Journal of Petrology Advance Access originally published online on November 26, 2004
Journal of Petrology 2005 46(3):473-503; doi:10.1093/petrology/egh084
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Petrology and Geochemistry of Intraplate Basalts in the South Auckland Volcanic Field, New Zealand: Evidence for Two Coeval Magma Suites from Distinct Sources
1 DEPARTMENT OF EARTH SCIENCES, THE UNIVERSITY OF WAIKATO, PRIVATE BAG 3105, HAMILTON, NEW ZEALAND
2 DEPARTMENT OF GEOLOGY, UNIVERSITY OF AUCKLAND, PRIVATE BAG 92019, AUCKLAND, NEW ZEALAND
3 DEPARTMENT OF EARTH SCIENCES, LA TROBE UNIVERSITY, BUNDOORA, VIC. 3083, AUSTRALIA
RECEIVED JUNE 1, 2003; ACCEPTED SEPTEMBER 28, 2004
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ABSTRACT |
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TOP
ABSTRACT
INTRODUCTIONThe South Auckland Volcanic Field is a Pleistocene (1·59–0·51 Ma) basaltic intraplate, monogenetic field situated south of Auckland City, North Island, New Zealand. Two groups of basalts are distinguished based on mineralogy and geochemical compositions, but no temporal or spatial patterns exist in the distribution of various lava types forming each group within the field: Group A basalts are silica-undersaturated transitional to quartz-tholeiitic basalts with relatively low total alkalis (3·0–4·6 wt %), Nb (7–29 ppm), and (La/Yb)N (3·4–7·6); Group B basalts are strongly silica-undersaturated basanites to nepheline-hawaiites with high total alkalis (3·3–7·9 wt %), Nb (32–102 ppm), and (La/Yb)N (12–47). Group A has slightly higher 87Sr/86Sr, similar
Nd, and lower 206Pb/204Pb values compared with Group B. Contrasting geochemical trends and incompatible element ratios (e.g. K/Nb, Zr/Nb, Ce/Pb) are consistent with separate evolution of Groups A and B from dissimilar parental magmas derived from distinct sub-continental lithospheric mantle sources. Differentiation within each group was controlled by olivine and clinopyroxene fractionation. Group B magmas were generated by <8% melting of an ocean island basalt (OIB)-like garnet peridotite source with high 238U/204Pb mantle (HIMU) and enriched mantle (EMII) characteristics possibly inherited from recycled oceanic crust. Group A magmas were generated by <12% melting of a spinel peridotite source also with HIMU and EMII signatures. This source type may have resulted from subduction-related metasomatism of the sub-continental lithosphere modified by a HIMU plume. These events were associated with Mesozoic or earlier subduction- and plume-related magmatism when New Zealand was at the eastern margin of the Gondwana supercontinent. KEY WORDS: continental intraplate basalts; geochemistry; HIMU, EMII; Sr, Nd, and Pb isotopes; South Auckland; sub-continental lithospheric sources
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INTRODUCTION |
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The South Auckland Volcanic Field is one of a number of Pliocene to Recent continental basaltic intraplate volcanic fields in North Island, New Zealand (Fig. 1). The field formed between 0·51 and 1·59 Ma (Briggs et al., 1994
) and is distinct from the younger Auckland Volcanic Field to the north. It consists of eroded basaltic lava flows, scoria cones, and tuff deposits, produced from 
100, predominantly monogenetic, volcanic centres over an area of 300 km2. Aspects of the geology of the field have been described by Rafferty & Heming (1979) and Briggs et al. (1994).
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Previous petrological studies of the South Auckland basalts have documented a wide range of alkalic basaltic magmas with contrasting compositions, erupted during the 1 Myr life of the field (Rafferty & Heming, 1979
; Briggs et al., 1994). Rafferty & Heming (1979) divided the basalts into two broad groups based primarily on their distinct geochemical differences: a predominantly hypersthene (hy)-normative subalkaline group and a predominantly nepheline (ne)-normative alkaline group, referred to here as Group A and Group B, respectively. In this paper we present and interpret new mineralogical, geochemical, and Sr, Nd and Pb isotopic data, in the range of basalt compositions that make up Groups A and B.
Numerous petrological studies of continental intraplate volcanic fields elsewhere in the world demonstrate that they commonly consist of a suite of alkalic basalts characterized by a spectrum of genetically related compositions, ranging from strongly silica-undersaturated to less silica-undersaturated compositions (e.g. Ewart et al., 1988; Zhi et al., 1990; Huang et al., 1997, 2000; Zhou & Mukasa, 1997; Cebriá et al., 2000; Zou et al., 2000; Zhang et al., 2001). Despite their wide range of geographic distribution, these basalts commonly share ocean island basalt (OIB)-like geochemical characteristics. These studies also demonstrate that the composition of the mantle sources involved typically varies on a regional scale, and these variations are commonly interpreted in terms of the mixing of two or more of the mantle reservoir end-members: depleted MORB mantle (DMM), high 238U/204Pb mantle (HIMU), and enriched mantle (EMI or EMII) of Zindler & Hart (1986).
Although the South Auckland basalts have geochemical compositions, and OIB-like geochemical characteristics (Cook, 2002), comparable with silica-undersaturated mafic magmas erupted in other continental intraplate tectonic settings, they are different from suites of alkalic basalts elsewhere because each group appears to have evolved as a set of distinct lineages. In addition, Group A and B basalts exhibit subtle but distinct variations in abundances of the large ion lithophile elements (LILE), high field strength elements (HFSE), rare earth elements (REE), and isotopic compositions. Such variations are commonly attributed to source heterogeneity, or variable degrees of partial melting, or both (e.g. Zhi et al., 1990; Zou & Zindler, 1996). These variations may also be explained by incompatible element enrichment as a result of mantle metasomatism (e.g. McDonough et al., 1985; O'Reilly & Zhang, 1995), geochemical processes in the crust–mantle system (e.g. Nakada et al., 1997), or modification by assimilation–fractional-crystallization processes (e.g. Price et al., 1997).
The monogenetic nature of eruptive events in the South Auckland Field and distinct geochemical characteristics of Group A and B lavas preclude the presence of a long-lived shallow magma reservoir beneath the field. The geochemical diversity is more consistent with a variety of parental magmas giving rise to a diverse range of derivative lava compositions, rather than derivation from a single discrete parental magma. Furthermore, as demonstrated by Briggs et al. (1994), there is no apparent temporal or spatial pattern between, or among, the lavas that make up each group (Fig. 2), and the distribution of each Group A- and B-type volcano appears to be random over the field.
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The diversity in the major and trace element and isotopic compositions of the South Auckland basalts, together with their close time–space relationships and proximity to an active convergent margin, raises some key questions regarding their origin and the nature of the mantle sources from which they were derived. The focus of this paper is on the petrogenesis and evolution of the alkalic basaltic lavas, with the aim of determining: (1) the nature of the mantle source (or sources); (2) the extent to which the compositions of the basalts are source- or process-dependent; (3) if there is a genetic link between the ranges of magma compositions erupted throughout the field; (4) what magmatic processes are involved in their genesis and evolution.
