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Journal of Petrology | Volume 44 | Number 11 | Pages 2053-2080 | 2003
© Oxford University Press 2003; all rights reserved
Phonolitic Diatremes within the Dunedin Volcano, South Island, New Zealand
1 SCHOOL OF SCIENCE AND TECHNOLOGY, THE UNIVERSITY OF WAIKATO, PRIVATE BAG 3105, HAMILTON, NEW ZEALAND
2 DEPARTMENT OF GEOLOGY, THE UNIVERSITY OF OTAGO, PO BOX 56, DUNEDIN, NEW ZEALAND
3 SCHOOL OF EARTH SCIENCES, UNIVERSITY OF MELBOURNE, VIC. 3010, AUSTRALIA
4 SCHOOL OF GEOSCIENCES, MONASH UNIVERSITY, CLAYTON, VIC. 3800, AUSTRALIA
* Corresponding author. E-mail: r.price{at}waikato.ac.nz
RECEIVED SEPTEMBER 13, 2002; ACCEPTED MAY 19, 2003
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONTHE PORT CHALMERS BRECCIAPETROGRAPHYMINERAL CHEMISTRYWHOLE-ROCK MAJOR AND TRACE...ISOTOPE GEOCHEMISTRYDISCUSSIONCONCLUSIONSREFERENCES |
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The Port Chalmers Breccia is a vent-filling, clastic volcanic unit exposed within the Miocene Dunedin Volcano of South Island, New Zealand. Clasts (up to in excess of 1 m but generally <20 cm) are supported in ash and fine lapilli of phonolitic (ne-benmoreite or tephro-phonolite) composition and the dominant clast type (55 to almost 100%) is also phonolitic. Less abundant lithologies include ne-normative basalt (basanite), hawaiite, mugearite and trachyandesite, syenites and microsyenites, coarse-grained mafic (gabbros) and ultramafic rocks (pyroxenites, hornblendites), schists and sediments. The breccias were emplaced as diatremes associated with localized, but highly explosive, eruptive events in which mantle-derived CO2 was an important component. The syenitic and ultramafic clasts could represent intrusive suites produced by crystal fractionation acting on parental ne-benmoreite magmas that may themselves have been derived by crystal fractionation from basanitic precursors. An alternative variation on this model is that the parental ne-benmoreites were generated through partial melting of an alkalic igneous underplate. Sr, Nd and Pb isotopic compositions are strikingly similar to those of intraplate igneous rocks, ranging in age from 100 to less than 10 Ma, from elsewhere in the South Island, and New Zealand's sub-Antarctic islands, the south Tasman Sea and the Ross Sea region. This regional, HIMU-influenced, isotopic signature is believed to be derived from within the lithospheric mantle.
KEY WORDS: phonolite; diatreme; nepheline syenite; Dunedin Volcano; alkalic rocks; fractional crystallization
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONTHE PORT CHALMERS BRECCIAPETROGRAPHYMINERAL CHEMISTRYWHOLE-ROCK MAJOR AND TRACE...ISOTOPE GEOCHEMISTRYDISCUSSIONCONCLUSIONSREFERENCES |
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The city of Dunedin on the SE coast of New Zealand's South Island (Fig. 1) is located within a complex alkalic volcano of Miocene age. Coombs et al. (1986)
defined the Dunedin Volcanic Group to include rocks of the volcano and other similarly aged alkalic eruptives and shallow intrusives in east Otago. Within the Dunedin district, rocks of the Dunedin Volcanic Group are exposed over 450–500 km2 with vertical relief of around 700 m. Most of the Dunedin Volcano was constructed between 13 and 10 Ma (McDougall & Coombs, 1973) with the earliest activity being shallow marine eruptions of silica-saturated trachytes and basalts and the most recent events the subaerial emplacement of phonolitic domes and flow domes.
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The vast majority of Dunedin volcanic rocks are nepheline-normative (ne-) basalts and phonolites (Benson, 1942
, 1968; Coombs et al., 1960; Coombs & Wilkinson, 1969; Allen, 1974; Price & Chappell, 1975). The volcanic stratigraphy within the volcano was divided by Benson (1959, 1968) into four phases (Initial, First, Second and Third), based on inferred regional correlations of inter-volcanic sedimentary sequences interpreted as representing periods of erosion and reduced volcanic activity.
The focus of this paper is a remarkable series of volcanic breccias collectively referred to as the Port Chalmers Breccia (PCB) (Fig. 2b). Compared with the rest of the Dunedin Volcano the PCB contains a wider variety of rock types in the form of fine- and coarse-grained fragments comprising basalt, basanite, ne-trachyandesite, ne-benmoreite and phonolite, hornblendites, gabbros and nepheline syenites, together with fragments of schist and non-volcanic clastic sediments. The igneous clasts preserve an extended petrological history of processes that occurred in the evolution of an undersaturated (basanite–phonolite) magmatic series and permit fuller isotopic comparisons with data for rocks from the distinctive Tasman– Balleny–New Zealand isotopic province (Lanyon et al., 1993; Baker et al., 1994).
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THE PORT CHALMERS BRECCIA |
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TOPABSTRACTINTRODUCTIONTHE PORT CHALMERS BRECCIAPETROGRAPHYMINERAL CHEMISTRYWHOLE-ROCK MAJOR AND TRACE...ISOTOPE GEOCHEMISTRYDISCUSSIONCONCLUSIONSREFERENCES |
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The PCB is exposed in vents aligned NW–SE from Port Chalmers, across the central Otago Harbour into Hoopers Inlet on the Otago Peninsula (Fig. 2a). The most extensive exposure (1·3 km x 2·5 km) occurs at Port Chalmers on the northern side of Otago Harbour (Fig. 2). Benson's 1968 map shows other roughly circular vents containing PCB located on Otago Peninsula and ranging in diameter from 160 to 650 m. Coombs (1965)
noted that the explosive eruptions that were associated with the emplacement of the PCB were regarded by Benson (1942) as occurring at the end of the First Main Eruptive Phase, preceding the main constructional event of the Second Main Eruptive Phase (Coombs et al., 1960; Benson, 1968). Our primary focus has been on the most extensive exposures of PCB around Port Chalmers (northern locality; Fig. 2) where we have examined and sampled two specific exposures (Leans Rock and Scott Memorial; Fig. 2b) in considerable detail. We have also collected data and samples from a locality on the Otago Peninsula (southern locality; Fig. 2a).
