JOURNAL OF PETROLOGY
VOLUME 41
NUMBER 6
PAGES 867–895
2000
Magma Origin and Evolution of White
Island (Whakaari) Volcano, Bay of Plenty,
New Zealand
J. W. COLE1∗, T. THORDARSON2 AND R. M. BURT1
1
DEPARTMENT OF GEOLOGICAL SCIENCES, UNIVERSITY OF CANTERBURY, PRIVATE BAG 4800, CHRISTCHURCH,
NEW ZEALAND
2
CSIRO, MAGMATIC ORE DEPOSIT GROUP, DIVISION OF MINING AND EXPLORATION, PRIVATE BAG PO WEMBLEY,
6014 W.A., AUSTRALIA
RECEIVED FEBRUARY 5, 1998; REVISED TYPESCRIPT ACCEPTED NOVEMBER 23, 1999
White Island is an active composite stratovolcano in the Bay of
Plenty, New Zealand, that comprises many small volume
(<0·1 km3) andesite–dacite lava flows and pyroclastic deposits
with phenocryst contents of >15–44%. Minor high-Mg basaltic
andesite explosive eruptions, such as those of 1976–1992, may
have occurred at intervals throughout the history of White Island,
but are rarely preserved. These alternate with major episodes of
andesite–dacite lava extrusion. The high-Mg magmas form by
hydrous melting of mantle, metasomatized by fluids from the
dehydrating slab at the slab–mantle wedge interface, that rise
rapidly to shallow magma chambers (2–7 km?) where limited
mixing and contamination occurs before eruption. Some of this
magma remains in the magma chamber where it interacts with the
crystal mush, from which it inherits phenocrysts, to form so-called
‘dirty’ lavas. Total phenocryst content of these lavas is correspondingly
higher. As more magma is intruded into the chamber, the heat flux
will increase and melt fraction will eventually rise to the surface to
form high-silica andesite–dacite magma (‘clean’ lavas) with fewer
inherited phenocrysts. Similar multi-magma chamber plumbing
systems, with complex evolution involving fractionation and contamination, probably occur in most andesite–dacite arc volcanoes.
New Zealand lies along an obliquely convergent plate
boundary between the Pacific and Australian plates. East
of the North Island, the Pacific Plate is being subducted
westwards under the Australian Plate at a rate of
>50 mm/yr to form the Taupo–Hikurangi arc–trench
system. Volcanic activity is now concentrated in the
Taupo Volcanic Zone (TVZ), the youngest expression
of 22 Ma Cenozoic continental margin arc volcanism in
the North Island of New Zealand (Cole, 1990). TVZ
(Fig. 1 inset) is at the southern end of the 2800 km
Tonga–Kermadec–Taupo volcanic arc system and the
Lau Basin–Havre Trough–Ngatoro Basin back-arc basin
system. Southern Havre Trough and TVZ are offset
sinistrally by 40–50 km and Wright (1992) believes the
offset to be accommodated by a series of dextrally oblique
and en-echelon bookshelf faults. The southwestern limit
of this fault system is placed <20 km north of White
Island and has NW–SE orientation (Fig. 1 inset).
TVZ extends offshore into the Bay of Plenty as a
graben structure of 45 km width (Fig. 1) bounded by the
Tauranga Fault Zone to the west and the White Island
Fault Zone to the east (Wright, 1992). The eastern side
of TVZ coincides with the current active volcanic front
and a broadly linear alignment of andesitic volcanoes
extends from Ruapehu in the southwest to Whakatane
seamount, located at the edge of the continental crust,
in the northeast (Gamble et al., 1993a; Fig. 1 inset).