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GEOLOGICAL SETTING |
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Cenozoic volcanism is a dominant feature in the geological development of northern North Island, New Zealand. North Island is located toward the eastern margin of the Indo-Australian Plate, bounded by the convergence and associated westward subduction of the Pacific Plate at the Hikurangi Trough (Fig. 1a). Since the early Miocene, the evolution of the Indo-Australian–Pacific plate boundary has resulted in subduction-related volcanism in Northland Peninsula, the Coromandel Peninsula, and the Taupo Volcanic Zone (TVZ) (e.g. Hayward et al., 2001
), producing large volumes of subduction-related magma varying compositionally from basalt through andesite to rhyolite. The geological record also indicates that volumetrically minor amounts of basaltic magma with alkalic to tholeiitic compositions were erupted in a number of distinct intraplate volcanic fields in northern North Island. From north to south, these fields are: Kaikohe–Bay of Islands and Whangarei (e.g. Huang et al., 2000) of the Northland Volcanic Province; and Auckland (e.g. Huang et al., 1997), South Auckland (e.g. Rafferty & Heming, 1979), Ngatutura (e.g. Briggs et al., 1990) and Okete (e.g. Briggs & Goles, 1984) of the Auckland Volcanic Province (Fig. 1b). Each field contains numerous, relatively small-volume, predominantly monogenetic volcanoes that produced silica-undersaturated lavas and tephra with olivine (ol)- and ne- or hy-normative compositions, typical of continental intraplate basalts. These intraplate volcanic products temporally overlap but, except for the Northland fields, are generally spatially separated from those derived from subduction-related processes.
The South Auckland Field is situated in an extensional tectonic environment (Spörli, 1980) about 160 km behind the active volcanic front of the TVZ. The seismic data of Adams & Ware (1977) and Anderson & Webb (1994) indicate that the NW-dipping Wadati–Benioff zone, which marks the upper boundary of the subducting Pacific Plate beneath North Island, can be tracked only to about 250 km depth, and that there is no seismic evidence for the existence of the slab beneath the South Auckland Field. Seismic evidence also indicates that there is a zone of low-velocity mantle material NW of the TVZ between 75 and 125 km depth (Hatherton, 1970).
The thickness of the continental crust in the South Auckland region ranges from 22 to 28 km (Stern & Davey, 1987). Volcanic centres west of the Drury Fault (Fig. 2) were generally erupted through marine siltstones, sandstones, mudstones, and volcaniclastic sequences of the Triassic–Cretaceous Murihiku Terrane (e.g. Ballance & Campbell, 1993; Middleton, 1993; Kamp & Liddell, 2000) and aquifer-bearing Tertiary formations (i.e. the Oligocene Te Kuiti Group; White & Waterhouse, 1993). Centres east of the Drury Fault erupted primarily through the greywackes, argillites, and volcaniclastic metasediments of the Triassic–Jurassic Waipapa Terrane (Spörli, 1978).
Many of the volcanic centres in the South Auckland Field are located either along or adjacent to faults, or inferred extensions of faults. Spörli & Eastwood (1997) argued that tensional stresses, inherited from Mesozoic tectonic events, could control the location of individual intraplate fields within the Auckland Province (Fig. 1b), as well as facilitate localized decompressional melting and thus the onset of volcanism.
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ANALYTICAL TECHNIQUES |
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Abundances of major elements and selected trace elements were determined by X-ray fluorescence (XRF) analysis for 93 samples at the Analytical Facility, Victoria University, Wellington, New Zealand, and 110 samples at the Department of Geology, University of Auckland, Auckland, New Zealand, on fused glass discs and pressed powder pellets using methods similar to those described by Norrish & Chappell (1977)
and Parker et al. (1993). Precision is better than 1% (1
) for each major element and 5% (1) for trace elements. Abundances of selected trace elements and REE were determined for a subset of 15 samples by spark source mass spectrometry (SSMS) at the Research School of Earth Sciences, Australian National University, Canberra, Australia, and a subset of 18 samples by inductively coupled plasma mass spectrometry (ICP-MS) at the Trace Element Analysis Laboratory, Monash University, Melbourne, Australia. Methods for SSMS have been described by Taylor & Gorton (1977). Methods for ICP-MS are similar to those described by Price et al. (1997). USGS standard BHVO-1 was used as the calibration standard and to monitor accuracy during ICP-MS and SSMS analysis. Precision for both analytical techniques is generally better than 5% at the 95% confidence level.
Sr–Nd–Pb isotope ratios for 18 samples were measured on a Finnigan-MAT 262 thermal ionization mass spectrometer at La Trobe University, Melbourne, Australia, using methods similar to those described by Kamenetsky & Maas (2002). Samples were selected from those analysed by SSMS and ICP-MS. Sr isotope ratios were normalized to 86Sr/88Sr = 0·1194 and adjusted to be compatible with NBS-987 87Sr/86Sr = 0·71024. Nd isotope ratios were normalized to 146Nd/144Nd = 0·7219 and adjusted to be compatible with La Jolla 143Nd/144Nd = 0·511860. Pb isotope data were corrected using a 204Pb–207Pb double spike, which produces highly accurate and precise data (usually 0·05%; see Woodhead et al, 1995). Four SRM-981 loads were also run with the double-spike and yielded the following double-spike-corrected ratios (±2): 206Pb/204Pb = 16·934 ± 0·004; 207Pb/204Pb = 15·489 ± 0·005; 208Pb/204Pb = 36·698 ± 0·012. These results agree very well with the recommended values (16·937, 15·491 and 36·700; Richards, 1986). No age correction was applied to the data because of the young age of the rocks.
Minerals from 52 samples were analysed using a modified JEOL-JXA(5A) electron probe micro-analyser at the Department of Geology, University of Auckland, with an electron gun accelerating voltage of 15 kV, 600 pA sample absorption current, and 3 µm spot diameter. Calibrations were performed using nine international mineral standards. Precision is 1% RSD for element oxides that make up >25 wt % of a given mineral phase to >10% RSD for element oxide contents that are <2 wt %.
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ROCK TYPE CLASSIFICATION |
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The South Auckland basalts are classified based on their CIPW-normative compositions in accordance with the classification scheme of Johnson & Duggan (1989)
. They are divided into two broad groups (A and B) based on differences in petrography and mineralogy that are marked by distinct geochemical characteristics (Table 1). Group A comprises alkali ol-basalts, transitional basalts, hawaiites, and ol- and quartz (qz)-tholeiitic basalts. They are predominantly subalkaline, hy-normative, and have relatively low light REE (LREE) and HFSE abundances. Group B comprises nephelinites, basanites, ne-hawaiites, alkali ol-basalts, and mugearites, which are alkaline, ne-normative, and have comparatively high LREE and HFSE abundances.
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PETROGRAPHY |
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The South Auckland basalts are porphyritic, predominantly holocrystalline, and poorly vesicular. Detailed petrographic descriptions for each rock type and modal mineralogy data have been given by Cook (2002)
. The petrographic features of each Group A and B rock type and phenocryst compositions are summarized in Table 2.
