JOURNAL OF PETROLOGY
VOLUME 53
NUMBER 10
PAGES 2139^2189
2012
doi:10.1093/petrology/egs050
The Anatomy of an Andesite Volcano: a
Time^Stratigraphic Study of Andesite
Petrogenesis and Crustal Evolution at
RuapehuVolcano, New Zealand
R. C. PRICE1*, J. A. GAMBLE2, I. E. M. SMITH3, R. MAAS4,
T. WAIGHT5, R. B. STEWART6 AND J. WOODHEAD4
1
FACULTY OF SCIENCE AND ENGINEERING, UNIVERSITY OF WAIKATO, PRIVATE BAG 3105, HAMILTON, NEW ZEALAND
2
DEPARTMENT OF GEOLOGY, NATIONAL UNIVERSITY OF IRELAND, UNIVERSITY COLLEGE CORK, CORK, IRELAND
3
SCHOOL OF ENVIRONMENT, UNIVERSITY OF AUCKLAND, AUCKLAND 1142, NEW ZEALAND
4
SCHOOL OF EARTH SCIENCES, UNIVERSITY OF MELBOURNE, PARKVILLE, VIC. 3052, AUSTRALIA
5
DEPARTMENT OF GEOGRAPHY AND GEOLOGY, UNIVERSITY OF COPENHAGEN, STER VOLDGADE 10, 1350,
COPENHAGEN, DENMARK
6
INSTITUTE OF NATURAL RESOURCES, MASSEY UNIVERSITY, PALMERSTON NORTH 4442, NEW ZEALAND
RECEIVED OCTOBER 3, 2011; ACCEPTED JUNE 27, 2012
ADVANCE ACCESS PUBLICATION AUGUST 21, 2012
Ruapehu, New Zealand’s largest active andesite volcano, is located at
the southern tip of the Taupo Volcanic Zone (TVZ), the main locus
of subduction-related volcanism in the North Island. Geophysical
data indicate that crustal thickness increases from 525 km within
theTVZ to 40 km beneath Ruapehu.The volcano is built on a basement of Mesozoic meta-greywacke, and geophysical evidence together
with xenoliths contained in lavas indicates that this is underlain by
oceanic, meta-igneous lower crust. The present-day Ruapehu edifice
has been constructed by a series of eruptive events that produced
a succession of lava flow-dominated stratigraphic units. In order
from oldest to youngest, these are the Te Herenga (250^180 ka),
Wahianoa (160^115 ka), Mangawhero (55^45 ka and 20^30 ka),
and Whakapapa (15^2 ka) Formations. The dominant rock types
are plagioclase- and pyroxene-phyric basaltic andesite and andesite.
Dacite also occurs but only one basalt flow has been identified.
There have been progressive changes in the minor and trace element
chemistry and isotopic composition of Ruapehu eruptive rocks over
time. In comparison with rocks from younger formations, Te
Herenga eruptive rocks have lower K2O abundances and a relatively
restricted range in major and trace element and Nd^Sr isotopic composition. Post-Te Herenga andesites and dacites define a Sr^Nd
isotopic array that overlaps with the field forTVZ rhyolites and basalts, but Te Herenga Formation lavas and the Ruapehu basalt have
higher 143Nd/144Nd ratios.The isotopic, and major and trace element
composition of Te Herenga andesite can be replicated by models involving mixing of an intra-oceanic andesite with a crustal component
derived from a meta-igneous composition. Post-Te Herenga andesites
show considerable variation in major and trace element and Sr and
Nd isotopic compositions (87Sr/86Sr ranges from 0·7049 to 0·7060
and 143Nd/144Nd from 0·51264 to 0·51282). The range of compositions can be modeled by assimilation^fractional crystallization
(AFC) involving meta-greywacke as the assimilant, closed-system
fractionation, or by mixing of intra-oceanic andesite or basalt and a
meta-greywacke crustal composition. Plagioclase and pyroxene compositions vary over wide ranges within single rocks and few of these
have compositions consistent with equilibration with a melt having
the composition of either the host-rock or groundmass. The 87Sr/86Sr
compositions of plagioclase also vary significantly within single
whole-rock samples. Glass inclusions and groundmasses of andesitic
rocks all have dacitic or rhyolitic major and trace element compositions. The application of various mineral geothermometers and geobarometers indicates pre-eruption temperatures between 950 and
*Corresponding author. Telephone: þ61353343811.
E-mail: [email protected]
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JOURNAL OF PETROLOGY
VOLUME 53
11908C and pressures ranging from 1 to 0·2 GPa.These pressure estimates are consistent with those obtained from xenolith mineral assemblages and geophysical information. Plagioclase hygrometry and
the paucity of amphibole are indications that melts were relatively
dry (54 wt % H2O). Magmas represented by Ruapehu andesites
were dacitic or rhyolitic melts carrying complex crystal and lithic cargoes derived from the mantle and at least two crustal sources. They
have evolved through a complex interplay between assimilation, crystal fractionation, crustal anatexis and magma mixing. Parental
magmas were sourced in both the mantle and crust, but erupted compositions very strongly reflect modification by intracrustal processes.
Geochemical variation in systematically sampled lava flow sequences
is consistent with random tapping of a complex plumbing system in
which magma has been stored on varying time scales within a
plexus of dispersed reservoirs. Each magma batch is likely to have
had a unique history with different sized magma storages evolving
on varying time scales with a specific combination of AFC and
mixing processes.
andesite; volcano; Taupo Volcanic Zone; New Zealand;
assimilation; fractional crystallization
KEY WORDS:
I N T RO D U C T I O N
Andesitic volcanic systems are notoriously unpredictable
with respect to duration of eruptive cycles and eruptive
volumes. This presents a major challenge to geophysicists
and volcanologists charged with forecasting or predicting
eruptions and their magnitudes. The May 1980 eruption
of Mount St Helens in the Cascades of the western USA
occurred after a repose period of over 100 years. In 1995,
Soufrie're Hills Volcano on Monserrat in the Lesser
Antilles arc reactivated after centuries of quiescence and
this eruption has continued to the present day.
Klyuchevskoy, in the Kamchatka Peninsula, erupts almost
continuously. Mount Ruapehu at the southern tip of the
Taupo Volcanic Zone (TVZ) in New Zealand is fed by a
magmatic system characterized, over at least the last two
millennia, by relatively frequent (decadal) small-scale
(50·05 km3 magma batches) eruptions lasting several
months Monitoring of arc-type volcanoes requires an
understanding of the magmatic plumbing systems that
feed them. A particularly effective petrological approach,
by which the behavior of a magmatic system prior to and
during volcanic eruptions can be understood, is to examine
fine-scale mineralogical and chemical changes in erupted
materials within a detailed temporal^stratigraphic framework (e.g. Turner et al., 2008a, 2008b).
The geochemical similarity between arc andesite and
continental crust has been recognized for more than
40 years (Taylor, 1967; Rudnick & Gao, 2005) and consequently andesite petrogenesis and arc^back-arc magmatism have been linked to crustal growth (e.g. Arculus,
1999; Tatsumi & Kogiso, 2003; Keleman et al., 2005;
NUMBER 10
OCTOBER 2012
Davidson & Arculus, 2006; Kodaira et al., 2007). In this
study detailed petrological information for a large and
long-lived andesite stratovolcano is used to construct an
integrated model for the petrogenesis of andesitic magmas
and to explore the linkage between the processes of andesite genesis and continental crust formation.
Over the past three decades an extensive and prolonged
international programme of research directed primarily
at oceanic arcs has led to a general consensus that
subduction-related magmas have an ultimate origin in the
mantle owing to melting caused by migration of fluids
and/or melts from the subducting plate (e.g. Hawkesworth
et al., 1979, 1993; Arculus & Powell, 1986; Grove & Kinzler,
1986; McCulloch & Gamble, 1991; Brenan et al., 1995;
Elliott et al., 1997; Kessel et al., 2005). Considerable effort
has been devoted to unravelling the complexities of andesite^dacite magmatic systems through detailed stratigraphic studies of: (1) specific eruptive episodes or single
volcanoes (e.g. Gamble et al., 1999; Dungan et al., 2001;
Hobden et al., 2002); (2) degassing and crystallization histories (Blundy & Cashman, 2001; Rutherford & Devine,
2003; Blundy et al., 2006); (3) crystal isotope stratigraphy
(Davidson et al., 2007); (4) phenocryst stratigraphy
(Humphreys et al., 2006); (5) the links between andesitic
and rhyolite volcanism in arcs in general (Price et al.,
2005; Reubi & Blundy, 2009). It is now recognized that
the interaction of mantle-derived magmas with the existing lower and/or middle crust is a significant, possibly the
dominant, factor in the evolution of continental,
subduction-related magmas (e.g. Price et al., 2005; Annen
et al., 2006; Ruebi & Blundy, 2009) and may also be important in magmatic processes taking place in intra-oceanic
subduction systems (e.g. Smith et al., 2010).
In this study we describe the magmatic evolution of
Mount Ruapehu (2797 m), which is the largest, currently
active volcano and highest mountain in the North Island
of New Zealand; the present-day edifice has a volume estimated at 150 km3 (Hackett & Houghton, 1989). The objectives of the work are to understand the anatomy and
unravel the complex phylogeny of andesitic magmas
erupted throughout the history of a single arc volcano and
to thereby gain insights into the role of subduction-related
magmatic processes in crustal evolution. A systematic
examination of the petrology and geochemistry of samples
taken within a well-defined temporal framework of volcano
growth and erosion (Graham & Hackett, 1987; Gamble
et al., 1999, 2003) provides the basis for a model for the generation of Ruapehu andesitic magmas and the evolution
of crust in a continental subduction setting. Ruapehu andesites are argued to represent complex, multi-sourced,
crystal-rich rhyolite or dacite melts [see Price et al. (2005)
and Reubi & Blundy (2009)] derived through an intricate
interplay between melting, mixing, mingling and differentiation processes that takes place largely in the crust
2140
PRICE et al.
ANDESITE PETROGENESIS, RUAPEHU VOLCANO
(Hildreth & Moorbath, 1988), with mantle-derived
magmas filtered through arc lithosphere (Smith et al.,
2003; Price et al., 2005; Annen et al., 2006). This subduction
factory scenario (Tatsumi & Kogiso, 2003; Tatsumi, 2005)
specifies an evolving lower crust in which mantle- and
crust-derived magmas interact, mix, mingle, homogenize,
crystallize and differentiate to form derivative melts carrying a complex cargo of crystals (phenocrysts, xenocrysts
and antecrysts), restite, and lithic fragments.
Subduction-related volcanism in
New Zealand
Volcanism in northern New Zealand is associated with
westward subduction of the Pacific Plate beneath the
Australian Plate along the Hikurangi^Kermadec Trench
system (Cole, 1979, 1986; Reyners et al., 2006; Stern et al.,
2010) (Fig. 1). The plate boundary extends from Tonga
southward into New Zealand and convergence is progressively more oblique to the south, with convergence rates
varying from in excess of 100 mm a1 near Tonga to less
than 50 mm a1 beneath North Island New Zealand
(Cole, 1979; De Mets et al., 1990). The age and composition
of the subducting plate is broadly similar along the arc
but from north to south there are geographically constrained differences. These include the following: (1) to the
north, the Louisville seamount chain has been and continues to be subducted along the Tonga arc; (2) in the
south, the Hikurangi Plateau (Mortimer & Parkinson,
1996) is at present subducting; (3) it is likely that the subducted sediment contribution to the mantle wedge increases from north to south along the arc (Gamble et al.,
1996; Wysoczanski et al., 2006, 2010; Todd et al., 2010). The
Kermadec Trench segment of the subduction boundary is
oceanic but the North Island is underlain by continental
crust and subduction along the Hikurangi Trough is beneath continental lithosphere (Cole, 1979, 1986; Gamble
et al., 1993b). The boundary between oceanic and continental crust of the Australian Plate lies to the north of New
Zealand (Fig. 1). In the oceanic segment of the subduction
system, magmatism is manifested as a chain of volcanic
seamounts and islands, whereas in the continental segment
the principal locus of magmatic activity is the TVZ.
The central TVZ, to the north of Ruapehu, is characterized by exceptionally high heat flow (700 mW m2 or
30 mW km1 of strike length; Stern, 1987; Hochstein et al.,
1993; Bibby et al., 1995; Stern et al., 2010), high rates of
regional extension that vary from around 19 mm a1 in
the north to around 7^8 mm a1 in the south (Darby &
Meertens, 1995; Beanland & Haines, 1998; Rowland &
Sibson, 2001; Rowland et al., 2010; Stern et al., 2010) and
relatively thin crust (15^20 km; Stern & Davey, 1987;
Harrison & White, 2006; Stern et al., 2010). The TVZ is
one of the most active and productive silicic magmatic systems on Earth (Houghton et al., 1995; Wilson et al., 1995).
The northern sector of the onshore TVZ is dominated by
a series of rhyolitic caldera centres associated with voluminous rhyolitic ignimbrite eruptions extending back to
2·0 Ma (Wilson et al.,1995).Volumetrically minor basaltic
cones pepper this region (Cole, 1990). The southern sector
comprises the andesite volcanoes of the Tongariro Volcanic
Centre (Cole, 1978, 1986) including Ruapehu and the
nearby volcanoes of Tongariro and Ngauruhoe. Ruapehu
lies within a graben (Fig. 1) that marks the southern
margin of the SW propagating tip of the TVZ (Stern
et al., 2006; Villamor & Berryman, 2006a, 2006b; Reyners,
2010).
F I E L D R E L AT I O N S ,
G E O C H RO N O L O G Y A N D
P E T RO G R A P H Y O F RUA P E H U
ANDESITES
Stratigraphic framework, geochronology
and magma flux
Ruapehu volcano has been built in a series of intense constructional events separated by periods of erosion, sector
collapse and low-level volcanic activity. Four major lava
flow formations have been identified (Hackett, 1985;
Hackett & Houghton, 1989). From oldest to youngest
they are the Te Herenga, Wahianoa, Mangawhero and
Whakapapa Formations (Fig. 1). Tanaka et al. (1997) and
Gamble et al. (2003) established a chronology for this stratigraphic framework dating back to 250 ka and thereby
demonstrated that growth of the volcano occurred during
relatively discrete periods of intense activity at around
250^180, 160^115, 55^45, 30^20 and 15^2 ka (Table 1).
The post-Holocene (515 ka) record of activity is best preserved in tephra sequences of the ring plain of lahar,
debris-avalanche and tephra deposits that surrounds the
Ruapehu edifice (Donoghue et al., 1995). Samples from the
most recent events (1945^1996) were described by Gamble
et al. (1999) and Nakagawa et al. (1999) but samples from
documented early 20th and 19th century events are not
available.
Flows making up the Whakapapa Formation are too
young to be dated by the Ar^Ar method but the details of
the Quaternary eruptive history of Ruapehu are partially
preserved in the tephras making up the ring plain that surrounds the mountain (Donoghue et al., 1995). Many of
these tephras have been dated by the 14C method or from
their stratigraphic positions relative to 14C-dated rhyolitic
tephras from the TVZ caldera volcanoes to the north.
Where andesitic and rhyolitic tephras occur within
Whakapapa lava flow sequences, they provide a stratigraphic framework within which the age of flows can be
estimated (Table 1). There are numerous cases where
tephra sequences are preserved on top of lava flows and
when these can be correlated with dated units on the ring
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VOLUME 53
NUMBER 10
OCTOBER 2012
Fig. 1. (a) The location of Ruapehu volcano at the southern end of the Taupo Volcanic Zone (TVZ) in New Zealand’s North Island. C/O is the
approximate line of transition from continental (to the south) to oceanic lithosphere. Inset shows location of North Island New Zealand relative
to the subduction boundary between the Australian and Pacific Plates. 1, 2, 3 and 4 are basalt or basaltic andesite eruption centres in the southern TVZ, discussed by Graham & Hackett (1987). 1, Ohakune Craters; 2, Hauhungatahi; 3, Tongariro basaltic centres, including Pukekaikiore,
Pukeonake and Red Crater; 4, Waimarino. (b) Map of the volcanic geology of Ruapehu volcano. Map includes new data as well as information
from Hackett (1985), Schneider (1995), Valente (1995), Beyer (1996), Chapman (1996), Nairn et al. (1998) and Waight et al. (1999). PR, Pinnacle
Ridge; Wh, Whakapapa skifield; Tr, Turoa skifield; Tk, Tukino skifield. (c) Schematic cross-section [A^B in (b)] showing the location of
Ruapehu within the Mt. Ruapehu graben [incorporates interpretations from Villamor & Berryman (2006a) and Cassidy et al. (2009)].
plain they establish a basis on which the age of the lavas
can be constrained.
The Whakapapa Formation overlies and is in direct contact with Te Herenga and Mangawhero Formation rocks
across the northern, western and southern flanks of the
volcano and there is clear field evidence that it was
emplaced after the last glacial maximum; for example, in
exposures in the headwaters of the Whakapapaiti stream
on western Ruapehu, glacially scoured and polished outcrops of the Te Herenga Formation are directly overlain
by unglaciated Whakapapa Formation flows showing original rubbly flow tops. This provides a first-order limit on
the age of the Whakapapa Formation.
The flow sequences that have been mapped within these
exposures (Table 1 and Fig. 2) appear to be related to a
linear array of young vents or vent systems (Fig. 2) that
Nairn et al. (1998) argued were the source for the Pahoka^
Mangamate tephra sequence. Using 14C dating, Nairn
et al. (1998) concluded that these tephras were emplaced at
10 ka BP during an intensive period (200^400 years) of
pyroclastic eruptive activity, which was presumably associated with and followed by lava flow emplacement. From
north to south, the principal sub-units of the Whakapapa
Formation discussed are: the Saddle Cone, Delta Corner,
Older Whakapapa, Sunset Ridge (East and West), and
Rangataua flows (Table 1 and Fig. 2). The Rangataua
flows have been emplaced during three separate events
and the unit is therefore subdivided into proximal, medial
and distal flows.
Estimates of relative age for the different Whakapapa
flow units, as determined from tephra stratigraphy, are
summarized in Table 1. The most extensive lava fields
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PRICE et al.
ANDESITE PETROGENESIS, RUAPEHU VOLCANO
Table 1: Flow stratigraphy for Ruapehu volcano (after
Hackett, 1985), and principal flow units of the Whakapapa
Formation, Ruapehu volcano (see Fig. 2)
Formation
Member
Age
Whakapapa
1945–1996
566 a
Gamble et al. (1999)
Delta Corner
2·5–3·3 ka
From tephra sequences
Older Whakapapa
3·3–10 ka
From tephra sequences
overlying flows
overlying flows
Sunset Ridge
55 ka
From tephra sequences
overlying flows
Rangataua
10–12 ka
From tephra sequences
overlying flows
Saddle Cone
410 ka
From tephra sequences
overlying flows
Mangawhero
Younger
20–30 ka
Older
Gamble et al. (2003)
45–55 ka
Gamble et al. (2003)
Wahianoa
115–160 ka
Gamble et al. (2003)
Te Herenga
180–250 ka
Gamble et al. (2003)
Flow unit,
Age constraints Source
Whakapapa Formation
Saddle Cone
510 ka
Delta Corner
2·5–3·3 ka
Older Whakapapa
3·3–9·5 ka
1
1
1, 2, 3
Sunset West
45 ka
4
Sunset East
45 ka
4
Rangataua Proximal
59·7 ka
1
Rangataua Medial
59·7 ka
1, 5
Rangataua Distal
411·9 ka
1, 5
method at 9540 100 a (Topping, 1973). Mapping of the
Whakapapa skifield has defined a complex series of lava
flows but for the purposes of this study these are simplified
and subdivided into two units. Older Whakapapa lava
flows are overlain by the Delta Corner Flows.
Over the past 2 kyr the eruptive history of Ruapehu has
been characterized by low-volume (50·05 km3) but frequent (25^30 years) phreatomagmatic eruptions occurring
through a crater lake at the summit (Donoghue et al.,
1995, 1997). Eruptions between and including 1945 and
1995^1996 are the most recent manifestation of this continuing low-level activity. The petrography and geochemistry of these magmatic products has been discussed by
Gamble et al. (1999). There is geophysical evidence that
recent eruptive activity is associated with the migration of
small magma batches at shallow levels in the subvolcanic
plumbing system (Gerst & Savage, 2004).
Assuming a total volume of 300 km3 for all material
(flows of the present-day edifice as well as tephra and reworked deposits) erupted during the lifespan of the volcano Gamble et al. (2003) calculated an average magma
flux for Ruapehu of 1·2 km3 ka1. There is considerable uncertainty associated with this overall average estimate.
For example, if a volume of 150 km3 [estimated for the
flows making up the present-day edifice by Hackett &
Houghton (1989)] is used then the average flux reduces to
0·6 km3 ka1. The magma fluxes calculated by Gamble
et al. (2003) for the single flow formations are, however,
consistent with the average range, with, in each case,
the magma production rate varying between 0·9 and
1·0 km3 ka1 and the highest rate of magma flux being calculated for the Wahianoa Formation. It is clear that the
flow formations represent a number of discrete pulses of
effusive activity of similar scale, with flow sequences in
each being emplaced over time intervals of the order of
10^40 kyr.
Basement geology and xenoliths
Data sources: 1, R. B. Stewart & R. C. Price (unpublished
data for tephra sequences); 2, Topping (1973); 3, Palmer
& Neall (1989); 4, Price et al. (2000); 5, Donoghue et al.
