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The relationship of Palaeozoic metamorphism and S-type magmatism on thepaleo-Pacific Gondwana margin
James Scott a,b,⁎, Janet Muhling c, Ian Fletcher d, Marco Billia a, J. Michael Palin a,Tim Elliot a, Christina Günter b
a Geology Department, University of Otago, Leith Street, PO Box 56, Dunedin, New Zealandb Insitut für Erd- und Umweltwissenschaften, Universität Potsdam, Karl-Leibknecht-Straβe 24, 14476 Potsdam, Germanyc Centre for Microscopy, Characterisation and Analysis, University of Western Australia, 35 Stirling Highway, Crawley WA 6009, Australiad Department of Applied Geology, Curtin University, GPO Box U1987, Perth, WA 6845, Western Australia
Article history:Received 25 April 2011Accepted 6 September 2011Available online 17 September 2011
Keywords:Geothermobarometric P–TMonazite growthPalaeozoic metamorphism New ZealandGondwana
A massive pulse of granitic magma was rapidly emplaced into the once contiguous West Antarctic and NewZealand segments of the palaeo-Pacific margin of the Gondwana supercontinent at ~371 Ma. In New Zealand,these Late Devonian S-type granitoids cover an areal extent of N 3400 km2, but the tectonic setting for crustalpartial melting has remained unclear because most of the exposure represents either emplacement-level, orrocks that have been reworked during Cretaceous orogenesis. New petrologic data indicate that aluminousparagneisses and orthogneisses in the Bonar Range represent a rare portion of Devonian middle crust thatpreserves evidence for the initiation of crustal melting. The investigated rocks outline the tail of a clockwiseP–T path that involved partial melting at peak conditions (~670 °C, 5.1 kb), deformation during the immedi-ately following near-isothermal decompression, and then partial re-equilibration under static conditions.Syn- to post-kinematic growth of zoned monazite establishes the timing of recrystallisation to a ~16 Maperiod that began at 373.4±4.1 Ma. This age overlaps with the initiation of regional Karamea S-type graniticmagmatism. Although estimated metamorphic conditions were insufficient for large amounts of melt to havebeen produced from Bonar Range pelites (calculated melt volumes are b10%), they do provide evidence con-sistent with widespread heating and partial melting in the deeper crust. This heating episode was contempo-raneous with partial melting in Fiordland (New Zealand) and West Antarctica, although Mesozoic thermaland deformational events complicate the Palaeozoic record in both those areas. Nevertheless, theapparent 1000 s km of along-strike crustal partial melting indicates that a continental-scale tectonic platemargin re-organisation took place at this time. The cause in the New Zealand segment was most likely, butnot unequivocally, an extensional tectonic regime with an elevated geothermal gradient caused by conduc-tive heating from a shallowed lithospheric mantle.
In terranes where S-type granites and migmatites are present, thecrustal P–T conditions must have been perturbed because typical geo-thermal gradients are seldom hot enough to melt continental crust(e.g. Brown, 2010; Clemens, 2003; Craven et al., 2011; Thompson,2001). End-member causes of melting commonly appealed to areadvective heating by emplacement of large volumes of hot plutonicrocks in continental arcs (e.g., Annen et al., 2006) or rapidly exhumedlower crust (e.g., O'Brien, 2000); near-isothermal decompression ofthe crust (e.g., Ring et al., 1999; Whitney et al., 2004); or conductiveheating caused by upwelling of asthenosphere during continentalextension (e.g., Collins and Richards, 2008).
A recent synthesis of Palaeozoic plutonism in New Zealand hasestablished that S-type granite magma was emplaced over ~3400 km2
of the Westland portion of the New Zealand Gondwana margin be-tween 371 and c. 340 Ma (Fig. 1) (Tulloch et al., 2009b). The timingand nature of crustal partial melting is strikingly similar to magmatisminWest Antarctica, whichwas contiguouswith theNew Zealandmicro-continent prior to Cretaceous separation (Bradshaw et al., 1997). InNew Zealand, coeval mafic rocks that could be indicators of heatingby basaltic intra- or under-plating are not abundant (Muir et al., 1996;Tulloch et al., 2009b) and neither Palaeozoic decompressional fabricsnor evidence for juxtaposition against hot orogenic crust has yet beenidentified. However, thesemechanisms cannot yet be ruled out becausethe exposedWestland geology is dominated by lowmetamorphic gradeturbidites and relatively undeformed upper crustal granitoids (Cooper,1989; Cooper and Tulloch, 1992), and the rare areas of exhumedmiddlecrust that have been investigated have mostly been reworked in theCretaceous (e.g., Hiess et al., 2010; Ireland and Gibson, 1998; Klepeis
Fig. 1. A. Location map of the area of study and nearby Karamea and Ridge Suite plutonic rocks. The Alpine Fault is the current Australia–Pacific plate boundary and truncates theWestern Province geology. Figure modified from Jongens (2006). B. Geological map of the Bonar Range, modified from Jongens (2006) to account for our own observations. Thetopographic rise from the Waitaha River to the centre of the Bonar Range is approximately 900 m.
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and Clarke, 2004; Sagar and Palin, 2011; Spell et al., 2000; Tulloch andKimbrough, 1989). Understanding the cause of New Zealand Palaeozoicmagmatism is important because this crust forms the central (and per-haps least well known) part to a 1000 s of km long orogen that oncestretched along the Antarctica–New Zealand–Australia Gondwana su-percontinent margin.
This paper reports an investigation of rocks of the little-visited BonarRange, rising out of the narrow coastal strip between the Southern Alpsmountain range (thrust up by the Australia–Pacific plate boundary)and the Tasman Sea on thewestern side of NewZealand (Fig. 1). Throughan integrated study of field, petrographic and geochronologic character-istics of aluminous paragneisses and orthogneisses throughout therange, the preserved fabrics are interpreted tohave formed in the Palaeo-zoic and been unaffected by Cretaceous deformation. These rocks thus
provide insights into the processes that were operating at an importanttime in the geological history of the palaeo-Pacific Gondwana margin.