In the vicinity of Port Chalmers, the PCB was emplaced through microsyenites, and dolerites that show hydrothermal alteration associated with abundant carbonate (calcite and ankerite) veins. Allen (1974) described exposures in a railway tunnel cutting the northern boundary of the Port Chalmers exposures, which provide evidence that at least some of the breccia was erupted onto an eroded surface. Dykes of basalt and trachyandesite cut the dolerites and microsyenites underlying the PCB but not the PCB itself (Fig. 2b), indicating that they were emplaced before the breccia-forming eruptions. In turn, the PCB was intruded by a later phase of basaltic dykes. At the southern locality, the emplacement occurred through trachytic tuffs of the Initial Eruptive Phase (Benson, 1968; Allen, 1974).
The PCB shows a range of facies types from the dominant massive, matrix-supported breccia to rare bedded units in which layers of fine lapilli vary in thickness from 2 cm to 1–2 m. Within the Port Chalmers exposure bedding dips to the west, with the angle of dip decreasing from around 40° at the margins of the vent near inferred contacts with underlying eroded flows to around 16° within the breccia pipe (Allen, 1974, and this work). Neither cross-bedding nor graded bedding have been observed in any of the PCB exposures.
The clasts in the PCB are matrix supported, with around 80% of a typical outcrop being matrix finer than 1 cm and the largest clasts up to 1·6 m. At the southern locality most of the clasts (almost 100%) are phonolitic (ne-benmoreite or phonolite). Phonolitic types are dominant (56%) at the northern locality but clasts of basalt and dolerite (12%), trachyandesite (9%), syenite (9%), gabbro and ultramafic cumulate (
1%), schist (12%), and Tertiary or Cretaceous sediment (1%) also occur. Most phonolitic clasts show thin (1–5 mm) bleached rims and the matrix of the PCB is cemented by carbonate, analcime, rare alkali feldspar and kaolinite, suggesting that hydrothermal alteration was pervasive following eruption and deposition. Clasts with the characteristics expected of juvenile material (e.g. glassy rims, strong vesiculation, jig-saw disaggregation) are rare. The obvious candidates are pale coloured devitrified glassy volcanic rocks containing abundant carbonate- and analcime-filled vesicles and occurring either as discrete clasts up to 1–2 cm in diameter or as occasional thin discontinuous rims on syenite fragments.
The level of exposure of the PCB is close to the base of the volcanic sequence, as inliers of the underlying sedimentary sequence are known in both localities (e.g. Coombs et al., 1960). Also, close to the breccia outcrops at Port Chalmers the country-rock trachytic tuffs are cut by thin, fine-grained dykes of foraminiferal limestone, forced up from underlying sediments by lithostatic loading (P. Gurney, personal communication, 1992). The clasts in the PCB are, however, dominantly phonolitic and rocks of this type are not a significant component among the earliest eruptives of the Dunedin Volcano, which are dominantly basalts and quartz-normative trachytes. Because juvenile clasts are rare, it is unlikely that the abundant phonolitic clasts in the breccias represent disrupted juvenile magma. If the phonolitic clasts are accidental then their abundance suggests that they have come from above the present levels of exposure in the PCB and this could only be the case if the breccias are significantly younger than has previously been supposed. Rather than being part of the First Eruptive Phase (Benson, 1968), the PCB must post-date at least the First and possibly the Second Main Eruptive Phase and have formed relatively late in the history of the Dunedin Volcano.
The garnet–albite zone metamorphic grade represented by breccia schist clasts is the same upper greenschist facies assemblage as that observed in basement exposures in the Dunedin district. Higher-grade assemblages are not observed.
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PETROGRAPHY |
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TOPABSTRACTINTRODUCTIONTHE PORT CHALMERS BRECCIAPETROGRAPHYMINERAL CHEMISTRYWHOLE-ROCK MAJOR AND TRACE...ISOTOPE GEOCHEMISTRYDISCUSSIONCONCLUSIONSREFERENCES |
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Petrographic aspects of volcanic rocks of the Dunedin Volcano and clasts from the PCB have been described in detail by Benson (1942)
, Coombs & Wilkinson (1969), Allen (1974) and Price & Chappell (1975). The petrography of representative rock types is summarized in Table 1. The unique feature of the PCB, in contrast to other units of the Dunedin Volcano, is the presence of coarse-grained rocks showing relatively extreme compositions. Ultramafic rocks have textures indicating a cumulate origin and this is also the case for some coarse syenites. Gabbroic rocks and some syenites appear, however, to be coarse-grained equivalents of basanitic and phonolitic volcanic rocks. They do not show textures or chemical compositions consistent with a cumulate origin. It should also be emphasized that the modal compositions of ultramafic cumulate rocks are dominated by pyroxene and amphibole. Olivine has not been observed in these rocks.
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MINERAL CHEMISTRY |
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TOPABSTRACTINTRODUCTIONTHE PORT CHALMERS BRECCIAPETROGRAPHYMINERAL CHEMISTRYWHOLE-ROCK MAJOR AND TRACE...ISOTOPE GEOCHEMISTRYDISCUSSIONCONCLUSIONSREFERENCES |
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Methods
Minerals were analysed in carbon-coated polished thin sections on the University of Otago's JEOL JXA-8600 electron microprobe using wavelength-dispersive techniques. Operating conditions used an accelerating voltage of 15 kV, a specimen current of 2 x 10-8 A, and a beam diameter ranging from 1–2 µm for stable minerals to broad beam for feldspathoids, feldspars and glass. Pure compounds and natural minerals were used as standards. Raw counts were corrected by ZAF procedures.
Pyroxene
Representative pyroxene compositions are listed in Table 2 and pyroxene compositional variation, in terms of Diopside–Hedenbergite–Acmite (Aegirine) components is illustrated in Fig. 3a.
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The pyroxenes of the mafic clasts are reasonably uniform in composition and, according to the classification scheme of Morimoto (1989)
, are augites, diopsides, or their titanian or aluminous equivalents (Wo48–50, En34–44, Fs6–17; TiO2 0·4–5·1 wt %, Al2O3 up to 7·7%). In contrast, phonolitic rocks (e.g. PCB20) contain a bimodal pyroxene compositional population (Fig. 3a), often with green and colourless microphenocrysts occurring side by side in a microcrystalline or devitrified glassy groundmass. Green phenocrysts, with patchy or concentric zoning, are acmitic (acm 18–22 mol %), low in Al and Ti, but contain appreciable Zr (ZrO2 0·42–0·52%). Coexisting acmite-poor phenocrysts are sector and concentrically zoned, with prism sectors characterized by an aluminium-bearing diopside (e.g. PCB25 No. 3, Table 2) and basal sectors by a low Al–Ti ferroan hedenbergite (e.g. PCB25 No. 2, Table 2). This type of disequilibrium crystallization feature has been reported in a wide variety of igneous rocks (e.g. Nakamura & Coombs, 1973; Dowty, 1976; Cooper, 1986; Shearer & Larsen, 1994) and is ascribed to control by crystallographic, proto-site configurations.