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Group A
Group A basalts are relatively coarse-grained and consist of olivine ± clinopyroxene ± plagioclase phenocrysts. Olivine is the dominant phase. Phenocrysts are typically subhedral and except for qz-tholeiitic basalts, occasionally contain small Cr-titanomagnetite inclusions. Clinopyroxene is generally subhedral, colourless, pale green, or light brown. It is commonly subordinate to plagioclase and occasionally absent in transitional and ol-tholeiitic basalts. Plagioclase is the only feldspar present and varies from euhedral to anhedral phenocrysts. Titanomagnetite and ilmenite occur as rare microphenocrysts. The groundmass is hyalopilitic, intersertal, or intergranular, and contains abundant plagioclase with subordinate clinopyroxene, olivine, titanomagnetite, and ilmenite. Apatite, K-feldspar, chlorite, and smectite are minor accessory groundmass phases, and calcite and zeolite are occasional amygdaloidal infillings.
Clusters of olivine ± clinopyroxene ± plagioclase, up to 6 mm across, are common. Some have disequilibrium textures, i.e. corroded margins lined with opaques, reaction rims of pyroxene grains, and crystals with kink-band metamorphic textures. Some olivine and clinopyroxene phenocrysts have similar textures, which suggests possible derivation from xenolith disaggregation. Rare partially resorbed quartz and feldspar xenocrysts rimmed by Fe–Ti oxides or pyroxene grains occur in some hawaiites and ol-tholeiitic basalts.
Group B
Group B basalts contain olivine phenocrysts and titaniferous clinopyroxene microphenocrysts. Olivine is the principal phase, subhedral or euhedral, and commonly contains small Cr-titanomagnetite inclusions. Clinopyroxene is purplish brown, euhedral or prismatic, and is commonly rimmed by dark, purplish pink, high-Ca titanaugite or diopside. Clinopyroxene is the dominant phase in some basanites and ne-hawaiites. Rare plagioclase microphenocrysts occur in some alkali ol-basalts and mugearites. The groundmass is fine-grained hyalopilitic, intersertal, intergranular, or pilotaxitic, and consists of plagioclase and K-feldspar microlites and laths, olivine, titanaugite, titanomagnetite, and rare ilmenite in mugearites. Nepheline is found in interstitial pools.
Some basanites, ne-hawaiites, and alkali ol-basalts contain rare ultramafic xenoliths and olivine and pyroxene xenocrysts, which occasionally have corroded margins lined with opaques or pyroxene grains and exhibit rare kink-band metamorphic textures. Sanders (1994) examined dunite, lherzolite, harzburgite, and wehrlite xenoliths hosted by several basanites and ne-hawaiites. Based on geochemical and geobarometric calculations, Sanders concluded that they represent accidental fragments equilibrated in the spinel peridotite stability field (P = 1·65–1·7 GPa), and were derived either from an enriched upper mantle, or from a refractory residue related to a metasomatized, strongly heterogeneous mantle. Some basanites and ne-hawaiites also contain rare quartzofeldspathic xenoliths, <5 mm across, and quartz and feldspar xenocrysts, generally with reaction rims of pyroxene.
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MINERAL COMPOSITIONS |
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Olivine
Compositions of olivine phenocrysts in Group A rocks (Fo82–59) span a similar range to those in Group B (Fo92–60), with the majority of phenocrysts in each group ranging between Fo81 and Fo65. Phenocrysts from both groups are commonly normally zoned and typically consist of a comparatively large, homogeneous Mg-rich core mantled by a relatively thin (<150 µm), less forsteritic rim. The measured decrease in Fo content of rims relative to cores commonly exceeds 10%. Calculated distribution coefficients for Group A samples are typically
, outside the range of olivine–host-rock equilibrium (i.e. 
; Roeder & Emslie, 1970; Herzberg & Zhang, 1996). In contrast, most Group B samples contain equilibrated and poorly equilibrated phenocrysts, a feature comparable with the ne-normative basalts from eastern Australia (see Ewart, 1989).
Clinopyroxene
Clinopyroxene phenocrysts in Group A are predominantly augite (Wo47–36En50–36Fs23–11; Fig. 3) with comparatively low TiO2 contents (0·6–2·3 wt %). They are commonly normally zoned, occasionally reversely zoned, and have rims typically within 2·5 wt % MgO, TiO2, Al2O3, and 4·0 wt % FeO* of their adjacent cores. In contrast, clinopyroxenes in Group B are predominantly diopside (Wo52–45En46–32Fs20–10) with TiO2 contents in the range 1·0–5·2 wt %. Most crystals in Group B lavas are normally zoned with cores enriched in MgO, Al2O3, and TiO2, up to 4·0, 6·9, and 3·7 wt %, respectively, and slightly depleted in FeO*, typically <2·0 wt %, relative to rims. Calculated distribution coefficients for the majority of Group A samples fall outside the range of clinopyroxene–host-rock equilibrium (i.e. 
; Liotard et al., 1988), whereas 
values for Group B indicate that most samples have mixed populations of equilibrated and poorly equilibrated crystals.
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Feldspar
Plagioclase phenocrysts in Group A are principally labradorite. They typically exhibit weak normal zoning (core–rim pairs are within the range of labradorite compositions), although some phenocrysts are strongly zoned (e.g. labradorite cores with bytownite, An72, to oligoclase, An30–19Or8–4, rims). Rare K-feldspar (Or43–17) occurs in the groundmass of some samples. Feldspar phenocrysts in Group B are rare and range from labradorite (An60–58) to sanidine (Or38).
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WHOLE-ROCK COMPOSITIONS |
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Major elements
Representative geochemical compositions for Group A and B lavas are presented in Tables 3 and 4 respectively. Group A lavas range from subalkaline to mildly alkaline (Na2O + K2O = 3·0–4·6 wt %), whereas those of Group B have mildly to strongly alkaline compositions (Na2O + K2O = 3·3–7·9 wt %; Fig. 4). Although lavas from both groups span a similar range of Mg-number (47–68) with most >60, the variation diagrams shown in Fig. 5 illustrate that there are large differences in major element oxide abundances at a given MgO content. Group A lavas generally have higher SiO2, lower TiO2, Na2O, K2O, and P2O5, and broadly similar
, Al2O3, and CaO contents compared with those from Group B. Group B rocks exhibit marked increases in Na2O, K2O, and P2O5 with decreasing MgO compared with the relatively weak trends of Group A rocks. In contrast, Al2O3 contents in each group show strong negative correlations with MgO. Group A rocks exhibit a weak trend of increasing CaO with decreasing MgO, with relatively large variations of CaO at a given MgO content, especially between 8 and 10 wt % MgO. In contrast, Group B rocks show a relatively strong positive CaO–MgO correlation. Group A rocks have Ca/Al (wt %) values in the range 0·5–0·7, whereas in Group B, Ca/Al ranges from 0·5 to near-chondritic values (i.e. 1·1; McDonough & Sun, 1995).
