JOURNAL OF PETROLOGY VOLUME 51 NUMBER 5 PAGES 1121^1141 2010 doi:10.1093/petrology/egq015 Chronology and Evolution of Caldera-forming and Post-caldera Magma Systems at Okataina Volcano, New Zealand from Zircon U^Th Model-age Spectra B. L. A. CHARLIER1 AND C. J. N. WILSON2 1 DEPARTMENT OF EARTH AND ENVIRONMENTAL SCIENCES, THE OPEN UNIVERSITY, WALTON HALL, MILTON KEYNES MK7 6AA, UK 2 SCHOOL OF GEOGRAPHY, ENVIRONMENT AND EARTH SCIENCES, VICTORIA UNIVERSITY, PO BOX 600, WELLINGTON 6040, NEW ZEALAND RECEIVED MAY 11, 2009; ACCEPTED MARCH 10, 2010 ADVANCE ACCESS PUBLICATION APRIL 16, 2010 U^Th disequilibrium model-age data are presented for zircons from four young eruptive units from Okataina volcano, New Zealand. These data highlight contrasts in the time-scales over which eruptible silicic magma bodies are generated and accumulated below a highly active rhyolite volcano prior to and following the 61ka caldera-forming Rotoiti eruption (80 ^ 120 km3 magma). The Rotoiti event was followed by 12 explosive eruptions of the Mangaone Subgroup between 45 and 30 ka. A change in eruptive styles between 30 and 25 ka brackets the major Oruanui eruption of Taupo volcano 80 km away; subsequently eight rhyolitic, lava-dominated and two basaltic explosive eruptions occurred between 25 ka and 1886 CE. We present (238U^230Th) zircon model-age data determined by secondary-ion mass spectrometry (SIMS) and thermal ionization mass spectrometry (TIMS) from the Rotoiti, Ngamotu (Unit B) (45 ka), Mangaone (Unit I) (33 ka) and Rotorua (15·4 ka) eruptive units. We couple these data with published and new compositional information to trace magma crystallization and storage patterns. Population density curves of SIMS model ages from zircons in the two pumice types (biotite-free and biotite-bearing) from the Rotoiti eruption differ. Zircons from the former yield a model-age peak coincident with eruption age, whereas zircons from the latter show a peak at 70^90 ka and little variation inTIMS model-age values from different crystal size fractions. Concentration weighted means of model ages from the two pumices, however, are the same within 1 SD error, and their Sr isotopic values overlap at 2 SD precision, suggesting that they are *Corresponding author.Telephone: (þ44) 1908 652558. Fax: (þ44) 1908 655151. E-mail: [email protected] genetically linked. Model-age spectra for the Ngamotu, Mangaone and Rotorua pumices are complex, indicating recycling of crystals from multiple older populations that largely pre-date the 61ka caldera-forming eruption. Superimposed on the older age spectra are variably developed younger pre-eruptive suites of ages reflecting varying amounts of crystallization prior to each eruption. A lack of commonality in these younger peaks in the post-caldera eruptive rocks, along with compositional and isotopic differences between (and sometimes within) the eruptive units, collectively precludes their origin from a single melt-bearing mush system. Okataina contrasts with Taupo, where comparable-age eruption deposits have simpler age spectra, consistent with larger-scale crystallization cycles and thermal events in the magma chambers there. When compared with other caldera-related silicic systems for which suitable data are available, Okataina differs in lacking a simple pre-eruptive history prior to its caldera-forming event and having a complex, non-coherent magmatic history of post-caldera eruptions. KEY WORDS: zircon; U-series dating; rhyolite; Taupo Volcanic Zone; Okataina Volcano I N T RO D U C T I O N A knowledge of the dynamics of large-scale silicic magmatism and associated volcanism provides important insights ß The Author 2010. Published by Oxford University Press. All rights reserved. For Permissions, please e-mail: journals.permissions@ oxfordjournals.org JOURNAL OF PETROLOGY VOLUME 51 into processes that build and modify the continental crust, as well as generating the basis for assessment of hazards and possible forecasting of future eruptions. Information from the timing, sizes and styles of volcanic outbursts can be combined with geochemical, isotopic and crystalspecific studies of the eruption products to build detailed pictures of the nature of silicic magma chambers in the lead-up to, during and following a climactic calderaforming eruption. In particular, single-crystal (using secondary-ion mass spectrometry; SIMS) and multiplecrystal (using thermal ionization mass spectrometry; TIMS) dating of zircons by U-series techniques (U^Th disequilibrium or U^Pb) has proved effective, when combined with other data, in showing over what time-scales parental magma chambers can produce small to vast volumes of eruptible (i.e. melt-dominant) silicic magma. Three aspects of large silicic magmatic systems are explored here: (1) the accumulation times for magma involved in the caldera-forming event; (2) the nature of the caldera-related chamber, in particular whether the thermal^magmatic events associated with the growth of the chamber impose a uniform pattern of age spectra; (3) how fast post-caldera activity is re-established and what aspects of the caldera-forming system are recognizable in post-caldera events. We present a study of four eruptions at Okataina volcano in the central Taupo Volcanic Zone of New Zealand, compare these data with published results from a parallel study at nearby Taupo volcano, and make comparisons and contrasts with other well-characterized silicic systems. Worldwide, the compositions, timings and zircon age spectra of post-caldera volcanic eruptive rocks highlight the diverse behaviour of large silicic magmatic systems. Several examples (e.g. Long Valley, Valles) show geologically rapid magmatic (and structural) resurgence within thousands to a few tens of thousands of years, and involve compositions related to the less-evolved dregs of the climactic magma chamber (Stix et al., 1988; McConnell et al., 1995; Hildreth, 2004; Phillips et al., 2007). Other systems show more complex post-caldera geochemical relationships over periods of hundreds of thousands of years, notably the Yellowstone system, where age and O-isotopic data from zircon in particular show evidence for recycling of crystals pre-dating the caldera-forming episode (Bindeman & Valley, 2001; Bindeman et al., 2001, 2008; Christiansen et al., 2007; Girard & Stix, 2009; Vazquez et al., 2009). In all these systems, however, the geochronological detail of eruptive events and magmatic processes is constrained by the 104^105 year uncertainties in 40Ar/39Ar and U^Pb systematics, respectively. In contrast, Okataina and Taupo have eruptive frequencies 1^2 orders of magnitude more rapid than the ‘classical’ caldera systems cited above. These eruption records are linked mostly to 14C constraints with 1 SD uncertainties of 102^103 years, and NUMBER 5 MAY 2010 the magmatic chronologies can be inferred by U^Th disequilibrium techniques on zircon with 103^104 year 1 SD uncertainties. There is thus an opportunity in the New Zealand examples to interrogate magma chamber processes at large silicic volcanoes to a greater level of temporal precision at more frequent intervals. At both Okataina and Taupo volcanoes the host magma chambers are demonstrably capable of generating and disposing of large volumes of melt-dominant silicic magma on exceptionally short time-scales. At Taupo, for example, a combination of geochemical and zircon model-age data imply that the 530 km3 of magma evacuated in the 27 ka Oruanui caldera-forming eruption was effectively assembled in less than 3000 years, and that 6000 years later, no compositional linkages can be discerned between the Oruanui magma and the first dacitic and subsequent rhyolitic post-caldera eruptive rocks (Sutton et al., 2000; Wilson & Charlier, 2009). At Okataina, the rhyolitic Rotoiti caldera-forming eruption occurred at 61ka, followed by at least 20 silicic plus two basaltic eruptions since then, with a wide range in compositional and mineralogical characteristics. Numerous studies (see next section) have considered these characteristics, but the only crystal-specific age information for Okataina is the zircon model-age data from deposits of the Rotoiti eruption and the immediately following, but independent, Earthquake Flat eruption (Nairn & Kohn, 1973; Charlier et al., 2003). O K ATA I N A VO L C A N O Okataina and Taupo are the world’s two most frequently active and productive rhyolitic volcanoes (Figs 1 and 2: Wilson et al., 1995, 2009; Nairn, 2002). Their rapid eruption frequencies yield snapshots of their magmatic processes at intervals typically an order of magnitude shorter than at other silicic caldera volcanoes worldwide (Fig. 3). When compared with Taupo, Okataina is somewhat less productive (3·5 km3/kyr vs 9·6 km3/kyr: volumes from Jurado-Chichay & Walker, 2000; Nairn, 2002; Wilson et al., 2009), but its eruptions have more even spacing and less widely variable eruption volumes (Fig. 3). Okataina has been active at least since the calderaforming eruption of the Matahina ignimbrite (Bailey & Carr, 1994) at 322 7 ka (1 SD) and may have seen earlier activity back to 4500 ka (new 40Ar^39Ar ages from G. S. Leonard & A. T. Calvert, personal communication, 2008). The probably complete record of recent activity (Howorth, 1975; Jurado-Chichay & Walker, 2000; Nairn, 2002; Fig. 3) began with the Rotoiti eruption (Nairn, 1972). Following a period of quiescence the 12, possibly more, eruptions of the Mangaone Subgroup occurred from 50 ka onwards (Howorth, 1975; Jurado-Chichay & Walker, 2000; Smith et al., 2002), generating small to large Plinian fall deposits, at least two of which were accompanied by ignimbrite emplacement. The largest of these, also 1122 CHARLIER & WILSON OKATAINA VOLCANO MAGMA SYSTEMS Fig. 1. Location maps. (a) Index map of the Taupo Volcanic Zone (TVZ) in New Zealand. (b) Index map of the rhyolite (R) dominated central Taupo Volcanic Zone, showing its boundaries with the andesite (A) dominated northern and southern Taupo Volcanic Zone segments. Named silicic volcanoes are those known or inferred to have been active in the last 61kyr. specifically named Mangaone (Howorth, 1975) occurred at 33 ka. Earlier eruption products of the Mangaone Subgroup are dacites to low-silica rhyolites and later eruptions involved moderate- to high-silica rhyolites (Huppert, 1981; Nairn, 2002; Smith et al., 2002; Shane et al., 2005a). No lava domes are preserved from this time period, but relatively high abundances of dense glassy juvenile clasts and/or lava lithic fragments in some deposits suggest that lavas, now destroyed or buried, may have been extruded in association with these or additional eruptions. Following the 27 ka Oruanui eruption from Taupo, 80 km to the SW, the style of eruptive activity changed markedly at Okataina. Eight rhyolitic eruptions (most of which involved a basaltic component: Leonard et al., 2002; Shane et al., 2007, 2008a) plus two basaltic pyroclastic eruptions have occurred (Nairn, 2002). All of these eruptions have involved multiple vents aligned along trends that imply that they were fissure-fed by dikes at depth. The rhyolitic eruptions were prolonged in duration (years to decades; e.g. Nairn et al., 2001), with numerous episodes of pyroclastic activity