(1995).
assigned to the Whakapapa Formation are found on the
Whakapapa skifield (Fig. 2). They were erupted from the
summit region and partially fill an extensive amphitheatre
on the northwestern flank of the mountain between
Pinnacle Ridge and the Whakapapaiti catchment. Their
emplacement could have followed a sector collapse event
that formed this amphitheatre and deposited an extensive
debris avalanche and lahar deposit on the northwestern
ring plainçthe Murimotu Formation of Palmer & Neall
(1989), which has been dated by the radiocarbon
The geology of the central North Island of New Zealand
is dominated by a NNE^SSW-trending axial range
of low-grade, Mesozoic (Jurassic to Cretaceous) metagreywacke and meta-argillite rocks with rare occurrences
of meta-basite. These are cut by a series of transcurrent
faults of the North Island Shear Belt. Based on geochronology, zircon inheritance, petrology and geochemistry
(Roser & Korsch, 1999; Adams et al., 2002, 2007) two distinct basement terranes have been identified in the North
Island. The Torlesse terrane, which is a Cretaceous greywacke^argillite sequence of felsic composition derived
from mature quartzo-feldspathic (granitoid) basement,
dominates the eastern ranges. In contrast, the Jurassic
Waipapa terrane of the western ranges was derived from
more mafic (volcanic arc) basement. The suture between
the two terranes is located to the west of Ruapehu
(Mortimer et al., 1997; Adams et al., 2007) and may well
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VOLUME 53
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OCTOBER 2012
Fig. 2. Map showing the distribution of principal flow groups of the Whakapapa Formation. Vents associated with the Pahoka^Mangamate
event are from Nairn et al. (1998). SC, Saddle Cone; WS, Whakapapa skifield; TuS, Turoa skifield; TkS, Tukino skifield.
have played a significant role in the spatial development of
volcanism through the central TVZ. It is likely that these
basement rocks rest on oceanic lithosphere of
the pre-Gondwanan breakup (see below). New estimates
of average compositions for the Torlesse and Waipapa terranes are provided as Supplementary Data (Electronic
Appendix A1, available for downloading at http://www
.petrology.oxfordjournals.org).
Unconformably overlying the Mesozoic basement are
sediments of the Late Cenozoic Wanganui Basin
(Mortimer et al., 1997), which include sands, silts, limestones and shell beds, with interbedded tuffs. In the vicinity of the Tongariro Volcanic Centre they form a thin
(a few tens to hundreds of metres; Fig. 1c) veneer separating the underlying Mesozoic basement from the younger
late Quaternary volcanic sequences (Cassidy et al., 2009).
Xenoliths contained in Tongariro andesites are dominantly samples of the deeper crust beneath the exposed
basement. They have been described in detail by Graham
(1987) and Graham et al. (1990), with additional information provided by Price et al. (2005, 2010). Summary compositional data are provided in the Supplementary Data
(Electronic Appendix A1). Most xenoliths have mineral assemblages and chemical compositions that are consistent
with a restitic origin; they have had melt extracted from
them and the mineral assemblage is largely refractory.
Small (55 cm) meta-sedimentary xenoliths, which are
likely to be high-grade metamorphic equivalents of the
Torlesse and Waipapa basement terranes, occur in many
Ruapehu andesites. The most abundant meta-sedimentary
xenolith type is small (2^50 mm) and fine grained
(52 mm) with a mineral assemblage of plagioclase (andesine^labradorite), Mg-orthopyroxene and magnetite.
Interstitial glass is common. Meta-igneous xenoliths are
widely distributed in Ruapehu andesites. The usual type is
a fine-grained (52 mm) granulite consisting of granoblastic plagioclase, orthopyroxene and ilmenite. Olivine and
clinopyroxene are common, and there are also samples
containing quartz, biotite, apatite, titanomagnetite, spinel
(Cr-spinel or pleonaste), or sulphide. Interstitial brown
glass is common and in some samples extensive (1^2 mm)
glass patches occur. Plagioclase in all Ruapehu xenoliths
(meta-sedimentary and meta-igneous) is largely unzoned
and plagioclase at the margins of xenoliths and in microxenoliths is sieve-textured with both a compositional and a
textural similarity to sieve-textured phenocrysts in the
host andesites (Price et al., 2005). Sr isotopic data indicate
that meta-igneous xenoliths could represent refractory
samples of an altered oceanic crustal component derived
from within the deep basement and underlying the
2144
PRICE et al.
ANDESITE PETROGENESIS, RUAPEHU VOLCANO
Mesozoic meta-greywackes of the Torlesse and Waipapa
terranes (Graham et al., 1990; Price et al., 2005).
Meta-sedimentary xenoliths are considered to represent
refractory samples of the latter.
Petrography
The original petrographic descriptions of Ruapehu
andesites were by Clark (1960) and Cole (1978); Graham
& Hackett (1987) first described the petrology in the
context of the volcanic stratigraphy, petrography and geochemistry. Their study included basalt and basaltic andesite samples from Hauhungatahi, Ohakune, Pukekaikiore,
Pukeonake, Red Crater and Waimarino (Fig. 1); these
eruptive centres were implied to be vents ‘related’ to or
‘parasitic’ on Ruapehu volcano. However, all are some distance from the Ruapehu edifice and at least one of them
was active well before the earliest eruptions occurred at
Ruapehu. Hauhungatahi, the closest of these vents, lies
12 km from Ruapehu’s present-day central crater (Fig. 1),
outside the Ruapehu ring plain, and it has an eruptive age
estimated at over 900 ka (Cameron et al., 2010).
Pukekaikiore, Pukeonake and Red Crater are generally
considered to be part of the Tongariro complex of
volcanoes to the north of Ruapehu, and the Waimarino
locality is almost 60 km to the north within the central
TVZ (Fig. 1). With the possible exception of Ohakune
Craters none of these vents can be considered to
have any direct connection with the Ruapehu magmatic
system.
Graham & Hackett (1987) identified six petrographic
groups (Types 1^6), with each group having distinctive
petrographic and, in some cases geochemical features.
Type 5 lavas are, however, not found on Ruapehu;
Graham & Hackett (1987) recognized them only among
eruptive rocks from the Hauhungatahi, Ohakune,
Pukekaikiore and Waimarino vents. According to the
Graham & Hackett (1987) classification, Type 1 lavas are
plagioclase^pyroxene phyric (Fig. 3), Type 2 are plagioclase phyric (Fig. 3), Type 3 and 4 andesites and dacites
are pyroxene-phyric, and Type 6 are olivine- and
pyroxene-phyric with features indicative of a hybrid
origin (namely, stronger than usual disequilibrium textures
and common reaction coronas surrounding phases such as
orthopyroxene and olivine).
Modal data for representative Ruapehu lavas are presented in Table 2. The vast majority (485%) are strongly
porphyritic Type 1 andesites with phenocryst abundances
averaging 35^55% (Fig. 3). Aphyric lavas (510% phenocrysts) are very rare. Phenocrysts range in size up to 1^
2 mm in maximum dimension. Magnetite is a ubiquitous
minor phase in Ruapehu lavas with modal abundances
ranging from 51% to 6% (mean 1%). Amphibole is rare
(Fig. 3g and h); it has been observed in fewer than a
dozen samples covering three of the four stratigraphic formations. In all formations except the Te Herenga, brown
glass is a common interstitial constituent and melt inclusions are abundant in plagioclase and pyroxene phenocrysts. The most common groundmass is a felted
microcrystalline aggregate of plagioclase, clinopyroxene
and magnetite, with or without brown glass (Fig. 3e).
Most lavas are weakly vesicular with vesicle abundance
averaging around 5% and ranging from 51% to, in rare
cases, 15^17%. Unlike flows of the younger formations Te
Herenga lavas have crystalline groundmasses (Fig. 3f). It
is probable that this represents glass that has devitrified
and recrystallized.
Type 2 andesites occur only within the Wahianoa
Formation on eastern Ruapehu (Unit C, Fig. 4). These
rocks contain 24^43% (mean 33%) plagioclase phenocrysts and relatively low modal abundances of pyroxene
(1^5%; mean 2%) and magnetite (51%). In contrast,
other Wahianoa flows have mean plagioclase and pyroxene
modal abundances of 23% and 12%, respectively.
The only basalt identified on Ruapehu (R04/04; 14 855
of Graham & Hackett, 1987) occurs in flows of the
Mangawhero Formation exposed on the northern slopes
of the volcano south of Saddle Cone (Fig. 1). It is a moderately porphyritic rock (30% phenocrysts) with plagioclase,
pyroxene and olivine phenocrysts up to 1mm across that
are clearly distinguishable in hand specimen. Olivine
makes up 4% of the mode. Plagioclase and clinopyroxene
(19% and 5% modal abundance respectively) are the
major phenocryst phases and orthopyroxene and magnetite are each 1% of the mode.
Dacites, which according to the Graham & Hackett
(1987) classification are Type 3 lavas, are relatively rare
among Ruapehu eruptive rocks. They occur among historical eruptive rocks and are also found in the Mangawhero
Formation on western Ruapehu. A typical dacite is porphyritic (24^48% phenocrysts) with approximately equal proportions of pyroxene and plagioclase phenocrysts (7^12%
and 8^18% respectively). Phenocryst phases in Ruapehu
andesites generally show complex zoning and evidence for
reaction and resorption. This is most evident in plagioclase, which commonly shows complex oscillatory zoning;
however, the plagioclase population in single samples can
include crystals showing reverse or normal zoning or no
compositional zoning at all. Sieve-textured plagioclase
crystals are ubiquitous (Fig. 3) and in many samples crystal
aggregates of plagioclase and pyroxene are common.
There does not appear to be a consistent or specific order
of crystallization for pyroxene relative to plagioclase;
plagioclase crystals can be found with pyroxene inclusions
and vice versa.
Xenolithic crustal fragments (see above) occur in most
Ruapehu lavas and are more common in the younger
lavas. They range in size from rare examples up to 10 cm
in maximum dimension down to crystal aggregates less
than 1mm across that can make up 1^2% of a thin section
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Fig. 3. Photomicrographs of Ruapehu andesites. (a, b) Type 1 plagioclase^pyroxene andesite (T6-15) from the Whakapapa Formation. (a) is
view under crossed polars; (b) is view in plane-polarized light. (c, d) Type 2 plagioclase andesite (W9/50) from the Wahianoa Formation. (c)
is view under crossed polars; (d) is view in plane-polarized light. (e, f) Very fine cryptocrystalline groundmass in Whakapapa andesite (R97/6)
(e), compared with more coarsely crystalline groundmass in a Te Herenga andesite (T6-7) (f). Both views in plane-polarized light. (g, h)
Partially resorbed amphibole in Type 1 plagioclase^pyroxene andesite (R97/56) from the Whakapapa Formation. (g) is view under crossed
polars and (h) is view in plane-polarized light. In all cases the scale bar represents 1mm. Pl, plagioclase; Cpx, clinopyroxene; Opx, orthopyroxene; Am, amphibole.
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Table 2: Representative modal phenocryst and groundmass data for Ruapehu volcano
Sample
Formation
Type
Lithology
SiO2 (wt %)
Phenocrysts
Plag
Cpx
Opx
Mt
Ol
Amph
vesicles
Gmass
T6-7
TH
1
andesite
56·32
39·7
3·7
7·5
2·5
1·5
45·1
T6-15
TH
1
andesite
56·12
35·1
4·8
8·7
3·2
6·0
42·2
T6-83
TH
1
bas. andesite
54·87
27·8
6·7
5·1
3·8
6·3
50·3
R96/18
WA
1
bas. andesite
54·95
28·0
7·7
7·3
1·0
R95/28
WA
2
andesite
57·84
30·7
1·2
0·3
R96/6
WA
1
andesite
59·14
35·2
3·3
6·2
2·5
R96/7
WA
1
andesite
60·48
24·3
3·3
6·0
1·7
R04/04
MA
1
basalt
52·79
19·4
5·1
1·1
0·8
R97/6
MA
1
andesite
58·75
29·3
5·7
4·7
2·7
R97/10a
MA
1
andesite
62·56
20·0
1·7
5·0
0·7
T5/28C
MA
4
andesite
57·25
39·7
5·3
2·9
1·9
T5-62
MA
1
andesite
59·88
13·3
2·6
8·7
T5-11
MA
1
dacite
64·47
16·9
6·9
4·4
T5-16
MA
3
dacite
63·90
8·3
8·7
R97/19
WH
1
andesite
59·08
30·7
R96/26
WH
1
andesite
58·25
T6-36
WH
1
andesite
57·84
R97/82
WH
1
andesite
R97/4
WH
1
R97/23
WH
R95/15
56·0
67·8
52·8
64·7
4·3
69·3
57·6
72·7
1·2
49·0
1·8
73·6
0·3
0·2
71·3
2·6
0·6
0·5
6·7
7·7
0·3
26·0
7·0
2·7
1·7
29·6
7·7
3·3
3·3
58·81
29·1
9·0
6·8
0·8
54·3
andesite
57·64
20·3
6·7
5·3
1·0
66·7
1
andesite
61·71
9·7
1·7
1·7
0·0
87·0
WH
1
andesite
58·95
27·7
6·0
3·0
1·3
62·0
R95/10
WH
1
andesite
59·60
35·7
8·0
1·7
0·0
54·7
R95/9
WH
1
andesite
59·65
30·7
6·7
3·7
1·0
58·0
79·3
54·7
62·7
0·6
55·5
TH, Te Herenga; WA, Wahianoa; MA, Mangawhero; WH, Whakapapa. ‘Types’ are from Graham & Hackett (1987). bas.
andesite is basaltic andesite. Plag, plagioclase; Cpx, clinopyroxene; Opx, orthopyroxene; Mt, magnetite; Ol, olivine;
Amph, amphibole; Gmass, groundmass. T5 sample data are from Valente (1995). T6 data are from Schneider (1995).
(Hackett, 1985; Graham, 1987; Graham & Hackett, 1987;
Graham et al., 1990; Price et al., 2005).
A N A LY T I C A L M E T H O D S
All whole-rock samples were crushed using a tungsten carbide ring mill. Contamination of trace elements during
the crushing process is restricted to W and Co; Ta and particularly Nb contamination is negligible (Roser et al.,
2003). For most samples abundances of major and minor
elements and selected trace elements were determined at
La Trobe University (Melbourne, Australia) by X-ray
fluorescence (XRF) analysis. Major and minor elements
(Si, Ti, Al, Fe, Mn, Mg, Ca, Na, K, P, and S) were determined as oxide components using methods similar to
those described by Norrish & Hutton (1969). In general,
precision for each major or minor element is better than
1% (1s) of the reported value. FeO abundances were
measured by direct titration using a standardized CeSO4
solution and H2O and CO2 by a gravimetric method.
Trace elements determined by XRF on pressed powder
pellets, using methods similar to those described by
Norrish & Chappell (1977), included Ba, Sr, Rb, Zr, Nb, Y,
Sc, V, Cr, Ni, Cu, Zn, and Ga. For these elements precision
is generally better than 1% for Sr and Zr, 1^3% for V, Cr,
Zn and Y, 3^5% for Ba, and 5^10% for Rb and Nb (1s).
Detection limits are 51ppm for Rb, Sr, Y, Zr and Nb,
1^2 ppm for Sc, V, Cr, Ni, Cu and Zn, and 5^10 ppm for
Ba. Accuracy was monitored by repeat analyses of
well-documented standard rocks.
The abundance levels of Pb, Th, U, and Nb are close
to XRF detection limits so these elements, along with Cs,
Hf and the rare earth elements (REE), were measured
at the Victorian Institute of Earth and Planetary
Sciences (VIEPS) Trace Element Laboratory at Monash
University (Melbourne, Australia) on selected samples
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Fig. 4. (a) Map of the flow units making up the Wahianoa Formation exposed between the Whangaehu and Wahianoa Rivers on eastern
Ruapehu above the Rangipo Desert. The rectangle shows the location of the section in (b) and (c). (b) Panoramic view from the north of the
Whangaehu Gorge section through the Wahianoa Formation. (c) Interpretation of the section showing the stratigraphic flow units within the
formation and sample locations (open circles).
using inductively coupled plasma mass spectrometry
(ICP-MS). Methods have been described in detail by
Price et al. (1999). Precision for elements analysed by
ICP-MS is typically better than 5%, with accuracy, based
on replicate analysis of BHVO-1, being for most elements
better than 5% at the 95% confidence level. Additional
trace element data were obtained for some samples by
ICP-MS at the University of Melbourne (Melbourne,
Australia) using methods adapted from Eggins et al. (1997)
and Kamber et al. (2005). For these samples 100 mg aliquots
of sample powder were digested in HF^HNO3 on a hotplate over several days. After re-dissolving the samples in
HNO3, a multi-isotope spike was added and equilibrated
with the sample solution. ICP-MS analysis was carried
out on a Varian quadrupole spectrometer using W-2 as a
calibration standard. Regular analyses of standard rocks
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PRICE et al.
ANDESITE PETROGENESIS, RUAPEHU VOLCANO
(BCR-2 and BHVO-2; see the Supplementary Data) indicate precisions of 1% (RSD) for the rare earth elements,
Sc, Nb, Hf, Pb, U and Th.
For most samples, Sr, Nd, and Pb isotopic data were obtained at La Trobe University by thermal ionization mass
spectrometry (TIMS) using a Finnigan-MAT 262 system
and methods described in detail by Price et al. (1999).
Chips were used for Pb whereas Sr^Nd isotope analyses
were carried out on rock powders. All samples were
acid-leached (6 M HCl, 1h, 1008C) and rinsed, followed
by digestion on a hotplate. Pb was extracted on small
(0·1ml) anion resin columns; total Pb blanks (50·1ng)
were negligible. Sr and Nd were extracted using standard
cation exchange and HDEHP chromatography. Mass bias
in Sr and Nd runs was corrected by normalizing to
86
Sr/88Sr ¼ 0·1194 and 146Nd/144Nd ¼ 0·7219, and corrected
data have typical in-run precisions (2SE) of 0·000020
and 0·000012, respectively. Minor instrumental bias
was eliminated by adjusting data to SRM987 ¼ 0·710230
and La Jolla Nd ¼ 0·511860. Results for USGS basalt standards BCR-1 and BHVO-1 averaged 0·70500 4 (n ¼ 6,
2SD) and 0·512634 18 (n ¼ 7, 2SD), and 0·70348 4
(n ¼13, 2SD) and 0·512989 13 (n ¼ 5, 2SD), respectively.
This indicates external precisions (2SD) of 0·000040 for
87
Sr/86Sr and 0·000020 for 143Nd/144Nd. Present-day
CHUR is 0·512638. Mass bias for Pb isotopes was corrected
using standard bracketing with SRM981, which provided
a fractionation factor of 0·109% a.m.u.1 for SRM981.
External precisions (2SD) for 78 runs of this standard
are 0·097% for 206Pb/204Pb, 0·130% for 207Pb/204Pb,
and 0·175% for 208Pb/204Pb.
For a subset of the samples, Sr, Nd and Pb isotope data
were obtained by multi-collector (MC) ICP-MS at the
University of Melbourne. Leaching, dissolution and Pb extraction protocols were identical to those described above
but Sr and Nd were purified with EICHROM SR-, REand LN-resins. Total analytical blanks were 50·1ng and
negligible. Isotopic analyses were carried out by MC-ICPMS on a NU Plasma system coupled to a CETAC Aridus
desolvating nebulizer (Maas et al., 2005). Mass bias in Sr
and Nd runs was corrected by normalizing to
86
Sr/88Sr ¼ 0·1194 and 146Nd/145Nd ¼ 2·0719425 (equivalent
to 146Nd/144Nd ¼ 0·7219; Vance & Thirlwall, 2002), using
the exponential law. Corrected data have in-run precisions
(2SE) of 0·000020 and 0·000008, respectively.
Results are reported relative to SRM987 ¼ 0·710230 and
La Jolla Nd ¼ 0·511860. The following standard results
were obtained: J-Nd-1 0·512109 16, BCR-1 0·705016 46,
0·512641 18, BHVO-1 0·703478 36, 0·512998 18, and
E&A Sr carbonate 0·708005 47 (all quoted errors are external precisions, 2SD, n ¼ 2^5 in all cases). This indicates
external precisions (2SD) near 0·000040 (Sr) and
0·000020 (Nd), similar to those obtained by TIMS (see
above). Mass bias during Pb isotope analysis was corrected
using the thallium-doping technique, which produces data
accurate to 0·03% (2SD), relative to the SRM981 composition reported by Woodhead (2002).
Trace element compositions of mineral phases and glass
inclusions were obtained by laser ablation (LA)-ICP-MS
or by solution-mode ICP-MS of small powder samples extracted with a dental drill. LA-ICP-MS data for glasses
and glass inclusion and groundmasses have been presented
by Price et al. (2005) and the technique, which follows that
of Eggins et al. (1998), has been described in that paper.
Micro-sampling of plagioclase and groundmass from polished rock surfaces was done with a small electric drill
tipped with a 0·6 mm carbide dental bur. Rock surfaces
were ultrasonicated upside down in distilled water between
samples, followed by drying under a heat lamp. Drill bits
were cleaned with dilute HCl and distilled water and inspected under a microscope between samples. For each
rock, six sub-samples of plagioclase or groundmass were
collected. Sample powders (53 mg) produced during drilling were removed from the rock surface by shaking the
harvested powder onto clean weighing paper. Trace element concentrations in the drilled sample powders, and
87
Sr/86Sr in a subset of these samples, were determined by
ICP-MS and MC-ICP-MS, respectively, at the University
of Melbourne, using the procedures described above.
In situ Sr isotope analyses of plagioclase were carried out
at the University of Melbourne, using a HELEX-excimer
193 nm laser ablation system coupled to a NU Plasma
MC-ICP-MS system (Woodhead et al., 2005; Paton et al.,
2007a, 2007b). Plagioclase was ablated for 60 s using spot
sizes of 120^150 mm (5 Hz, 55 J cm2) resulting in total Sr
signals of (0·9^1·3) 1011 A from targets with 500^
600 ppm Sr. As noted in other studies (e.g. Waight et al.,
2002; Ramos et al., 2004), plagioclase is a relatively
straightforward matrix for laser ablation Sr isotope analysis, provided that Rb interference can be corrected accurately. In this study the 87Rb/85Rb ratio used in online Rb
interference corrections was optimized using several
in-house feldspar standards before application to unknowns with Rb/Sr in the range 0·02^0·06. Data reduction
was done using the on-board NU Instruments software.