In phonolites (e.g. PCB20) and syenites (e.g. PCB53) some diopside grains carry a pale green, rounded, Na–Fe-enriched core. In the analysed syenite, the green core is a ferrian magnesian hedenbergite (PCB53 No. 3, Table 2) rimmed by diopside (PCB53 No. 6) and then by the more acmitic ‘groundmass’ phase, which is a ferroan magnesian aegirine–augite (PCB53 No. 7). These ‘green-cored’ pyroxenes are widely reported in nepheline-normative mafic and intermediate rocks of southern New Zealand (Price, 1973; Cooper, 1979; Brodie & Cooper, 1989) and from elsewhere (e.g. Brooks & Printzlau, 1978; Bédard et al.,1988; Neumann et al., 1999).
Pyroxenes in feldspathoidal syenite have an irregular or sub-ophitic texture, with colour zoning in places truncated at grain boundaries. Compositions in the feldspathoidal syenite clasts of the PCB show a much broader range than has been observed in the rest of the Dunedin Volcanic Group (Fig. 3b). They are typically ferroan aegirine–augites (PCB1 No. 12, Table 2). These cores are overgrown by oscillatory zones characterized by an overall chemical trend towards increasingly more sodic and Zr-rich compositions (e.g. PCB1 No. 16). The fractionation trend culminates in a brown-coloured patchy overgrowth of a Mn-rich aegirine (PCB1 No. 8, Table 2). Although Zr contents correlate crudely with the abundance of the acmite component, the maximum ZrO2 content of 2·48 wt % is attained in a pyroxene with an acmite content of a little over 80% (PCB48 No. 6, Table 2), and some of the most acmitic pyroxenes are pale green and low in ZrO2 (e.g. PCB48 No. 9, Table 2). Aegirines, although generally low in Ti, can contain appreciable TiO2 (up to 2·15 wt %; e.g. PCB9 No. 35, Table 2).
The pyroxenes of the clasts in the PCB collectively define an extended trend, similar to that observed in peralkaline undersaturated intrusive suites in south Greenland (Qôroq—Stephenson, 1972; Ilímaussaq—Larsen, 1976; Marks & Markl, 2001) and at Chinduzi, Chilwa (Woolley & Platt, 1988).
Amphibole
Representative amphibole compositions are listed in Table 3. From mafic to felsic clast types there is a progressive change in amphibole chemistry from calcic to sodic. Within the mafic cumulates and gabbros the amphiboles are generally compositionally homogeneous, varying with rock type from kaersutites to ferro-kaersutites and titanian ferroan pargasites [classification of Leake et al. (1997)]. TiO2 abundance ranges from 2·71 to 5·78 wt %, Na2O is typically 2·90% and K2O 1%. Mg number [= Mg/(Mg + Fe)] ranges from 0·670 to 0·390. NaB values are low, with a maximum content of 0·22 cations per formula unit (p.f.u.).
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In the phonolites, amphiboles are kaersutites, titanian ferroan pargasites and potassian titanian hastingsites. Some of the ultramafic cumulates contain amphibole with similar compositions to that in phonolite, suggesting cumulates may relate to crystallization from evolved magmas. However, the extended range in phonolite amphiboles, principally towards lower Mg number and TiO2, probably reflects continued fractionation beyond the stage recorded in amphiboles from analysed cumulates.
Syenite amphiboles range from ferro-kaersutites, hastingsites and pargasites in PCB5, PCB42 and OU 24259, through kataphorites in PCB51 and PCB53, to ferric-ferronyobite in OU 242273. The spectrum is marked by a decrease in Mg number from 0·477 to 0·092 and TiO2 from 5·13 to 0·25 wt %, and increases in SiIV (from 6·027 to 7·505 cations p.f.u.), Na2O (up to 8·68 wt %), K2O (up to 1·86 wt %), and NaB (from 0·274 to 1·000). Weak chemical zoning in individual grains of the most evolved ferric-ferronyobites of OU 24273 mimics the general trend described above, with the exception that there is a slight increase in Mg number from core to rim. MnO is markedly enriched in these evolved ferric-ferronyobites, with contents reaching 3·57 wt %, reflecting the behaviour of Mn across the amphibole spectrum, and resulting in an analogous enrichment in Mn to that described for late-stage acmitic pyroxenes in evolved Dunedin nepheline syenites.
The compositional trends defined by amphibole compositions of the Dunedin Volcano are very similar to those in undersaturated plutons of the Gardar Province, Greenland (see Mitchell, 1990).
Biotite
A trioctahedral biotite mica occurs as a minor phase in mafic and phonolitic clasts and is abundant in some of the nepheline syenites. Compositionally, micas are phlogopite–annite solid solutions, there being no octahedral Al to form an eastonite or siderophyllite component (Table 4). Many of the biotites show marked deficiencies in the tetrahedral site (up to 0·4 cations p.f.u., Table 4) and significant Fe3+ is presumably required to fill the site (Rieder et al., 1998). This is a characteristic of biotites from undersaturated rocks (e.g. Cooper, 1979). As with pyroxenes and amphiboles, there is a general decrease in the Mg number of biotite from the gabbros (0·423) to nepheline syenites (0·029). There is a correlation between Ti (cations p.f.u.) and Mg number, with Ti being most abundant (up to 6·84 wt % TiO2) in biotites of the mafic rocks (Fig. 4a).
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In Fig. 4b, the compositions of PCB biotites have been compared in terms of Al, Mg and Fe2+ contents with those from syenites from elsewhere. Biotites from Port Chalmers nepheline syenite clasts are relatively iron rich but they extend the trends defined by biotite compositions from syenites of Malawi (Fig. 4b). Al enrichment trends observed in biotites of the Ilomba and Ulindi nepheline syenites of the North Nyasa alkaline province are not present in the biotites of the PCB nepheline syenite clasts.
