Lithos 126 (2011) 1–21

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The Longwood Igneous Complex, Southland, New Zealand: A Permo-Jurassic,intra-oceanic, subduction-related, I-type batholithic complex

Richard Price a,⁎, Carl Spandler b, Richard Arculus c, Anthony Reay d

a Faculty of Science and Engineering, University of Waikato, Hamilton, New Zealandb Earth and Environmental Sciences, James Cook University, Townsville, Queensland 4811, Australiac Research School of Earth Sciences, Australian National University, Canberra, ACT 0200, Australiad Department of Geology, University of Otago, Dunedin, New Zealand

⁎ Corresponding author at: 13 Parkwood St., AlfredTel.: +61 3 5334 3811.

E-mail address: r.price@waikato.ac.nz (R. Price).

0024-4937/$ – see front matter © 2011 Elsevier B.V. Aldoi:10.1016/j.lithos.2011.04.006

a b s t r a c t

a r t i c l e i n f o

Article history:Received 10 November 2010Accepted 23 April 2011Available online 1 May 2011

Keywords:GabbroDioriteI-type graniteEnclavesSubduction-related magmatismNew Zealand

The Longwood Igneous Complex (LIC) is located in Southland, New Zealand on the eastern side of theCarboniferous to Cretaceous, I-type, Median Batholith. Intrusives of the Complex range in age from Permian toJurassic and show trace element characteristics typical of subduction-related magmas. Gabbro, gabbroicdiorite and basaltic dyke rocks show trace and minor element patterns and isotopic compositions indicatingthat they represent magmas generated in an intra-oceanic subduction system. Radiometric ages decreaseacross the LIC from 254 Ma in the east to 142 Ma in the west and mineral chemistry and mineral phaserelationships indicate emplacement at depths between 15 and 25 km. Thus the petrology and geochemistry ofthe LIC provides the basis for evaluating the composition of lower–middle crust assembled above a long livedintra-oceanic subduction system and we estimate this to be andesitic and similar to bulk continental crust.Rocks of the LIC range in composition from troctolite and gabbro through diorite to trondhjemite and granite.All of the ultramafic rocks and most of the gabbros have petrographic and geochemical features consistentwith a cumulate origin and mineral chemistry shows similarities with arc cumulate sequences fromelsewhere. Few of the plutonic rocks making up the LIC have direct analogues among modern intra-oceanicvolcanic rocks. The latter are the end products and the former the leftovers from magmatic processes thatincluded fractional crystallisation, crustal assimilation andmagmamixing andmingling. Longwood intrusionsdo not represent magma chambers. They formed as crystal cumulates andmushes left over from the processesthat generated magmas erupted at the contemporary volcanic arc.A correlation between decreasing age of emplacement and Sr and Nd isotopic compositions and inheritance inzircons dated by ion probe are indications of crustal recycling. The generation of felsic rocks in the Longwoodintra-oceanic arc involved crustal anatexis and, over the 100 million year history of the arc, the crust evolvedtowards a composition similar to bulk continental crust and average andesite.Dioritic rocks of the LIC contain abundant mafic enclaves, which are argued to represent fragments of maficmagma, derived by fractional crystallisation from basalt, which was intruded into a hot but solid or near soliddiorite. Heating and remobilisation of the dioritic host disrupted and disaggregated the intruding maficmagma to form enclaves and zones of intrusion breccia that show every variation from liquid–liquid to liquid–solid mingling and mixing. They were then further modified chemically and mineralogically by diffusion ofH2O, Na, P, Ba, REE, and, to a lesser extent, Rb.Mafic dykes occur throughout the Complex and a number of these are composite with compositions rangingfrom dolerite through andesite to dacite. The components of composite dykes do not define unequivocallinear mixing trends and hybridisation processes that took place within them have only localised significance;mingling and hybridisation in the composite dykes do not appear to have controlled geochemical variationamong the major intrusive units of the Complex.

ton, Victoria 3350, Australia.

l rights reserved.

© 2011 Elsevier B.V. All rights reserved.

1. Introduction

The Longwood Igneous Complex (LIC) in the southern South Islandof New Zealand comprises intrusive igneous rocks, ranging in

composition from ultramafic to granite (Bignall, 1987; Challis andLauder, 1977; Cowden et al., 1990; Price and Sinton, 1978; Rombouts,1994), which were emplaced in a long-lived intra-oceanic magmaticarc (Mortimer et al., 1999a). Exposures of intrusive rocks fromsubduction-related magmatic systems are relatively unusual and conse-quently they provide a rare opportunity to study magma generation andevolution in a subduction setting. More specifically, the Longwoodintrusives provide insights into the processes by which I-type granitic

2 R. Price et al. / Lithos 126 (2011) 1–21

batholiths are constructed and an opportunity to evaluate the relativecontributions of fractional crystallisation, crystal accumulation, magmamixing andmingling and crustal recycling to geochemical variation in anintra-oceanic, subduction-related I-type intrusive suite. Collectively, theLongwood intrusive suite represents across-section throughsub-arc crustand consequently ameans of estimating the composition of crust formedwithin a long-lived oceanic subduction system.

The geology of the South Island of New Zealand is dominated bytwo major tectonic provinces with contrasting metamorphic andstructural histories and different lithological assemblages (Fig. 1). TheWestern Province, which is composed of Cambrian–Silurian agedmeta-sedimentary rocks cut by Devonian to Carboniferous granitoids,is considered to represent a piece of lithosphere from the Palaeozoiccontinental margin of Gondwana (Mortimer et al., 1999b). TheEastern Province formed as the result of convergent margin processesand contains arc-volcanic rocks, arc-derived sedimentary sequences,and accretionary complexes of Permian–Cretaceous age (e.g. Adamsand Graham, 1996; Adams et al., 1998; Mortimer et al., 1999b; Roserand Korsch, 1999). The boundary between the two provinces wasoriginally considered to be a linear suture between adjoining or pairedmetamorphic belts and consequently it was termed the MedianTectonic Line (Landis and Coombs, 1967). Further work saw thesuture redefined as a broad zone (Bradshaw, 1993; Frost and Coombs,1989; Kimbrough et al., 1993, 1994; Muir et al., 1998; Williams andHarper, 1978) and most recently Mortimer et al. (1999b) suggestedthe boundary is stitched by a Median Batholith comprising intrusiverocks ranging in age from Carboniferous to Early Cretaceous (Fig. 1).Price et al. (2006) highlighted the difficulty of differentiating intrusiverocks of the Median Batholith from those of the Eastern and WesternProvinces; within each of these tectonic terranes and in the MedianBatholith, igneous rocks have the petrological characteristics expectedof those formed in subduction settings (e.g. Coombs et al., 1976;Houghton and Landis, 1989; Muir et al., 1998; Spandler et al., 2003).Detailed geochronology provides the only means of assigning specificintrusions to a particular province or to the Median Batholith.

1.1. The geology of the Longwood Igneous Complex (LIC)

The LIC lies within the southern and eastern margin of the MedianBatholith, approximately 40 km to the west of Invercargill in the farsouth of the South Island. It forms a low range of hills (maximumelevation of 804 m at Bald Hill) extending ~35 kmnorthwards from the

Fig. 1. (A) location of the Longwood Igneous Complex (LIC) in relation to the Median(B) Interpretive map showing distribution of plutonic units of the LIC (Mortimer et al., 199

coast (Fig. 1). The geology, lithologies and age relationships of the LIChave been described in detail by Bignall (1987), Mortimer et al. (1999a),Rombouts (1994) and Price et al. (2006). Inland, the Complex is heavilyforested and rocks deeply weathered but Mortimer et al. (1999a) usedavailable outcrop and petrological and geochronological information todevelop an interpretative map, which is reproduced here as Fig. 1B. Tothe east, the oldest intrusives of the Complex are interpreted to havebeen emplaced into volcanic and volcaniclastic sediments of the PermianBrook Street Terrane with successive intrusive episodes becomingprogressively younger to the west. Mortimer et al. (1999a) grouped theintrusive rocks of the Complex into four broad geochronological andpetrologicalunits (Fig. 1B): (a) Early toMiddlePermian felsic rocks of thePourakino Trondhjemite; (b) Late Permian to Early Triassic gabbros anddiorites of the Hekeia Gabbro; (c) Middle and Late Triassic diorites andgranites of the Holly Burn Intrusives; and (d) Late Triassic–Early Jurassicgabbros of the Pahia Intrusives. The age progression from east to westacross the Complex is summarised in Fig. 2A, which incorporates datafromMortimer et al. (1999a) andPrice et al. (2006). The LIC represents abatholith that was assembled sequentially from east to west beneath anintra-oceanic subduction-related volcanic arc on a 100 million year timescale.

The focus of this paper is on excellent exposures found along thecoast between Oraka and Pahia Points (Figs. 1 and 3). These provide a20 km cross-section through the LIC incorporating all of the intrusivegroupings recognised by Mortimer et al. (1999a) except the PourakinoTrondhjemite. At the eastern extremity of the section, at Oraka Point,Late Permian to Early Triassic (245 Ma; Price et al., 2006) gabbroic rocksare bordered by granite to the east and diorite to the south–east andwest. The contacts are marked by a mixed or mingled zone, implyingthat mafic and more felsic magmas co-existed. The middle of the crosssection, between Wakaputa Point and Mullet Bay is dominated bydioritic rocks but gabbro also occurs and, as atOraka Point, contacts tendto be transitional rather than sharp. Diorites from this segment of thecross section give Triassic ages (215–227 Ma; Price et al., 2006). At thewestern end of the section from Mullet Bay through Pahia Point to TeWae Wae Bay, gabbroic rocks dominate although ultramafic rocks anddiorite also occur and contacts between felsic and mafic rocks aremarked by mixed or mingled zones. Pahia Point rocks give Late Triassicto Jurassic ages (203–211 Ma; Price et al., 2006). A leuco-gabbro dyke atPahia Point gives a SHRIMP zircon age of 142±2 Ma and appears to berelated to the Anglem Complex on Stewart Island to the south of theLongwood coast across Foveaux Strait (Price et al., 2006).

Batholith of South Island, New Zealand (Mortimer et al., 1999a; Muir et al., 1996).9a).

Fig. 2. Age (A) and initial 87Sr/86Sr isotopic ratio (B) of plutonic rocks across theLongwood Igneous Complex (LIC) from east to west. Zircon SHRIMP ages are U–Pb agesdetermined on zircon using sensitive high resolution ion microprobe and Zircon U/Pbages were determined using conventional methods on individual zircon grains orzircon separates.Data are from Devereux et al., 1968; Kimbrough et al., 1994; Mortimer et al., 1999a;Price et al., 2006; Tulloch et al., 1999.