External precision for 87Sr/86Sr based on plagioclase standards and scaled to comparable signal sizes suggests reproducibility for the Ruapehu plagioclase results of
0·00018 (2SD).
The major element compositions of mineral phases in
Ruapehu lavas were analysed in representative samples
from all four mapped formations and from xenolith samples. For this study, the majority of the analyses were carried out by electron probe micro-analysis (EPMA) using
a Jeol JXA-840A instrument at the University of
Auckland. A relatively small proportion (1% of a total of
over 1400 analyses) of the data is from earlier studies
(Hackett, 1985; Graham & Hackett, 1987). Analytical data
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JOURNAL OF PETROLOGY
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obtained at the University of Auckland were gathered
using a Princeton Gamma Tech Prism 2000 Si (Li) EDS
X-ray detector, a 2 mm focused beam, an accelerating voltage of 15 kV, a beam current of 800 pA and 100 s live
count time. For Si, Ti, Al, Fe, Mg, Ca and K precision,
determined by repeated analysis of standard materials, is
generally better than 2% (1SD) and for Na it is of the
order of 4%. Accuracy can be gauged from repeated analysis of standard materials. For basaltic glass NKT-1G, analyses for all major elements are generally within 3% of
recommended values (1SD).
M I N E R A L C H E M I S T RY
Although the numbers of analyses representing mineral
compositions in each formation is variable and depends
on sample availability, the range of compositions is considered to be representative for each formation. In general
the analytical strategy was to analyse phenocryst cores
and rims together with groundmass crystals for each selected sample.
Plagioclase
A compilation of all analysed plagioclase compositions is
provided as Supplementary Data (Electronic Appendix
A2). Figures 5 and 6 illustrate the range in the composition
of plagioclase in each of the four main stratigraphic units
that make up Ruapehu volcano; in Fig. 5 these are compared with plagioclase compositions in meta-igneous and
meta-sedimentary xenoliths. In Wahianoa, Mangawhero
and Whakapapa samples plagioclase phenocryst rims,
smaller phenocrysts and groundmass crystals tend to be
less calcic than crystal cores, but there is considerable overlap between rim and core compositions and wide compositional variations in each analysed crystal population
(Fig. 6). The compositional range for plagioclase from the
Te Herenga Formation is An40^80. The Wahianoa and
Whakapapa stratigraphic units include plagioclase compositions that range to both higher and lower values
(An33^87 for Wahianoa and An23^87 for Whakapapa plagioclase). Composition is crudely correlated with crystal
type, with higher anorthite contents in larger and resorbed
crystals and the lowest values among microphenocryst
and groundmass plagioclase. A relatively small number of
analyses is available for the Mangawhero Formation, but
these suggest compositional variation comparable with
that of the Wahianoa and Whakapapa Formations.
Plagioclase compositions in the meta-igneous xenoliths encompass a wide range (An48^100) with a peak comparable
with that of plagioclase in the lava units, but a marked
tail toward extremely calcic compositions. A small
number of plagioclase analyses from a meta-sedimentary
xenolith are andesine in composition (An43^48).
The wide ranges in plagioclase compositions from single
samples indicate that the crystals originated from a variety
NUMBER 10
OCTOBER 2012
of sources or crystallized under a spectrum of compositional and physical conditions. The interdependence between plagioclase composition, melt composition and
H2O content, as well as temperature and pressure, is
demonstrated in experimental studies of plagioclase^melt
equilibria (Kudo & Weill, 1970; Baker & Eggler, 1987;
Housh & Luhr, 1991; Sisson & Grove, 1993; Takagi et al.,
2005). An empirical equation to describe the relationship
for equilibrium near-liquidus plagioclase in an anhydrous
melt for pressures of 0·5 and 1·0 GPa and temperatures of
940^13408C has been derived for basaltic compositions by
Panjasawatwong et al. (1995); this has been used to calculate
equilibrium plagioclase compositions expected for the
range of groundmass and groundmass glass compositions
that have been analysed in Ruapehu andesites (five samples) and in the Ruapehu basalt. These compositions vary
from basaltic andesite through andesite to dacite but the
range of Al# and Ca# values is similar to that of the
Panjasawatwong et al. (1995) experiments.
The plagioclase compositions calculated using the equilibrium equation can be compared with those actually
observed in Ruapehu rocks (Fig. 6). For each of the six
groundmass compositions, the equilibrium plagioclase
composition predicted by the Panjasawatwong et al. (1995)
equation has been calculated assuming magma temperatures of 10008C and 11008C and pressures of 0·5 GPa and
1GPa. These are likely to represent the range of
mid-crustal conditions in which most Ruapehu andesites
have evolved (see below). The relationships observed in
Fig. 6 indicate that the plagioclase contained in the
Ruapehu basalt must have crystallized from a melt close
in composition to the host-rock, but each of the Ruapehu
andesite samples contains a much more variable suite of
plagioclase crystal compositions with, in each case, crystal
rims or groundmass crystals having compositions approaching those expected for equilibrium with the groundmass melt composition. In each andesite sample,
plagioclase crystallization appears to have occurred over
a range of physical conditions and/or crystals have equilibrated with a variety of melt compositions. Plagioclase
compositions in Ruapehu andesites extend toward the
high An values observed in the xenoliths and this is particularly the case for plagioclase in the Wahianoa and
Whakapapa Formations (Fig. 5).
Lange et al. (2009) have shown that, if temperature can
be independently estimated, plagioclase compositions can
be used to estimate the H2O contents of equilibrium melts
or, if melt H2O content is known, the plagioclase compositions can be used to calculate equilibrium temperatures.
Pressure has a relatively limited effect on this hygrometer^
thermometer (Lange et al., 2009; compare with Putirka,
2005). The hygrometer has been applied using plagioclase
rim and groundmass compositions for eight post-Te
Herenga andesites (one sample each from the Wahianoa
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PRICE et al.
ANDESITE PETROGENESIS, RUAPEHU VOLCANO
Fig. 5. Histograms illustrating the range and frequency of plagioclase compositions in the major stratigraphic units and the xenoliths suite of
RuapehuVolcano. It should be noted that the frequency values on the y-axis are actual numbers of analyses and these vary from diagram to diagram. (a) Te Herenga Formation; (b) Wahianoa Formation; (c) Mangawhero Formation; (d) Whakapapa Formation; (e) xenoliths. Analyses
from crystal cores and rims are differentiated.‘Other’ includes groundmass crystals, microphenocrysts and plagioclase crystals included in pyroxene phenocrysts.
and Mangawhero Formations and six from the
Whakapapa Formation). In each case it was assumed that
the plagioclase crystal rims had equilibrated with melt
having the composition of the groundmass glass or bulk
groundmass. Temperatures used in the calculations
are those obtained from pyroxene geothermometry
(see below). Estimates of H2O contents range from 0·7 to
3·6 wt % and average 1·8 wt %.
Pyroxene and olivine
Typically, ortho- and clinopyroxene make up 10^20% of
the phenocryst assemblage of the Ruapehu andesites.
The compositions of the pyroxenes are illustrated in Fig. 7
and the complete dataset is available as Supplementary
Data (Electronic Appendices A3 and A4). For the Te
Herenga Formation ortho- and clinopyroxene show compositional ranges of En63^75 Fs23^33 Wo2^4 and En38^47
Fs11^24 Wo35^44 respectively (where En is enstatite, Fs is
ferrosilite and Wo is wollastonite component). The
Wahianoa, Mangawhero and Whakapapa Formations
show a much wider range of orthopyroxene compositions
(En45^85 Fs12^52) and the pyroxenes of the Wahianoa
Formation show a scatter to sub-calcic clinopyroxene.
Clinopyroxene compositions for all formations are comparatively restricted, with the more magnesian compositions being analyses of larger crystals and phenocryst
cores. In comparison clinopyroxene composition in the
meta-igneous xenoliths have a much wider compositional
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Fig. 6. Variation of plagioclase composition (An %) vs SiO2 content of the host-rock for the four main stratigraphic units of Ruapehu Volcano.
Analyses from crystal cores and rims are differentiated. ‘Other’ includes groundmass crystals, microphenocrysts and plagioclase crystals
included in pyroxene phenocrysts. The curves trace the composition of plagioclase that would be in equilibrium with the groundmass and
groundmass glass compositions analysed in five Ruapehu andesites and the Ruapehu basalt (RB). The groundmass and glass compositions
range from basaltic andesite to dacite. Model equilibrium plagioclase compositions have been calculated at pressures of 0·5 and 1·0 GPa and temperatures of 1000 and 11008C respectively using the equation An ¼ [41·836ln(Ca#)] þ [33·435ln(Al#)] ^ [63970/T] ^ [2575·3(P/T)] ^ 164·1
(after Panjasawatwong et al., 1995). Ca# is 100[Ca/(Ca þ Na)] and Al# is 100[(Al/(Al þ Si)] (atom %).
range (En11^74 Fs15^41 Wo35^50). The composition of olivine
crystals observed in the Wahianoa, Mangawhero and
Whakapapa Formations and in the xenoliths is also illustrated in Fig. 7 and the data are reported in Electronic
Appendix A5. In general the forsterite (Fo) content of olivine overlaps the En range of the coexisting pyroxene.
The Mg# [100 mol. MgO/(MgO þ FeO)] of pyroxene that could exist in equilibrium with melt having the
composition of the groundmass and groundmass glass
(Fig. 8) has been calculated using the equation of Smith
et al. (2010) (following Sisson & Grove, 1993); these results
are compared with the analysed compositions in Fig. 8.
For these calculations, the whole-rock Fe2O3/FeO ratio is
assumed to have been 0·25. Fe2O3/FeO ratios were directly
measured on 350 whole-rock samples to obtain a mean
value of 0·37 and a range of 0·11^0·87. The value of 0·25 in
the original melts has been chosen on the assumption that
whole-rock samples are likely to be more oxidized than
the original magmas; oxidation during cooling and incipient alteration on the surface, along with oxidation during
rock crushing, are all likely to raise the Fe2O3/FeO ratio
above original values. The KD(xstal^liq) [(Fe/Mg)mineral/
(Fe/Mg)liquid] values used to calculate the equilibrium
Mg#s of the mafic phases are assumed to be 0·27 for
clinopyroxene and 0·31 for orthopyroxene and olivine. The
equilibrium compositions predicted from the calculations
have Mg#s that approach or overlap with the highest
Mg#s observed among the pyroxenes and olivine in each
andesite (Fig. 8), but in each case these represent only
a small proportion of all the analysed pyroxenes and
olivines. In each andesite, pyroxenes and olivines have a
spectrum of Mg#s that extend from the calculated
groundmass melt equilibration curve down to lower
values. Each of the analysed pyroxene and olivine populations appears to include crystals that have equilibrated
with a range of melt compositions. These patterns contrast
with those observed for plagioclase. In the latter case the
compositions approaching those expected to have equilibrated with melts having the composition of the groundmass or groundmass glass are less calcic, whereas only
those pyroxenes with high Mg#s plot close to the equilibrium curves in Fig. 8.
Various geothermometers have been developed using
pyroxene compositions (e.g. Lindsley, 1983; Brey &
Ko«hler, 1990; Putirka et al., 2003) and application of these
to Ruapehu pyroxenes provides an estimate of the
pre-eruption temperatures of the Ruapehu andesitic
magmas. Data used in these calculations and results are
summarized in the Supplementary Data (Electronic
Appendix A6). The thermometer of Lindsley (1983) gives
temperature ranges of 1000^11208C for a Te Herenga lava
and 950^11008C for four post-Te Herenga andesites.
Application of the Brey & Ko«hler (1990) two-pyroxene
thermometer gives similar results, with temperature estimates of 9808C for the Te Herenga andesite and a range
of 900^11308C for the four post-Te Herenga andesites.
The Putirka et al. (2003) clinopyroxene geothermometer
requires an estimate of both pyroxene and melt
2152
PRICE et al.
ANDESITE PETROGENESIS, RUAPEHU VOLCANO
Fig. 7. Pyroxene compositions in Ruapehu volcanic rocks and xenoliths plotted in the pyroxene quadrilateral (Wo^En^Fs). Also shown on lines
below relevant quadrilateral diagrams are olivine compositions. Analyses from crystal cores and rims are differentiated. ‘Other’ includes
groundmass crystals, microphenocrysts and pyroxene crystals included in other phases. (a) Te Herenga Formation; (b) Wahianoa Formation;
(c) Mangawhero Formation; (d) Whakapapa Formation; (e) xenoliths.
compositions. Temperatures obtained by this method,
using groundmass or groundmass glass and pyroxene rim
compositions from seven post-Te Herenga andesites, range
from 1085 to 13458C. Temperature estimates have also
been obtained for pyroxenes from meta-igneous xenoliths.
For five meta-igneous xenoliths, temperature ranges estimated by the various methods are 750^11008C (Lindsley,
1983), 900^9508C (Brey & Ko«hler, 1990), and 960^9908C
(Putirka et al., 2003).
Estimation of pressure is more problematic because it is
likely that the mineral compositions have equilibrated
over a range of pressures (see below). Using the Putirka
et al. (2003) clinopyroxene geobarometer, pressure estimates obtained for equilibration of groundmass or phenocryst clinopyroxene with melts having groundmass or
groundmass glass compositions range from 0·2 to 1GPa
for seven post-Te Herenga andesites and 2 GPa for the
Ruapehu basalt.
Amphibole
Amphibole has been observed in three Wahianoa, five
Mangawhero and two Whakapapa samples and analyses
are available for one sample from the Mangawhero and
one from the Whakapapa Formation (Electronic
Appendix A7). With one exception, all analysed amphiboles are pargasites (Leake et al., 1997). The exception is a
sodic^calcic amphibole composition obtained for an inclusion within a plagioclase crystal in a Mangawhero andesite. In silicic rocks, Al in hornblende can be used as a
geobarometer and, although the full mineral assemblage
required is not present in hornblende-bearing andesites
from Ruapehu, the amphibole compositions may provide
a broad estimate of the pressure conditions prevailing
during amphibole equilibration. Application of the four
geobarometers developed by Hammarstrom & Zen (1986),
Hollister et al. (1987), Johnston & Rutherford (1989)
and Schmidt (1992), respectively, to Ruapehu hornblende
2153
JOURNAL OF PETROLOGY
VOLUME 53
NUMBER 10
OCTOBER 2012
Fig. 8. Mg# [100Mg/(Mg þ Fe)] for clinopyroxene (a), orthopyroxene (b) and olivine (c) compositions from Ruapehu stratigraphic units
plotted against Mg# for the host-rock composition. Analyses from crystal cores and rims are differentiated.‘Other’ includes groundmass crystals, microphenocrysts and pyroxene and olivine crystals included in other phases. The lines trace the composition of pyroxene or olivine that
would be in equilibrium with the groundmass and groundmass glass compositions analysed in five Ruapehu andesites and the Ruapehu
basalt. The groundmass and glass compositions range from basaltic andesite to dacite. Equilibrium compositions were calculated using the
equation Mg#Cpx ¼ (Mg#host 100)/[(Mg#host þ KD(100 ^ Mg#host)], where KD [(Fe/Mg)mineral/(Fe/Mg)liquid] is assumed to be 0·27 for
clinopyroxene and 0·31 for olivine and orthopyroxene. Host-rock Fe2O3/FeO is assumed to be 0·25.
compositions gives pressures of 0·5^0·8 GPa for a
Mangawhero andesite and 0·5^0·9 GPa for a Whakapapa
andesite.
Iron oxides and spinels
Magnetite is ubiquitous in the Ruapehu volcanic rocks,
averaging 1·5% of the modal composition. Ilmenite is
rare in the andesites but is relatively common in the
meta-igneous xenoliths. Similarly, although aluminous
and chrome spinels are rare in Ruapehu volcanic rocks,
aluminous spinel is more common in the meta-igneous
xenoliths and some of these also contain chrome spinel.
Compositional variation in magnetite is illustrated in
Fig. 9 and data for spinel and ilmenite are reported in
Electronic Appendices A8 and A9.
In the Te Herenga Formation andesites the modal abundance of magnetite averages 2%, with some samples
containing up to 3%. Magnetites are titaniferous (Fig. 9)
2154
PRICE et al.
ANDESITE PETROGENESIS, RUAPEHU VOLCANO
Fig. 9. Spinel compositions in Ruapehu basalt and andesite and xenoliths plotted in terms of Fe2TiO4^(MgFe)Al2O4^Fe3O4 [u«lvospinel (Usp),
spinel/hercynite (Sp/Hc) and magnetite (Mt)]. (a) Te Herenga Formation; (b) Wahianoa Formation; (c) Mangawhero Formation; (d)
Whakapapa Formation; (e) xenoliths. Filled symbols in (c) are analyses of magnetite from the Ruapehu basalt.
with u«lvospinel component averaging 30% (range is 20^
47%) and Al2O3 abundance between 0·6 and 3·7 wt %.
Magnetite makes up 1% of the average Wahianoa andesite with modal abundance up to 4%. Wahianoa magnetites
are titanomagnetites with all but one of the analysed compositions containing between 22 and 53% (average 36%)
u«lvospinel component (Fig. 9). Al2O3 abundance in
Wahianoa magnetites is between 0·7 and 4 wt %. In
Mangawhero lavas, titanomagnetite with 9^80% u«lvospinel component (Fig. 9) is the dominant opaque phase,
making up, on average 1·3% of the mode. In rare cases
magnetite is more abundant (up to 6% of the rock).
Al2O3 abundance in Mangawhero magnetites varies from
0·7 to 11wt %. Magnetite is generally 56% of the mode
in Whakapapa andesites; one sample contains 10% but
the average modal abundance is 2%. Whakapapa magnetites are titaniferous; the u«lvospinel component ranges
from 23 to 73% with an average of 39% (Fig. 9). Al2O3
abundance in Whakapapa magnetites is between 1·0 and
7·2 wt % and Cr2O3 is up to 5·2 wt %.
Aluminous spinel was analysed in andesites from the
Wahianoa and Whakapapa Formations; it also occurs as
inclusions in plagioclase in Mangawhero andesite and
olivine in the Ruapehu basalt (Fig. 9). In the Mangawhero
spinel, the (Mg þ Fe)Al2O4 component accounts for
90% of the composition and the spinel and hercynite
components are present in approximately equal proportions. In the Whakapapa spinel, the (Mg þ Fe)Al2O4 component is 82% and spinel dominates over hercynite.
Chrome spinel occurs in a Mangawhero and a
Whakapapa andesite. Those analysed contain between 65
and 79% (Mg þ Fe)Cr2O4 and have Mg#s between 39
and 42%.
Ilmenite has been analysed in three andesites; one from
each of the Wahianoa, Mangawhero and Whakapapa
Formations. Compositions range from Ilm74Hm26 to
Ilm90Hm10 (where Il is ilmenite and Hm is the hematite
component). Coexisting ilmenite and magnetite have
been analysed in an andesite sample from the Wahianoa
Formation and these compositions have been used to
obtain an equilibration temperature of 9608C and log fO2
of ^10·5. The calculations used the algorithms of Lepage
(2003) and methods of Powell & Powell (1977) and
Spencer & Lindsley (1981). Oxidation conditions would
therefore appear to have been above the nickel^nickel
oxide (NNO) buffer, which is consistent with estimates for
2155
JOURNAL OF PETROLOGY
VOLUME 53
andesitic rocks from elsewhere (e.g. Arculus, 1978; Ghiorso
& Sack, 1991) and with the more general expectation that
subduction-related magmas should be relatively oxidized
(e.g. Parkinson & Arculus, 1999).
Iron oxide and spinels are abundant in the meta-igneous
xenoliths and these show considerably more compositional
variability than is observed in the host lavas. Spinel group
minerals form three distinct compositional groups.
Magnetite contains between 16 and 69% u«lvospinel and
51 to 28% (Mg,Fe)Al2O4 (Fig. 9). Mg# averages 16%.
Spinel contains 430% (Mg,Fe)Al2O4 (Fig. 9); the range is
between 30 and 87% with the average value being 75%.
Mg# averages 44%.The third compositional group, which
was identified in two meta-igneous xenoliths, is chromite.
Spinels of this group contain between 26 and 39%
(Mg,Fe)Cr2O4 with Mg# being between15 and 22%.
Ilmenite is common in all the meta-igneous xenoliths.
Compositions range from pure ilmenite to Ilm65Hm35,
with the average composition being Ilm85Hm15. The
ilmenite^magnetite geothermometer and oxygen geobarometer (Powell & Powell, 1977; Spencer & Lindsley, 1981;
Lepage, 2003) has been applied using coexisting ilmenite
and magnetite compositions from four Ruapehu metaigneous xenoliths. Temperature estimates range from 880
to 9108C and log fO2 from ^11·0 to ^13·7, which is consistent with the estimates obtained from coexisting ilmenite
and magnetite in one Wahianoa lava (see above).
Summary of physical conditions prevailing
at Ruapehu during magma storage and
transport
Coexisting ilmenite and magnetite in a Wahianoa andesite
give an equilibration temperature of 9608C, which is at
the lower end of the range of estimates obtained for
post-Te Herenga andesites using various pyroxene
geothermometers (generally in the range 950^11908C).
