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@
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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
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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).
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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.
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JOURNAL OF PETROLOGY
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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.
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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
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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
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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’
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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.
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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).
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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
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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
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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
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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
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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.
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