., 1987;Harley and Black, 1997). Numerous studies, usinga variety of isotopic systems across the NapierComplex, report ages similar or identical to thisvalue and these have been interpreted as repre-senting isotopic disturbance during D3.

However, the above chronological scheme forthe evolution of D1–D3 has not been universallyaccepted. While there is little dispute concerningthe antiquity of various protoliths across theNapier Complex, a number of authors have ex-pressed an alternative view for the timing of theUHT metamorphism and structural evolution(Grew and Manton, 1979; Grew et al., 1982;Grew, 1998; DePaolo et al., 1982). These authorsmaintain that the timing of UHT metamorphismcan be readily explained in terms of a short-livedmetamorphic event at �2450–2480 Ma (Grew,1998), in contrast to the much earlier D1–D2

tectonothermal evolution indicated above (Harleyand Black, 1987, 1997). This alternative viewevolved primarily over differing interpretations ofsyn-tectonic pegmatites in the Casey Bay region(Fig. 1). The syn-D3 pegmatite dated by Black etal. (1983a); 2456 +8/−5 Ma) is instead inter-


Fig. 2. Sample localities on Tonagh Island, including the sample site of Shirashi et al. (1997). Lithological subdivision boundaries(dashed and labelled) after Osanai et al. (1999).

preted to date UHT metamorphism based on thepreservation of UHT mineral assemblages in struc-turally similar pegmatites (Grew, 1998). Sandifordand Wilson (1984) also suggest that such peg-matites were emplaced during D2 boudinage, dur-ing UHT conditions. In this contribution, wedocument new zircon ion probe U–Pb data froma felsic orthogneiss exposed on central TonaghIsland and provide additional constraint on thetiming of UHT metamorphism and of subsequentevents that affected the Napier Complex.

3. Tonagh Island orthogneiss

Tonagh Island, the site of this study, is locatedin Amundsen Bay, south-western Napier Complex(Fig. 1). The geology of Tonagh Is. has beendescribed in detail by Osanai et al. (1999) and, likethe majority of the Napier Complex, comprises avariety of felsic to intermediate orthogneisses, mi-nor mafic and ultramafic granulites and subordi-nate layered paragneiss sequences that includequartzites and ironstones. Osanai et al. (1999) also

subdivided the island into five lithological subdivi-sions (units I–V). These units represent coherentblocks of rock associations, the boundaries ofwhich are marked by steeply dipping to verticalshear zones of unknown age.

One sample (sample 27) for U–Pb zircon analy-sis was taken from an orthopyroxene-bearingquartzo-feldspathic orthogneiss (after the map unitof Osanai et al., 1999; hereafter called the TonaghIsland orthogneiss in this contribution), the loca-tion of which is shown in Fig. 2. This unit iscommon on Tonagh Island and is light brown- tobuff-coloured, homogeneous in appearance, con-taining an intense S1 gneissosity defined by 1–10mm scale subtle discontinuous, quartzo-feldspathicrich segregations. Typically the unit contains an-hedral quartz (�35% mode), K-feldspar (�14%)and plagioclase (�22%) and large (�2–3 mmdiameter) subhedral mesoperthite (10–20%), or-thopyroxene (3–5%) and minor clinopyroxene (�1%), with accessory ilmenite-magnetite, apatite,zircon and rare monazite (typically as mantles onapatite). A typical chemical analysis is listed inTable 1 (sample 27) and is further discussed below.


In addition to sample 27, multiple sampleswere collected from an alteration selvedge orzone immediately adjacent to an EarlyPalaeozoic pegmatite. The location of samplescollected from the alteration selvedge (30, 28/6,28/1 and 28/2) are illustrated in Fig. 3. Withinthe alteration selvedge, the pyroxene-bearing or-thogneiss had been extensively re-hydrated andrecrystallised by aqueous fluids that accompa-nied pegmatite emplacement, resulting in re-placement of pyroxene with biotite–hornblende–epidote assemblages at upper amphibolite facies

conditions (Carson et al., 2002). These sampleswere collected originally in order to assess theeffect of Early Palaeozoic aqueous fluid influxon the host orthogneiss zircon population, theresults of which have been presented in Carsonet al. (2002). However, additional data and fur-ther assessment of the zircon U–Pb data ofCarson et al. (2002) is presented here within thecontext of the regional chronological evolutionof the Napier Complex.

4. Analytical procedures

Zircons were extracted from bulk hand sam-ples (�5 kgs) and were separated using stan-dard heavy liquid and magnetic techniques,mounted in epoxy, together with zircon standardAS3, (near concordant TIMS 206Pb/238U age of1099.1�0.5 Ma; Paces and Miller, 1993). Themounted grains were equatorially sectioned, pol-ished and Au-coated. Backscattered electron(BSE) imaging was conducted to assess internalzircon structure and to facilitate selection ofanalysis sites (Fig. 4a–h) and were obtained us-ing the JEOL JXA-8600 electron microprobe lo-cated at Yale University. Zircon U–Pb ionmicroprobe analyses were conducted on aCAMECA IMS 1270, located at the Departmentof Earth and Space Sciences, University of Cali-fornia, Los Angeles, USA. The analytical proto-cols for zircon analysis are summarized inQuidelleur et al. (1997). All isotopic ratios werecorrected, where necessary, for common lead us-ing measured 204Pb. Common Pb ratios used are206Pb/204Pb=16.2, 207Pb/204Pb=15.35. Uraniumand Th concentrations were estimated semi-quantitatively using U/94Zr2O and Th/94Zr2O inunknowns relative to that of the AS3 standardand normalised against published U and Thconcentrations (Paces and Miller, 1993). Datareduction and processing were performed usingZIPS v2.4.1 (C.D. Coath 2000, unpublishedsoftware) and ISOPLOT v2.3 of Ludwig (1999).

Zircon U–Pb isotope results have been pre-viously published in part (Carson et al., 2002;alteration zone samples 28/6, 28/1 and 28/2;unaltered protolith, sample 27; Fig. 3).

Table 1Chemical composition of Tonagh Island orthogneiss

Sample 27Oxides

71.08SiO2

TiO2 0.44Al2O3 14.05Fe2O3 3.19MnO 0.04MgO 0.90CaO 2.86Na2O 3.29K2O 3.67P2O5 0.14LOI 0.32

99.98Total

TraceBa 1355

6Nb69Rb

257Sr13Y

167ZrCe (0.1) 90.0

0.4U (0.1)Th (0.1) 5.8Pb (2) 14

220.7K/RbCeN/YN 15.5Th/U 14.5

Bulk rock geochemistry was conducted using XRF, conductedby X-ray assay laboratories (XRAL) in Don Mills, Ont.,Canada. Major elements were determined on fused glass discs,trace elements (Ba, Nb, Rb, Sr, Y and Zr) on pressed powderpellets and Ce, U, Th and Pb determined by ICPMS tech-niques (detection limits shown). Results shown are average oftwo samples for major and trace elements, from one samplefor ICPMS elements.


Fig. 3. Photograph and sketch map of sample site within the Tonagh Island orthogneiss adjacent to an Early Palaeozoic pegmatite.Note conspicuous ‘bleach’ zone adjacent to pegmatite (see also Carson et al., submitted for publication). The visible outer boundaryof alteration zone indicates pyroxene-out isograd in the host orthogneiss. Samples 28/6, 28/1 and 28/2, 30 are 10, 160, 300 and 1100mm respectively from left margin of the pegmatite; sample 27, 20 m to the left of pegmatite out of photograph. Photographer facingeast, geologist for scale, sample location 67°05�37.3�S, 050°17�09.8�E.

We now evaluate these data presented by Carsonet al. (2002) within the context of the regionalgeochronological evolution of the Napier Com-plex, and present additional isotopic data (Tables2–4). Errors ellipses shown in Figs. 5–9 are atthe one sigma level, weighted mean and interceptages quoted are at the two sigma level and in-clude U decay constant uncertainties.

5. Results

Zircons from the protolith and the alterationselvedge have similar morphologies. Most com-mon are euhedral to subhedral, clear to honeycoloured grains, �100–250 �m in length, thathave aspect ratios of �2–3. Under back scat-tered electron imaging, sectioned grains may ex-hibit fractured well defined cores which mayshow micron-scale oscillatory zoning (Fig. 4a

and b). Typically this core region is overgrownby a volumetrically dominant mantle, which al-though generally homogeneous (Fig. 4b and c),may exhibit oscillatory zoning at the micronscale (Fig. 4d). The mantle may crosscut the coreregions (Fig. 4a) but more typically shows a con-formable relationship with them. Another majorpopulation consists of anhedral spherical glassy‘gem-like’ pinkish grains, up to 20–500 �m indiameter (Fig. 4g and h), which display minimalstructural development, and lack a clearlydefined euhedral core or micron scale oscillatoryzoning, although some coarse (10–20 �m) zoningmay be present (Fig. 4g). A narrow (�20 �m),low U outer rim is typically present on mostzircons (Fig. 4a–g) and may crosscut delicateoscillatory zoning present in the mantle (Fig. 4cand d). In total, 133 U–Pb analyses were con-ducted on 70 zircon grains; Tables 2–4 sum-marise the isotopic data.


Fig. 4. BSE images of representative zircon morphologies. All annotated ages are 207Pb/206Pb ages, which may differ from upperintercept ages quoted in the text depending on degree of discordance, scale bar in all images=100 �m: (a) sample number 28/2,grain 3, (disc 1), note zoned �2600 Ma core, overgrown by faintly zoned mantle and outer rim. Analysis point with 2526 Ma agerejected due to overlap with significantly older core; (b) sample number 28/2, grain 6 (disc 1), finely zoned �2600 Ma core,overgrown by wide homogenous mantle which grades into dark outer rim; (c) sample number 28/3 grain 14 (disc 1), unzoned �2600Ma core, overgrown by zoned mantle and outer rim with �2550 Ma ages. Note the transgressive nature of the outer rim at thearrowed location; (d) sample 28/1, grain 5 (disc 7). Finely zoned core (reset to �2423 Ma?) overgrown by mantle and outer rim,outer rim appears to truncate mantle (e) sample 28/1, grain 7 (disc 7), coarsely zoned �2550 Ma euhedral core over grown by outerrim; (f) sample number 28/2, grain 1 (disc 7) (g) sample number 27, grain 14 (disc 1), auhedral grain with coarse concentric zoning.(h) sample 28/6, grain 15 (disc 7). Homogenous anhedral grain with negligible zoning.