Pyroxenes in meta-igneous xenoliths indicate a temperature range of 750^11008C. Pressure estimates obtained
from amphibole in two post-Te Herenga andesites are in
the range 0·5^0·9 GPa, which is well within the very wide
range obtained using the Putirka et al. (2003) pyroxene^
melt geobarometer. This approach gives a pressure range
of 0·2^1GPa for post-Te Herenga andesites and 2 GPa for
the Ruapehu basalt. The wide variation in pressure estimates could be a reflection of both the mixed character of
the pyroxene population carried in each andesite and the
vertically dispersed nature of the magma storage and
plumbing system. Coexisting ilmenite and magnetite in a
Wahianoa andesite and the meta-igneous xenoliths give
similar estimates of log fO2 of ^10·5 and ^11 to ^13·7 respectively, above the NNO buffer. Plagioclase groundmass
crystals, phenocryst rims and microphenocrysts have been
used to estimate the H2O contents of the melt the in
NUMBER 10
OCTOBER 2012
post-Te Herenga andesites, with values of between 0·7 and
3·6 wt % (average 1·8 wt %) being obtained.
W H O L E - RO C K M AJ O R
A N D T R AC E E L E M E N T
G E O C H E M I S T RY
Major and trace element and Sr, Nd, and Pb isotopic data
for representative whole-rock samples from Ruapehu
are shown in Table 3. A comprehensive compilation of
whole-rock geochemical data is available as a supplementary data file (Electronic Appendix A10).
Major and trace element variations
Variation in major elements for Ruapehu whole-rock
samples is illustrated in Fig. 10 using silica variation diagrams. According to the classification of Gill (1981) most
Ruapehu rocks are medium-K, low- and high-silica andesites (Fig. 10g). Only two samples (both from the same lava
flow) classify as basalts, with a small number of Te
Herenga, Wahianoa, and Mangawhero Formation rocks
being basaltic andesites. Dacites have been sampled in
the Mangawhero Formation and some of the 1945^1996
eruptive rocks are also dacitic. Te Herenga rocks show relatively restricted distributions on the element^element
plots, Wahianoa rocks display well-defined linear arrays,
whereas Mangawhero rocks show a broader scatter.
Whakapapa lavas define a series of fields suggestive of
local chemical heterogeneity, possibly associated with particular source vents (see below).
Trace element variations are demonstrated using MgO
variation diagrams (Fig. 11) and mid-ocean ridge basalt
(MORB)-normalized multi-element plots (Fig. 12).
Chondrite-normalized REE patterns are illustrated in Fig.
13. With the exception of samples from the Te Herenga
Formation and the basalt from the Mangawhero
Formation, all the Ruapehu samples have similar chondrite-normalized REE patterns. Light REE (LREE) are
enriched relative to heavy REE (HREE) and negative
Eu anomalies are ubiquitous. On normalized extended
trace element plots (Fig. 12), all the Ruapehu rocks
show patterns with features that are characteristic of
subduction-related magmas (e.g. Pearce, 1982; Tatsumi
et al., 1986; McCulloch & Gamble, 1991; Hawkesworth
et al., 1993; Keleman et al., 2005) or continental crust (e.g.
Rudnick & Gao, 2005). Cs, Rb, Ba, K, and the LREE are
enriched relative to Y, Zr, Hf, Ti and the HREE, which
have depleted normal (N)-MORB-like abundances. Nb is
depleted relative to K and Pb is enriched relative to Ce.
In common with most other subduction-related volcanic
rocks, including basalts and andesites from intra-oceanic
arcs such as the Tonga^Kermadec arc to the north of New
Zealand, Ruapehu lavas show low abundances of Ni and
Cr. The Ruapehu basalt has 136 ppm Ni and 366 ppm Cr
2156
6·11
0·15
4·87
FeO
MnO
MgO
2157
13
5·04
11·50
1·76
7·92
2·42
La
Ce
Pr
Nd
Sm
20
Y
2·23
1·72
2·1
Hf
53
Zr
Nb
50
2
U
2·32
8·07
1·75
11·92
5·09
19
2·0
0·38
1·13
2·48
208
3·69
Sr
5
Rb
99·24
173
Pb
14
206
Ba
0·10
0·18
0·03
0·08
0·65
3·34
7·73
4·53
0·15
5·30
2·87
17·01
Th
99·84
179
Total
0·19
0·35
CO2
H2O/LOI
0·03
P2O5
H2Oþ
0·67
0·08
K2O
7·93
2·24
Fe2O3
3·25
16·94
Al2O3
Na2O
0·70
TiO2
CaO
56·60
56·32
SiO2
0·68
TH(1)
Formation: TH (1)
2
T6-20
Sample no.: T6-7
1
2·22
8·40
1·92
11·92
5·49
18
2·3
2·01
53
1
3
4·66
253
17
347
100·20
0·50
0·22
0·02
0·09
0·82
3·38
7·93
4·95
0·14
5·05
2·95
17·36
0·67
56·12
TH(6)
T6-15
3
2·47
9·25
2·11
15·03
6·97
18
2·1
1·96
53
0·57
1·71
5·18
235
17
239
99·66
0·40
0·51
0·02
0·09
0·78
3·36
7·57
4·72
0·14
4·47
3·42
17·50
0·66
56·02
TH(9)
T6-19
4
2·07
7·33
1·69
11·46
4·94
19
1·4
1·58
55
0·41
1·22
4·08
277
45
213
99·33
0·10
0·31
0·01
0·07
0·67
3·08
8·23
5·38
0·15
5·70
2·83
17·27
0·65
54·87
TH(10)
T6-83
5
2·73
10·96
2·67
20·13
9·19
20
2·7
2·81
92
0·95
2·95
9·60
243
38
293
100·01
0·19
0·34
0·05
0·11
1·19
3·32
7·34
4·40
0·13
4·82
2·68
17·32
0·68
57·44
WA(A)
R96/22
6
2·71
10·15
2·44
17·82
7·72
21
3·0
2·59
78
0·69
2·18
6·13
226
27
258
99·99
0·14
0·38
0·06
0·11
0·94
2·96
8·59
5·37
0·14
5·63
2·85
17·12
0·75
54·95
WA(B)
R96/18
7
Table 3: Representative whole-rock analyses for Ruapehu Volcano
8
2·78
11·26
2·82
20·35
9·13
20
3·3
3·02
105
0·92
2·55
8·55
301
33
357
99·57
0·13
0·22
0·31
0·11
1·19
3·88
7·44
2·34
0·09
5·97
19·22
0·68
57·99
WA(C)
R95/18
9
2·99
13·30
3·32
25·13
11·90
20
4·3
3·02
107
1·05
3·19
8·54
329
36
373
99·22
0·13
0·27
0·02
0·12
1·23
3·78
7·73
2·38
0·09
5·52
19·81
0·66
57·48
WA(C)
R95/28
10
3·21
14·25
3·52
27·89
12·68
23
4·6
3·74
129
1·23
4·23
11·34
267
57
401
99·63
0·24
0·50
0·05
0·14
1·61
3·39
6·44
3·42
0·11
4·90
1·71
17·23
0·75
59·14
WA(D)
R96/6
11
3·59
16·29
4·17
31·54
15·98
26
5·0
3·53
130
1·36
4·54
12·58
246
60
409
99·69
0·12
0·20
0·06
0·13
1·69
3·55
6·23
2·87
0·10
4·07
1·99
17·49
0·71
60·48
WA(D)
R96/7
12
2·02
7·69
1·64
10·42
4·97
18
1·8
1·73
57
0·33
1·18
3·10
201
12
202
100·43
–0·23
0·10
0·57
2·64
9·63
9·01
0·15
9·60
15·50
0·66
52·79
MA
R04/4
13
2·29
8·66
1·91
14·59
5·92
20
2·8
2·15
73
0·52
1·96
4·87
229
19
255
100·10
0·19
0·12
0·79
2·97
8·51
5·03
0·16
9·10
17·89
0·69
54·67
MA
R04/3
14
3·13
14·01
3·46
27·04
12·91
21
4·8
3·23
122
1·35
4·83
11·74
229
63
346
99·78
0·10
0·40
0·06
0·14
1·55
3·29
6·94
3·53
0·14
5·35
1·56
17·25
0·72
58·75
MA
R97/6
15
3·88
18·36
4·89
38·61
17·51
22
6·4
4·12
160
1·72
6·54
16·85
302
83
499
99·89
0·04
0·47
0·06
0·20
2·14
3·88
5·25
2·48
0·11
3·95
1·45
16·50
0·80
62·56
MA
R97/10a
16
3·71
16·43
4·13
31·04
13·79
22
5·9
4·47
148
1·78
6·52
14·40
231
73
434
99·80
0·12
0·54
0·10
0·17
1·88
3·16
6·39
5·17
0·13
4·91
1·24
15·13
0·79
60·07
MA
R97/65a
17
3·19
14·44
3·56
27·28
12·74
22
4·9
4·15
121
1·36
5·12
12·57
229
62
352
99·99
0·11
0·56
0·12
0·14
1·55
3·45
6·93
3·57
0·14
5·22
1·60
17·19
0·72
58·70
MA
18
3·52
15·93
4·11
32·85
15·12
23
5·4
4·12
137
1·80
5·98
13·92
259
75
368
100·96
1·63
0·34
0·06
0·14
1·87
3·21
6·32
4·21
0·11
4·20
1·52
16·11
0·69
60·55
MA
X1-16
(continued)
R97/65b
PRICE et al.
ANDESITE PETROGENESIS, RUAPEHU VOLCANO
0·493
2·977
0·711
1·857
1·816
Tb
Dy
Ho
Er
Yb
2158
45
74
16
Ni
Cu
Zn
Ga
208
207
206
Pb 18·746
Pb 15·578
204
204
0·512903
Pb/204Pb 38·518
Pb/
Pb/
Nd/144Nd
29·6
38·634
15·610
18·770
0·512905
0·704869
16
68
28
20
42
210
38·636
15·608
18·779
0·512926
0·704904
17
66
36
29
44
188
25·3
0·288
1·819
1·712
0·562
2·617
0·452
2·449
0·781
TH(6)
4
38·682
15·619
18·809
0·51288
0·70506
17
65
80
24
57
175
25·3
0·265
1·703
1·665
0·599
2·712
0·437
2·652
0·822
TH(9)
T6-19
5
38·653
15·613
18·789
0·512916
0·704799
16
74
39
25
61
217
34
0·291
1·911
1·771
0·630
2·770
0·439
2·653
0·712
TH(10)
T6-83
6
0·512756
0·705490
18
72
55
31
54
205
25·7
0·316
1·988
1·834
0·675
3·075
0·504
3·092
0·778
WA(A)
R96/22
7
38·711
15·623
18·822
0·512780
0·705384
19
78
54
25
74
259
27·4
0·309
1·965
1·882
0·685
3·109
0·495
3·047
0·811
WA(B)
R96/18
8
38·709
15·625
18·825
0·512774
0·705209
18
67
65
14
10
194
4·3
0·280
1·749
1·738
0·632
2·827
0·485
3·059
0·822
WA(C)
R95/18
9
38·664
15·611
18·815
0·512784
0·705102
18
57
51
19
26
181
14·6
0·300
2·060
1·964
0·686
3·149
0·514
3·266
1·039
WA(C)
R95/28
10
0·705722
19
70
31
12
36
187
20·6
0·315
1·989
1·926
0·680
3·232
0·552
3·373
0·884
WA(D)
R96/6
11
38·660
15·606
18·815
0·512712
0·705728
19
72
38
12
26
166
18·7
0·337
2·082
2·082
0·753
3·655
0·600
3·826
0·914
WA(D)
R96/7
12
38·608
15·603
18·769
0·512931
0·704746
17
77
76
136
366
256
32·2
0·235
1·551
1·626
0·582
2·523
0·401
2·272
0·748
MA
R04/4
13
0·705115
18
78
54
22
49
236
25·1
0·255
1·703
1·732
0·608
2·606
0·419
2·476
0·797
MA
R04/3
14
0·512691
0·705717
23
69
25
14
49
169
22·7
0·330
1·971
1·975
0·692
3·192
0·552
3·480
0·908
MA
R97/6
15
0·512689
0·705494
23
59
25
10
41
125
15·4
0·306
2·058
1·995
0·737
3·463
0·576
4·180
1·093
MA
R97/10a
16
38·625
15·600
18·795
0·512749
0·705149
22
59
19
80
243
164
23·9
0·343
2·204
2·255
0·791
3·593
0·605
4·187
1·063
MA
R97/65a
17
18
38·740
15·633
18·832
0·512680
0·705648
18
67
52
33
110
148
21·6
0·302
1·978
1·874
0·693
3·280
0·560
3·560
0·907
MA
X1-16
(continued)
0·512684
0·705641
23
70
30
15
53
172
22·7
0·339
2·062
2·106
0·738
3·273
0·560
3·672
0·893
MA
R97/65b
NUMBER 10
143
0·704830
23
Cr
Sr/86Sr
49
V
0·293
1·827
1·814
0·638
2·926
0·459
2·679
0·795
3
T6-15
VOLUME 53
87
28·0
235
Sc
0·295
2·793
Gd
Lu
0·806
Eu
TH(1)
Formation: TH (1)
2
T6-20
Sample no.: T6-7
1
Table 3: Continued
JOURNAL OF PETROLOGY
OCTOBER 2012
2159
6·88
15·14
2·19
9·06
2·20
Ce
Pr
Nd
Sm
17
Y
La
2·37
2·8
63
Zr
Hf
2
U
Nb
7·78
2
Sr
Pb
23
228
Rb
Th
99·72
233
Ba
0·94
Total
0·48
H2O
0·01
CO2
H2Oþ/LOI
0·09
P2O5
4·07
MgO
0·86
0·13
MnO
K2O
4·21
7·32
2·82
Fe2O3
FeO
3·37
17·60
Al2O3
Na2O
0·57
TiO2
CaO
57·25
SiO2
3·31
16·58
4·42
31·57
14·89
20
5·5
4·53
136
2
7
14·95
214
69
365
99·52
0·43
0·17
0·01
0·13
1·78
3·12
6·42
6·17
0·11
4·62
1·21
14·76
0·71
59·88
MA
Formation: MA
20
T5/62
Sample no.: T5/28c
19
Table 3: Continued
21
2·42
7·92
5·74
45·71
21·65
25
8·9
3·06
236
3·46
12·68
20·74
202
129
520
100·15
0·34
0·16
0·01
0·17
2·99
3·39
4·81
3·46
0·08
3·88
0·89
14·72
0·77
64·47
MA
T5-11
22
4·40
21·38
5·62
43·89
20·58
24
8·3
6·27
229
4
10
20·00
202
125
503
99·92
0·43
0·19
0·02
0·16
2·89
3·33
5·03
3·49
0·09
3·92
0·80
14·94
0·73
63·90
MA
T5-16
23
3·72
16·34
4·24
34·88
16·32
20
5·9
3·81
148
1·74
6·30
15·84
268
81
433
100·13
0·08
0·56
0·09
0·15
1·61
3·13
7·12
4·68
0·14
4·48
2·33
15·99
0·70
59·08
WH
R97/19
24
3·25
14·00
3·60
28·53
13·18
24
4·7
3·91
110
1·55
5·12
11·88
235
62
358
99·69
0·40
0·11
0·07
0·13
1·59
3·10
7·15
4·60
0·13
5·30
1·67
16·48
0·71
58·25
WH(SC)
R96/26
25
2·96
13·29
3·36
23·47
10·85
19
4·1
3·06
93
2
5
10·41
281
45
323
99·94
0·17
0·18
0·02
0·11
1·38
3·21
7·58
5·14
0·13
5·69
1·24
16·57
0·68
57·84
WH(DC)
T6-36
26
3·12
13·56
3·31
23·83
10·96
19
4·4
2·95
93
4
10
10·24
279
46
324
98·97
0·13
0·38
0·01
0·11
1·38
3·10
7·47
5·10
0·13
5·12
1·87
16·50
0·68
56·98
WH(DC)
T6-65
27
28
R97/82
29
T5-87
30
R97/4
31
R97/23
3·67
16·24
4·07
31·59
14·25
20
5·1
3·48
122
1·61
4·93
9·28
269
58
380
99·79
0·14
0·41
0·05
0·15
1·61
3·44
6·99
4·11
0·14
5·21
1·16
17·04
0·67
58·67
3·43
15·85
3·97
31·19
14·27
20
5·1
4·08
124
1·55
4·90
10·05
273
61
374
99·65
0·13
0·43
0·06
0·14
1·59
3·45
6·98
4·16
0·14
5·11
1·25
16·72
0·67
58·81
3·29
14·79
3·92
28·62
13·36
20
5·0
4·13
116
1
5
12·97
266
56
367
99·40
0·50
0·14
0·01
0·12
1·60
3·38
6·80
4·47
0·12
4·70
1·56
16·40
0·66
58·94
3·30
14·29
3·55
27·58
12·60
21
4·7
3·14
1·23
4·58
11·29
248
57
332
99·53
0·10
0·29
0·05
0·15
1·47
3·19
7·46
4·11
0·15
5·34
1·96
16·86
0·77
57·64
3·67
16·95
4·25
33·37
15·41
20
5·7
3·87
144
1·55
5·77
15·13
298
71
420
100·06
0·04
0·53
0·08
0·20
2·07
3·72
5·80
2·58
0·11
4·69
0·86
16·85
0·83
61·71
WH(OW) WH(OW) WH(OW) WH(SW) WH(SE)
R97/78
32
3·60
16·52
4·07
31·82
13·73
23
5·7
3·90
146
1·59
6·51
10·78
239
64
431
99·86
0·27
0·32
0·40
0·13
1·71
3·34
6·46
3·37
0·12
7·03
17·03
0·73
58·95
WH(RP)
R95/15
33
34
R95/9
3·85
17·69
4·41
34·61
15·31
22
5·3
3·74
159
1·69
6·86
11·02
288
68
462
99·78
0·10
0·30
0·03
0·14
1·84
3·24
6·65
4·18
0·12
6·92
15·92
0·74
59·60
3·80
18·00
4·40
35·28
15·49
21
6·2
3·90
162
1·69
6·98
11·32
287
68
473
100·11
0·25
0·48
0·04
0·14
1·84
3·25
6·52
4·06
0·12
6·87
16·14
0·75
59·65
WH(RM) WH(RD)
R95/10
35
1·3
3·3
0·7
3·8
1·3
10
0·3
0·4
12·7
0·1
0·1
1·2
152
4·2
49
36
2·65
12·70
14·9
5·73
19
3
1·89
79
0·10
0·51
2
350
4
70
99·05
0·20
0·14
0·27
2·49
11·16
9·12
0·13
6·99
1·05
18·09
0·89
48·52
Kak
TVZ15
(continued)
99·64
0·03
–0·33
0·05
0·11
0·93
13·33
10·30
0·18
11·08
15·55
0·56
47·85
Ker
7135
PRICE et al.
ANDESITE PETROGENESIS, RUAPEHU VOLCANO
2160
19
36
69
16
Ni
Cu
Zn
Ga
Pb 18·827
Pb 15·621
204
204
0·512820
38·608
15·595
18·799
0·51276
38·711
15·622
18·825
0·512744
0·705402
15
52
55
43
116
128
38·609
15·594
18·802
0·512755
0·705301
16
51
39
43
114
130
15·1
0·364
2·363
2·323
0·797
3·981
0·709
4·643
1·023
MA
23
0·512654
0·705918
23
56
27
22
71
126
18·2
0·285
1·911
1·811
0·661
3·216
0·544
3·677
0·974
WH
R97/19
24
38·445
15·578
18·772
0·512719
0·705321
18
63
31
27
84
196
24·4
0·338
2·179
1·999
0·721
3·381
0·564
3·448
0·842
WH(SC)
R96/26
25
38·604
15·592
18·801
0·512772
0·705154
17
65
49
39
74
182
23·5
0·308
1·945
1·870
0·622
3·124
0·549
3·219
0·959
WH(DC)
T6-36
26
38·623
15·597
18·804
0·512777
0·705137
16
66
56
42
73
185
23·8
0·301
1·859
1·876
0·642
3·146
0·536
3·213
0·948
WH(DC)
T6-65
27
28
R97/82
29
T5-87
30
R97/4
31
R97/23
38·653
15·607
18·811
0·512759
0·705325
23
59
38
33
72
166
24·6
0·313
2·014
1·943
0·707
3·335
0·582
3·679
0·982
0·512730
0·705316
23
63
52
33
74
159
25·8
0·315
2·067
1·932
0·708
3·355
0·549
3·578
0·964
38·556
15·579
18·791
0·512755
0·705298
16
66
53
37
114
160
20·5
0·334
2·130
1·978
0·693
3·329
0·573
3·570
0·974
0·512728
0·705383
22
71
37
14
58
199
25·5
0·312
2·004
2·009
0·699
3·252
0·542
3·614
0·951
32
0·691
38·726
15·626
18·833
0·512716
0·705346
24
62
34
6
12
167
17·3
0·305
1·992
2·002
38·678
15·614
18·825
0·512689
0·705743
16
74
24
9
29
183
21·5
0·341
2·208
2·371
0·798
3·543
0·593
0·571
3·228
3·449
0·981
WH(RP)
R95/15
3·914
1·064
WH(OW) WH(OW) WH(OW) WH(SW) WH(SE)
R97/78
33
34
R95/9
38·626
15·598
18·811
0·512697
0·705644
17
63
29
14
63
171
23·3
0·312
2·016
2·156
0·744
3·429
0·589
3·503
0·996
38·719
15·626
18·838
0·512683
0·705680
16
67
30
17
59
177
22·5
0·305
1·996
2·169
0·752
3·442
0·582
3·456
0·985
WH(RM) WH(RD)
R95/10
35
38·342
15·571
18·648
0·51306
0·70342
11
51
73
98
228
353
52·6
0·2
1·2
1·1
0·4
1·9
0·3
1·5
0·5
Ker
7135
36
38·613
15·600
18·808
0·512913
0·703878
14
70
56
113
85
191
28
0·23
1·60
0·40
0·91
Kak
TVZ15
Values in italics are ICM-MS data. Other data were obtained by XRF. 34 from Smith et al. (2010); 35 from Gamble et al. (1993a, 1996). TH, Te Herenga; WA,
Wahianoa; MA, Mangawhero; A–D, flow units in Wahianoa Formation; WH, Whakapapa; SC, Saddle Cone; DC, Delta Corner; OW, older Whakapapa; SE and
SW, Sunset East and West; RP, RM, RD, proximal, medial and distal Rangataua; Ker, Kermadec basalt; Kak, TVZ basalt–Kakuki basalt; LOI, loss on ignition.