Fe–Ti oxides
Representative oxide compositions are shown in Table 5. Magnetite is ubiquitous in clasts of the PCB and in most rocks it is homogeneous and does not show exsolution. Phonolites and nepheline syenite clasts show a similar range of magnetite compositions from Usp63Mt37 to Usp10Mt90 (Fig. 5). Magnetites in the mafic rocks show a much more limited compositional range and they tend to be more Ti rich (Usp44Mt56 to Usp94Mt6). Chromium and MgO contents are generally low. MnO abundances are consistently higher than MgO contents, with magnetites in nepheline syenites containing up to 4·4 wt % and those in ultramafic cumulates up to 2·3 wt % MnO. The pattern of compositional variation in magnetites from clasts in the PCB reflects that shown within the Dunedin Volcanic Group. Dunedin phonolites contain magnetites that are generally less Ti rich than those observed in the basanites and basalts (Price, 1973).
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Exsolved ilmenite in the nepheline syenites is Ti rich (Ilm90–96) and application of the Ghiorso & Sack (1991)
oxide geothermometer and oxygen geobarometer to a coexisting spinel–ilmenite pair in syenite PCB42 gives oxygen conditions close to the FMQ (fayalite–magnetite–quartz) buffer (918°C and logfO2 = -12·2). This estimate is consistent with the more oxidized end of the range for nepheline syenites from elsewhere (e.g. Ilímaussaq, Larsen, 1976; Marks & Markl, 2001), with conditions from FMQ to more reducing being commonly estimated. The presence of aenigmatite in phonolitic clasts of the PCB is also an indication of relatively low fO2 (Thompson & Chisholm, 1969; Lindsley, 1970; Hodges & Barker, 1973).
Feldspar
Feldspars occur in all clasts of the PCB except a few of the more extreme ultramafic cumulate compositions. Compositional variation in the feldspars is illustrated in Fig. 6. Plagioclase in a dolerite (PCB57) ranges in composition from calcic cores (An65Ab34Or1) to sodic rims (An11Ab78Or21). Alkali feldspar (An1Ab39Or60) is present in the outermost rim of zoned grains and in the groundmass. The feldspars in phonolitic rocks are dominantly anorthoclase (e.g. PCB25: An3Ab54Or43; PCB20: An4Ab63Or33) but a few intermediate plagioclase phenocrysts (e.g. PCB20: An28Ab68Or4) are present in most phonolites and are reasonably common in ne-benmoreites. The compositional range among feldspars in phonolite clasts is wider than is the case for feldspars from phonolitic rocks of the Dunedin Volcanic Group (Fig. 6), with some extremely potassic (PCB20: An0Ab2Or98) and sodic feldspars (PCB20: An5Ab95Or0) resulting from perthitic exsolution. Such compositions do not occur in rocks of the Dunedin Volcanic Group (Fig. 6).
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The feldspars of the syenitic rocks show much the same range of variation as that observed among the feldspars of the phonolitic clasts, with alkali feldspars dominating (PCB53: An0Ab72Or28 to OU 24273: An0Ab8Or91) and plagioclase rare (PCB42: An15 Ab82Or3 to PCB5: An34Ab60Or6). Compositional zoning is subdued and trends are often inconsistent; some grains become enriched in Na from core to rim (e.g. PCB9: core An0Ab57Or43, rim An0Ab61Or39), whereas in other rocks zoning trends are obscured by the effects of perthite exsolution.
Nepheline, sodalite and analcime
Representative nepheline, sodalite and analcime compositions are given in Table 6. Nepheline occurs as a groundmass and phenocryst phase in phonolite and ne-trachyandesite, and is common in syenite clasts. Nepheline compositions are plotted in the system Q–Ne–Ks in Fig. 7, along with temperature limits on nepheline solid solution determined by Hamilton (1961). In nepheline–sodalite syenite, PCB9, nephelines are consistently zoned, with phenocryst core compositions (Na2O 17·45%, K2O 4·20%) evolving through rims (Na2O 16·52%, K2O 5·19%) to a more potassic groundmass phase (Na2O 16·52%, K2O 5·87%). Similar potassic enrichment with fractionation is observed in phonolite PCB20. The nepheline compositions in phonolitic rocks are consistent with temperatures in the range 1000 to 700°C, with syenitic rocks showing a wider compositional range that is perhaps indicative of a wider temperature range (>1068°C to 500°C).
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Sodalite occurs in many of the syenite clasts. In PCB9 it forms primary, euhedral, octahedral phenocrysts that crystallized along with feldspar and nepheline. Analysed sodalites are unzoned; all are low in calcium and there is no detectable solid solution with a potassic end-member (Table 6).
Analcime is also common in syenites, forming interstitial patches and cross-cutting veinlets associated with calcite and white mica. In syenite, PCB9, analcime forms a marginal replacement of sodalite. In some ultramafic cumulates analcime occurs filling vesicles in interstitial glass, and in a similar, petrogenetically late mode, as a cement in the PCB.
Glass
Spherical glass inclusions (0·01–0·05 mm diameter) occur within pyroxenes, amphiboles and apatites of mafic cumulate rocks, and interstitial and devitrified glass patches are common in gabbro and clinopyroxenite clasts. Average glass compositions of the spherical inclusions, obtained using electron microprobe analysis, are shown in Table 7. Glasses in the ultramafic cumulates and gabbros are strongly undersaturated and phonolitic (ne-benmoreite or tephro-phonolite).
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WHOLE-ROCK MAJOR AND TRACE ELEMENT GEOCHEMISTRY |
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TOPABSTRACTINTRODUCTIONTHE PORT CHALMERS BRECCIAPETROGRAPHYMINERAL CHEMISTRYWHOLE-ROCK MAJOR AND TRACE...ISOTOPE GEOCHEMISTRYDISCUSSIONCONCLUSIONSREFERENCES |
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Methods
Major and minor elements (Si, Ti, Al, Fe, Mn, Mg, Ca, Na, K, P and S) were determined by X-ray fluorescence (XRF) at La Trobe University using methods similar to those described by Norrish & Hutton (1969)
. For these elements, precision is generally better than ±1% (1
). FeO abundances were measured by direct titration using a standardized CeSO4 solution, and H2O and CO2 by a gravimetric method. Selected trace elements were determined by XRF on pressed powder pellets (Norrish & Chappell, 1977) and, for these elements, theoretical detection limits are of the order of 1–2 ppm and reproducibility is better than ±5% (1).