within single eruptions and multiple lava extrusions (Nairn, 2002). The rhyolitic eruption vents are concentrated along two alignments, and lava extrusions have accumulated to form two major elongate dome complexes (Haroharo and Tarawera), with an additional minor dome complex to the SW (at 25 and 15·4 ka, the latter being studied here) and other isolated domes (Cole, 1970; Nairn, 2002; Fig. 2). The two major dome complexes show separate mineralogical and geochemical lineages (Ewart, 1968; Cole, 1970; Smith et al., 2005) and there is no evidence for eruptions at one focus influencing the timing, volumes, or styles of activity at any other. For example, many Haroharo eruptive rocks contain cummingtonite and lack biotite, whereas Tarawera eruptive rocks lack cummingtonite and most have biotite (Ewart, 1968; Ewart et al., 1975). In addition, most post-25 ka pyroclastic products contain two or more rhyolite compositions that are inferred to have undergone separate evolutionary pathways in the crust and to have met only shortly prior to or during eruption (e.g. Nairn et al., 2004; Smith et al., 2004, 2006; Shane et al., 2007, 2008a, 2008b). D E P O S I T S ST U D I E D In the light of the geochemical complexities in many of the Okataina eruption products, we focused on four representative units to investigate the magmatic chronology of Okataina in and following the caldera-forming Rotoiti event. As well as investigating the zircon model-age spectra of the eruption products to compare with the eruptive chronology, we also determined major and trace elements and Sr-isotopic ratios on aliquots from the single- or bulk-pumice samples used for zircon extraction (Table 1). This was done to provide first-order constraints on the 1123 JOURNAL OF PETROLOGY VOLUME 51 NUMBER 5 MAY 2010 Fig. 2. Simplified geological map of Okataina volcano. The Haroharo caldera is a composite feature resulting from collapse in association with the 322 ka Matahina and 61ka Rotoiti eruptions, together with possible contributions from two or three other eruptions (Nairn, 1981, 2002). Post-caldera activity over the past 25 kyr has constructed two large composite accumulations of lava and pyroclastic deposits: Haroharo (four rhyolite eruptions) and Tarawera (one basaltic and four rhyolite eruptions). The Rotoiti and Earthquake Flat ignimbrites are inferred to have been erupted from multiple vents aligned along NE^SWand NNW^SSE trends, respectively; the first (from Nairn,1981) is hypothetical because of burial by later lavas, and the second (from Nairn, 2002) is represented by crater forms at the present day. Generalized vent areas for the Ngamotu (Unit B) and Mangaone (Unit I) eruptions are from Jurado-Chichay & Walker (2000), and mapped thick accumulations of the Mangaone ignimbrite are marked from Nairn (2002). The area of the domes and vents for the 15·4 ka Rotorua eruption are also indicated (from Nairn, 1980). 1124 CHARLIER & WILSON OKATAINA VOLCANO MAGMA SYSTEMS Fig. 3. Plot of approximate eruption volumes (magma, after Jurado-Chichay & Walker, 2000; Nairn, 2002; Wilson et al., 2009) versus time for eruptions at Okataina and in the Taupo^Maroa area for the past 61kyr. Labelled units in the Okataina panel are those studied here, and the units discussed by Wilson & Charlier (2009) are labelled in the Taupo^Maroa panel. compositions and consanguinity or otherwise of any contrasting magma compositions (see Wilson & Charlier, 2009, fig. 4). Rotoiti eruption deposits The Rotoiti eruption is the largest known from Okataina in the last 300 kyr, with magma volume estimates ranging from 80 to 120 km3 (Nairn, 2002; Shane et al., 2005b; Wilson et al., 2009). The eruption generated the non-welded Rotoiti ignimbrite and the coeval, widespread Rotoehu fall deposit, with collapse of the northern sector of the composite Haroharo caldera occurring towards the end of the eruption (Nairn, 1972, 2002). The eruption age is 61·0 1·4 ka (1 SD) from ages of lavas bracketing the fall deposit on Mayor Island volcano (Wilson et al., 2007). Vent locations are poorly constrained, largely because of the widespread, even distribution and fine-grained nature of the fall deposits (Walker, 1979; Newnham et al., 2004); Nairn (1981) proposed multiple vents aligned along a fissure coincident with the vent alignment for the post-27 ka Haroharo dome complex (Fig. 2). Petrological studies (Davis, 1985; Schmitz & Smith, 2004; Shane et al., 2005b) indicate two pumice types in the ignimbrite. The dominant composition (‘T1’ of Shane et al., 2005b) has cummingtonite4hornblende in the ferromagnesian fraction and lacks biotite. The second pumice type, erupted later, contains biotite plus a second hornblende composition, but is itself inferred to represent a mixture of the ‘T1’ magma with a biotite- and hornblende-dominant, more evolved magma (‘T2’ of Shane et al., 2005b). Samples P1567 (multiple clasts) and 123/1 (single clast) analysed by Charlier et al. (2003) were biotite-bearing pumices sampled near the top of the Rotoiti ignimbrite. Shane et al. (2005b), on the basis of their recognition that the biotite-bearing pumices were mixtures of two independently sourced magmas, contended that the model-age data presented by Charlier et al. (2003) were from unrepresentative, mixed populations of zircons. To test our earlier results, we sampled biotite-free pumices (P1747) for zircons from the earlier, biotite-free Rotoiti ignimbrite at the Maungarangi Road site sampled by Shane et al. (2005b), at a locality at U15/ 050622 (grid reference to the nearest 100 m in the New 1125 JOURNAL OF PETROLOGY VOLUME 51 NUMBER 5 MAY 2010 Table 1: Summary of geochemical and isotopic data for the units studied Sample no.: Unit: Age (ka): SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Rb Sr Y Zr Nb Ba Pb Th U Sc V Cr Co Ni Cu Zn Ga 87 Sr/86Sr* 2 SD (ppm) U (ppm) Th (ppm) (234U/238U) 2 SD (238U/232Th) 2 SD (230Th/232Th) 2 SD Zircon thermometry M T (8C) Fe–Ti oxide thermometry T (8C) P1568 Earthquake Flat 61·0 1·4 123/1 Rotoiti (bi-bearing) 61·0 1·4 73·73 0·315 13·99 2·33 0·075 0·50 2·11 3·94 2·95 0·061 1·88 100·5 148 18·7 141 7·2 776 13 11 2 4 19 7 2 3 3 34 14 0·705287 9 2·3 9·6 1·010 0·003 0·725 0·001 0·748 0·011 70·21 0·260 12·34 1·65 0·062 0·51 1·87 4·82 2·40 0·054 3·47 77·3 156 17·9 134 6·1 765 10 8 2 3 12 4 2 3 3 30 12 0·705364 8 1·8 7·6 1·005 0·005 0·735 0·004 0·738 0·010 74·77 0·32 13·83 2·22 0·09 0·51 1·89 3·90 2·44 0·02 2·22 75·9 157 19·2 155 7·1 779 9 9 3 4 12 7 3 6 3 35 12 — 1·40 775 1·63 754 1·29 791 702–8051 7662 P1567 Rotoiti (bi-bearing) 61·0 1·4 1·7 7·0 1·004 0·004 0·735 0·006 0·722 0·005 7662 P1747 Rotoiti (bi-free) 61·0 1·4 P1839 Unit B P1565 Mangaone 45 2 33 1 P1810 Rotorua (bi-free) 15·4 0·3 P1816 Rotorua (bi-bearing) 15·4 0·3 77·02 0·26 12·39 1·71 0·07 0·38 1·64 3·89 2·59 0·05 2·01 75 130 20·3 150 6·4 826 11 8 2 3 17 1 1 2 3 29 11 0·705379 8 — — — — 0·683 0·008 0·730 0·001 69·36 0·61 16·04 3·81 0·12 0·96 3·10 3·87 1·97 0·17 4·87 58 215 32·2 194 7·7 693 10 6 1 10 28 3 1 3 3 62 16 0·705198 10 — — — — 0·740 0·011 0·742 0·002 74·63 0·273 14·51 1·73 0·093 0·26 1·32 4·40 2·75 0·033 4·06 82·7 121 35·2 205 10·0 854 14 9 2 3 7 2 1 0 3 56 14 0·705340 7 2·1 9·0 1·009 0·004 0·702 0·003 0·701 0·011 74·56 0·33 13·47 2·20 0·07 0·47 1·97 4·07 2·78 0·07 1·82 92 152 26·6 222 7·2 784 14 9 2 5 17 2 1 3 2 41 14 0·705338 10 — — — — 0·755 0·011 0·744 0·002 73·79 0·28 14·15 2·04 0·07 0·42 1·88 4·13 3·19 0·06 2·91 122 140 22·4 169 6·6 828 16 12 3 3 15 1 1 3 4 42 14 0·705416 10 — — — — 0·710 0·009 0·735 0·001 1·36 783 1·36 806 1·25 820 1·42 814 1·40 790 7122 9253 7953 8354 7504 Whole-rock powders were prepared using an agate ball mill, and analysed on a ARL8420 þ WDS XRF instrument. Sr isotope ratio determinations were carried out using the techniques detailed by Charlier et al. (2006). The Neptune MC-ICP-MS instrument was operated in static collection mode using a normalizing value of 86Sr/88Sr ¼ 0·1194 and the exponential law relationship for fractionation correction. Repeated analysis of NBS 987 on the day of analysis gave 87 Sr/86Sr ¼ 0·710239 13 (18·8 ppm) 2 SD (n ¼ 9). To allow direct comparison with other published data, the Sr isotope data were normalized to 87Sr/86Sr ¼ 0·710250 for NBS 987. U-series data for P1568, 123/1, P1567 and P1565 were measured by thermal ionization mass spectrometry on a Finnegan MAT262 mass spectrometer using an RPQ-II energy filter. All errors are 2s in-run errors. Activity ratios were calculated from isotopic ratios using the following decay constants: 232 ¼ 4·94752 10–11 a–1; 230 ¼ 9·1577 10–6 a–1; 238 ¼ 1·55125 10–10 a–1. The (230Th/232Th) external reproducibility was monitored by repeat analyses of an in-house standard, Th’U’std, which gave 230Th/232Th ¼ 6·173 0·038 10–6 (0·7% 2 SD: n ¼ 22) during this study. This probably underestimates the sample external reproducibility and, based on replicate zircon samples from this study and others carried out in the Open University laboratory over the same period of time, we estimate the external reproducibilities as 51·5% for (238U/232Th), 52% for (230Th/232Th) and53·5% for (234U/238U). External reproducibilities for (230Th/232Th) (run unspiked on the Neptune MC-ICP-MS system) and (238U/232Th) (from ICP-MS) were 50·5% and 51·5% respectively. Sr, U and Th procedural blanks were always 5100 pg for each element, which are negligible compared with the 100 ng of sample analysed. Zircon saturation thermometry was calculated using the equations presented by Watson & Harrison (1983), whereas Fe–Ti oxide thermometry values are taken from 1Molloy et al. (2008), 2Shane et al. (2005b), 3Smith et al. (2002) and 4Smith et al. (2004). Likely uncertainties on both methods of thermometry are 208C. LOI, loss on ignition. 