3R. Price et al. / Lithos 126 (2011) 1–21

2. Methods

2.1. Mineral chemistry

Major element analyses of minerals from samples with prefixes“PP”, “WA” and “OR” are from Price and Sinton (1978) and methodsare described therein. All other analyses were carried out at theElectron Microscopy Unit, Australian National University (ANU) usingpolished thin sections of rock samples and a JEOL 6400 scanningelectron microscope equipped with an energy-dispersive spectrom-eter (EDS). Acceleration voltage, beam current and counting timeswere set to 15 kV, 1 nA and 100 s respectively. Element concentra-tions were standardised against silicate mineral standards of known

Fig. 3. Geological map of coastal section through the Longwood Igneous Complex betweendescribed in this paper.

composition produced by Astimex Scientific Limited. ZAF correctionswere applied to all analyses.

2.2. Bulk-rock geochemistry

Major and selected trace element analyses were determined byX-ray fluorescence spectroscopy at La Trobe University (Price andSinton, 1978) and the ANU using lithium metaborate fused glass discsand pressed powder pellets and methods described by Norrish andHutton (1969) and Norrish and Chappell (1977). Precision for majorelements is generally better than ±1% (1σ). For trace elements,theoretical detection limits are of the order of 1 ppmand reproducibilityis better than ±5% (1σ). FeO abundances were determined by directtitration using standardised CeSO4 solution.

Rare earth element (REE), Cs, Pb, Th, U, Nb, Hf, Y and Sc data wereobtained at Monash University using solution techniques andinductively coupled plasma source mass spectrometry (ICP-MS) andat the ANU using laser ablation (LA) ICP-MS.

Sampleswithprefixes “PP”, “WA” and “OR”were analysedby ICP-MSanalysis at theMonashUniversity followingmethods described by Priceet al. (1997). 100 mg aliquots of sample powder were digested inHF–HNO3 in Teflon pressure vessels over several days. The solute wasthen refluxed in HNO3 and taken up in 50 ml of 2% HNO3. Samplesolutions were diluted and spiked with an internal standard (100 ppbiridium). ICP-MS analysis was carried out on a VG PlasmaQuad PQ2+ inpeak jumping mode using AGV-1 and BHVO-1 as calibration standards.Analytical blanks were b10 ppb for all elements. Repeat analyses ofBHVO-1 indicate precision of better than 5% at the 95% confidence level.

Samples analysed by LA-ICP-MS at the ANU were prepared as fusedglass discs (Li-metaborate flux: sample=5:1), which were ablated andanalysed using an EXCIMER laser system, operating in the ultra-violetspectrum at awavelength of 193 nm (Eggins et al., 1998) and coupled toa Fisons PQ2 STE ICP-MS. Spot size was 90–120 mμ. The standard NISTSRM 612 glass was run after every 15 analyses. Si concentrationsobtained by XRF were used as the internal standard to account for anyvariation in ablation yield between samples and calibration standardsand the ‘matrix effect’ of variations between counts per second and ppmof different elements. During the course of the analysis period, a BCR-2 gglass standard was also analysed (every 30 analyses) to provide an

Te Waewae Bay and Colac Bay (Price and Sinton, 1978) showing location of samples

Table 1Representative whole rock analyses from the Longwood Igneous Complex.

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23

BB34A BB35A BB38 BB39 PP21 PP22 WA6A WA31 BB31 BB41E WA1A WA1XC WA1XE WA22X1 BB40 BB41B BB41C BB41D PP12A OR8 OR3 OR7B P2837

Troct. Troct. Gabbro Gabbro Diorite Diorite G-diorite G-diorite Diorite Diorite Diorite Enclave Enclave Enclave Dyke Dyke Dyke Dyke Dyke Gabbro Diorite Granite Trondh.

SiO2 42.41 42.81 44.74 44.95 60.12 60.30 52.47 54.87 58.35 60.62 64.93 53.65 53.07 55.15 48.65 55.34 52.89 60.18 41.46 48.93 58.39 76.02 70.59TiO2 0.29 0.39 0.29 0.22 0.89 0.98 1.23 0.80 0.74 0.87 0.62 1.27 1.20 1.14 0.90 0.88 0.95 0.69 1.91 1.32 1.25 0.21 0.19Al2O3 8.03 8.28 23.32 19.68 16.74 17.07 16.63 17.97 19.44 17.11 16.90 18.20 18.82 19.12 16.55 17.72 16.88 18.24 18.79 18.66 15.48 12.41 16.17Fe2O3 4.05 3.87 2.38 1.51 2.05 2.14 3.78 2.25 2.27 2.50 1.80 3.56 3.81 3.08 2.81 2.60 2.95 2.88 6.20 4.40 2.24 0.57 1.45FeO 10.68 10.41 4.15 4.58 3.94 3.89 5.96 5.46 2.95 3.21 2.22 4.62 4.72 3.76 6.24 4.41 4.60 3.24 7.47 5.75 5.27 0.81MnO 0.20 0.19 0.10 0.08 0.11 0.13 0.15 0.14 0.11 0.09 0.09 0.23 0.22 0.23 0.13 0.12 0.12 0.10 0.16 0.16 0.15 0.03 0.05MgO 23.59 22.62 8.19 11.80 2.56 2.30 5.18 4.13 2.15 3.03 1.43 3.56 3.34 2.35 9.49 4.81 6.31 2.34 7.41 5.57 3.14 0.42 0.31CaO 7.88 8.00 13.88 14.52 5.34 4.83 7.61 8.14 5.37 5.38 3.74 6.81 6.48 6.05 9.90 6.77 5.86 5.59 12.53 10.31 5.75 1.01 3.08Na2O 0.57 0.98 1.10 0.81 4.23 4.22 3.89 3.50 5.71 2.92 4.87 5.31 5.45 6.42 1.99 2.72 2.38 3.13 2.17 3.35 4.13 2.92 5.37K2O 0.10 0.10 0.05 0.05 2.26 2.62 1.01 1.25 1.28 1.86 2.55 1.24 1.48 0.91 0.59 1.72 3.49 1.45 0.29 0.26 2.24 5.11 1.04P2O5 0.04 0.06 0.04 0.04 0.19 0.19 0.34 0.17 0.28 0.27 0.18 0.22 0.34 0.39 0.19 0.27 0.37 0.31 0.35 0.19 0.34 0.01 0.04H2O- 1.12 1.21 1.46 1.35 0.99 0.90 0.99 1.22 0.82 0.90 0.58 0.86 0.95 0.64 1.65 1.57 2.17 0.76 1.34 0.95 1.33 0.53H2O+ 0.13 0.15 0.19 0.22 0.03 0.13 0.11 0.14 0.35 0.14CO2 0.50 0.30 0.25 0.20 0.10 0.11 0.10 0.10 0.37 0.06 0.07 0.14 0.15 0.04 0.07 0.07 0.07 0.08 0.43 0.04 0.06S 0.02 0.04 0.03 0.03 0.06 0.02 0.10 0.02 0.03 0.05LOI 0.75Total 99.61 99.41 100.17 100.04 99.52 99.68 99.34 100.00 99.93 98.97 99.98 99.67 100.03 99.28 99.37 99.16 99.42 99.18 100.51 99.85 99.75 100.11 99.04Corr. S 99.60 99.39 100.16 100.03 99.52 99.68 99.34 100.00 99.90 98.96 99.98 99.67 100.03 99.28 99.32 99.15 99.41 99.16 100.51 99.85 99.75 100.11 99.04Cs 0.1 0.1 0.2 0.1 4.9 6.4 1.0 3.5 1.6 3.3 3.5 3.4 4.4 1.2 0.7 2.0 3.1 2.9 0.1 0.3 14.3 6.3Ba 53 78 41 42 412 444 303 221 433 382 591 157 127 125 176 428 728 348 40 82 240 561 32Rb 3 2 3 3 74 94 25 39 b1 57 97 62 85 21 15 61 149 41 3 3 72 162 16Sr 402 500 1044 1106 508 467 643 501 627 649 451 501 454 560 680 833 732 687 769 715 310 109 561Pb 2.6 2.3 1.4 1.5 15.0 18.1 9.9 14 11.4 12.6 19 9.1 9.3 11.9 7.0 12.2 22 12.6 2.9 2.8 32.0 20.7 12Th 0.24 0.48 0.10 0.08 1.56 8.00 1.93 5.50 4.78 6.04 10.40 6.43 6.92 1.71 0.56 2.61 2.0 6.04 0.12 0.16 4.79 10.28U 0.09 0.16 0.03 0.02 2.57 1.39 0.53 0.50 1.26 1.34 3.00 2.19 3.45 1.25 0.19 0.76 1.35 0.07 0.08 1.00 2.25Zr 34 42 72 77 208 260 142 99 170 175 229 121 72 203 80 110 99 147 36 37 343 149 62Hf 0.3 0.4 0.7 0.2 6.1 4.2 1.5 1.9 4.1 4.4 4.8 3.6 2.3 4.6 1.6 2.0 1.2 4.4 1.1 1.4 2.9 5.7Ta 0.08 0.10 0.47 0.17 0.75 1.10 0.43 0.42 0.68 0.86 0.91 0.45 0.24 0.36 0.68 0.37 0.47 0.93 0.76Nb 0.6 0.4 0.7 0.1 4.7 6.6 4.0 2 3.8 4.9 6 9.8 10.8 9.4 2.0 2.4 2 4.9 1.6 1.7 7.7 1.3 3Y 2.4 2.9 1.8 2.3 20.9 23 22.7 16 12.9 16.8 21 33.5 32.2 36.0 11.7 10.6 9 16.8 14.5 14.3 32 12 4La 1.8 2.6 1.2 1.0 3.6 17.1 15.7 11.2 17.0 15.6 19.4 23.8 21.5 20.4 7.7 11.1 13.0 15.6 3.4 3.6 18.1 9.4 8Ce 3.9 5.5 2.3 2.3 26.1 36.1 37.6 26.5 33.5 34.5 48.0 61.7 54.2 52.1 18.5 22.9 27.0 34.8 10.8 11.1 47.3 19.5Pr 0.5 0.7 0.3 0.4 1.0 5.7 5.5 4.0 4.5 8.4 7.5 8.3 2.6 2.9 4.5 1.9 1.9 6.8 2.3Nd 2.1 3.1 1.4 1.8 3.6 22.1 24.0 15.4 15.9 18.3 25.0 33.4 30.5 37.0 11.1 11.9 14.0 18.3 10.0 9.2 29.0 8.0Sm 0.6 0.8 0.4 0.6 0.7 5.0 5.5 4.2 3.3 4.1 4.8 7.3 6.6 8.2 2.5 2.7 2.5 4.1 3.0 2.7 7.0 1.5Eu 0.25 0.31 0.35 0.29 0.21 1.15 1.42 1.10 1.37 1.08 1.04 1.62 1.39 0.83 0.89 0.92 0.95 1.08 0.99 0.96 1.41 0.33Gd 0.53 0.70 0.42 0.54 0.73 4.70 5.15 4.30 2.67 3.41 4.70 6.73 6.15 7.58 2.39 2.23 2.80 3.41 3.15 2.84 6.95 1.24Tb 0.08 0.06 0.10 0.72 0.76 0.69 0.68 0.76 0.96 0.92 1.13 0.35 0.45 0.47 0.43 1.07 0.18Dy 0.46 0.57 0.34 0.46 0.54 4.02 4.00 2.21 2.97 5.49 5.23 6.12 1.97 1.88 2.97 2.60 2.41 6.10 1.06Ho 0.08 0.06 0.10 0.83 0.79 0.80 0.80 0.95 1.12 1.05 1.24 0.39 0.50 0.53 0.51 1.26 0.23Er 0.24 0.28 0.16 0.23 0.28 2.24 1.91 1.21 1.55 2.97 2.75 3.21 1.05 1.01 1.55 1.34 1.29 3.26 0.70Tm 0.03 0.02 0.04 0.35 0.27 0.46 0.41 0.45 0.15 0.18 0.19 0.49 0.12Yb 0.19 0.23 0.12 0.17 0.27 2.25 1.68 1.89 1.17 1.45 2.35 2.98 2.63 2.78 0.97 0.92 0.71 1.45 1.13 1.20 3.15 0.90Lu 0.02 0.03 0.02 0.03 0.04 0.36 0.25 0.30 0.19 0.22 0.33 0.47 0.37 0.40 0.16 0.13 0.10 0.22 0.17 0.18 0.47 0.16Sc 26 26 20 28 3.9 15.1 7.9 23.0 12 14 6.2 20.6 17.8 10.6 26 21 18.0 10 35.9 23.8 21.5 2.3 6V 138 159 87 73 129 143 295 206 105 123 63 187 148 104 216 181 181 82 477 366 160 15 22Cr 1601 1351 473 580 17 15 143 43 16 17 13 9 8 5 640 42 162 4 53 52 19 9 5Ni 583 524 182 190 6 7 39 11 0 21 5 9 4 1 176 46 104 9 30 26 14 5 5Cu 17 34 14 10 28 29 30 40 25 15 13 64 58 17 105 71 84 84 248 220 180 9 101Zn 122 120 51 43 61 85 99 85 81 82 66 133 111 93 86 88 101 84 136 112 109 24 48Ga 10 11 473 580 21 24 17 21 24 640 42 162 4 21 23 19 13 18(La/Yb)n 6.65 7.85 7.11 4.00 9.41 5.46 6.68 4.25 10.39 7.72 5.92 5.72 5.86 5.26 5.67 8.60 13.13 7.70 2.15 2.15 4.11 7.51(Eu/Eu*)n 1.422 1.241 2.655 1.611 0.871 0.719 0.800 0.784 1.373 0.866 0.661 0.694 0.658 0.316 1.098 1.119 1.093 0.866 0.992 1.059 0.614 0.735