Tab

le2

U–P

bag

esan

dis

otop

icco

mpo

siti

ons

ofse

lect

edzi

rcon

core

sfr

omT

onag

hIs

land

,A

mun

dsen

Bay

Age

(Ma)

��

%20

6 Pb*

207 P

b*/23

5 UA

ge(M

a)�

Age

(Ma)

206 P

b*/23

8 U�

207 P

b*/20

6 Pb*

��

Th/

UD

isc.

%U

Th

Spot

Dat

e�sa

mpl

elo

cati

on(p

pm)

206 P

b/23

8 Uc

.dis

cc

.gra

inc

.spo

tc(p

pm)

207 P

b/20

6 Pb

207 P

b/23

5 U

10.5

2594

99.7

11.8

700.

146

25.5

0.49

830.

0059

30.

1728

0.00

1080

0.4

100.

974

826

0733

211

.620

00-2

Dec

�25

85D

efine

d28

1.d1

.gr1

2.sp

1co

re14

.725

8999

.411

.800

0.22

433

.90.

4870

0.00

783

0.17

5825

580.

0015

6017

.80.

397

.976

522

120

00-2

Dec

�26

14D

efine

d28

1.d1

.gr1

5.sp

1co

re15

.526

3914

.699

.311

.960

0.19

80.

4859

0.00

588

0.17

850.

0015

700.

196

.752

368

2000

-2D

ec�

Defi

ned

25.5

2553

2601

281.

d1.g

r16.

sp1

core

15.9

30.2

99.5

2584

11.9

2025

980.

216

16.9

0.49

310.

0070

00.

1753

0.00

1670

0.9

99.0

384

360

2000

-2D

ec�

2609

Defi

ned

281.

d1.g

r8.s

p1co

re11

.825

2199

.410

.970

0.22

937

.20.

4445

0.00

833

0.17

9023

710.

0012

7019

.40.

589

.741

023

720

00-2

Dec

�26

44D

efine

d28

2.d1

.gr3

.sp1

core

26.4

2634

12.6

99.6

12.1

000.

340

0.49

290.

0137

00.

1780

0.00

1350

0.8

98.1

568

479

2000

-2D

ec�

Defi

ned

59.1

2583

2612

282.

d1.g

r6.s

p1co

re46

.526

1424

.225

5499

.011

.370

0.30

924

780.

4688

0.01

060

0.17

590.

0025

500.

694

.824

915

120

00-2

Dec

�25

.4D

efine

dco

re28

2.d1

.gr8

.sp1

10.2

2672

99.6

12.9

000.

139

20.7

0.51

980.

0048

70.

1799

0.00

1100

0.4

2698

101.

710

.283

839

920

00-2

Dec

�26

52D

efine

d28

3.d1

.gr1

4.sp

1co

re

%20

6 Pb*

deno

tes

perc

enta

geof

radi

ogen

ic20

6 Pb;

unce

rtai

ntie

sar

egi

ven

atth

e1�

leve

l.U

and

Th

cont

ents

are

sem

iqua

ntit

ativ

ees

tim

ates

base

don

U/94

Zr 2

Oan

dT

h/94

Zr 2

Ore

lati

veto

that

ofm

easu

red

valu

esof

AS3

stan

dard

and

publ

ishe

dU

and

Th

conc

entr

atio

ns.

Dis

c.%

isth

ein

divi

dual

disc

orda

nce=

(206 P

b/23

8 Uag

e/20

7 Pb/

206 P

bag

e−1×

100)

+10

0,10

0%=

conc

orda

ntan

alys

is.


Tab

le3

U–P

bag

esan

dis

otop

icco

mpo

siti

ons

ofse

lect

edzi

rcon

sfr

omT

onag

hIs

land

,A

mun

dsen

Bay

�A

ge(M

a)�

Th

�A

ge(M

a)%

206 P

b*20

7 Pb*

/235 U

�20

6 Pb*

/238 U

�20

7 Pb*

/206 P

b*�

Th/

UD

isc.

%Sp

otU

Dat

e�sa

mpl

eA

ge(M

a)(p

pm)

loca

tion

c.d

isc

c.g

rain

c.s

potc

207 P

b/23

5 Ua)

207 P

b/20

6 Pb

(ppm

)20

6 Pb/

238 U

)

6.35

99.9

11.9

400.

124

2548

0.51

250.

0054

20.

1690

0.00

0640

0.5

9.72

104.

726

6711

2161

720

00-3

0Nov

�23

.1D

efine

d26

00co

re28

2.d7

.gr1

1.sp

325

2920

.79.

5199

.711

.410

0.11

425

580.

4953

0.00

480

0.16

710.

0009

471.

110

2.5

685

813

2000

-30N

ov�

9.34

Defi

ned

2593

282.

d7.g

r12.

sp2

core

6.54

99.5

11.8

100.

128

2590

0.50

670.

0047

50.

1690

0.00

0660

20.3

1.0

2643

103.

765

470

020

00-3

0Nov

�10

.1D

efine

d25

48co

re28

2.d7

.gr1

6.sp

29.

0525

407.

9199

.811

.660

0.11

326

260.

5029

0.00

520

0.16

820.

0007

940.

510

3.4

655

331

2000

-30N

ov�

2578

Defi

ned

22.3

282.

d7.g

r6.s

p2co

re25

5019

.410

.799

.510

.950

0.14

625

190.

4695

0.00

443

0.16

920.

0010

800.

197

.356

272

2000

-30N

ov�

12.4

Man

tle

2481

283.

d7.g

r1.s

p113

.798

.311

.240

0.17

014

.10.

4833

0.00

595

0.16

870.

0013

8025

430.

325

4599

.932

797

2000

-1D

ec�

2542

Defi

ned

25.9

281.

d7.g

r7.s

p1co

re17

.725

3915

.899

.410

.900

0.20

724

850.

4703

0.00

763

0.16

810.

0015

800.

197

.968

258

2000

-2D

ec�

2515

Man

tle

33.4

281.

d1.g

r3.s

p38.

7299

.511

.480

0.15

824

.00.

4850

0.00

552

0.17

170.

0008

960.

499

.025

4989

725

6343

920

00-2

Dec

�12

.9M

antl

e25

7428

3.d1

.gr1

4.sp

211

.299

.411

.520

0.21

217

.20.

4963

0.00

764

0.16

840.

0011

3025

660.

132

.910

2.2

500

5720

00-2

Dec

�25

98D

efine

d25

42co

re28

6.d1

.gr1

0.sp

125

2636

.114

.299

.510

.170

0.19

024

500.

4420

0.00

808

0.16

680.

0014

100.

393

.457

521

320

00-2

Dec

�17

.3M

antl

e23

6028

2.d1

.gr3

.sp2

%20

6 Pb*

deno

tes

perc

enta

geof

radi

ogen

ic20

6 Pb;

unce

rtai

ntie

sar

egi

ven

atth

e1�

leve

l.U

and

Th

cont

ents

are

sem

iqua

ntit

ativ

ees

tim

ates

base

don

U/94

Zr 2

Oan

dT

h/94

Zr 2

Ore

lati

veto

that

ofm

easu

red

valu

esof

AS3

stan

dard

and

publ

ishe

dU

and

Th

conc

entr

atio

ns.

Dis

c.%

isth

ein

divi

dual

disc

orda

nce=

(206 P

b/23

8 Uag

e/20

7 Pb/

206 P

bag

e−1×

100)

+10

0,10

0%=

conc

orda

ntan

alys

is.


Tab

le4

U–P

bag

esan

dis

otop

icco

mpo

siti

ons

ofse

lect

edzi

rcon

sfr

omT

onag

hIs

land

,A

mun

dsen

Bay

Age

(Ma)

Age

(Ma)

��

%20

6 Pb*

207 P

b*/23

5 UA

ge(M

a)�

�20

6 Pb*

/238 U

�20

7 Pb*

/206 P

b*�

Th/

UD

isc.

%U

Spot

Th

Dat

e�sa

mpl

ec

.dis

cc

.gra

inc

.spo

tc(p

pm)

loca

tion

(ppm

)20

6 Pb/

238 U

207 P

b/20

6 Pb

207 P

b/23

5 U

81.7

24.4

99.0

10.3

400.

442

2465

0.46

500.

0186

00.

1612

0.00

233

3.2

99.8

7827

720

00-3

0Nov

�39

.6O

uter

2462

2468

282.

d7.g

r1.s

p3ri

m97

.59.

256

0.51

050

.50.

4340

0.01

280

0.15

470.

0063

03.

223

6496

.923

9876

267

2000

-30N

ov�

69.3

Out

er23

2457

.7ri

m28

2.d7

.gr1

1.sp

197

.99.

798

0.37

048

.50.

4510

0.01

090

0.15

760.

0053

11.

024

0098

.824

1611

111

420

00-3

0Nov

�34

.8O

uter

2430

57.1

282.

d7.g

r11.

sp2

rim

97.7

10.2

900.

481

43.3

0.46

830.

0176

00.

1594

0.00

386

2.8

2461

101.

124

4965

196

2000

-30N

ov�

41.0

Out

er24

7677

.2ri

m28

2.d7

.gr1

6.sp

124

7646

.997

.810

.940

0.51

825

180.