1126 CHARLIER & WILSON OKATAINA VOLCANO MAGMA SYSTEMS Zealand metric map grid). A single pumice from the P1747 suite along with powder from the biotite-bearing single pumice (123/1) were also analysed for Sr isotopes. Table 2: Summary of zircon samples analyzed Sample Nature Eruption unit Ngamotu (Unit B) eruption deposits The earliest post-Rotoiti fall deposit was named the Ngamotu Tephra by Howorth (1975). Recognition of a weakly developed palaeosol within this tephra by Jurado-Chichay & Walker (2000) led them to divide it into Units A and B, and they found an additional, new deposit above Unit B and labelled it Unit C. Unit B represents the greater part of the Ngamotu Tephra of Howorth (1975). It is of moderate volume (1·2 km3, magma) and has Plinian dispersal characteristics. A broadly defined source area was proposed by Jurado-Chichay & Walker (2000) in the SW sector of the Haroharo caldera (Fig. 2), aligned with the lineament of multiple vents that have created cumulatively the post-27 ka Tarawera dome complex. Unit B pumices are moderately crystal-rich and rhyodacitic (69· 4% SiO2, volatile-free: Table 1), broadly similar to the succeeding crystal-poor Maketu (Unit D) and Hauparu (Unit F) rhyodacites (Smith et al., 2002; Shane et al., 2005a). The age of Unit B is not constrained by radiocarbon dating, but is inferred by us to be 45 ka, based on stratigraphic relationships and the development of palaeosols between this deposit and others above (Units D and F at 38 ka: Jurado-Chichay & Walker, 2000) and below (Rotoiti at 61ka: Wilson et al., 2007). A reasonable error estimate on this age is thus likely to be 2 kyr (1 SD). A suite of 1^4 cm diameter pumices (P1839) was sampled from Unit B from a section at V16/421386. The slightly older Unit A was not studied because it is weathered and finer grained at the sites accessible to us, and gave inadequate yields of small zircon crystals. Mangaone (Unit I) eruption deposits This eruption, the largest of the Mangaone Subgroup, produced a Plinian fall deposit (4· 2 km3, magma) and an extensive, thin non-welded ignimbrite (1·8 km3, magma; Jurado-Chichay & Walker, 2000). More recently, however, a partially welded ignimbrite that locally exceeds 50 m thickness, and that had been mapped to the NE of Okataina and previously been assumed to be 240 kyr old (Beresford & Cole, 2000), has been correlated with the Mangaone eruption (K. Spinks, personal communication, 2001; Nairn, 2002). Its volume is poorly constrained because of uncertainty over the extent of co-eruptive caldera collapse (with the possibility of significant volumes of intracaldera tuff), but it is inferred here to be of the order of 10 km3, magma. The eruption age is estimated at 33 calibrated ka (Jurado-Chichay & Walker, 2000), with uncertainties constrained by radiocarbon age determinations on older and younger units and inferred to be 1kyr (1 SD). The vent is sited in the SE portion of Haroharo caldera, co-aligned with vents for the succeeding Grains Number of analysed analyses with finite ages Rotorua (15·4 0·3 ka) P1810 Multiple Fall deposit top 47 47 (46) Fall deposit base 48 48 (48) Fall deposit base 55 53 (55) Multiple Fall deposit (coarsest 44 38 (43) pumices part) Upper ignimbrite (Re3) 45 40 (45) Middle ignimbrite (Re2) 26 24 (26) pumices P1816 Multiple pumices Mangaone (33 1 ka) P1565 Multiple pumices Unit B (45 2 ka) P1839 Rotoiti (61·0 1·4 ka) P1747 Multiple pumices P1567 Multiple pumices Details of samples and eruption units are given in the text. Crystals with ‘finite ages’ denote those that plot at least 1 SD below the equiline. Numbers in parentheses denote those used in concentration-weighted isotopic means and used for plotting PDF curves. Tarawera dome complex and the vent area for Unit B (Jurado-Chichay & Walker, 2000; Fig. 2). Juvenile pumices are rhyolitic (74^76% SiO2; Nairn, 2002; Smith et al., 2002; Table 2) and crystal poor. Pumices (P1565) were collected for zircon extraction from the lower part of the Plinian fall deposit at a site at V16/418318. To provide average model ages for the zircon population and for comparison with information from the SIMS analyses, we analysed three aliquots of 563 mm and two of 63^250 mm zircon size fractions by TIMS. Zircons extracted from the 125^250 mm sieve fraction were analysed by SIMS. Rotorua eruption deposits The Rotorua eruption occurred at 15· 4 0· 3 ka (calibrated radiocarbon age, 1 SD uncertainty; Lowe et al., 2008). The eruption was from a group of closely spaced, NNW^SSE-aligned vents on the SW extension of the vent lineament proposed for the Rotoiti eruption and the multiple post-27 ka vents that extruded lavas to build the Haroharo dome complex (Fig. 2). A Plinian pumice fall deposit with a magma volume of 0· 4 km3 was followed by extrusion of a 0· 6 km3 compound lava dome with minor pyroclastic activity (Nairn, 1980; Smith et al., 2004; Kilgour & Smith, 2008). The eruption began with a crystal-poor rhyolite then, in turn, increasing amounts of 1127 JOURNAL OF PETROLOGY VOLUME 51 crystal-rich biotite-bearing rhyolite are observed in the fall deposits, whereas the lava and its accompanying pyroclastic deposits are composed of the crystal-rich, biotite-bearing rhyolite. Smith et al. (2004) proposed that these two rhyolites were independently generated and stored prior to eruption. We sampled crystal-poor, biotitefree (P1810) and crystal-rich, biotite-bearing (P1816) pumice clasts from the lower and upper parts, respectively, of the fall deposits at the type locality (at map reference U16/017313) previously described by Smith et al. (2004) and Kilgour & Smith (2008). A N A LY T I C A L T E C H N I Q U E S Samples used for zircon separation are listed in Table 1; the separation techniques used were those of Charlier et al. (2005). For SIMS U^Th analyses by SHRIMP-RG (sensitive high-resolution ion microprobe^reverse geometry) using the joint USGS-Stanford instrument, techniques used were based on those in Lowenstern et al. (2000) and Charlier et al. (2005). Zircons were mounted in epoxy resin, polished, photographed in reflected light, and imaged by cathodoluminescence. The mount was acid rinsed, coated with 100 nm of Au and left in the sample chamber overnight to reach full vacuum. Using a 34 nA 16 ^ O or 16 O 2 primary ion beam, a 50 mm 50 mm square region was rastered for 2 min to remove the Au coat and any surface contamination. A flat-floored elliptical pit 2 mm 25 mm 37 mm was then excavated into the zircon during analysis. This liberated approximately 4^6 ng of sample that was sent as positive secondary ions to the mass spectrometer. Data were collected in 10 scans per point for 90Zr216O, 230Th16O, 232Th, 232Th16O, 238U and 238 16 U O. Dwell times ranged from 2 to 40 s for each peak. In particular, to allow for the generally low U contents of the zircons, 230Th16O and the background (at mass 246·16) were measured for 40 s each. Additionally, a 4 s measurement at mass 244 (‘ThC’) was used as a check for the beam impinging on the epoxy resin in the mount. If the count rate for mass 244 significantly exceeded the background count rate the analysis was discarded, as inevitably the resulting meaningless data lay in the sector of Th-excess [i.e. with (238U/230Th) 41] on the equiline diagram. In all cases we were attempting to determine rim ages, with grains being chosen that had adequate U contents (as judged by the greyscale of the cathodoluminescence image). A U^Th fractionation factor was empirically determined through the repeated analysis of several zircon standards run from the same mounts as those for the unknowns. Samples and standards were run at a ratio of 4:1. The standards included, at various stages: (1) MAD (concentration standard), a 555 Ma gem-quality zircon from Madagascar with U ¼ 4196 ppm and Th ¼1166 ppm NUMBER 5 MAY 2010 (F. K. Mazdab, personal communication, 2006), (2) CZ3 (concentration standard: 550 ppm U), (3) R33, (4) VP10 and (5) AS57 (see Ireland & Williams, 2003; Charlier et al., 2005; Lowenstern et al., 2006). Given their antiquity, 238 U and 230Th activities in these zircons are at secular equilibrium and therefore, after the application of a U^Th fractionation factor, the calculated (230Th/238U) should equal unity. This was determined on a mount-bymount basis using the measured 230Th16O þ /238U16O þ ratios. To achieve a weighted mean (230Th/238U) of unity for the standards, fractionation factors varied between 1·04 and 1·11 over the analysis period, the reciprocals of which were applied in the calculation of the (238U/232Th) of the Okataina zircons for the relevant day of their analysis. Repeated analysis of the standard allowed us to arrive at a conservative best-estimate 1s error of 3% on the fractionation factor, which translates into a corresponding 3% error on the (238U/232Th) values. This uncertainty is substantially larger than that calculable from the count statistics alone and hence is inferred to give a more robust limit to total errors in the (238U/232Th) values (see Charlier et al., 2005). For the (230Th/232Th) values, only the uncertainties derived from the count statistics could be measured and incorporated in the model-age determinations. This is because, globally, there is no age standard with constant (230Th/232Th) with which to quantify the external reproducibility. In some cases, therefore, for some grains with highTh contents (low U/Th ratios) the uncertainties on the (230Th/232Th) ratios may be underestimated and yield anomalously precise ages. When close to the eruption age, this may lead to apparent small inconsistencies between the crystallization and eruption ages. Values of the (238U/232Th) and (230Th/232Th) ratios for whole-rock samples P1568, 123/1, P1567 and P1565, and for various zircon size fractions in sample P1565 were obtained by isotope dilution TIMS techniques and reported by Charlier (2000). Whole-rock determinations for samples P1747, P1839, P1810 and P1816 were obtained by a different approach, using instruments at The Open University. The (230Th/232Th) was determined on unspiked dissolutions by multicollector inductively coupled plasma mass spectrometry (MC-ICP-MS) using a Thermo-Finnigan Neptune instrument equipped with a retarding potential quadrupole, whereas the (238U/232Th) was calculated from the U and Th elemental ratio by referencing to a set of solutions of known U/Th ratio using quadrupole (Q)-ICP-MS. Repeat analysis of one whole-rock powder (P1209) using both TIMS, and MC- and Q-ICP-MS approaches yielded data that were essentially identical within analytical uncertainty. The external reproducibility of the (230Th/232Th) using either approach was always better than 1% (2 SD). Sr isotopic values (values and details are given in Table 1) were determined using the methods of Charlier et al. (2006). 1128 CHARLIER & WILSON OKATAINA VOLCANO MAGMA SYSTEMS 230 Th^238U isochron ages were calculated as two-point model ages by referencing each of the zircon analyses in turn to the respective whole-rock (WR) analyses on the equiline diagram. The zircons are considered to have the same initial (230Th/232Th) as the host rocks; that is, they plot in a horizontal array on the equiline diagram at the time of their crystallization. Samples that are older than 5 times the half-life of 230Th (i.e. 350 kyr) plot within error of the equiline at a (230Th/232Th) value that reflects the (238U/232Th) value. The only age information that can be obtained from these samples is that their age is older than 350 ka. To compare the overall model-age spectra from the SIMS analyses within and between samples, we calculated concentration-weighted isotopic mean ages by weighting all the SIMS (230Th/232Th) and (238U/232Th) values according to the U and Th concentrations of each spot analysis (to the nearest 100 ppm). Z I RC O N M O D E L - A G E A N D S r I S O T O P I C DATA Rotoiti eruption deposits range of model ages, but P1747 has a more pronounced cluster that peaks around the inferred 61ka eruption age. Concentration-weighted mean ages for the SIMS model ages for P1747 and P1567 are the same within error (Table 3) and both are older than, but still within error of, the TIMS analyses for samples P1567 and 123/1 reported by Charlier et al. (2003) and summarized in Table 4. The Sr isotopic values for the biotite-free pumices (P1747) and biotite-bearing pumice (123/1) overlap at the 2 SD level (Table 1). Ngamotu (Unit B) eruption deposits Forty-four grains were analysed, of which 38 gave finite ages (Fig. 6; Electronic Appendix 2). The model ages show a broad spectrum, with peaks from 60 to 80 ka and around 140 ka, whereas the concentrationweighted mean of all the analyses with finite ages is 160 þ 28/^22 ka (Table 3). The Sr-isotopic value for this unit is the lowest