2. Geological setting
The Bonar Range occurswithin theWestland portion of theWesternProvince of New Zealand (Fig. 1A). The Western Province comprisesLate Cambrian–Devonian meta-sedimentary and meta-igneous rocksof the elongate and parallel Takaka and Buller terranes (Cooper, 1989;Laird, 1972; Münker and Cooper, 1999). The eastern Takaka Terranecontains a Cambrian arc and Early Ordovician passive margin sequence(Cooper, 1989; Münker and Cooper, 1999). The western Buller Terrane—the focus of this paper—is dominated bymetamorphosed turbidites ofthe Ordovician Greenland Group (Cooper, 1975; Gage, 1948; Laird,
Fig. 2. A representative field image of paragneiss in the Bonar Range. The pelitic layerstypically contain sheared felsic leucosomes with a top-to-the-NE sense of shear, but thepsammitic layers show no macroscopic evidence for having been partially melted. Thehammer handle is approximately 1 m long.
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1972; Laird and Shelley, 1974). Both terranes have similar distinctivedetrital mineral signatures indicative of derivation from Grenville(~1100 Ma) and Ross-Delamerian (~600–550 Ma) orogenic phases(Adams and Kelley, 1998; Hiess et al., 2010; Ireland, 1992; Jongens etal., 2003; Nebel-Jacobson et al., 2010), and were deposited on thepaleo-Pacific margin of Gondwana (Adams and Kelley, 1998; Cooper,1989; Gibson and Ireland, 1996).
The Greenland Group is dominated by a monotonous series ofquartzofeldspathic argillite and greywacke turbiditic beds that haveundergone low-grade greenschist facies metamorphism during up-right folding (Laird, 1972; Laird and Shelley, 1974; Shelley, 1975).This unit has attracted attention for over 150 years because it con-tains economic shear-hosted gold mineralisation (e.g. Bierlen et al.,2004; Gage, 1948). However, the timing of deformation is poorlyknown due to the paucity of high-quality radiometric dating andsome partial resetting by Cretaceous events. Limited K–Ar and Rb–Sr whole rock age data suggest that deformation probably occurredin the Palaeozoic (Adams, 2004; Adams et al., 1975) but the accuracyof the results is compromised by the presence of radiogenic Ar in de-trital grains and incomplete homogenisation during metamorphism(Adams and Kelley, 1998). The c. 368 Ma Barrytown Granite, whichintrudes the Greenland Group (Tulloch et al., 2009b), provides theonly firm lower age bracket to deformation.
A variety of Palaeozoic plutonic suites have intruded the GreenlandGroup. The main suite is the S-type granite Karamea Suite (371–368 Ma), followed by the smaller volume I-type Paringa Suite (369–362 Ma), S-type Ridge Suite (354–340Ma), and I-type Tobin Suite(350–342 Ma) (Tulloch et al., 2009b). The two S-type granite suites con-tain abundant inherited zircon with age spectra similar to that of theGreenland Group, as well as evolved Sr, Nd, Pb and O isotope signatures(Tulloch et al., 2009b). These properties indicate the source of theS-type granitoids was, at least in part, melted continental crust (Tullochet al., 2009b).
Westland has been variably affected by magmatism and tectonismrelated to the Late Cretaceous break-up of New Zealand and Australia(Laird and Bradshaw, 2004; Spell et al., 2000; Tulloch and Kimbrough,1989). Early Cretaceous granites were emplaced between ~125 and100 Maduring the transition from convergent to divergent tectonics im-mediately prior to opening of the Tasman Sea (Sagar and Palin, 2011;Tulloch et al., 2009a;Waight et al., 1998). Evidence for this period of de-formation is mainly preserved in scattered extensional mylonitic shearzones that juxtapose low-grade Greenland Group and undeformedmid-Palaeozoic and Cretaceous plutons against gneisses and deformedmid-Palaeozoic and Cretaceous plutons (Spell et al., 2000; Tulloch andKimbrough, 1989). These shear zones have been interpreted to repre-sent the deeper levels of extensional detachment faults to Cretaceousmetamorphic core complexes (Klepeis et al., 2007; Spell et al., 2000;Tulloch and Kimbrough, 1989). Hiess et al. (2010) have demonstratedby zircon and monazite geochronology that there are areas of gneissicrock in Westland that record Palaeozoic±Cretaceous metamorphicevents. As the metamorphic conditions of these two metamorphicevents are broadly similar (Ireland and Gibson, 1998; Spell et al.,2000; White, 1994), it is extremely difficult to determine the ages ofspecific fabrics without the aid of geochronology.
3. Geology of the Bonar Range
The geology of the Bonar Range is summarised in Fig. 1B. It com-prises inter-layered psammitic and pelitic paragneiss (originally grey-wacke and argillite) and abundant aluminous porphyroclasticorthogneiss (Jongens, 2006). Exposure is generally restricted to creekbeds and landslides due to dense vegetation cover. Pelitic layers containdeformed leucosomes, whereas interleaved psammitic layers tend tolack macroscopic evidence for leucosomes (Fig. 2). Boudinaged and ro-tated porphyroclastic textures in the gneisses consistently show a dex-tral top-to-the-northeast sense of shear. Undeformed lamprophyric and
basaltic dikes of up to several metres inwidth crosscut gneissic foliationwithin the orthogneisses and paragneisses. These dikes are tentativelyinferred to be Cretaceous or younger due to their similarity to otherlamprophyric dike swarms in Westland (e.g., Tulloch and Kimbrough,1989), although there are as yet no radiometric ages to confirm thisinterpretation. No other mafic lithologies were observed in place or inthe float.