Pb/204Pb 38·700
208
Pb/
207
Pb/
206
Nd/144Nd
0·70538
13
61
63
109
306
165
14·6
0·355
2·315
2·297
0·845
2·977
0·493
2·793
0·806
MA
22
T5-16
NUMBER 10
143
0·70526
30
Cr
Sr/86Sr
171
V
20·4
0·328
2·186
2·033
0·672
21
T5-11
VOLUME 53
87
20·1
Sc
1·951
Yb
0·310
1·742
Er
Lu
0·580
Ho
3·364
0·593
0·482
2·771
Tb
Dy
3·596
2·349
Gd
0·856
0·749
Eu
MA
Formation: MA
20
T5/62
Sample no.: T5/28c
19
Table 3: Continued
JOURNAL OF PETROLOGY
OCTOBER 2012
PRICE et al.
ANDESITE PETROGENESIS, RUAPEHU VOLCANO
Fig. 10. Variation of selected major element oxides vs SiO2 (wt %) for Ruapehu volcanic rocks. Fields shown in (g) are from Gill (1981).
2161
JOURNAL OF PETROLOGY
VOLUME 53
NUMBER 10
OCTOBER 2012
Fig. 11. Variation of selected trace elements (ppm) vs MgO (wt %) for Ruapehu volcanic rocks.
but only nine of over 440 analysed andesites have Ni
4100 ppm and only 18 have Cr 4250 ppm. The mean Ni
and Cr abundances for all Ruapehu volcanic rocks are
26 ppm and 75 ppm respectively.
Pb^Sr^Nd isotopic compositions
Strontium, Nd and Pb isotopic variations for the Ruapehu
samples are shown in Fig. 14. On the Sr^Nd isotopic diagram (Fig. 14a and b), the field defined by post-Te
2162
PRICE et al.
ANDESITE PETROGENESIS, RUAPEHU VOLCANO
Fig. 12. MORB-normalized extended trace element plots for selected Ruapehu volcanic rocks. (a) AD 1945^1996 eruptive rocks of the
Whakapapa Formation; (b) prehistoric Whakapapa Formation with sub-units identified: SC, Saddle Cone; DC, Delta Corner; WS, older
Whakapapa skifield; Rang., Rangataua [proximal (P), medial (M), distal (D)]; (c) Mangawhero Formation basalt and andesite; (d)
Mangawhero Formation dacites; (e) Wahianoa Formation with sub-units identified; (f) Te Herenga Formation. Grey field in (a)^(e) is field of
Te Herenga Formation from (f). Abundances have been normalized to N-MORB values from Sun & McDonough (1989).
Herenga Ruapehu rocks extends to higher 87Sr/86Sr and
lower 143Nd/144Nd than the array defined by TVZ basalts,
overlapping the field of TVZ rhyolites. With 87Sr/86Sr of
0·70475 and 143Nd/144Nd of 0·51293, the Ruapehu basalt
plots to the right (higher 87Sr/86Sr) of the TVZ array.
Ruapehu eruptive rocks show a restricted range in Pb
isotopic composition (e.g. 206Pb/204Pb ranges from 18·769
to 18·865) with the Ruapehu basalt having the least radiogenic composition (206Pb/204Pb ¼18·769, 207Pb/204Pb ¼
15·603, 208Pb/204Pb ¼ 38·608). On Pb^Pb isotope diagrams
(Fig. 14) the basalt composition marks the lower end of an
array that begins below the low 206Pb/204Pb end of the
TVZ basalt field extending across and overlapping with
the fields defined by TVZ basalts and rhyolites.
2163
JOURNAL OF PETROLOGY
VOLUME 53
NUMBER 10
OCTOBER 2012
Fig. 13. Chondrite-normalized REE plots for selected Ruapehu volcanic rocks. (a) AD 1945^1996 eruptive rocks of the Whakapapa Formation;
(b) prehistoric Whakapapa Formation with sub-units identified: SC, Saddle Cone; DC, Delta Corner; WS, older Whakapapa skifield; Rang.,
Rangataua [proximal (P), medial (M), distal (D)]; (c) Mangawhero Formation basalt and andesite; (d) Mangawhero Formation dacites; (e)
Wahianoa Formation with sub-units identified; (f) Te Herenga Formation. Darker grey field in (a)^(e) is field of Te Herenga Formation from
(f). Lighter grey field in (d) is the field of groundmass and groundmass glass compositions from Ruapehu andesites. Abundances have been normalized to the chondrite average of Sun & McDonough (1989).
Price et al. (2005) noted that 87Sr/86Sr ratio (and K2O
abundance) in Ruapehu lavas shows a general increase
from older to younger eruptive rocks. With time there is a
tendency for more evolved magmas to be erupted.
Variability is most limited in the Te Herenga flows and
most extreme in the Mangawhero Formation. 87Sr/86Sr
ratio also shows a crude positive correlation with SiO2
abundance (Graham & Hackett, 1987).
The Te Herenga Formation
Te Herenga samples show comparatively limited compositional variability. SiO2 and MgO abundances are in the
range of 54·9^58 wt % and 5·4^4·4 wt % respectively.
Te Herenga andesites are relatively aluminous
(Al2O3 ¼16·74^17·95) and they are distinctly less potassic
than other Ruapehu lavas, with compositions plotting into
the low-K andesite field (Fig. 10). P2O5 abundances are
low relative to other Ruapehu rocks. Thirteen lavas flows
have been systematically sampled and analysed from a sequence exposed along the lower slopes of Pinnacle Ridge
with the objective being to assess temporal geochemical
variation. All the samples from this sequence show a limited range in composition and systematic, stratigraphically
controlled variation is not observed.
The trace element compositions are characterized by
low abundances of REE and relatively flat chondritenormalized REE patterns [(La/Yb)n ¼1·31^2·78; Fig. 13f]
and low Rb (10^20 ppm) and Zr (50^63 ppm) contents
2164
PRICE et al.
ANDESITE PETROGENESIS, RUAPEHU VOLCANO
Fig. 14. (a) 143Nd/144Nd vs 87Sr/86Sr for Ruapehu volcanic rocks and xenoliths compared with fields for regional Kermadec^New Zealand volcanic rocks. H, Havre Trough; K, Kermadec islands and seamounts; Ba, Taupo Volcanic Zone (TVZ) basalts; R, TVZ rhyolites; RB, Ruapehu
basalt (R04/04). Regional data from Ewart & Hawkesworth (1987), Gamble et al. (1993a, 1996), McCulloch et al. (1994) and Sutton et al. (1995).
T and W are average compositions for Torlesse and Waipapa basement terranes of the North Island (data from this study). MIX is an average
meta-igneous xenolith composition (this study and Graham et al., 1990). (b) is an enlargement of part of (a) to show the details of the
Ruapehu data. (c) and (d) 207Pb/204Pb and 208Pb/204Pb vs 206Pb/204Pb for Ruapehu volcanic rocks and comparisons with fields for regional
Kermadec^New Zealand volcanic rocks. Shadings, lettering and symbols are as in (a) and (b). NHRL is the Northern Hemisphere Reference
Line of Hart (1984).
compared with other Ruapehu andesites (Fig. 11). Te
Herenga andesites have K2O, Rb and LREE contents
that are comparable with or only slightly higher than
those observed in TVZ basalts and they generally have
lower abundances of Zr [see Table 3 and basalt data of
Gamble et al. (1993a)]. Their Ba contents are generally
higher than the abundances observed in TVZ basalts and
they show considerable variability; with the exception of
one sample, which has a Ba content of 347 ppm, Te
Herenga lavas have Ba abundances ranging from 167 to
280 ppm. However, compared with other Ruapehu rocks
Te Herenga andesites have distinctly lower Ba abundances.
Sr abundances are also relatively low (average 210 ppm)
and chondrite-normalized REE patterns manifest more
subdued Eu anomalies (Eu/Eu* ranges from 0·91 to 1·04).
Sc, V, and Ni abundances are similar to those observed in
other Ruapehu andesites but Cr contents are lower.
Te Herenga Sr and Nd isotopic compositions are
also distinct (Fig. 14). 143Nd/144Nd ratios (0·51287^0·51293)
are higher than those in post-Te Herenga lavas
(0·51264^0·51282) and the Sr^Nd data define an array that
parallels the overall post-Te Herenga Ruapehu trend but
is displaced to lower 87Sr/86Sr (0·7048^0·7052 vs 0·7050^
0·7061). The only post-Te Herenga rock with Sr^Nd isotope
ratios similar to the Te Herenga Formation is the Ruapehu
basalt (Fig. 14b). Te Herenga Pb isotopic compositions are
generally similar to those observed in other Ruapehu
rocks although 206Pb/204Pb tends to be lower than in most
of the younger eruptive rocks. Together with the Ruapehu
basalt, Te Herenga Pb isotope ratios define slightly shallower trends than the post-Te Herenga volcanic rocks
(Fig. 14c and d).
The Wahianoa Formation
Wahianoa andesites tend to have lower K2O, Ba, Rb and
Zr abundances than their counterparts in younger flow formations (Figs 10 and 11). K2O abundances are broadly
intermediate between those observed in the Te Herenga
2165
Fig. 15. Variations of SiO2, Al2O3, MgO and K2O abundances and 87Sr/86Sr as functions of stratigraphic position within the Wahianoa Formation. Samples (numbers shown at left) are arranged
in stratigraphic order from oldest at the bottom to youngest at the top. Letters A^E refer to flow units mapped and sampled on eastern Ruapehu in the Wahianoa and Whangaehu River catchments (see Fig. 4). O is a suite of samples from a section in the Ohinepango River (Fig. 1). Samples connected by tie-lines are believed to have been erupted sequentially without significant
breaks in time. Numbers at the right are Ar^Ar ages from Gamble et al. (2003).
JOURNAL OF PETROLOGY
VOLUME 53
2166
NUMBER 10
OCTOBER 2012
PRICE et al.
ANDESITE PETROGENESIS, RUAPEHU VOLCANO
Formation and those of the Mangawhero and Whakapapa
Formations. Compared with the Te Herenga samples,
Wahianoa lavas have chondrite-normalized REE patterns
(Fig. 13) that show stronger enrichment of LREE over
HREE [Fig. 13; (La/Yb)n ¼ 2·67^5·43] and more strongly
developed negative Eu anomalies (Eu/Eu* ¼ 0·75^1·07).
Plagioclase-phyric andesites [similar to the Type 2 andesites of Graham & Hackett (1987)] occur in the middle
of the Whangaehu Gorge and Wahianoa sections (Fig. 15).
In contrast to other andesites these show comparatively
elevated Al2O3, CaO and Sr and lower FeO* and MgO
abundances (Figs 10 and 11), which is consistent with
slightly elevated modal abundances of plagioclase and
lower modal abundances of pyroxene (see above). Type 2
lavas also show a subtle difference in Eu anomaly; Eu
anomalies are only slightly negative with Eu/Eu* averaging 0·95 compared with 0·84 for all other Wahianoa
lavas (Fig. 13).
87
Sr/86Sr and 143Nd/144Nd ratios in the Wahianoa
Formation show a range that covers most of the spectrum
of isotopic compositions observed in post-Te Herenga
Ruapehu lavas (Fig. 14). Only the Mangawhero Formation
shows as extensive a range in isotopic composition.
87
Sr/86Sr ranges from 0·7049 to 0·7059 and 143Nd/144Nd
between 0·51269 and 0·51281. The Wahianoa Formation
shows the widest variations in Pb isotopic composition
of any of the Ruapehu flow formations. 206Pb/204Pb
values vary between 18·77 and 18·650, 207Pb/204Pb from
15·597 to 15·673 and 208Pb/204Pb from 38·601 to 38·870
(Fig. 14).
Detailed sampling has been carried out through sections
of the Wahianoa Formation exposed in the Whangaehu
Gorge (Fig. 4b and c), the upper Wahianoa valley and in
the Ohinepango stream (Fig. 1), and the distribution of
flows between the Whangaehu and Wahianoa Gorges has
been mapped in detail (Fig. 4a). Data from each of these
sections have been combined into a single stratigraphic
column with age control provided by Ar^Ar dating
(Gamble et al., 2003) and five flow groups (Units E^A in
Fig. 4a) have been recognized. The objective of this detailed sampling was to evaluate short-term (flow-to-flow)
geochemical variation and to use this information to develop a temporal model for geochemical evolution in the
magmatic system operating beneath the volcano at the
time the Wahianoa Formation was emplaced (see Gamble
et al., 2003). Selected results from this work are shown in
Fig. 15, which summarizes the variation for SiO2, Al2O3,
MgO, K2O and 87Sr/86Sr for the Wahianoa flow sequence
on eastern Ruapehu. Variations are not consistently systematic, either in single flow sequences or in the sampled
stratigraphic section overall, although some of the flow
units show distinctive compositional characteristics. For example, unit C comprises flows of Type 2 andesite with
higher Al2O3 and lower MgO abundances and Unit E
has higher MgO and lower Al2O3 contents than other
flows in the sequence. From Unit O to Unit C there appears to be an overall subtle trend of increasing SiO2 and
K2O and decreasing MgO abundance but this impression
is largely created by the very distinctive end-member
effect of Unit C. From Unit O to Unit B these trends are
not obvious; the ranges of SiO2, K2O and MgO abundances are similar in Units O, A and B.
Within the Wahianoa^Whangaehu and Ohinepango sequences groups of 2^3 flows have been recognized that
appear to have been erupted in quick succession; there is
no evidence for erosion of, or soil development on, the
stratigraphically lower flows, and the flows within each
group are not separated by pyroclastic units. Rapidly deposited flows of this type (five sets in the Wahianoa^
Whangaehu section and one in the Ohinepango sequence)
are flagged by tielines in Fig. 15 and may provide insights into relatively short-term fluctuations in magma
chemistry. In each case, SiO2 and K2O abundances are
lower in the younger flow whereas MgO content is
higher. In all but one case 87Sr/86Sr decreases from oldest
to youngest flow. Al2O3 shows limited variation with a decrease from oldest to youngest flow apparent in three of
the groups.
The Mangawhero Formation
The Mangawhero Formation includes the complete spectrum of compositional variation from basalt to dacite.
Most lavas show K2O abundances and isotopic compositions that are similar to those observed in other post-Te
Herenga andesites but a low-K group exposed in the
upper Whakapapaiti valley has some major and trace
element similarities to the Te Herenga Formation flows.
With the exceptions of the Ruapehu basalt and one of
the low-K series andesites, Mangawhero samples have Sr
and Nd isotopic compositions plotting within the array
defined by post-Te Herenga lavas (Fig. 14). The most
evolved Sr and Nd isotopic compositions observed on
Ruapehu occur in high-silica andesites from this formation
but Mangawhero Formation dacites (87Sr/86Sr50·7057) do
not share these characteristics. Two samples of low-K
series andesite also have relatively unevolved Sr and
Nd isotopic compositions (87Sr/86Sr ¼ 0·70526^0·70528;
143
Nd/144Nd ¼ 0·51277^0·51282) with one of these samples
having a composition intermediate between the Te
Herenga and the post-Te Herenga arrays. Pb isotope
ratios in Mangawhero andesites and dacites show limited
variation (206Pb/204Pb 18·795^18·832, 207Pb/204Pb 15·594^
15·633, 208Pb/204Pb 38·608^38·740) and plot within the
main Ruapehu array (Fig. 14c and d).
The Whakapapa Formation
Chemical compositions inWhakapapa Formation lavas are
almost as diverse as those observed in andesites of the
Mangawhero Formation. They are dominantly Type 1
2167
JOURNAL OF PETROLOGY
VOLUME 53
NUMBER 10
OCTOBER 2012
Fig. 16. Variation of Al2O3, MgO, K2O and Zr abundances, (La/Yb)n and 87Sr/86Sr vs SiO2 (wt %) for volcanic rocks from the Whakapapa
Formation of Ruapehu volcano. The sub-units of the formation (see Fig. 2) are differentiated (see legend). The shaded grey field is for AD 1945^
1996 eruptive rocks of the Whakapapa Formation (Gamble et al., 1999). ‘Prox.’, ‘Med.’, and ‘Dist.’ are proximal, medial and distal Rangataua
flows, respectively.
plagioclase^pyroxene andesites with SiO2 ranging from
55·5 to 61·7 wt %, MgO from 5·41 to 2·29 wt % and K2O
from 1·23 to 2·14 wt %. Major and trace element compositional ranges overlap those in Mangawhero andesites and
are only slightly more restricted. 87Sr/86Sr ranges from
0·7051 to 0·7059 in all analysed Whakapapa lavas, slightly
less than the total range for all post-Te Herenga andesites
(0·7048^0·7061) (Fig. 14).
Detailed stratigraphic sampling reveals contrasting patterns of geochemical variability within and between
Whakapapa flow groups (Fig. 16). The relatively young
Delta Corner flows on the Whakapapa skifield are less
evolved than other Whakapapa andesites and are characterized by relatively high MgO and lower SiO2, Al2O3,
K2O, and Zr abundances. They also have relatively unevolved Sr and Nd isotopic compositions and low
(La/Yb)n (Fig. 16).
Sunset East flows are more evolved with elevated SiO2,
Al2O3, K2O, and Zr abundances, higher (La/Yb)n ratios
and higher 87Sr/86Sr compared with the Delta Corner
flows. The most evolved Sr^Nd isotope compositions
(0·70583^0·70564, 0·51268^0·51272) in the Whakapapa
Formation occur in the Rangataua flows, but these
evolved isotopic signatures are associated with SiO2,
MgO, K2O and Zr abundances in the middle of the
range shown by all samples from the formation. Al2O3
abundances in the Rangataua flows vary widely (15·9^
18 wt %, Fig. 16).
2168
6·92 0·80
2169
0·33 0·14
Cl
3·3
0·82
3·27
3·51
1·95
Sm
Eu
Gd
Dy
Er
30·5 4·7
26·7
14·2
Ce
Nd
2·16 0·34
3·77 0·59
3·61 0·45
0·95 0·14
3·7 0·6
16·4 2·5
4·18
7·31
6·54
1·73
7·3
34·7
69·5
33·6
6·95 2·25
11·64 3·00
10·73 2·82
1·70 0·28
10·7 2·3
47·8 9·5
88·9 18·2
41·5 8·4
2·91
5·04
5·20
0·87
5·4
24·1
49·0
22·9
2·78
14·7 2·3
8·8 3·8
12·7
5·24
La
1·98 0·28
1·75
208
U
31·1 13·6
194 9
114
9·5
18·0
295
203 40
598
LA
3·19
4·61
2·36
0·20
18·2
6·8 0·9
189 25
233
918 162
LA
0·10 0·03
2·66 0·57
2·49 1·08
3·98 1·82
0·40 0·02
2·45
6·0
162
Sr
91 19
944
LA
0·12 0·04
3·00 0·23
1·74 0·39
2·49 0·45
0·52 0·04
3·59 0·80
0·17 0·10
Th
77
Rb
425 77
LA
0·31 0·12
0·15 0·04
2·71 0·26
4·82 0·72
3·73 0·71
0·15 0·04
Pb
368
Ba
LA
Trace elements (ppm)
0·15 0·05
P2O5
3·03 0·47
K2O
2·58 0·90
3·46 0·80
3·64 0·61
CaO
Na2O
0·17 0·09
1·08 0·41
MgO
6·24 1·29
6·24 1·29
4·94 1·06
0·13 0·09
FeO
MnO
26·5
1·88
3·41
3·19
0·86
3·3
14·7
2·89
4·93
5·18
1·00
5·2
23·6
46·6
21·7
2·79
1·69
12·6
10·0
16·1
195
100
554
LA
3·16 0·10
4·57 0·10
3·57 0·25
0·69 0·23
3·34 0·73
5·6
185
77
371
LA
0·15 0·04
2·58 0·18
3·49 0·19
5·06 0·48
2·06 0·43
0·17 0·06
14·72 0·72 16·03 0·54
0·80 0·12
14·45 0·17 14·63 1·26 14·11 0·79 15·60 2·17 15·39
0·63
Al2O3
0·94 0·23
1·15 0·20
1·29 0·28
63·69 1·34 67·83 0·73
1·04 0·14
0·98 0·14
Av. GI-Opx Av. Gms Gls Av. Gms
Whakapapa
R97/82
7
67·82 2·03 63·46 1·74 72·85 0·57 70·07 3·63 71·17
GI-Plag
P57536
6
Whakapapa 1996 July
R96/82
5
TiO2
GI-Plag
Waihianoa
R95/28
4
SiO2
Major components (wt %)
GI-Plag
GI-Plag
description:
Sample
R96/6
3
Whakapapa Whakapapa Waihianoa
R96/27
2
Fm/date:
Sample no.:R95/82
1
2·91
4·76
4·72
1·05
4·6
19·6
39·2
17·8
2·28
8·2
12·7
256
92
696
MD
Gms
Whakapapa
R96/27
8
Table 4: Major trace, trace element and isotopic data for Ruapehu glasses and groundmasses
3·24
5·31
5·41
1·06
5·4
23·0
44·6
20·5
2·11
8·3
16·2
198
83
515
LA
3·14 0·15
4·11 0·11
3·45 0·31
0·79 0·11
3·57 0·30
15·27 0·56
0·67 0·09
69·00 0·47
Av. Gms
Waihianoa
R96/6
9
2·46
4·22
4·62
1·08
5·0
24·0
54·0
27·4
1·56
11·0
6·9
322
58
402
MD
Gms
Waihianoa
R95/28
10
2·90
4·89
5·21
0·59
5·3
24·0
49·4
22·4
3·05
10·9
23·1
48
139
533
LA
0·12 0·03
0·34 0·04
4·46 0·96
3·66 0·69
2·13 0·08
0·18 0·05
2·12 0·45
14·33 1·64
1·08 0·18
71·89 1·95
6·28 0·82
11·20 1·44
11·37 1·44
0·74 0·04
12·7 1·6
56·1 6·9
122·0 14·0
50·3 5·7
15·52 2·00
67·1 9·4
16·6 3·7
73 3
80 11
526 53
LA
0·12 0·14
0·37 0·59
4·88 01·08
3·99 0·69
1·17 1·10
0·43 0·19
2·98 0·77
15·30 1·44
1·03 0·20
69·76 2·10
xen (MS)
Av. glass
xen (MI)
Whakapapa
104X
12
Av. glass
Whakapapa
R97/92X1
11
dacite
Taupo
P1106.4
14
0·84
2·36 0·09
3·69 0·49
3·85 0·79
0·95 0·10
4·0 1·0
18·2 6·9
44·7 1·3
20·3 1·3
3·3 0·2
12·4 0·3
18·1 0·6
202 1
128 2
510 6
0·16 0·01
2·97 0·05
3·35 0·03
4·84 0·12
3·42 0·08
0·08
4·65 0·04
1·13
4·1
19·3
41·9
17·9
2·0
7·40
10·0
242
67
474
0·11
1·94
3·68
4·44
1·73
0·11
5·48
14·86 0·13 15·72
0·76 0·02
64·30 0·24 65·94
dacite
Ruapehu
Average
13
(continued)
1·25
5·6
25·8
53·8
24·7
3·0
10·00
17·0
155
95
634
0·05
2·82
4·42
1·75
0·35
0·10
2·33
13·63
0·31
74·24
rhyolite
Taupo
P1174
15
PRICE et al.