of any of the units considered in this paper (Table 1). Mangaone (Unit I) eruption deposits Analyses from 26 grains (24 with finite ages) were presented by Charlier et al. (2003) from the biotite-bearing sample P1567, and an additional 45 analyses (40 with finite ages) obtained from the biotite-free sample P1747 (Table 2; Figs 4 and 5; see also Electronic Appendix 1, which is available for downloading at http://www .petrology.oxfordjournals.org/). Both samples show a wide Fifty-five grains were analysed, of which 53 yielded finite model ages (Fig. 7; Electronic Appendix 3). In contrast to the earlier samples, the Mangaone model-age spectrum is dominated by a peak centred on 47 ka, with 34 out of 53 model ages between 36 and 67 ka. The concentrationweighted mean for the grains with finite model ages is 77 þ13/^12 ka (Table 3). Model ages for eachTIMS analysis gave a narrow range of values (42·5 1·1 to 50· 6 1·2 ka; Fig. 4. Zircon model-age data from the biotite-free Rotoiti sample P1747. (a) SIMS (230Th/232Th) vs (238U/232Th) equiline diagram for single zircons. Reference isochrons indicate the eruption age (61ka), 100, 200 and 300 ka. Error ellipses represent 1s analytical uncertainties on (230Th/232Th) and a standard 3% error on (238U/232Th) (see Charlier et al., 2005). An isochron slope and age was determined by referencing this isotopic average datum point to the whole-rock values (Table 1) to generate a two-point model age. (b) Probability density function (PDF) curve (from Isoplot: Ludwig, 2003) and histogram based on isochron slopes derived from two-point whole-rock-zircon SIMS determinations. The PDF line is based on the two-point isochron slopes rather than the ages determined from them, as the slope uncertainty is symmetrical with respect to the slope value, whereas the absolute age is not. The concentration-weighted isotopic mean age (93 þ19/^16 ka; Table 3) is plotted here as a shaded band representing the envelope of the 1 SD uncertainties. Analytical data and model ages are given in Electronic Appendix 1. 1129 JOURNAL OF PETROLOGY VOLUME 51 NUMBER 5 MAY 2010 Fig. 5. Zircon model-age data from the biotite-bearing Rotoiti sample P1567 (from Charlier et al., 2003). Model ages and analytical data are given in Electronic Appendix 1. (a) SIMS (230Th/232Th) vs (238U/232Th) equiline diagram for single zircons, with the TIMS analyses of bulk zircon fractions marked (see also Fig. 10, below). (b) Cumulative probability density curve (with maxima marked), histogram, and a grey band denoting the envelope of the concentration-weighted mean value 1 SD uncertainties (104 þ 25/^21ka) from SIMS data on grains that returned finite ages. Other details as in Fig. 4. Table 3: Summary of concentration-weighted mean model ages of zircons in the samples analyzed in this study Eruption Material (238U/ Error 232 (230U/ 232 Th) Error Slope Error Age Error Error slope (years) þ age age 0·399 0·032 55370 5950 5640 0·360 0·028 48530 4870 4660 0·506 0·057 76690 13330 11880 0·770 0·052 159830 27870 22170 0·573 0·068 92550 18860 16070 0·617 0·080 104370 25480 20630 Th) Rotorua P1810 P1816 Zircons 2·37 0·07 1·39 0·04 Whole-rock 0·76 0·01 0·74 0·00 Zircons 2·26 0·07 1·29 0·04 Whole-rock 0·71 0·01 0·74 0·00 Zircons 2·40 0·07 1·56 0·09 Whole-rock 0·70 0·00 0·70 0·01 Zircons 3·19 0·10 2·63 0·10 Whole-rock 0·74 0·01 0·74 0·00 Mangaone (Unit I) P1565 Ngamotu (Unit B) P1839 Rotoiti P1747 P1567 Zircons 2·30 0·07 1·66 0·10 Whole-rock 0·68 0·01 0·73 0·00 Zircons 2·95 0·09 2·09 0·17 Whole-rock 0·74 0·00 0·72 0·00 Errors on the slopes are calculated as 1s a priori uncertainties. Uncertainties on the activity ratios are 3% for (238U/232Th) and 1s internal error (within-run) for (230Th/232Th) for the single data points as well as the calculated weighted means and calculated slopes. Uncertainties on the ages are also 1s. Table 4), which correspond to the peak in the probability density function (PDF) curve from SIMS analyses (star symbols in Fig. 7: see also Fig. 10, below). In contrast to the Oruanui zircon TIMS data (Charlier & Zellmer, 2000), there is no systematic age variation with crystal size fraction. The Sr-isotopic value (Table 1) is marginally lower than that for the Rotoiti magmas, distinctly higher than that for the Ngamotu, but very similar to that for the biotite-free Rotorua magma. 1130 CHARLIER & WILSON OKATAINA VOLCANO MAGMA SYSTEMS Table 4: Summary of data from multi-grainTIMS analyses of zircon separates Sample Eruption Size Th U fraction (ppm) (ppm) (234U/238U) (238U/232Th) (230Th/232Th) Age (ka) 1 SD (mm) P1565 Mangaone 563 198 187 1·006 0·006 2·874 0·018 1·404 0·008 45·2 0·5 P1565 Mangaone 563 199 184 1·004 0·004 2·817 0·016 1·420 0·005 45·2 0·5 P1565 Mangaone 563 194 185 0·997 0·005 2·892 0·022 1·431 0·006 44·1 0·5 P1565 Mangaone 63–250 172 161 1·006 0·007 2·852 0·017 1·501 0·007 50·6 0·6 P1565 Mangaone 63–250 172 171 1·007 0·007 3·029 0·019 1·465 0·011 43·3 0·6 P1567* Rotoiti 563 213 205 1·002 0·009 2·932 0·017 1·788 0·014 72·2 0·8 P1567* Rotoiti 63–125 264 234 1·015 0·009 2·571 0·073 1·644 0·012 75·8 2·4 P1567* Rotoiti 63–125 272 220 1·009 0·009 2·456 0·013 1·547 0·011 71·0 0·8 P1567* Rotoiti 125–250 189 194 1·022 0·009 3·109 0·021 1·839 0·005 69·1 0·5 P1568* EQF 63–250 462 395 1·019 0·009 2·590 0·015 2·233 0·009 173·0 2·7 Analytical techniques are given in the text. EQF, Earthquake Flat ignimbrite. *Age previously published by Charlier et al. (2003). Fig. 6. Zircon model-age data from the Ngamotu (Unit B) sample P1839. Model ages and analytical data are given in Electronic Appendix 2. (a) SIMS (230Th/232Th) vs (238U/232Th) equiline diagram for single zircons. (b) Cumulative probability density curve (with maxima marked), histogram, and a grey band denoting the envelope of the concentration-weighted mean value 1 SD uncertainties (160 þ 28/^22 ka) from SIMS data on grains that returned finite ages. Other details as in Fig. 4. Rotorua eruption deposits A total of 95 grains were analysed from the Rotorua pumices, 47 from the crystal-poor, biotite-free lithology (P1810), and 48 from the crystal-rich, biotite-bearing lithology (P1816) (Electronic Appendices 4 and 5). All analyses yielded finite ages, although analysis P1810-1.1 was rejected from the dataset because it is clearly an outlier from the main population on account of its extraordinarily high (238U/232Th) and (230Th/232Th) ratios (28· 68 and 19· 09 respectively). The two pumice types share a common peak with a PDF maximum around 28^29 ka, then their patterns diverge with increasing age (Figs 8 and 9). Zircons from P1810 have an additional broad peak centred on 51ka, then a tail-off to older ages with no clear modes. Sample P1816 shows three additional modes in the model ages, reflected also in the PDF peaks, around 40, 63 and 91ka, and a pronounced lack of model ages to match those controlling the 50 ka PDF peak from P1810. Concentration weighted means are, however, similar, at 55 6 ka (P1810) and 49 5 ka (P1816). Strontium isotopic 1131 JOURNAL OF PETROLOGY VOLUME 51 NUMBER 5 MAY 2010 Fig. 7. Zircon model-age data from the Mangaone (Unit I) sample P1565. Model ages and analytical data are given in Electronic Appendix 3. (a) SIMS (230Th/232Th) vs (238U/232Th) equiline diagram for single zircons, with the TIMS analyses of bulk zircon fractions marked (see also Fig. 10, below). (b) Cumulative probability density curve (with maxima marked), histogram, and a grey band denoting the envelope of the concentration-weighted mean value 1 SD uncertainties (160 þ 28/^22 ka) from SIMS data on grains that returned finite ages. Other details as in Fig. 4. Fig. 8. Zircon model-age data from the Rotorua biotite-free sample P1810. Model ages and analytical data are given in Electronic Appendix 4. (a) SIMS (230Th/232Th) vs (238U/232Th) equiline diagram for single zircons. (b) Cumulative probability density curve (with maxima marked), histogram, and a grey band denoting the envelope of the concentration-weighted mean value 1 SD uncertainties (55 6 ka) from SIMS data on grains that returned finite ages. Other details as in Fig. 4. characteristics for the two pumice types are markedly different (Table 1). DISCUSSION Controls on zircon age spectra at Okataina The zircon (model-)age spectra presented here (as in many other examples globally) reflect potentially complex patterns of crystal growth and dissolution during the history of a magma chamber or chambers, and growth of the melt-dominant portion that is erupted and sampled. In particular, the dominance of a young, immediately pre-eruptive peak in model ages versus older age modes reflects both the late-stage evolution of the magma (i.e. whether zircon saturation is achieved and abundant growth of new zircons swamps any earlier or inherited population) and the earlier history of contributing sources (i.e. whether they have undergone temperature^composition pathways that did not crystallize zircon or would have dissolved pre-existing grains). The mass yields of zircons reported by Charlier et al. (2005) implied that the age spectra reported at Taupo reflected the relative 1132 CHARLIER & WILSON OKATAINA VOLCANO MAGMA SYSTEMS Fig. 9. Zircon model-age data from the Rotorua biotite-bearing sample P1816. Model ages and analytical data are given in Electronic Appendix 5. (a) SIMS (230Th/232Th) vs (238U/232Th) equiline diagram for single zircons. (b) Cumulative probability density curve (with maxima marked), histogram, and a grey band denoting the envelope of the concentration-weighted mean value 1 SD uncertainties (49 5 ka) from SIMS data on grains that returned finite ages. Other details as in Fig. 4. abundance of late-stage crystallization. Samples that were neutral or undersaturated contained a high proportion of old grains, but had a much lower overall mass abundance of zircon. The Okataina examples presented here range from undersaturated to oversaturated in zircon (Table 1), using published eruption temperatures from coexisting Fe^Ti oxides (Smith et al., 2002, 2004; Shane et al., 2005b) and estimating the zircon saturation temperature for each melt using the method of Watson & Harrison (1983). At one extreme, the Ngamotu (Unit B) melt composition is strongly zircon undersaturated. The model-age spectrum (Fig. 6) is therefore interpreted to largely reflect inherited grains. An analogous situation exists for the immediately post-caldera eruptive rocks at Taupo (Sutton et al., 2000; Charlier et al., 2005), where only inherited grains are present. In contrast, all the other samples are indicated to be near-neutral to strongly oversaturated (the latter being biotite-bearing magmas in the Rotoiti and Rotorua deposits). The zircon model-age patterns do not simply reflect the degrees of oversaturation, however. The Mangaone (Unit I) sample shows the strongest development of a pre-eruptive peak in model ages (Fig. 7), yet is only mildly oversaturated, whereas the strongly oversaturated biotite-bearing magmas of the Rotoiti and Rotorua eruptions have only modest development of young age peaks (Figs 5 and 9). One feature that is more prominent in the Okataina data than at Taupo is the presence of crystals recording model ages that are younger than the proposed eruption age. This is seen in the Rotoiti, Mangaone and