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Compatible and incompatible trace elements
The Ni and Cr contents of rocks from each group span a similar range and exhibit relatively strong positive correlation with MgO (Fig. 6). Nickel contents vary from 324 ppm for the most primitive samples (i.e. Group B SAB179) to 51 ppm for those with evolved compositions (i.e. Group A SAB162), whereas Cr contents range from 396 ppm to 89 ppm, respectively, for these samples. The abundances of Sc and V in both groups generally overlap, but in Group A, Sc–MgO and V–MgO correlations are poor, whereas in Group B, Sc and V contents generally decrease with increasing MgO (not shown).
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Group A is characterized by relatively low LILE, HFSE, and LREE abundances that show little variation with MgO content (Fig. 6). In this respect, Group A lavas are similar to the hawaiites from the Ngatutura Field to the south (Briggs et al., 1990
). In contrast, Group B exhibits a relatively large range of incompatible element concentrations that increase systematically with decreasing MgO, similar to the Okete basanites (Briggs & Goles, 1984). The incompatible element patterns of Group B lavas (Fig. 7b) have strong OIB-like characteristics (e.g. Sun & McDonough, 1989); Nb and Ta enrichment relative to the LILE (Rb, K, and Ba) and an overall decrease in enrichment in HFSE and heavy REE (HREE) abundances. These features are similar to those for alkaline basalts from the Okete, Ngatutura, and Auckland fields of the Auckland Province, and the eastern Australia fields (e.g. O'Reilly & Zhang, 1995; Zhang & O'Reilly, 1997; Zhang et al., 2001). However, they contrast with those of some basalts from the Northland Province, which have negative Nb anomalies (Huang et al., 2000). The patterns for Group A lavas are also OIB-like (Fig. 7a) but they are less enriched than Group B, especially from LREE to middle REE (MREE) (La to Eu), a feature similar to the low-Mg hawaiites and ol- and qz-tholeiitic basalts from eastern Australia (e.g. Ewart et al., 1988; O'Reilly & Zhang, 1995). In addition, the enrichment of Nb and Ta in most samples is less prominent than in Group B rocks, but abundances of Ti, Y, and Yb in both groups are similar to average OIB (Sun & McDonough, 1989). The absence of strong negative Sr anomalies in Groups A and B suggests that Sr behaved incompatibly for the entire suite of South Auckland lavas.
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Group B lavas have higher Ce (60–210 ppm) and U (0·87–3·3 ppm) contents than those in Group A (19–51 ppm) and (0·10–0·66 ppm), respectively, whereas Pb contents in Group B (2·8–7·4 ppm) overlap those of Group A lavas (1·2–5·7 ppm). Group B lavas have relatively high Ce/Pb (18–42) and U/Pb (0·27–0·62) values comparable with similar basalts with HIMU–OIB-like characteristics from the Auckland Field (Huang et al., 1997
), eastern Australia (e.g. Zhang et al., 2001), and SE China (e.g. Zou et al., 2000). In contrast, the relatively low Ce/Pb (4–22) and U/Pb (0·03–0·29) values of Group A are similar to those of the transitional to qz-tholeiitic basalts in the Newer Volcanic Province of southeast Australia (Price et al., 1997).
All Group A and B lavas are enriched in LREE relative to HREE, a feature typical of alkalic intraplate basalts. Significant negative or positive Eu anomalies (0·96 < Eu/Eu* < 1·13) are absent. The chondrite-normalized REE patterns for Groups A and B (Fig. 8) are comparable with those of similar rock types from the Okete, Ngatutura, and Auckland fields, and also the eastern Australia fields (e.g. Zhang & O'Reilly, 1997; Price et al., 1997), although some Group A lavas exhibit small negative Ce anomalies. The negative Ce anomalies could have resulted from localized incipient weathering (e.g. Price et al., 1991), although no petrographic evidence for weathering has been detected. Group B rocks have high to very high LREE and low HREE, high (La/Yb)N values (12–47), and steep normalized patterns, whereas Group A rocks have lower LREE, and are relatively enriched in MREE and HREE, resulting in lower (La/Yb)N values (3·4–7·6) and flatter normalized patterns that cross over those of Group B.
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Radiogenic isotopes
Sr–Nd–Pb isotopic ratios for representative Group A and B lavas are given in Table 5. Despite their generally distinct major and incompatible element compositions, the two groups have very similar isotopic compositions.
Nd values for Group A (+6·0 to +6·7) and Group B (+6·1 to +6·7) are identical, and considerable overlap exists for 87Sr/86Sr values (Fig. 9). However, most Group B samples (0·70274–0·70285) are more tightly clustered and tend to be less radiogenic than the more diverse Group A samples (0·70281–0·70318). Sr–Nd isotope ratios for South Auckland lavas are similar to those for basalts from the Northland (Huang et al., 2000), Auckland (Huang et al., 1997), and Okete fields (Briggs & McDonough, 1990), although the last have slightly higher 87Sr/86Sr and lower Nd values. In contrast, the South Auckland basalts have Sr–Nd isotopic compositions that are distinct from most eastern Australian Cenozoic basalts, including the Newer Volcanic Province (McDonough et al., 1985; Price et al., 1997), and basalts from the TVZ (Gamble et al., 1993). In terms of modern mantle components, Sr–Nd isotope ratios for Groups A and B fall between the fields for DMM (i.e. Pacific MORB) and HIMU-OIB.
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, where (143Nd/144Nd)CHUR = 0·512638.
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Pb isotope ratios for Group A and B lavas are also similar, although Group A lavas are more diverse and range to slightly lower 206Pb/204Pb and higher 207Pb/204Pb values (Fig. 10a). Group A lavas with the lowest 206Pb/204Pb values overlap the Okete Field, whereas those with higher 206Pb/204Pb values, and the tightly clustered Group B lavas, overlap the Auckland Field. All three volcanic fields show only limited Pb isotopic overlap with basalts from the Northland fields. These relationships suggest a range of mantle source Pb isotope characteristics in the region. The South Auckland basalts have higher 206Pb/204Pb values than those of Pacific MORB, Newer Basalts, the TVZ, most eastern Australian Cenozoic basalts, and the DMM, EMI, and PREMA mantle components of Zindler & Hart (1986)
. Pb isotope ratios for Groups A and B plot close to the Northern Hemispheric Reference Line (NHRL); in detail, Group A lavas have 207Pb/204Pb values above, and some 208Pb/204Pb values below the NHRL, whereas Group B lavas form a cluster straddling the NHRL in both Pb isotope ratio plots (Fig. 10a and b). Unlike intraplate continental basalts from eastern Australia (e.g. Zhang et al., 2001), eastern China (e.g. Zou et al., 2000, and references therein), and SE Asia (e.g. Zhou & Mukasa, 1997), none of the South Auckland samples possess the elevated 
7/4 and 8/4 values characteristic of southern hemispheric DUPAL-type OIB (Hart, 1984), and typical of Indian Ocean MORB and plumes.