4899

0.01

740

0.16

200.

0045

02.

110

3.8

6013

620

00-3

0Nov

�25

70O

uter

44.0

75.2

rim

282.

d7.g

r8.s

p124

5725

.398

.411

.310

0.32

825

490.

5122

0.01

460

0.16

010.

0024

01.

210

8.5

154

209

2000

-30N

ov�

2666

Out

er27

.162

.128

2.d7

.gr9

.sp2

rim

97.5

10.6

100.

632

55.3

0.49

460.

0212

00.

1555

0.00

613

4.0

2489

107.

624

0847

206

2000

-30N

ov�

67.0

Out

er25

9091

.6ri

m28

3.d7

.gr1

.sp2

2504

25.9

98.4

11.7

600.

262

2692

0.51

820.

0068

90.

1646

0.00

253

1.5

107.

516

327

120

00-3

0Nov

�25

86O

uter

29.2

20.9

rim

283.

d7.g

r10.

sp3

98.8

10.7

000.

366

31.8

0.47

690.

0143

00.

1627

0.00

229

2497

1.0

62.6

101.

216

717

920

00-3

0Nov

�25

14U

G24

8423

.828

3.d7

.gr1

6.sp

143

.560

.897

.69.

254

0.44

723

630.

4195

0.00

957

0.16

000.

0057

63.

392

.094

338

2000

-30N

ov�

44.3

Out

er22

5824

55ri

m28

3.d7

.gr1

8.sp

224

1629

.498

.79.

622

0.25

623

990.

4465

0.00

939

0.15

630.

0027

03.

498

.594

347

2000

-30N

ov�

2380

UG

24.5

41.8

283.

d7.g

r19.

sp1

97.7

9.80

70.

531

2474

0.43

980.

0157

00.

1617

0.00

528

2.6

49.9

95.0

55.1

9426

720

00-3

0Nov

�23

50O

uter

70.2

2417

rim

283.

d7.g

r3.s

p268

.255

.297

.710

.390

0.53

124

710.

4622

0.01

550

0.16

310.

0053

43.

698

.476

303

2000

-30N

ov�

47.3

Out

er24

4924

88ri

m28

3.d7

.gr4

.sp2

98.2

10.4

200.

442

39.3

0.47

900.

0130

00.

1578

0.00

350

2473

2.9

56.8

103.

710

031

420

00-3

0Nov

�25

23M

antl

e24

3337

.628

3.d7

.gr5

.sp2

98.1

10.7

200.

409

61.9

0.47

980.

0142

00.

1620

0.00

353

2.3

2526

102.

024

9978

193

2000

-30N

ov�

35.5

Man

tle

2477

36.7

283.

d7.g

r7.s

p297

.010

.800

0.68

624

770.

4834

0.01

600

0.16

210.

0070

06.

510

2.6

59.0

4972

.835

220

00-3

0Nov

�25

42C

ore

69.5

2506

ofU

G28

6.d7

.gr1

0.sp

1


Tab

le4

(con

tinu

ed)

�A

ge(M

a)T

hU

�%

206 P

b*A

ge(M

a)20

7 Pb*

/235 U

�A

ge(M

a)�

206 P

b*/23

8 USp

ot�

207 P

b*/20

6 Pb*

�T

h/U

Dat

e�sa

mpl

eD

isc.

%20

7 Pb/

235 U

(ppm

)(p

pm)

c.d

isc

c.g

rain

c.s

potc

207 P

b/20

6 Pb

loca

tion

206 P

b/23

8 U

81.4

96.1

9.75

20.

613

0.45

570.

0170

075

.10.

1552

0.00

744

2421

5.8

2412

100.

736

228

2000

-30N

ov�

57.9

Out

er24

0428

6.d7

.gr1

0.sp

2ri

mof

UG

4.6

99.9

10.6

500.

094

0.48

650.

0042

024

930.

1587

0.00

043

18.2

0.1

2555

104.

616

1713

520

00-3

0Nov

�8.

2O

uter

2442

rim

286.

d7.g

r5.s

p123

.024

2719

.899

.310

.220

0.25

40.

4710

0.01

010

2488

0.15

730.

0018

40.

810

2.5

214

178

2000

-30N

ov�

2455

Out

er44

.128

2.d7

.gr1

2.sp

1ri

m7.

124

326.

010

0.0

10.4

800.

080

0.48

180.

0031

225

350.

1577

0.00

056

0.1

104.

223

6732

820

00-3

0Nov

�13

.6R

eset

2478

283.

d7.g

r10.

sp1

core

22.4

99.3

10.2

500.

275

0.47

340.

0130

024

570.

1570

0.00

208

56.7

1.1

24.8

103.

120

525

120

00-3

0Nov

�24

24O

uter

2499

283.

d7.g

r10.

sp4

rim

19.3

2450

19.5

99.0

10.3

400.

215

0.47

030.

0075

924

850.

1594

0.00

184

1.4

101.

416

825

420

00-3

0Nov

�24

66O

uter

33.3

283.

d7.g

r5.s

p3ri

m8.

324

335.

099

.910

.010

0.09

00.

4600

0.00

381

2440

0.15

790.

0004

70.

110

0.3

1611

242

2000

-30N

ov�

16.8

Res

et24

36co

re28

3.d7

.gr7

.sp1

33.9

98.5

11.1

500.

336

0.49

840.

0090

60.

1623

39.0

0.00

326

2607

2.2

105.

217

541

520

00-1

Dec

�O

uter

2536

28.1

2479

281.

d7.g

r20.

sp2

rim

78.2

96.3

10.7

000.

735

0.49

490.

0229

025

920.

1568

0.00

722

2.5

107.

124

2158

63.8

155

2000

-1D

ec�

2497

Out

er98

.9ri

m28

1.d7

.gr5

.sp3

2465

42.0

31.4

98.2

10.4

300.

285

0.47

010.

0095

824

740.

1609

0.00

299

2.0

100.

811

524

620

00-1

Dec

�25

.3O

uter

2484

281.

d7.g

r7.s

p3ri

m70

.224

8521

.698

.810

.050

0.39

40.

4478

0.01

580

2440

0.16

280.

0020

81.

396

.017

024

620

00-1

Dec

�23

85U

G36

.228

2.d7

.gr8

.sp2

81.3

96.7

10.7

600.

797

0.48

040.

0216

068

.90.

1624

0.00

783

2502

5.9

2481

101.

957

365

2000

-1D

ec�

2529

UG

93.9

286.

d7.g

r15.

sp1

2448

44.1

52.7

98.1

10.2

700.

437

0.46

200.

0120

024

590.

1612

0.00

421

1.0

99.2

188

202

2000

-2D

ec�

39.4

Out

er24

6828

1.d1

.gr1

5.sp

2ri

m50

.498

.010

.910

0.53

30.

4846

0.01

300

2516

0.16

330.

0048

956

.62.

225

4710

2.3

137

325

2000

-2D

ec�

45.5

Out

er24

90ri

m28

1.d1

.gr1

6.sp

322

.899

.09.

973

0.26

70.

4495

0.01

090

48.3

0.16

090.

0021

723

930.

724

3297

.125

518

520

00-2

Dec

�24

.7O

uter

2465

281.

d1.g

r8.s

p3ri

m65

.996

.210

.370

0.71

20.

4789

0.02

350

2468

0.15

700.

0061

010

2.0

4.7

63.6

104.

156

294

2000

-2D

ec�

2423

Man

tle

2523

286.

d1.g

r10.

sp2


Tab

le4

(con

tinu

ed)

�A

ge(M

a)A

ge(M

a)D

isc.

%�

��

%20

6 Pb*

207 P

b*/20

6 Pb*

207 P

b*/23

5 U�

�20

6 Pb*

/238 U

Age

(Ma)

UT

h/U

Th

Dat

e�sa

mpl

eSp

ot(p

pm)

(ppm

)20

7 Pb/

206 P

bc

.dis

cc

.gra

inc

.spo

tclo

cati

on20

7 Pb/

235 U

206 P

b/23

8 U

25.8

98.6

9.84

70.

280

0.44

3923

680.

1609

45.5

0.00

245

0.01

020

2.1

96.1

144

325

2000

-2D

ec�

Out

er24

2126

.224

6528

6.d1

.gr1

3.sp

3ri

m67

.096

.910

.300

0.63

30.

4588

0.01

760

56.9

0.16

280.

0064

73.

624

8597

.977

.778

311

2000

-2D

ec�

2462

Man

tle

2434

286.

d1.g

r14.

sp2

32.7

2487

30.4

98.6

10.0

300.

355

0.44

650.

0131

00.

1630

0.00

294

1.5

95.7

195

315

2000

-2D

ec�

Man

tle

2380

2438

58.4

286.

d1.g

r4.s

p270

.997

.210

.760

0.61

00.

4894

0.01

370

2502

0.15

940.

0066

81.

924

4910

4.9

2568

132

278

2000

-2D

ec�

59.3

UG

52.7

30.d

1.gr

15.s

p125

0745

.726

.199

.011

.040

0.30

20.

4855

0.01

050

2527

0.16

490.

0025

60.

610

1.8

393

261

2000

-2D

ec�

25.4

UG

2551

30.d

1.gr

18.s

p124

6353

.238

.197

.910

.390

0.37

90.

4688

0.01

210

2470

0.16

070.

0036

32.

510

0.6

101

277

2000

-2D

ec�

33.8

UG

2478

30.d

1.gr

18.s

p2r

44.2

97.6

9.16

30.

358

0.43

800.

0141

035

.80.

1517

0.00

393

2354

7.1

2365

99.0

6449

720

00-2

Dec

�23

42U

G63

.330

.d1.

gr19

.sp1

r54

.724

3327

.098

.89.

844

0.28

80.

4523

0.01

230

2420

0.15

780.

0025

24.