For selected samples the rare earth elements (REE) and a selection of other trace elements were analysed by inductively coupled plasma mass spectrometry (ICPMS) at the VIEPS Trace Element Laboratory at Monash University using methods described by Price et al. (1997). Precision for these elements is typically better than 5% and accuracy, based on analysis of BHVO-1, better than 5% at the 95% confidence level.
Major and trace element variations
Representative major and trace element data for samples from the PCB are presented in Table 8. All analysed rocks are nepheline-normative (ne-) and, using the total alkalis vs SiO2 classification scheme of Le Maitre et al. (1989), the fine-grained clasts range from basanitic and basaltic through trachyandesitic to phonolitic (Fig. 8). Phonolitic clasts [tephri-phonolites and phonolites under the scheme of Le Maitre et al. (1989)] have been classified on the basis of normative compositions and differentiation index (Coombs & Wilkinson, 1969; Price & Chappell, 1975) as ne-benmoreites and phonolites.
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The bulk composition of the matrix of the PCB is ne-benmoreite. The major element variations defined by clasts from the PCB mirror those observed in the Dunedin Volcanic Group (Fig. 8). Most of the syenitic clasts are chemically similar to the phonolites but several have compositions resembling those of Dunedin trachyandesites. The majority of the syenite clasts do, however, have lower K2O contents than is the case for phonolites of the Dunedin Volcanic Group including the PCB (Fig. 8). Trace element variations among PCB clasts are broadly similar to those observed in the Dunedin Volcanic Group (Fig. 9) but syenitic clasts tend to have higher Sr contents and a few have significantly higher Zr abundances (up to 5236 ppm). As expected, ultramafic rocks and gabbros show higher abundances of MgO, FeO, TiO2, CaO, Cr, Ni and V, and lower Al2O3, Na2O, K2O, Ba, Rb, Sr, Nb and Zr contents. Two ultramafic rocks have higher Al2O3, Ba and Sr abundances than other cumulates and they also show relatively elevated Rb (10–12 ppm compared with 4–7 ppm for other cumulates) and Nb (160–376 ppm compared with 29–72 ppm) abundances. In both of these clasts interstitial glass is common.
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Chondrite-normalized REE abundance patterns for representative PCB clasts are shown in Fig. 10, along with patterns for representative samples from the Dunedin Volcanic Group. Phonolitic and syenitic patterns are all broadly similar with relatively flat heavy REE (HREE) to middle REE (MREE), enrichment of light REE (LREE) over MREE and HREE [(La/Yb)n = 22–60] and moderate depletions in Eu relative to Sm and Gd (Eu/Eu* = 0·53–0·75). PCB syenites and phonolites show patterns that are similar to those observed in Dunedin Volcano ne-benmoreites but they lack the more extreme Eu depletions of the most felsic Dunedin phonolites (Fig. 10). Gabbro and basanite clasts show very similar normalized rare earth patterns with enrichment of LREE over HREE [(La/Yb)n = 13–15] and weak negative Eu anomalies (Eu/Eu* = 0·7). They are in many respects compositionally similar to basanites of the Dunedin Volcanic Group. Ultramafic cumulate clasts show more variability than the gabbro clasts. PCB30 has MREE abundances similar to those observed in other ultramafic and mafic rocks but shows depletion of LREE relative to MREE and lower abundances of HREE (Fig. 10).
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In major, trace and minor element terms, syenite and phonolite clasts are very similar to ne-benmoreite compositions from the Dunedin Volcanic Group. Relative to basanite, these rocks are enriched in large ion lithophile elements such as Rb, K and the REE, and some high field strength elements such as Zr and Nb, but Ba, Sr, Eu and Ti are relatively depleted. Ultramafic rocks generally have trace and minor element abundance patterns complementary to those of the felsic rocks. None of the phonolitic or syenitic rocks of the PCB shows the more extreme depletions in Ba, Sr, Eu and Ti observed in some of the phonolites of the Dunedin Volcanic Group.
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ISOTOPE GEOCHEMISTRY |
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TOPABSTRACTINTRODUCTIONTHE PORT CHALMERS BRECCIAPETROGRAPHYMINERAL CHEMISTRYWHOLE-ROCK MAJOR AND TRACE...ISOTOPE GEOCHEMISTRYDISCUSSIONCONCLUSIONSREFERENCES |
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Methods
Strontium, Nd and Pb isotopic data were obtained in the VIEPS isotope laboratory at La Trobe University using a seven-collector Finnigan-MAT 262 spectrometer and methods described in detail by Price et al. (1999)
. Reproducibility on Pb isotopic compositions for SRM981 (n = 78, 2) is ±0·097% for 206Pb/204Pb, ±0·130% for 207Pb/204Pb, and ±0·175% for 208Pb/204Pb. The average fractionation factor during runs was 0·091% (2 = ±0·034%), close to the empirical value of 0·109%.
For Sr, instrumental mass fractionation was corrected by normalizing to 86Sr/88Sr = 0·1194. Typically, 5–7 blocks of 10 x 8 s integrations produced in-run precision (2) of ±0·003%. 87Sr/86Sr (±2) for SRM987 (n = 100) is 0·71023 ± 7 (0·01%), for BCR-1 (n = 6) 0·70500 ± 4, and for BHVO-1 (n = 19) 0·70348 ± 6.
For Nd, fractionation was corrected by normalizing to 146Nd/144Nd = 0·7219. Typically, 5–7 blocks of 10 x 8 s integrations produced in-run precisions (2) of ±0·0025%. 143Nd/144Nd (±2) for La Jolla (n = 100) is 0·511860 ± 16, for BCR-1 (n = 7) 0·512634 ± 18, and for BHVO-1 (n = 5) 0·512989 ± 13. Present-day CHUR was taken as 0·512631.
Stable isotope ratios were measured at the VIEPS stable isotope facility at Monash University using a Finnigan MAT 252 mass spectrometer. CO2 was extracted from calcite by reaction with H3PO4 at 25°C for 12–18 h in sealed vessels (McCrea, 1950). 
18O and 13C values are expressed relative to V-SMOW and V-PDB, respectively. Internal calcite standard ISACC analysed at the same time as the samples yielded 13C and 18O values within 0·1
of its long-term average. This standard was calibrated using IAEA-CO-1 and its long-term average 13C and 18O values are 6·37 ± 0·06 and 12·68 ± 0·13. Based on replicate analyses, reproducibility is estimated as ±0·1 for both O and C.
