On the onset and evolution of the Ross-orogeny magmatism in North Victoria Land — Antarctica

Chemical Geology 240 (2007) 103 – 128 www.elsevier.com/locate/chemgeo

On the onset and evolution of the Ross-orogeny magmatism in North Victoria Land — Antarctica F. Giacomini a,⁎, M. Tiepolo b , L. Dallai c , C. Ghezzo a b

a Dipartimento di Scienze della Terra, Università di Siena, via Laterina 8, 53100, Siena, Italy CNR–Istituto di Geoscienze e Georisorse (IGG) Unità di Pavia, via Ferrata 1, 27100, Pavia, Italy c CNR–Istituto di Geoscienze e Georisorse (IGG) Unità di Pisa, Via Moruzzi 1, 56124 Pisa, Italy

Received 22 December 2005; received in revised form 10 January 2007; accepted 3 February 2007 Editor: S.L. Goldstein

Abstract An extensive geochemical (major-, trace-element and oxygen isotope) and geochronological (U/Pb geochronology on zircon) characterisation of Ol-bearing ultramafic–mafic cumulates and gabbro-diorites recently discovered in the southwestern region of the Wilson Terrane (North Victoria Land — Antarctica) was undertaken in order to constrain source characteristics, magma evolution and emplacement history in this sector of the Cambro–Ordovician Ross Orogeny. Textural and chemical data indicate disequilibria among mineral phases of these Ol-bearing cumulate rocks. Corroded Cr-rich clinopyroxene with high Mg# (0.82) and Ca-rich plagioclase likely represent xenocrysts from an old magmatic system, and are not in equilibrium with the host cumulate assemblage (Ol + Opx + Amp + Bt + Pl). The calculated liquid in equilibrium with the xenocrysts (“melt 1”) exhibits exceptionally high La/Yb ratios and Th–U concentrations, suggesting a strong sediment influx in the mantle source, possibly from the subducted slab. Because of the extensive fractional crystallisation (Fo70; Mg# = 76) and crustal contamination (e.g., δ18OOpx = 7.13–7.47‰) the trace element composition of the liquid in equilibrium with the cumulate assemblage (“melt 2”) does not reveal the nature of the mantle source not the differentiation processes. The gabbroic parental liquid for the main cumulate assemblage was not produced by assimilation and fractional crystallisation (AFC) of the equilibrium “melt 1” calculated from the xenocrystic paragenesis. Rather, it likely represents a new pulse of magma originated from a different mantle source, which then evolved through AFC into gabbro-diorites. Zircon U–Pb dating of one gabbro-diorite yields a crystallisation age of 489 ± 4 Ma. Zircon in the Ol-bearing cumulates yielded two ages populations at 521 ± 2 Ma and 502 ± 3 Ma, which we interpret to represent the actual age difference between the two magmatic systems. This age difference suggests that poorly differentiated melts with adakite-type signature intruded the crust prior to the generation of the large volumes of gabbroic and dioritic magmas. These early pulses represent up to now the oldest proof of subduction related mantle melts in North Victoria Land and predate the diffuse igneous activity dominated by intermediate- to felsic products. © 2007 Elsevier B.V. All rights reserved. Keywords: Zircon geochronology; Ultramafic-cumulates; Antarctica; Adakite-type mantle melts; Laser ablation; Equilibrium

⁎ Corresponding author. Tel.: +39 0577 233 802; fax: +39 0577 233 938. E-mail address: [email protected] (F. Giacomini). 0009-2541/$ - see front matter © 2007 Elsevier B.V. All rights reserved. doi:10.1016/j.chemgeo.2007.02.005

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1. Introduction Chemical changes in primary magma compositions throughout the evolution of an active margin are thought to result from variations in collision dynamics (age of the slab, subduction rate, plate direction or obliquity; e.g., Vanek et al., 1994; Macdonald et al., 2000). It is therefore important to constrain the chemical evolution of magmatism in orogenic areas, particularly of primary magmas occurring in response of subduction, and characterise the nature of their mantle sources in order to understand the original geometry of the active margin. Unfortunately, mafic magmatism occurring in active continental margins is often characterised by significant crustal contamination at depth, so that the original mantle signature is obscured if not lost completely. Although mafic and ultramafic cumulates represent the early differentiation products of the most primitive subduction related magmas (De Bari, 1997), their occurrence is extremely limited. In the fortunate cases where these rocks are available, the composition of mineral phases, their zoning and the presence of disequilibria among phases may provide unique information on the petrogenetic processes that occurred during the orogenic event (cf. Blundy and Shimizu, 1991; Tiepolo et al., 2002; Tiepolo and Tribuzio, 2005). As the bulk rock composition only provides an average picture of the system, and all information on local disequilibria is lost, the micro-analytical investigation of the major, trace element and isotope composition of the minerals is the only approach that allows deciphering all the petrogenetic records within these rocks. In this work chemical disequilibria within mafic and ultramafic cumulates were used to better constrain timing and affinity of primary magmas at the onset of the Ross Orogen in Antarctica, i.e. the Palaeozoic subduction of the palaeo-Pacific plate under Gondwana. Although the main frame of the Ross orogenic event – extending in the Transantarctic Mountains from North Victoria Land to the Pensacola Mountains – is nowadays delineated, the chronology of the tectonometamorphic evolution and of the magmatic activity is still debated (cf. Stump et al., 2006). The main emplacement phase of the “Cordilleran type” igneous rocks, known as “Granite Harbour Intrusives” (Gunn and Warren, 1962; Borg et al., 1987) is constrained along the different sectors of the Orogen between about 510 and 480 Ma (Allibone and Wysoczanski, 2002; Stump et al., 2002; Bomparola et al., 2007, and references therein). In the Central Transantarctic Mountains and in South Victoria Land

however, granitic to tonalitic rocks, locally with gneissic texture, yielded ages as old as ∼ 550 Ma (Goodge et al., 1993; Cottle and Cooper, 2006a; Stump et al., 2006, and references therein). These old subalkaline intrusions are almost coeval with the plutonic rocks of the Koettlitz Alkaline Province a large area of south Victoria Land, which is proposed as having formed in the time span 551–530 Ma (Rowell et al., 1993; Encarnación and Grunow, 1996; Mellish et al., 2002; Cottle and Cooper, 2006b; Foden et al., 2006; Veevers et al., 2006). Multiple deformation events and transpressional/transtensional phases during the development of the orogen are invoked to explain the occurrence of both alkaline and calc-alkaline igneous products in the same region (ibid.). Up to now, the presence of magmatic rocks older than 510 Ma has not been unequivocally demonstrated in North Victoria Land: in such a complex framework, the study of the age and chemical affinity of the intrusive products and particularly of the least differentiated igneous rocks is essential to evidence and explain the possible diachronic onset of the magmatic activity along the belt and thus to reconstruct the geometry of the active margin. We thus carried out an extended geochemical and geochronological microanalytical investigation of recently discovered mafic/ultramafic rocks and associated evolved rocks cropping out at Teall Nunatak (Prince Albert Mountains — North Victoria Land). Results suggest that subduction related magmatic precursors of the late- to post collisional Granite Harbour Intrusives occurred in the North Victoria Land up to 520 Ma. 2. Geological setting and field relations The Cambro–Ordovician Ross–Delamerian Orogeny began in the early Palaeozoic with the subduction of oceanic lithosphere beneath the palaeo-Pacific margin of the supercontinent Gondwanaland (Bradshaw and Laird, 1983; Kleinschmidt and Tessensohn, 1987; Borg and De Paolo, 1991; Goodge, 1997). The belt is over 4000 km long extending from southeastern Australia to the margin of the East-Antarctic Craton (Gunn and Warren, 1962; Borg and De Paolo, 1991; Goodge et al., 1993). In Antarctica (Fig. 1a), the Ross Orogen spans the whole width of the continent and the Transantarctic Mountains represent its uplifted and exposed basement (Stump, 1995). The Granite Harbour Intrusives represent the products of magmatic activity in the active Gondwanaland margin (Wilson Terrane, Fig. 1b) during Palaeozoic subduction (Gunn and Warren, 1962; Bradshaw and

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Fig. 1. a) Sketch map of Antarctica and location of the Transantarctic Mountains; b) schematic geologic map of North Victoria Land, with location of the study area (Teall Nunatak). Asterisks show the known outcrops of mafic/ultramafic enclaves in North Victoria Land.

Laird, 1983; Kleinschmidt and Tessensohn, 1987; Kleinschmidt et al., 1992; Tessensohn and HenjesKunst, 2005). Geochronological investigations of these plutonic rocks in different sectors of the orogen have revealed the presence of distinct magmatic pulses separated in both time and space. In the central Transantarctic Mountains and South Victoria Land magmatic activity with calc-alkaline to adakite affinity, evidence of active convergent tectonic activity, covers a time span of 70 Ma from about 550 to about 480 Ma (cf. Goodge et al., 1993; Rowell et al., 1993; Allibone et al., 1993a,b; Allibone and Wysoczanski, 2002; Cottle and Cooper, 2006a; Stump et al., 2006; Bomparola et al., 2007). In North Victoria Land, the northernmost sector of the incipient arc, the occurrence of such old intrusions is questionable. Comparable, though poorly constrained ages of 530 ± 57 and 522 ± 17 Ma (Rb/Sr whole rock isochron and U/Pb on zircon populations, respectively) were obtained for a strongly deformed monzo-granite orthogneiss (Tonarini and Rocchi, 1994; Rocchi et al., 2004). A mean concordia age of 544 ± 4 Ma has been

determined by Black and Sheraton (1990) in zircons of a diatexite (Schlüsser et al., 2006) from the Rennick Glacier: this age has been interpreted as evidence of the granitic magma production during an early stage of the Ross Orogen, predating of about 70 Ma the emplacement of the Granite Harbour Intrusives. The magmatic climax of the Ross Orogeny in North Victoria Land has been dated between about 510 and 480 Ma (Borg et al., 1987; Armienti et al., 1990; Rocchi et al., 2004; Bomparola et al., 2007; Schlüsser et al., 2006). Intrusive rocks emplaced during this time interval have diverse compositions, varying from calcalkaline through high-K calc-alkaline to shoshonitic (Ghezzo et al., 1987; Armienti et al., 1990; Rocchi et al., 1998; Rocchi et al., 2004). Intermediate and felsic intrusives dominate the Granite Harbour Intrusives in North Victoria Land: however, recent field reports indicate that meter-scale enclaves of mafic/ultramafic rocks with cumulate texture occasionally occur within some plutonic complexes in the Prince Albert Mountains and in the Deep Freeze Range area (Fig. 1b). One of the most preserved and best exposed examples of

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these mega-enclaves crops out at Teall Nunatak, a 4 km2 plutonic complex at the confluence of the Reeves and Priestley Glaciers (Fig. 2). Teall Nunatak is a composite intrusive body with coarse-grained monzo- to syenogranite and medium-grained biotite-bearing gabbrodiorite to quartz-diorite as the dominant lithologies. The mafic and intermediate rocks intruded the granitoids and strongly interacted with them. The cumulate mega-enclaves at Teall Nunatak are associated with the

Fig. 3. a) Mafic/ultramafic cumulate enclave within gabbro-diorites at Teall Nunatak (Prince Albert Mountains); b) similar cumulate enclave within a monzogranite sill from the south-eastern flank of Capsize Glacier (Deep Freeze Range); c) typical mingling relationships between felsic and mafic dykes at Teall Nunatak.

Fig. 2. Detailed geologic map of Teall Nunatak and its panoramas at two different scales.

gabbro-diorite (Fig. 3a); in the Mt. Nansen and in the Deep Freeze Range areas, similar cumulates are hosted within monzogranite rocks (Fig. 3b). Coarse-grained ultramafic rocks are generally preserved in the core of the enclaves and contain clino- and orthopyroxene,

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olivine and large poikilitic amphibole and phlogopite crystals; the enclave rims mainly consist of coarsegrained biotite- and amphibole-bearing cumulate gabbro. The entire intrusive sequence at Teall Nunatak is crosscut by a subvertical dyke network consisting in fine-grained biotite diorite to Qtz-diorite and mediumto fine-grained syenogranite with net-veining relationships indicative of co-mingling (Fig. 3c).