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MAGMA GENERATION IN THE SOUTH AUCKLAND VOLCANIC FIELD |
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Primary magmas
The geochemical features of Group A and B basalts suggest qualitatively that they evolved as discrete volcanic lineages that are not related by processes such as variable degrees of melting of a common mantle source or by fractional crystallization of a common parental magma. Although none of the samples studied possess all the requisite features to be considered a primary magma in equilibrium with its mantle source (i.e. cognate peridotite xenoliths, Mg-number
68, MgO 11 wt %, and Ni >300 ppm; e.g. Frey et al., 1978), Frey et al. (1978) and Price et al. (1997) have shown that compositions similar to Group A (SA14, SA69, SAB150, SAB174, and SAB187; Table 3) and Group B (SA05, SA55, SA60, SAB135, and SAB179; Table 4) can be derived largely by olivine fractionation of primary partial melts generated from a peridotite source with Mg-number 88–89 and containing olivine of Fo88–90. Therefore, to evaluate the generation and evolution of Group A and B lavas quantitatively, we first estimated primary magma compositions for the above samples by adding olivine Fo90 to each composition, while maintaining a constant 
, until a primary composition with Mg-number 68 and olivine Fo89–91 as its liquidus phase was reached. The calculated primary compositions (Table 6) are used in later sections to estimate: (1) melt segregation pressures of discrete magma batches, which help constrain the mineralogical characteristics of the upper mantle source for Groups A and B; (2) the degree of partial melting required to generate primary Group A and B magmas; (3) the composition of the source (or sources) for Groups A and B.
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Mineralogical characteristics of the South Auckland source region
The strong HREE depletion relative to LREE and relatively high (La/Yb)N values observed in the Group B lavas suggests that garnet was a residual phase in their source (e.g. Langmuir et al., 1977
; Frey et al., 1978). Residual garnet in the source regions of alkalic basalts has also been suggested for the neighbouring Okete, Ngatutura, Auckland, and Northland fields (Briggs & Goles, 1984; Briggs & McDonough, 1990; Briggs et al., 1990; Huang et al., 1997, 2000). In contrast, Group A lavas have lower LREE, generally higher or overlapping HREE, less fractionated REE patterns, and lower (La/Yb)N values than Group B. These features suggest that primary Group A magmas could have been derived from either (1) a LREE-enriched garnet-bearing source, in which garnet was totally consumed during relatively large degrees of melting, thus enriching the melt in HREE, or (2) a source with comparatively low LREE abundances in which garnet was absent.
High-pressure melting experiments show that garnet is a solidus phase at >2·5 GPa for anhydrous conditions (e.g. Herzberg & Zhang, 1996) or >2·0 GPa in the presence of small amounts of water (i.e. 0·2 wt % H2O; Green, 1973b). Calculated magma segregation pressures (Table 6) based on the empirical relationship [ln P(GPa) = {[5·04 MgO/(SiO2 + MgO)] – 0·12 SiO2 + 7·47}/10] from Albarède (1992) indicate that Group A primary magmas were generated outside the garnet stability field at pressures ranging from 1·3 to 1·9 GPa, whereas those from Group B were generated within the range of garnet stability, from 2·5 to 4·7 GPa. In addition, a number of melting experiments involving garnet and spinel peridotite compositions (e.g. Green 1973a, 1973b; Jaques & Green, 1980; Hirose & Kushiro, 1993; Hirose & Kawamoto, 1995; Falloon et al., 2001) have demonstrated that melts generated below the spinel–garnet transition will be hy-normative and less silica-undersaturated (i.e. Group A), whereas strongly silica-undersaturated, ne-normative, basanitic and nephelinitic liquids are generated at higher pressures, above the spinel–garnet transition (i.e. Group B). These studies also demonstrate that alkalic basalts, similar to primary Group B compositions, can form by <2 to 15% melting of a garnet peridotite source, whereas basalts similar to primary Group A compositions can be generated by 7–30% melting of a spinel peridotite source. Hence, from the above data, a garnet peridotite source is suggested for Group B, and a spinel peridotite source is suggested for Group A.
Magmatic processes
Fractional crystallization
The distinctive geochemical trends exhibited by the suites of Group A and B lavas (Figs 5 and 6), for example, the decrease in Ni and Cr, are qualitatively consistent with evolution by olivine and clinopyroxene fractionation of distinct parental magmas along separate lineages. However, attempts to quantitatively model fractional crystallization processes for each lineage using the least-squares method of Stormer & Nicholls (1978) for major elements and mass balance calculations for selected trace elements (Arth, 1976), failed to yield acceptable results. Generally, these models resulted in large residual major element contents, large differences between observed and calculated trace element concentrations, and yielded mineralogically unacceptable results where the removal of a mineral phase from the parental magma was inconsistent with the observed modal abundance. Hence these quantitative models are not presented here. Furthermore, evidence that the variations within Group A and B lavas were not solely related by fractional crystallization is supported by (1) a lack of clear differentiation trends for the incompatible elements and REE (Fig. 6), and (2) a lack of uniform incompatible element and REE ratios.
Evaluation of crustal contamination, assimilation–fractional crystallization (AFC), and late Cenozoic arc magmatism
The compositions of the South Auckland basalts indicate that Group A and B magmas have not been contaminated by continental crust or modified by crustal assimilation or AFC processes. Quartzofeldspathic xenoliths armoured by a reaction rim of clinopyroxene occur rarely in some Group A and B lavas, but there is no geochemical or isotopic evidence that they have modified the magmas. Evidence for this includes their low 87Sr/86Sr (
0·7032) and ‘high’ Nd (>+6), and lack of strong positive 87Sr/86Sr–K and 87Sr/86Sr–Rb correlations that would be expected had assimilation of K- and Rb-bearing crustal material accompanied fractional crystallization. Furthermore, most Group B lavas have Ce/Pb (26 ± 7) and Nb/U (39 ± 6) values similar to those in basalts derived from N-type MORB or OIB-like sources unaffected by crustal contamination or sediment recycling (Ce/Pb 25 ± 5, Nb/U 47 ± 10, Hofmann et al., 1986; Sun & McDonough, 1989; Sims & DePaolo, 1997), which contrasts with the lower Ce/Pb (13 ± 6) but similar Nb/U (42 ± 10) values of Group A lavas (Tables 3 and 4). However, binary mixing models for South Auckland basalts failed to demonstrate that the low Ce/Pb values of Group A result from the incorporation of a Pb-enriched upper-crustal component such as the Murihiku (Pb 19 ± 4 ppm, Middleton, 1993) or Waipapa (Pb 18 ± 5 ppm, N. Mortimer, unpublished data, 2002) basement terrane metasediments into primitive Group A magmas.
Despite the relative close proximity of the South Auckland Field to the Coromandel and Taupo volcanic zones, the geochemical characteristics of Group A and B lavas show no evidence of an arc component in their source associated with the subduction of the Pacific Plate during the late Cenozoic. Evidence for this includes the absence of negative Nb and Ta anomalies [compare the intraplate alkalic basalts of the Chugoku district, SW Japan (Iwamori, 1992), the Whangarei Field, Northland, New Zealand (Huang et al., 2000) and the back-arc TVZ basalts (Gamble et al., 1996)], low Ba/Nb, Ta/Yb and Th/Yb values consistent with intraplate basalts (Fig. 11), and binary mixing models that fail to establish that the low U/Pb and Ce/Pb values for Group A lavas result from the recycling of Pb-enriched sediments such as those offshore from the TVZ (Pb 19 ± 4 ppm; Gamble et al., 1996).