Pb, Sr and Nd isotopic data
Pb, Sr and Nd isotopic data for PCB samples are presented in Table 9. The Sr data have been age corrected but, given the relatively young ages (10–13 Ma), the differences between initial and present-day ratios are minor.
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In 207Pb/204Pb vs 206Pb/204Pb and 208Pb/204Pb vs 206Pb/204Pb diagrams (Fig. 11), data for PCB clasts form a tight cluster on the Northern Hemisphere Reference Line (Hart, 1984
), with 207Pb/204Pb ratios ranging from 15·624 to 15·711 and 208Pb/204Pb ratios from 39·155 to 39·547. Sr and Nd isotopic compositions also show very little variation (Fig. 11), with present-day 87Sr/86Sr ratios ranging from 0·702794 to 0·703410 and 143Nd/144Nd from 0·512866 to 0·512926. Lead isotopic data are not available for other rocks from Dunedin Volcano, but Sr and Nd isotopic data (Price & Compston, 1973; Coombs et al., 1986; McDonough et al., 1986) indicate that the clasts of the PCB are isotopically indistinguishable from Tertiary East Otago and Banks Peninsula volcanics (Fig. 11; see Fig. 1 for geographical distribution).
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Regional comparisons illustrate that Dunedin volcanic rocks are isotopically very similar to volcanic and intrusive rocks of various ages from elsewhere in southern New Zealand, the Tasman Sea and the sub-Antarctic region to the south (Fig. 11; see Fig. 1 for geographical distribution) and these similarities have been recognized for some time. McDonough et al. (1986)
reported that east Otago volcanic rocks have 87Sr/86Sr ratios ranging from 0·7029 to 0·7032 and 
Nd from +4·7 to +6·3, and noted that this limited isotopic range was very similar to those observed among rocks of the New Zealand sub-Antarctic islands and from Banks Peninsula. Tertiary lamprophyric dykes of the West Coast of New Zealand's South Island (the Alpine dyke swarm—Cooper, 1986; Barriero & Cooper, 1987) also show a very close isotopic similarity to Dunedin volcanic rocks (Fig. 11). Lanyon et al. (1993) ascribed the isotopic characteristics of the southern New Zealand igneous rocks to the regional influence of a mantle plume and broadened the region showing these isotopic features to include Tertiary basalts in Tasmania, seamounts of the southern Tasman Sea, and volcanics of the Balleny Islands and Marie Byrd Land in Antarctica (Fig. 1). Most recently, Baker et al. (1994) noted the similarity between the isotopic composition of the Alpine dyke swarm and the Tapuaenuku igneous complex, a layered alkalic intrusive complex of Cretaceous age (90–100 Ma) in the Inland Kaikoura Ranges of north Canterbury (Fig. 1), thereby extending the isotopic regional province into the northern South Island and back in time to the late Mesozoic.
The New Zealand–Tasman–Antarctic isotopic province has Pb, Sr and Nd compositions that partially overlap with those observed in the Marquesas and Samoan seamount chains (Fig. 11), but the Pb isotopic compositions extend towards more radiogenic values and Lanyon et al. (1993) and Baker et al. (1994) argued that an enriched high U/Pb (high µ or HIMU) component was required in the mantle source.
Stable isotope data
Four samples were analysed for carbon and oxygen isotopic composition. Three of these are carbonate veins and the fourth is a sample of carbonate separated from a fenitized schist clast. Results are presented in Table 10. Two of the vein samples have relatively primitive compositions with 18Osmow in the range 7·2–7·5 and 13C between -5·8 and -6·0. The fenitized schist carbonate sample also shows primitive values (8·4 and -7·2). The third vein sample is significantly different (18O = 13·4 and 13C = 1·7) and has a composition that is more like that expected for a sedimentary rock. The more primitive isotopic compositions have carbonatitic affinity (e.g. Deines, 1989; Reid & Cooper, 1992) and they indicate that magmatic CO2 was a significant volatile during the crystallization and emplacement of the PCB.
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DISCUSSION |
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TOPABSTRACTINTRODUCTIONTHE PORT CHALMERS BRECCIAPETROGRAPHYMINERAL CHEMISTRYWHOLE-ROCK MAJOR AND TRACE...ISOTOPE GEOCHEMISTRYDISCUSSIONCONCLUSIONSREFERENCES |
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Emplacement of the Port Chalmers Breccia
The outcrops of PCB across the central Otago Harbour (Fig. 2) have all the characteristics of diatremes (Lorenz, 1986
) and appear to represent feeder structures to what were probably small maar volcanoes.
The Laacher See volcano in the East Eifel district of Germany could provide an analogue for the volcanoes that lay above the Port Chalmers diatremes. An estimated 5·3 km3 of phonolitic pyroclastic material (Wörner & Schminke, 1984a; Freundt & Schminke, 1986) was erupted at Laacher See to form a dispersed tuff-ring around a central crater that has a present-day diameter of 2–3 km. Wörner & Schminke (1984a, 1984b) concluded that eruptions tapped a shallow (3–6 km) zoned magma chamber and they attributed the explosive nature of the eruptive activity to interaction between groundwater and magma. The PCB pipes also have similarities to phonolitic breccia pipes of the Cripple Creek district of Colorado. Kelley et al. (1998) proposed that Cripple Creek phonolitic magmas evolved in the lower crust or upper mantle and then rose rapidly to the surface along faults. Some of the Cripple Creek Breccia pipes appear to have been generated when magmas encountered groundwater, but others may have formed during explosive emplacement arising from vapour saturation (H2O and CO2) in rapidly rising magma.
In cases where diatreme and maar formation is argued to involve interaction between magma and groundwater (e.g. Wörner & Schminke, 1984a; Lorenz, 1986), depth of diatreme growth is generally related to water-table depth or, very commonly, to the intersection of rising magma with specific aquifers. For the PCB, the deepest potential aquifer is the basal unit of the Cretaceous–Tertiary cover sequence, the Taratu Formation (McKellar, 1966) at a depth of 800–900 m (Coombs, 1965). Clasts from this unit are extremely rare but schist clasts are common, indicating that brecciation occurred in the basement below the cover sequence. Brecciation of cumulate and syenitic blocks must have also taken place during explosive disruption of a melt cumulate zone but the metamorphic grade of the schists places a relatively shallow limit (2–3 km?) on the depth at which this took place.