biotite-bearing Rotorua samples (see Electronic Appendices 1, 3 and 5). Although the accepted age for the Rotorua eruption is 15· 4 0· 3 ka (Lowe et al., 2008), two crystals yielded model ages of 8· 0 2· 4 and 12·5 þ 2·1/^2· 2 ka (all errors are 1 SD). Similarly, the Mangaone eruption is dated at 33 1ka (Jurado-Chichay & Walker, 2000; Table 4), but one crystal yielded an age of 20· 3 þ10· 5/^9· 7 ka. Four crystals from the biotite-free Rotoiti sample (P1747) record ages that pre-date eruption to as young as 40 6 ka, and one crystal from the biotite-bearing sample (P1567) yielded 29 þ 21/^17 ka. Contamination, however, can be excluded as a cause of this discrepancy for two reasons. First, the method of sample handling and zircon separation used by us should absolutely preclude sample contamination at the 1^10% level. Second, no material of an age suitable to supply any 8· 0 and 12·5 ka crystals in the Rotorua sample was being handled at any stage. The best explanation for the anomalously young ages is that the (230Th/232Th) uncertainties may have been underestimated for these grains because we have no option but to use the count statistics alone for this ratio in the calculation of the model ages (see Analytical Techniques section, above). Implications of zircon age spectra at Okataina First-order observations of the Okataina model-age spectra are the absences of any single dominant age peak, of continuity in age spectra between the deposits analysed and of any signs of changes in model-age spectra related to caldera formation. The caldera-forming Rotoiti eruption tapped two contrasting magmas and the compositional differences are mirrored in their zircon age spectra and PDF peaks (Figs 4 and 5), although their concentrationweighted isotopic mean ages are similar (Fig. 5; Table 3). A consequent inference is that no ‘magma residence time’ 1133 JOURNAL OF PETROLOGY VOLUME 51 NUMBER 5 MAY 2010 Fig. 10. Equiline diagram with a compilation of the TIMS data for various size fractions of zircons from the Earthquake Flat ignimbrite (P1568, Charlier et al., 2003), biotite-bearing Rotoiti ignimbrite pumices (P1567, Charlier et al., 2003) and the Mangaone (Unit I) sample P1565. Shape and size of the symbols denote 2 SD uncertainty envelopes. A noteworthy feature is the distribution of the Rotoiti data along a near-isochronous trend that is orthogonal to the trend that would be expected were there to have been significant mixing with zircons from the Earthquake Flat ignimbrite source (see Molloy et al., 2008). (See text for discussion.) for the melt-dominant bodies that were erupted can be established from the age data, for two reasons. First, no pre-caldera eruption units are known that might have served to indicate temporal changes in crystal model-age populations and so chart the growth of the melt-dominant magma body (as at Taupo: Charlier et al., 2005; Wilson & Charlier, 2009). Second, there is no clear dominance of a pre-eruptive peak in model ages that could be taken to indicate crystallization during the physical growth of the magma bodies erupted in the climactic event. Our model-age and available Sr isotopic data (Table 1; Schmitz & Smith, 2004) in turn imply that any distinction drawn on the grounds of mineralogy, or major-element pumice and glass chemistry, between the Rotoiti biotitebearing and biotite-free magma compositions is, however, minor (see Shane et al., 2005b). If the biotite-bearing pumice is a hybrid between the biotite-free composition represented by P1747 and a hypothetical end-member biotite-bearing composition (Schmitz & Smith, 2004; Shane et al., 2005b), then our data imply that the latter does not have a unique model-age spectrum that can be separately distinguished. The closely similar values of average model ages yielded by TIMS analyses of different size aliquots of zircons in the biotite-bearing Rotoiti pumices (Table 4; Fig. 10; Charlier et al., 2003) show also that there is no mixing relationship between populations of contrasting ages (in contrast to the Oruanui rhyolite: Charlier & Zellmer, 2000). In addition, the spreads in (230Th/232Th) and (238U/232Th) values reported in Rotoiti TIMS data (Fig. 10; Charlier et al., 2003) define a linear array that is perpendicular to the TIMS average model age for Earthquake Flat ignimbrite zircons. This observation and the Sr isotopic values (Table 1) preclude any significant mixing relationship between the Rotoiti and Earthquake Flat magmas, despite the negligible time gap between the two eruptions (see Molloy et al., 2008). The earlier post-caldera eruptions, represented here by the Ngamotu (Unit B) and Mangaone (Unit I) deposits, both involve a single magma type, but contrast greatly in their age patterns (Fig. 11). The Ngamotu data are consistent with an inferred 45 ka eruption age and show three modes (Figs 5 and 11). The youngest mode is broadly comparable with the average model ages of zircons and other mineral phases from the Rotoiti TIMS analyses (Fig. 10; Table 4; Charlier et al., 2003), but the older Ngamotu modes have no analogues in the other age spectra. 1134 CHARLIER & WILSON OKATAINA VOLCANO MAGMA SYSTEMS Fig. 11. Summary diagram to show overall features of the zircon model-age spectra for the units detailed in this study. PDF curves from each of Figs 4^9. Bold dotted line denotes the eruption age. (See text for discussion.) The Ngamotu source is geographically closest of any of the Mangaone Subgroup eruptions to the vents for the 61ka Earthquake Flat ignimbrite and its associated fall deposits (Fig. 2). Thus the 60^70 ka peak in the PDF curve of the Ngamotu model ages could be interpreted to represent a crystallization event associated with the thermal anomaly that remobilized the otherwise stagnant Earthquake Flat magma (Molloy et al., 2008). A second peak in the Ngamotu PDF curve around 100^120 ka matches a comparable peak in the limited dataset of SIMS model ages from the Earthquake Flat ignimbrite (Charlier et al., 2003, fig. 2C) and may represent another crystallization event that is common to the two sources. However, contrasting Sr isotopic values and bulk chemistries for each unit (Table 1; Davis, 1985; Smith et al., 2002; Molloy et al., 2008) show that they were from independent magma chambers. In contrast, the model-age spectrum for the Mangaone eruption is unimodal and dominated by crystals younger than the 61ka caldera-forming Rotoiti eruption. The sub-63 mm fractions (crystals with needle-shaped morphology) analysed by TIMS have higher U and Th contents, lower (230Th/232Th) and similar U/Th ratios compared with the larger size fraction. Both fractions yield model ages around 43 ka that are younger than the inferred Rotoiti eruption age and closely match the peak in the PDF curve derived from the SIMS data (Fig. 7). The larger size fraction of crystals yielded two model ages, one within the range of the sub-63 mm data (43 ka), and one slightly older (51ka), implying that these crystals do not represent a mixed population of widely varying mean model ages (see Oruanui: Charlier & Zellmer, 2000; Charlier et al., 2005). Thus most of the crystallization of zircons in the Mangaone magma post-dated the 61ka caldera collapse; however, grains from older sources were neither entirely stripped out nor swamped by younger grains during genesis of the Mangaone rhyolite. The Mangaone is the first of three large rhyolitic eruptions (units I, K and L of Jurado-Chichay & Walker, 2000) from 33 to 31ka that collectively represent 57% of the total Mangaone Subgroup volume and were erupted from the SE sector of the Haroharo caldera at Okataina. We interpret these three large eruptions to be the surface expression of a more vigorous rhyolite magma-producing system that was active between 61 and 33 ka. However, this system did not extend into the northern half of the Haroharo caldera, as venting of the voluminous, high-temperature dacites of units D and F around 36^38 ka (Huppert, 1981; Shane et al., 2005a) from that geographical area (Jurado-Chichay & Walker, 2000) would preclude the contemporaneous generation of rhyolite in the crust beneath those vent sites. The later post-caldera eruptive rocks (post-27 ka) at Okataina show the most diversity in magma types within single eruptions, as documented by other workers cited elsewhere in this paper. In the Rotorua eruption products, the spectra of ages from the two pumice types, coupled with contrasts in Sr isotopic compositions (Table 1), are consistent with proposals that the two magmas involved in this eruption were separately generated (Smith et al., 2004). However, the youngest PDF peak at 28^29 ka is common to both magma types (Fig. 11) and implies that they shared a common crystallization history (i.e. thermal history, given the evolved nature of both magmas) for 15 kyr prior to their co-eruption (see Smith et al., 2004). The storage areas for the two magmas may thus have been in close physical proximity, although staying compositionally independent. The older peaks in the PDF curves for both magma typoes notably do not match (Fig. 11), and this is taken to show that in the aftermath of the Rotoiti eruption, the source regions for the two Rotorua magmas were both chemically and thermally independent. The vents for the Rotorua eruption form a short NNW^ SSE-aligned lineament. When compared with the 410 km long lineaments associated with most of the other post-27 ka eruptions at Okataina the close spacing of the vents feeding the two Rotorua magma types implies that the magma bodies were vertically (rather than laterally) separated in the crust, reflected also by contrasts in inferred magma temperatures (Smith et al., 2004). 