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interstices. The poikilitic brown mica and amphibole are often intergrown suggesting contemporaneous crystallisation. Poikilitic amphibole is zoned with brown cores and green rims. Zircon and ilmenite are accessory minerals. Locally, chlorite partially replaces amphibole and mica.

3. Petrography 3.1. Gabbro-diorites Gabbro-diorites are coarse- to medium-grained and usually isotropic. They occasionally exhibit a slight preferred orientation of minerals near the contacts with the host granite or have developed mylonitic fabrics within isolated, meter-scale late-magmatic shear zones. The ferromagnesian minerals include clinopyroxene, orthopyroxene, biotite, green to colourless amphibole and accessory ilmenite. Plagioclase, quartz and subordinate K-feldspar are the main nonferromagnesian minerals. Clinopyroxene is occasionally zoned and shows skeletal texture, with small amphibole and biotite intergrowths. Biotite occurs as large, unaltered and often poikilitic crystals containing ilmenite, pyroxene and plagioclase. The modal abundance of green amphibole varies considerably from sample to sample. It replaces clinopyroxene and grades to colourless amphibole (actinolite–cummingtonite) towards the rim. Ilmenite is commonly anhedral and included within biotite. Plagioclase is euhedral to subhedral and often displays complex patchy zoning. A few crystals have a normal concentric zoning. Orthoclase, microcline and quartz are interstitial phases in the most differentiated domains; quartz is always interstitial. Acicular apatite, and subhedral to euhedral allanite, titanite and zircon are common accessory minerals. 3.2. Mafic/ultramafic cumulate enclaves Ol-bearing ultramafic cumulates (Fig. 4a) are coarse grained and consist of orthopyroxene, poikilitic amphibole and brown mica, olivine and minor amounts of plagioclase. Clinopyroxene occurs as relics with corroded boundaries within the large poikilitic amphibole and brown micas and is present only in minor amounts (Fig. 4b, c). Olivine is enclosed within orthopyroxene, amphibole or mica, but shows no relationship with clinopyroxene. Subhedral plagioclase with corroded boundaries is found as inclusions in orthopyroxene, brown amphibole and mica. Anhedral, strongly zoned plagioclase occurs in the

Fig. 4. Photomicrographs of one selected cumulate sample (DR22). a) Cumulus olivine and plagioclase within poikilitic brown amphibole; b), c) partially resorbed clinopyroxene relics within poikilitic brown amphibole.

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In the mafic cumulates olivine disappears, while plagioclase, brown mica and amphibole modal amounts increase. Textural relations resemble those observed in Ol-bearing rocks. 3.3. Granitoids Granitoid rocks are coarse-grained and span in composition from Qtz-monzonite through monzogranite. They can be slightly foliated, especially near the contacts with diorites. The foliation is magmatic and outlined by the preferred orientation of K-feldspar megacrysts and biotite-rich schlieren. The mafic minerals are mainly biotite and green amphibole. Accessory phases in decreasing order of abundance are ilmenite, titanite (often rimming ilmenite), allanite, apatite and zircon. 3.4. Dyke network The dyke network is made up of fine-grained porphyritic diorite (and Qtz-diorite) and syenogranite net-veining each other, which suggests contemporaneous emplacement. In the mafic-intermediate dykes, biotite is the most abundant ferromagnesian phase, followed by green amphibole, often preserving diopside relics at core. Plagioclase is usually subhedral, and large porphyritic crystals with a typical sieved texture defined by micro-inclusions of clinopyroxene, biotite and ilmenite are common. The modal abundance of titanite may reach 5 vol.% in some samples. Accessory amounts of apatite and ilmenite needles are found mainly within plagioclase. The abundance of the ferromagnesian phases decreases when increasing mingling/mixing with the felsic domains. Dykes with syenogranite composition have finegrained textures dominated by plagioclase and microcline crystals. The ferromagnesian phases consist of biotite (± green hornblende), whereas titanite and euhedral, zoned allanite are minor constituents. Mafic enclaves in form of small fragments, pillows and clots are common and result from mingling with the associated diorites. 4. Analytical methods Bulk rock major and trace-element compositions were determined by ICP-AES at SARM CRPG–CNRS of Nancy, France. The analytical accuracy for major and trace elements is tested on more than eight reference materials (N 25 analyses each) during routine analytical sessions over a period of more than six months. Accuracy for major elements varies between 1 and 3% for

concentrations above 5 wt.% and between 2 and 10% for concentrations between 1 and 5 wt.%. Accuracy for trace elements varies between 5 and 15% for concentrations above 1 ppm and between 10 and 25% for concentrations below 1 ppm. Electron microprobe analyses were performed at the CNR–Istituto di Geologia Ambientale e Geoingegneria (IGAG), Rome, Italy. Mineral compositions were measured using a Cameca SX 50 equipped with five WDS spectrometers. Operating conditions were set at 15 kV accelerating voltage, 15 nA beam current on the sample, 10 μm beam diameter, and 20 s counting time. Mineral standards where analysed prior and after each analytical run. Standards used were: wollastonite (SiK, TAP; CaK, PET), periclase (MgK, TAP), corundum (AlKα TAP), jadeite (NaKα, TAP), orthoclase (KKα, PET), magnetite (FeKα, LIF), rutile (TiKα, LIF), Mn metal (MnKα LIF), fluorphlogopite (FKα, TAP) and sylvite (ClKα, TAP). Estimated errors are about 1% rel. for elements with concentration above 2 wt.% and about 3–5% rel. for Na and elements with concentration below 2 wt.%. The trace element mineral composition was determined by laser ablation (LA)-ICP-MS at the CNR– Istituto di Geoscienze e Georisorse — Unità di Pavia, Italy. The reader is referred to Tiepolo et al. (2003) for the analytical details. The LA-ICP-MS instrument couples a Nd:YAG laser working at 213 nm with a double focusing sector field ICP mass spectrometer type Element I from Thermo. The laser was operated at a repetition rate of 10 Hz, and the spot diameter was from 20 to 40 μm and a respective pulse energy of about 0.01 and 0.03 mJ. Data reduction was performed using the software package “Glitter” (van Achterbergh et al., 2001). The sample NIST SRM 612 (assumed composition reported by Pearce et al., 1997) was used for an external standard. 44 Ca or 29 Si was adopted for an internal standard depending on the mineral composition previously determined by microprobe analyses. During each analytical run the BCR-2 USGS (assumed composition reported by Wilson, 1997) reference glass has been analysed as unknown to assess precision and accuracy levels: these resulted always better than 6% relative. Zircon grains for U–Pb geochronology were separated using standard techniques, mounted in epoxy resin and then polished, with the final step using 1/4 μm diamond paste. In-situ U–Pb geochronology for the LC4 dioritic sample was carried out at CNR–Istituto di Geoscienze e Georisorse — Unità di Pavia, Italy with the same laser ablation (LA)-ICP-MS instrument previously described for trace element determinations and strictly adopting the analytical method described in

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Tiepolo (2003). Instrumental and laser induced U/Pb fractionation was corrected using the 1065-Ma zircon standard 91500 (Widenbeck et al., 1995). Zircon 02123 (295 Ma; Ketchum et al., 2001) was analysed together with the unknown samples for quality control (results reported in Table 7). Spot size was set to 20 μm and output energy at 2.5 mW. Data reduction was carried out with the software package “Glitter” (van Achterbergh et al., 2001), whereas for age calculations and concordia plots the Isoplot 3.0 software by Ludwig (2000) was used. The reproducibility of the external standard within the same analytical run was propagated to each determination for all isotope ratios according to the method reported in Horstwood et al. (2003). After this correction each age estimate is retained accurate within quoted errors. Zircons from the Ol-bearing cumulate DR22 were dated at the CNR–Istituto di Geoscienze e Georisorse — Unità di Pavia using an ArF excimer laser microprobe working at 193 nm (Geolas200Q-Microlas) coupled to the Element ICP-MS previously described. The spot size was set to 20 μm and laser fluency at 12 J/cm2. The analytical method used for this sample was analogous to that previously described with the only exception that the 235U signal is calculated from 238U on the basis of the natural ratio 238 U/235U = 137.88 (cf. Giacomini et al., 2007). The 18 O/16 O ratio of mineral separates was measured by conventional laser fluorination at the CNR– Istituto di Geoscienze e Georisorse — Pisa, following the procedure described by Sharp (1995). Pure F2 desorbed from K3NiF7 salt (Asprey, 1976) was used as reagent, and O2 was the analyte measured with a Finnigan Delta XP MS. Precision and accuracy of the analyses were monitored by measuring aliquots of laboratory quartz standards (QMS, δ18 O = 14.05‰) during each set of analyses, yielding an average reproducibility of ± 0.12‰ (2σ; n = 9). The NBS-28 (δ18 O = 9.60‰) measured in Pisa yielded an average δ18 O = 9.54 ± 0.15‰ (2σ, n = 4). No data correction was necessary, and results are reported in the standard per mil delta notation vs. SMOW. 5. Whole rock chemistry Harker-type diagrams show trends with continuous distributions of points for covariation between SiO2 and other major elements from the gabbro-dioritic to granitoids rocks (e.g. Fig. 5). The mafic/ultramafic enclaves deviate in these trends: the lower CaO, TiO2, Al2O3, Na2O, K2O and higher MgO and MnO concentrations at SiO2 concentrations comparable with those of gabbrodiorites are indicative of their cumulus origin (Table 1).

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Fig. 5. a) MgO against SiO2 and b) Y against SiO2 for the samples studied compared to the fields for the Abbott Unit and the Vegetation Unit using data from Di Vincenzo and Rocchi (1999) and Dallai et al. (2003).

The mafic and felsic end-members from the dyke network are compositionally similar to the host coarsegrained intrusives. The chondrite-normalised REE pattern of cumulate enclaves (Fig. 6a) is characterised by a strong Light Rare Earth Elements (LREE) enrichment relative to Heavy Rare Earth Elements (HREE), which are at about 10 times C1 chondrite (Sun and McDonough, 1989). The C1-normalised REE patterns of gabbro-dioritic rocks parallel those of the enclaves, but at higher concentrations (HREE concentrations are about 50-fold those of C1 chondrite). HREE concentrations in granitoids are comparable to those in the cumulates, but LREE are about one order of magnitude more enriched. The (La/ Yb)N ratio steadily increases with the silica concentration of the rock. The N-MORB-normalised (Sun and McDonough, 1989) incompatible trace element patterns (Fig. 6b) of the mafic cumulate and gabbro-diorite are almost parallel, and are characterised by a strong enrichment in Large Ion Lithophile elements — LILE (up to 100 times N-MORB) compared to High Field Strength Elements (HFSE) and REE. Pairs of Nb–Ta and Zr–Hf show negative and positive anomalies with

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Table 1 Representative whole-rock major and trace element analyses of samples from Teall Nunatak Sample

DR22

F8

F8b

C12

DR9

LC6

Z7

LC4

AF13

C 11

AF7

LZ58

Z3

Lithotype

Ol-cum

Ol-cum

Cum

Gbdr

Gbdr

Gbdr

Gbdr

Gbdr

Gbdr

Gbdr

Gbdr

Gbdr

Gbdr

SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO Na2O K2O P2O5 LOI Sum Ba Rb Sr Cs Ga Ta Nb Hf Zr Y Th U Cr Ni Co V Cu Pb Zn La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

44.40 0.81 8.56 2.71 12.63 0.19 18.15 6.50 1.45 1.62 0.28 2.58 99.88 370 56 323 0.99 10.2 0.39 6.01 2.22 123 23.0 3.00 1.00 434 318 69 125 17 10 98 17.0 36.0 3.9 19.0 3.50 0.97 3.32 0.51 2.81 0.63 1.49 0.19 1.26 0.23