On the basis of fission track data and similarities to the nearby Creta-ceous Paparoa metamorphic core complex, the Bonar Range shear fab-rics were tentatively interpreted to have formed in the footwall of anEarly Cretaceous detachment fault (Jongens, 2006). However, a recentapatite and zircon fission track study of samples collected along theNE-trending stretching lineation (Ring and Bernet, 2010), coupledwith older fission track analyses from White and Green (1986) andKamp et al. (1992), found that zircon fission track ages progressivelydecrease in age from ~125 to 75Mawith increasing elevation. Althoughthese ages overlap with a period of regional extension (110–90 Ma;Gibson and Ireland, 1995; Spell et al., 2000), Ring and Bernet (2010)interpreted the rocks to represent an enclave that cooled slowly duringthe Cretaceous because the calculated cooling rates are an order ofmag-nitude less than commonly expected in extensional detachment faultsystems. A simple assessment of their data using 2 standard deviationerrors indicates the minimum and maximum rates of vertical exhuma-tion were somewhere between ~15 and ~220 mMa−1. The geody-namic significance of the gneissic–mylonitic fabric within the BonarRange therefore remains unclear.
4. Sample descriptions
4.1. Paragneiss
Three representative paragneiss samples are described (P67589, OU79917 and OU 79886). P67589 and OU 79917 contain foliation-parallel
felsic leucosomes comprising alkali feldspar and quartz. OU 79886 is abiotite- and sillimanite-rich gneiss without any leucosomes. Becauseof the complex history, minerals are designated garnet1, garnet2, bio-tite1, biotite2, etc. according to their paragenesis. Representative min-eral analyses of the metamorphic minerals obtained by wavelengthdispersive spectroscopy on an electron microprobe are discussedbelow and the data presented in Data Supplementary 1. Whole rock
Fig. 3. Representative photomicrographs ofmineral assemblages in pelitic gneisses. A. Planepolbiotite1, quartz1 and fine sillimanite1 prisms (see also Fig. 4). Sample P67589. B. Backscatteredbeen overgrown statically by quartz2 and prisms of sillimanite2. Sample P67589. C. Backscatt(plagioclase2, quartz2 and biotite2). Sample P67589. D. A backscattered electron image an alkaliscatter images of cordierite1 and coarse biotite1 partially replaced by sillimanite2 and quartz2
compositions were measured by X-ray fluorescence and are presentedin Data Supplementary 2.
Samples P67589 and OU 79917 are composed of foliation-parallelelongate garnet grains within a matrix of biotite1, alkali feldspar1,quartz1 and plagioclase1 (Fig. 3A,B), with accessory muscovite, zircon,monazite, ilmenite and apatite. The cores of the garnet grains inP67589 are almandine-rich but show a decrease in Mg and Ca and
arised light petrographic image of a porphyroblastic garnet grain. Garnet contains elongateelectron image of themargin of a garnet grain in sample OU 79917. Thematrix biotite1 hasered electron image of an alkali feldspar1 grain partially pseudomorphed by myrmekitefeldspar1 inclusion in garnet being replaced bybiotite2–quartz2 and plagioclase2. E. Back-. Sample OU 79886. Garnet and cordierite were never identified in the same rock.
Fig. 4. Representativewavelength dispersive elementmaps of porphyroblastic garnet grains. Fe (A) and Ca (B) elementmaps are from the garnet grain shown in Fig. 3A (P67589). Ca hasdiffused out along a crack in the grain, and at grainmargins. Fe increases towards themargins. Elementmaps ofMg (C) andMn (D) are from themargins of a porphyroblastic garnet in OU79917. Both Mg and Mn increase towards rims and cracks.
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sympathetic increase of Fe and Mn toward the rims (e.g., core–rim:Alm67–74Prp25–18Grs5–4Sps2–3; XMg27–20) (Fig. 4). Garnet grains in OU79917 show a similar rim-ward increase in Fe and Mn and a decreasein Mg (core–rim=Alm70–73Prp22–17Sps5–7Grs3; XMg24–18) (Fig. 4). Gar-net in both samples contains elongate inclusions of alkali feldspar1, bio-tite1, sillimanite1 and quartz1 (Figs. 3A, 4). These inclusions are matrix-parallel, and there is no evidence for garnet rotation. Alkali feldspargrains within the leucosomes contain plagioclase exsolution lamellaetoo fine to be analysed; the bulk composition is Or85–89Ab11–15An0. Bio-tite1 has a Ti content of 0.12–0.19 atoms per 12 O and is Mg-rich(XMg0.65–0.58). Plagioclase1 shows a subtle compositional difference be-tween samples: P67589 contains andesine (Ab60–56An43–39Or1), where-as plagioclase1 in OU 79917 is slightly more sodic (e.g., Ab68An32Or1).This is also reflected in thewhole rock composition, with OU79917 con-taining more SiO2 and slightly more Na2O than P67589, which in turncontains more Al2O3, Fe2O3(t), MgO and K2O than OU 79917. Monaziteis dispersed widely throughout P67589, occurring within garnet and inthe matrix. Monazite morphologies vary from round to elongate, andmany grains have grown parallel to grain boundaries and along cleavageplanes within biotite.
The garnet1–biotite1–alkali feldspar1–plagioclase1–sillimanite1–quartz1 assemblages have been partially statically overgrown. Coarsealkali feldspar1 within the matrix has been partially to totally replacedby myrmekite (plagioclase2+quartz2)±fine-grained biotite2 (Fig. 3C).Alkali feldspar1 inclusions in garnet have also been partially replaced bymyrmekite±fine-grained biotite2 where the garnet has fractured(Fig. 3D). Compared with matrix biotite2, biotite2 within garnet has anenrichedMg component (XMg66) and virtually no Ti. Symplectite plagio-clase2 is more Na-rich (Ab68–66An33–31Or1 in P67589) than plagioclase1.Plagioclase1 within the matrix locally exhibits an up to 30 μm widesodic-rich rim (Ab83–84An16–15Or1–0). Coarse foliation-defining biotite1grains have been overgrown by quartz2 and slender needles of fibroliticsillimanite2. Many of these prisms have developed along cleavage planesbut are not necessarily oriented parallel to the lineation within the rock.