098

.814

966

120

00-2

Dec

�24

05U

G27

.030

.d1.

gr19

.sp2

r34

.998

.49.

633

0.30

40.

4361

0.00

973

29.1

0.16

020.

0033

124

001.

643

.794

.919

633

820

00-2

Dec

�23

33M

antl

e24

5830

.d1.

gr2.

sp1

58.5

2473

65.0

96.9

9.47

00.

528

0.42

490.

0129

023

850.

1616

0.00

622

1.6

92.3

181

323

2000

-2D

ec�

2283

Out

er51

.230

.d1.

gr2.

sp3

rim

27.9

2477

10.6

99.5

10.2

800.

151

0.46

020.

0063

324

600.

1620

0.00

101

0.4

98.5

616

239

2000

-2D

ec�

2440

Cor

e13

.630

.d1.

gr4.

sp1

ofU

G18

.699

.411

.200

0.25

50.

4934

0.00

848

2540

0.16

460.

0018

236

.60.

621

.310

3.2

531

326

2000

-2D

ec�

2504

Out

er25

8530

.d1.

gr4.

sp2

rim

ofU

G39

.724

7749

.798

.210

.580

0.45

30.

4736

0.01

200

2499

0.16

210.

0047

71.

710

0.9

146

269

2000

-2D

ec�

2487

Man

tle

52.4

30.d

1.gr

6.sp

193

.794

.29.

526

0.87

70.

4333

0.02

650

84.6

0.15

950.

0088

323

907.

111

9.0

94.7

3930

920

00-2

Dec

�23

21O

uter

2450

rim

30.d

1.gr

6.sp

2r51

.497

.39.

091

0.45

40.

4095

0.01

540

70.7

0.16

100.

0049

022

132.

223

4789

.713

632

220

00-2

Dec

�45

.7O

uter

2466

30.d

1.gr

6.sp

3rri

m27

.999

.010

.310

0.34

00.

4624

0.01

10.

1617

0.00

267

0.6

48.4

99.0

2463

376

236

2000

-2D

ec�

30.5

Out

er24

7424

50ri

m28

1.d1

.gr1

2.sp

3


Tab

le4

(con

tinu

ed)

�A

ge(M

a)A

ge(M

a)U

Th/

U�

�%

206 P

b*�

207 P

b*/23

5 U20

7 Pb*

/206 P

b*�

206 P

b*/23

8 U�

Age

(Ma)

Dis

c.%

Th

Dat

e�sa

mpl

eSp

ot20

7 Pb/

235 U

206 P

b/23

8 U(p

pm)

(ppm

)c

.dis

cc

.gra

inc

.spo

tclo

cati

on20

7 Pb/

206 P

b

99.3

9.82

90.

194

0.44

280.

0066

70.

1610

0.00

196

0.7

29.8

95.8

2419

459

350

2000

-2D

ec�

18.2

Out

er24

6623

6320

.5ri

m28

6.d1

.gr4

.sp3

2442

19.2

99.3

9.70

10.

167

0.44

320.

0052

30.

1587

0.00

180

0.9

96.8

378

370

2000

-2D

ec�

2365

Out

er23

.424

0715

.828

1.d1

.gr3

.sp2

rim

98.2

10.3

400.

441

0.45

190.

0166

0.16

590.

0044

56.

039

.595

.524

0464

420

2000

-2D

ec�

73.8

Man

tle

2517

45.1

2465

27.d

1.gr

1.sp

142

.116

.099

.39.

970

0.23

10.

4413

0.00

941

0.16

380.

0015

61.

594

.427

043

820

00-2

Dec

�24

32U

G21

.323

5724

9627

.d1.

gr1.

sp2

2389

60.0

97.0

9.75

20.

499

0.45

970.

0163

0.15

390.

0054

25.

710

2.1

6037

420

00-2

Dec

�71

.9O

uter

2438

2412

47.2

rim

27.d

1.gr

14.s

p299

.310

.490

0.19

60.

4586

0.00

743

0.16

580.

0013

21.

417

.396

.724

3336

155

420

00-2

Dec

�25

16U

G24

7932

.813

.427

.d1.

gr3.

sp1

58.6

43.0

98.4

9.92

70.

386

0.45

240.

0132

0.15

910.

0040

41.

598

.311

618

620

00-2

Dec

�24

28M

antl

e35

.924

0624

4727

.d1.

gr3.

sp2

99.2

10.3

300.

280

0.45

930.

0115

0.16

310.

0019

71.

297

.927

324

3635

950

.720

00-2

Dec

�24

64U

G25

.124

8820

.327

.d1.

gr9.

sp1

99.7

10.4

700.

221

0.47

480.

0101

0.15

990.

0010

90.

910

2.1

2454

384

19.6

394

2000

-2D

ec�

2477

UG

11.5

2505

44.0

27.d

1.gr

14.s

p123

9999

.431

.99.

866

0.18

00.

4509

0.00

718

0.15

870.

0013

40.

898

.234

931

020

00-2

Dec

�24

22U

G16

.924

4214

.327

.d1.

gr2.

sp1

%20

6 Pb*

deno

tes

perc

enta

geof

radi

ogen

ic20

6 Pb;

unce

rtai

ntie

sar

egi

ven

atth

e1�

leve

l;U

G,

anhe

dral

unst

ruct

ured

grai

n.U

and

Th

cont

ents

are

sem

iqua

ntit

ativ

ees

tim

ates

base

don

U/94

Zr 2

Oan

dT

h/94

Zr 2

Ore

lati

veto

that

ofm

easu

red

valu

esof

AS3

stan

dard

and

publ

ishe

dU

and

Th

conc

entr

atio

ns.

Dis

c.%

isth

ein

divi

dual

disc

orda

nce=

(206 P

b/23

8 Uag

e/20

7 Pb/

206 P

bag

e−1×

100)

+10

0,10

0%=

conc

orda

ntan

alys

is.


A subgroup of analyses (eight analyses fromeight grains) have 207Pb/206Pb ages between �2600–2650 Ma (Fig. 5; Table 3). These analysesare all from texturally well-defined zircon cores(e.g. Fig. 4a–c), and have U and Th contentsbetween �250–840 ppm and �70–480 ppm re-spectively, with Th/U of between 0.1–0.9. Analy-ses from this population range from slightlyreversely discordant (102%, one analysis), to nor-mally discordant (90%). The data approximatesa linear trend (MSWD=4.7) with an upper con-cordia intercept of 2626�28 Ma and a veryimprecise lower intercept of 164�1200 Ma.

Another small subgroup of analyses (10 analy-ses from 10 grains) has 207Pb/206Pb ages between�2529–2575 Ma (Fig. 6; Table 3). Analysis sitesof this population are well defined variably-zoned cores (Fig. 4e) or mantles that overgrowcores of �2600 Ma age (Fig. 4c). Uranium andTh contents are variable, �320–1100 ppm and�60–810 ppm respectively, with Th/U ranging

from 0.1–1.1. This population displays a rangefrom reversely to normally discordant with amaximum of 104–93%, respectively. An errorweighted regression (MSWD=2.1) gives an up-per concordia intercept of 2546�13 Ma, withan imprecise lower intercept of 45�640 Ma.

Most analyses form a population with 207Pb/206Pb ages around �2450–2470 Ma (Fig. 7;Table 4 and including previously published datafrom Carson et al., 2002). Many of these analy-ses are from mantles and low-U outer rim sites,although several are from defined cores andmay represent isotopically reset magmatic cores(Fig. 4d). No concordant age could be calculatedfor these data (Ludwig, 1998), but they forma near linear trend, with some excess scatter,for which an error weighted regression (n=115,MSWD=4.0) gives an upper concordia interceptof 2476.9�9 Ma with a lower interceptof 531�200 Ma. In further considering thesedata, we subdivided these analyses into ei-

Fig. 5. Conventional concordia diagram showing zircon U–Pb isotope data from a population of old well-defined cores (Fig. 4a–c).Line is an error-weighted regression with an upper intercept of 2626�28 Ma and a poorly defined lower intercept of 164�1200Ma.


Fig. 6. Conventional concordia diagram showing zircon U–Pb isotope data from a selected population of analyses that form anerror-weighted regression with an upper intercept of 2546�13 Ma and a poorly defined lower intercept of 45�640 Ma.

ther high Th/U types or outer rims (which mayinclude analyses with high Th/U). Analyses withhigh Th/U (�1.2) define a linear trend with anerror weighted regression with no excess scatter(n=52, MSWD=1.03) and an upper concordiaintercept of 2483.0�9 Ma (Fig. 8). Common Pbis typically higher in analyses with elevated Th/U,and with lower U contents (Table 2). High Th/Uanalysis sites are commonly broad mantles, butmay also include occasional narrow outer rimsand, less commonly, indistinct unzoned cores. Alllow U outer rims (with ‘dark’ BSE response; Fig.4b–g), regardless of Th/U, were regressed to-gether and show a linear trend with no excessscatter (n=24, MSWD=0.67) and an upper in-tercept of 2455�16 Ma (Fig. 9). It is tempting tosuggest that the narrow outer rims may representthe later stage of a progression of zircon growth(�2455 Ma) which commenced with the growthof the high Th/U group at �2483 Ma. Treatmentas separate populations is supported by a MSWDof �1 for each group (high Th/U and the outerrim groups) when regressed separately.