1135 JOURNAL OF PETROLOGY VOLUME 51 The zircon model-age data we have obtained are consistent with the ideas from geochemical arguments (e.g. Smith et al., 2005) that the melt-dominant magma bodies at Okataina are transient and are derived from chemically and isotopically distinct domains below the volcano. The model-age data, however, suggest that some of these domains shared some common aspects in their zircon crystallization histories and hence thermal histories. In addition, the lack of any one event that reset the magmatic system below the volcano, and the overlap of age spectra between eruptions of contrasting chemical and isotopic characteristics, precludes simple geochemical linkages being drawn between successive eruptions from the volcano. Thus any evolutionary trends derived from plots of geochemical parameters with time may be misleading as, for example, the Mangaone rhyolite was not derived from a source in common with the earlier Ngamotu dacite. Comparisons and contrasts in silicic magma systems Okataina versusTaupo Taupo and Okataina are individually the two most frequently active and productive rhyolite volcanoes on Earth. Despite their closely similar settings, they show distinct contrasts in their magmatic histories, as discussed below, and volcanic histories and eruptive styles (Wilson, 1993; Nairn, 2002). In comparison with the record from Taupo (Charlier et al., 2005; Wilson & Charlier, 2009), the Okataina model-age spectra are notable for their wide diversity of model ages and associated peaks in the PDF curves generated in Isoplot. In general, these observations suggest that growth of zircon and the generation and crystallization of silicic melts has been intermittently occurring under Okataina for at least the past 150^250 kyr, despite a lack of evidence for any significant volcanism for several tens of thousands of years prior to the Rotoiti eruption at 61ka (Nairn, 2002). In contrast, many Taupo magmas show evidence for a more limited number of crystallization events, most notably centred on 95 ka and 40 ka in the voluminous Oruanui magma body and its precursors and successors. It does not appear as though the crustal sources for the Okataina silicic eruptive rocks have ever undergone a single major episode of zircon crystallization, but instead the sources for all the eruptions we have studied have tapped a variety of domains with contrasting age patterns. Such diversity in model-age patterns mimics that seen in the chemical and mineralogical characteristics of the rhyolites (Smith et al., 2002, 2004, 2005, 2006; Nairn et al., 2004; Schmitz & Smith, 2004; Shane et al., 2005b, 2007, 2008a) and implies that a single ‘Okataina magma chamber’ does not exist. Over the period covered by the eruptions considered here, there are several other distinct features of the NUMBER 5 MAY 2010 magmatic systems at Okataina volcano that contrast with Taupo, as follows. The frequency with which mafic magmas have been intercepted by silicic melt-dominant bodies, or erupted as discrete deposits, is higher at Okataina. At Okataina these mafic magmas are sometimes basaltic (e.g. Leonard et al., 2002; Shane et al., 2008a), whereas at Taupo they are always basaltic andesite to high-silica andesite. At both volcanoes, however, these mafic magmas have undergone some significant degree of contamination by crustal components and in most cases undergone additional evolution by crystal fractionation prior to meeting the rhyolites (e.g. Wilson et al., 2006). The feeder systems for the mafic roots to both volcanoes are thus in themselves physically complex and do not permit uncontaminated mantle-derived magmas to reach the rhyolite magma bodies or the surface. At Okataina, over the past 25 kyr, contemporaneous but geographically separated and compositionally distinctive eruptive foci have been present, marked by the Haroharo and Tarawera dome complexes, both contained within the one composite caldera. At Taupo over the same time period, compositionally and isotopically distinct magma batches have been erupted from sources focused below Lake Taupo in systematic temporal succession; however, the envelopes for these vents overlap geographically (Wilson, 1993; Sutton et al., 2000) over a total area about half that of the combined vent envelopes for Haroharo and Tarawera. At Taupo there is an additional active magma system just NE of that volcano that has erupted sporadically (as recently as 16 ka) biotite-bearing rhyolite of contrasting chemistry and zircon model-age characteristics to the ‘mainstream’ Taupo eruptive rocks (Sutton et al., 1995; Wilson & Charlier, 2009). Okataina is more diverse in its model-age populations of zircons that yield finite ages in the U^Th system. Taupo has more zircon crystallization going on in the source region over shorter spans of time and, in most examples, old zircons are effectively stripped out or swamped to levels of the order of 1^10% or less. Okataina has more crystals coming through to eruption of a great variety of model ages, suggesting less vigour in rhyolite generation (though still unusually productive on a global scale) and less drastic changes in thermal conditions to drive the crystallization of zircon (and other mineral phases). These features presumably reflect greater rates of heat input and loss below Taupo, which serve to drive the extraordinarily high magma generation, crystallization and eruption rates. Global comparisons and contrasts Despite the abundance of studies on silicic systems worldwide, there are few direct comparisons between the magmatic histories of caldera-forming eruptions and their ‘post-caldera’ successors that can be made between our datasets from Okataina and Taupo and other comparably sized systems because of a dearth of zircon model-age 1136 CHARLIER & WILSON OKATAINA VOLCANO MAGMA SYSTEMS data that cover pre-, syn- and post-caldera eruption products. Deposits from caldera-forming rhyolites such as the Bishop Tuff (Long Valley) and Lava Creek Tuff (Yellowstone) have been investigated in some detail, but the uncertainties on the age determinations are typically much greater than those using U^Th techniques on the New Zealand rocks. The resulting temporal changes are thus less clearly discernible, although some distinctions are apparent. First, for the caldera-forming Rotoiti eruption at Okataina there was at no stage a single meltdominant magma body developed with a coherent history of crystallization, as seen at Taupo, Long Valley and Yellowstone (see Bindeman et al., 2001; Simon & Reid, 2005; Wilson & Charlier, 2009). Second, at Okataina, the caldera-forming magma chambers did not crystallize enough zircon in any plutonic mush zone to leave a distinctive signal on the subsequent post-caldera activity. In contrast, at Yellowstone, zircon age data from post-640 ka (i.e. post-Lava Creek caldera) eruption units imply that deep-seated residues from earlier eruptive episodes were remobilized and incorporated into the younger ‘post-caldera’ domes (Bindeman & Valley, 2001; Bindeman et al., 2001, 2008). The nature of post-caldera activity varies widely. Rapid resumption of post-caldera activity is seen at Okataina (after 15 kyr or less; Wilson et al., 2009), Taupo (after 6 kyr: Wilson & Charlier, 2009) and Long Valley (within uncertainty of the Bishop Tuff eruption age itself: Mankinen et al., 1986). In contrast, at Yellowstone, following the Huckleberry Ridge Tuff and Lava Creek caldera-forming eruptions, the earliest post-caldera activity does not occur until 220 and 120 kyr later, respectively (Christiansen, 2001). At Long Valley, the ‘post-caldera’ stage lasted in effect for 100 kyr (Hildreth, 2004), and at Yellowstone since the Lava Creek Tuff eruption for 380 kyr (Bindeman et al., 2008; Vazquez et al., 2009) before there were major changes in the vigour (and nature) of the magmatic systems, diminishing at Long Valley and increasing at Yellowstone. At both Okataina and Taupo, however, the ‘post-caldera’ activity has maintained a high productivity with no signs of waning. Indeed, the definition of what constitutes ‘post-caldera’ eruptions at Long Valley and Yellowstone encompasses time periods that are comparable with the average recurrence interval of caldera-forming eruptions in the Taupo Volcanic Zone (Wilson et al., 2009). Compositions of the post-caldera eruptive rocks at large silicic volcanoes also reflect complex processes of recycling versus new generation of magmatic components (melts and crystals). Age data from Taupo (Wilson & Charlier, 2009) and Yellowstone (Bindeman et al., 2008, for overview) show that zircons are being recycled, but at Okataina the age spectra of zircons from the calderarelated system are not distinctive enough for them to be uniquely identifiable in the post-caldera eruptive rocks that we have analysed. No comparable data are available for the Long Valley system. However, the compositional characteristics of the Early Rhyolites at Long Valley suggest strongly that they represent rejuvenated portions of the deeper parts of the Bishop Tuff magma chamber (Hildreth, 2004). At Taupo, the compositional and isotopic trends shown by the post-caldera eruptive rocks are wholly independent of the caldera-forming system (despite the presence of recycled zircons) and no appreciable proportion of caldera-related melt has survived to influence the composition of eruptions only 6^25 kyr after the climactic eruption (Sutton et al., 2000). At Okataina, the compositional characteristics of the post-caldera eruptive rocks do not suggest any significant involvement of remobilized Rotoiti material; instead, a wide range of magmagenerating sources appears to have been tapped, sometimes within single eruptions (e.g. Shane et al., 2007, 2008a, 2008b). There are also marked contrasts in the zircon model-age spectra in small- and large-scale eruptive deposits from long-lived silicic centres (whether caldera-related or not). For Okataina and Taupo there simply is no relationship between the volumes of the eruptions and the time differences between model-age PDF peaks and eruption ages. At Long Valley, the young, small-volume rhyolites display ‘residence times’ for their zircon populations that are comparable with or longer than the lead-in to the large-volume Bishop Tuff (Reid et al., 1997; Reid & Coath, 2000; Simon & Reid, 2005). Similarly at the Coso volcanic field, small-volume high-silica rhyolite domes show variable degrees of ‘residence time’ in their zircon model-age populations (Miller & Wooden, 2004; Simon et al., 2009) that are inferred to reflect the vigour of the magmatic system. However, the consequent relationships between the ‘pre-eruption crystallization interval’ and eruption volume illustrated by, for example, Simon et al. (2009, fig. 10) are meaningless unless the origins of the zircons carrying the various model ages can be inferred. Reactivation of a crystal mush and recycling will introduce into the meltdominant magma body older antecrystic or xenocrystic zircons that cannot be identified on age data alone. The model ages of these zircons may bear no simple relationship to the residence time of the melt-dominant body itself that eventually is tapped to generate the eruption products. In turn, the apparent contrast between long ‘pre-eruption crystallization intervals’ for small eruptions and vice versa, remarked on by Vazquez & Reid (2002) and Simon et al. (2008, 2009) and presented against the longstanding view that eruption volume is proportional to repose period (e.g. Smith, 1979; Trial & Spera, 1990; Reid, 2008), need not arise, for three reasons. (1) Newer studies and associated age information show that the ‘repose period’ used to construct such relationships 1137 JOURNAL OF PETROLOGY VOLUME 51 can be exceedingly variable and in many cases is limited by burial or destruction of evidence for the precursors to large eruptions. At the voluminous extremes, for supereruptions of magma volumes of 500 km3 or more, the ages of the youngest precursor eruptions can be very close to those of the climactic outburst. For example, the 2500 km3 Huckleberry Ridge Tuff has only one documented precursor eruption (Snake River Butte dome), the published age of which is within error of the age of the tuff itself (Christiansen, 2001). At Long Valley, the youngest lava eruption at the Glass Mountain complex was only 30 kyr prior to the Bishop Tuff (Metz & Mahood, 1985), and at Taupo, the youngest precursor eruption occurred only 3 kyr before the Oruanui eruption itself (Lowe et al., 2008). (2) As emphasized from studies at Taupo and Okataina, we distinguish between the time-scales inferred or presumed to be needed to generate the chemistry of silicic crystal-poor magmas by some combination of fractionation (e.g. Smith, 1979; Trial & Spera, 1990; Bachmann & Bergantz, 2004) or crustal melting (e.g. Annen et al., 2006; Streck & Grunder, 2008), versus the time-scales for physical assembly of the melt-dominant magma body that eventually is erupted. The latter time-scale will control the periodicity of eruptions, and in areas of active rifting such as in the Taupo Volcanic Zone (Rowland et al., 2010) the frequency of eruptions of any given size may be chaotic because of non-linear relationships between the time-scales over which eruptible magma bodies are generated and the frequency of rifting events that may trigger eruptions. Such triggering may occur through changes in the regional stress distribution and/or by accompanying mafic dike injection that can trigger the silicic magma eruption (e.g. Leonard et al., 2002; Nairn et al., 2004, 2005; Shane et al., 2007, 2008a; Wilson et al., 2009). (3) From our data presented here and elsewhere (Charlier et al., 2005; Wilson & Charlier, 2009) we suggest that there are two reasons why larger eruptions carry zircon populations that are on average closer to the eruption age than for small eruptions. The first reflects the observation discussed in the previous section that larger eruptions are the products of magma chambers with correspondingly large inputs and losses of heat to drive melt generation and differentiation of the magmas. Similar inferences have been made at Long Valley, with an increase in thermal flux and evidence for greater involvement of mantle-derived components in the lead-up to the Bishop Tuff eruption (Simon et al., 2007). The second reason why larger-volume rhyolites might have apparently shorter ‘pre-eruption crystallization intervals’ is simply because the majority of those examples studied are zirconsaturated. Thus, in any sample, younger crystals formed in the melt in the immediate lead-up to the eruption will tend to outnumber any older grains, leading to younger NUMBER 5 MAY 2010 average crystal populations and peaks in the PDF curve that are closer to the eruption age than in samples that are just saturated or undersaturated with respect to zircon. CONC LUSIONS The complexity of the magma storage systems at Okataina, previously recognized from mineralogical and geochemical studies, is reflected in the zircon model-age spectra from the deposits studied here. Zircons from six samples from four eruption units [61ka Rotoiti biotite-free; Rotoiti biotite-bearing; 45 ka Ngamotu (Unit B); 33 Mangaone (Unit I); 15· 4 ka Rotorua biotite-free; Rotorua biotite-bearing] show multiple zircon populations within the time limits of the U^Th disequilibrium technique. Pumices with contrasting mineralogies and compositions in the Rotoiti deposits have different age spectra (implying that there was no single melt-dominant body for the caldera-forming event) but have similar Sr isotopic characteristics. Post-caldera eruptions at 45 ka (Ngamotu) and 33 ka (Mangaone) have distinct compositions and show no clear inheritance of material from the caldera-forming magma system. In contrast, the two pumice types in the 15· 4 ka Rotorua eruption products are clearly isotopically different and from different sources, but share a common 28^29 ka PDF peak in their model-age spectra that suggests that the two magmas underwent a similar crystallization (i.e. thermal) history in the 15^20 kyr prior to the eruption. The rhyolite magma generation zone below Okataina is inferred to be much more heterogeneous than the corresponding volume at Taupo, and there has been no large-scale crystallization (thermal) event or events comparable with those (represented by dominant model-age peaks) below Taupo volcano. Both Okataina and Taupo are at a stage where their post-caldera eruptions represent the generation of small- to moderatevolume magma batches (up to 35 km3 erupted volume) from diverse sources, without development of a single coherent magma chamber at either volcano. At what point ‘post-caldera’ becomes ‘pre-caldera’ activity is not clear. AC K N O W L E D G E M E N T S We thank Joe Wooden, Frank Mazdab, Bettina Weigand and Brad Ito for their technical support and guidance with our U-Th dating studies at the USGS-Stanford ion probe facility. Ilya Bindeman, and two anonymous reviewers are thanked for their helpful comments. Simon Turner is also thanked for editorial handling. FU NDI NG We thank the Marsden Fund administered by the Royal Society of New Zealand (C.J.N.W.) and the UK Natural Environment Research Council (B.L.A.C.) for financial 1138 CHARLIER & WILSON OKATAINA VOLCANO MAGMA SYSTEMS support. The University of Auckland Research Committee contributed to the ion-probe costs. S U P P L E M E N TA RY DATA Supplementary data for this paper are available at Journal of Petrology online. R EF ER ENC ES Annen, C., Blundy, J. D. & Sparks, R. S. J. (2006). The genesis of intermediate and silicic magmas in deep crustal hot zones. Journal of Petrology 47, 505^539. Bachmann, O. & Bergantz, G. W. (2004). On the origin of crystal-poor rhyolites: extracted from batholithic crystal mushes. Journal of Petrology 45, 1565^1582. Bailey, R. A. & Carr, R. G. (1994). Physical volcanology and eruptive history of the Matahina Ignimbrite, Taupo Volcanic Zone, New Zealand. New Zealand Journal of Geology and Geophysics 37, 319^344. Beresford, S. W. & Cole, J. W. (2000). Kawerau Ignimbrite: a 0· 24 Ma ignimbrite erupted from the Okataina caldera complex, Taupo Volcanic Zone, New Zealand. New Zealand Journal of Geology and Geophysics 43, 109^115. Bindeman, I. N. & Valley, J. W. (2001). Low-18O rhyolites from Yellowstone: magmatic evolution based on analyses of zircons and individual phenocrysts. Journal of Petrology 42, 1491^1517. Bindeman, I. N., Valley, J. W., Wooden, J. L. & Persing, H. M. (2001). Post-caldera volcanism: in situ measurement of U^Pb age and oxygen isotope ratio in Pleistocene zircons from Yellowstone caldera. Earth and Planetary Science Letters 189, 197^206. Bindeman, I. N., Fu, B., Kita, N. T. & Valley, J. W. (2008). Origin and evolution of silicic magmatism at Yellowstone based on ion microbrobe analysis of isotopically zoned zircons. Journal of Petrology 49, 163^193. Charlier, B. L. A. (2000). U^Th isotopic constraints on the preeruptive dynamics of large-scale silicic volcanism: examples from New Zealand, Ph.D. thesis, The Open University, Milton Keynes. Charlier, B. L. A. & Zellmer, G. F. (2000). Some remarks on U^Th mineral ages from igneous rocks with prolonged crystallisation histories. Earth and Planetary Science Letters 183, 457^469. Charlier, B. L. A., Peate, D. W., Wilson, C. J. N., Lowenstern, J. B., Storey, M. & Brown, S. J. A. (2003). Crystallisation ages in coeval silicic magma bodies: U^Th disequilibrium evidence from the Rotoiti and Earthquake Flat eruption deposits, Taupo Volcanic Zone, New Zealand. Earth and Planetary Science Letters 206, 441^457. Charlier, B. L. A., Wilson, C. J. N., Lowenstern, J. B., Blake, S., van Calsteren, P. W. & Davidson, J. P. (2005). Magma generation at a large, hyperactive silicic volcano (Taupo, New Zealand) revealed by U^Th and U^Pb systematics in zircons. Journal of Petrology 46, 3^32. Charlier, B. L. A., Ginibre, C., Morgan, D., Nowell, G. M., Pearson, D. G., Davidson, J. P. & Ottley, C. J. (2006). Methods for the microsampling and high-precision analysis of strontium and rubidium isotopes at single crystal scale for petrological and geochronological applications. Chemical Geology 232, 114^133. Christiansen, R. L. (2001).The Quaternary and Pliocene Yellowstone Plateau Volcanic Field of Wyoming, Idaho, and Montana. US Geological Survey, Professional Papers 729-G, 1^143. Christiansen, R. L., Lowenstern, J. B., Smith, R. B., Heasler, H., Morgan, L. A., Nathenson, M., Mastin, L. G., Muffler, L. J. P. & Robinson, J. E. (2007). Preliminary assessment of volcanic and hydrothermal hazards in Yellowstone National Park and vicinity. US Geological Survey, Open-File Report 2007^1071, 1^94. Cole, J. W. (1970). Structure and eruptive history of the Tarawera Volcanic Complex. New Zealand Journal of Geology and Geophysics 13, 879^902. Davis, W. J. (1985). Geochemistry and petrology of the Rotoiti and Earthquake Flat pyroclastic deposits. M.Sc. thesis, University of Auckland, New Zealand. Ewart, A. (1968). The petrography of the central North Island rhyolitic lavas. Part 2çregional petrography including notes on associated ash-flow pumice deposits. New Zealand Journal of Geology and Geophysics 11, 478^545. Ewart, A., Hildreth, W. & Carmichael, I. S. E. (1975). Quaternary acid magma in New Zealand. Contributions to Mineralogy and Petrology 51, 1^27. Girard, G. & Stix, J. (2009). Magma recharge and crystal mush rejuvenation associated with early post-collapse Upper Basin Member rhyolites, Yellowstone caldera, Wyoming. Journal of Petrology 50, 2095^2125. Hildreth, W. (2004). Volcanological perspectives on Long Valley, Mammoth Mountain, and Mono Craters: several contiguous but discrete systems. Journal of Volcanology and Geothermal Research 136, 169^198. Howorth, R. (1975). New formations of Late Pleistocene tephras from the Okataina volcanic centre, New Zealand. New Zealand Journal of Geology and Geophysics 18, 683^712. Huppert, F. (1981). Petrology of late Pleistocene tephra deposits from the Okataina Volcanic Centre, New Zealand. M.Sc. thesis, University of Auckland, New Zealand. Ireland, T. R. & Williams, I. S. (2003). Considerations in zircon geochronology by SIMS. In: Hanchar, J. M. & Hoskin, P. W. O.(eds) Zircon. Mineralogical Society of America, Reviews in Mineralogy and Geochemistry 53, 215^241. Jurado-Chichay, Z. & Walker, G. P. L. (2000). Stratigraphy and dispersal of the Mangaone Subgroup pyroclastic deposits, Okataina Volcanic Centre, New Zealand. Journal of Volcanology and Geothermal Research 104, 319^383. Kilgour, G. N. & Smith, R. T. (2008). Stratigraphy, dynamics, and eruption impacts of the dual magma Rotorua eruptive episode, Okataina Volcanic Centre, New Zealand. New Zealand Journal of Geology and Geophysics 51, 367^378. Leonard, G. S., Cole, J. W., Nairn, I. A. & Self, S. (2002). Basalt triggering of the c. AD 1305 Kaharoa rhyolite eruption, Tarawera Volcanic Complex, New Zealand. Journal of Volcanology and Geothermal Research 115, 461^486. Lowe, D. J., Shane, P. A. R., Alloway, B. V. & Newnham, R. M. (2008). Fingerprints and age models for widespread New Zealand tephra marker beds erupted since 30,000 years ago: a framework for NZ-INTIMATE. Quaternary Science Reviews 27, 95^126. Lowenstern, J. B., Persing, H. M., Wooden, J. L., Lanphere, M. A., Donnelly-Nolan, J. & Grove, T. L. (2000). U^Th dating of single zircons from young granitoid xenoliths: new tools for understanding volcanic processes. Earth and Planetary Science Letters 183, 291^302. Lowenstern, J. B., Charlier, B. L. A., Clynne, M. A. & Wooden, J. L. (2006). Extreme U^Th disequilibrium in rift-related basalts, rhyolites and granophyric granite and the timescale of rhyolite generation, intrusion and crystallization at Alid volcanic center, Eritrea. Journal of Petrology 47, 2105^2122. Ludwig, K. R. (2003). Isoplot/Ex version 3.41, A geochronological toolkit for Microsoft Excel. Berkeley Geochronology Center Special Publication 4. Mankinen, E. A., Gromme, C. S., Dalrymple, G. B., Lanphere, M. A. & Bailey, R. A. (1986). Paleomagnetism