44.78 0.65 9.01 1.97 10.33 0.19 21.87 5.14 1.29 1.18 0.23 2.95 99.59 369 52 293 1.09 10.8 0.40 5.74 2.38 95 14.5 2.49 0.70 314 313 88 95 19 6 114 16.1 33.0 4.3 17.7 3.57 1.02 3.05 0.46 2.70 0.52 1.45 0.21 1.39 0.22

48.24 0.84 14.04 1.68 7.92 0.14 12.52 8.79 1.97 1.52 0.28 1.97 99.91 445 60 438 1.16 15.4 0.48 6.87 3.01 116 19.4 2.43 0.74 388 136 53 137 23 7 96 19.6 40.7 5.4 22.2 4.62 1.37 4.12 0.61 3.56 0.69 1.90 0.28 1.84 0.28

46.49 2.56 18.74 0.09 10.93 0.15 4.17 9.28 2.88 1.82 0.91 2.00 100.02 1761 82 916 1.75 25.5 0.88 14.08 3.23 142 32.7 2.10 0.62 b D.L. b D.L. 29 241 13 7 137 36.5 87.2 11.4 49.4 10.05 3.07 8.61 1.16 6.42 1.13 3.04 0.40 2.47 0.37

47.56 2.53 17.74 2.87 8.21 0.19 4.25 8.33 3.13 2.82 1.1 1.49 100.22 2502 109 979 0.97 26.7 1.38 25.54 4.13 184 47.5 1.94 0.46 b D.L. b D.L. 27 214 15 11 161 51.8 129.3 17.4 75.7 15.76 3.89 12.89 1.77 9.52 1.69 4.44 0.58 3.43 0.50

47.95 2.23 19.68 0.86 9.35 0.18 3.62 7.77 3.42 2.81 0.84 1.73 100.44 3045 95 1100 0.76 24.2 0.90 15.96 1.10 34 32.7 1.38 0.33 b D.L. b D.L. 17 106 8 7 128 44.8 108.2 13.1 52.1 9.65 3.71 7.71 1.07 6.24 1.13 3.23 0.47 2.98 0.42

49.26 2.55 15.77 1.98 8.02 0.15 5.22 8.15 2.69 2.49 1.11 2.00 99.39 907 82 686 2.34 21.0 1.63 24.07 6.13 267 36.1 4.55 1.00 81 33 29 162 25 18 142 59.0 128.3 16.2 66.1 12.59 3.03 9.61 1.33 7.19 1.25 3.38 0.45 2.92 0.42

49.90 1.16 19.20 0.22 9.47 0.19 4.58 7.76 3.29 2.26 0.18 1.80 100.01 728 72 649 1.47 20.7 1.21 18.98 5.87 258 28.9 4.32 0.74 209 49 24 117 21 16 115 48.2 105.3 13.3 53.2 10.16 2.27 8.00 1.10 5.80 1.01 2.74 0.38 2.38 0.35

50.62 2.35 14.55 0.80 11.34 0.21 5.01 8.06 2.83 1.84 0.34 1.89 99.84 342 109 293 5.74 25.8 1.13 13.58 6.03 235 53.8 5.61 2.22 50 12 42 260 19 12 150 30.2 74.0 9.5 40.5 9.96 2.58 10.00 1.57 9.74 1.92 5.42 0.78 5.21 0.76

51.6 2.27 15.74 0.02 8.51 0.14 5.22 8.81 2.69 2.27 1.02 1.79 100.08 917 82 751 1.07 23.3 1.65 25.01 7.49 324 42.8 5.71 1.12 61 30 27 162 14 21 143 67.0 151.9 19.1 78.0 15.23 3.24 11.90 1.62 8.65 1.51 3.90 0.54 3.34 0.48

51.72 2.02 15.08 1.22 9.84 0.19 4.93 7.98 2.92 1.74 0.4 1.83 99.87 537 120 343 6.89 26.2 1.29 15.85 7.47 309 53.1 4.53 1.10 81 20 37 197 22 14 158 30.4 74.2 9.6 41.8 9.96 2.62 9.80 1.55 9.35 1.84 5.32 0.76 5.33 0.78

51.91 1.45 17.25 1.44 5.87 0.11 4.92 8.38 2.59 2.49 0.52 3.04 99.97 742 94 666 3.00 21.9 1.16 18.07 7.32 331 24.5 4.30 1.05 192 47 20 110 20 15 107 46.0 95.9 11.9 46.1 8.49 2.10 6.35 0.91 4.71 0.83 2.32 0.34 2.24 0.33

52.02 2.03 16.12 2.29 6.90 0.13 4.45 7.04 3.00 2.78 0.86 2.38 100.00 1078 91 657 1.88 21.9 1.74 25.17 6.31 271 35.2 6.27 1.26 72 30 24 128 22 18 133 57.5 127.8 15.7 61.9 11.80 3.01 9.04 1.30 6.93 1.24 3.27 0.45 2.89 0.45

Ol-cum: olivine-bearing cumulate; Cum: cumulate; Gbdr: gabbro-diorite; Dr: diorite; Mz: monzonite; Mzgr: monzogranite; Sygr: syenogranite; Q-: quartz-. *Composite dyke of intermediate composition due to extensive mingling at the micro-scale. b D.L.: below detection limits.

respect to the neighbouring elements. Incompatible trace element concentrations are generally higher in the gabbro-diorite than in the mafic/ultramafic cumulate enclaves. The negative Sr anomaly is pronounced in the

gabbro-diorite. Incompatible trace element patterns similar to those of gabbro-diorite are also observed in granitoid rocks, but with higher LILE concentrations and more pronounced negative Nb–Ta, Sr and Ti anomalies.

F. Giacomini et al. / Chemical Geology 240 (2007) 103–128

LZ56

Z4

LZ59

LZ57

F6

Gbdr

Gbdr

Gbdr

Gbdr

52.03 1.44 17.08 0.88 6.67 0.12 5.64 9.02 2.64 1.94 0.53 1.98 99.97 723 65 673 1.23 20.7 1.09 17.08 6.45 275 28.9 3.81 0.74 165 48 24 113 16 15 111 45.8 100.4 12.5 50.7 9.67 2.26 7.51 1.07 5.72 1.02 2.69 0.39 2.41 0.36

52.11 2.06 16.14 1.47 7.75 0.13 4.41 6.97 3.08 2.95 0.86 1.93 99.86 1090 90 662 1.47 21.6 1.69 25.20 8.12 360 35.6 6.15 1.30 72 29 24 129 22 18 136 61.0 131.9 16.1 64.2 11.80 2.90 9.39 1.31 6.95 1.21 3.24 0.45 3.00 0.47

52.20 1.54 16.82 1.35 6.34 0.12 5.50 8.87 2.61 2.07 0.56 1.96 99.94 762 75 672 1.52 21.3 1.25 19.46 6.11 268 30.5 4.45 0.80 221 50 25 122 22 17 120 50.2 110.4 13.7 55.6 10.30 2.35 8.14 1.15 6.03 1.06 2.75 0.37 2.45 0.36

52.32 53.52 1.44 2.08 17.11 15.85 1.47 1.60 6.12 7.21 0.11 0.13 5.57 4.06 8.94 6.72 2.65 3.14 2.01 3.27 0.52 0.92 1.68 1.42 99.94 99.92 717 1078 66 119 684 615 0.99 1.37 20.9 22.7 1.14 1.92 17.51 30.85 5.58 9.07 242 426 27.9 42.4 4.03 7.15 0.70 1.28 160 59 48 22 23 21 114 133 15 12 14 19 108 147 47.1 57.5 101.1 132.4 12.8 18.0 50.7 73.5 9.83 14.48 2.38 3.06 7.41 11.35 1.07 1.57 5.61 8.35 0.98 1.47 2.58 3.83 0.36 0.54 2.28 3.44 0.34 0.51

Gbdr

C8

DR13

F3

F9C

F1

F4b

111

DR14

F6b

Sy-gr dyke 71.63 0.33 14.12 0.76 1.48 0.03 0.36 1.60 2.70 6.02 0.10 0.98 100.11 803 281 161 2.76 21.0 1.39 19.23 8.66 357 37.8 23.26 1.65 11 b D.L. 3 15 – 30 53 88.1 175.5 19.3 65.4 11.06 1.00 8.41 1.27 6.96 1.28 3.56 0.51 3.33 0.49

Q-dr

Q-Mz

Mzgr

Dr dyke

Dr dyke

Dr dyke

Dr dyke

Gdr dyke⁎

59.13 1.85 14.75 0.50 7.62 0.12 2.68 5.26 2.51 3.64 0.54 1.34 99.94 1056 156 334 1.70 25.6 1.83 27.01 12.43 527 58.7 14.95 1.87 35 9 19 130 12 25 142 72.2 162.2 20.2 82.0 16.30 2.75 13.45 1.96 11.19 2.10 5.65 0.80 5.32 0.79

63.74 0.73 16.84 0.85 3.55 0.06 1.09 3.04 3.25 5.59 0.18 1.12 100.04 1900 206 407 3.81 23.0 1.17 16.79 13.24 561 30.2 15.10 1.72 8 b D.L. 6 22 5 24 69 55.5 113.6 13.3 50.2 9.59 1.91 7.21 1.05 5.61 1.05 2.88 0.40 2.61 0.41

68.23 0.52 15.55 0.92 2.32 0.05 0.91 2.83 3.45 4.59 0.20 0.80 100.37 1357 119 433 1.37 20.2 0.76 14.28 6.83 295 20.0 7.34 0.74 10 b D.L. 6 32 – 18 77 54.1 106.5 11.8 43.6 7.50 1.59 5.67 0.78 4.16 0.72 1.86 0.25 1.62 0.25

51.62 1.78 15.85 1.72 6.71 0.14 5.15 7.69 2.94 3.09 0.62 1.67 98.98 774 145 517 2.98 21.7 1.46 21.87 7.18 332 34.7 6.34 1.39 173 51 26 147 23 14 124 53.4 116.8 14.6 58.4 11.25 2.62 8.86 1.26 6.84 1.20 3.18 0.45 2.92 0.43

52.19 1.78 16.02 1.82 6.80 0.14 5.32 7.74 2.98 3.26 0.66 1.74 100.45 752 144 511 2.99 21.6 1.43 21.63 7.03 315 34.7 6.07 1.35 171 52 26 146 20 14 136 53.2 116.2 14.4 57.9 11.12 2.60 8.86 1.25 6.64 1.18 3.16 0.44 2.85 0.43

52.48 1.87 15.89 1.94 6.73 0.12 4.88 7.35 3.03 3.15 0.72 1.83 99.99 1008 139 645 2.22 22.3 1.63 25.15 7.66 349 35.6 6.22 1.39 153 43 25 135 23 17 134 61.9 132.5 16.4 64.9 12.06 2.74 9.47 1.30 6.88 1.23 3.28 0.45 2.87 0.43

55.44 1.51 15.84 0.86 6.77 0.11 4.64 6.79 3.03 3.08 0.5 1.39 99.96 962 145 639 2.58 24.2 1.53 20.57 7.26 311 41.9 14.12 1.76 153 44 26 143 16 18 120 66.8 147.6 17.5 67.4 12.75 2.53 9.51 1.37 7.42 1.35 3.60 0.50 3.27 0.50

65.15 0.66 14.93 1.21 2.81 0.06 1.54 2.62 2.92 5.69 0.28 1.32 99.19 977 187 287 1.89 21.6 1.55 21.47 9.90 430 38.5 18.45 1.72 43 14 7 37 8 29 79 72.2 149.6 17.2 62.3 11.64 1.86 8.97 1.32 7.14 1.32 3.61 0.53 3.37 0.51

As shown in Figs. 5 and 6, the whole rock major and trace element compositions of the mafic and felsic intrusives of Teall Nunatak closely resemble those of the Abbott Unit in the Terra Nova Intrusive Complex

F5

described by Di Vincenzo and Rocchi (1999), Dallai et al. (2003) and Rocchi et al. (2004). Strong compositional similarities were found also between the mafic dykes from Teall Nunatak and the Vegetation

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F. Giacomini et al. / Chemical Geology 240 (2007) 103–128

Fig. 6. Diagrams for whole rock REE (a) and incompatible trace element (b) normalised to C1 chondrites and N-MORB, respectively. Chondrite patterns are characterised by a strong LREE enrichment relative to HREE; incompatible trace element patterns are characterised by a strong enrichment in LILE with respect to HFSE and REE. The C1 chondrite and N-MORB values are from Sun and McDonough (1989).