6. Discussion

Prior to discussing the zircon isotopic databelow, it is appropriate to deliberate on the chem-ical characteristics of the Tonagh Island or-thogneiss (Table 1). Sheraton and Black (1983)made two broad distinctions of felsic to interme-diate orthogneiss units within the Napier Com-plex, ‘undepleted’ and ‘depleted’ varieties, thelatter defined by strong depletion of a variety ofelements and also on the basis of various elemen-tal ratios, including Th/U and chondrite nor-malised CeN/YN. Sheraton and Black (1983)concluded that depleted varieties acquired theirparticular geochemical signature primarily as aresult of partial melting of a garnet-bearing maficsource, rather than resulting from UHT metamor-phism, though this effect was also considered.This geochemical discrimination was used bySheraton and Black (1983) to characterise intru-sive suites that pre-dated UHT metamorphism.However, it should be stressed that chemical sig-


natures alone cannot provide a robust criteria forassessing orthogneiss emplacement relative toUHT metamorphism. The purpose of the follow-ing geochemical information is to further describethe nature of the orthogneiss under investigation.In presenting this data, we do not imply any timingrelationships of orthogneiss emplacement relativeto UHT conditions.

The Tonagh Island orthogneiss illustrates manyof the defining characteristics of ‘depleted’ or-thogneisses. This unit has low values of Y relativeto TiO2 (cf. Fig. 10 of Sheraton and Black (1983)),low Th (5.8 ppm) and U (0.4 ppm) and very highTh/U (14.5; Fig. 13 of Sheraton and Black (1983)).Relative to SiO2, Ce, Nb, Ti, Zr, Rb and Y all arecomparable with the depleted orthogneiss fieldsshown in Fig. 5 of Sheraton and Black (1983).Based on these geochemical characteristics, we

conclude that the Tonagh Island orthogneiss be-longs to the ‘depleted’ class of felsic orthogneissesdefined by Sheraton and Black (1983).

The small population of clearly defined zirconcores (n=8) with an upper intercept age of2626�28 Ma (Fig. 5) has Th/U values typicallyaround 0.3–0.5 (with three analyses at 0.1, 0.8 and0.9), values typical of magmatic zircon (Maas etal., 1992). It is also significant to note that thereare no ages older than �2650 Ma in 133 analysesof 70 zircon grains from this orthogneiss, preclud-ing any significant prior geological history. Themagmatic values of Th/U, together with the well-defined, micron-scale oscillatory zoned nature ofthe cores suggests an igneous origin. We concludethat the emplacement of this voluminous ‘de-pleted’ felsic intrusive on Tonagh Island occurredat 2626�28 Ma.

Fig. 7. Conventional concordia diagram showing all remaining zircon U–Pb isotope data. These data approximate a linear tread(n=115; MSWD=4.0) and an error-weighted regression yield an upper intercept of 2476.9�9 Ma and a lower intercept of531�200 Ma. The diagram is visually dominated by 59 analyses with large error ellipses; such analyses typically include high Th/Uand low U outer rims, that have elevated common Pb levels and have correspondingly large Pb/U errors. The high Th/U and lowU outer rim analyses are separately treated in Figs. 8 and 9 respectively. The inset at lower right shows a selected subset (n=56)of U–Pb analyses with Pb* contents �99% and correspondingly smaller errors.


Fig. 8. Conventional concordia diagram showing zircon U–Pb isotope data from high Th/U (�1.2) analyses, a subdivision of thedata presented in Fig. 7. These analyses form an error-weighted regression, with no excess scatter (n=52; MSWD=1.03) with anupper intercept of 2483.0�9 Ma and a poorly defined lower intercept of 370�350 Ma.

The population of ages with an upper concor-dia intercept of 2546�13 Ma (Fig. 6) is some-what problematic. Shirashi et al. (1997) also noteda small population of zircons with 207Pb/206Pbages �2550 Ma in their analysis of a quartz-feldspar gneiss from layered paragneiss sequenceon northern Tonagh Island (Fig. 2), but wereunclear on the significance of these. They specu-lated that the population may represent a episodeof new zircon growth. Recent SHRIMP analysison zircons extracted from a leucosome fromMcIntyre Is (Fig. 1) also yield U–Pb ages of2560–2590 Ma (S.L. Harley; unpublished data).The growth of new zircon is also suggested at thistime (S.L. Harley personal communication).These interpretations are consistent with our ob-servations in that analysis sites that yield �2550Ma represent either cores (Fig. 4e) or overgrowths(Fig. 4c) on older �2600 Ma cores. We acknowl-edge that there was a period of new zircon growthat 2546�13 Ma, with formation of new grainsand overgrowths on �2600 Ma cores, but main-

tain that the geological significance of this periodof zircon growth is currently unclear given thelimited geological and isotopic information onthis event.

6.1. Significance of the 2450–2470 Ma ages andthe timing of UHT metamorphism

The majority of zircon U–Pb analyses fromthis study form a population that have an upperconcordia intercept �2450–2470 Ma, statisticallyidentical with many isotopic ages from across theNapier Complex. Partial resetting of old zirconU–Pb systems from orthogneiss protoliths resultsin (generally imprecise) lower intercepts at �2400–2500 Ma (Black and James, 1983; Williamset al., 1984; Black et al., 1986a,b; Harley andBlack, 1997), interpreted to indicate a period ofmajor isotopic disruption. Upper intercepts ofother zircon populations also commonly indicateages of �2450–2470 Ma, which has been inter-preted as a period of zircon growth or of the


complete resetting of pre-existing zircon (e.g.Black et al., 1983a, 1986a,b). Rb–Sr and Sm–Ndwhole-rock isochrons ages also indicate a periodof extensive isotopic re-equilibration at this time(e.g. Black and James, 1983; Black et al., 1983a;Black and McCulloch, 1987; Owada et al., 1994;Tainosho et al., 1994, 1997). The interpretation ofthis strong isotopic signature at �2450–2480Ma, recorded from a wide variety of rock typesand isotopic systems is critical to models of thegeochronological-metamorphic framework of theNapier Complex.

Black et al. (1983a), for example, maintainedthat the widespread and strong isotopic signatureat �2460 Ma is due to the pervasive effects of D3

and recrystallisation associated with the develop-ment of S3, and that the D1–D2 and UHT evolu-tion occurred �400 Ma earlier, during the mid tolate Archaean. In contrast, Grew (1998) acknowl-edges that while all studies seem to agree thatisotopic systems were highly disturbed and resetat �2450–2480 Ma, he questions the assertion

that UHT metamorphism occurred �400 Maearlier. Grew (1998) maintains that the timing ofUHT metamorphism can be readily explained interms of a short-lived metamorphic event at �2450–2480 Ma in contrast to the much earlierD1–D2 tectonothermal evolution preferred byHarley and Black (1987, 1997). This alternativeinterpretation is largely based on the preservationof UHT assemblages within ‘syn-D3’ pegmatites(Grew, 1998) and syn-D2 UHT pegmatites (Grewet al., 1982; Sandiford and Wilson, 1984; Sandi-ford and Powell, 1986b) both of which yield up-per intercept ages of �2500 Ma (Grew et al.,1982). These contrasting interpretations must beaddressed within the context of the presentlyavailable data and take into account the followingfacts.

(1) In the Field Islands (Fig. 1) the effects of D3

are reported to be locally intense (Black et al.,1983a) and proceeded at upper-amphibolite faciesconditions (650–720 °C and 5–8 kbar; Harleyand Black, 1987). This D3 event is thought to

Fig. 9. Conventional concordia diagram showing zircon U–Pb isotope data from low U outer rim analyses only, a subdivision ofthe data presented in Fig. 7. These analyses form an error-weighted regression, with no excess scatter (n=24; MSWD=0.67) withan upper intercept of 2455�16 Ma and a poorly defined lower intercept of 628�580 Ma.


Fig. 10. Summary concordia diagram of the U–Pb isotopic results from two pre-UHT orthogneiss units (Tonalitic Gneiss, MtSones, Black and James, 1983; Williams et al., 1984; Harley and Black, 1997; and a charnockitic leucogneiss from Fyfe Hills, Blackand James, 1983). Of importance here is the lack of disturbance at the timing of UHT metamorphism (�2820–2850 Ma) favouredby Harley and Black (1997); indicated by the solid star). Although the lower intercept indicated by the study of Williams et al. (1984)is somewhat younger than �2500 Ma (2030�170 Ma), that study concluded the lower intercept was the likely result of disturbanceduring an event at �2400 Ma, but did not elaborate on the nature of that event.

have been responsible for extensive resetting ofRb–Sr whole-rock and U–Pb zircon systematicsat �2450–2480 Ma. However, intensity of D3

across the Napier Complex is quite varied, yet thisstrong isotopic imprint at �2450–2480 Ma isalmost always present. At the sample location forthis study (Fig. 3), for example, there is no localpenetrative fabric development or any deforma-tion that can be attributed to D3, yet the upperintercept age for the majority of analyses pre-sented here ranges between 2455–2483 Ma.

(2) The sampled orthogneiss on Tonagh Islandcontains a pervasive, generally flat-lying gneissos-ity, S1, that parallels a S1 gneissosity that isdefined by UHT assemblages in adjacent parag-neiss sequences. The orthogneiss also containsabundant (10–20% modal) mesoperthite. These

observations suggest that the orthogneiss was em-placed into the crust prior to UHT conditions,and subsequently endured the entire D1–D2 struc-tural evolution, including UHT metamorphism.Since orthogneiss emplacement, based on the datapresented here, occurred at 2626�28 Ma, UHTmetamorphism and D1 cannot be older than this.

(3) A more general observation from the nu-merous isotopic studies is that a number of zir-con U–Pb systems, which date the emplacementof pre-D1/UHT orthogneiss units, show little orno isotopic disturbance at the timing of UHTconditions favoured by either Black et al. (1983a)or Harley and Black (1997) (Black and James,1983, Fig. 3; Black et al., 1983b, 1986b, Fig. 6;Williams et al., 1984, Fig. 3; Fig. 4; Harleyand Black, 1997, Fig. 3a and b). These units


instead show discordia with (albeit imprecise)lower intercepts consistent with disturbance at�2400–2500 Ma (Fig. 10).