Analyses in italics are ICP-MS data. REE, Cs, Hf, Th, U and Sc for sample BB41C (column 17) by INNA. All other data are XRF. 1–6 Pahia Pt to Mullet Bay.7–11 Ruahine Hill to Wakaputa Pt. 12–14 enclaves. 15–19 dykes. 20–22 Oraka Pt. 23 fromMortimer et al. (1999b). Troct. = troctolite.G-diorite = gabbroic diorite. Trondh. = trondjhemite. 16–18 are phases of composite dyke at locality BB41. 19 is a hornblendite dyke.

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independent assessment of accuracy and precision. The replicateanalyses of this standard rock indicate that LA-ICP-MS data for Cs, Hf,Nb, Pb, Th, U and the rare earth elements are within 5% of publishedreference values.

For three samples (including BB41C, listed in Table 1) the REE, Cs,Hf, Th, U and Sc were analysed by instrumental neutron activation(INAA) using methods described by Chappell and Hergt (1986). Forthese elements, INAA gives precision and accuracy (based on replicateanalyses of BHVO-1) better than ±5% (2σ).

3. Petrography and mineral chemistry

3.1. Petrography

Ultramafic rocks outcrop on the Longwood Tops and at Pahia Point.They are coarse-grained, feldspathic peridotites and troctolites, consist-ing of cumulus olivine (Fo74–82) and calcic plagioclase (An80–97).

Fig. 4. Photomicrographs of rocks from the Longwood Igneous Complex. (A–B): Sample BB35amphibole (Am). (C–D): Sample BB35B from Pahia Point. Olivine (Ol) cumulate with poikilitishowing weak cumulate texture. Cpx = clinopyroxene, Pl = plagioclase, Bi = biotite. (G–H)are 0.5 mm. Left panels (A,C,E,G) are views under crossed polars and right panels (B,D,F,H)

Intercumulus minerals include clinopyroxene, orthopyroxene, horn-blende and phlogopite. In the Pahia Point examples large (2–3 cm)poikilitic brownhornblende and biotite oikocrysts enclose the cumulateolivine and plagioclase (Fig. 4). Chromian-magnetite, pyrite, pyrrhotiteand chalcopyrite occur in trace amounts within the Longwood Topsultramafic rocks. In some of these samples and others from Pahia Pointaluminous green spinel forms coronae between plagioclase and olivine.Alteration minerals include actinolite, cummingtonite, anthophyllite,and chlorite. The ultramafic unit at Pahia Point contains inclusions oflayered, cumulate-textured, hornblende-bearing, noritic gabbro up to10 cm in maximum dimension.

Gabbroic rocks are well exposed at Pahia Point, Ruahine Hill,Wakaputa Point and Oraka Point and they are also reported as layersassociated with ultramafic rocks on the Longwood Tops. Mostcommonly they are coarse to fine-grained olivine gabbro, gabbro-noriteand hornblende gabbro (Fig. 4). Olivine gabbro consists of cumulusplagioclase (An45–97) and up to 15% cumulus olivine (Fo73–79) with

A from Pahia Point. Plagioclase (Pl)—olivine (Ol) cumulate with poikilitic intercumulatec intercumulate phlogopite (Phl). (E–F) Sample PP5A from Pahia Point. Pyroxene gabbroEnclave WA22X3 fromWakaputa Point. Pl = plagioclase, Am = amphibole. Scale barsare views in polarised light.

6 R. Price et al. / Lithos 126 (2011) 1–21

variable modal abundances (up to 35%) of postcumulus clinopyroxene[Mg#=79–86; Mg#=100*mol. Mg/(Mg+Fe*)]. Magnetite, amphi-bole and pleonaste spinel occur as minor phases and pyrite andpyrrhotite are rare. Alteration minerals include chlorite, epidote andtitanite. Gabbro-norite consists of 50–70% plagioclase (An45–90), 20–30%clinopyroxene (Mg#=68–70), up to10%orthopyroxene (Mg#=56–70),with b5% magnetite and minor hornblende. Hornblende gabbrocomprises coarse, possibly cumulate, brown hornblende and plagioclase(An45–55). Anorthosite has been reported from the Longwood Tops(Challis and Lauder, 1977; Mortimer et al., 1999a).

Overall, the crystallisation sequence that led to development of theultramafic and mafic cumulate rocks was olivine→plagioclase→clinopyroxene+orthopyroxene+magnetite→ hornblende→phlogopite.

Diorite and quartz diorite are of two textural varieties (Price andSinton, 1978); foliated diorites are relatively rich in mafic enclavesand non-foliated types are more homogenous and enclave poor.Diorite crops out extensively across Wakaputa Point, on the easternside of Pahia Point and on the western coast of Oraka Point. A typicalexample consists of 45–50% medium to coarse grained (up to 4 mm),mottled and oscillatory zoned plagioclase (An30–48), quartz (up to20%), biotite (15–20%), hornblende (10–15%) and alkali feldspar(b5%). Alkali feldspar is microcline and microperthite. Biotite ispleochroic from dark chocolate brown to straw yellow and horn-blende from pale blue-green to dark brown. Minor phases include (inapproximate order of abundance) magnetite (up to 2%), apatite,clinopyroxene, orthopyroxene, titanite, zircon and allanite. Chlorite iscommon as an alteration mineral within biotite and epidote isassociatedwith alteration of plagioclase and hornblende. Rare patchesof cummingtonite occur within hornblende-biotite aggregates and areprobably replacing orthopyroxene. The order of crystallisation wasapatite and magnetite followed by plagioclase and pyroxene, thenhornblende and biotite.

Enclaves are common in foliated Wakaputa diorite. They range insize from 50 cm down to microscopic mafic aggregatesb1 mm acrossand in the macroscopic varieties there is a crude correlation betweenshape and size with the larger enclaves being more angular. Mostmafic enclaves are rounded and spherical or ovoid in cross-section and

Fig. 5. (A) Swarm of mafic enclaves in diorite below Ruahine Hill. (B) Hornblendite dyke slocality BB41, Mullet Bay. “h” is host diorite and “fel” is felsic component of the dyke. (D) Cmafic components. Arrow indicates cuspate margin on mafic blob. “xen” is a xenolith of ho

many have cuspate margins. Enclaves tend to be concentrated inswarmsor zones near the contactswith gabbroic rocks (Fig. 5). A typicalmafic enclave is medium to fine grained (1–2 mm) and consists of anequigranular, polygonal aggregate of plagioclase, green-brown horn-blende and dark chocolate brown biotite with minor interstitial alkalifeldspar and, in some cases, quartz (Fig. 4). Magnetite, apatite, titaniteand zircon are minor phases. Plagioclase crystals commonly showcomplex oscillatory zoning (An26–44) with corroded and altered coresand they contain inclusions of amphibole and apatite. Pyroxene has notbeen observed in any mafic enclaves but in some examples fineaggregates of colourless amphibole within hornblende aggregates mayrepresent replaced pyroxene. Apatite is abundant as fine (b0.5 mm)needles and occurs as inclusions in all other phases. Some enclaves haveamoreporphyritic texturewith slightly larger plagioclase crystals (up to3 mm) and poikilitic biotite plates 2–3 mm across.

Granite occurs as dyke-like intrusions on the eastern side of OrakaPoint and at Cosy Nook to the east of Mullet Bay. A typical exampleconsists of a medium grained (1–2 mm), equigranular aggregate ofquartz, perthitic alkali feldspar, minor plagioclase (An18–27), reddish-brown biotite, fine (b0.5 mm) magnetite, rare apatite and zircon.Plagioclase shows weak oscillatory zoning and crystal cores are alteredto fine grained aggregates of muscovite. Quartz and alkali feldspar formmyrmekitic intergrowths. Biotite is commonly altered to pale greenchlorite.

Mortimer et al. (1999a) inferred the existence of an extensive areaof trondhjemite (the Pourakino Trondhjemite) making up the westernside of the inland Longwood Range but this rock type is not representedin the coastal section and outcrop is sparse inland.