Unit in the Terra Nova Intrusive Complex (Di Vincenzo and Rocchi, 1999; Rocchi et al., 2004). 6. Mineral chemistry Fresh samples representative of the Ol-bearing cumulate (samples DR22 and F8), of the Ol-free cumulate (sample F8b), of gabbro-diorites (sample LZ56) and mafic dykes from the dyke network (sample F9c) were selected for major and trace element analysis of minerals. Tables 2–5 report the results of major and trace element analyses on representative minerals. 6.1. Olivine and clinopyroxene Olivine from the ultramafic cumulates (sample DR22 and F8) is relatively homogeneous with compositions of Fo67–72 and negligible Ni concentrations. Clinopyroxene relics from the cumulate DR22 have a Ca-rich augitic composition and relatively high Mg# (0.82–0.83). Its

chondrite-normalised REE pattern is characterised by a slight LREE enrichment with respect to the almost flat HREE (LaN/YbN =3.5–4.3). HREE values are about 7 times those of C1-chondrite (Fig. 7a, Table 3). The negative Eu anomaly is negligible. The chondrite-normalised incompatible trace element pattern (Fig. 8a) reveals that HFSE (Nb, Ta, Zr, Hf, Ti) are depleted relative to the REE and actinides (ThN/UN ratio 1.6–2.0). Cr concentrations are high (around 4000 ppm), whereas Li, V and Sc are relatively low at around 5, 240 and 80 ppm, respectively. The clinopyroxene in gabbro-dioritic rocks is occasionally zoned. Its composition ranges from diopside at the core to augite towards the rim with significantly lower Mg# (0.70–0.65) than in clinopyroxene from the cumulate rocks. The chondritenormalised REE pattern also differs from that of clinopyroxene in Ol-bearing cumulates. REE concentrations are significantly higher (up to 200 times C1 chondrite), LREE show a convex upward pattern and Eu is markedly depleted with respect to the neighbouring REE. Cr concentrations are lower (500–150 ppm), and Li, Zr and Hf concentrations are higher than in clinopyroxene from cumulates. The composition of clinopyroxene in the mafic dyke (F9c) is not significantly different from that of clinopyroxene in the coarsegrained gabbro-diorites, except for slightly lower Zr and Hf concentrations. 6.2. Orthopyroxene Orthopyroxene in the Ol-bearing cumulate is En74–78. No significant zoning was observed and the Mg# varies between 0.74 and 0.78. The chondrite-normalised REE pattern is characterised by a steady enrichment from LREE to HREE (at about 1–2 times the concentration in C1 chondrite) with a strong REE fractionation (LaN/YbN = 0.07–0.11; Fig. 7b, Table 3). The incompatible trace element pattern (Fig. 8b) shows marked negative Sr, Pb and Ti anomalies with respect to the neighbouring elements. Th and U concentrations are comparable to those of the LREE, and Li ranges between 1.5 and 1.7 ppm. Cr concentrations are low (3.2–5.4 ppm). The orthopyroxene from the gabbro-diorites has Mg# = 0.68, significantly lower than that of orthopyroxene from the Ol-bearing cumulate. The C1-normalised REE pattern of orthopyroxene of the gabbro-diorite is similar to that of orthopyroxene from the Ol-cumulate, but is shifted towards values that are about one order of magnitude greater. LREE and HREE are respectively at about 3 and 10 times C1 chondrite. The incompatible trace element pattern (Fig. 8b) parallels that of orthopyroxene from Ol-bearing cumulate, except for the presence

F. Giacomini et al. / Chemical Geology 240 (2007) 103–128

of a negative Zr–Hf anomaly and a slight positive Ti anomaly. 6.3. Amphibole The brown amphibole rimming and replacing clinopyroxene crystals in the Ol-bearing cumulate is a Ti-rich pargasite (TiO2 = 1.47–2.32) with Mg# ranging from 0.72 to 0.74. The chondrite-normalised REE pattern (Fig. 7c, Table 4) is similar to that of the associated clinopyroxene and is characterised by a steady decrease in values from LREE (at about 100 times C1) to HREE (at about 15 times C1). LaN/YbN values range between 5.7 and 9.5. The chondrite-normalised incompatible trace element pattern (Fig. 8c) reveals Rb, Sr and Pb depletion relative to the neighbouring elements. Ba and Nb–Ta are, respectively, significantly and slightly depleted with respect to LREE. Th is enriched compared to U (ThN/UN = 1.8–2.8) and Cr concentrations range from 760 to 6000 ppm. The composition of poikilitic brown amphiboles ranges from Ti-rich pargasite at the core (Mg# 0.75) to Mg-hornblende at the rim (Mg# 0.75–0.71) with very low Ti concentrations (0.1 atoms per formula unit). Chondrite-normalised REE patterns (Fig. 7c) differ from those of Ti-rich pargasite replacing clinopyroxene in that they show convex upward LREE and nearly flat HREE patterns with values about 30–40 times those of C1 chondrite. The LaN/YbN ratio ranges between 2.6 and 1.9 and there is occasionally a weak negative Eu anomaly. Incompatible trace element patterns (Fig. 8c) show that Ba and Sr concentrations are higher than those of brown amphibole replacing clinopyroxene. Cr concentrations are significantly lower (265–588 ppm), whereas the ThN/UN ratio is highly variable (0.4–1.2). Sc and V concentrations are 43.6–71.4 and 318–746 ppm, respectively. Amphibole in the gabbro-diorite (LZ56) and in the diorite dyke (not reported in the figure) is a Mghornblende compositionally equivalent to the one found in the rims of poikilitic amphibole in cumulates. The chondrite-normalised REE pattern (Fig. 7c) is characterised by marked LREE enrichment compared to the HREE (LaN/YbN = 5.3–5.5), and a negative Eu anomaly. The incompatible trace element pattern (Fig. 8c) reveals that Ba, Sr and Ti concentrations are significantly lower than those in brown amphiboles from cumulates. A higher Nb/Ta fractionation (NbN/TaN = 0.3–0.5) is also observed. 6.4. Plagioclase The corroded plagioclase included into orthopyroxene and pargasitic amphibole in the Ol-bearing cumulate has a

113

calcic composition of An82–84. The anhedral plagioclase in the interstices is more sodic with An50–52. The calcic plagioclase shows a chondrite-normalised REE pattern (Fig. 7d, Table 5) with LREE concentrations about 10 times those of C1 chondrite and LaN/SmN values at 4.6–5.7. Ba and Pb concentrations are 55–68 and 24–28 ppm, respectively. The more sodic plagioclase in the interstices shows LREE concentrations up to one order of magnitude higher than the aforementioned one (100 times C1 chondrite) and also higher LaN/SmN values (47–58). Sr, Ba, U and Pb concentrations (Table 5) are also significantly higher (up to one order of magnitude) than in calcic plagioclase. The LREE concentrations of the An-rich plagioclase in the Ol-free cumulate range between 17 and 37 times those of C1 chondrite and the LaN/SmN value is 8.9–17. The Ba concentration is extremely variable (85–3570 ppm), whereas Sr and Pb concentrations are 911–1670 and 13–32 ppm, respectively. Plagioclase in the gabbro-dioritic rocks is commonly andesine to labradorite in composition (An35–53) and displays complex patchy zoning, sometimes preserving relics of calcic labradorite composition (An65). A few crystals have normal concentric zoning. The composition of plagioclase from the mafic dyke (F9c) resembles that of plagioclase from the coarse-grained diorite, both having high LREE concentrations (up to 119 times that of C1 chondrite) and showing strong fractionation of LREE over HREE (LaN/SmN is up to 61). 6.5. Brown mica The poikilitic brown mica in the cumulates is characterised by relatively high Mg# (0.80–0.82) and occasionally high Ti concentrations (up to 0.12 apfu); it shows high Ba, Rb, Sr and appreciable Nb, Ta and Pb concentrations. Th–U and LREE are significantly depleted relative to the neighbouring elements in the incompatible trace element pattern (Fig. 8d, Table 5). Brown mica shows significantly lower Mg# (0.58–0.56), Sr and Pb concentrations in the gabbro-diorite than in the cumulate rocks. Higher Li, Nb, Ta and Ti concentrations were instead observed. The composition of brown mica in the dioritic dyke is analogous to that of biotite from the coarse-grained intrusives. 7. Oxygen isotope mineral composition Olivine, orthopyroxene and plagioclase separated from the Ol-bearing cumulates and orthopyroxene and plagioclase from the gabbro-diorite were measured for δ18O (Fig. 9, Table 6). The δ18Oolivine in the cumulates (sample DR22 and F8) are 7.03 and 7.19‰, whereas the δ18Oorthopyroxene in the same rocks are 7.13 and 7.47‰.

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F. Giacomini et al. / Chemical Geology 240 (2007) 103–128

Table 2 Representative major-element analyses of the main mineral phases Sample

DR22

F8

DR22

LZ56 F9c

Lithotype Ol-cum Ol-cum Ol-cum GDr

DR22

F8

F8b

Dr dyke Ol-cum Ol-cum Cum

Mineral

Olivine

Clinopyroxene

SiO2 TiO2 Al2O3 Cr2O3 FeO MnO MgO CaO Na2O K2O NiO F Cl Sum Si Ti Al Cr Fe3+ Fe2+ Mn Mg Ca Na K Ni F Cl

37.38 0.01 – 0.04 28.09 0.38 34.44 0.03 0.01 0.01 0.03 – – 100.36 1.00 – – – – 0.63 0.01 1.37 – – – – – – Fo%

38.08 – – – 25.39 0.28 36.36 0.08 0.02 – – – – 100.15 1.00 – – – – 0.56 0.01 1.43 – – – – – –

53.27 0.18 1.40 0.31 6.00 0.18 15.67 22.31 0.44 – – – – 99.77 1.96 0.01 0.06 0.01 0.03 0.16 0.01 0.86 0.88 0.03 – – – – Mg#

52.07 0.37 1.34 0.06 11.58 0.30 13.26 20.42 0.30 – – – – 99.63 1.96 0.01 0.06 – 0.02 0.35 0.01 0.75 0.82 0.02 – – – –

52.54 0.09 0.35 0.03 11.12 0.55 11.83 22.80 0.27 – – – – 99.58 1.99 – 0.02 – 0.02 0.34 0.02 0.67 0.93 0.02 – – – –

54.52 0.17 1.20 0.03 15.41 0.37 27.31 0.91 0.02 – – – – 99.87 1.97 0.01 0.07

69

72

0.82

0.67

0.65

LZ56 DR22

DR22

DR22

GDr Ol-cum Ol-cum Ol-cum

Orthopyroxene

F8b

LZ56 F9c

Cum

GDr Dr dyke

Amphibole In cpx

0.02 0.45 0.01 1.47 0.04 – – – – –

54.82 0.12 1.12 0.03 15.39 0.41 27.46 0.93 0.03 – – – – 100.28 1.97 0.01 0.05 – 0.01 0.45 0.01 1.47 0.04 – – – – –

52.57 0.12 1.27 0.03 20.15 0.61 24.03 0.79 0.02 – – – – 99.60 1.94 – 0.05 – 0.04 0.58 0.02 1.32 0.03 – – – – –

52.47 0.73 1.43 0.05 20.26 0.44 23.30 1.14 0.01 – – – – 99.83 1.94 0.02 0.06 – 0.03 0.61 0.01 1.29 0.05 – – – – –