OU 79886 is a biotite–cordierite–sillimanite–quartz-bearing peliticgneiss. The main minerals are coarse biotite1, cordierite1 and quartz1,and together these define the prominent penetrative fabric. Accessoryminerals are ilmenite and zircon. Biotite1 has a lower XMg (0.51–0.54) and slightly less wt.% K2O than biotite1 within the garnet-bearinggneisses. Cordierite1 has XMg=0.67–0.68 and 1.22–0.57 wt.% Na2O.Cordierite1 and biotite1 have been partially replaced and overgrownby numerous fine prisms and aggregates of fibrolite and quartz(Fig. 3E). The biotite- and sillimanite-rich composition is reflected inthe whole rock analysis, which shows it to be poor in SiO2 but rich inAl2O3 and Fe2O3(t), and to have no CaO and very little Na2O. This isalso consistent with the absence of feldspar.
4.2. Orthogneisses
The orthogneisses throughout the Bonar Range are granitic to tona-litic in composition, and commonly highly aluminous (Jongens, 2006).Representative samples OU 79915 and OU 79956 contain porphyroclas-tic plagioclase grains up to ~7 mm in diameter, set amongst red–brownbiotite, muscovite and fine-grained plagioclase and quartz. In addition,OU 79915 contains alkali feldspar porphyroclasts up to 1 cm diameter.The porphyroclasts sometimes contain coarse inclusions of biotite andquartz that are randomly oriented and may therefore be relict igneousphases. The porphyroclasts are rotated winged varieties that havebeen flattened and are slightly elongate parallel to the matrix biotitec-axes and to polycrystalline quartz bands. Together these fabrics definea gneissic–mylonitic fabric. Biotite is commonly partially replaced bymuscovite, and some muscovite grains enclose tiny sillimanite prisms.Other accessory phases are apatite, epidote, zircon, monazite and opa-que oxide. Myrmekite partially replaces earlier formed alkali feldsparin OU 79956 (e.g., Fig. 3C).
5. Metamorphism
5.1. Metamorphic assemblages and P–T path
The garnet1–alkali feldspar1–biotite1–plagioclase1–quartz1–sillimanite1 assemblages in pelites OU 79917 and P67589 define apenetrative fabric characteristic of formation at high-T and low-P dur-ing partial melting. Leucosomes are deformed by this fabric, but theirpresence indicates a prior peak metamorphic T in excess of ~650 °C inthe systemNaO–K2O–FeO–MgO–Al2O3–SiO2–H2O (NKFMASH). The ab-sence of orthopyroxene provides an approximate maximum T of~750 °C (Spear et al., 1999). The garnet–biotite (GB) geothermometersolved iteratively with the garnet–aluminosilicate–silica–plagioclase(GASP) geobarometer of Holdaway (2000, 2001), using the averaged ac-tivity models for garnet, biotite and plagioclase suggested by Holdaway(2000, 2001), returned results of 670±25 °C and 5.1±0.8 kb forP67589. The assigned errors are the values endemic to the particulargeothermometer and geobarometer because the intra-sample scatterwas less than the precision. The P–T estimates, coupled with the pres-ence of partialmelt textures, plot within suitable space for the observedmineral assemblages in the NKFMASH petrogenetic grid for aluminouspelites (Fig. 5A). Geothermobarometric estimates for OU 79917returned slightly lower values of 647±25 °C and 3.2±0.8 kb, asexpected from the lower XMg and Ca in garnet1 and higher Na inplagioclase1.
There are caveats with the accuracy of the P–T estimates. For exam-ple, the Holdaway (2000, 2001) GB-GASP method assumes that ferric
Fig. 5. A. P–T diagram for the calculated GB-GASP-derived geothermobarometric estimates. The reactions shown in pale grey are from NKFMASH for an average alumina-rich pelite(Spear et al., 1999). B. Calculated partial melt generated for P67589, with P–T estimates.
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iron comprises 11.3% of total FeOmeasured in biotite and 3% of total FeOmeasured in garnet. This may not always be valid and any departurewill affect T calculations. Nevertheless, a degree of confidence is placedin the results because the Ti-in-biotite geothermometer, which can beused as an independent test for garnet–biotite disequilibrium if the ap-propriate buffering minerals are present (Henry et al., 2005), yieldedaveraged results of 696±24 °C for P67589 and 677±24 °C for OU79917, both of which overlap with the GB-GASP-derived temperatures.The errors quoted are those endemic to the geothermometer. Biotite1 inthe cordierite-bearing gneiss, OU 79886, yielded a similar Ti-in-biotitetemperature of 660±24 °C.
The penetrative assemblages have been partially replaced under ret-rograde conditions. Partial re-equilibration is recorded by the decreas-ing XMg and increasing Mn at the rims and cracks within garnet(Fig. 4), the formation of myrmekite (plagioclase2, quartz2±biotite2)after alkali feldspar1 (Fig. 3C), and the replacement of biotite1 by silli-manite2–quartz2 aggregates (Fig. 3D). These textures are explainedby the following open-system reactions:
Coupled reactions (1) and (2) predict that garnet, alkali feldspar andhydrous fluid will be consumed to produce sillimanite2, biotite2, andquartz2. Biotite2 has a higher XMg than biotite1 because Fe has back-diffused into garnet and Mg was taken up by the newly grown biotite2.The compositional difference between plagioclase1 and plagioclase2 ismost pronounced where myrmekite has formed on alkali feldspar1 in-clusions in garnet. In this textural setting, infiltration of a hydrous
fluid infiltrated along cracks but the myrmekite composition wascontrolled by local domainal equilibrium; that is, the low XAn ofplagioclase2 in myrmekite within garnet may be explained by themain Ca-reservoir being the adjacent low-Ca garnet, which suggeststhat Ca net transfer should have locally taken place between thesetwo phases in this textural position. Similarly, the slightly higher NaandMgof biotite2within garnet compared to biotite2within thematrixis also probably a function of domainal equilibrium within garnet andrestricted element diffusion from the matrix. Biotite2 within garnetcontains no significant Ti because garnet is not a suitable titanianphase, unlike matrix biotite2, which is comparatively Ti-rich because itis probably in equilibriumwith ilmenite. The complete replacement of al-kali feldspar1 by myrmekite within OU 79917 is an indicator of furtherprogress of reaction (1) compared to P67589 because reaction (3) re-quires a fluid phase as a catalyst in order to relocate cations within therock matrix but does not consume H2O. Therefore, if reaction (1) stopsdue the exhaustion of fluid, reaction (3) will effectively cease also. Reac-tion (2)will stopwhen the garnet–biotite Fe–Mgexchange system closes.