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Partial melting
The predominantly monogenetic nature of volcanism in the South Auckland Field and lack of an apparent spatial or temporal correlation in composition between Group A and B lavas suggests that (1) relatively small magma batches were generated from different parts of the mantle source regions for both groups at similar times, and (2) melting of these sources occurred near-continuously during the life of the field.
Rafferty & Heming (1979) argued that the subalkaline (Group A-type) magmas formed by larger degrees of melting of the same source relative to their alkaline (Group B-type) counterparts. However, based on the findings of the experimental and petrological studies discussed above, it is suggested that primary Group A magmas were derived from a spinel peridotite source [e.g. similar to KLB-1 of Hirose & Kushiro (1993) with 58% olivine, 25% orthopyroxene, 15% clinopyroxene, and 2% spinel] by larger degrees of melting than was the case for primary Group B magmas, which represent relatively small degree melts of a garnet peridotite source [e.g. similar to 60% olivine, 18% orthopyroxene, 14% clinopyroxene, 8% garnet, of Green (1973a)].
The degrees of melting (F) required to produce primary magma compositions for the South Auckland basalts were determined by the dynamic melting inversion (DMI) method of Zou & Zindler (1996), modified by Zou (1998), and using the equation F = 
+ (1 – )X of Zou et al. (2000), where X is the mass fraction of the melt extracted from the initial solid, derived by the DMI method, and = 8·5 x 10–3, the mass porosity as defined by Zou & Zindler (1996). The criteria for the DMI method have been summarized by Zou et al. (2000). Basanite SAB179 (Group B) and ol-tholeiitic basalt SAB187 (Group A) were chosen for modelling purposes because their primitive compositions required the least amount of olivine addition (<1% and 3%, respectively) to derive putative primary compositions. In addition, these samples have the lowest incompatible element and REE abundances of their respective groups and presumably represent the highest degree of melting for their groups. Hawaiite SA69 from Group A, and basanite SA05 and nephelinite SAB135 from Group B, were also chosen for modelling because they combine a relatively primitive character (i.e. requiring the addition of 11%, 8%, and 9% olivine, respectively, to obtain Mg-number of 68) and comparatively high incompatible element and REE abundances. These compositions could have formed by lower degrees of partial melting and thus provide a lower limit to the degree of melting involved in South Auckland magma genesis.
The modelling results (Table 7) indicate that F estimates for Group A and B primary magmas are likely to vary over the range of 5–12% and 3–8%, respectively. The F estimates for basanite SA05 and SAB179, nephelinite SAB135, and ol-tholeiitic basalt SAB187 are consistent with the F values predicted by the experimental studies of Green (1973a, 1973b), Jaques & Green (1980), Hirose & Kushiro (1993), Hirose & Kawamoto (1995), and Falloon et al. (2001). The F = 5% estimate for hawaiite SA69 is lower than estimates for the hawaiites from the Okete Field (9–12%; Briggs & Goles, 1984), or for similar melts produced by 11% melting in the experiments of Jaques & Green (1980). However, the melting experiments of Hirose & Kawamoto (1995) show that such melts can be produced at F = 7% at 1·0 GPa by adding 0·3 wt % H2O to the starting lherzolite material.
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Evaluation of mantle source geochemical characteristics
The concentrations of selected incompatible trace elements in the model mantle source (Co) for Groups A and B were calculated from the relationship Co = CLX/[1 – (1 – X)]1/[
+(1–)]D of Zou et al. (2000), where CL is the trace element concentration in the primary partial melt and D is the bulk distribution coefficient (Table 7). Co values for Group A SA69 and SAB187 and Group B SA05, SAB135, and SAB179 are presented in Table 7. The results show that the sources for each group are LILE-, HFSE-, and LREE-enriched relative to chondrites (Co)N, although the Group A source is generally less enriched than the source for Group B. Co for SA69 and SA05 is lower in LILE and HFSE than other samples in their respective groups. These variations could be related to source heterogeneity, although large-scale mantle heterogeneity is not reflected in the isotopic compositions of either group.
The modelling results also indicate that LREE abundances in the Group A source are 2·0–3·7 times chondrites, HREE abundances are generally lower than chondrites, and (La/Yb)N = 3·7–4·4 (Table 7; Fig. 12). These features are comparable with the HREE-depleted model source for some ol-tholeiites from SE Australia, considered by Frey et al. (1978) as indicative of a garnet-free source. In contrast, LREE abundances in the Group B source are 3·4–6·3 times chondrites, HREE are 1·9–3·7 times chondrites, and (La/Yb)N = 2·3–3·2, features comparable with those of garnet-bearing sources proposed for similar basalts from other continental intraplate fields [compare SE Australia (Frey et al., 1978, fig. 5), Okete Volcanics (Briggs & Goles, 1984) or SE China (Zou et al., 2000)].
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Hart & Allègre (1980)
have shown that LILE/HFSE and LREE/HFSE ratios do not strongly fractionate during partial melting of a homogeneous peridotite source with olivine, orthopyroxene, clinopyroxene, garnet, and spinel residues, or by fractional crystallization of these phases because of the highly incompatible nature of LILE, HFSE, and LREE in these minerals (i.e. Kd << 1; Halliday et al., 1995). In addition, Sr, Nd, and Pb isotope ratios are not fractionated during partial melting or fractional crystallization (Hart & Allègre, 1980). Therefore, the observed ratio of these elements and isotopes probably reflects that of the source. Based on the modelling results and the arguments discussed above, we conclude that the enrichment of Group B lavas in LILE (i.e. Ba, Th, and K), HFSE (i.e. Zr and Nb), and LREE relative to those of Group A (Figs 6–8), together with their generally lower Sr/Nb, K/Nb, Zr/Nb, and higher Zr/Y values compared with those of Group A (Tables 3 and 4), can be attributed to contrasting source characteristics. Furthermore, because LREE/HREE ratios of both groups are higher than primitive mantle, each group requires sources that are geochemically enriched.
HIMU and EMII signatures in Group A and B basalts
Recent studies on continental intraplate basalts with OIB-like characteristics similar to the South Auckland basalts attribute their enriched geochemical features to a combination of DMM or HIMU, and EMI or EMII mantle components (e.g. Cebriá et al., 2000; Zou et al., 2000; Zhang et al., 2001). Group A and B lavas have Nd values that are more similar to HIMU than DMM (i.e. Pacific MORB; Fig. 9), and Sr–Nd isotopic compositions that are higher than HIMU and lower than EMI and EMII. However, they lack the broad range of Sr and Nd isotopic compositions and negative Nd vs 87Sr/86Sr correlation that would be expected had mixing between the DMM, HIMU, and EM end-members occurred. Nevertheless, Groups A and B have trace element and isotopic characteristics that suggest their respective mantle source region had HIMU and EMII signatures. Each group has incompatible element ratios that overlap known HIMU- and EM-OIB values (Fig. 13), and Pb isotopic compositions that overlap EMII (Fig. 10a and b). Group B lavas generally plot along the NHRL between DMM and HIMU, whereas Group A lavas trend from the NHRL towards EMII. In addition, Group B lavas have relatively high U/Pb, Ce/Pb, and low K/Nb values (Tables 1 and 4), features consistent with intraplate basalts derived from sources considered to have a HIMU signature (e.g. Hofmann et al., 1986; Huang et al., 1997; Zhang et al., 2001). In contrast, Group A lavas have relatively low U/Pb, Ce/Pb, and high K/Nb values (Tables 1 and 3), comparable to similar basalts generated from sources dominated by the EM-type component (e.g. Zhang et al., 2001).