43.81 1.93 11.07 0.85 9.69 0.09 14.84 11.53 2.55 0.72 – 0.12 0.06 97.21 6.38 0.21 1.90 0.10 0.37 0.81 0.02 3.22 1.80 0.72 0.13 – 0.05 0.02

0.76

0.77

0.68

0.68 0.73

Poik. core Poikilitic 44.97 0.38 10.71 0.46 8.98 0.22 16.08 11.13 2.54 0.36 – 0.14 – 95.96 6.50 0.04 1.83 0.05 0.81 0.28 0.03 3.47 1.72 0.71 0.07 – 0.06 –

43.13 3.61 11.03 0.05 8.87 0.21 15.17 11.40 2.66 0.81 – 0.14 0.03 97.09 6.29 0.40 1.90 0.01 0.26 0.83 0.03 3.30 1.78 0.75 0.15 – 0.07 0.01

48.71 0.82 7.13 0.08 10.40 0.21 15.92 11.42 1.16 0.39 – 0.08 0.05 96.35 7.01 0.09 1.21 0.01 0.68 0.57 0.03 3.41 1.76 0.32 0.07 – 0.04 0.01

49.98 0.63 5.06 0.05 12.77 0.21 14.49 12.76 0.80 0.48 – 0.28 – 97.50 7.31 0.07 0.87 0.02 1.03 1.05 0.03 3.16 2.01 0.45 0.09 – 0.13 –

49.33 0.52 5.13 0.05 14.12 0.39 13.70 11.80 0.98 0.60 – 0.65 – 97.29 7.26 0.06 0.89 0.01 0.37 1.37 0.05 3.01 1.86 0.28 0.11 – 0.30 –

0.76

0.75

0.73

0.67 0.63

Ol-cum: olivine-bearing cumulate; Cum: cumulate; GDr: gabbro-diorite; poik.: poikilitic; inerst.: interstitial.

These data define Δ18Oopx-ol values (δ18Oorthopyroxene − δ18Oolivine) clearly below the expected 0.5‰ equilibrium fractionation between these phases. The very high δ18Oolivine values indicate that magmatic 18O/16O ratios were possibly perturbed by fluid-assisted processes, such as interaction with a hydrous magma. Deviation from the δ18Oolivine − Mg#olivine trend was observed (Dallai et al., 2003) in the amphibole-bearing cumulate gabbros of the nearby Mt. Abbott, where cumulitic olivine gabbro rocks interacted with granitic melts (Fig. 9b). In contrast the δ18Oorthopyroxene values are consistent with those of less contaminated amphibole-free gabbros, suggesting that they represent the O-isotope composition of the anhydrous magmatic phase (Dallai et al., 2003). Orthopyroxene shows slightly lower δ18O values (6.74‰) in the Olfree cumulate than in the Ol-bearing samples. The δ18 O values of orthopyroxene and plagioclase from the gabbro-diorite (LC4) are 7.6‰ and 8.8‰, respectively.

8. Zircon U–Pb geochronology Zircon grains were separated from one Ol-bearing cumulate (DR22) and from one gabbro-diorite sample (LC4). Prior to U–Pb geochronological determinations, zircon were characterised by cathodoluminescence (CL) to evidence the internal zoning and to assess the history of growth, which potentially could include inherited components, inclusions and metamict portions. This allowed the selection of the most suitable domains for in situ U–Pb geochronology. Analysis results are reported in Table 7. 8.1. DR22 Ol-bearing cumulate Zircon grains from the DR22 ultramafic cumulate (Fig. 10a–e) show prismatic and acicular habits. Both zircons types are characterised by well-developed to

F. Giacomini et al. / Chemical Geology 240 (2007) 103–128

115

Table 2 Representative major-element analyses of the main mineral phases DR22

DR22

F8

F8

LZ56

LZ56

F9C

DR22

F8

LZ56

F9c

Ol-cum

Ol-cum

Ol-cum

Ol-cum

GDr

GDr

Dr dyke

Ol-cum

Ol-cum

GDr

Dr dyke

Plagioclase

Biotite

relic

interst.

interst.

relic

46.18 – 33.79 0.02 0.00 – – 16.56 1.77 – – – – 98.30 2.15 – 1.86 – – – – – 0.83 0.16 – – – – An%

53.88 – 28.43 – 0.13 – 0.01 10.41 5.40 0.02 – – – 98.27 2.47 – 1.54 – – – – – 0.51 0.48 – – – –

54.62 – 28.67 0.01 0.04 – – 10.24 5.74 0.01 – – – 99.34 2.48 – 1.53 – – – – – 0.50 0.50 – – – –

46.55 0.02 33.81 – 0.06 – 0.02 16.76 1.99 0.01 – – – 99.23 2.15 – 1.84 – – – – – 0.83 0.18 – – – –

55.01 0.06 28.27 0.01 0.19 – – 10.42 5.67 0.21 – – – 99.85 2.49 – 1.51 – 0.01 – – – 0.50 0.50 0.01 – – –

58.58 0.04 26.16 – 0.20 – 0.01 7.63 7.01 0.26 – – – 99.88 2.62 – 1.38 – 0.01 – – – 0.37 0.61 0.01 – – –

58.33 – 24.81 0.01 0.14 – 0.00 7.54 7.85 0.07 – – – 98.74 2.65 – 1.33 – 0.01 – – – 0.37 0.69 – – – –

37.98 1.12 16.69 0.03 8.45 0.14 20.20 0.05 1.56 7.41 – 0.18 – 93.61 2.77 0.06 1.43 – 0.51 0.01 2.19 0.01 0.22 0.69 – 0.04 – – Mg#

37.83 2.17 16.75 – 8.11 0.01 20.15 – 1.89 6.84 – 0.29 – 93.92 2.74 0.12 1.43 – 0.49 – 2.17 – 0.27 0.63 – 0.07 – –

36.10 5.44 13.78 0.10 17.11 0.06 12.64 0.09 0.26 9.12 – 0.36 – 94.89 2.74 0.31 1.23 0.01 1.09 0.01 1.43 0.01 0.04 0.88 – 0.09 – –

36.98 3.41 13.41 0.11 17.52 0.21 12.54 0.12 0.08 9.17 – 1.45 – 94.39 2.81 0.20 1.20 0.01 1.11 0.01 1.42 0.01 0.01 0.89 – 0.35 – –

84

52

50

82

50

37

35

81

82

57

56

faint oscillatory zoning, typical of growth under igneous conditions (Hoskin, 2000). A texturally different core, most likely of xenocrystic origin, was found in one prismatic zircon fragment (Fig. 10e). It is unlikely to be of metasedimentary origin due to the absence of rounded boundaries. Among the separated and mounted zircons, eleven grains without fractures and inclusions were selected for U–Pb geochronology. Four of the 16 analytical spots yielded discordant ages that cluster close to the concordia (Fig. 11a). Small amount of common Pb, a weak metamictisation or small inclusions not detectable by the time-resolved signal are most likely responsible for the discordance. These data were thus excluded from age calculations. In prismatic zircons the domains with a clear igneous texture yield a mean concordia age of 521 ± 2 Ma (1σ) (Fig. 11b). The xenocrystic core yields a concordant age at 546 ± 12 Ma. Acicular zircons yield a significantly younger age

(Fig. 11c) of 502 ± 3 Ma (1σ) which is statistically different from the previous one. 8.2. LC4 gabbro-diorite Zircons from the gabbro-diorite (Fig. 10f–h) are mostly elongated with prismatic to acicular habits. Although no well-developed oscillatory zoning is observed, the occasional presence of a brighter rim in continuity with the morphology of the inner portion can be reasonably interpreted as a zoning. The inner sector of the crystals shows a homogeneous composition with no inclusions or evidence of xenocrystic cores. Eight grains with no fractures and inclusions were selected for U–Pb geochronology from among the separated and mounted zircons. All zircons yield concordant ages giving a mean concordia age at 489 ± 4 Ma (95% confidence, Fig. 11d). Note that the analysis of a brighter rim

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F. Giacomini et al. / Chemical Geology 240 (2007) 103–128

Table 3 Trace element analyses of clinopyroxene from melagabbros and diorites Sample

DR22

DR22

F8

F8

LZ56

LZ56

LZ56

F9

Lithotype

Ol-cum

Ol-cum

Ol-cum

Ol-cum

GDr

GDr

GDr

Dr dyke

Mineral

Cpx

Cpx

Opx

Opx

Cpx

Cpx

Opx

Cpx

(ppm) Li B Sc Ti V Cr Co Rb Sr Y Zr Nb Cs Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Ta Pb Th U

5.13 6.95 87.0 2450 234 4050 33.9 0.27 42.71 16.3 46.3 0.16 0.012 – 7.85 18.9 2.51 10.8 3.02 0.8 3.61 0.52 3.03 0.66 1.76 0.21 1.26 0.19 1.65 b D.L. 1.41 0.69 0.1

4.27 7.35 82.5 2224 241 3976 33.4 0.2 46.41 13.3 48.5 b D.L. 0.03 0.55 7.26 18.1 2.23 10.6 2.96 0.92 3.40 0.51 2.02 0.52 1.3 0.22 1.44 0.22 1.59 b D.L. 0.9 0.59 0.11

1.74 – 4.9 19.5 1.6 5.38 61.3 b D.L. 0.09 1.1 6.60 b D.L. b D.L. 0.03 0.034 0.135 0.024 0.090 0.052 0.016 0.063 0.017 0.128 0.030 0.154 0.027 0.323 0.067 0.195 0.004 0.008 0.006 0.001

1.51 – 4.3 13.2 0.8 3.17 61.7 b D.L. 0.08 0.5 2.76 b D.L. b D.L. b D.L. 0.020 0.072 0.016 0.053 0.017 0.006 0.041 0.009 0.059 0.020 0.075 0.015 0.125 0.031 0.087 0.006 b D.L. b D.L. 0.001

15.59 – 87.3 2418 278 432 36.9 b D.L. 33.25 130 154 2.28 b D.L. 0.11 27.2 135 24.1 116 29.1 2.10 23.6 4.22 26.4 4.68 11.6 1.72 9.19 1.24 5.35 0.114 1.88 0.123 0.079

18.01 – 80.7 1891 254 400 37.2 0.38 31.76 79 131 0.15 b D.L. 0.10 21.2 95.6 19.4 96.7 29.1 2.74 26.5 3.47 21.0 3.59 9.89 1.30 9.50 1.09 7.11 0.013 1.59 0.135 0.084

10.24 – 38.0 3942 167 416 71.0 b D.L. 0.52 7.0 10.1 0.02 b D.L. 0.31 0.34 1.29 0.205 1.27 0.529 0.151 0.960 0.136 1.34 0.367 1.052 0.171 0.810 0.280 0.355 b D.L. 0.269 0.028 0.008

13.62 – 110.3 1115 221 223 37.3 b D.L. 16.68 77 21.4 0.20 b D.L. 0.54 15.4 67.6 12.3 68.4 18.4 1.03 16.3 2.46 15.5 2.86 7.65 1.08 6.87 0.936 1.26 0.011 1.16 0.280 0.030

Ol-cum: olivine cumulate; GDr: gabbro-diorite; Dr: diorite. b D.L.: below detection limits.

(crystal C22, Fig. 10f) yields a concordia age that is indistinguishable from that of the darker cores (Fig. 10h). 9. Discussion 9.1. Evidence for disequilibrium in Ol-bearing cumulates The xenocrystic clinopyroxene with high Cr concentrations (4000 ppm) and Mg# (0.82) is the most primitive phase present in the Ol-bearing cumulate enclave. Tirich pargasite rimming and partially replacing clinopyroxene also has high Cr concentrations and a trace element signature similar to that of the clinopyroxene.