In the Oraka to Pahia Point coastal section the contacts betweengabbroic and dioritic rocks are marked by narrow (up to 30 m) tran-sitional or hybrid zones of weakly porphyritic gabbroic diorite. Thesehybrid rocks consist of plagioclase, hornblende, biotite, alkali feldspar,clinopyroxene, magnetite, apatite and zircon. Plagioclase occurs ascomplexly zoned (An30–85) phenocrysts (up to 3 mm) with, in somecases, alkali feldspar rims. Hornblende occurs in clots or aggregateswithbiotite and someof these contain cores of cummingtonite thatmay havereplaced orthopyroxene. Clinopyroxene was observed in one of thesehybrid rocks and another contains quartz.

howing comb or harrisitic texture and felsic core—Pahia Point. (C) Composite dyke atomposite dyke at locality BB41, Mullet Bay showing contact between felsic (“fel”) andst diorite in the felsic phase.

Fig. 6. Histogram showing Longwood Igneous Complex (LIC) plagioclase compositionsby geographic area.

7R. Price et al. / Lithos 126 (2011) 1–21

Mafic dykes are common throughout the coast section betweenPahia and Oraka Points. Most are less than one metre in width and thedominant rock types are basalt (dolerite) and basaltic andesite. A typicalexample is medium grained (up to 1–2 mm) and consists of oscillatoryzoned plagioclase (An27–77), green-brown hornblende and dark brown-yellow biotite. Hornblende occurs as needle-like prisms and also inlarger (1–2 mm) clots or aggregateswith biotite. Fine remnant grains oforthopyroxene occur within some of these aggregates. Some of theandesite dykes contain rare interstitial grains of alkali feldspar and, insome cases, quartz.Minor phases includemagnetite, apatite and titaniteand, rarely olivine. Alteration is commonwith chlorite, epidote and finemuscovite replacing biotite and feldspar.

Hornblendite dykes also occur. These range from near-monomi-neralic varieties in which fine to medium grained hornblende prismswith or without plagioclase define a foliation, to spectacular comb-textured or harrisitic varieties with large (up to 5 cm) hornblendecrystals growing in from the walls to felsic cores (Fig. 5).

Composite dykes occur on the eastern side of Pahia Point nearMulletBay (Fig. 5) where they are hosted in quartz diorite. We have studiedone of these dykes, at locality BB41, in detail (Figs. 3 and 5) and thisintrusion has also been the subject of a BSc (Honours) project (Laird,2008). Textures in composite dykes are generally interpreted topreserve magmamingling and/or mixing relations between juxtaposedmafic and felsic components (e.g. Wiebe and Ulrich, 1997). Mafic blobswithin the composite dyke shown in Fig. 5 have pillow-like forms andfiner grained margins, suggesting that mafic magma chilled againstcooler felsic magma. The most mafic phase is a fine-grained (1 mm)basaltic andesite consisting of plagioclase, ortho- and clino-pyroxene,green-brownhornblende and biotite withminormagnetite and apatite.Alteration minerals include chlorite, epidote, titanite and muscovite.The most felsic component of the dyke is a fine grained (1 mm) graniteconsisting of alkali feldspar, quartz and biotite. Alteration mineralsinclude muscovite, chlorite, epidote and titanite.

Fig. 7. Longwood hornblende compositions plotted in terms of cations of Al and Tiversus Mg# [100* Mg/(Mg+Fe)] data points coded by rock type.

3.2. Mineral compositions

Representativemineral compositions are shown in SupplementaryTables 1–6 and selected projections in Figs. 6–8.

Plagioclase shows a wide range in composition both across thewhole Complex (Fig. 6) and within particular rock types (see above).Themost calcic plagioclase occurs in gabbroic rocks (An92 in an olivinegabbro from the Longwood Tops) and troctolite (up to An97) butplagioclase in most gabbro is zoned with rim compositions as low asAn45. Plagioclase in diorite shows oscillatory zoning with composi-tions varying from An50 to as sodic as An26. Plagioclase in the graniticrocks is andesine and oligoclase (An18–27). Dyke rocks have highlyvariable plagioclase compositions (An47–82).

Primary amphibole is brown hornblende (classification of Leake,1978) with Mg# ranging from 73 to 82 in ultramafic rocks, 58 to 76 ingabbros and 50 to 58 in diorites (Fig. 7). In dyke rocks, Mg#s forhornblendes are between 67 and 70. Aluminium (cations) in brownamphiboles varies between 1.79 and 2.30 in ultramafic rocks, 1.25 and2.18 in gabbro and 1.21 and 1.57 in diorite. Titanium (cations) is in therange 0.01 to 0.42 for ultramafic rocks, 0.05 to 0.41 for gabbro and 0.12to 0.21 for diorite. The highest Al (cations) and relatively high Mg(cations) are observed in hornblendes from Pahia Point (averageMg=3.25 and average Al=2.05). The average Mg (cations) and Al(cations) for Wakaputa and Oraka hornblendes are 2.74 and 1.37 and3.29 and 1.55 respectively.

Green amphiboles are hornblendes with Mg#s similar to thoseobserved in brown amphiboles (54–77) but Al and Ti are compara-tively low; Al is in the range 0.88 to1.36 in most rocks but up to 2.24 inone of the troctolites and Ti is in the range 0.01 to 0.17.

Enclaves contain relatively iron rich brown hornblendes withMg#s between 44 and 61 but green amphiboles in enclaves have

highly variable Mg#s (59–94). Amphiboles in enclaves have Al(cations) in the range 1.20 to 1.55 and Ti is relative low (0.07-0.14).

Micas are common in Longwood rocks and range in composition frombiotite to phlogopite with the latter occurring exclusively in ultramafic

Fig. 8. Longwood spinel compositions plotted in terms of Fe3+/(Cr+Al+Fe3+) versus Fe2+/(Mg+Fe2+), Cr/(Cr+Al) versus Fe2+/(Mg+Fe2+) and TiO2 (wt.%) versus Fe3+/(Cr+Al+Fe3+).Open circles aremagnetite andCr–magnetite and shaded circles pleonaste. Fields are fromBarnes andRoeder (2001). “OB” is oceanfloor basalt, “AB” isfield of island arc rocks and “A” isfield oflayered Alaskan ultramafics.

8 R. Price et al. / Lithos 126 (2011) 1–21

rocks. For biotites, Mg#s vary between 58 and 68 in gabbroic rocks, 49and 54 in quartz diorite and 58 and 62 in the Oraka granite. Titanium(cations) ranges from 0.30 to 0.51 and Al (cations) from 2.75 to 3.02.PhlogopitehasMg# ranging from85 to89, Ti (cations) 0.08 to0.38andAl(cations) 3.07 to 3.35. Mica in enclaves is biotite with Mg# varyingbetween 43 and 56, Ti (cations) 0.3 to 0.4 andAl (cations) from2.9 to 3.0.

Olivine occurs in ultramafic rocks, in a few gabbros and rarely insome of the dykes. In ultramafic rocks the olivine composition rangesfrom Fo74 to Fo82, which is similar to the range observed in olivinegabbro (Fo73–Fo79). An analysed olivine in a dyke rock (PP12B) has acomposition of Fo73.

Pyroxenes are common in ultramafic rocks and gabbro but in moreevolved rocks they are usually replaced by amphibole (hornblende andcummingtonite) and biotite. Orthopyroxene in ultramafic rocks ishypersthenewithMg#s ranging from76 to81. In gabbro orthopyroxenehas awider compositional rangewithMg#s varying between 56 and 80.In one of the dyke rocks orthopyroxene compositions have a narrowrange in Mg# from 72 to 74. A single orthopyroxene analysis from atrondhjemite from the inland Longwood Range has a Mg# of 59.Clinopyroxene is diopside or augite. In ultramafic rocks clinopyroxenehasMg# between 82 and 86. The range for gabbro is wider (66–86) andin one of the diorites remnant clinopyroxene hasMg#s between 63 and68. Clinopyroxene relics in anenclavehaveMg#s between63and64. AllLongwood clinopyroxene has Al (cation)b0.18 and Ti (cation) b0.02. In

Fig. 9. Longwood Igneous Complex (LIC) whole rock analyses plotted on (

most clinopyroxene Cr abundance is b0.05 wt.% but in ultramafic rocksCr contents are in the range 0.31 to 0.86 wt.% and in clinopyroxene froman olivine gabbro they reach 0.69 wt.%.

Spinel is ubiquitous in Longwood intrusive rocks. In gabbro anddiorite the primary spinel is magnetite whereas in ultramafic rocksthe composition is Cr–magnetite (Fig. 8). The compositional range ismost similar to that observed in zoned Alaskan mafic and ultramaficcomplexes (Fig. 8; Barnes and Roeder, 2001). Green pleonaste spineloccurs as a secondary mineral in metamorphic reaction coronae introctolite from the Longwood Tops where it is interpreted to haveformed through the reaction:

olivineþ plagioclase ¼ aluminousspinelðpleonasteÞ þ clinopyroxeneþ orthopyroxene:

(Frost, 1976; Kushiro and Yoder, 1966; Price and Wallace, 1976).

3.3. Physical conditions: estimates of temperature, pressure and oxygenfugacity

Oxygen fugacity and re-equilibration (blocking) temperatureshave been estimated for cumulate rocks of the LIC using thecompositions of coexisting olivine and spinel (Ballhaus et al., 1991)and pyroxenes (Lindsley, 1983). All oxygen fugacity calculations give

Na2O+K2O)–FeO*–MgO (AFM) ternary diagram by geographic area.

9R. Price et al. / Lithos 126 (2011) 1–21

values above the fayalite-magnetite-quartz (FMQ) synthetic oxygenbuffer and are similar to other estimates for arc cumulate rocks (e.g.Ballhaus et al., 1991). There is, however a significant difference betweenthe values obtained from Pahia Point and those calculated for the inland

Fig. 10. Selected major element Harker variation diagrams for the Longwood Igneou

Longwood rocks. The latter givehigher oxygen fugacity values consistentwith the higher Fe3+ contents of the magnetite in those rocks. There isalso a significant difference between the olivine-spinel and two-pyroxene re-equilbration temperature data for the inland Longwood

s Complex (LIC) with data points coded by geographic location and rock type.

Fig. 11. Selected trace element Harker variation diagrams for the Longwood Igneous Complex (LIC) with data points coded by geographic location and rock type.

10 R. Price et al. / Lithos 126 (2011) 1–21

Range cumulates. In contrast, re-equilibration temperatures estimatedfor Pahia Point are similar for both the olivine-spinel and two-pyroxenemethods.

The depth of crystallisation of the LIC can be approximatelyconstrained by field relationships and from the composition of primaryand secondary mineral phases. Mortimer et al. (1999a) used Al-in-

hornblende geobarometry (Anderson and Smith, 1995; Holland andBlundy, 1994; Schmidt, 1992) to obtain estimates of emplacementdepths of 4–10 km (0.17 and 0.33 GPa). These results are consistentwithfield relationships, which indicate that the granitic and dioritic rocks arein intrusive contact with the Brook Street Terrane; the maximummetamorphic grade of the Brook Street Terrane is lower greenschist

11R. Price et al. / Lithos 126 (2011) 1–21

facies, constraining the load pressure to less than 0.5 GPa (Houghton,1982; Sivell, 1984; Williams, 1978).