Rhyolite volcanism dominates central TVZ and is associated with eruptions of minor high-Al basalt (Cole,
1990). Andesite volcanism dominates both ends of TVZ
and is also recognized as lithic fragments in rhyolitic
∗Corresponding author. Telephone: 64-3-364-2766. Fax: 64-3-3642769. e-mail: [email protected]
Extended data set can be found at: http://www.petrology.
oupjournals.org
 Oxford University Press 2000
andesite petrogenesis; magma evolution; high-Mg magmas;
Taupo Volcanic Zone; White Island
KEY WORDS:
INTRODUCTION
JOURNAL OF PETROLOGY
VOLUME 41
NUMBER 6
JUNE 2000
Fig. 1. Map of the Bay of Plenty. Offshore faults and main structural elements are from Wright (1992). VR, Volckner Rocks. Inset map of
Taupo Volcanic Zone shows location of White Island and other andesitic volcanoes mentioned in the text. Thick dashed line in inset map
represents the southern limit of inferred en-echelon cross-fracture structures (Wright, 1992).
ignimbrites erupted from calderas in central TVZ (Cole
et al., 1998).
White Island, together with the sea stacks of Club
Rocks and Volckner Rocks (Figs 1 and 2), is the emergent
summit of a larger (16 km × 18 km) submarine structure
(Duncan, 1970) hereafter referred to as White Island
Volcano. White Island Volcano has a total volume of
>78 km3, and is located at the southern end of a chain
of seamounts, which form the Ngatoro and White Island
ridges (Fig. 1).
White Island is New Zealand’s most active volcano
and in historical times (the last 150 years) has been
characterized by sporadic eruptive episodes featuring
small phreatic, phreatomagmatic and strombolian eruptions, associated with extensive fumarolic activity. The
last eruptive episode on White Island began in 1976,
with numerous small phreatomagmatic and strombolian
eruptions (Houghton & Nairn, 1989). Olivine-bearing
basaltic andesite bombs and blocks were erupted in
March 1977 (Cole & Graham, 1987), and it is these
samples that are compared and contrasted with the
prehistoric lavas exposed in outcrops on Ngatoro and
Central cones. The most recent eruptive episode ceased
in 1992, although small phreatic explosions continue,
and the level of activity is now (1999) again increasing.
This paper reports new field, petrographic, mineralogical, chemical and isotopic data from lavas and
pyroclastic rocks of White Island Volcano. Unlike earlier
studies on the island, the geochemical data have been
fitted into a stratigraphic succession. This has allowed
changes in magma composition with time to be evaluated.
We compare these new data with studies of surrounding
basalts, andesites and dacites, including nearby Motuhora
Island (Fig. 1), and discuss implications for magma origin
and evolution.
GEOLOGY OF WHITE ISLAND
868
The geology of White Island Volcano is dominated by
two overlapping stratocones (Black, 1970; Duncan, 1970).
Remnants of the extinct Ngatoro Cone crop out to the
west of the currently active Central Cone (Fig. 2). These
cones, together with the outlying lavas of Volckner Rocks,
Troup Head and Club Rocks, can be divided into 22
eruption units (see Appendix for a list of these units, their
locations on the island and sample numbers). A schematic
stratigraphy of the Ngatoro lava succession is shown in
Fig. 3 and that of Central Cone in Fig. 4. Much of
Ngatoro Cone was removed before the construction of
COLE et al.
MAGMA ORIGIN AND EVOLUTION OF WHITE ISLAND
Fig. 2. Location map of White Island. Boundary of Ngatoro Cone and Central Cone from T. Thordarson et al. (unpublished data. 1999)
although it should be noted that eruptions of lava and tephras from Central Cone have partially covered much of Ngatoro Cone. Arrow depicts
dip of volcanic layers in Ngatoro Cliffs. Dashed lines along main crater floor depict boundaries of sub-craters whereas those in the bays adjacent
to Pinnacle Head and Troup Head locate hydrothermal explosion craters. RW- and TRW- numbers depict sample locations. Inset displays
relative position of Club and Volckner Rocks relative to White Island.
Central Cone. This is especially true of the northern
sector of Ngatoro, whereas the southern slopes appear
to be concordant with the overlying Central Cone. This
suggests that while erosion was occurring on one part of
the island, lava was being erupted at another, and that
the stratigraphic succession is at least semi-continuous
through time. Like Ngatoro Cone, Central Cone also
comprises alternating lava flows and pyroclastic and
epiclastic deposits. The main crater comprises at least
three coalescing sub-craters (Fig. 2). The sub-craters are
aligned NW–SE and the southeastern crater wall is
breached by the sea in three places to form Crater,
Wilson and Shark Bays, which are separated from each
other by Troup Head and Pinnacle Head (Fig. 2). These
869
JOURNAL OF PETROLOGY
VOLUME 41
three bays are believed to have been formed by longlived, semi-continuous, weak, hydrothermal explosions,
similar to those occurring in the Western sub-crater
today. The present crater floor is less than 30 m above
sea level and is largely covered by material from the
1914 debris avalanche, which killed 11 sulphur miners.