ANDESITE PETROGENESIS, RUAPEHU VOLCANO
2
3
16·1
Sc
2170
MD
66
11
1·5
142
15·5
8·0
210
1
7
3
4
2
2
4
3
2
Whakapapa
3
1
0·705269 26
0·705292 45
MD
85
58·5
206
29·5
4·6
164
26·4
2·83
Gms
9
3
2
4
0·705722 54
0·705709 31
MD
73
6
1·5
88
19·0
7·0
191
31·7
3·14
Av. Gms
Waihianoa
R96/6
10
3
1
0·705102 15
0·705188 30
MD
64
34
16·5
208
18·2
8·1
3·4
124
22·8
2·40
Gms
Waihianoa
R95/28
11
12
2
16
0·706234 16
50
9
0·4
79
9·4
9·6
234
29·8
9
13
0·707225 14
30 6
0·3 0·5
3·0 8·2
28 8
8·0 1·0
26·9 3·6
334 40
64·8 8·5
5·71 0·75
Av. glass
Av. glass
2·74
xen (MS)
Whakapapa
104X
xen (MI)
Whakapapa
R97/92X1
13
14
2·5
dacite
Taupo
P1106.4
4
4
5
0·70537 5
51 1
42 1
111 6
127 3
18·0 0·5
8·5 0·2
5·4 1·6
233 3
0·70597 1
71
6
11
97
15
8·2
4·8
188
24·5 0·1·0 25·7
2·28 0·09
dacite
Ruapehu
Average
15
0·70601 1
71
5
6
6
7
9·5
5·7
217
34·1
3·3
rhyolite
Taupo
P1174
Fm/date, Flow Formation or date of eruption; Av, average; GI, glass inclusion; Gls, glass; Px, pyroxene; Pl, plagioclase; Opx, orthopyroxene; Gms, groundmass; xen (MI),
meta-igneous xenolith; xen (MS), meta-sedimentary xenolith. Columns 14 and 15 are Taupo whole-rock samples (Sutton et al., 1995). Column 13 is average Mangawhero dacite
(this work). Columns 1, 3, 6, 7, 9, 11, 12 include data from Price et al. (2005). All major element data normalized to 100% on volatile-free basis. indicates 1s. Standard deviation
has not been calculated for trace elements where sample numbers are less than four. LA, laser ablasion ICP-MS analysis; MD, micro-drilled sample analysed by conventional
ICP-MS and multi-collector mass spectrometry.
Sr isotopes
14
Traces
8
R96/27
NUMBER 10
Majors
Analyses
0·705310 37
12
18·4
203
18·5
4·9
134
2·80
29·6
0·705316 22
2
1·86
20·6
groundmass
6
Whakapapa
Whole-rock
Micro-drilled
Sr/86Sr
4
10·0 7·7
49
7·7
39·1 6·8
208
8·0
520 147
7
R97/82
VOLUME 53
87
8
2·78
30·3
25·4 9·8
19·9 7·4
710 282
Whakapapa 1996 July
42
3·2
217
14·9
14·8
10·5
396
6·46 2·32
69·8 21·7
6
P57536
Av. GI-Opx Av. Gms Gls Av. Gms
Zn
6·6 3·0
221 27
17·6 2·7
5·8 0·8
4·2 0·6
158 21
4·03
43·4
Waihianoa
GI-Plag
5
R96/82
1
3·3
2·10 0·25
22·2 3·0
GI-Plag
4
R95/28
Ni
Cr
193
5·1
Nb
V
3·9
Hf
142
Y
Zr
1·84
20·3
Yb
description:
GI-Plag
Whakapapa Whakapapa Waihianoa
GI-Plag
R96/6
Sample
R96/27
Fm/date:
Sample no.:R95/82
1
Table 4: Continued
JOURNAL OF PETROLOGY
OCTOBER 2012
PRICE et al.
ANDESITE PETROGENESIS, RUAPEHU VOLCANO
Table 5: Major and trace element data for Ruapehu plagioclase and pyroxene
EPMA data
SiO2
Plag
Plag
Plag
Plag
Cpx
Opx
Opx
Opx
Opx
96/6
97/82
97/92X1
97/104X
R96/6
R96/6
97/82
R97/92X1
R97/104X
1s
n ¼ 26
n¼3
n ¼ 53
52·07
1s
2·95
n ¼ 48
51·34
1s
3·42
n ¼ 53
1s
50·43
2·63
n¼5
55·98
1s
n ¼ 30
0·95
TiO2
Al2O3
1s
n ¼ 53
1s
n ¼ 23
51·36
0·93
0·56
0·16
52·87
0·72
53·29
0·54
52·93
0·96
51·80
0·24
0·10
0·26
0·13
0·35
0·09
0·29
1s
29·32
2·14
29·68
2·47
30·63
1·82
28·22
0·43
2·46
0·63
1·59
1·03
1·35
0·87
2·43
1·83
4·00
0·61
0·21
0·70
0·26
0·62
0·17
0·41
0·09
10·63
1·50
19·71
2·40
18·55
1·38
18·24
1·51
16·88
MnO
0·25
0·07
0·46
0·12
0·42
0·07
0·44
0·11
0·31
MgO
14·47
0·72
22·68
2·27
23·78
0·99
23·60
2·44
25·81
0·26
FeO
CaO
12·90
2·10
13·39
2·29
14·21
1·97
9·89
0·54
19·10
1·10
1·84
2·38
1·59
0·35
1·52
0·52
Na2O
3·99
1·13
3·48
0·98
3·27
1·11
5·86
0·32
0·36
0·16
0·29
0·51
0·22
0·02
0·18
0·72
0·26
0·24
0·31
0·33
0·15
0·15
0·40
0·07
0·05
0·03
0·11
0·17
K2O
ICP-MS data
Ba
Rb
Sr
n¼7
73
1·7
585
n¼6
11·3
1·4
31·3
120
2·2
617
n¼2
21
2·8
66
32
0·1
895
n¼2
807
3·2
766
n¼2
n¼1
n¼2
n¼4
4
50·1
50·1
0·6
0·8
50·1
50·1
50·1
1·4
50·4
50·4
2·0
1·7
50·01
20
Pb
2·76
0·60
3·29
1·05
3·18
12·20
0·33
50·01
50·01
0·1
Th
0·04
0·05
0·15
0·23
50·05
2·30
0·08
50·05
50·05
50·05
50·05
1·05
0·02
50·05
50·05
50·05
1·6
3·7
6·6
50·05
50·05
50·05
3·5
7·5
9·8
n¼2
16
1·75
U
50·05
Zr
2·3
1·9
3·2
5·0
8·3
4·2
Nb
0·11
0·12
0·20
0·23
50·05
1·19
Y
0·6
0·33
0·8
0·8
1·0
9·7
La
2·16
0·20
3·12
0·66
5·36
58·86
1·68
50·01
0·30
0·06
0·67
Ce
3·88
0·42
5·32
1·34
7·86
116·15
7·24
0·05
0·79
0·23
3·44
Nd
1·53
0·25
1·97
0·65
2·56
41·75
8·60
0·11
0·73
0·41
1·63
Sm
0·23
0·15
0·29
0·16
0·41
6·34
3·28
0·09
0·37
0·3
0·98
Eu
0·53
0·04
0·71
0·14
1·96
2·98
0·78
0·03
0·06
0·09
0·06
Gd
0·16
0·09
0·25
0·11
0·33
4·07
4·18
0·22
0·46
0·66
2·09
Dy
0·17
0·09
0·22
0·14
0·19
2·24
4·61
0·46
1·08
1·36
5·71
Er
0·06
0·05
0·06
0·08
0·09
0·75
2·65
0·46
0·85
1·26
6·10
Yb
0·03
0·04
0·05
0·08
0·05
0·39
2·39
0·7
1·42
1·66
Sc
0·7
0·4
0·5
0·3
0·3
0·5
119·0
41·8
47·9
V
3·8
3·4
3·7
3·9
1·1
0·6
430·4
117·7
158·2
348·5
116·1
Cr
1·4
3·0
50·4
0·5
319·7
12·8
77·4
608·0
111·9
Ni
1·3
1·5
0·3
0·2
0·3
0·3
42·8
29·9
170·5
262·7
46·6
Zn
13·1
6·9
8·1
2·4
3·3
2·3
99·8
248·5
360·9
265·2
228·0
0·06
29·0
0·08
25·0
61·76
0·85
39·5
0·89
45·0
9·18
100·3
EPMA, electron probe micro-analysis; ICP-MS, laser ablation inductively coupled plasma source spectrometry; Plag,
plagioclase; Cpx, clinopyroxene; Opx, orthopyroxene; samples R96/6 and R97/82 are from the Wahianoa and
Whakapapa Formations respectively; R97/92X1 and R97/104X are meta-igneous and meta-sedimentary xenoliths
respectively.
Like the Wahianoa Formation, the Whakapapa
Formation is a collection of sub-units, each consisting of
stratigraphically and spatially associated lava flows and
each showing subtle differences in major and trace element
and isotopic composition. They are interpreted to be representative of single magma batches, which evolved
separately within distinct storage reservoirs. This is clearly
reflected in the very wide range of magma compositions
that has been sampled by historical eruptions (shaded
field in Fig. 16). All Whakapapa lavas erupted over the
past 65 years are porphyritic andesites or dacites but their
geochemistry and petrology have fluctuated widely over
2171
JOURNAL OF PETROLOGY
VOLUME 53
NUMBER 10
OCTOBER 2012
short time intervals (Gamble et al., 1999; Nakagawa et al.,
1999). They have major and trace element and Sr, Nd and
Pb isotopic compositions that cover almost the full range
observed in the prehistoric lavas and they include samples
with higher SiO2 abundances.
M E LT I N C L U S I O N , C RY S TA L A N D
G RO U N D M A S S T R AC E E L E M E N T
A N D I S O T O P E DATA
The major and trace element compositions of glassy
groundmass and glass inclusions in minerals within
Ruapehu andesites are broadly similar to those of TVZ
rhyolites and dacites (Price et al., 2005) and to Ruapehu
dacite (Table 4). LA-ICP-MS analyses for glass and
groundmass and mineral phases in andesites and ICP-MS
data for micro-drilled andesite samples and mineral trace
element compositions in xenoliths are presented in Tables
4 and 5, and glass and groundmass trace element compositions are compared in Table 4 with an average Ruapehu
dacite composition and TVZ rhyolite and dacite. In an
examination of mineral-scale isotopic relationships,
87
Sr/86Sr was measured in micro-drilled plagioclase and
groundmass in four andesite samples; these data were augmented by LA-MC-ICP-MS 87Sr/86Sr analyses for plagioclases from two of these andesites (Fig. 17). The results
shown in Fig. 17a indicate that 87Sr/86Sr ratios in the
groundmass are similar to, or slightly higher than, the
whole-rock 87Sr/86Sr. In each sample, the micro-drilled
plagioclase phenocrysts show a spread of 87Sr/86Sr ratios
around the whole-rock and/or associated groundmass composition. In three of the four samples, the bulk-rock composition lies within the isotopic compositional range for
micro-drilled plagioclase. Bulk-rock and plagioclase isotopic values are roughly correlated, which is expected as
plagioclase is the dominant phenocryst phase in each of
the samples (52^95% of phenocrysts) and it contains concentrations of Sr that are generally higher than in the
bulk-rock (e.g. 267 ppm for whole-rock sample R96/6 and
585 ppm in the plagioclase phenocrysts it contains;
273 ppm in R97/82 and 617 ppm in the plagioclase).
Among micro-drilled plagioclase the largest variations in
Sr isotopic compositions were observed in aType 1 andesite
from the Wahianoa Formation (R96/6) whereas the narrowest range occurs in a Type 1 andesite from the
Whakapapa Formation (R97/82).
Laser ablation microsampling (at the 50·2 mm scale) of
plagioclase in a Wahianoa and a Whakapapa andesite confirms the existence of resolvable Sr isotopic inhomogeneity
within each of the plagioclase phenocryst populations
(Fig. 17b and c). Single crystals show significant variability
in Sr isotopic composition, well outside experimental
error. For example, crystal 10 (Fig. 17b) in sample R96/6
shows a range in 87Sr/86Sr from 0·70529 7 to 0·70562 8,
Fig. 17. (a) 87Sr/86Sr for micro-drilled samples of plagioclase (open
circles) and groundmass (open squares) compared with whole-rock
data for four Ruapehu andesites: two from the Whakapapa
Formation and two from the Wahianoa Formation. For each of the
four andesite samples, three separate plagioclase crystals (labelled
1^3) and groundmass material were drilled from a polished slab
of rock and the collected powders analysed by inductively coupled,
plasma source, multi-collector mass spectrometry (ICP-MC-MS)
after conventional ion column chemical separation of Sr. The wholerock data were obtained by thermal ionization mass spectrometry following conventional ion column chemical separation of Sr. Vertical
dashed lines mark whole-rock compositions. (b) Comparison of
87
Sr/86Sr micro-drill data [data shown in shaded areas summarize information in (a)] with laser ablation (LA)-ICP-MC-MS data for
plagioclase in two Ruapehu andesites. Data for single crystals are
separated by dashed lines and alphanumeric codes on the y-axes each
indicate crystal and analysis number. ‘s’, sieve-textured crystal;
‘Gmass’, groundmass; ‘m’, microphenocryst; ‘d’, duplicate analysis of a
drilled powder. Filled circles are plagioclase phenocryst cores and
open circles are phenocryst rims or groundmass grains. Vertical lines
provide a reference to the whole-rock compositions.
and in crystal 7 in sample R97/82 the ratio varies from
0·70516 8 to 0·70568 13 (Fig. 17b). Some of the highest
87
Sr/86Sr ratios are obtained for sieve-textured crystals
(e.g. crystal 10a in sample R96/6 and crystals 4 and 11 in
2172
PRICE et al.
ANDESITE PETROGENESIS, RUAPEHU VOLCANO
R97/82; Fig. 17b), which are common in Ruapehu andesites
and in the margins of micro-xenoliths.
DISCUSSION
Ruapehu eruptive rocks are dominantly porphyritic andesites with trace element abundance patterns typical of
subduction-related magmas. LREE, Ba, Sr, Rb, Th and K
are all enriched relative to N-MORB, HREE, Ti, Zr and
Y abundances are similar to or more depleted than
N-MORB, Nb is depleted relative to K, and Pb and Sr
are enriched relative to Ce (Figs 12 and 13). The scarcity
of amphibole and estimates of water content obtained
from plagioclase^melt hygrometry indicate that the
magmas were relatively dry (54 wt % H2O) and fO2 conditions are estimated to have been slightly above the
NNO buffer.
The minor and trace element characteristics that typify
the Ruapehu volcanic eruptive rocks have commonly been
interpreted to have an ultimate origin in the mantle
wedge, immediately above the subducting slab (Arculus
& Powell, 1986; Tatsumi et al., 1986; McCulloch &
Gamble, 1991; Hawkesworth et al., 1993; Brenan et al., 1995;
Elliott et al., 1997; Kessel et al., 2005). Primary,
subduction-related, basaltic magmas are argued to be generated where slab fluids or melts interact with depleted
mantle, but basalts with compositions consistent with
direct derivation from peridotitic mantle are rare in
subduction-related volcanic arcs and most primary
magmas have physical properties such as crystallinity,
density, rheology and viscosity that mean they are very
likely to have been trapped and modified within the crust
(e.g. Marsh, 1981; Kent et al., 2010). Basalts in the TVZ
show typical subduction-related trace element patterns
similar to those observed in basalts from the intra-oceanic
Kermadec arc to the north of New Zealand. Titanium, Zr,
Nb and Ta are low in abundance (relative to MORB) and
Ba, Rb, K and LREE are enriched (Gamble et al., 1993a).
TVZ basalts tend to have higher Zr abundances than
their counterparts in the Kermadec arc and the Sr and
Nd isotopic compositions observed in TVZ basalts vary
significantly from the MORB-like values of the Kermadec
arc to relatively radiogenic ratios that have been interpreted as either a subducted sediment or crustal contamination signal (Gamble et al., 1993a, 1996). It would seem
probable that mantle-derived magma feeding into the continental lithosphere beneath the TVZ has an isotopic and
trace element composition that is generally similar to
those generated beneath the southern Kermadec arc.
The most significant difference between the oceanic and
continental segments of the arc is the lithosphere and particularly the crust. In North Island New Zealand, the shallow crust is represented by Mesozoic basement greywacke
of the Waipapa and Torlesse terranes (Electronic
Appendix A1) and by meta-sedimentary xenoliths that
occur within the lavas. The nature of the deeper crust is
more difficult to resolve. The available geophysical data
do not allow precise estimates of the thickness of either
the greywacke basement or the substrate on which these
were deposited. Meta-igneous xenoliths found in North
Island andesites have been mineralogically equilibrated at
pressures of 0·7^1GPa, suggesting depths of at least 20 km
(Graham et al., 1990), and on the basis of their major and
trace element chemistry and isotopic compositions they
have been argued to represent samples of the oceanic
lower crust on which the greywackes of the North Island
basement were deposited (Graham et al., 1990; Price et al.,
2005). This interpretation is consistent with crustal structure profiles developed from seismic velocity information;
for example, Stern et al. (2010) showed profiles with a crustal thickness of 20 km with data obtained from deeper
levels interpreted to indicate basaltic underplating to
25 km. Meta-igneous xenoliths may be representative of
the lower crust beneath Ruapehu, but most have chemical
and mineralogical compositions indicating that they are
restites derived from melting of the original protoliths
(Graham et al., 1990; Price et al., 2005, 2010). Consequently,
the composition of the original deep crustal protolith and
more particularly the material involved in lower crustal
processes affecting the generation and evolution of andesite
magma can only be inferred from xenolith bulk-rock and
mineral compositions.
Geochemical variation in the Ruapehu
andesite suite: the role of open- and
closed-system fractional crystallization
The long-term variability in the chemistry of Ruapehu
magmas has been argued to reflect increased fractionation
and crustal involvement over the long history of the magmatic system beneath the volcano (e.g. Price et al., 2005).
This variability, although complex in detail, does show certain regularities. K2O concentrations are lowest in the Te
Herenga Formation. K2O content is higher in the younger
formations, peaking in the Mangawhero and Whakapapa
Formations (Fig. 10g; also Price et al., 2005, 2007). Overall
compositional variability also increases, being lowest in
the Te Herenga lavas and greatest in the Mangawhero and
Whakapapa Formations, which have the widest ranges in
SiO2, MgO and K2O abundances and 87Sr/86Sr values.
REE patterns also vary with time. Te Herenga Formation
lavas have relatively flat REE patterns with La/Yb being
markedly lower than in all other Ruapehu lavas except
the Ruapehu basalt and one other Mangawhero
Formation sample (Fig. 13). La/Yb ratios are higher in the
Wahianoa Formation and peak in the Mangawhero and
Whakapapa Formations and Wahianoa Formations. Te
Herenga andesites are also texturally different from those
of the younger formations. They tend to be more strongly
porphyritic (average phenocryst content is 48%
2173
JOURNAL OF PETROLOGY
VOLUME 53
compared with averages of 35^44% for the younger formations) and they have crystalline rather than glassy
groundmass.
The Ruapehu sample suite as a whole shows scattered
positive correlations between SiO2 and Na2O, K2O, Ba
and Rb contents and 87Sr/86Sr, and negative correlations
between SiO2 and Al2O3, FeO*, MgO, CaO, Sc, V, Cr
and Ni abundances. These patterns of variation have previously been interpreted (e.g. Graham & Hackett, 1987;
Graham et al., 1995) to reflect the influence of fractional
crystallization or combined assimilation^fractional crystallization (AFC) and some of the variation among the
post-Te Herenga andesites can be quantitatively modeled
by these processes; as an example, the trace element and
isotopic composition of high-SiO2 andesite R97/19 can be
approximated by an AFC model in which high-MgO andesite (R95/20) is the parental magma and Torlesse
meta-sediment the assimilant (Table 6 and Figs 18 and 19).