If the explosive emplacement of the PCB was not driven by interaction with groundwater it is likely that exsolution of juvenile gas was involved. The abundance of amphibole and biotite indicates that magmas were water bearing but there is geological and petrological evidence that CO2 was also a significant volatile component (e.g. carbonate veins and fenitization of schist clasts). Stable isotope data indicate that the CO2 involved was largely of magmatic origin. CO2 solubility is generally higher in silica-undersaturated than in silica-saturated or -oversaturated magmas (e.g. Mysen et al., 1975; Spera & Bergman, 1980; Holloway & Blank, 1994) and decreasing pressure favours decreased CO2 and H2O solubility (e.g. Burnham, 1967; Spera & Bergman, 1980; Stolper et al., 1987; Blank & Brooker, 1994). H2O and CO2 also lower melt viscosity and melt density (Burnham, 1967; Blank & Brooker, 1994). The bulk chemistry and the volatile contents of PCB clasts are therefore consistent with a magmatic system in which parental melts were relatively low density, low viscosity, and CO2 and H2O rich. Conceivably, fracturing along a NW–SE fault could have released such magmas rapidly and explosively from shallow crustal sources.
The generation and evolution of the magmas represented by the clasts of Port Chalmers Breccia
The only clasts found in the PCB that can be interpreted as representing juvenile material are small, originally highly vesicular, analcime- and calcite-impregnated, devitrified glassy fragments found within the breccia matrix. Other clasts show evidence of being transported within, or physically infiltrated by a very similar gas-charged magma. Compositions of fresh glassy inclusions in minerals indicate that this magma was phonolitic (ne-benmoreite or phono-tephrite). Price & Chappell (1975) argued that Dunedin ne-benmoreites were derived from basanitic parental magmas by crystal fractionation involving olivine, pyroxene, amphibole and magnetite (see also Coombs & Wilkinson, 1969). Similar hypotheses have been put forward to explain geochemical variation in basanite– phonolite associations elsewhere (e.g. Baker, 1969; Nash et al., 1969; Wörner & Schmincke, 1984b; Price et al., 1985; Le Roex et al., 1990).
The ne-benmoreite clasts of the PCB are relatively mafic when compared with the more felsic phonolites flows and domes of the Dunedin Volcanic Group and they do not exhibit the extreme relative depletions in Mg, Ca, Sr and Eu commonly observed in the felsic phonolites. The differences could reflect the levels at which crystal fractionation has taken place, with the ne-benmoreites generated by crystal fractionation in the deep crust or upper mantle (Irving & Price, 1981) and more strongly fractionated phonolites, with their relatively low Sr contents and Eu-depleted REE patterns, by shallow-level, feldspar-dominated crystal fractionation.
An alternative, crustal anatexis model has been suggested for the origin of phonolitic rocks in the Kenya Rift (e.g. Bailey, 1964; Williams, 1970; Hay & Wendlandt, 1995). Proponents of a crustal melting origin for phonolites argue that volume considerations, the uniformity of plateau-type flood phonolites and the paucity of intermediate compositions are all factors that present problems for a fractional crystallization origin for these rocks. Hay & Wendlandt (1995) used high-pressure experiments to demonstrate that Kenya rift flood phonolites have equilibrated under lower-crustal conditions and that it is feasible to generate phonolites by partial melting of an alkali basaltic composition in the lower crust. They supported their experimental conclusions with an analysis of geochemical data for Kenya rift phonolites (Hay et al., 1995).
An analogous two-stage model could apply to the ne-benmoreites of the PCB (Fig. 12). During the first stage, mantle melting under near water-saturated conditions could have generated basanitic magmas that underplated and intruded the lower crust. Some of this material could have crystallized in the lower crust to form alkali gabbroic assemblages consisting of plagioclase, clinopyroxene, amphibole, biotite and magnetite. Other batches of magma moved upwards towards the surface, undergoing varying degrees of crystal fractionation and forming small temporary magma reservoirs or crystallizing to form sills, dykes and small intrusions distributed throughout the crust (Fig. 12). Reilly (1971) interpreted gravity data for the Dunedin Volcano to indicate that an extensive volume of the crust underlying the volcano is composed of dense intrusive rocks. At the second stage, rising geotherms associated with continued injection of mafic magma could cause partial melting of lower-crustal mafic intrusives to produce ne-benmoreite magmas.
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The textures, mineralogy and chemistry of the coarse-grained clasts in the PCB indicate that they constitute a disrupted assemblage of coarse-grained basanitic rocks (alkali gabbros), ultramafic and felsic cumulates, and differentiated nepheline syenites. Amphiboles in some ultramafic cumulates in the PCB are very similar compositionally to those in the ne-benmoreites and some of the syenites, and this might indicate that at least some amphibole cumulates have crystallized from relatively evolved magmas. Olivine is not present in ultramafic cumulate rocks, which have modal compositions dominated by amphibole and pyroxene. It would therefore appear feasible that much of the ultramafic and possibly some of the syenitic cumulates in PCB were derived from crystal fractionation of relatively felsic magmas (e.g. ne-benmoreite).
The processes of transfer of magmas from the mantle, melting of mafic crust and differentiation of mantle- and crust-derived magmas are envisaged to have continued throughout much of the history of the Dunedin Volcano. Emplacement of the Port Chalmers diatremes appears, however, to have been a unique event triggered by faulting, which caused the rapid ascent and degassing of ne-benmoreite magma with consequent brecciation and entrainment of a wide range of lithologies between source and surface. The presence of clasts of Cretaceous sediment means that the depth of the source has to be at least 1 km but the metamorphic grade of the schist clasts indicates that it was located in the upper (1–3 km?) rather than middle or lower crust. We envisage that, before disruption, the source region for the PCB consisted of a mush of near volatile-saturated ne-benmoreite melt within a complex of gabbros, syenites, and ultramafic and felsic cumulates.
SiO2-undersaturated basaltic and phonolitic magmas were continuously emplaced throughout the history of the Dunedin Volcano but the more extremely fractionated felsic phonolites become progressively more common among the younger eruptives. We interpret these patterns to indicate that, both before and after the PCB event, mantle- and lower-crustal-derived magmas continued to feed into a dispersed magmatic system in which crystal fractionation continued to take place in dykes, sills and other small magma reservoirs throughout the crust. Virtually all the mafic rocks in the Dunedin volcano have compositions indicating that they have undergone crystal fractionation. The felsic low-Mg, low-Sr phonolitic rocks were derived from ne-benmoreite precursors by more extreme fractionation at relatively shallow levels (Price & Chappell, 1975).