Al-in hornblende geobarometry is however not strictly applicable tothe cumulate rocks of the LIC although some of the quartz-bearingdioritic rocks have appropriate mineral assemblages. For these rocks,various hornblende geobarometers (Hammarstrom and Zen, 1983;Hollister et al., 1987; Johnston and Rutherford, 1989; Schmidt, 1992)give pressure estimates of 0.2 to 0.5 GPa for brown hornblende and 0.03to 0.3 GPa for secondary green hornblendes. Averaging all estimates bythe fourmethods gives pressure ranges for Pahia Point of 0.5 to 0.7 GPa,0.2 to 0.4 Gpa for Wakaputa Point and 0.3 to 0.4 GPa for Oraka Pointintrusives; these averages and ranges reflect the Al contents of thehornblendes in the three suites (see above).

The pleonaste spinel-bearing reaction coronae noted in cumulateultramafic rocks from the Longwood Tops and Pahia Point provide anadditional constraint on the depth of emplacement. The reaction (seeabove) by which the spinel was formed can be used as an igneousgeobarometer (Frost, 1976; Kushiro and Yoder, 1966). Using thephase relationships described by Frost (1976), the compositions ofcoexisting olivine and pleonaste were used to estimate pressure overa temperature range of 500 to 800 °C. The full range of values obtainedfrom this approach is 0.45 to 0.7 GPa, which, with the estimatesobtained from the hornblende compositions of dioritic rocks, can beused to deduce emplacement at depths between 13 and 25 km.

4. Whole rock geochemistry

Representative whole rock analyses for the LIC are shown inTable 1 with major element variation shown in Figs. 9 and 10 andtrace element variation and patterns in Figs. 11–13. Whole rock dataare differentiated in the diagrams on the basis of geography, whichequates approximately with age, and rock type.

The Oraka, Wakaputa and Pahia suites all show similar trends ofiron enrichment on total alkali–FeO–MgO (AFM) plots (Fig. 9); this iscommonly termed a ‘tholeiitic’ trend. On silica variation diagramscumulate ultramafic rocks are clearly distinguished by relatively highFeO*, MgO, Ni and Cr contents and comparatively lower abundancesof TiO2, Al2O3, CaO, Sr, and V (Figs. 10, 11 and 14). Cumulate gabbroshave relatively low FeO* and V contents and elevated Al2O3, CaO andSr abundances.

N-MORB normalised extended element plots are shown for repre-sentative rocks as Fig. 12 and chondrite normalised rare earth element(REE) plots are shown as Fig. 13.

The primary striking feature of all analysed samples, includingmaficenclaves is that they have trace element patterns that are characteristicof subduction-related magmas (Fig. 12D). Furthermore, there is asimilarity between the N-MORB-normalised patterns of the ultramaficrocks (Fig. 12B) and various dykes (Fig. 12F andG). The latter are clearlycloser tomelts rather thancrystal-dominated cumulates of some type soif the ultramafic and gabbroic rocks were strictly adcumulates, thevariable degrees of compatibility of partitioning between constituentphases and melt should be reflected to some extent in their traceelement patterns. For example, in wehrlitic and gabbroic cumulatesclinopyroxene is the dominant host of the REE and typically has a“hump-shaped” abundance pattern reflecting the preferential partition-ing of the middle as opposed to light and heavy REE. There is howeverlittle indication in most of these bulk-rock trace element patterns ofdominantly cumulate-modulated controls, except possibly with thehornblendite dykes. These are discussed below.

Chondrite-normalised REE patterns are similar for most gabbroand diorite samples with enrichment of light over heavy REE andvarying degrees of Eu depletion. Cumulate gabbros show light REEenrichment but they also manifest positive Eu anomalies (Fig. 14).

Cs, Rb, Ba, Th, U, K and the light REE are enriched relative to N-MORBwhereas themiddle and heavy REE are at or belowN-MORB levels. Nb isdepleted relative to K and Th and Pb is enriched relative to Ce. In

ultramafic cumulates fromPahia Point the heavyREE andY are depletedrelative to middle REE (Figs. 12B and 13B). This feature is also observedin some of the dyke rocks (Figs. 12G, H, 13F and G). The Oraka gabbosand some of the dyke rocks are more depleted in large ion lithophileelements than is the case for other Longwood samples (Fig. 12E, F, andH) and in this respect they show some similarities with modern intra-oceanic, subduction-related basaltic rocks (e.g. Raoul Island basalts andbasaltic andesites from the Kermadec arc; see Fig. 12F).

Enclaves from theWakaputa and Pahia diorites havemajor and traceelement compositions that are broadly similar in composition to the hostdiorites except that they show relative enrichment in Na2O, P2O5, Zr, REEand Y (Figs. 10, 11 and 13); they have REE abundances that are generallyhigher thanall other rocks except theOrakagranite. Theyhave the lowestNi abundances and show themost negative Eu anomalies (Fig. 14). Theseaspects of the trace element characteristics of the enclaves suggest thatthey are among the most fractionated rocks in the LIC.

Dyke rocks have a range of compositional variation. Excluding thefelsic and hybrid components of the composite dyke at locality BB41,dykes have elevated Al2O3 abundances (16.55–19.09 wt.%) and SiO2

contents that vary from 41.5 to 53.5 wt.% with the lowest values beingobserved in hornblendite dykes. The hornblendite dykes also havedistinctly different convex-upwards, chondrite normalised REE pat-terns. The light REE are depleted relative to the middle REE and thepatterns are those expected of amphiboles. A dyke from locality BB30has relatively high K, Ba and Rb abundances, a feature in commonwithmafic phases from locality BB41 composite dyke.

The composite dyke at locality BB41 is hosted in quartz diorite(Fig. 5). It is ~1 m wide, trends north–south and, although it does notappear to have been sheared, it was emplaced long a left lateral faultthat has offset earlier generation mafic dykes by ~5 m. The contacts ofthe composite dyke with host diorite are sharp and marked by a felsicselvedge that anatomises into and net veins the mafic core. Xenolithsof the host are surrounded by this felsic phase, which also containssmall (1–2 cm) inclusions of mafic material. The mafic componentvaries texturally from plagioclase-phyric to aphyric. Fig. 15 comparesthe compositions of three components of the composite dykewith thehost diorite, other Pahia Point diorites and basaltic and basalticandesite dykes from the Pahia and Wakaputa segments of the coastalsection. BB41A is the felsic component, BB41B the aphyric maficcomponent, and BB41C is a sample of the plagioclase-phyric portion ofthe dyke. In geochemical terms the relationships are complex; thethree components do not consistently form a linear array and the hostdiorite and other diorites do not lie on a mixing line between felsicand most mafic component on all variation diagrams (Fig. 15). Themost mafic component has K, Ba and Rb abundances significantlyhigher than those observed in Pahia rocks with similar SiO2 contents.The felsic phase has lower REE abundances than other components ofthe composite dyke and significantly lower total REE and particularlyheavy REE than the host diorite.

5. Discussion

Longwood Igneous Complex gabbroic rocks and mafic dykes haveminor and trace element compositions that are characteristic ofprimitive subduction-related magmas. These include “arc” type traceelement patterns on normalised extended element plots and “tholeiitic”trends onAFMdiagrams. Rocks of theComplexhaveprimitive Sr andNdisotopic compositions with initial 87Sr/86Sr isotopic ratios varyingbetween0.70303and0.70394and initial εNd ranging from+2.6 to+7.2(Price et al., 2006).

Mineral compositions are also consistent with a subductionsetting. The intermediate chromian magnetite is most similar tospinel compositions observed in Alaskan layered ultramafic intrusions(Barnes and Roeder, 2001) and calcic plagioclase is typical of islandarc cumulates (Fig. 16); the very calcic compositions reflect the highmagmatic water contents expected in subduction related magmas

Fig. 12. N-MORB normalised extended element plots for the Longwood Igneous Complex with data points coded by geographic location and rock type. (A) Diorite and gabbro fromPahia Point. (B) Ultramafic rocks from Pahia Point. (C) Diorite and gabbro from Wakaputa Point/Ruahine Hill. (D) Mafic enclaves. (E) Oraka Point rocks. (F) Oraka Point gabbrocompared with average Raoul Island basalt/basaltic andesite (Ave Raoul B/BA). (G) Phases of composite dyke (BB41A-C) and adjacent mafic dyke (BB41D) from locality BB41, MulletBay. (H) Dyke rocks. PP12A and PP51003 are hornblendite dykes. Shaded field in diagrams B–G is Pahia Point diorite pattern from diagram A. Normalising values are from Sun andMcDonough (1989). Data for Raoul Island basalts and basaltic andesites from Smith et al. (2010).

12 R. Price et al. / Lithos 126 (2011) 1–21

(Arculus and Wills, 1980; Panjasawatwong et al., 1997; Sisson andGrove, 1993).

Distinctive features of the cumulate rocks of the LIC are theabundance of troctolitic cumulate rocks and intercumulous amphi-

bole. Troctolites have been documented elsewhere in cumulatesequences that have been interpreted to have originated in “arc”settings (e.g. Claeson and Meurer, 2004). They also occur as xenolithsor inclusions in some intra-oceanic arc volcanic rocks (e.g. the

Fig. 13. Chondrite-normalised rare earth element plots for the Longwood Igneous Complex with data points coded by geographic location and rock type. (A) Diorite and gabbro fromPahia Point. (B) Ultramafic rocks from Pahia Point. (C) Diorite and gabbro fromWakaputa Point/Ruahine Hill. (D) Mafic enclaves. (E) Oraka Point rocks. (F) Phases of composite dyke(BB41A-C) and adjacent mafic dyke (BB41D) from locality BB41, Mullet Bay. (G) Dyke rocks. PP12A and PP51003 are hornblendite dykes. Shaded field in diagrams B–G is Pahia Pointdiorite pattern from diagram A. Normalising values are from Sun and McDonough (1989).

13R. Price et al. / Lithos 126 (2011) 1–21

Kermadec arc; see below) and Bindeman and Bailey (1999) used traceelement geochemistry to argue that anorthite megacrysts in Kurile arcvolcanic rocks indicated that calcic plagioclase was the first or second(after olivine) phase to appear on the liquidus of high-Al basaltic

magmas. The conditions favouring the early appearance on the liquidusof plagioclase with olivine would appear to be a magma compositionthat is both tholeiitic and aluminous and a critical window of pressureandwater content. For example, in experiments to determine the phase

Fig. 14. Ni abundance (A) and Eu anomaly (B) versus MgO content for LongwoodIgneous Complex (LIC) with data points coded by geographic locality and rock type.(Eu/Eu*)n=Eu/[(Sm+Gd)/2] with abundances normalised to chondritic values of Sunand McDonough (1989).

14 R. Price et al. / Lithos 126 (2011) 1–21

relations of high alumina basalt, Foden and Green (1992) observed thatplagioclase and olivine were the liquidus phases at pressures below 0.5Gpa with water contents up to 5 wt.%. Fractional crystallisation ofanhydrous olivine and plagioclase assemblages from a parental meltwith 2–5 wt.% H2O could result in residual melts that would crystalliseamphibole. The implications are that most of the ultramafic cumulatesof the LIC were generated at relatively low pressures (i.e. within thecrust) from hydrous melts.