and K^Ar ages of volcanic 1139 JOURNAL OF PETROLOGY VOLUME 51 rocks from Long Valley caldera, California. Journal of Geophysical Research 91, 633^652. McConnell, V. S., Shearer, C. K., Eichelberger, J. C., Keskinen, M. J., Layer, P. W. & Papike, J. J. (1995). Rhyolite intrusions in the intracaldera Bishop Tuff, Long Valley caldera, California. Journal of Volcanology and Geothermal Research 67, 41^60. Metz, J. M. & Mahood, G. A. (1985). Precursors to the Bishop Tuff eruption: Glass Mountain, Long Valley, California. Journal of Geophysical Research 90, 11121^11126. Miller, J. S. & Wooden, J. L. (2004). Residence, resorption and recycling of zircons in Devils Kitchen rhyolite, Coso volcanic field, California. Journal of Petrology 45, 2155^2170. Molloy, C., Shane, P. & Nairn, I. A. (2008). Pre-eruption thermal rejuvenation and stirring of a partly crystalline rhyolite pluton revealed by the Earthquake Flat Pyroclastics deposits, New Zealand. Journal of the Geological Society, London 165, 435^447. Nairn, I. A. (1972). Rotoehu Ash and the Rotoiti Breccia Formation, Taupo Volcanic Zone, New Zealand. New Zealand Journal of Geology and Geophysics 15, 251^261. Nairn, I. A. (1980). Source, age, and eruptive mechanisms of Rotorua Ash. New Zealand Journal of Geology and Geophysics 23, 193^207. Nairn, I. A. (1981). Some studies of the geology, volcanic history, and geothermal resources of the Okataina volcanic centre, Ph.D. thesis, Victoria University, Wellington, New Zealand. Nairn, I. A. (2002). Geology of the Okataina Volcanic Centre, scale 1:50,000. Institute of Geological & Nuclear Sciences geological map 25, Lower Hutt: Institute of Geological & Nuclear Sciences Limited, 1 sheet þ156 pp. Nairn, I. A. & Kohn, B. P. (1973). Relation of the Earthquake Flat Breccia to the Rotoiti Breccia, central North Island, New Zealand. New Zealand Journal of Geology and Geophysics 16, 269^279. Nairn, I. A., Self, S., Cole, J. W., Leonard, G. S. & Scutter, C. (2001). Distribution, stratigraphy, and history of proximal deposits from the c. AD 1305 Kaharoa eruptive episode at Tarawera volcano, New Zealand. New Zealand Journal of Geology and Geophysics 44, 467^484. Nairn, I. A., Shane, P. R., Cole, J. W., Leonard, G. J., Self, S. & Pearson, N. (2004). Rhyolite magma processes of the AD 1315 Kaharoa eruption episode, Tarawera volcano, New Zealand. Journal of Volcanology and Geothermal Research 131, 265^294. Nairn, I. A., Hedenquist, J. W., Villamor, P., Berryman, K. R. & Shane, P. A. (2005). The AD1315 Tarawera and Waiotapu eruptions, New Zealand: contemporaneous rhyolite and hydrothermal eruptions driven by an arrested basalt dike system? Bulletin of Volcanology 67, 186^193. Newnham, R. M., Lowe, D. J., Green, J. D., Turner, G. M., Harper, M. A., McGlone, M. S., Stout, S. L., Horie, S. & Froggatt, P. C. (2004). A discontinuous ca. 80 ka record of Late Quaternary environmental change from Lake Omapere, Northland, New Zealand. Palaeogeography, Palaeoclimatology, Palaeoecology 207, 165^198. Phillips, E. H., Goff, F., Kyle, P. R., McIntosh, W. C., Dunbar, N. W. & Gardner, J. N. (2007). The 40Ar/39Ar age constraints on the duration of resurgence at the Valles caldera, New Mexico. Journal of Geophysical Research 112, B08201, doi:10.1029/2006JB004511. Reid, M. R. (2008). How long does it take to supersize an eruption? Elements 4, 23^28. Reid, M. R. & Coath, C. D. (2000). In situ U^Pb ages of zircons from the Bishop Tuff: no evidence for long crystal residence times. Geology 28, 443^446. Reid, M. R., Coath, C. D., Harrison, T. M. & McKeegan, K. D. (1997). Prolonged residence times for the youngest rhyolites NUMBER 5 MAY 2010 associated with Long Valley caldera: 230Th^238U ion microprobe dating of young zircons. Earth and Planetary Science Letters 150, 27^39. Rowland, J. V., Wilson, C. J. N. & Gravley, D. M. (2010). Spatial and temporal variations in magma-assisted rifting, Taupo Volcanic Zone, New Zealand. Journal of Volcanology and Geothermal Research 190, 89^108. Schmitz, M. D. & Smith, I. E. M. (2004). The petrology of the Rotoiti eruption sequence, Taupo Volcanic Zone: an example of fractionation and mixing in a rhyolitic system. Journal of Petrology 45, 2045^2066. Shane, P., Smith,V. C. & Nairn, I. A. (2005a). High temperature rhyodacites of the 36 ka Hauparu pyroclastic eruption, Okataina Volcanic Centre, New Zealand: Change in a silicic magmatic system following caldera collapse. Journal of Volcanology and Geothermal Research 147, 357^376. Shane, P., Nairn, I. A. & Smith,V. C. (2005b). Magma mingling in the 50 ka Rotoiti eruption from Okataina Volcanic Centre: implications for geochemical diversity and chronology of large volume rhyolites. Journal of Volcanology and Geothermal Research 139, 295^313. Shane, P., Martin, S. B., Smith, V. C., Beggs, K. F., Darragh, M. B., Cole, J. W. & Nairn, I. A. (2007). Multiple rhyolite magmas and basalt injection in the 17·7 ka Rerewhakaaitu eruption episode from Tarawera volcanic complex, New Zealand. Journal of Volcanology and Geothermal Research 164, 1^26. Shane, P., Nairn, I. A., Smith, V. C., Darragh, M., Beggs, K. F. & Cole, J. W. (2008a). Silicic recharge of multiple rhyolite magmas by basaltic intrusion during the 22· 6 ka Okareka eruption episode, New Zealand. Lithos 103, 527^549. Shane, P., Nairn, I. A., Martin, S. B. & Smith, V. C. (2008b). Compositional heterogeneity in tephra deposits resulting from the eruption of multiple magma bodies: implications for tephrochronology. Quaternary International 178, 44^53. Simon, J. I. & Reid, M. R. (2005). The pace of rhyolite differentiation and storage in an ‘archetypical’ silicic magma system, Long Valley, California. Earth and Planetary Science Letters 235, 123^140. Simon, J. I., Reid, M. R. & Young, E. D. (2007). Lead isotopes by LA-MC-ICPMS: tracking the emergence of mantle signatures in an evolving silicic magma system. Geochimica et Cosmochimica Acta 71, 2014^2035. Simon, J. I., Renne, P. R. & Mundil, R. (2008). Implications of preeruptive magmatic histories of zircons for U^Pb geochronology of silicic extrusions. Earth and Planetary Science Letters 266, 182^194. Simon, J. I.,Vazquez, J. A., Renne, P. R., Schmitt, A. K., Bacon, C. R. & Reid, M. R. (2009). Accessory mineral U^Th^Pb ages and 40 Ar/39Ar eruption chronology, and their bearing on rhyolitic magma evolution in the Pleistocene Coso volcanic field, California. Contributions to Mineralogy and Petrology 158, 421^446. Smith, R. L. (1979). Ash-flow magmatism. In: Chapin, C. E. & Elston, W. E. (eds). Ash-flow Tuffs. Geological Society of America, Special Papers 180, 5^27. Smith, V. C., Shane, P. A. R. & Smith, I. E. M. (2002). Tephrostratigraphy and geochemical fingerprinting of the Mangaone Subgroup tephra beds, Okataina Volcanic Centre, New Zealand. New Zealand Journal of Geology and Geophysics 45, 207^219. Smith, V. C., Shane, P. & Nairn, I. A. (2004). Reactivation of a rhyolitic magma body by new rhyolitic intrusion before the 15· 8 ka Rotorua eruptive episode: implications for magma storage in the Okataina Volcanic Centre, New Zealand. Journal of the Geological Society, London 161, 757^772. Smith, V. C., Shane, P. & Nairn, I. A. (2005). Trends in rhyolite geochemistry, mineralogy, and magma storage during the last 50 kyr at Okataina and Taupo volcanic centres, Taupo Volcanic Zone, 1140 CHARLIER & WILSON OKATAINA VOLCANO MAGMA SYSTEMS New Zealand. Journal of Volcanology and Geothermal Research 148, 372^406. Smith, V. C., Shane, P., Nairn, I. A. & Williams, C. M. (2006). Geochemistry and magmatic properties of eruption episodes from Haroharo linear vent zone, Okataina Volcanic Centre, New Zealand during the last 10 kyr. Bulletin of Volcanology 69, 57^88. Stix, J., Goff, F., Gorton, M. P., Heiken, G. & Garcia, S. R. (1988). Restoration of compositional zonation in the Bandelier silicic magma chamber between two caldera-forming eruptions: geochemistry and origin of the Cerro Toledo Rhyolite, Jemez Mountains, New Mexico. Journal of Geophysical Research 93, 6129^6147. Streck, M. J. & Grunder, A. L. (2008). Phenocryst-poor rhyolites of bimodal tholeiitic provinces: the Rattlesnake Tuff and implications for mush extraction models. Bulletin of Volcanology 70, 385^401. Sutton, A. N., Blake, S. & Wilson, C. J. N. (1995). An outline geochemistry of rhyolite eruptives from Taupo volcanic centre, New Zealand. Journal of Volcanology and Geothermal Research 68, 153^175. Sutton, A. N., Blake, S., Wilson, C. J. N. & Charlier, B. L. A. (2000). Late Quaternary evolution of a hyperactive rhyolite magmatic system: Taupo volcanic centre, New Zealand. Journal of the Geological Society, London 157, 537^552. Trial, A. F. & Spera, F. J. (1990). Mechanisms for the generation of compositional heterogeneities in magma chambers. Geological Society of America Bulletin 102, 353^367. Vazquez, J. A. & Reid, M. R. (2002). Time scales of magma storage and differentiation of voluminous high-silica rhyolites at Yellowstone caldera, Wyoming. Contributions to Mineralogy and Petrology 144, 274^285. Vazquez, J. A., Kyriazis, S. F., Reid, M. R., Sehler, R. C. & Ramos, F. C. (2009). Thermochemical evolution of young rhyolites at Yellowstone: Evidence for a cooling but periodically replenished postcaldera magma reservoir. Journal of Volcanology and Geothermal Research 188, 189^196. Walker, G. P. L. (1979). A volcanic ash generated by explosions where ignimbrite entered the sea. Nature 281, 642^646. Watson, E. B. & Harrison, T. M. (1983). Zircon saturation revisited: temperature and composition effects in a variety of crustal magma types. Earth and Planetary Science Letters 64, 295^304. Wilson, C. J. N. (1993). Stratigraphy, chronology, styles and dynamics of late Quaternary eruptions from Taupo volcano, New Zealand. Philosophical Transactions of the Royal Society of London, Series A 343, 205^306. Wilson, C. J. N. & Charlier, B. L. A. (2009). Rapid rates of magma generation from contemporaneous magmatic systems at Taupo volcano, New Zealand: insights from zircon model-age spectra. Journal of Petrology 50, 875^907. Wilson, C. J. N., Houghton, B. F., McWilliams, M. O., Lanphere, M. A., Weaver, S. D. & Briggs, R. M. (1995). Volcanic and structural evolution of Taupo Volcanic Zone, New Zealand: a review. Journal of Volcanology and Geothermal Research 68, 1^28. Wilson, C. J. N., Blake, S., Charlier, B. L. A. & Sutton, A. N. (2006). The 26· 5 ka Oruanui eruption, Taupo volcano, New Zealand: development, characteristics and evacuation of a large rhyolitic magma body. Journal of Petrology 47, 35^69. Wilson, C. J. N., Rhoades, D. A., Lanphere, M. A., Calvert, A. T., Houghton, B. F., Weaver, S. D. & Cole, J. W. (2007). A multipleapproach radiometric age estimate for the Rotoiti and Earthquake Flat eruptions, New Zealand, with implications for the MIS 4/3 boundary. Quaternary Science Reviews 26, 1861^1870. Wilson, C. J. N., Gravley, D. M., Leonard, G. S. & Rowland, J. V. (2009). Volcanism in the central Taupo Volcanic Zone, New Zealand: tempo styles and controls. In: Thordarson, T., Self, S., Larsen, G., Rowland, S. K. & Hoskuldsson, A. (eds). Studies in Volcanology: The Legacy of George Walker. Special Publications of IAVCEI 2, 225^247. 1141
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