Based on our data and observations, UHTmetamorphism must have post-dated emplace-ment of the Tonagh Island orthogneiss at 2626�28 Ma. There are two periods of zircon growthsubsequent to 2626�28 Ma identified in thisstudy, one at 2546�13 Ma and again at 2476.9�9 Ma (taking the weighted mean of all analyses of�2450–2480 Ma as shown in Fig. 7); UHT meta-morphism must have occurred during one of theseepisodes. Given the intense isotopic disturbance ina number of isotopic systems across the entireNapier Complex at �2450–2480 Ma and theabsence of any significant D3 effects at the samplelocality (point 1), we propose that UHT metamor-phism and D1–D2, at least in the Tonagh Islandregion, occurred at �2450–2480 Ma, supportingthe assertion of Grew (1998). We acknowledgethat the growth of zircon at 2546�13 Ma mayindicate a metamorphic episode at this time. How-ever, we discount the possibility that UHT meta-morphism occurred at this time, given the paucityof data of �2550 Ma from presently availablechronological studies of the Napier Complex.

Our conclusion contrasts with the structurally-constrained chronological framework presentedby, for example, Harley and Black (1997) whichargues for UHT metamorphism at 2850–2820Ma. The timing of syn-D1 felsic intrusives, inparticular the Dallwitz Nunatak orthogneiss (�2850 Ma), is central to the calibration of therevised chronological framework of Harley andBlack (1997). In order to reconcile these differ-ences, one might argue for multiple episodes ofUHT metamorphism across the Napier Complex.An episode of UHT metamorphism that peaked inthe eastern periphery of the Tula Mountains at2850 Ma (Harley and Black, 1997) may have beenpost-dated by a younger episode of UHT meta-morphism at �2470 Ma in the Amundsen Bayregion, and which also exerted a strong isotopicoverprint across the terrane and cumulated in theregional development of upright D3 fold struc-tures at upper amphibolite grade. This suggestionwould imply that the D1–D2 structural featuresbetween the eastern Tula Mountains and Amund-

sen Bay are not time equivalents, although thispresents no particular geological difficulty.

Indeed, the possibility of high-grade polymeta-morphism in the Napier Complex has been previ-ously explored by Hensen and Motoyoshi (1992)and Harley and Black (1997). In one proposal,Hensen and Motoyoshi (1992) suggested that sep-arate overprinting high temperature events mayoffer a better explanation for petrological featuresassociated with D1 and D2 at Mt Riiser-Larsen(Fig. 1), considering that the time span originallyproposed for the D1–D2 evolution (�3072–2850Ma) is considered too long to sustain tempera-tures in excess of 900 °C (Harley and Hensen,1990). Harley and Black (1997) also address thepossibility of multiple tectonothermal events.Those authors propose that a high-grade fabricforming event at �2980�9 Ma, synchronouswith their revised timing of emplacement of theProclaimation Island orthogneiss, occurred innorthern regions of the terrane. This fabric form-ing event was superseded by D1–D2 and UHTmetamorphism in the Tula-Scott mountains re-gion, at 2850–2820 Ma, the time of UHT fa-voured by Harley and Black (1997). While thesediscussions do not directly relate to the suggestionoffered here, in that UHT metamorphism in theAmundsen Bay region at �2470 Ma post-datedan earlier episode UHT metamorphism in theTula-Scott mountains region at 2850–2820 Ma,they do indicate willingness of researchers to ad-dress the possibility of polymetamorphic events inorder to account for inconsistencies in the petro-graphic, chronological and structural frameworkfor the Napier Complex.

In spite of the abundance of isotopic data col-lected in the Napier Complex, there remains in-sufficient data to fully test and assess thepossibility of two UHT events, one at 2850–2820Ma and another at �2450–2480 Ma in theAmundsen Bay area. We believe that anotherdifficulty with this suggestion is that it fails toexplain point 3 above, in that some lithologieslocated in the eastern Tula Mountains, such as theTonalitic Gneiss at Mt Sones and the Gage Ridgeorthogneiss, show no clear evidence of isotopicdisturbance at the timing of UHT metamorphismfavoured by Harley and Black (1997). It would


seem unusual that UHT metamorphism at �2850 Ma (Harley and Black, 1997), an event ofextreme metamorphic conditions (�1100 °C)and intense pervasive deformation, across a con-siderable region, left an comparatively minorisotopic imprint and only in some rock types. Inaddition, one might argue that it seems equallyunusual that D3, a relatively lower grade event(�700 °C), with marked heterogenous struc-tural distribution, resulted in a widespread, re-markably consistent, strong isotopic imprint at�2450–2480 Ma in virtually all rock types andisotopic systems (e.g. Sandiford and Wilson,1984; Sheraton et al., 1987).

6.2. High Th/U zircons

An intriguing feature of the zircon data ofBlack et al. (1986b), and of this study, is thepresence of a near concordant zircon populationwith an unusually high Th/U (�1.2). Many zir-con analyses of the volumetrically dominantmantle and outer rims in this study have highTh/U (Table 4; Carson et al., 2002) which, ifregressed as a separate population, yield an up-per intercept age of 2483.0�9 Ma (n=52, Fig.8). This age is identical to that reported byBlack et al. (1986b) for a near concordant pop-ulation of high Th/U (�1.2) ‘structureless crys-tals’ which yielded a ‘preferred’ 207Pb/206Pbmean age of 2479�23 Ma. Although exhaustivestudies of the processes and conditions con-trolling zircon Th/U remain to be conducted, ingeneral, zircons grown under metamorphic con-ditions frequently have low Th/U, typically �0.05–0.1 (Williams and Claesson, 1987; Maas etal., 1992). Indeed, reported Th/U values greaterthan unity are quite uncommon, and the highTh/U zircon population presented here, and inBlack et al. (1986b), clearly indicates the forma-tion of zircon in an unusual chemical environ-ment. Black et al. (1986b) argued that thegrowth of zircon with such unusual chemistry isdue to crystallisation in an environment depletedin U relative to Th. Those authors also arguethat depletion must have occurred prior to crys-tallisation of this distinctive zircon population,during UHT metamorphism at 3070 Ma.

However, in a zircon U–Pb ion probe studyof the Scandinavian Caledonides, Williams andClaesson (1987) concluded that the growth ofzircon, with characteristically low Th/U, oc-curred during the transitional amphibolite–gran-ulite Caledonian orogeny, at which time Th andU were mobilised. They discounted the alterna-tive possibility that loss of Th and U from thehost rock occurred prior to the growth of thelow Th/U zircon. They based this on the obser-vation that a spectrum of pre-existing Precam-brian zircon grains, with ‘‘normal’’ Th/U(typically 0.1–0.5), have Caledonian age lowerintercepts, a period of metamorphism duringwhich new zircon growth occurred, characterisedby low Th/U (�0.1).

Similarly, following Williams and Claesson(1987), we propose that the zircon populationwith high Th/U described here characterises zir-con growth during UHT metamorphism. Al-though the example of Williams and Claesson(1987) specifically addresses the growth of lowTh/U zircon during Palaeozoic metamorphismin the Scandinavian Caledonides, the logic oftheir conclusion is relevant here. We suggestthat the actinides were mobilised during UHTmetamorphism, with preferential extraction of Uover Th from the terrane, providing a chemicalenvironment during peak conditions that facili-tated the growth of high Th/U zircon. Discor-dance trends in much older, �2470 Ma,pre-existing zircon populations reported in otherU–Pb studies from the Napier Complex fre-quently have (albeit imprecise) lower interceptsthat coincide with the growth of the high Th/Uzircon population at �2470 Ma (see point 3above). We reaffirm, therefore, that the upperintercept age of 2483.0�9 Ma for the high Th/U analyses (Fig. 8) may indicate the best esti-mate for the timing of UHT metamorphism.

The low U outer rims (Fig. 9) yield an upperintercept of 2455�16 Ma, which is significantlydistinct from the high Th/U zircon population(2483.0�9 Ma). Zircon from the low U popula-tion are restricted to the outer margins of zircongrains (e.g. Fig. 4b–g) and are generally charac-


terised by low U (and relatively low Th) con-tents. The low U outer rims may result fromzircon growth during continued U depletion ofthe terrane resulting from UHT metamorphism.Based on the textural context of the low U zircon(as rim material), we suggest that the low Uouter rim zircon represents the final period ofzircon growth that occurred during waning UHTmetamorphism.

6.3. Re�erse discordance

It should be noted that a number of the U–Pbanalyses shown in Fig. 7 are reversely discordant.Although discussion of the significance of theobserved reverse discordance for this datasetis presented elsewhere (Carson et al., 2002), abrief summary is offered here. Reverse discor-dance is considered to be either genuine, resultingfrom intragrain net excess of Pb* due to redistri-bution during a geological disturbance (e.g.Williams et al., 1984; Harley and Black, 1997)or an analytical artefact due to differential ioni-sation efficiencies between standard and un-known resulting in enhanced Pb ion yields(e.g. McLaren et al., 1994; Wiedenbeck, 1995).Carson et al. (submitted for publication) con-cluded that reverse discordance exhibited by asubset of the zircon population illustrated in Fig.7 is genuine, based on a non-zero lower intercept,as defined by the spread of both normaland reversely discordant data, that coincides witha known episode of geological disturbance(Section 6.4), and evidence of sub-micron scalePb and U heterogeneities. An analytical artefactrationale for the observed reverse discordancewas discounted, primarily on the absence ofa zero lower intercept; as a zero intercept isa characteristic feature of differing ionisationefficiencies between the standard and a set ofunknowns (e.g. McLaren et al., 1995; Wieden-beck, 1995). It should be noted, however,that regardless of the nature of reverse discor-dance, real or analytical artefact, the upper inter-cept age is still considered robust (Wiedenbeck,1995), and the geochronological informationderived from upper intercepts in this study re-main valid.