Mineral chemical and geological data indicate that the LIC wasemplaced at depths between 13 and 25 km and geochronology byvarious methods suggests emplacement ages ranging from 250 to140 Ma (Fig. 2; Price et al., 2006). Given the uncertainties relating toequilibriummineral assemblages and the cumulate nature of many ofthe rocks, pressure estimates obtained from hornblende compositionsshould be treated with some caution but the data suggest that PahiaPoint hornblendes may have crystallised at higher pressures than wasthe case for older components (Wakaputa and Oraka) of the LIC furtherto the east. This is consistent with the suggestion that the three suitesrepresent different arcs active at different times and representative ofdifferent stages of crustal evolution. Pahia Point hornblendes appear tohave crystallised deeper suggesting thicker crust or deeper mantlesources.

Oraka and Pahia Points are spatially separated by less than 20 kmbut the rocks exposed at each locality span ~100 million years of

magmatic activity (Fig. 2). Thus the LIC provides a time-integratedsample of development of the lower to middle crust in an intra-oceanicsubduction setting. The temporal evolution of the Longwood intrusivesis therefore a template for the petrologic evolution of sub-arc crustformedandmodifiedduringprolonged subduction and the geochemicalcomposition of the Complex provides a generally applicable insight intothe nature of material that is added to the crust during present day andpast intra-oceanic arc magmatism.

5.1. Batholiths and subduction-related volcanic rocks

There is an emerging view that intrusive rocks forming batholiths donot all havedirect volcanic analogues (e.g. Glazner et al., 2008). Kent et al.(2010)haveproposed that andesitic volcanic rocks erupt in preference tothe magmas that produced them and Bachmann and Bergantz (2004,2008) argue that granitoids in batholiths are in fact the preserved crystalmushes or “the leftovers” from felsic melt extraction. These hypothesesare derived from consideration of continental rhyolites, andesites andgranitoids but our observations on the LIC suggest that these types ofmodels are also directly applicable tomagmatic systems in intra-oceanicarcs. Ultramafic rocks, most of the gabbro and some of the diorite of theComplex are demonstrably cumulate in origin; they have cumulatetextures andAl2O3,MgO, CaO, Sr, Cr, andNi abundance patterns (Figs. 10,11, and 14), positive Eu anomalies (Figs. 13 and 14) and mineralcompositions (Figs. 8 and 16) that are all consistent with a cumulatederivation. Most hornblendites, even those occurring as dykes, appear torepresent crystal residues from which melt was extracted, possibly bypressure filtering although the comb-textured varieties (e.g. Fig. 5) mayhave crystallised from water-saturated, residual melt pockets.

Few of the Longwood rocks appear to have direct analogues amongmodern intra-oceanic volcanic rocks. The exception is Oraka Pointgabbro, which has trace element abundances and normalised traceelement patterns that resemble those observed in intra-oceanic basaltsand basaltic andesites from the Kermadec intra-oceanic arc (Fig. 12F).There are however significant mineralogical differences. Kermadecbasalts have phenocryst assemblages that include olivine, plagioclase,orthopyroxene, and clinopyroxene but the Oraka gabbro containshornblende and biotite. Hornblende is not observed in Kermadec lavasbut it is found in coarse grained cumulate blocks (Ewart et al., 1977;Smith et al., 2010) in felsic tephras of Raoul Island. Texturally andmineralogically the Raoul xenolith suite of troctolite and hornblendegabbro is similar to the Longwood cumulate rocks. Davidson et al.(2007) have postulated that cryptic amphibole fractionation may be afactor determining geochemical variation in volcanic rocks that containlittle or noamphibole and they suggest that the crust beneathmany arcsmust contain amphibole-bearing cumulates. The presence of amphibolecumulates in the LIC provides some confirmatory evidence in support oftheDavidsonet al. “amphibole sponge” concept.Weconclude thatmanyof the rocks of the LIC represent cumulates and crystal mushesassembled in the lower or middle crust beneath the Permo-Cretaceousintra-oceanic arc of New Zealand. The only directly comparable rocksfound inmodern intra-oceanic arcs are coarse grained xenoliths carriedup from depth in lavas. There are no direct compositional andmineralogical analogues for Longwood rocks among modern arceruptives and this is very probably the case for other primitive I-typeintrusives elsewhere.

5.2. Magmatic processes: fractional crystallisation and crystal accumulation

In qualitative terms there is evidence that fractional crystallisationwasa significantprocess controllinggeochemical variation in Longwoodrocks. Many show textural and geochemical evidence that they formedas crystal cumulates. On silica variation diagrams and in the range 48 to60 wt.% SiO2 analysed samples define approximately curvilinear trendson AFM (Fig. 9) and Harker variation (Figs. 10 and 11) diagrams. FeO,MgO, CaO, Sr, Sc, Cr and V abundances decrease andNa2O, K2O, P2O5, Ba,

Fig. 15. Selected Harker variation diagrams for components of composite dyke at locality BB41, Mullet Bay. “A” is the felsic phase, “B” and “C”mafic components. Also shown are datafor the host diorite, Pahia Point diorites and Pahia and Wakaputa dyke rocks. The line connects the felsic and most mafic components of the dyke.

15R. Price et al. / Lithos 126 (2011) 1–21

Rb, Zr and REE contents increase with increasing wt.% SiO2. Systematicchanges in whole rock geochemistry are reflected in mineral compo-sitions (e.g. Fig. 7). These trends can be explained qualitatively in termsof fractional crystallisation involving combinations of olivine, plagio-clase, amphibole, pyroxene and spinel and they are consistent with themake-up of cumulate rocks found throughout the Complex.

Mathematical models for fractional crystallisation, with or withoutcrustal assimilation, assume rock compositions represent points on liquidlines of descent (e.g. Davidson et al., 2007; Ewart and Hawkesworth,1987; Graham et al., 1995; Grove et al., 1982; MacPherson et al., 2010;Price et al., 1999). The universal validity of this assumption isquestionable. For example, Arculus et al. (1995) compared glass andcrystal-bearing lavas for the Izu-Bonin-Mariana arcs and were able toshow that the former correspond unequivocally to liquids and coincidemuch more closely with experimentally-determined liquid lines ofdescent than is the case for the rocks.

For rocks of the LIC andother primitive I-type batholiths, quantitativemathematical modelling of fractional crystallisation or assimilationcrystal fractionation processes is, for a number of reasons, problematic.

Firstly, most rocks represent cumulates or crystal mushes rather thanmelts and theassumptionsof consanguinity and liquid lineof descent arenot appropriate. Secondly, there is evidence in most rocks that originalliquidus phases such as olivine and pyroxene have been replaced byamphibole and biotite as melts become progressively more water rich.Consequently, it is difficult to determine the liquidus phases andappropriate compositions to be used for these in the models. Thirdly,although the major element compositions can be modelled using majorphases, in intermediate and felsic, silica-saturated rocks, trace elementbehaviour is strongly influenced by accessory phases such as apatite,zircon, titanite, monazite and allanite (e.g. Gromet and Silver, 1983;Hoskin et al., 2000;Michael, 1988). Partition coefficients for these phasesare not well established and are dependent on melt composition andstructure and their modal abundances are generally poorly constrained.Finally, there is evidence in some rocks (e.g. chlorite–epidote–titanitealteration and lower pressure amphibole compositions; unusualenrichments in K2O, Ba and Rb) that subsolidus alteration has modifiedrockandmineral compositions. For these reasons andafter anexhaustiveevaluation of the efficacy of a variety of mathematical models we have

Fig. 16. Compositions of coexisting olivine and plagioclase from the Longwood IgneousComplex. Fields are from Beard (1986).

16 R. Price et al. / Lithos 126 (2011) 1–21

concluded that they cannot be applied to quantitatively verify fractionalcrystallisation as an explanation for geochemical variation in the LIC. Thefield relationships, geochemical variation, and textures observed in theLongwood intrusives suggest a complex interplay ofmagmatic processesthat included fractional crystallisation and assimilation fractionalcrystallisation, crustal anatexis and magma mingling and localisedmixing. A further implication of these events is that lower–middle crustwas hot for a long time.

5.3. Magmatic processes: crustal recycling

Mortimer et al. (1999a) and Price et al. (2006) noted that the Srand Nd isotopic compositions changed systematically from east towest across the LIC. In Fig. 2B this trend is illustrated using 87Sr/86Srratios and is interpreted to indicate that older crust (earlier portions ofthe arc) was recycled during magmatic processes that formed theintrusive rocks of the younger arc. Inheritance in zircons dated bySHRIMP ion probe confirms this interpretation (Price et al., 2006). Thedata support the hypothesis that as intra-oceanic subduction systemsdevelop and mature, crustal recycling becomes increasingly importantin the processes by which magmas are generated and evolve. Crustalrecycling has been argued to be amajor factor in the generation of largevolume felsic eruptions in the intra-oceanic arcs (Brophy, 2008;Handley et al., 2008; Smith et al., 2003, 2006; Tamura and Tatsumi,2002; Tamura et al., 2009). Smith et al. (2010) have also suggested thatcrustal assimilation is significant in the evolution of basaltic andesitesand andesites in intra-oceanic subduction settings. It is possible thatcrustal recycling is as important in long-lived intra-oceanic arcs as it is incontinental tectonic settings (e.g. Davidson et al., 2005; Dungan et al.,2001; Price et al., 2005) but the primitive nature of the lower crustmeans that the impact on rock compositions may not be as obvious(Smith et al., 2010).

By analogy with modern volcanic rocks (e.g. Davidson et al., 2005;Dungan et al., 2001; Graham et al., 1995; Price et al., 1999) crustalrecycling in the LIC is likely to have involved assimilation withfractional crystallisation (De Paolo, 1981) but the possibility that felsicmagmas were generated by direct melting of crust (e.g. Smith et al.,2003, 2006) cannot be precluded. The granitic rocks at Oraka Pointand the felsic component of the composite dyke at locality BB41 arerelatively depleted in REE (Fig. 12E and F) and Y (Fig. 10E) and at thelatter location enclaves of host diorite occur within the felsic phase ofthe dyke (Fig. 5). These features might indicate that felsic rocks were

formed during partial melting of diorite with REE being held back in arefractory phase.