Mantle sources and tectonic setting
The Pb, Sr and Nd isotopic data indicate that magmas represented by the rocks of the PCB were generated from mantle sources with compositions between HIMU and a depleted mantle component. This particular mantle source appears to be very widespread over a region that includes the South Island of New Zealand, the South Tasman Sea, Tasmania, the New Zealand sub-Antarctic islands, the Balleny Islands and parts of mainland Antarctica (Lanyon et al., 1993). Tertiary volcanics in eastern Australia (McDonough et al., 1985) and seamounts of the central and north Tasman Sea (Eggins et al., 1991) are isotopically different and appear to have derived from mantle sources with compositions involving depleted and enriched (EM1) components.
Lanyon et al. (1993) noted that the isotopic similarities of mantle source compositions across the region were supported by trace and minor element data, and argued that these regional HIMU signatures arise from the effects of two separate plume sources that are now located in the Balleny Islands and Marie Byrd Land, respectively. They concluded that all intraplate association igneous rocks in the region showing the influence of a HIMU mantle source are related directly to plume activity, which was also the driving force initiating continental rifting and separation before 83 Ma BP. The work of Baker et al. (1994) on the Tapuaenuku igneous complex of Cretaceous age (90– 100 Ma) in the Inland Kaikoura Ranges of the South Island has, however, illustrated that the HIMU influence on regional mantle isotopic signatures was already present around 20 Myr before the initial rifting associated with Tasman Sea opening.
The problem for the plume model is that a specific isotopic signature is manifested over a vast area (Fig. 1), in a wide range of intraplate basaltic rocks of various ages (from virtually the present day to 100 Ma). The seamounts of the South Tasman Sea and the Balleny Islands show systematic progression in age that can be reconciled with a hotspot trace but across the rest of the region there appears to be no systematic spatial–temporal relationship (see Adams, 1981). One possibility is that the influence of the Balleny and Marie Byrd plumes was very widespread before continental separation so that the lithospheric mantle across the region was chemically preconditioned. Magmas produced during any subsequent thermal events in the region are either derived directly from the lithospheric mantle or are contaminated by it. Such a model would require a substantial lead time (>20 Myr) between the arrival of a plume at the base of the lithosphere and the first significant rifting in the Tasman Sea. An alternative to the plume model would be along the lines proposed by Baker et al. (1994), whereby the characteristic isotopic signature of the post-Mesozoic intraplate igneous rocks reflects a shallow regional mantle reservoir that was generated during prolonged subduction throughout the Mesozoic and before continental separation. The implication is that the Dunedin Volcanic Group magmas contain a significant component that originated in the sub-continental, lithospheric mantle.
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CONCLUSIONS |
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TOPABSTRACTINTRODUCTIONTHE PORT CHALMERS BRECCIAPETROGRAPHYMINERAL CHEMISTRYWHOLE-ROCK MAJOR AND TRACE...ISOTOPE GEOCHEMISTRYDISCUSSIONCONCLUSIONSREFERENCES |
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Construction of the Dunedin Volcano began around 13 Myr ago with the submarine emplacement of basalts and quartz-normative trachytes, and continued over a period of 3 Myr with subaerial eruption of dominantly silica-undersaturated basaltic and phonolitic magmas. Outcrops of PCB are found only in contact with older units of the volcanic sequence but the clast assemblages in the breccias indicate that they were emplaced after significant volumes of phonolitic magma had been erupted and a substantial subaerial volcanic complex constructed. The PCB outcrops represent diatremes and proximal pyroclastic deposits associated with maar volcanoes aligned along a NW–SE fault system. They were emplaced when faulting released volatile-rich ne-benmoreite magma from melt–cumulate zones located 1–3 km beneath the volcano.
The PCB is phonolitic in composition (ne-benmoreite or tephro-phonolite) and consists of abundant clasts of ne-benmoreite, phonolite, basanite, ne-trachyandesite, syenite, gabbro, ultramafic cumulate, altered schist and Cretaceous or Tertiary sediment in a sand and silt matrix with a bulk composition of ne-benmoreite.
Mineralogically, the syenites of the PCB show variations similar to those observed in nepheline syenites from elsewhere (e.g. South Greenland, East Africa) but some of the mineralogical compositional variation is more extreme in the PCB syenite clasts. Pyroxenes are dominantly aegirine and aegirine–augite, and biotite is abundant and Ti rich. Alkali amphibole is present within syenite clasts but Ti-rich calcic amphiboles are abundant in all coarse-grained rocks and dominate ultramafic cumulates and gabbros. Nepheline is ubiquitous in syenitic blocks and sodalite is common.
Major and trace element variations defined by the clast suite from the PCB are similar to those observed in the Dunedin Volcanic Group, although the extreme depletions in Mg, Ca and Sr and strong negative Eu anomalies shown by the most fractionated Dunedin phonolites do not feature among the phonolite clasts of the PCB and are uncommon among the syenite clasts. The clast assemblage of the PCB could be interpreted to represent a consanguineous suite related by crystal fractionation of a basanitic parent in a crustal magmatic system. Ne-benmoreites could be crystal fractionation products from basanite but we believe it is more likely that they were generated in a two-stage process involving partial melting of deep crustal intrusive rocks. Highly fractionated phonolites of the Dunedin Volcanic Group were derived by feldspar-dominated crystal fractionation from ne-benmoreite precursors.
Parental magmas of the Dunedin Volcanic Group magmas came from an isotopically distinct mantle source with a Sr, Nd and Pb isotopic composition between HIMU and depleted mantle. The composition matches closely those of Mesozoic and Tertiary intraplate igneous rocks across southern New Zealand, the South Tasman Sea, Tasmania, the New Zealand sub-Antarctic islands, the Balleny Islands and Marie Byrd Land. The pattern indicates the presence of a distinctive, long-lived and extensive lithospheric mantle reservoir in the region.
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ACKNOWLEDGEMENTS |
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The technical support of Ian McCabe, Jorg Metz and Allen Jacka is gratefully acknowledged. John Gamble, Steve Weaver and Joel Baker very generously gave access to unpublished isotopic data for Auckland Islands, Banks Peninsula and Tapuaenuku samples. The development of the project benefited significantly from discussions with Professor Doug Coombs. Constructive and thorough reviews of an earlier version of the paper by Anton Le Roex and Brian Upton are gratefully acknowledged, as is the editorial assistance of Richard Arculus. This research was funded by research grants from La Trobe University and the University of Otago.
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