The Amph/CpxD values calculated for trace elements, HREE, Zr, Hf and transitional elements (e.g., Co, Zn, Cr) yield a nearly flat pattern with values close to unity (around 1.5). Slightly higher values (around 2) are obtained for LREE. These solid/solid partitioning values (Fig. 12) are comparable to those reported in literature (e.g., Tiepolo et al., 2000, 2001 and references therein, Rivalenti et al., 2004) for close to equilibrium conditions. Thus, Ti-rich amphibole and clinopyroxene most likely approach equilibrium conditions. Olivine shows a relatively low Fo concentration (0.70) compared to the Mg# of the clinopyroxene. Evidence for disequilibrium between olivine and clinopyroxene comes from the application of the Ol-LiqKdFe–Mg and Cpx-LiqKdFe–Mg

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The trace element signature of poikilitic amphibole is significantly different from that of clinopyroxene and the highly fractionated Amph/Cpx DREE pattern with relatively high values (up to 5) is not representative of equilibrium conditions. When applied to orthopyroxene, the Opx-Liq KdFe–Mg experimentally determined by Gudfinnsson and Presnall (2000) yields a liquid with a Fe/Mg ratio of 1.1, i.e., nearly twice that computed for clinopyroxene. The different Fe/Mg ratio of the liquid in equilibrium with orthopyroxene and olivine suggests

experimentally determined by Gudfinnsson and Presnall (2000). Assuming low-pressure conditions (around 0.5 GPa), olivine yields a liquid with a Fe/Mg ratio of about 1.4, whereas clinopyroxene yields a liquid with a ratio of about 0.6. There are no significant changes in the Fe/Mg ratio when pressure is varied. Crystallisation from a more differentiated liquid is suggested for the poikilitic amphibole and orthopyroxene in the Ol-bearing cumulates. Both minerals show lower Mg# than clinopyroxene (0.76) and negligible Cr concentrations. Table 4 Trace element analyses of amphibole from melagabbros and diorites Sample

DR22 DR22

Lithoype Olcum Mineral

(ppm) Li B Sc Ti V Cr Co Ni Zn Rb Sr Y Zr Nb Cs Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Ta Pb Th U

Olcum

DR22

DR22

DR22

F8

F8

F8

F8

F8

F8b

F8b

F8b

Lz56

Lz56

Lz56

Ol-cum

Ol-cum

Ol-cum

Olcum

Olcum

Olcum

Olcum

Olcum

Cum Cum Cum GDr

GDr

GDr

Amph Amph

Amph-rim Amph-rim Amph-rim Amph Amph Amph Amph Amph Amph Amph Amph Amph Amph Amph Cpx Cpx Cpx

poik.

poik.

rim

rim

core

poik.

poik.

poik. poik. poik.

0.72 – 63.0 27902 591 126 47.7 – – 6.68 382 59.5 94.8 14.0 0.04 390 14.7 56.3 10.4 55.6 15.8 4.29 14.3 2.17 12.6 2.51 6.60 0.79 4.86 0.63 4.01 0.76 6.53 0.58 0.19

0.45 – 63.5 28787 606 142 48.2 – – 6.12 402 60.7 94.9 15.0 b D.L. 401 16.5 62.0 11.2 60.7 16.7 4.60 13.9 2.17 13.1 2.53 6.38 0.86 5.04 0.65 3.68 0.75 7.62 0.79 0.26

1.16 12.9 68.7 8340 238 894 65.2 218 68.5 1.38 132 26.5 210 16.0 b D.L. 31.9 28.0 64.1 7.1 26.5 5.37 1.42 5.44 0.83 5.56 1.03 2.92 0.39 2.91 0.39 3.91 0.83 4.45 2.84 0.35

1.19 15.1 69.1 7957 228 766 88.9 255 118 1.49 132 26.1 214 15.8 0.07 26.2 27.9 62.1 7.0 26.2 6.17 1.65 5.91 0.84 5.53 1 2.70 0.4 2.38 0.3 4.12 0.83 4.52 3.88 0.38

1.13 9.07 90.9 13169 441 6573 50.9 144 58.5 6.80 129 29.9 126 9.04 0.04 184 20.6 47.4 6.0 26.2 6.38 1.80 7.32 1.00 6.17 1.12 2.78 0.46 2.83 0.38 2.88 0.53 5.27 1.82 0.28

3.66 – 53.6 28845 438 366 57.8 – – 8.26 351 49.0 92.6 15.9 0.114 368 17.4 63.3 10.6 53.9 13.7 3.54 11.9 1.64 10.3 2.01 5.69 0.76 5.29 0.66 3.75 0.81 7.7 0.640 0.312

3.68 – 65.2 26840 626 332 54.4 – – 2.79 255 47.1 104 14.6 b D.L. 155 19.7 68.1 10.9 54.4 14.2 3.47 13.0 1.79 11.0 2.11 5.78 0.72 5.02 0.62 3.86 0.76 6.48 1.414 0.327

3.34 2.15 1.69 1.17 1.18 1.58 2.76 – – – – – – – 43.6 66.4 67.5 75.0 80.8 87.9 51.4 23782 31116 29594 28087 19575 18797 3043 319 747 712 660 549 568 132 289 588 433 163 573 607 193 54.7 60.5 53.8 50.6 53.2 53.3 41.0 – – – – – – – – – – – – – – 6.23 5.26 5.04 6.35 5.17 6.03 2.09 266 334 334 378 177 161 32.9 42.4 54.9 56.5 71.4 55.8 61.5 61.7 141 83.3 87.9 109 113 128 59.9 18.1 14.9 15.0 18.0 13.8 14.3 8.27 b D.L. b D.L. b D.L. b D.L. b D.L. b D.L. 0.018 432 345 343 345 349 413 1.49 23.0 17.2 16.4 15.5 26.3 28.8 40.9 75.91 62.8 66.08 60.33 79.75 84.72 136 11.9 11.0 11.6 10.9 12.3 13.1 19.6 55.6 60.1 61.3 60.7 62.2 67.3 82.6 12.39 16.26 16.5 16.41 17.25 17.56 15.3 3.11 4.11 4.05 4.45 3.69 3.59 1.98 10.4 15.9 14.7 15.3 15.5 17.5 11.2 1.43 2.10 2.02 2.17 2.27 2.47 1.65 8.53 13.2 13.2 14.0 14.2 15.3 11.0 1.59 2.52 2.52 2.59 2.77 2.93 2.13 4.34 7.15 6.75 6.79 7.17 7.81 5.53 0.59 0.95 0.89 0.92 0.97 1.02 0.68 3.99 5.65 5.83 5.82 6.34 6.42 4.2 0.61 0.72 0.7 0.75 0.86 0.87 0.71 4.73 4.06 3.87 4.04 4.76 4.49 3.48 0.9 0.7 0.69 0.68 0.82 0.79 1 7.01 6.31 6.2 5.77 4.84 4.73 1.89 1.411 0.623 0.616 0.506 1.838 2.370 0.540 0.400 0.323 0.247 0.291 0.440 0.424 0.158

Ol-cum: olivine-bearing cumulate; Cum: cumulate; GDr: gabbro-diorite; poik.: poikilitic. b D.L.: below detection limits.

b 0.54 – 66.5 3752 199 302 47.3 – – 2.84 31.0 65.4 72.6 8.36 0.103 3.76 65.1 177 22.1 93.4 19.77 2.61 15.5 2.5 15.7 2.74 7.67 0.97 5.79 0.98 4.74 1.83 4.61 0.525 0.189

3.89 – 57.5 3175 175 94.7 33.9 – – 1.85 23.6 66.0 40.4 6.79 b D.L. 1.84 74.8 213 25.5 88.6 19.19 2.05 15.4 2.5 13.9 2.61 6.77 1.00 5.6 0.83 2.26 1.31 4.11 1.382 0.270

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Table 5 Trace element analyses of plagioclase and biotite Sample

F8

F8

F8

F8

F8b

F8b

F9

F9

F8

F8

F8b

F8b

F8b

LZ56

LZ56

Mineral

Pl

Pl

Pl

Pl

Pl

Pl

Pl

Pl

Bt

Bt

Bt

Bt

Bt

Bt

Bt

relic

relic

interst.

interst.

Lithotype Ol-cum Ol-cum Ol-cum Ol-cum Cum

Cum

Dr dyke Dr dyke Ol-cum Ol-cum Cum

Cum

Cum

GDr

GDr

(ppm) Li Sc Ti V Cr Co Rb Sr Y Zr Nb Cs Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Ta Pb Th U

0.96 1.98 103 0.29 1.05 0.43 0.06 912 0.11 b D.L. 0.01 b D.L. 84.7 4.00 6.55 0.523 1.70 0.155 0.609 0.140 0.002 0.043 0.009 0.022 b D.L. b D.L. b D.L. 0.007 b D.L. 13.0 0.000 b D.L.

b D.L. 2.27 64.6 0.08 b D.L. 0.13 0.95 741 0.89 b D.L. b D.L. b D.L. 165 28.0 34.9 2.52 7.67 0.955 1.60 0.454 0.046 0.144 0.031 0.104 0.004 0.048 0.002 0.182 b D.L. 20.2 0.091 0.020

8.62 2.33 897 15 50.8 70.2 213 74.1 0.025 2.24 5.15 5.80 701 0.022 0.001 0.002 b D.L. b D.L. 0.020 b D.L. b D.L. b D.L. b D.L. b D.L. b D.L. b D.L. b D.L. 0.080 0.458 13.7 0.006 0.004

9.7 2.15 939 13 87.9 82.9 268 89.7 0.019 6.62 4.24 8.17 847 0.010 0.005 b D.L. 0.007 b D.L. 0.038 0.041 b D.L. b D.L. b D.L. 0.015 b D.L. 0.008 b D.L. 0.153 0.546 14.4 0.009 0.008

84.3 12.0 19938 329 299 58.9 386 2.62 0.077 10.9 82.5 2.56 362 0.018 0.010 0.003 0.035 b D.L. b D.L. b D.L. b D.L. b D.L. b D.L. b D.L. 0.006 b D.L. 0.007 0.484 3.540 3.25 b D.L. b D.L.

112 12.2 40154 402 383 70.9 462 4.66 0.109 10.3 44.2 4.15 3068 0.082 0.044 0.004 0.067 0.043 b D.L. 0.256 b D.L. 0.043 0.005 b D.L. b D.L. 0.053 b D.L. 0.480 2.400 4.32 0.022 b D.L.

b D.L. 2.89 136 0.84 1.98 0.22 0.150 1463 0.16 b D.L. b D.L. b D.L. 55.4 2.46 3.94 0.423 1.50 0.273 0.612 0.111 0.018 0.010 0.002 0.019 b D.L. b D.L. 0.008 b D.L. b D.L. 28.8 b D.L. b D.L.

b D.L. 2.77 139 0.65 b D.L. b D.L. 0.15 1042 0.15 0.09 b D.L. b D.L. 68.4 1.90 3.19 0.301 1.34 0.260 0.489 0.103 0.015 0.060 0.003 0.021 b D.L. 0.027 b D.L. b D.L. b D.L. 24.5 0.005 0.005

b D.L. 5.23 14.9 b D.L. b D.L. 0.34 b D.L. 3214 0.07 b D.L. b D.L. b D.L. 190 19.3 16.1 0.748 1.76 0.206 0.969 b D.L. b D.L. b D.L. b D.L. b D.L. b D.L. b D.L. b D.L. b D.L. b D.L. 60.7 b D.L. 1.07

b D.L. 6.11 97.1 0.24 4.53 b D.L. b D.L. 3233 0.17 b D.L. b D.L. b D.L. 317 11.6 9.26 0.584 1.19 0.154 0.981 b D.L. b D.L. b D.L. b D.L. b D.L. b D.L. b D.L. b D.L. b D.L. 0.019 61.3 0.005 5.23