The composition of sample R95/20 can in turn be modeled
by AFC, using a primitive intra-oceanic basalt (Kermadec
basalt 7135; Smith et al., 2010) as the parental magma
and Torlesse meta-sediment as the assimilant (Table 6 and
Fig. 19).
There is also evidence that some of the more fractionated Ruapehu samples, such as the Mangawhero dacites,
may have derived from andesites by essentially
closed-system fractional crystallization. Close compositional similarities exist between these dacites and andesite
groundmass and groundmass glass (Table 4 and Fig. 13d),
and on a 87Sr/86Sr vs Zr plot (Fig. 18c) the dacites (with
200 ppm Zr) lie close to extrapolations of tie lines connecting groundmass and host whole-rock compositions,
indicating that closed-system fractionation or mechanical
separation of the phenocryst assemblage from a post-Te
Herenga andesite would drive residual melt compositions
towards dacite.
Type 2 plagioclase andesites of Unit C in the
Wahianoa Formation (Figs 3c, d and 15) appear to be rare
examples of fractionated andesites in which plagioclase
has accumulated. Compared with other Wahianoa lavas,
this andesite type is characterized by slightly elevated
modal abundances of plagioclase and lower modal abundances of pyroxene and by relatively higher abundances of
Al2O3, CaO and Sr, less well-developed Eu anomalies in
chondrite-normalized REE patterns and lower MgO
abundances. All of these characteristics are consistent
with a small degree of pyroxene-dominated fractional
crystallization and plagioclase accumulation (see Graham
& Hackett, 1987). For example, least-squares mixing modeling involving a parental magma with the composition of
a Type 1 Wahianoa andesite from Unit A (R96/22), 11%
plagioclase addition and removal of 6% orthopyroxene,
3% clinopyroxene and 1% magnetite gives a good fit for
major elements (sum of squares of residuals ¼ 0·024) to
NUMBER 10
OCTOBER 2012
Type 2 andesite R95/28 from Unit C of the Wahianoa
Formation.
Open- or closed-system fractional crystallization can
explain some aspects of the geochemical relationships
among the Ruapehu andesites and dacites but for many
Ruapehu andesites these types of processes do not provide completely satisfactory models. This is because
most of the andesites do not appear to represent melts
on simple lines of liquid descent; there is clear evidence
that they are complex mixtures of melts and crystals.
There is also considerable uncertainty with respect to
identifying appropriate parental magma compositions to
use in fractional crystallization and AFC models. For example, the Ruapehu basalt is the least evolved composition in the sample suite but it has a Nd^Sr isotopic
composition such that if it is used as a parental magma
in AFC models, with either meta-igneous xenolith or
Torlesse basement as the assimilant, it cannot reproduce
the isotopic composition of any of the post-Te Herenga
andesites (Fig. 18). Crustal assimilation and fractional
crystallization have clearly had a general influence on
geochemical variation at Ruapehu but for most andesites
they represent only part of a complex petrogenetic
history.
Magma supply, magma storage and
magmatic plumbing at Ruapehu
Collectively major and trace element and isotopic behaviour for the whole Ruapehu sample suite suggest a role for
assimilation and fractional crystallization but the
small-scale (within formation) compositional variations
are not consistent with a single line of liquid descent, or
processes taking place within a single magma chamber.
The nature of the magma storage and plumbing system is
reflected in the variation observed in the Wahianoa flow sequence in the Whangaehu and Wahianoa valleys (Fig. 15)
and by the differences in geochemical variation observed
in the lava flow groups of the Whakapapa Formation
(Fig. 16), including the very youngest eruptive rocks (AD
1945^1996).
From base to top, the complete Wahianoa flow sequence
exposed on eastern Ruapehu (Fig. 4) does not show systematic temporal trends in geochemical behavior that might
indicate progressive tapping of a single, evolving magma
chamber. Within the sequence, groups of two to four stratigraphically related lava flows, which appear to have
erupted sequentially without major time breaks, do show
some systematic variations and each of these could indicate
sequential egress of magma from a specific, small-volume
(50·5 km3), heterogeneous magma reservoir. In these
related flow groups, the earliest flows generally have the
highest SiO2 and K2O and the lowest MgO abundances
(Fig. 15), which could indicate eruption triggered by an
influx of new hotter, more mafic magma. These variations
are not in all cases correlated with changes in 87Sr/86Sr,
2174
PRICE et al.
ANDESITE PETROGENESIS, RUAPEHU VOLCANO
Table 6: AFC and mixing models for trace elements and isotopic ratios in Ruapehu lavas
a, AFC model: post-Te Herenga high-SiO2 andesite
r ¼ 0·4; F ¼ 0·8
Cpx ¼ 0·53, Opx ¼ 0·17, Pl ¼ 0·30
R95/20
K2O
TORL
0·96
Model
2·88
R97/19
1·66
1·61
Rb
28
120
55
81
Sr
198
302
223
268
Zr
82
206
132
148
La
7·4
34·6
14·9
16·3
Ce
17·7
66·7
32·8
34·9
Nd
10·9
30·8
18·2
16·3
Sm
2·7
5·9
4·1
3·7
Eu
0·81
1·02
1·03
0·97
Gd
2·83
4·92
4·03
3·68
Yb
1·82
2·68
2·52
1·91
87/86Sr
0·70497
0·70986
0·70592
0·70592
143/144Nd
0·51281
0·51245
0·51271
0·51265
b, AFC model: Post-Te Herenga, high-MgO andesite
c, Mixing model: Post-Te Herenga, high-MgO andesite
r¼ 0·2; F ¼ 0·57
Ol ¼ 0·03, Cpx ¼ 0·33, Opx ¼ 0·17, Pl ¼ 0·42, Mt ¼ 0·05
7135
TORL
2·88
Model
R95/20
X¼
46358
0·16
0·84
0·48
R95/20
K2O
0·11
4
120
29
28
Rb
120
7
25
28
Sr
152
302
155
198
Sr
302
187
205
198
Zr
13
206
59
82
Zr
206
50
74
82
0·96
4·47
Model
Rb
0·72
K2O
TORL
1·11
0·96
La
1·3
34·6
8·5
7·4
La
34·6
2·7
7·7
7·4
Ce
3·9
66·7
18·7
17·7
Ce
66·7
7·5
16·9
17·7
Nd
3·3
30·8
11·1
10·9
Nd
30·8
7·0
10·7
10·9
Sm
1·3
5·9
3·1
2·7
Sm
5·9
2·2
2·8
2·7
Eu
0·48
1·02
0·81
0·81
Eu
1·02
0·82
0·85
0·81
Gd
1·51
4·92
3·22
2·83
Gd
4·92
2·86
3·19
2·83
Yb
1·24
2·68
2·41
1·82
Yb
2·68
2·53
2·55
1·82
87/86Sr
0·70342
0·70986
0·70497
0·70497
87/86Sr
0·70985
0·70349
0·70497
0·70497
143/144Nd
0·51306
0·51245
0·51275
0·51281
143/144Nd
0·51244
0·51305
0·51277
0·51280
(continued)
2175
JOURNAL OF PETROLOGY
VOLUME 53
NUMBER 10
OCTOBER 2012
Table 6: Continued
d, AFC model: Ruapehu basalt
e, Mixing model: Ruapehu basalt
r ¼ 0·4; F ¼ 0·69
R97/92X1
Ol ¼ 0·03, Cpx ¼ 0·34, Opx ¼ 0·15, Pl ¼ 0·42, Mt ¼ 0·06
Melt
Pl
Opx
MIX A
7135
0·6
0·25
0·15
0·35
0·65
7135
Ave MIX
Model
0·35
R04/04
K2O
0·11
Rb
2
13
7
12
Sr
152
378
186
201
Zr
13
84
42
57
0·26
X ¼
K2O
0·57
4·47
0·11
R04/04
0·10
0
139
0
0
83
4
32
12
Sr
48
895
2
253
152
187
201
Zr
234
8
7
143
13
59
57
Rb
2·71
Model
1·02
0·57
La
1·3
8·0
4·2
5·0
La
22·4
5·4
0·1
14·8
1·3
6·0
5·0
Ce
3·9
18·5
10·8
10·4
Ce
49·4
7·9
0·2
31·6
3·8
13·5
10·4
Nd
3·3
8·3
7·0
7·7
Nd
24·0
2·6
0·4
15·1
3·3
7·4
7·7
Sm
1·3
2·3
2·4
2·0
Sm
5·3
0·4
0·3
3·3
1·3
2·0
2·0
Eu
0·48
1·22
0·87
0·75
Eu
0·59
1·96
0·10
0·86
0·50
0·63
0·75
Gd
1·51
3·99
3·10
2·27
Gd
5·21
0·33
0·66
3·31
1·50
2·13
2·27
Yb
1·24
2·62
2·37
1·55
Yb
2·74
0·05
1·66
1·91
1·20
1·45
1·55
87/86Sr
0·70342
0·70655
0·70475
0·70475
87/86Sr
0·70623
0·70623 0·70623
0·70623
0·70342
0·70475
0·70475
143/144Nd
0·51306
0·51285
0·51299
0·51293
143/144Nd
0·51274
0·51274 0·51274
0·51274
0·51306
0·51283
0·51293
f, AFC model: Te Herenga andesite
g, Mixing model: Te Herenga andesite
r ¼ 0·2; F ¼ 0·34
R97/92X1
Ol ¼ 0·13, Cpx ¼ 0·41, Pl ¼ 0·39, Mt ¼ 0·07
Melt
Pl
Opx
MIX B
46358
0·25
0·5
0·25
0·28
0·72
7135
Ave MIX
0·35
Model
T6/7
X¼
K2O
0·11
Rb
4
13
12
14
Rb
Sr
152
378
187
206
Zr
13
84
70
53
0·48
0·67
K2O
4·47
2·71
0·48
Model
T6/7
0·10
0
139
0
0
83
7
15
14
Sr
48
895
2
253
187
263
206
Zr
234
8
7
143
50
54
53
0·67
0·68
La
1·3
8·0
7·2
5·0
La
22·4
5·4
0·1
14·8
2·7
4·2
5·0
Ce
3·9
18·5
18·9
11·5
Ce
49·4
7·9
0·2
31·6
7·5
10·0
11·5
Nd
3·3
8·3
12·3
7·9
Nd
24·0
2·6
0·4
15·1
7·0
7·1
7·9
Sm
1·3
2·3
4·2
2·4
Sm
5·3
0·4
0·3
3·3
2·2
2·0
2·4
Eu
0·48
1·22
1·27
0·81
Eu
0·59
1·96
0·10
0·86
0·82
0·91
0·81
Gd
1·51
3·99
5·05
2·79
Gd
5·21
0·33
0·66
3·31
2·86
2·52
2·79
Yb
1·24
2·62
3·91
1·82
Yb
2·74
0·05
1·66
1·91
2·53
2·14
1·82
87/86Sr
0·70342
0·70655
0·70483
0·70483
87/86Sr
0·70623
0·70623 0·70623
0·70623
0·70349
0·70483
0·70483
143/144Nd
0·51306
0·51285
0·51300
0·51290
143/144Nd
0·51274
0·51274 0·51274
0·51274
0·51305
0·51296
0·51290
AFC models: AFC is assimilation–fractional crystallization (De Paolo, 1981); F, fraction of liquid remaining; r, ratio of
assimilation to fractional crystallization; Ol, olivine; Cpx, clinopyroxene; Opx, orthopyroxene; Pl, plagioclase; Mt, magnetite; partition coefficient data used in AFC models are from Dunn & Sen (1984), Ewart & Hawkesworth (1987), Halliday
et al. (1995) and Blundy & Wood (2003); 7135 is a primitive Kermadec basalt.
Mixing models: X, weight fraction; R97/92X1 is a meta-igneous xenolith; MIX is the mixture of Melt (Glass) þ Pl þ Opx;
Ave MIX is an average meta-igneous xenolith composition; 7135 is a primitive Kermadec basalt; 46358 is a Kermadec
andesite; TORL is average Torlesse. Model and real compositions with which they are compared are shown in italics at
the right of each table.
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PRICE et al.
ANDESITE PETROGENESIS, RUAPEHU VOLCANO
Fig. 18. (a) 143Nd/144Nd vs 87Sr/86Sr and (c) 87Sr/86Sr vs Zr showing comparison between data for Ruapehu volcanic rocks and assimilation^
fractional crystallization models (AFC) involving a parental intra-oceanic arc basalt (7135) or high-MgO Ruapehu andesite (R95/20) and
Torlesse basement meta-greywacke or meta-igneous xenolith (MIX) assimilant compositions. r, ratio of material assimilated to material crystallized (DePaolo, 1981). For all the AFC model curves F, the fraction of melt remaining, is marked in 20% steps. RB, Ruapehu basalt (R04/04).
(b) and (d) 143Nd/144Nd vs 87Sr/86Sr and 87Sr/86Sr vs Zr showing the comparison between data for Ruapehu volcanic rocks and mixing
models involving intra-oceanic arc basalt (7135) or andesite (46358) and meta-igneous xenolith melt þ plagioclase þ orthopyroxene (in MIX A
the proportions of melt:plagioclase:orthopyroxene are 0·6:0·25:0·15; in MIX B the proportions are 0·25:0·5:0·25) or Torlesse meta-greywacke
(Table 6). Mixing trajectories are marked off in steps of x (proportion of MIX) ¼ 0·1. Models are explained in detail in the text and in Table 6.
Intra-oceanic arc basalt and andesite data are from Smith et al. (2010). Continuous-line, near horizontal, arrows in (c) connect analysed groundmass or groundmass glass (Gmss) and host whole-rock compositions. Shaded areas in (c) and (d) indicate the range of dacite compositions.
which could indicate that each of the mafic and felsic components in each related flow group has a subtly different
evolutionary history. Unit C stands out as a distinctive
magma batch erupted during a specific time interval.
Similarly, although each of the various flow sequences of
the prehistoric Whakapapa Formation manifests some
degree of systematic geochemical variation, collectively
they show very scattered distributions on variation diagrams (Fig. 16). The range in SiO2, MgO and K2O abundances and 87Sr/86Sr within the very youngest eruptive
rocks (AD 1945^1996) is almost as broad as that for the
whole of the Whakapapa Formation and shows no consistent variation with time, even over time scales of days to
months.
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NUMBER 10
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Fig. 19. Normalized extended trace element plots comparing Ruapehu basalt and andesite data with compositions derived from assimilation^
fractional crystallization (AFC) and mixing models. The models are explained in detail in the text and in Table 6. Intra-oceanic arc basalt
(7135) and andesite (46358) data are from Smith et al. (2010). Shaded field labeled ‘TH’ is the complete range of analysed compositions for the
Te Herenga Formation. In AFC models, F is the fraction of melt remaining and r is the ratio of material assimilated to material crystallized
(DePaolo, 1981). Analysed whole-rock compositions are shown with filled symbols and model compositions with open symbols. (a) Model for
derivation of a high-SiO2 Whakapapa andesite (R97/19) from a high-MgO Wahianoa andesite (R95/20) by AFC involving Torlesse
meta-greywacke as the assimilant. (b) Model for derivation of a high-MgO Wahianoa andesite (R95/20) from an intra-oceanic arc basalt
(7135) by AFC involving Torlesse meta-greywacke as the assimilant or by mixing between Torlesse meta-greywacke and an intra-oceanic arc andesite (46358). (c) Model for the derivation of Ruapehu basalt (R04/04) from an intra-oceanic arc basalt (7135) by AFC involving average
meta-igneous xenolith (ave MIX) as the assimilant or by mixing between meta-igneous xenolith melt and minerals and an intra-oceanic arc
basalt (7135). In the mixing model, the meta-igneous component (MIX A) is a mixture of melt (glass), plagioclase (pl) and orthopyroxene
(opx) from meta-igneous xenolith R97/92X1 in the proportions 60:25:15. (d) Model for the derivation of Te Herenga andesites from an
intra-oceanic arc basalt (7135) by AFC involving average meta-igneous xenolith (ave MIX) as the assimilant or by mixing between
meta-igneous xenolith melts and minerals and an intra-oceanic arc andesite (46358). In the mixing model, the meta-igneous component
(MIX B) is a mixture of melt (glass), plagioclase and orthopyroxene from meta-igneous xenolith R97/92X1 in the proportions 25:50:25.
For both the Wahianoa and the Whakapapa Formations
the intra-formational geochemical variations are consistent
with magmatic evolution within a dispersed and complex
magma storage and plumbing system. Magma batches
have been tapped from single storage systems at particular
times and the geochemistry of each batch has been determined by the size, longevity and state of evolution of each
of these magma reservoirs at the time of eruption. The mobilization of particular reservoirs at specific times may be
related to the movement of magma from deeper in the
plumbing system, which would mean mixing of magmas
from different levels, adding an additional aspect of complexity to the geochemical variation.
The locus and geometry of the shallow (51km) subsurface magma supply system beneath Ruapehu can be
inferred from field (Hackett, 1985; Houghton et al., 1987)
and seismic information (Bryan & Sherburn, 1999;
Sherburn et al., 1999; Jolly et al., 2010). On Pinnacle Ridge
on NW Ruapehu, lava flows and pyroclastic deposits of
the Te Herenga Formation are cut by a plexus of andesitic
dikes and shallow intrusions that provide analogues for
the situation immediately beneath the present-day crater
lake. Results from the permanent seismic network and
short-term seismometer deployments have yielded useful
information on the degassing history and shallow (55 km)
movement of magma associated with recent eruptions but
rather limited information on the deeper structure of the
magma storage and plumbing system; magma volumes
associated with recent eruptions are small and the presence
of the crater lake complicates the interpretation of the
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PRICE et al.
ANDESITE PETROGENESIS, RUAPEHU VOLCANO
seismic data (Bryan & Sherburn, 1999; Sherburn et al.,
1999). Seismic anisotropy data obtained before, during
and after the 1995^1996 Ruapehu eruptions indicate that
magma immediately beneath the crater lake could be
stored and transported in small-volume (50·1km3) dykes
(Miller & Savage, 2001).
The long-term variation in andesite geochemistry at
Ruapehu is likely to reflect large-scale fluctuations in
magma flux from the mantle and consequent thermal and
compositional changes taking place in the crust as the
magmatic system has evolved. The shorter-term variability
reflects open-system replenishment, fractionation, crustal
assimilation, magma mingling and mixing taking place in
a dispersed, dynamic and vertically extensive system of
small magma chambers, dykes and sills (e.g. Gamble
et al., 1999; Hobden et al., 1999; Price et al., 2000, 2005, 2007).
The Ruapehu basalt
Basalts are extremely rare among Ruapehu eruptive rocks.
One Mangawhero-age basalt flow was identified by
Hackett (1985) and this has been re-collected and
re-analysed (R04/4 in Table 3). The Ruapehu basalt is
olivine-bearing, has an Mg# of 64·2, and Ni and Cr contents of 136 and 366 ppm respectivelyçfeatures that
could be taken to indicate that it is close to a primary
basalt composition. It is not strongly porphyritic (31%
phenocrysts) and it has the lowest 87Sr/86Sr and highest
143
Nd/144Nd observed in the whole Ruapehu suite (RB in
Fig. 14), similar to the most primitive Sr^Nd isotope ratios
in the Te Herenga suite and very different from other
Mangawhero-age eruptive rocks.
Mineral compositions in the Ruapehu basalt are not in
all cases those expected if the minerals were in equilibrium
with a liquid having either the whole-rock or glass inclusion composition (Fig. 8). The calculated Mg# for olivine
in equilibrium with a melt having the host-rock composition is 86 [method of Sisson & Grove (1993)], but analysed
olivines have Mg#s of 68^78. In contrast, plagioclase
compositions appear to approach those expected for equilibrium with melt inclusion glass compositions in the
Ruapehu basalt (Fig. 6). It would therefore seem likely
that the genesis of the Ruapehu basalt has involved
magma mingling and/or crystal fractionation with possibly
some crustal assimilation; it cannot be regarded as an unmodified near primary melt.
Relative to other TVZ basalts (e.g. the Kakuki basalt;
Gamble et al., 1993a) the Ruapehu basalt is depleted in Zr
and LREE and enriched in K, Ba and Rb. The isotopic
composition is distinctly different; both TVZ and
Kermadec basalts have lower 87Sr/86Sr at comparable
143
Nd/144Nd ratios. This distinctive isotopic composition
can be explained as a consequence of a contribution from
lower crust having an isotopic composition similar to the
meta-igneous xenoliths contained in central North Island
andesites (e.g. Price et al., 2005; Lee et al., 2008).
Two models for the origin of the Ruapehu basalt have
been tested. Both involve interaction of a mantle-derived,
primary basaltic composition with a crustal meta-igneous
protolith. An intra-oceanic arc basalt has been used as the
primary, mantle-derived magma (7135 from the
Kermadec Islands; Smith et al., 2010; see Table 3); TVZ basalts such as the Kakuki basalt (TVZ15 of Gamble et al.,
1993a; see Table 3) are significantly more enriched in some
incompatible trace elements (K2O 0·27 wt %, Zr 79 ppm
and La 5·7 ppm in TVZ15 compared with K2O 0·11wt %,
Zr 13 ppm and La 1·3 ppm in 7135). The crustal component
is assumed to have had a composition that is represented
among Ruapehu meta-igneous xenoliths and their minerals. In the first model an attempt has been made to produce the Ruapehu basalt composition by AFC, and in
the second the basalt composition has been derived
by mixing components of a meta-igneous xenolith with
an intra-oceanic arc basalt The AFC models are, in
mathematical terms, those of DePaolo (1981); in thermal
terms, bulk assimilation is assumed to be directly related
to crystallization of fractionating phases. The assimilant
is assumed to have the composition of an average metaigneous xenolith.
The theoretical basis for the crust^mantle magma
mixing models is that emplacement of mantle-derived
magmas causes partial melting of the deep crust and development of ‘crustal hot zones’ (Dufek & Bergantz, 2005;
Annen et al., 2006) in which crustal melts, restite and
magmas, derived by fractional crystallization from
mantle-derived parents, mix and mingle (Hildreth &
Moorbath, 1988; Price et al., 2005; Reubi & Blundy, 2009).