As the above discussion indicates garnet2, biotite2, plagioclase2, sil-limanite2 and quartz2 co-crystallised in an open-system, the P–T condi-tions for the strain-free assemblages in P67589 and OU 79917 can becalculated using standard geothermobarometric methods. The GB-GASP method for second generation phases yielded results of 580±25 °C and 2.6±0.8 kb for P67589, and an overlapping 602±25 °C and3.4±0.8 kb for OU 79917. Ti-in-biotite geothermometry for matrix bio-tite2 yields temperatures of 603±24 °C (P67589) and 639±24 °C (OU79917) (Fig. 5A). Due to sluggish reaction kinetics and only partiallyhydrous conditions during retrogression, net-transfer reactions in thisrock appear to have been frozen somewhere between 500 °C and625 °C along the cooling–exhumation path. In summary, the GB-GASPestimates and mineral assemblages for the peak and retrograde assem-blages outline a portion of a clockwise P–T path that began with near-isothermal exhumation followed by near isobaric cooling (Fig. 5A).AFMdiagrams (not shown) for the inferred P–T path suggest that cordi-erite in OU 79886 formed during exhumation andwas not a peakmeta-morphic mineral.
5.2. Partial melting
The inferred P–T path outlined above and shown on Fig. 5 indicatesthat partial melt formed at peak conditions, although the degree ofmelting did not exceed several volume percent because muscovite1was not completely consumed during dehydration. This inference is
supported by thermodynamic modelling using Perple_X 07 (Connolly,2009; Connolly and Kerrick, 1987) with the '04 revision of the Hollandand Powell (1998) thermodynamic database and additional solutionmodels of Mica(CHA) (Auzanneau et al., 2010; Coggon and Holland,2002), Bio(HP) (Powell andHolland, 1999), Gt(HP) (Holland and Powell,1998) and hCrd and melt(HP) (Holland and Powell, 2001; White et al.,2001). At peak T (670±25 °C) in P67589, partial melting producedb10 vol.% melt (Fig. 5B). Fig. 5B also shows that the melt volume likelyremained constant during the very early decompression because theP–T path is sub-parallel to the isopleths for melt abundance. However,this diagram also indicates that by the time the static assemblages haddeveloped, melt was no longer being produced at all, which is consistentwith petrographic observations.
6. Isotope geochronology
6.1. Zircon
Analyses were obtained from 75 zircon grains extracted from peliticparagneiss sample OU79917 (data and systematics are reported in Data
Fig. 6. A, B and C. Zircon age spectra and cathodoluminescence images. Populations on the procence images refer to the analyses in Data Supplementary 3. Standard errors are shown as 2σ.
Supplementary 3). Of these, 14 analysesweremore than 10%discordantor had internal MSWD greater than 10. The remaining analyses yieldeda spread of dates ranging from 317 to 1877 Ma. There are resolvablepopulations at: 485, 572, 672, 760 and 963 Ma (Fig. 6A). The threeyoungest analyses yielded dates of 317.3±18.7 Ma, 339.8±19.2 and355.1±21.2 Ma, and have distinctively low Th/U (0.0–0.04). The detri-tal zircon peaks are very similar to those from the Greenland Group(Ireland, 1992), and the youngest detrital peak of 485 Ma is compatiblewith the Ordovician sedimentation age of the Greenland Group(Cooper, 1975).
Twenty-two analyses were obtained from 20 zircon grains inorthogneiss OU 79915 (Data Supplementary 3). Virtually all grainshave narrow dully luminescent overgrowths visible under cathodolu-minescence that truncate the internal morphologies (Fig. 6B). The zir-con spectrum contains distinct populations at 505, 598, and 686 Ma.With three exceptions (analyses 7, 10 and 18) most overgrowths weretoo narrow to be targeted with the laser beam, although the laser suc-cessfully bored from the cores to the rims of two grains (analyses 15and 17). These youngest 5 analyses have ages of 436 and 330 Ma anddistinctly low Th/U (0.18–0.00). With exclusion of the 436 Ma analysis,
bability density diagrams are discussed in the text. The numbers on the cathodolumines-
the remaining four analyses yield a pooled age of 342.1±17.8 Ma(n=4, MSWD=2.8).
Twenty-two analyses of 20 zircon grainswere conducted on orthog-neiss OU 79956, but 8 were too discordant or had MSWD values toohigh to be used in the age calculations. The ages of acceptable datesrange from 336 to 1314Ma (Data Supplementary 3). There are distinctpopulations at 493, 686 and 964 Ma (Fig. 6C). During analysis, the laserbored through the cores and into the rims in two grains, resolving agesof 336.1±20.8 Ma and 397.5±25.4 Ma. These two analyses, and onesuccessfully targeted rim (397.7±22.4 Ma), have low Th/U (0.12–0.02).