1.36 2.87 69.1 0.17 0.670 0.75 0.130 1342 0.15 b D.L. b D.L. b D.L. 103 7.02 9.40 0.798 2.61 0.258 1.05 0.224 0.004 0.013 0.003 b D.L. b D.L. b D.L. 0.004 b D.L. b D.L. 20.0 0.060 0.014

b D.L. 2.2 45.9 0.15 b D.L. 0.31 0.97 848 0.17 b D.L. b D.L. b D.L. 191 27.8 33.7 2.29 5.77 0.286 2.19 0.123 0.012 0.052 0.018 b D.L. 0.005 b D.L. b D.L. 0.018 b D.L. 21.0 b D.L. 0.015

12.9 5.60 14570 159 107 77.3 367 99.6 0.118 3.11 10.7 5.45 2250 0.086 0.044 0.002 0.021 0.010 0.049 b D.L. 0.002 0.014 b D.L. 0.007 b D.L. b D.L. b D.L. 0.124 0.546 21.5 0.008 0.005

5.75 4.06 11587 107 79.7 78.2 348 126 0.099 4.18 13.7 4.50 2466 0.084 0.096 0.004 0.051 0.024 0.045 b D.L. b D.L. 0.009 0.001 0.013 b D.L. b D.L. 0.002 0.109 0.838 20.8 0.012 0.014

9.22 2.79 1005 18 66.6 82.7 225 83.1 0.116 6.49 5.4 5.62 796 0.203 0.269 0.039 0.094 b D.L. 0.034 b D.L. 0.006 b D.L. 0.006 0.014 0.016 0.005 0.003 0.239 0.483 15.4 0.058 0.025

Ol-cum: olivine-bearing cumulate; Cum: cumulate; GDr: gabbro-diorite; interst.: interstitial. b D.L. : below detection limits.

that these two phases never attained complete equilibrium. Oxygen isotopes lead to the same conclusion, because the Δ18OOpx-Ol values of 0.10 in sample DR22 and of 0.28 in sample F8 are lower than those reported by Mattey et al. (1994) for equilibrium conditions (0.5 ± 0.18). The very low Cr concentrations and textural relations reveal that brown mica also crystallised from a liquid more differentiated than that in equilibrium with clinopyroxene. Although the Mg# of brown mica is rather high, Zanetti et al. (1996) have demonstrated that even higher Mg# would be expected for equilibrium with the clinopyroxene. Chemical data suggest that clinopyroxene relics (and possibly the high Ti-pargasite) in the Ol-cumulates

formed in a different magmatic system not in equilibrium with the melt that gave rise to the dominant cumulate assemblage (olivine + orthopyroxene + poikilitic amphibole + mica + plagioclase). The corroded boundaries of clinopyroxene crystals are also evidence that minerals derived from the first magmatic system reacted with a later melt. It cannot be excluded that the Ca-rich plagioclase cores also belong to the original clinopyroxene-forming magmatic system. Accordingly, the Olbearing cumulates bear evidence of at least two different melts being involved in their petrogenetic process: a “melt 1” that formed the clinopyroxene xenocrysts ± Carich plagioclase, and a “melt 2” in equilibrium with the dominant gabbroic assemblage.

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Fig. 7. REE compositions of minerals in the mafic cumulate assemblage and in the diorite normalised to C1 chondrite. The most primitive clinopyroxene and amphibole from the cumulate exhibit different patterns compared to those of the same minerals in the more evolved rocks.

9.2. “Melt 1” — liquid in equilibrium with the xenocrystic paragenesis Considering the relatively high Mg#, Cr concentrations and absence of negative Eu anomaly in clinopyroxene, “melt 1” is mantle derived and did not undergo extensive fractional crystallisation. Its trace element composition was determined by applying the Amph/LD and Cpx/LD values for the basaltic system reported in Tiepolo et al. (2002) for an average composition of clinopyroxene and brown amphibole. Calculated liquids from the two mineral phases are similar (Fig. 13), characterised by high U–Th concentrations (up to 950 times N-MORB) and UN/ThN values of around 0.5. LREE are strongly enriched relative to HREE (LaN/YbN around 50), which have N-MORB-like concentrations. Nb and Ta show negative anomalies relative to the neighbouring

elements (NbN/LaN = 0.32–0.39), and Zr is slightly enriched relative to Sm (SmN/ZrN = 0.3–0.5). The liquid composition calculated using amphibole differs from the one based on clinopyroxene only for a slight negative Sr and Ti anomalies and slightly lower Ba concentrations. The computed trace element composition of “melt 1” has strong LILE enrichment with respect to HREE and HFSE, and negative Nb–Ta anomalies. However, the LILE, HFSE and LREE concentrations as well as LREE/ HREE fractionation (LaN/YbN up to 50) are inconsistent with those of typical island arc magmas (e.g., McCulloch and Gamble, 1991). An unrealistically high S/LD, inconsistent with the present knowledge on S/LD variation (e.g., Sisson, 1994; Tiepolo et al., 2000), would be required to lower the trace element concentration of the computed melt to that of a typical island arc basalt (IAB). The sub-continental nature of the mantle wedge under the

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Fig. 8. Incompatible trace element composition of selected mineral from the mafic cumulate assemblage normalised to C1 chondrite.

Wilson Terrane may be responsible for the selective Nb, Th, U and LREE enrichments, similar to the primary mantle melts of the Adamello batholith (Italian Alps) reported by Tiepolo et al. (2002). However, compared to the primary mantle melts from the Alpine orogen, the computed liquid for the xenocrystic mineral assemblage of the Teall Nunatak intrusion is still significantly more enriched (up to 10 times) in incompatible trace elements. This peculiar trace element signature could be related to a crustal contribution. In this case the Opx + Grt ± Crd metasedimentary granulites described by Talarico et al. (1995) in the southern Wilson Terrane are assumed to be representative of the assimilated crust. They are characterised by marked LREE enrichment relative to HREE (LaN/YbN = 7.8–17.2), strong Nb depletion and relatively high Ba, Rb and K concentrations. The generally high U and Th concentrations of sediments and particularly their low UN/ThN ratios (Plank and Langmuir, 1998) would also fit with the signature of computed liquids from Teall Nunatak. Nevertheless, the bulk assimilation of the above granulites is unlikely. Even starting from a slightly enriched island arc basalt composition (Tiepolo et al., 2002) and assuming a garnet residuum in the assimilated granulites, the extremely high LaN/YbN ratios and U and Th concentrations of the computed liquid require rela-

tively high degrees of crustal assimilation that contrast with the Cr-rich clinopyroxene and its relatively high Mg#. The extensive assimilation of a metasedimentary sequence could also result in the presence of some detrital zircons, even if zircon could dissolve easily in mafic, Zrundersaturated melts. In the absence of extensive crustal contamination, the sediment signature of “melt 1” is likely inherited from the subducted slab through partial melting of the sedimentary cover. The high Cr and Mg# of “melt 1” would result from mixing and equilibration of the slab melts with liquids originated in the mantle wedge (Tatsumi, 2001). Similar liquids, assumed to have originated by partial melting of the subducted slab's sedimentary cover, are referred to as adakite-type melts and have been reported for the Alpine belt by Tiepolo and Tribuzio (2005). The presence of adakite-type rocks in the Ross Orogen is not a novelty. Small intrusive bodies with adakite signature, dated between 516 ± 10 and 531 ± 10 Ma, were found in southern Victoria Land by Allibone and Wysoczanski (2002). Partial melting of a subducting slab during the initiation of the Ross subduction has been linked (ibid.) to the relatively young and hot nature of the slab due to the presence of a spreading ridge near the margin of the East Antarctic Craton at about 530 Ma.

F. Giacomini et al. / Chemical Geology 240 (2007) 103–128

Fig. 9. a) Oxygen isotope composition of orthopyroxene and olivine from the ultramafic cumulate enclaves at Teall Nunatak (and from Mt. Abbott gabbro for comparison). The low Δ18Oopx-ol suggests disequilibrium between the mineral phases and the high δ18O values are indicative of significant crustal contamination; b) δ18O olivine vs. MgO of olivine plot, showing the possible effects of interaction with hydrous melt on the δ18O value of olivine. Data for the Abbott gabbros are from Dallai et al. (2003).

121

Nunatak (Rocchi et al., 1998). Because olivine, orthopyroxene, brown mica and plagioclase hold just negligible trace element concentrations – leading to major uncertainties in their solid/liquid partition coefficient – only the poikilitic brown amphibole was considered in the liquid computation, for which the Amph/L D for the basaltic system reported by Tiepolo et al. (2002) were adopted. Comparison between “melt 1” and “melt 2” shows that the latter has considerably lower U and Th concentrations (up to 300 times N-MORB), UN/ThN ratio of about 1.3 (melt 1 has UN/ThN = 0.5) and significantly lower LREE/HREE fractionation (LaN/YbN = 16) (Fig. 13). Other major differences are the less pronounced depletion in Nb–Ta relative to LREE and LILE (NbN/LaN = 0.68) and the higher HREE concentrations (2–3 times N-MORB). No Zr enrichment with respect to Sm (SmN/ZrN = 1.8) is apparent. Judging from the abundance of mica and amphibole, the liquid in equilibrium with the gabbroic assemblage must have also been strongly hydrated and K-rich. Because of the demonstrated extensive fractional crystallisation and crustal assimilation, the trace element signature of the calculated liquid assumed to be in equilibrium with the cumulate assemblage cannot provide first-order information about the mantle source. However, its petrologically relevant ratios (e.g., La/Yb, Zr/Sm) and trace element patterns are similar to those of the whole rock analyses of the more evolved gabbro-diorites. There are also strong similarities with the whole rock composition of the Ol-bearing gabbro from Mt. Abbott reported by Dallai et al. (2003). If the trace element composition is considered, the differences in Sr, Eu, Ti and Ba between amphibole, mica and orthopyroxene from the gabbro-diorite and the cumulate rocks can be ascribed to the fractional crystallisation of mica + plagioclase + amphibole. An exotic input is required only to explain the higher LREE enrichment over HREE observed in amphibole from

9.3. “Melt 2” — liquid in equilibrium with the dominant cumulate assemblage Based on the chemical composition of olivine and orthopyroxene (e.g., Fo70; Mg#76), the “melt 2” in equilibrium with the dominant cumulate assemblage must have experienced a high degree of fractional crystallisation and significant crustal contamination. The high δ18O values for orthopyroxene (7.13–7.47‰) and olivine (7.03–7.19‰) confirm that crustal contamination occurred, as also supported by the high 87Sr/86Sr (0.70901–0.70945) and low εNd (from − 6.9 to − 7.4) values reported for gabbro-dioritic rocks from Teall

Table 6 Oxygen isotope analyses of mineral separates Sample

DR22

F8

F8b

LC4

Lithotype

Ol-cum

Ol-cum

Cum

GDr

7.03 7.13 9.2

7.19 7.47 8.8

– 6.74 9.0

– 7.60 8.8

(‰) δ18OOl δ18OPx δ18OPl

Ol-cum: olivine-bearing cumulate; Cum: cumulate; GDr: gabbrodiorite.