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Late Cretaceous to Recent intraplate alkalic volcanic rocks with HIMU isotopic and trace element signatures are widespread over a region in the South Pacific Ocean that includes North Island, South Island, and the sub-Antarctic islands of New Zealand; Marie Byrd Land and the Balleny Islands, Antarctica; and Tasmania, Australia (Briggs & McDonough, 1990
; Lanyon et al., 1993; Baker et al., 1994; Weaver et al., 1994; Huang et al., 1997; Price et al., 2003). Lanyon et al. (1993) attributed the HIMU signature to widespread plume activity that began prior to Cenozoic continental fragmentation of Gondwana associated with the opening of the Southern Ocean and the Tasman Sea (see Weaver et al., 1994; Fig. 1). However, there is no geochemical evidence to suggest that the HIMU signature observed in the South Auckland basalts was directly related to late Cenozoic asthenospheric plume-related magmatism. For example, the 206Pb/204Pb values of Group A and B lavas, as indeed in the alkaline basalts of the Okete and Auckland fields, are lower than in typical HIMU-OIB (e.g. St Helena: 206Pb/204Pb = 20·896; Sun & McDonough, 1989). Furthermore, there is no evidence for hotspot activity in the South Auckland region because the direction of the relative plate motion contrasts with the age trend of the four volcanic fields in the Auckland Volcanic Province (Fig. 1). This suggests that the HIMU signature, and by inference the EMII signature, observed in the South Auckland basalts could be a feature of relatively shallow lithospheric mantle source regions (see Huang et al., 1997; Zhang et al., 2001).
Numerous studies have concluded that the range in trace element and isotopic compositions between HIMU- and EM-type OIBs, such as those observed in Group A and B lavas, reflects the relative contribution of subducted oceanic crust and sediment in the source region (e.g. Palacz & Saunders, 1986; Weaver et al., 1986; Dupuy et al., 1988; Weaver, 1991a, 1991b; Woodhead, 1996). Weaver (1991a, 1991b) argued that the characteristic positive Nb and Ta anomaly of OIB-like magmas, such as those exhibited by Group A and B lavas, is an artefact of ancient (Mesozoic or earlier) subduction-related dehydration processes that retain these elements in a subducted slab so that they are recycled into the mantle. Long-term storage and isotopic evolution of recycled oceanic crust can reproduce the Sr, Nd, and Pb isotopic characteristics of a HIMU source (Chauvel et al., 1992). This process occurs deep in the mantle and may take between 1 and 2 billion years (Weaver, 1991b; Chauvel et al., 1992; Woodhead, 1996). However, the relatively high, positive Nd values for Groups A and B (Nd > +6) imply that source enrichment in the South Auckland region has been ‘recent’ (e.g. Zou et al., 2000). Long-term storage would have resulted in Nd isotopic compositions evolving to negative EM-like Nd values.
Huang et al. (1997) interpreted the low Pb isotopic compositions (e.g. 206Pb/204Pb = 19·321) of the Auckland Field basalts, relative to HIMU, as evidence that their HIMU feature is relatively young and that the source region for the Auckland Field will develop St. Helena-like Pb isotopic values in 500 Myr. The source region for Group A and B lavas, which have similar Sr–Nd–Pb isotopic compositions to the Auckland Field basalts, could have been affected in a similar way by an ancient (Mesozoic or earlier) subduction-related event. One possibility is modification of the sub-continental lithospheric mantle by processes associated with a prolonged period of subduction (e.g. Baker et al., 1994) when northern New Zealand was situated at the eastern margin of Gondwana during the Mesozoic, or an even older subduction-related event.
In the case of Group B, their relatively high, HIMU-like, U/Pb and Ce/Pb values could be due to a Mesozoic (or earlier) metasomatic enrichment of U and Ce in their source region [compare the Cooktown nephelinites, Australia (Zhang et al., 2001)]. Fluids generated by partial melting of the subducting slab could have caused metasomatism of the sub-continental lithosphere, thus enriching this region in LILE, HFSE, and LREE, which are now reflected in Group B lava compositions. However, this process could also account for some of their EMII-like incompatible element and isotopic characteristics (e.g. Woodhead, 1996). In contrast, if the Group A source was affected in this way then their HIMU signature has been partially masked. Weaver et al. (1986) and Weaver (1991a, 1991b) argued that the EM-type source trace element characteristic is due largely to mixing between HIMU-OIB and terrigenous sediment, which has been subducted with oceanic crust. Alternatively, Hawkesworth et al. (1986) proposed that EM-type geochemical features could result from the delamination of metasomatized-enriched sub-continental lithosphere. Because northern New Zealand was located at the eastern margin of Gondwana prior to Cenozoic continental fragmentation, the EMII component of Group A lavas may have been inherited from the supercontinent. The comparatively low LREE enrichment (Fig. 8), U/Pb, and Ce/Pb of Group A, together with the lack of evidence for crustal contamination probably reflects derivation from a source with low U and Ce but enriched in Pb. The low U/Pb and Ce/Pb values could be the effect of U–Pb and Ce–Pb decoupling during subduction-related metasomatism (e.g. Miller et al., 1994; Brenan et al., 1995; Gamble et al., 1996). However, the lack of evidence for source modification by late Cenozoic arc magmatism (e.g. Pb enrichment) requires the metasomatic enrichment to have occurred earlier, possibly during the Mesozoic.
Processes associated with plume-related magmatism that may have begun in mid-Cretaceous time (Lanyon et al., 1993; Weaver et al., 1994) and subsequent melting could also have imparted an EM-type signature into the Gondwana sub-continental lithosphere. Price et al. (2003) proposed that widespread plume activity during the Mesozoic prior to continental fragmentation could chemically precondition the lithosphere. Sanders (1994) argued that the upper mantle and lower crust of northern North Island have been metasomatized, and cited the presence of volatile-bearing minerals (i.e. amphibole, mica, apatite, carbonate, and sulphide globules) in a variety of upper-mantle xenolith types from the alkalic basalts in the Okete, Ngatutura, and South Auckland fields as evidence for this. Therefore, Group A basalts could have resulted from the partial melting of the remnants of a metasomatized Gondwana sub-continental lithosphere with HIMU and EMII signatures.