6.4. Significance of post-2460 Ma lower interceptages

The present study also permits some discussionof the significance of post-2460 Ma lower inter-cepts reported in many U–Pb studies of theNapier Complex. A number of imprecise U–Pblower intercepts at �1000 Ma have been docu-mented from the Field Islands, near theboundary with the NeoProterozoic Rayner Com-plex (Fig. 1). These lower intercepts have beenattributed to isotopic disturbance in response to�1000 Ma granulite-facies tectonothermal activ-ity in the adjacent Rayner Complex (Grew et al.,1982; Black et al., 1983a, 1986b; Sandiford andWilson, 1984). However, Shirashi et al. (1997)re-examined the chronological evolution of thewestern-most regions of the Rayner Complex andshowed that, immediately adjacent to the NapierComplex, the Rayner Complex experienced high-grade metamorphism during the Early Palaeozoic(�500–550 Ma). Shirashi et al. (1997) proposedthat the western extremities of the Rayner Com-plex instead represent the eastward extension ofthe Early Palaeozoic Lutzow–Holm Complexand that the western Rayner Complex may infact represent a collage of NeoProterozoic andEarly Palaeozoic terranes, including �770 Mafelsic intrusive bodies. This scenario clearly intro-duces some complexity in the interpretation andsignificance of the �1000 Ma lower interceptages reported from within the Napier Complex.

Although, imprecise due to extended extrapola-tions to concordia, the U–Pb zircon analysespresented here (Fig. 7) indicate an EarlyPalaeozoic lower intercept (see also Grew andManton, 1979). Carson et al. (2002) concludedthat, for the orthogneiss studied here, zircon dis-cordance was directly related to disturbance dueto aqueous fluid infiltration, at upper-amphibolitegrade (�8 kb and �670 °C; Carson et al., sub-mitted for publication), accompanying emplace-ment of an Early Palaeozoic pegmatite. Thisraises the possibility that post-2450 Ma concor-dia lower intercepts may result, at least in part,from isotopic disturbance due to aqueous fluidinfiltration into the Napier Complex during theEarly Palaeozoic, particularly for lithologies


where there is demonstrable hydration due toproximity to Early Palaeozoic pegmatites. Forexample, based on the description of the biotite-bearing ‘grey gneiss’ by Black et al. (1983a) fromAyatollah Island, Casey Bay, that unit is clearlyrelated to the proximity of a large pegmatite bodyof Early Palaeozoic age (Rb–Sr whole-rock age of522�10 Ma; Black et al., 1983a). Almost cer-tainly, the formation of this ‘grey gneiss’ resultedfrom hydration and recrystallisation of the anhy-drous ‘pink gneiss’ by influx of aqueous fluidsaccompanying pegmatite emplacement (Grew andManton, 1979; Carson et al., 2002). This conclu-sion is supported by a Rb–Sr whole-rockisochron from this ‘grey gneiss’ which yields asomewhat imprecise age of 650�145 Ma (Blacket al., 1983a). Although the lower intercept basedon conventional U–Pb zircon analyses from thisunit is 1183 +433/−446 Ma (Black et al.,1983a), a likely cause of zircon discordance is theinflux of magmatic fluids due to the emplacementof the nearby Early Palaeozoic pegmatite. Al-though, some well-defined lower intercepts areclearly related to the effects of tectonothermalactivity of Rayner age (Black et al., 1984), theinterpretation of other less precise lower inter-cepts and ‘geologically meaningless’ lower inter-cepts (Black et al., 1986a, Fig. 7b and c) acrossthe Napier Complex are problematic in this light,and should address the possibility of superim-posed effects of fluid influx into the Napier Com-plex, during the Early Palaeozoic.

7. Conclusions

The zircon U–Pb data presented from TonaghIsland provide additional constraints on the tim-ing of UHT metamorphism within the NapierComplex, East Antarctica, which has been a topicof debate for two decades. We conclude that:(a) the emplacement of a major felsic orthogneiss

body on Tonagh Island, occurred at 2626�28 Ma, prior to the onset of UHT metamor-phism and intense D1–D2 deformation.

(b) Two episodes of zircon growth occurred sub-sequent to orthogneiss emplacement. A minorepisode of growth occurred at 2546�13 Ma,

followed by a major episode of growth com-mencing with the high Th/U mantles at2483.0�9 Ma and terminating with outerrim development at 2455�16 Ma.

(c) UHT metamorphism occurred, at least in theTonagh Island region, synchronous with thegrowth of high Th/U zircon at 2483.0�9Ma, with zircon growth continuing at leastuntil 2455�16 Ma as represented by thegrowth of low-U outer rims. We find nocompelling evidence that supports UHTmetamorphism at either �2840 or �3070Ma.

(d) Post �2460 Ma zircon lower intercepts re-ported from the Napier Complex are com-monly considered to result from isotopicdisturbance during adjacent Rayner Complextectonothermal activity (�1000 Ma). Westress that such lower intercepts may also becomplicated by isotopic disturbance due toaqueous fluid influx, at elevated conditions,into the Napier Complex which accompaniedEarly Palaeozoic pegmatite emplacement at�500 Ma.

Acknowledgements

CJC acknowledges participation in the 40thJapanese Antarctic Research Expeditions and theSEAL (Structure and Evolution of East AntarcticLithosphere) program, at the kind invitation of K.Shiraishi and Y. Motoyoshi of the National Insti-tute of Polar Research, Toyko. CJC also wishes tothank the expeditioners of JARE-39 and -40, par-ticularly the Tonagh Island field party, and theofficers and crew of the icebreaker Shirase, for thehospitality extended to him during JARE-40. As-sistance by, and discussion with, T.M Harrison,M. Grove, K. McKeegan and C.M. Breedingduring ion probe sessions at the Department ofEarth and Space Sciences, UCLA, and with J.Eckert during electron microprobe sessions atYale University is appreciated. E.S. Grew pro-vided the photograph in Fig. 3 and comments onearly version of the manuscript. The manuscriptbenefited from valuable and constructive reviewsfrom Simon Harley and an anonymous reviewer.


CJC acknowledges receipt of the Damon Wellsfellowship at Yale University; analytical costswere supported by National Science Foundationgrant EAR-0001084 and -9810089.

References

Black, L.P., James, P.R., 1983. Geological history of theArchaean Napier Complex in Enderby Land. In: Oliver,R.L., James, P.R., Jago, J.B. (Eds.), Antarctic Earth Sci-ence. Cambridge University Press, Cambridge, pp. 11–15.

Black, L.P., James, P.R., Harley, S.L., 1983a. Geochronologyand geological evolution of metamorphic rocks in the FieldIslands area, East Antarctica. J. Metamorphic Geol. 1,277–303.

Black, L.P., James, P.R., Harley, S.L., 1983b. The geochronol-ogy, structure and metamorphism of early Archaean rocksat Fyfe Hills, Enderby Lands, Antarctica. PrecambrianRes. 21, 197–222.

Black, L.P., Fitzgerald, J.D., Harley, S.L., 1984. Pb isotopiccomposition, colour, and microstructure of monazites froma polymetamorphic rock in Antarctica. Contrib. MineralPetrol. 85, 141–148.

Black, L.P., Sheraton, J.W., James, P.R., 1986a. Late Ar-chaean granites of the Napier Complex, Enderby Land,Antarctica: a comparison of Rb–Sr, Sm–Nd and U–Pbisotopic systematics in a complex terrain. PrecambrianRes. 32, 343–368.

Black, L.P., Williams, I.S., Compston, W., 1986b. Four zirconages from one rock: the history of a 3930 Ma old granulitefrom Mt Sones, Enderby Land, Antarctica. Contrib. Min-eral Petrol. 94, 427–437.

Black, L.P., McCulloch, M.T., 1987. Evidence for isotopicre-equilibration of Sm–Nd whole-rock systems in EarlyArchaean crust of Enderby Land, Antarctica. Earth PlanetSci. Lett. 82, 15–24.

Black, L.P., Harley, S.L., Sun, S.S., McCulloch, M.T., 1987.The Rayner Complex of east Antarctica, Complex isotopicsystematics within a Proterozoic mobile belt. J. Metamor-phic Geol. 5, 1–26.

Black, L.P., 1988. Isotopic resetting of U–Pb zircon andRb–Sr and Sm–Nd whole-rock systems in Enderby Land,Antarctica: implications for the interpretation of isotopicdata from polymetamorphic and multiply deformed ter-rains. Precambrian Res. 38, 355–365.

Carson, C.J., Ague, J.J., Grove, M., Coath, C.D., Harrison,T.M., 2002. U-Pb isotopic behaviour of zircon duringupper-amphibolite facies fluid inflitration in the NapierComplex, east Antarctica. Earth and Plantary Science Let-ters, in press.

Dallwitz, W.B., 1968. Coexisting sapphirine and quartz ingranulites from Enderby Land, Antarctica. Nature 219,476–477.

DePaolo, D.J., Manton, W.I., Grew, E.S., Halpern, M., 1982.Sm–Nd, Rb–Sr and U–Th–Pb systematics of granulite

facies rocks from Fyfe Hills, Enderby Land, Antarctica.Nature 298, 614–618.

Ellis, D.J., 1980. Osumilite–sapphirine–quartz granulites fromEnderby Land, Antarctica: P–T conditions of metamor-phism, implications for garnet–cordierite equilibria andthe evolution of the deep crust. Contrib. Mineral Petrol.74, 201–210.

Ellis, D.J., Sheraton, J.W., England, R.N., Dallwitz, W.B.,1980. Osumilite–sapphirine–quartz granulites from En-derby Land, Antarctica—mineral assemblages and reac-tions. Contrib. Mineral Petrol. 72, 123–143.

Grew, E.S., 1980. Sapphirine+quartz association from Ar-chaean rocks in Enderby Land, Antarctica. Am. Mineralo-gist 65, 821–836.

Grew, E.S., 1982. Osumilite in the sapphirine–quartz terraneof Enderby Land, Antarctica: implications for osumilitepetrogenesis in the granulite-facies. Am. Mineralogist 67,762–787.