For both AFC and mixing models a reasonable fit can be
obtained for the 87Sr/86Sr and 143Nd/144Nd isotopic composition of the basalt and in each case several aspects of
the trace element compositions can also be matched
(Table 6 and Fig. 19). The AFC models have a significant
shortcoming in that, regardless of the ratio assumed
for assimilant to material crystallized [r of DePaolo
(1981)] the fractionation required to obtain a reasonable
approximation for major elements by least-squares mixing
is substantially higher than is needed to fit the trace elements and the isotopes. For example, at r ¼ 0·4 a reasonable
fit is obtained for major elements at F ¼ 0·4 but at this r
value the isotopes and trace elements are matched at
F ¼ 0·7.
It is likely that the Ruapehu basalt is a hybrid composition derived from a depleted primary magma by a combination of AFC (60^70% fractionation with r ¼ 0·1^0·2)
and mixing (30^40%) with melt and restite derived from
a meta-igneous lower crust.
The petrogenesis of Te Herenga Formation
andesites
Te Herenga andesites constitute a distinctive geochemical
group of Ruapehu eruptive rocks. They are characterized
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JOURNAL OF PETROLOGY
VOLUME 53
by relatively low K2O, Rb and Zr contents and REE patterns are less fractionated but they have SiO2 and MgO
abundances and low Ni and Cr contents that are similar
to those observed in many younger andesite lavas (Figs
10^13). Relative to younger eruptive rocks the Te Herenga
Formation has higher 143Nd/144Nd isotopic ratios and on
a Nd^Sr isotopic diagram (Fig. 14) samples from the formation form a separate linear array displaced to higher
87
Sr/86Sr values. Zr abundance is low and shows little
change, despite significant variation in SiO2 abundance
and isotopic composition. Modal compositions are relatively uniform but differ from those of most post-Te
Herenga lavas.
The offset away from the post-Te Herenga array of the Te
Herenga group, to higher 87Sr/86Sr at low 143Nd/144Nd,
has been argued to reflect the influence of a lower crustal
component with an isotopic composition similar to the
Ruapehu meta-igneous xenoliths (e.g. Price et al., 2005;
Lee et al., 2008). Despite this evidence for crustal input, Te
Herenga andesites have very low abundances of K2O, Zr
and REE compared with TVZ basalts, implying that both
the mantle-derived and crustal components involved in
their formation must have been extremely depleted in
these elements. Intra-oceanic arc basalt and andesite from
Raoul Island in the Kermadec group appear to be some
of the very few eruptive rocks in the Tonga^Kermadec^
New Zealand arc that have suitably depleted compositions
and these have therefore been used to model the generation
and evolution of Te Herenga compositions.
All attempts to model by AFC processes the generation
of Te Herenga compositions from either a Ruapehu basalt
or primitive TVZ or Kermadec basalts have produced unsatisfactory results (Table 6). Many of the Te Herenga
lavas have incompatible trace element abundances that
are similar to or even more depleted than those observed
in the Ruapehu basalt and cannot therefore be derived
from it by open- or closed-system fractionational crystallization. AFC models involving a primitive, intra-oceanic
(Kermadec) basalt, high degrees of fractionation (50^
70%) and significant amounts of assimilation can simulate
the major element compositions and 87Sr/86Sr isotopic
ratio of the Te Herenga lavas but they result in Rb, Zr
and REE abundances that are too high (Table 6 and Fig.
19). Furthermore, although the 87Sr/86Sr compositions of
the Te Herenga lavas can be approximated by these
models, the model 143Nd/144Nd values are generally higher
than the actual compositions (Table 6 and Fig. 18). The alternative, crust^mantle mixing hypothesis for the origin
of Te Herenga andesites is based on evidence that most
Ruapehu andesites are complex mixtures of crystals and
melts from both mantle and crustal sources (see below).
The mixing models explored here involve simple mixing
between an intra-oceanic arc andesite (46358 from Raoul
island; Smith et al., 2010), representing an evolved mantle
NUMBER 10
OCTOBER 2012
component and melt þ plagioclase þ pyroxene from a
meta-igneous xenolith, to represent the crustal component.
The mixing model produces results that have several advantages over the AFC models: Sr and Nd isotopic compositions are more closely approximated (Fig. 18b); K2O,
Rb, Zr and REE abundances show a better fit (Table 6
and Fig. 18c); and the variation of 87Sr/86Sr relative to
Zr abundance obtained from the model is a close fit
to the actual variation observed in the Te Herenga suite
(Fig. 18d).
On the basis of comparison between mixing and AFC
models and bearing in mind petrographic evidence for
complex mixing it appears likely that andesites of the Te
Herenga formation formed through magma mixing in a
deep crustal hot zone where mantle-derived magma was
fractionating to form andesitic magma while contemporaneously mixing with melt and restitic crystals formed by
anatexis of meta-igneous lower crust. Te Herenga andesites
represent magmas that were mixtures of dacitic melts,
crustal restite and crystals derived from both mantle and
crustal sources.
The petrogenesis of post-Te Herenga
andesites
With the exception of the Ruapehu basalt and one other
Mangawhero-aged sample, all Sr^Nd isotope results for
post-Te Herenga lavas form a linear cluster that is separate
from and has a slightly different slope from the Te
Herenga data array (Fig. 14). The least evolved
(high-MgO) post-Te Herenga andesites have isotopic compositions that can be quantitatively approximated by AFC
involving a primitive intra-oceanic arc basalt and Torlesse
meta-sedimentary basement (see above), and the derivation of more evolved isotopic compositions (high-SiO2
andesites and dacites) from less evolved andesitic precursors can be modeled by closed- or open-system crystal fractionation (see above). However, these models do not
provide an exclusive or comprehensive explanation for the
complete range of geochemical variation observed among
the whole of the post-Te Herenga sample suite. Other features of the petrology of the post-Te Herenga andesites indicate that AFC is likely to be only one of a number of
processes involved.
The characteristic petrographic feature of andesitic volcanic rocks is a porphyritic texture and the presence of
complex phenocryst assemblages, manifesting both compositional and textural disequilibrium (e.g. Eichelberger,
1978; Gill, 1981). Ruapehu andesites are in every respect
typical of continental andesites. They are strongly porphyritic (35^55%) and the phenocryst assemblage is dominated by plagioclase, which has a very wide
compositional range across the whole sample suite, within
each formation and in single andesite samples (Fig. 5).
Within samples plagioclase crystals show a complex range
of textural types from unzoned through oscillatory zoned,
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PRICE et al.
ANDESITE PETROGENESIS, RUAPEHU VOLCANO
to normally and reversed zoned. Complex clusters or glomerocrysts are common and many crystals show distinct
resorption zones, whereas others are sieve-textured.
Plagioclase, pyroxene and olivine phenocrysts show major
element compositional variations indicating that the crystal cargo carried in each andesite equilibrated with melts
that were compositionally different from their present
host-rock and from the groundmass or groundmass glass
of the host (Figs 6 and 8). Within single andesite samples,
although there is a general similarity between the isotopic
composition of host-rock and plagioclase (Fig. 17), the
plagioclase phenocryst population in each rock shows significant variation (well outside experimental error) in Sr
isotopic composition and in many cases this variability is
present within discrete crystals (Fig. 17). The implication
is that the phenocryst population in each andesite sample
contains a mixture of crystals with different magmatic histories or crystals derived from different sources. Many
plagioclase phenocrysts have textures that closely resemble
those observed in plagioclase in crustal xenoliths and the
range of plagioclase compositions observed in the andesites
overlaps the compositional spectrum observed in the xenoliths (Fig. 5).
Similar results have been obtained from the nearby
Ngauruhoe Volcano, where small-scale isotopic sampling
has revealed considerable isotopic variability between coexisting plagioclase and clinopyroxenes and encompassing
groundmass, between crystals and whole-rock compositions, between coexisting plagioclase crystals of different
sizes (and therefore crystallization histories) and over relatively short time scales (c. 100 years) (Davidson et al., 2007).
Estimates of equilibration temperatures and pressures
obtained from mineral compositions provide an additional
indication of the complex petrogenetic history of Ruapehu
andesite phenocryst populations. Pyroxenes in post-Te
Herenga andesites indicate eruption temperatures between
950 and 11908C and equilibration pressures that may have
ranged from 1 to 0·2 GPa. These pressure estimates are consistent with those inferred from the mineral assemblages
in xenoliths (e.g. Graham et al., 1990, 0·7^1GPa for
meta-igneous xenoliths; Graham, 1987, 0·4^0·7 GPa for
meta-sedimentary xenoliths) and those obtained from
amphibole compositions. They imply that the crystal and
lithic cargoes contained in the post-Te Herenga andesites
represents material crystallized and equilibrated over a
25^30 km crustal section. This is consistent with geophysical interpretations of crustal structure in the central
North Island of New Zealand (e.g. Salmon et al., 2011).
The similarities between the compositions of melt inclusions and groundmass material in post-Te Herenga andesites and TVZ rhyolites and dacites (Table 4), the rarity of
aphyric andesites on Ruapehu and the absence of melt inclusions with andesite composition can be interpreted to
indicate that many Ruapehu andesitic magmas are in fact
crystal-rich dacitic or rhyolitic melts carrying lithic and
crystal fragments derived from both crustal and mantle
sources (see Price et al., 2005; Reubi & Blundy, 2009). Both
bulk andesites and their dacitic groundmass (see
micro-drill Sr isotope data, Fig. 17) have Sr isotope ratios
that are intermediate between those of primitive
intra-oceanic subduction-related volcanic rocks and North
Island crustal values and they overlap with those of TVZ
rhyolites; this implies that, as is the case for the crystal cargoes, the melt component of the andesites also has a complex origin involving both crustal and mantle
contributions (see Eichelberger, 1978).
It has been demonstrated that mixing of melts and
crystals from crustal and mantle sources is likely to have
been involved in the generation of the Ruapehu basalt
and Te Herenga andesites so it is also reasonable to
expect that such processes also had some role in the generation of post-Te Herenga magmas. Meta-igneous xenoliths, presumed to have been derived from the lower
crust, are abundant in most post-Te Herenga andesites,
so it can be assumed that this component continued to
contribute to geochemical variation throughout the magmatic history of the volcano. Compositions within the
less evolved part of the post-Te Herenga andesitic spectrum can be approximated by AFC models (see above)
but they can also be modeled (Table 6 and Fig. 19d)
as the outcome of a mixing between fractionating
intra-oceanic basalt or andesite magmas and at least two
crustal components: a lower crustal component represented by the meta-igneous xenoliths and a shallower
crustal component represented by the Torlesse metasedimentary basement.
The detailed geochemical studies of the Wahianoa and
Whakapapa Formations (see above) provide clear evidence
that each post-Te Herenga andesite magma batch, now
represented by single flows or by groups of flows, had a
unique and complex evolutionary history.
Each post-Te Herenga andesite is the outcome of a
unique blend of processes taking place over a wide range
of crustal and mantle conditions and affecting magmas
derived from both crustal and mantle sources
Implications for crustal evolution
Crustal thickness in the central TVZ is generally believed
to be 15^25 km (e.g. Stern et al., 2010; Salmon et al., 2011),
with the crust underplated and intruded by
mantle-derived magmas (Rowland et al., 2010; Stern et al.,
2010). Ruapehu lies within an extensional graben at the
southern tip of the TVZ above thicker (40 km) crust (e.g.
Villamor & Berryman, 2006a, 2006b; Salmon et al., 2011).
Prior to the establishment of the Ruapehu magmatic
system, the crustal section would appear to have been
meta-igneous, oceanic crust, which is represented by
meta-igneous xenoliths in central North Island eruptive
rocks, overlain by Mesozoic greywacke and a thin veneer
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JOURNAL OF PETROLOGY
VOLUME 53
of Cenozoic sediments. If the interpretation of the geophysical data for the central TVZ is applied to the region
below Ruapehu, magmatic activity beneath the volcano is
likely to have been initiated by underplating of the crust
by mantle-derived magmas, with subsequent thermal
interaction between these and the lower crust.
NUMBER 10
OCTOBER 2012
The petrological, stratigraphic and geochemical interpretations explored above can be integrated and used
as the basis for a generalized model for the temporal
development of the Ruapehu magmatic system and for the
evolution of North Island crust over the past 2^300 kyr
(Fig. 20). Older, Te Herenga Formation andesite magmas
Fig. 20. (a, b) Schematic diagrams explaining the development of the Ruapehu magmatic system and the temporal evolution of the crust beneath the volcano, with geochemical variation at each stage illustrated in (c) and (d) 87Sr/86Sr vs Zr diagrams (see Fig. 18). Underplating by
mantle-derived, basaltic magma [M in (c) and (d)] heats the crust, and interaction between mantle-derived, fractionating magmas and the
lower crust results in the development of a mixed hot zone in the lower crust. Te Herenga magmas are generated by mixing of partially
melted lower, meta-igneous crustal rocks (lower crust^melt and plagioclase/orthopyroxene restite) and fractionated mantle-derived magmas.
Over time magmas derived from the deep crustal melt zone become dispersed in storage systems throughout the crust, where they evolve at
various depths and over different time scales through fractional crystallization (FC) involving plagioclase, clinopyroxene, orthopyroxene and
magnetite (Pl þ Cpx þ Opx þ Mt) and assimilation (AFC) of Torlesse meta-greywacke basement. Movement of magmas from deeper to
higher crustal levels results in variable amounts of mixing and mingling of magmas from different storages within the shallower crust.
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PRICE et al.
ANDESITE PETROGENESIS, RUAPEHU VOLCANO
were generated in a deep crustal, anatectic zone where
fractionating, mantle-derived magmas mixed with crustal
melts and their complementary restite. On a regional
scale this zone could be analogous to the MASH zone of
Hildreth & Moorbath (1988) or the deep crustal hot zone
of Annen et al. (2006). Subsequently, magmas from this
deep crustal hot zone rose into the overlying crust where,
by the time of emplacement of the Wahianoa Formation, a
dispersed and complex plumbing system of magma reservoirs and sills and dykes had become established. Post-Te
Herenga andesites evolved within this storage and plumbing system, with AFC and mixing processes taking place
on different volume scales, at varying depths and on variable time scales. With time the plumbing and storage
system and the magmas within it became progressively
more complex. Magmas moving through from deeper storage levels interacted and mixed with evolving magmas
stalled at higher levels.
The development of a deep crustal hot zone and the evolution of the overlying magmatic system have undoubtedly
modified the crust. The deep crust has been intruded
by mantle-derived magmas with consequent anatexis.
Magmas have been dispersed throughout the crust where
they have interacted with meta-greywacke basement and
crystallized to form intrusions, cumulates and crystal
mushes. Seismic data for the central North Island have
been interpreted to indicate the presence of extensive intrusive units below 6 km beneath the TVZ (Sherburn
et al., 2003).
The transition from deep to shallower crustal processes
with time reflects progressive evolution of the magmatic
system as mantle-derived magmas moved into the lower
crust and derivative magmas moved to higher levels, but
it is likely that the regional structural framework was also
changing. Differences between Te Herenga and post-Te
Herenga magmas could be related to a fundamental
change in the regional stress regime (Lee et al., 2008).
After the emplacement of the Te Herenga lavas, the rate
and direction of extension changed as the southern tip of
the TVZ extensional zone propagated southwards, and
this may have been accompanied by a change in magma
flux into and through the crust.
Over the life of the TVZ, the development of andesitic
magmatic systems similar to Ruapehu appears to have
been a precursor to rhyolitic magmatism, which is associated with rapid extension and lithospheric thinning.
Extension, crustal thinning, thermal erosion by upwelling
mantle and large-scale underplating by mantle-derived
magmas may be the major thermal drivers leading to crustal melting and rhyolitic magmatism, but compositional
and thermal preconditioning of the crust by andesitic magmatism must also be a factor affecting rhyolite production
and geochemistry (Price et al., 2005). Crustal preconditioning and thermal maturation may be a general influence
on the production of rhyolite in subduction-related magmatic systems, reflecting the temporal evolution and maturity of the arc system even in intra-oceanic tectonic
settings (e.g. Tamura & Tatsumi, 2002; Smith et al., 2006;
Brophy, 2008).
CONC LUSIONS
Ruapehu volcano is a complex stratovolcano that has been
constructed over 250 kyr in a series of time-restricted
eruptive events. The volcano is located at the southern,
propagating tip of the TVZ where crustal thickness increases from 25 km in the TVZ to the north to 40 km beneath Ruapehu. The dominant magmas are andesitic and
geochemical, petrographic and mineralogical data indicate that these represent complex, crystal-rich mixtures of
dacite and rhyolite that have been derived from parental
magmas generated both in the mantle and from a spectrum of crustal sources. They contain abundant crystals
derived from multiple sources (xenocrysts, antecrysts and
xenocrysts). The crystal cargo of a typical Ruapehu andesite reflects partial melting, mixing and AFC processes
that have taken place over a range of crustal and mantle
pressures with different magma batches evolving on varying temporal and spatial scales. The Ruapehu case demonstrates that, in continental andesite volcanoes, whole-rock
compositions are not necessarily direct analogues for melt
compositions (e.g. Gamble et al., 1999, 2003; Hobden et al.,
1999; Dungan et al., 2001; Price et al., 2005, 2007; Kent
et al., 2010); most andesites do not represent liquids on
simple lines of descent that connect parent and daughter
magmas and controlled exclusively by fractional crystallization or AFC (e.g. Eichelberger, 1978). At Ruapehu,
each erupted andesite is a unique blend of crystals and
melts derived from a variety of sources through polybaric
fractional crystallization, crustal assimilation and mixing
and mingling.
The Ruapehu andesites are representative of many
subduction-related volcanic rocks, in both intra-oceanic
and continental tectonic settings (e.g. Gamble et al., 1999,
2003; Smith et al., 2003, 2010; Price et al., 2005, 2007, 2010).
The interpretations developed here may be more generally
applicable and, like those from Ruapehu, many
intra-oceanic and intra-continental andesites may be the
outcome of an interplay of polybaric processes taking
place on highly variable time scales, involving fractional
crystallization, crustal anatexis and magma mingling and
mixing (e.g. Hildreth & Moorbath, 1988; Price et al., 2005;
Annen et al., 2006; Reubi & Blundy, 2009).
The complexity of the magma plumbing system and the
intricate nature of the processes by which andesitic
magmas have evolved in the Ruapehu system are demonstrated by detailed chemo-stratigraphic studies and the
variability in magma compositions erupted during the
period 1945^1996; the latter cover the entire spectrum of
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VOLUME 53
post-Te Herenga compositional variations. Flow sequences
such as those of the Wahianoa Formation exposed on eastern Ruapehu do not show coherent, systematic changes
with time. On a relatively small scale (2^4 sequential lava
flows) there are cases where stratigraphically controlled
changes in chemistry can be demonstrated. This relatively
uncommon pattern is generally characterized by a change
from relatively evolved SiO2- and K2O-rich rocks to
higher MgO and lower SiO2 and K2O abundances indicating sequential tapping of a small-volume magma reservoir. The overall dominant pattern is, however, one of
unsystematic and unpredictable changes in composition
between sub-units of sequential lava flows, suggesting
that separate dispersed magma batches have been tapped
and erupted to construct these large-volume lava flow
sequences.
The isotopic differences between the Te Herenga
Formation lavas can be explained in terms of a progressive
change in the nature of the dominant crustal component.
The isotopic composition of the Te Herenga magmas appears to have been strongly influenced by a deep crustal,
meta-igneous component, but, although this component is
likely to have continued to contribute to the evolution of
the post-Te Herenga magmas, the isotopic compositions of
the latter indicate the additional and dominant involvement of shallower crust with a composition similar to basement meta-greywacke.
The development of Ruapehu-like andesitic systems appears to be a precursor to the large-scale rhyolitic magmatism that dominates the present-day, central TVZ (e.g.
Wilson et al., 1995; Saunders et al., 2010). The development
of the andesitic systems involves underplating, crustal anatexis with the formation of a deep crustal hot zone and
evolution of a dispersed system of magma storage, crystal
mushes and cumulates throughout the crust. These andesitic magmatic systems are associated with incipient extension as the TVZ has propagated and developed, and their
formation preconditions the crust, facilitating later rhyolitic magmatism.
AC K N O W L E D G E M E N T S
We thank Madeleine Humphreys, Yoshihiko Tamura and
Michael Dungan for thorough and constructive reviews,
and Simon Turner for his editorial guidance through the
submission and review process. We particularly acknowledge the pioneering mapping of Bill Hackett, whose
work, together with Ian Graham’s investigations of petrology and geochemistry, provided the robust foundation
on which our mapping and sampling were carried out.
BSc (Honours) projects by Eloise Beyer, John Chapman,
Janet Schneider and Dianne Valente were crucially important for the mapping and sampling aspects of the project. Bernhard Spo«rli’s knowledge of the geology of the
Whakapapa skifield assisted our interpretation of
NUMBER 10
OCTOBER 2012
stratigraphic relationships in the Whakapapa Formation.
Steve Eggins and Craig Cook are thanked for their contributions to trace element and mineral analysis respectively.
Nick Mortimer and Barry Roser provided samples and/or
data that are the basis for our compilation of basement
greywacke chemistry. Jorg Metz, Ian McCabe and
Gordon Holm are thanked for their technical support.
Harry Keys and staff of the Department of Conservation
have provided long-term, sustained and enthusiastic logistic support and advice.
FU NDI NG
R.C.P., J.A.G. and I.E.M.S. acknowledge support from
Australian Research Council (grant number A39531624),
the New Zealand Foundation for Research Science and
Technology (grant number MAUX0401), the Marsden
Fund of New Zealand (grant number UOW106) and
Science Foundation Ireland.
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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