5.4. Magmatic processes: magma mingling and composite dykes

Textures and structures observed in coastal exposures of the LICsuggest thatmafic and relatively felsic magmas coexisted andmingledor mixed as the Complex was assembled. Composite dykes, theoccurrence of hybrid zones along gabbro/diorite contacts and thedevelopment of enclave-rich zones near these contacts are featuresindicating that mafic and more felsic components were in contactwhile both phases were capable of deforming plastically. Magmamingling structures and textures of this type occur elsewhere in theMedian Batholith (e.g. Cook, 1988; Turnbull et al., 2010) and in otherI- and A-type granitoids around the world where they have beenstudied in detail to develop models for magma chamber developmentin I-type plutons (e.g. Miller and Miller, 2002; Wiebe, 1974, 1993;Wiebe and Collins, 1998). Generally these models explain mixed(hybrid) rocks, interstratified mafic and felsic rock types and maficenclave swarms as arising from the injection of mafic magma intofelsic magma chambers (e.g. Collins et al., 2006; Turnbull et al., 2010;Wiebe, 1974; Wiebe and Collins, 1998). For these models, geochem-ical variation is explained by magma mixing with fractional crystal-lisation assigned an importance that varies between different models(e.g. Collins et al., 2006; Metcalf et al., 1995; Wiebe, 1993).

Clearly magmamixing textures and structures are present in the LICas they are elsewhere in the Median Batholith, but it is debatablewhether magma mixing has been a significant process controllingwhole rock geochemical variation within each of the Longwoodgeochronological suites. The influence of magma mixing and minglingis certainly a contested issue among those studying I-type granitic rocksof the Bega Batholith in south-eastern Australia (Collins et al., 2006compared with Chappell, 1996) and Coleman et al. (2004) concludedthat the plutonsmaking up the Tuolumne suite in California did not co-exist as liquid-rich magmas so the chemical evolution of the suite couldnot have involved fractional crystallisation or magma mixing.

We have shown that, although fractional crystallisation, crustalassimilation and localised crustal anatexis were involved in thegeneration and evolution of Longwood arc magmas, most of theintrusive rocks represent cumulates and crystal mushes left over fromthese processes. The mingling textures and structures observed withinthe LIC represent interaction of crystal-rich magmas in which the meltcontent is likely to have been in some cases b10%.

Composite dykes within the Pahia Point–Mullet Bay segment ofthe coastal section show evidence for magma mingling and for thedevelopment of hybrid magmas by mixing. Their geochemistry canalso be used to test whether these localised mingling processes arereflected in the geochemical variation defined by the Pahia–Wakaputasuites collectively. If the various phases of a composite dyke havecompositions that are related by simple mixing then they shoulddefine linear arrays on two component variation diagrams. Similarly,if diorites are derived by chemical and/or mechanical mixing betweenmagmas represented by themafic and felsic components of compositedykes then they should lie on the samemixing line. For the compositedyke at Pahia Point/Mullet Bay locality BB41 the variation is notstraightforward. The three analysed phases of the composite dykedefine an approximately linear array on MgO and Zr versus SiO2

diagrams but on the K2O, Ba, Rb and La versus SiO2 variation diagramsthey are not aligned (Fig. 15). On MgO, K2O and Rb versus SiO2

diagrams, Pahia diorites have compositions that plot near amixing linebetween the felsic andmostmafic phases of the dyke but this is not thecase for Ba, Zr and La. The hybridisation that occurred during theformation of the composite dykes involved more than simple mixingand there is no clear link between the processes taking place in thecomposite dyke and those controlling variation at the pluton orbatholith scale within the Complex.

17R. Price et al. / Lithos 126 (2011) 1–21

Composite dykes are generally interpreted to have formed whenbasaltic magma intrudes partially crystallised more felsic magmawitha consequent thermal and mechanical interaction. Wiebe and Ulrich(1997) interpret composite dykes in the Goldsbouro Granite of centralMaine to have formed when basalt dykes ruptured a chamber contain-ing liquid granitic magma. They argued that the felsic magma flowedinto these fractures mingling with and breaking up the basaltic magmawithin the dykes. This explanation does not appear to be applicable totheBB41 composite dyke. In the latter case contactswith thehost dioriteare sharp and mafic “pillows” or blobs are hosted in the felsic phase ofthedyke. Enclavesof thehostdiorite are present in the felsic phaseof thedyke (Fig. 5) and the host and the felsic component of the dyke havedistinctly different compositions. A possible explanation is that the hostdioritewas a crystalmushwhen the dykewas emplaced. Heating by thebasaltic dyke caused thehost diorite adjacent to thedyke tomelt and it isthis second generation felsic melt that is represented by the felsic phaseof the composite dyke. By the same reasoning, the mafic componentmight be expected to chill against the felsic component, therebypreserving mafic pillows. We conclude that hybridisation in thecomposite dykes is localised and does not relate to the larger scaleprocesses bywhich dioritic magmaswere generatedwithin the LIC. Thedykes do however demonstrate that the emplacement of mafic magmacan lead to localised crustal anatexis even in the waning stages of arcdevelopment.

5.5. Mafic enclaves

Microgranularmafic enclaves are a featureofmany I-typegranitoids.Theyhavebeen interpreted to representblobsofmaficmagmathat havebeen entrained into a more felsic host (e.g. Barbarin and Didier, 1992;Blundy and Sparks, 1992; Vernon, 1983; Vernon et al., 1988), refractoryrestite (Chappell and White, 1991) or dispersed synplutonic orcomposite dykes (Pitcher, 1991).

The mafic enclaves found in dioritic rocks of the LIC have featuresin common with I-type mafic enclaves found elsewhere. They aregenerally rounded or blob-like, they occur in swarms or clusters andthey have mineralogy similar to their hosts but different grain-sizeand different modal proportions. They contain abundant needle-likeapatite and they do not show cumulate textures. All analysed enclaveshave similar geochemical characteristics including a narrow compo-sitional range (SiO2 is between 53.5 and 57.6 wt.%) and in manyrespects they closely resemble gabbroic diorites and diorites of theComplex; they have TiO2, Al2O3, FeO*, MgO, CaO, and K2O abundancesthat are similar to those observed in gabbroic diorites and dioriteswith similar SiO2 contents (Fig. 10). There are however aspects oftheir geochemical variation that are unique. They show relativelyelevated abundances of Na2O, P2O5, Zr, Y, the REE and, to a lesserextent Rb and they are relatively depleted in Sr and Ni (Figs. 10, 11and 14). Their REE patterns are different from those of all otherLongwood rocks. Total REE abundances are higher and they showsome of the most well developed negative Eu anomalies. They havedistinctive flat middle to light REE patterns. In a study of maficenclaves of the Adamello Massif in northern Italy, Blundy and Sparks(1992) recorded very similar geochemical characteristics. Theyinterpreted the Adamello mafic enclaves to have formed after thehost pluton had consolidated but was still hot. Mafic intrusionsemplaced at this stage heated and remobilised the host granitoid andenclaves formed through interaction of the remobilised felsic host andintruding mafic magma. Diffusion between host and enclave signif-icantly modified aspects of the trace element composition of thelatter.

We interpret the geochemistry of the Longwood enclaves in terms ofa three stage process similar to that proposed by Blundy and Sparks(1992) for the Adamello inclusions. The first stage involved fractionalcrystallisationof awet basalticmagma. Fractionationof calcic plagioclasedepleted the magma in Sr and Eu and fractionation of olivine and

pyroxene lowered Ni abundances. The second stage of the process wasthe emplacement of this magma into a relatively felsic crystal mush.Heating and remobilisation of thismaterial resulted inmingling ofmaficmagma into the host to produce enclaves. In the third stage of theprocess enclaveswere chemicallymodified by diffusion from the host ofH2O, Na, P, Ba, REE, and Rb.

5.6. The composition of sub-arc crust: implications for crustal growthand evolution

It has been long recognised that continental crust has on average anandesitic composition (Rudnick and Gao, 2005) and consequently anumber of hypotheses developed to explain the growth and evolution ofcontinental crust have focused onprocesses that take place in subductionsystems (e.g. Arculus, 1999; Drummond and Defant, 1990; Eichelberger,1978; Ellam and Hawkesworth, 1988; Kelemen, 1995; Rudnick, 1995;Taylor, 1967, 1977). Alternative models include recycling of crustthroughmantle plumes and the formation of continental crust in oceanicplateaux (e.g. Albarède, 1998). One of the challenges formodels that seekto generate new continental crust in subduction systems is that whereasthe mantle-derived flux may be basalt, the net average crustalcomposition of intra-oceanic arcs is andesite (Davidson and Arculus,2006; Kelemen et al., 2005; Kodaira et al., 2007). It is also highly likelythat arc volcanic rocks do not directly represent themagmas fromwhichthey are derived (Kent et al., 2010) so that the volcanic sample iscompositionally biased; only magmas with a specific range of physicalproperties (crystallinity, density, rheology, viscosity)will be able to leavethe crust and erupt at the surface (Marsh, 1981).

The datawehave obtained for the LIC can beused to provide anotherperspective on this problem. It can be unequivocally demonstrated thattheLongwoodrocks originatedover a longperiodof time (100 Ma) in anintra-oceanic subduction system where they were emplaced into themiddle to lower crust andarenowamajor componentof the continentalcrust of Zealandia. Collectively, the Longwood rocks record theevolutionof a segment of crust over a 100 million year time span. The averagecomposition of the LIC should therefore provide an estimate of thecomposition of crust that can be developed over a prolonged geologicalperiod above an intra-oceanic subduction system. The temporalevolution of the arc is also recorded in the changes that occurred fromeast towest. The olderOraka gabbros to the east are the Longwood rocksmost similar to depleted oceanic basalts and the youngerWakaputa andPahia diorites further west provide the clearest evidence for crustalrecycling.

Table 2 shows average compositions for the principal rock types ofthe LIC along with crustal compositions calculated using theseaverages and estimates of their relative volumes. Three differentestimates are shown. The first two are for the rocks exposed betweenPahia and Oraka Points and the third [(c) in Table 2], which is limitedin terms of elements for which data are available, is based on relativevolumes estimated from the interpretive map of the inland LongwoodRange and trondhjemite data published by Mortimer et al. (1999a).The first estimate [(a) in Table 2], based on the relative abundance ofthe different rock types exposed along the coastline is a basalticandesite composition whereas the second [(b) in Table 2] is andesitic.The second (b) and third (c) calculations give similar results for mosttrace elements.

Also shown in Table 2 is the estimated bulk continental crustalcomposition of Rudnick and Gao (2005) and the (a) and (b) Longwoodaverages are compared graphically with this bulk crustal estimate inFig. 17. Although andesitic in average composition, Longwood crust hasa lower SiO2 and higher Mg and Al2O3 contents than bulk continentalcrust but for many elements (Ti, Mg, Ca, Na, K, P, Rb, Ba, Nd, Sm, Zr, Hf,Eu, V, Ni, Zn andGa) the abundances are comparable. Th, U, Pb, Sr and Crare enriched and Nb, the LREE, HREE, Y and Sc depleted in Longwoodcrust relative to bulk continental crust.

Table 2Estimates of average composition of Longwood Igneous Complex crust.