The two investigated porphyroclastic orthogneisses contain inher-ited zircon cores of similar age to the detrital zircon in the paragneiss(and Greenland Group; Ireland and Gibson, 1998; Hiess et al., 2010),as well as overgrowths of Late Devonian–Early Carboniferous age. Ifthe outermost overgrowths in these orthogneisses are attributed tometamorphism, then the age of emplacement is not immediatelyclear. Nevertheless, an age of intrusion can still be estimated becausethe laser drilled from an inherited core through the overgrowth andinto the epoxy in several grains. In these cases, these rims should becomposed of first a magmatic overgrowth and then a metamorphicovergrowth. As there are no resolvable internal age differences withinthe rims, the intrusion age is tentatively concluded to be within errorof the timing of metamorphism, and the high degree of inheritancedue to a lowmagmatic temperature and low solubility of Zr in the felsicmelt (e.g., Scott et al., 2011). Hiess et al. (2010) obtained a Devonian ageof 386.5±7.5 Ma from granodiorite on the southern side of the BonarRange. Although these authors suggest that this weakly deformedgranodiorite intrudes gneissic rocks, gneissic to mylonitic textures inthe northern portion of the dated granodiorite suggest the contrastingdeformation styles are the result of strain partitioning.
6.2. Monazite
Monazite grainswere extracted from a polished thin section of sam-ple P67589. The monazite element maps for Y show complex zoningpatterns (Fig. 7). The cores of the grains have zones of high Y (Y2O3
~2.1 wt.%) with irregular to cuspate boundaries, generally surroundedby zones with low Y (Y2O3 ~0.4) (Fig. 7A). Most grains have completeor partial rims with Y2O3 ~1.2%. Nineteen individual grains weredated by SHRIMP and 23 analyses obtained (Data Supplementary 4).The method of resolving the age populations is discussed in detail inData Supplementary 1, but is outlined briefly as follows. The analyseswere divided into UN~4500 ppm cores, Ub~3000 ppm mantles. TheU-rich cores occur in the centres of grains (Fig. 7A) and also as inclusionsin garnet. The 5 U-rich cores from grains in the sample matrix yieldedidentical pooled Th–Pb and U–Pb ages of 373.9±6.1 Ma (MSWDb1.0).Since the Pb–Pb and U–Pb ages were independently calibrated, theyare combined to yield an age of 373.5±4.3 Ma. If the two analysesfromgarnet inclusions are included in this population, the age is a slight-ly younger (but analytically indistinguishable) 373.4±4.1 Ma. Three an-alyses with intermediate U contents represent mixed populations.Eleven U-poor zones correspond to areas outside of the U-rich cores(Fig. 7B), and have Pb–U and Pb–Th ages 357.5±5.2 Ma(MSWD=1.1) and 359.5±5.4 Ma (MSWD=1.0), respectively. Thepooled age of this population is 358.5±3.7 Ma. The outermost rimswere too narrow to be successfully targeted (e.g., Fig. 7C) but there is
Fig. 7. A and B. Y-distribution and backscattered electron images of monazite illustratinginternal zoning and textural setting of dated specimens. The brighter the colour, thehigherthe Y content. The labelled circles correspond to analytical sites reported in Data Supple-mentary 4. C. Backscattered electron image of a monazite grain with a narrow moder-ate-Y overgrowth (element map not shown) extending along biotite1 cleavage. D.Monazite age versus analytical location diagram. The key feature is that ages for analysesthat contain both low/intermediate-U zones and overgrowths (indicated by a box aroundthat data point) do not show a mix of a ~355 Ma age and distinctly young age. Therefore,the age of the overgrowths must be very close to ~355 Ma.
Fig. 8. Schematic P–T path for the shallow crustal Greenland Group and middle crustalBonar Range. Greenland Group conditions are from Shelley (1975), Bonar Range pathis constructed from this current study, and Fiordland metamorphic conditions are fromIreland and Gibson (1998) and Scott et al. (2009a,b).
530 J. Scott et al. / Lithos 127 (2011) 522–534
no indication of a significantly younger age where the beam overlappeda rim (Fig. 7D).
As the closure temperature for Pb-diffusion in monazite is high(≥700 °C) and so generally does not re-equilibrate during slow cooling(Cherniak et al., 2004), the textural relations between monazite andbiotite1 points to most of the monazite having grown at the sametime as biotite1. However, some post-biotite1 monazite growth is indi-cated by the presence of overgrowths extending along biotite cleavageplanes. Y zoning in monazite usually reflects dissolution and growthof monazite and other Y-bearing phases, xenotime and garnet beingmost common (e.g., Spear and Pyle, 2002). Xenotime occurs only asrare inclusions in garnet and was not found in thematrix. As the garnetrecords a multistage history, the monazite domains likely reflect devel-opment related to garnet growth and breakdown. Monazite composi-tions are therefore interpreted to reflect initial growth and uptake of Yresulting in Y-rich cores, followed by simultaneous growth with garnetthat favourably partitioned Y and resulting in Y-poor zones, and finallymonazite growth during garnet breakdown that resulted in over-growths with moderate Y concentrations. The 373.4±4.1 Ma age istherefore interpreted to date the initiation of exhumation and meltingof the Bonar Range (i.e., garnet1 growth), whereas the 358.5±3.7 Maovergrowths and thin rims are interpreted to record subsequent staticre-equilibration. The monazite ages overlap with the age of most mea-sured overgrowths on the zircon grains.
7. Discussion
The currentWestland erosion level exposes several areas of gneissicrock amongst a region that is dominated by the low-grade OrdovicianGreenland Group. However, unlike the Bonar Range, geochronologicaldata have conclusively shown that most of these other gneissic areasunderwent recrystallisation in the Cretaceous during either regionalmetamorphism or shearing in the footwalls of Early Cretaceous faults(Hiess et al., 2010; Ireland and Gibson, 1998; Sagar and Palin, 2011;Spell et al., 2000; Tulloch and Kimbrough, 1989).Westlandmust, there-fore, comprise a collage of different crustal levels juxtaposed during orafter the Early Cretaceous, with the Bonar Range representing a relictPalaeozoic fragment free of subsequent tectonism. This is an importantconclusion because although the Bonar Range lies within, or very closeto, the exhumed portion of a Cretaceous arc that formed on the Gond-wana margin, the metamorphic, U–Pb geochronology and fission trackdata do not show evidence for these rocks having experienced orrecorded Cretaceous crustal metamorphism or core complex-style ex-tensional exhumation that is widely observed elsewhere in the region(Ireland and Gibson, 1998; Spell et al., 2000; Tulloch and Kimbrough,1989). The Bonar Range therefore provides insights into the evolutionof Palaeozoic processes that have elsewhere been obscured.