Grew, E.S., 1998. Boron and beryllium minerals in granulite-facies pegmatites and implications of beryllium pegmatitesfor the origin and evolution of the Archaean Napier Com-plex of East Antarctica. Memoirs of National Institute ofPolar Research, Special Issue, 53, 74–92.

Grew, E.S., Manton, W.I., 1979. Archaean rocks in Antarc-tica: 2.5-billion-year uranium–lead ages of pegmatites inEnderby Land. Science 206, 443–445.

Grew, E.S., Manton, W.I., Sandiford, M., 1982. Geochrono-logical studies in East Antarctica: age of pegmatites inCasey Bay, Enderby Land. Antarct. J. US 17, 1–2.

Harley, S.L., 1985a. Garnet–orthopyroxene bearing granulitesfrom Enderby Land, Antarctica: metamorphic pressure–temperature–time evolution of the Archaean Napier Com-plex. J. Petrol. 26, 819–856.

Harley, S.L., 1985b. Paragenetic and mineral-chemical rela-tionships in orthoamphibole-bearing gneisses from En-derby Land, east Antarctica: a record of Proterozoic uplift.J. Metamorphic Geol. 3, 179–200.

Harley, S.L., 1987. Pyroxene-bearing granulites from TonaghIsland, Enderby Land, Antarctica: further evidence forvery high temperatures (�950 °C) Archaean regionalmetamorphism in the Napier Complex. J. MetamorphicGeol. 5, 241–356.

Harley, S.L., 1998. An appraisal of peak temperatures andthermal histories in ultra high-temperature (UHT) crustalmetamorphism: the significance of aluminous orthopyrox-ene. Memoirs of National Institute of Polar Research,Special Issue, 53, 49–73.

Harley, S.L., Black, L.P. 1987. The Archaean geological evolu-tion of Enderby Land, Antarctica. In: Park, R.G., Tarney,J. (Eds.), Evolution of the Lewisian and Comparable Pre-cambrian High-grade Terrains. Geological Society SpecialPublication No. 27, pp. 285–296.

Harley, S.L., Hensen, B.J., 1990. Archaean and Proterozoichigh-grade terranes of East Antarctica (40°E–60°E): a casestudy of the diversity in granulite facies metamorphism. In:Ashworth, J.R., Brown, M. (Eds.), High TemperatureMetamorphism and Crustal Anatexis. Unwin Hyman,London, pp. 320–370.


Harley, S.L., Black, L.P., 1997. A revised Archaean chronol-ogy for the Napier Complex, Enderby Land, fromSHRIMP ion-microprobe studies. Antarct. Sci. 9, 74–91.

Harley, S.L., Motoyoshi, Y., 2000. Al zoning in orthopyrox-ene in a sapphirine quartzite: evidence for �1120 °CUHT metamorphism in the Napier Complex, Antarctica,and implications for the entropy of sapphirine. Contrib.Mineral Petrol. 138, 293–307.

Hensen, B.J., Motoyoshi, Y., 1992. Osumilite-producing re-actions in high temperature granulites from the NapierComplex, East Antarctica: tectonic implications. In:Yoshida, Y., Kaminuma, K., Shirashi, K. (Eds.), RecentProgress in Antarctica Earth Science. Terra SciencificPublishing Company, Tokyo, pp. 87–92.

Kelly, N.M., Clarke, G.L., Carson, C.J., White, R.W., 2000.Thrusting in the lower crust: evidence from the OygardenIslands, Kemp Land, East Antarctica. Geol. Mag. 137,219–234.

Lovering, J.F., 1979. The evidence for (4000 m.y. crustalmaterial in Archaean times. J. Geol. Soc. Aust. 26, 268abstract.

Ludwig, K.R., 1998. On the treatment of concordant ura-nium-lead ages. Geochim. Cosmochim. Acta 62, 315–318.

Ludwig, K.R., 1999. User’s manual for Isoplot/Ex, v2.3, ageochronological toolkit for Microsoft Excel. BerkeleyGeochronological Centre Special Publication No. 1a, pp.54.

McLaren, A.C., Fitzgerald, J.D., Williams, I.S., 1994. Themicrostructure of zircon and its influence on the age de-termination from Pb/U isotopic ratios measured by ionmicroprobe. Geochim. Cosmochim. Acta 58, 993–1005.

Maas, R., Kinny, P.D., Williams, I.S., Froude, D.O., Comp-ston, W., 1992. The earth’s oldest known crust: ageochronological and geochemical study of 3900–4200Ma old detrital zircons from Mt. Narryer and Jack Hills,Western Australia. Geochim. Cosmochim. Acta 56,1281–1300.

Miyamato, T., Grew, E.S., Sheraton, J.W., Yates, M.G.,Dunkley, D.J., Carson, C.J., Yoshimura, Y., Motoyoshi,Y., 2000. Lamproite dykes in the Napier Complex atTonagh Island, Enderby Land, East Antarctica. PolarGeosci. 13, 41–59.

Osanai, Y., Toyoshima, T., Owada, M., Tsunogae, T.,Hokada, T., Crowe, W.A., 1999. Geology of Ultrahigh-Temperature metamorphic rocks from Tonagh Island inthe Napier Complex, East Antarctica. Polar Geosci. 12,1–28.

Owada, M., Osanai, Y., Kagami, H., 1994. Isotopic equili-bration age of Sm–Nd whole-rock system in the NapierComplex (Tonagh Island), East Antarctica. Proc. NIPRSymp. Antarct. Geosci. 7, 122–132.

Paces, J.B., Miller, J.D., 1993. Precise U–Pb ages of DuluthComplex and related mafic intrusions, Northeastern Min-nesota: geochronological insights to physical, petroge-netic, paleomagnetic, and tectonomagnetic processesassociated with the 1.1 Ga midcontinental rift system. J.

Geophys. Res. 98, 13997–14013.Quidelleur, X., Grove, M., Lovera, O.M., Harrison, T.M.,

Yin, A., Ryerson, F.J., 1997. The thermal evolution andslip history of the Renbu Zedong Thrust, southeasternTibet. J. Geophys. Res. 102, 2659–2679.

Sandiford, M., 1985. The origin of retrograde shear zones inthe Napier Complex: implications for the tectonic evolu-tion of Enderby Land, Antarctica. J. Struct. Geol. 7,477–488.

Sandiford, M., Wilson, C.J.L., 1984. The structural evolu-tion of the Fyfe Hills-Khmara Bay region, EnderbyLand, East Antarctica. Aust. J. Earth Sci. 31, 403–426.

Sandiford, M., Powell, R., 1986a. Pyroxene exsolution ingranulites from Fyfe Hills, Enderby Land, Antarctica:evidence for 1000 °C metamorphic temperatures in Ar-chaean continental crust. Am. Mineralogist 71, 946–954.

Sandiford, M., Powell, R., 1986b. The origin of Archaeangneisses in the Fyfe Hills region, Enderby Land; fieldoccurrence, petrology and geochemistry. PrecambrianRes. 31, 37–68.

Sheraton, J.W., Black, L.P., 1981. Geochemistry andgeochronology of Proterozoic tholeiite dykes of EastAntarctica: evidence for mantle metasomatism. Contrib.Mineral Petrol. 78, 305–317.

Sheraton, J.W., Black, L.P., 1983. Geochemistry of Precam-brian gneisses: relevance for the evolution of the EastAntarctic Shield. Lithos 16, 273–296.

Sheraton, J.W., Tingey, R.J., Black, L.P., Offe, L.A., Ellis,D.J., 1987. Geology of Enderby Land and western KempLand, Antarctica. Bureau of Mineral Resources, Aus-tralia, Bulletin, no. 223, pp. 51.

Shirashi, K., Ellis, D.J., Fanning, C.M., Hiroi, Y., Kagama,H., Motoyoshi, Y., 1997. Re-examination of the meta-morphic and protolith ages of the Rayner Complex,Antarctic: evidence for the Cambrian (Pan-African) re-gional metamorphic event. In: Ricci, C.A. (Ed.), TheAntarctic Region: Geological Evolution and Processes.Terra Antarctica Publication, Siena, pp. 79–88.

Sobotovich, E.V., Kamenev, Y.N., Komaristyy, A.A., Rud-nik, V.A., 1976. The oldest rocks of Antarctic (EnderbyLand). Int. Geol. Rev. 18, 371–388.

Tainosho, Y., Kagami, H., Takahashi, Y., Iizumi, S., Os-anai, Y., Tsuchiya, N., 1994. Preliminary result for theSm–Nd whole-rock age of the metamorphic rocks fromMount Pardoe in the Napier Complex, East Antarctica.Proc. NIPR Symp. Antarct. Geosci. 7, 115–121.

Tainosho, Y., Kagami, H., Hamamoto, T., Takahashi, Y.,19947. Preliminary result for the Nd and Sr isotope char-acteristics of the Archaean gneisses from Mount Pardoe,Napier Complex, East Antarctica. Proc. NIPR Symp.Antarct. Geosci. 10, 92–101.

Wiedenbeck, M., 1995. An example of reverse discordanceduring ion microprobe zircon dating—an artifact of en-hanced ion yields from a radiogenic labile Pb. Chem.Geol. 125, 197–218.


Williams, I.S., Compston, W., Black, L.P., Ireland, T.R.,Foster, J.J., 1984. Unsupported radiogenic Pb in zircon: acause of anomalously high Pb–Pb, U–Pb and Th–Pbages. Contrib. Mineral Petrol. 88, 322–327.

Williams, I.S., Claesson, S., 1987. Isotopic evidence for thePrecambrian provenance and Caledonian metamorphismof high grade paragneisses from the Seve Nappes, Scandi-navian Caledonides. Contrib. Mineral Petrol. 97,205–217.

u–pb geochronology from tonagh island, east antarctica ... · island, enderby land, east...

Documents