Ave Ave Ave OR7B P2837 Ruapehu Raoul Is. Crust

U/M Gabbro Diorite Granite Trodjh (a) (b) (c) andesite andesite

SiO2 41.00 47.76 60.81 76.02 70.59 54.92 56.89 57.20 58.73 56.52 60.60TiO2 0.27 0.85 0.83 0.21 0.19 0.80 0.79 0.70 0.71 0.90 0.72Al2O3 8.24 18.88 17.40 12.41 16.17 17.76 17.42 17.78 16.68 16.41 15.90Fe2O3 10.67 4.88 2.09 0.57 1.45 3.49 3.12 3.18 7.20 11.10 6.71MnO 0.18 0.12 0.11 0.03 0.05 0.11 0.11 0.10 0.12 0.20 0.10MgO 26.35 8.64 2.32 0.42 0.31 5.62 4.83 4.65 4.34 3.25 4.66CaO 5.34 10.49 4.97 1.01 3.08 7.49 6.62 6.98 7.04 8.94 6.41Na2O 0.71 2.52 4.47 2.92 5.37 3.39 3.64 3.81 3.28 2.33 3.07K2O 0.18 0.42 2.11 5.11 1.04 1.41 1.68 1.14 1.52 0.40 1.81P2O5 0.06 0.13 0.22 0.01 0.04 0.16 0.17 0.14 0.13 0.12 0.13Mg# 83.03 77.79 68.72 59.34 29.75 76.13 75.42 74.37 54.41 36.72 57.90Ba 78 150 429 561 321 294 335 284 346 128 456Rb 4 9 80 162 16 48 59 35 55 7 49Sr 453 806 540 109 561 648 603 661 255 176 320Pb 3.36 5.81 13.04 20.69 5 9.75 10.85 8 10.52 2.59 5.6Th 0.60 1.13 5.61 10.28 3.58 4.27 4.69 0.28 0.93U 0.22 0.34 2.32 2.25 1.32 1.60 1.39 0.17 0.89Zr 21 51 213 149 62 128 150 110 115 41 132Hf 0.43 1.52 3.94 5.66 2.79 3.15 3.6 1.6 3.7Nb 0.6 1.5 5.0 1.3 5 3.1 3.5 3 4.7 0.8 8Y 1.89 8.41 17.96 12.00 4 12.78 14.03 11 21.3 26.6 19La 3.09 5.75 15.68 9.42 8 10.34 11.67 10 12.35 3.07 20Ce 4.92 12.20 38.29 19.54 32 24.14 27.62 26 26.89 8.62 43Pr 0.76 2.66 4.23 2.32 3.32 3.51 3.49 1.50 4.90Nd 3.10 11.63 17.86 7.96 14.13 14.85 13.91 8.28 20Sm 0.72 2.98 3.68 1.46 3.19 3.25 3.21 2.88 3.9Eu 0.28 0.99 0.99 0.33 0.95 0.94 0.92 1.07 1.1Gd 0.62 2.95 3.33 1.24 3.01 3.02 3.40 3.75 3.7Tb 0.08 0.51 0.56 0.18 0.51 0.51 0.54 0.68 0.6Dy 0.51 2.15 2.56 1.06 2.26 2.29 3.24 4.69 3.6Ho 0.08 0.58 0.66 0.22 0.59 0.59 0.70 1.04 0.77Er 0.26 1.08 1.37 0.70 1.18 1.21 1.94 2.79 2.10Tm 0.03 0.18 0.21 0.12 0.19 0.19 0.30 0.47 0.28Yb 0.22 1.14 1.52 0.90 1.29 1.33 2.00 3.28 1.90Lu 0.03 0.17 0.23 0.16 0.20 0.20 0.31 0.47 0.30Sc 13.3 21.2 10.0 2.3 6 15.2 13.4 14 23.2 32.7 22V 100 203 112 15 22 152 137 133 178 286 138Cr 2362 408 17 9 32 238 196 191 86 14 135Ni 924 122 8 5 5 75 65 57 30 15 59Cu 34 96 24 9 101 59 48 72 46 82 27Zn 122 74 77 24 48 74 74 70 67 73 72Ga 9 17 22 13 18 19 20 19 19 14 16(La/Yb)n 10.25 3.62 7.40 7.51 5.77 6.31 4.43 0.67 7.55(Eu/Eu*)n 1.248 1.013 0.847 0.735 0.923 0.899 0.842 0.994 0.872

Columns 1–3 are average compositions for ultramfic/troctolite (U/M), gabbro and diorite respectively.Column 4 is a granite composition and 5 is a trondhjemite (Trodhj) composition from Mortimer et al. (1999b).

(a) Crustal composition estimated from exposure at coastline: 1% ultramafic/troctolite, 49% gabbro, 45% diorite, 6% granite.(b) Crustal composition estimated from areas exposed in southern Longwoods between Pahia and Oraka Pts: 2% ultramafic/troctolite, 34% gabbro, 58% diorite, 6% granite.(c) Crustal composition estimated from areas shown in interpretive map of Mortimer et al. (1999b): 44% gabbro, 35% diorite, 21% trondhjemite.

Ruapehu andesite average is from data in Graham and Hackett (1987) and R.C. Price (unpublished data).Raoul Island andesite average is from Smith et al. (2010).Crust is composition of total crust from Rudnick and Gao (2005).

18 R. Price et al. / Lithos 126 (2011) 1–21

Comparisons are also made in Table 2 and Fig. 17 with continentaland oceanic andesite compositions from the present day Kermadec-New Zealand arc. Longwood crustal compositions are similar in manyrespects to the average composition of andesite from Ruapehuvolcano, which is located on continental crust at the southern endof the Taupo Volcanic Zone in New Zealand's North Island (Fig. 17Aand Table 2). The average andesite composition from Raoul volcano,from the intra-oceanic portion of the subduction system is distinctlydifferent, showing very strong relative depletions in large ionlithophile elements and Nb and Zr. Although they have SiO2

abundances similar to that calculated for the Longwood crust, theRuapehu and Raoul andesites are significantly more fractionated, withmuch lower Mg#s and lower abundances of Ni and Cr. Given that theLongwood crust includes cumulate rocks, this is not unexpected.Compared with both Raoul and Ruapehu andesites the Longwood

crustal averages show relative depletion of heavy relative to light REEand this could reflect involvement of garnet in the generation ofLongwood primary magmas (c.f. Kelemen et al., 2005).

The estimates of Longwood crustal composition demonstrate thatalthough basalt may dominate magmas reaching the surface in anintra-oceanic arc, over time the deeper crust is assembled fromcumulates, crystal mushes, crustal melts and fractionated magmasthat collectively approach andesitic compositions not dissimilar tobulk continental crust. The similarity between Longwood crustalcompositions and continental andesite and the contrasts with oceanicandesite would appear to reinforce the view that over a period of100 million years, the crust within an intra-oceanic arc will, throughrepeated magmatic contributions from the slab and sub-arc mantleand recycling of the maturing lower crust, develop and evolvetowards a bulk continental crustal composition.

Fig. 17. Bulk crust normalised extended element plots for estimated averagecompositions of Longwood Igneous Complex (LIC) crust and comparisons with averageRuapehu (Graham and Hackett, 1987; R.C. Price unpublished data) and Raoul Islandandesites (Smith et al., 2010). Calculation of crustal averages is explained in the textand in Table 2. (A) is estimated from exposure at coastline: 1% ultramafic/troctolite, 49%gabbro, 45% diorite, 6% granite. (B) is estimated from areas exposed in the southern LICbetween Pahia and Oraka Pts: 2% ultramafic/troctolite, 34% gabbro, 58% diorite, 6%granite. Normalising values are from Rudnick and Gao (2005). Shaded area is range formiddle and lower crust (Rudnick and Gao, 2005).

19R. Price et al. / Lithos 126 (2011) 1–21

6. Conclusions

The LIC is made up of I-type plutonic suites that preserve a sectionthrough the lower or middle crust that evolved in the southwest Pacificover a 100 million year period (late Palaeozoic toMesozoic) beneath anintra-oceanic, subduction-related magmatic arc. The average composi-tion of this crustal segment approximates andesite, demonstrating thatalthough volcanic rocks in intra-oceanic arcs are dominantly basalt orbasaltic andesite, over a long enough period of time, the crust that isassembled beneath the arc has a more evolved composition.

The plutons making up the LIC do not have direct analogues amongmodern intra-oceanic arc volcanic rocks. The Longwood rocks indicatethat sub-arc crust is largely composed of cumulates and crystal mushesrepresenting the leftovers from fractional crystallisation, crustalanatexis and assimilation and mixing and mingling processes thatgenerated magmas erupted at the surface. Volcanic rocks are the endproducts whereas the plutonic rocks preserve aspects of the processesthat for volcanic rocks can only be inferred through the interpretation ofwhole rock andmineral compositions. For example, the LIC confirms theimportance of amphibole fractionation and amphibole accumulation inthe evolution of arc volcanic rocks and sub-arc crust respectively despitethe fact that amphibole is rare as a liquidus phase in most intra-oceanicarc volcanic rocks.

Field and petrographic observations, isotopic data and zircongeochronology indicate that, as the Longwood crust evolved, crustalrecycling began to take place; early formed segments of sub-arc crustwere heated and began to melt, transient magmas were modified bycrustal assimilation that accompanied fractional crystallisation. Thisobservation adds weight to the proposal that crustal anatexis isimplicated in the generation of large volumes of felsic magma eruptedin some mature intra-oceanic arcs; if the arc has a long enough life thesub-arc crust will mature compositionally and thermally to a pointwhere relatively newly created crust will begin to melt and be recycled.

The LIC manifests physical evidence for magma mingling andlocalised magma mixing but these processes do not scale up to explainthe overall patterns of geochemical variationwithin the Complex. Theseare better explained in terms of crystal accumulation and magmaticevolution involving fractional crystallisation, with and without crustalassimilation, and crustal anatexis.

The geochemistry and petrology of mafic enclaves is consistentwith models such as those developed by Blundy and Sparks (1992). Inthe Longwood intrusives, mafic enclaves were formed when fraction-ated basaltic magma was emplaced into and thermally remobilisedhot but solid more felsic magma causing physical disaggregation ofthe former. The geochemistry of the enclaves is a reflection of severalprocesses: extensive fractional crystallisation of mafic magma, somedegree of mixing of crystals and melt between felsic host and themafic enclave magma, and sub-solidus diffusive modification.

As is the case for enclaves, composite dykes show physical evidencefor magma mingling and localised hybridisation but the geochemicaland petrological relationships between the various components ofcomposite dykes and their hosts do not indicate that the processestaking place during the formation of composite dykes are thosecontrolling the geochemical variation observed for the whole Complex.

Supplementarymaterials related to this article can be found onlineat doi:10.1016/j.lithos.2011.04.006.

Acknowledgements

This research has been funded by through an Australian ResearchCouncil research grant to RCP and RA administered by La TrobeUniversity. The technical support of Stephen Eggins, Roland Maas, andGordon Holm was crucial for the successful completion of this projectand Stephen Eggins is also thanked for his contribution to field workundertaken by CS. An earlier draft of the paper benefited considerablyfrom a review by John Gamble. John Foden and Philip Leat are thankedfor their constructive and insightful reviews. The support of LaTrobe, theAustralian National and Otago Universities is gratefully acknowledged.

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