The P–T data for the Bonar Range define a clockwise exhumationpath (e.g., Fig. 5A). The earlier stages of the P–T path remain poorlydefined because prograde garnet growth zoning is not preserved. Thisobservation points to an extended peak metamorphic event that homo-genised prograde garnet growth zoning. As new zircon age data showthat the Bonar Range paragneiss protolith is likely to have been theOrdovician Greenland Group, then there must have been Early to mid-Palaeozoic burial of this portion of the Westland Gondwana margincrust to at least ~19 km (P=5.1 kb; 1 kb representing 3.7 km depth).This result means that the Bonar Range paragneisses departed fromthe typical regional Greenland Group P–T trajectory, which reached anapproximate peak temperature of 400 °C during imposition of themac-roscopic folds (Shelley, 1975) (Fig. 8).
7.2. Timing of partial melting
The integration of monazite composition and age data with meta-morphic assemblages provides key insights into the evolution of thiscomplex crustal segment of the Palaeozoic Gondwana margin. Firstly,
the new data show that the initiation of crustal partial melting, at leastin the Bonar Range, occurred at 373.4±4.1 Ma. This date overlapswith the initiation of generation and emplacement of the large volume(N3400 km3) S-type Karamea Suite (371–368 Ma) inWestland (Tullochet al., 2009b). Although a crustal component to the Karamea Suite hasbeen identified previously (Muir et al., 1996; Tulloch et al., 2009b), theBonarRange paragneisses provide direct petrological and geochronolog-ical support for that theory. Nevertheless, the peak T estimates are toolow (~670 °C) for significant volumes of melt generation (e.g., Patino-Douce and Johnson, 1991), consistent with thermodynamically calculat-ed smallmelt volume (b10 vol.%). Thus, the exposed rocks cannot be thedirect source of the voluminous granitic S-typemagmas, but rather theyrepresent an indication of what was likely taking place at deeper andhotter crustal levels.
7.3. Extent of high-T, low-P 371 Ma metamorphism
Together withWest Antarctica and southeast Australia, theWesternProvince of New Zealand formed part of the Palaeozoic Gondwana con-tinental margin. New Zealand's Western Province, probably the leastwell known of these three areas, occupies a central position betweenAntarctica and Australia (Fig. 9). Within New Zealand, the WesternProvince has been dissected by Late Cenozoic movement along theAlpine Fault, with Fiordland dextrally offset by 480 km. Garnet-bearingparagneisses in portions of central Fiordland record high-T low-Pmeta-morphism (~640–680 °C, 3–5 kb) and partial melting at c. 370–360 Ma(Gibson, 1992; Ireland and Gibson, 1998) contemporaneouswithmeta-morphism and partial melting in the Bonar Range (Fig. 9). One distinctdifference is that large areas of Fiordland were affected by thickeningand equilibration at kyanite grade conditions (~6–8 kb, 620–650 °C)at c. 350–340 Ma (Daczko et al., 2009; Gibson, 1992; Ireland andGibson,1998; Scott et al., 2009a,b) (Fig. 8), and then extensively modified dur-ing Early Cretaceous granulite facies metamorphism and deformation(Bradshaw, 1989; Clarke et al., 2000; Daczko et al., 2001a,b, 2002a,b).The record of Palaeozoic high-Tmetamorphism in Fiordland is thus pre-sent but difficult to resolve.
There are striking lithological similarities between Westland–Fiordland and parts of West Antarctica. In Marie Byrd Land, the Swan-son Formation is equivalent to the Greenland Group, as is the Ford
Fig. 9. A reconstruction of the Palaeozoic paleo-Pacific Gondwana margin and a summary of relevant geochronology and tectonic events. The hatched area in the upper diagramindicates the inferred area of Palaeozoic partial melting. The U–Pb data for orthogneiss A557 is from Hiess et al. (2010); plutonic suite information is from Tulloch et al.(2009b); and Fosdick Range data is from Siddoway et al. (2004) and Siddoway and Fanning (2009).
531J. Scott et al. / Lithos 127 (2011) 522–534
Granodiorite to the Karamea Suite (Bradshaw et al., 1997; Pankhurst etal., 1998; Tulloch et al., 2009b). The Palaeozoic metamorphic history ofthe Fosdick Mountains in Marie Byrd Land also records a metamorphicevent of high-T low-P character thatwas associatedwith partialmeltingat c. 370 Ma (Siddoway and Fanning, 2009). However, like Fiordland,Cretaceous deformation and partial melting complicate the Palaeozoicgeological story (Korhonen et al., 2010; McFadden et al., 2010; Siddo-way et al., 2004). The cause of ~371 Ma high-T low-P metamorphismand associated S-type granite emplacement along 1000 s of km of theGondwana continental margin must be related to some form of criticallarge-scale tectonic re-organisation at this time.
7.4. Cause of crustal melting in the New Zealand Devonian Gondwanasegment
The geothermal gradient for Palaeozoic metamorphism in the BonarRange (36 °C km−1) is higher than typical for a stable crustal geother-mal gradient (≤30 °C km−1). Therefore, a thermal perturbation is re-quired to have disturbed the crust enough to promote partial melting.This could have been achieved via isothermal decompression; advectiveheating from intra-/under-plated mafic magmas or juxtaposed hotlower crust/upper mantle; conductive heating by lithosphere removaland replacement by hot asthenosphere; conductive heating by shal-lowed mantle during extensional tectonism; or some combination ofthese processes.