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MODEL FOR GENERATION OF SOUTH AUCKLAND VOLCANIC FIELD MAGMAS |
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Our preferred model for the generation of the South Auckland Volcanic Field Group A and B magmas is presented in Fig. 14. In this model, intraplate tensional forces occurring 160 km NW of the active TVZ arc triggered partial melting of lithospheric fragments from the Gondwana supercontinent. There is no geochemical evidence for any arc component in the South Auckland Volcanic Field magmas. Therefore, the partial melting must have taken place in the sub-continental lithospheric mantle west of the present subducted slab. The inferred presence of garnet in the Group B source requires melt generation at pressures >2·0 GPa (c. 60 km). Group B magmas were generated in the sub-continental lithosphere within a zone of low-velocity mantle material. Group A magmas formed at shallower depths in a spinel peridotite source at <2·0 GPa. The temporal and spatial randomness of the lavas that make up each group indicates coeval magma generation in the respective source regions for Groups A and B, and contemporaneous ascent of Group A and B magmas to the surface. This contrasts with an earlier model by Rafferty & Heming (1979)
, which required Group A-type magmas to be generated at depths greater than those for Group B and that the eruption of Group A-type magmas preceded Group B compositions.
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This model also requires mantle metasomatism as the cause for the distinct incompatible element and REE characteristics observed in Group A and B lavas, possibly the result of a protracted Mesozoic subduction-related metasomatic event when northern New Zealand was situated at the eastern margin of Gondwana. A problem with this model is the generation of subalkaline (Group A-type) melts in the sub-continental lithospheric mantle at the high degrees of melting (i.e. F = 20–30%) proposed by Frey et al. (1978)
and McDonough et al. (1985) for similar basalts from eastern Australia and the experimental studies previously cited. A possible solution is to produce Group A magmas under hydrous conditions. Falloon et al. (2001) have demonstrated that the melting of a hydrous, spinel peridotite mantle source would allow ol-tholeiitic magmas with compositions similar to those of Group A primary magmas to be generated at lower temperatures and by relatively lower degrees of partial melting (as implied by the dynamic melting models for Group A). Supporting evidence for possible hydrous melting conditions in the Group A source comes from the presence of amphibole in some of the spinel peridotite xenoliths (found in Group B basanites and ne-hawaiites) described by Sanders (1994). The large scatter in incompatible elements and LREE, and lack of any systematic geochemical trends in Group A lavas, is probably due to a combination of various magmatic processes and source heterogeneity. |
CONCLUSIONS |
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The South Auckland basalts can be divided into two broad groups (A and B) based on their contrasting petrography, mineralogy, and major and trace element compositions. Group A comprises silica-undersaturated, hy-normative, transitional to qz-tholeiitic basalts. Group B comprises strongly silica-undersaturated, ne-normative, basanites to ne-hawaiites. Group A basalts have notably lower LILE, HFSE, and LREE abundances than those from Group B, and have relatively flat chondrite-normalized REE patterns that cross over the steep REE patterns of Group B. Although Group A and B lavas have similar Sr–Nd–Pb isotopic compositions that range between the DMM, HIMU, and EMII mantle reservoirs, Group A lavas exhibit relatively weak and variable OIB-like incompatible element characteristics, whereas Group B lavas have a strong OIB-like signature.
Group A and B basalts evolved as a set of distinct volcanic lineages that do not appear to be related by processes such as variable degrees of melting of a common mantle source or by fractional crystallization of a common parental magma, and there is no apparent temporal or spatial pattern between or among the various lava types that make up each group. There is no petrological, isotopic, or other geochemical evidence to suggest that any of the basalts from either group were extensively contaminated by continental crust, or modified in any way by late Cenozoic subduction-related processes. The distinct geochemical differences between Group A and B basalts reflect differences in their respective mantle sources, which is supported by contrasting incompatible element ratios, such as K/Nb, Sr/Nb, and Zr/Nb. The estimated incompatible element and REE abundances for Group A and B sources indicate that Group A magmas were derived from an enriched upper-mantle source with relatively low concentrations of LILE, HFSE, and LREE, and low LREE/HREE ratios, whereas the Group B source had high incompatible element abundances and LREE/HREE ratios.
The strong positive correlation between decreasing Ni and Cr with MgO indicates that fractionation of olivine and clinopyroxene was qualitatively an important but not sole process in the evolution of the suites of Group A and B lavas. However, results of fractional crystallization models and mass balance calculations suggest that Group A and B basalts are not genetically related to a common parental magma. Also, the modelling results strongly suggest that although fractional crystallization played a role in the evolution of the lineages that define Groups A and B, it cannot account for the broad range of compositions exhibited by each group. Therefore, Group A and B basalts probably are derived from discrete magma batches formed by variable degrees of partial melting of distinct sub-continental lithospheric mantle sources and evolved independently by variable degrees of fractional crystallization.
Results of partial melting models, together with estimated melt segregation pressures, indicate that the primary magmas for Group A basalts were derived by 5–12% partial melting of a spinel peridotite source at depths of <2·5 GPa (i.e. <75 km). In contrast, the primary magmas for Group B were derived by comparatively smaller degrees of melting (i.e. 3–8%) of a garnet peridotite at depths >2·5 GPa (>75 km). This is supported by the relatively high (La/Yb)N values (12–47) of the Group B basalts, indicative of magma generation in the presence of residual garnet, which contrasts with the low (La/Yb)N values (3·4–7·6) of Group A basalts that would be expected for magmas generated from a garnet-free source.
Group A and B basalts are associated with partial melting of metasomatized sub-continental lithospheric mantle remnants following fragmentation of the Gondwana supercontinent. Extensional tectonic events affecting the South Auckland region could have created conditions that promoted adiabatic decompression melting of these widely dispersed lithospheric fragments. Each source region for Groups A and B had HIMU- and EMII-OIB signatures. However, there is no evidence for a HIMU asthenospheric plume source for the South Auckland basalts, nor is there isotopic evidence to suggest that mixing between two or more of the mantle end-members DMM, HIMU, or EM occurred. Each group has Pb isotopic characteristics indicative of a relatively young HIMU-OIB signature that may have resulted from metasomatic processes related to Mesozoic or earlier subduction when New Zealand was at the eastern margin of Gondwana.
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ACKNOWLEDGEMENTS |
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L. Frick, K. Palmer, R. Sims, and J. Wilmshurst provided technical support for this project. We thank Nick Mortimer of the Institute of Geological & Nuclear Sciences, New Zealand, for permission to use his unpublished data on the Murihiku and Waipapa Terranes. R. Price is thanked for his constructive review of the manuscript. Roger Briggs would like to thank S. R. Taylor for his co-operation and support while at the Research School of Earth Sciences, Australia National University. Financial support for this project was provided by the Department of Earth Sciences, The University of Waikato. The thorough reviews of the manuscript and constructive comments by R. Arculus and D. Gust are gratefully acknowledged.
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FOOTNOTES |
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Present address: School of Earth Sciences, University of Melbourne, Parkville, Vic. 3010, Australia. 
* Corresponding author. Telephone: 064-07-8384466, ext. 6415. Fax: 064-07-8560115. E-mail: c.cook{at}waikato.ac.nz
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