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Table 7 LAM-ICP-MS U–Pb isotope analyses and calculated ages of zircons Sample

Lithotype

Zircon#

207

Pb/235U

1σ

206

Pb/238U

1σ

Concordant age

2σ

DR22 DR22 DR22 DR22 DR22 DR22 DR22 DR22 DR22 DR22 DR22 DR22 DR22 DR22 DR22 DR22 02123 LC4 LC4 LC4 LC4 LC4 LC4 LC4 LC4

Ol-cum Ol-cum Ol-cum Ol-cum Ol-cum Ol-cum Ol-cum Ol-cum Ol-cum Ol-cum Ol-cum Ol-cum Ol-cum Ol-cum Ol-cum Ol-cum Std GDr GDr GDr GDr GDr GDr GDr GDr

1-2a 1-2b 4-1b 2-2 4-5a 1-5 4-3 2-11 4-1a 2-4a 3-3b 2.4b 3-3a 4-2a 4-2b 4-5b – c23 c5 c22 c30 c28 c20 c34 c19

0.7070 0.6860 0.6691 0.6663 0.6709 0.6764 0.6634 0.6620 0.6564 0.6375 0.6425 0.6385 0.6605 0.7264 0.6538 0.6568 0.3389 0.6290 0.5902 0.6330 0.6463 0.6186 0.5914 0.6490 0.6131

0.0104 0.0104 0.0111 0.0102 0.0114 0.0103 0.0099 0.0099 0.0100 0.0112 0.0104 0.0107 0.0102 0.0119 0.0098 0.0103 0.0060 0.0227 0.0246 0.0312 0.0225 0.0198 0.0234 0.0238 0.0190

0.0890 0.0846 0.0845 0.0848 0.0843 0.0838 0.0844 0.0844 0.0838 0.0821 0.0806 0.0802 0.0812 0.0859 0.0846 0.0851 0.0478 0.0802 0.0804 0.0794 0.0786 0.0785 0.0783 0.0781 0.0780

0.0010 0.0010 0.0010 0.0010 0.0010 0.0010 0.0010 0.0010 0.0010 0.0010 0.0009 0.0010 0.0010 0.0010 0.0010 0.0010 0.0006 0.0007 0.0006 0.0008 0.0006 0.0007 0.0007 0.0006 0.0006

546 526 523 522 522 521 520 519 516 507 501 498 – – – – 299 497 497 493 488 487 486 485 484

12 11 12 12 12 11 11 11 11 12 11 12 – – – – 6 8 7 10 7 8 8 7 7

Ol-cum: olivine-bearing cumulate; GDr: gabbro-diorite; Std: standard.

diorite. The gabbro-diorites and the dominant assemblage in the cumulate enclaves most likely derive from the same parental liquid, but represent different degrees of differentiation through a process of crustal assimilation and fractional crystallisation (AFC). In summary the main mafic intrusive sequence at Teall Nunatak comprises variably differentiated, and contaminated mantle-derived rocks. Similarly to the intrusives of the neighbouring Terra Nova Intrusive Complex (Rocchi et al., 2004), their bulk rock composition reflects as a whole those of high-K calc-alkaline sequences of active continental margins, where mantle-derived melts are likely to strongly interact with almost coeval crustalderived products. 9.4. Evidence for a change in the style of magma production (implications for the evolution of the Ross Orogeny) The chemical disequilibria observed in the ultramafic cumulate enclaves from Teall Nunatak point to the involvement of different melts in their genesis and suggest that the evolution of the magmatic activity during the Ross orogenic event in the region is characterised by multiple mantle-derived magmatic pulses.

The parental melts computed from the two mineral assemblages in the enclaves (the xenocrystic paragenesis and the dominant cumulate assemblage) show some compositional similarities. However, simple fractional crystallisation of olivine + amphibole + brown mica + orthopyroxene + plagioclase cannot account for the difference in LREE concentrations and in particular for the lower La/Yb ratios and the very different U/Th ratio of the “melt 2”, compared with “melt 1”. Indeed, the crystallisation of the above mineral phases would lead to higher La/Yb ratios in the “melt 2”. Furthermore, S/LD similarities for Th and U cannot account for the threefold difference in Th at almost constant U concentrations. Starting from “melt 1”, and assuming a process of assimilation and fractional crystallisation (AFC), the contaminant must have been strongly depleted in LREE with respect to HREE and must have had a high U/Th ratio to produce the “melt 2” in equilibrium with the cumulate assemblage. This is in contrast with the composition of the different granulite types reported by Talarico et al. (1995), considered to be representative of the crust in the Wilson Terrane. If the two liquids are not akin to one another, then they must represent distinct magmatic pulses generated from different sources.

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Fig. 10. Cathodoluminescence images of zircons from the two samples selected for U–Pb geochronology. Spot location and age results are reported for each zircon right on the images. The cited errors for each age determination are given as 2σ.

Combining the geochemical and geochronological data presented here, we propose the following model for the evolution of the Ross magmatic activity in this region. An early magmatic event produced mantle melts with a strong sediment fingerprint, likely inherited from the

subducted slab (“melt 1”). The older prismatic zircons in the Ol-bearing cumulates clustering around a mean age of 521 ± 2 Ma constrain the timing for this magmatic pulse. Up to now, such old mafic magma pulses have no equal in North Victoria Land and could represent

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Fig. 11. a) Concordia diagram of zircons from the DR22 cumulate: the four discordant points (light grey) plot close to the concordia, but were excluded from mean age calculations. b), c) Concordia diagrams with the mean concordia ages of pooled zircon fractions from the DR22 cumulate. d) Concordia diagrams with the mean concordia age of zircons from the LC4 diorite. Error ellipses are shown at 2σ.

analogues of the small gneissic intrusions with adakite affinity spanning in age from 516 ± 10 to 531 ± 10 Ma and interpreted by Allibone and Wysoczanski (2002) as predating the main “Cordilleran style” magmatism. This early mantle magmatism indicates a regional-scale process and marks the inception of the oceanic subduction under the palaeo-Pacific margin of Gondwanaland. The mantle melts likely provided a substantial amount of heat to the crust, triggering the subsequent granitedominated igneous activity. These early plutonic products were intruded and dismembered by younger batches of mantle magma (“melt 2”): the original geochemical signature of the younger melts was strongly modified by the extensive fractionation and crustal contamination. The study of equilibrium liquids allows to exclude an adakite-type signature for

“melt 2” and suggests an origin from a different and more “conventional” source. The 502 ± 3 Ma age of acicular zircons in Ol-bearing cumulates is interpreted to represent the initiation of the second magmatic pulse. It is likely that through crustal assimilation, fractional crystallisation (AFC), enhanced by interaction with almost coeval crustal melts, the liquids with “melt 2” affinity evolved in wet conditions towards the more differentiated diorites. The emplacement age at shallow crustal levels for this second potassic mafic episode is constrained at 489 ± 4 Ma, in agreement with the radiometric ages defined by Allibone and Wysoczanski (2002) and Bomparola et al. (2007) on similar high-K calc-alkaline to shoshonitic mafic intrusives in a large region extending from the Dry Valleys (South Victoria Land) to the Deep Freeze Range (North Victoria Land).

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Fig. 12. Solid/solid partitioning between the two different amphibole (rimming clinopyroxene and poikilitic) in Ol-bearing cumulates and clinopyroxene. The comparison with equilibrium values from the literature (Tiepolo et al., 2000, 2001; Rivalenti et al., 2004 and references therein) suggests that Ti-rich pargasite and clinopyroxene approach equilibrium conditions.

The occurrence of old intermediate-felsic intrusions with ages in the range 550–530 Ma, common in South Victoria Land and in the Central Transantarctic Mountains, is still matter of debate for the North Victoria Land region. A possible diachronism of the magma production between North Victoria Land and the other sectors of the belt, could be related to local microplate accretion preceding the main continental collision (e.g. the “Beardmore microcontinent” of Borg et al., 1990; Stump et al., 2006) or to oblique plate convergence. However, further investigations are necessary to exclude the occurrence in North Victoria Land of old granitoids, and to test this hypothesis.

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2) The clinopyroxene from the first-generation melt (“melt 1”) has high Cr concentrations and a relatively high Mg# suggesting equilibration with a liquid that had not been altered by crustal processes. This liquid composition, computed by applying S/LD to clinopyroxene, exhibits exceptionally high La/Yb ratios and Th–U concentrations, inconsistent with a typical island arc magma. Partial melting of the subducted sediments from the slab was most likely involved in the petrogenetic process of these liquids. 3) The oxygen isotope signature and the composition of the mineral phases in the dominant cumulate assemblage suggest that the parental liquid was significantly contaminated by the crust and underwent extensive fractional crystallisation. Its trace element compositions therefore are incapable of providing much useful information about the mantle source. Chemical data on the bulk rock and mineral phases suggest that the cumulate assemblage of the mafic/ultramafic enclaves and the dominant gabbro-diorites were derived from the same parental liquid through a process of crustal assimilation and fractional crystallisation (AFC). 4) The parental liquid (“melt 2”) of the cumulate assemblage was not derived from the liquid in equilibrium with the xenocrystic mineral assemblage (“melt 1”). Different magmatic pulses must have originated from different sources, implying a change in the style of magma production through time during the Ross orogenic event.

10. Summary and conclusions The Teall Nunatak plutonic complex in North Victoria Land consists of gabbro-dioritic and granitoid rocks. The gabbro-diorites preserve mega-enclaves of mafic/ultramafic cumulates with mineralogy that reveals a great deal about the chemical characteristics of the earliest liquids to form in this area in relation to oceanic subduction in the Ross orogen. The major oxide, trace element and oxygen isotopes compositions of minerals in the system as well as the zircon U–Pb geochronology support the following major conclusions. 1) In Ol-bearing cumulates, corroded clinopyroxene (± corroded Ca-rich plagioclase) represents relics of a mineral assemblage that crystallised from an early melt, different from the one that generated the dominant gabbroic assemblage (Amph + Bt + Opx + Pl).

Fig. 13. Incompatible trace element composition of calculated equilibrium liquids normalised to N-MORB. The melts in equilibrium with the most primitive mineral assemblage are significantly enriched in Th and U and strongly depleted in HREE with respect to those calculated from the more evolved minerals. This suggests the involvement of two different magmatic sources in the petrogenesis of the cumulate enclaves. Dotted and dashed lines represent the trace element compositions of one analysed mafic dyke and of the gabbrodioritic rocks (average of 15 samples), respectively.

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5) Zircon U–Pb geochronology shows that the first magmatic pulse at 521 ± 2 Ma predates by about 20 Ma the gabbro-diorites emplaced between 489 ± 4 Ma and 502 ± 8 Ma. This finding testifies for the occurrence in North Victoria Land of mantle melts related to the initiation of the Ross subduction. We suggest that the change from adakite-type to more conventional K-rich compositions is associated with: (a) the evolution from a relatively young and hot subducting slab to a more mature and cold one, and (b) to the increased interactions of the mantle magma with crustal derived melts, whose main emplacement phase occurred 510–490 Ma. 6) The occurrence of gneissic intrusions older than 520 Ma also in North Victoria Land, suggested by few and not well constrained geochronological data, raises the problem of a potential pre-“Granite Harbour” magmatism: a stimulating target for further research in the region. Acknowledgments Thanks are due to Riccardo Tribuzio for the helpful suggestions in the field and during manuscript preparation. Alberto Zanetti and Rosa Maria Bomparola are acknowledged for the constructive discussion and Marco Palenzona for his technical assistance during analysis with the LA-ICP-MS. Samuel Mukasa, an anonymous reviewer and the Editor Steven Goldstein are gratefully acknowledged for their careful and constructive reviews: all remaining mistake is down to us. This work was carried out with the financial and technical support of the Italian Programma Nazionale Ricerche in Antartide (P.N.R.A.). References Allibone, A., Wysoczanski, R., 2002. Initiation of magmatism during the Cambrian–Ordovician Ross orogeny in southern Victoria Land, Antarctica. Geological Society of America Bulletin 114, 1007–1018. Allibone, A.H., Cox, S.C., Graham, I.J., Smillie, R.W., Johnstone, R.D., Ellery, S.G., Palmer, K., 1993a. Granitoids of the Dry Valleys region, southern Victoria Land, Antarctica: geochemistry and evolution along the early Paleozoic Antarctic margin. New Zealand Journal of Geology and Geophysics 36, 299–316. Allibone, A.H., Cox, S.C., Smillie, R.W., 1993b. Granitoids of the Dry Valleys region, southern Victoria Land, Antarctica: plutons, field relations and isotopic dating. New Zealand Journal of Geology and Geophysics 36, 281–297. Armienti, P., Ghezzo, C., Innocenti, F., Manetti, P., Rocchi, S., Tonarini, S., 1990. Isotope geochemistry and petrology of granitoid suites from Granite Harbour intrusives of the Wilson Terrane, North Victoria Land, Antarctica. European Journal of Mineralogy 2, 103–123. Asprey, L.B., 1976. The preparation of very pure F2 gas. Journal of Fluorine Chemistry 7, 359–361.

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