Anumber of crustalmeltingmechanisms can be tentatively ruled out.Firstly, isothermal decompression is not evident in the P–T record or rel-ict mineralogy of the Bonar Range, and available data suggests probablynot in Fiordland either (e.g., Ireland and Gibson, 1998). Similarly, there isno evidence for rapid exhumation of Devonian lower crust/uppermantleand associated advective heat loss against Bonar Range or Fiordland li-thologies. There is a large amount of ~370–380 Ma S-type graniticorthogneiss in the Bonar Range (Jongens, 2006) that could have heatedthe area, although no plutons of this age have yet been identified inFiordland (Allibone et al., 2007, 2009a, b). However, as the S-type gran-itoids themselves require a crustal thermal perturbation to be generated,invoking them as the cause of crustal melting is circular and does not ad-dress the question of a regional melt trigger. There is no evidence forlarge volumes of Devonian mafic intra-plating in the deepest WesternProvince (exposed in Fiordland) (Allibone et al., 2007, 2009a, b), andseismic profiles across Fiordland (Eberhart-Phillips and Reyners, 2001)are not consistent with the presence of mafic under-plate prior to theCretaceous.
Evidence for lithosphere replaced by asthenosphere is also limited. I-type plutons that have intruded through Western Province crust con-tain Cambrian to Proterozoic bulk rock Sr and Nd and Hf-in-zircon de-pleted mantle model ages (e.g., Bolhar et al. 2008; Scott et al., 2009b;Tulloch et al., 2009b). Tertiary lamprophyres and basalts acrossNew Zealand likewise point to the presence of ≥Cambrian-agedmantle source at their time of eruption (Barreiro and Cooper,
1987; Coombs et al., 1986; Hoernle et al., 2006; Timm et al., 2010).These data do not support the idea of removal of lithosphere andreplenishment by hot asthenosphere beneath the WesternProvince in the Devonian, or emplacement and then removal of aDevonian under-plate.
On the other hand, several of the issues raised can be addressed ifthe middle and lower crust were thinned during Devonian continen-tal extension and conductively heated by a shallowed lithosphericmantle. Firstly, this relationship would provide a source for heat aswell as an explanation for the lack of evidence of Devonian mantle re-plenishment event. Secondly, because rapidly upwelled astheno-sphere should generate large volumes of melt, a sub-continentallithospheric mantle heat source would provide a simple explanationfor the limited extent of Devonianmafic magmas. Thirdly, a regionallyshallowed lithospheric mantle could have simultaneously generatedlarge volumes of felsic melt over the 1000 s of km from New Zealandto West Antarctica. Fourthly, a shallowed lithospheric mantle shouldcause protracted crustal partial melting because the heat source wasconductive, which is implied by the homogenisation of peak meta-morphic garnet in the Bonar Range (there is insufficient data todraw any conclusions of the zoning of Devonian garnet in Fiordland).This is not to say that the available evidence rules out a component ofasthenospheric heating, but rather that its contribution was, if at all,probably minor. A tectonic cause for a shallowed lithospheric mantlecould have been due to the halt of subduction or slab rollback alongthis Gondwana continental margin; however, further constraintson exactly when subduction ceased are required before a unifyingmodel can be proposed.
The gneisses of the Bonar Range were exhumed several kbar fol-lowing partial melting. U–Pb monazite ages from sample P67589and a typical crustal density value of 2.71 g/cm3 yields an error-prop-agated cooling rate of 5.6±3.0 °C Ma−1 and an exhumation rate of600±400 m/Ma−1. These are minimum rates because it is possiblethat the younger monazite ages do not represent the exact point ofinflection to slow cooling on the P–T path. Nevertheless, an importantconsideration is the clear presence of partial melt segregations. Aspartially molten rocks are more buoyant than crystalline crust (e.g.,Whitney et al., 2004), initial exhumation of the Bonar Range couldhave, at least in part, been related to a density contrast caused bythe presence of in-situ melt.
8. Conclusions
(1) Careful analysis of monazite textural relationships and associa-tions, composition and age data is capable of resolving multiplemetamorphic episodes. In this context, the Bonar Range of NewZealand is shown to not record Early Cretaceous deformationbut rather mid-Palaeozoic crystallisation.
(2) The P–T path andmetamorphic reactions for the studied gneissicBonar Range rocks outlines a history of crustal partial melting atan elevated geothermal gradient followed by decompression andcooling. The crustal partial melting event was contemporaneouswith the initiation of rapid emplacement of the large volume S-type Karamea Suite at 371 Ma in Westland as well as metamor-phism in Fiordland and West Antarctica.
(3) The striking similarities between the Devonian tectonics of theBonar Range, Fiordland, and West Antarctica indicate that thepalaeo-Pacific Gondwana margin experienced high heat flowand regional crustal melting in the mid-Palaeozoic. A majorasthenospheric heat source is tentatively ruled out, and a processwhereby crustal extensional led to conductive heating by a shal-lowed sub-continental mantle lithosphere is proposed.
Supplementary materials related to this article can be found onlineat: doi:10.1016/j.lithos.2011.09.008.
Acknowledgements
We acknowledge with thanks: the Western Australian SHRIMPfacilities for monazite U–Th–Pb dating—which were operated by a WAuniversity–government consortium with ARC support; the AustralianMicroscopy andMicroanalysis Research Facility at theCentre forMicros-copy, Characterisation and Analysis, The University ofWestern Australia—a facility that is funded by the Australian University, State and Com-monwealth Governments; the University of Otago for research grantsto JMP for undertaking zircon analyses; the Research School of EarthSciences at the Australian National University for permitting zirconanalyses,whichwere conducted under the guidance of CAllen; R. Jongensfor providing sample P67589; A. Cooper, N. Eby, B. Rasmussen, U. Ring,and A. Tulloch for comments on drafts; N. Daczko and an anonymous per-son for journal reviews.
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