Lavas forming Troup Head and Pinnacle Head are
collectively referred to as Troup Head lavas. Troup Head
lavas have previously been considered to have erupted
from a parasitic vent located on the eastern flank of
Central Cone (Hamilton & Baumgart, 1959). However,
lavas of Troup Head, Pinnacle Head and the southern
main crater wall comprise a single flow dipping towards
the southeast and are believed to have been erupted from
a vent in the area now occupied by Central Cone subcraters. Furthermore, lavas of Troup Head lie above the
earliest lavas of Central Cone in the vicinity of Shark
Bay but are found below lavas of Central Cone near
Otaketake Point (Fig. 2). Consequently, we believe the
Troup Head lavas to be part of the Central Cone volcanic
succession and not erupted from a parasitic cone. Troup
Head lavas differ from those of Central Cone in their
mineral assemblage and chemistry. They have one of
the most chaotic mineral assemblages of all lavas on
White Island, suggesting a complex mixing history.
Club Rocks lie 800 m south of Otaketake Point (Fig.
2 inset) and comprise a set of four sea stacks rising to
>20 m above sea level. These lavas also show a complex
mixing history, similar to those of Troup Head. They
are unconformably overlain by carapace breccias and
stratified tuffaceous sandstones and volcanic breccias,
which may represent the remnants of highstand sea-level
beach. Analysis of in situ offshore samples, collected at
depths of 20–30 m from vertical lava walls, in conjunction
with side-scan sonar observations, shows that Club Rocks
rest on lavas (E6; Otaketake Point lava I; Fig. 4), exposed
above sea level at Otaketake Point (T. Thordarson & G.
Kurras, unpublished data, 1999). Consequently, Club
Rocks are not part of an older cone as suggested by
Hamilton & Baumgart (1959) but an integral part of
Central Cone.
Five kilometres northwest of White Island are four sea
stacks collectively known as Volckner Rocks (Fig. 2 inset),
three of which rise precipitously from the sea floor
(>100 m depth) to a maximum height of 113 m above
sea level. The fourth is an eroded stump. The southeasternmost stack reveals well-jointed columnar lava with
abundant enclaves and is overlain by volcanic breccias
forming a roof carapace. The four stacks are considered
to be the eroded remnants of a lava dome, which may
represent the largest volume of any single magma body
of White Island Volcano. The age of Volckner Rocks
relative to the rest of White Island Volcano is unknown.
Numerous fumaroles and hot springs discharge onto
the floor of the main crater and represent the surface
NUMBER 6
JUNE 2000
expression of the White Island geothermal system (Giggenbach & Sheppard, 1989). Fumaroles and hot springs
are not restricted to the crater floor and have been
observed along much of the coastline and offshore. Despite the widespread nature of hydrothermal activity,
most lavas exposed in the crater walls are fresh and
exhibit little or no alteration, with fresh volcanic glass
not uncommon (see Table 2, below). Significantly, on 21
October 1992 (C. P. Wood, personal communication,
1996) a hot scoriaceous block with fresh glass, no alteration and chemically identical to prehistoric Central
Cone lavas, was ejected from the active crater. The
location of the sample in relation to the vent and crater
walls precludes the incorporation of this block into the
ejecta after falling from Mt Gisborne. Consequently, we
believe that it represents conduit wall material entrained
within the 21 October 1992 eruption and can be used to
demonstrate that unaltered volcanic material is available
within the geothermal field of White Island and that
alteration is neither as pervasive nor as extensive as
previously believed (Wilson, 1959).
Lavas of Central Cone, Ngatoro Cone, Troup Head,
Club Rocks and Volckner Rocks are collectively referred
to as prehistoric lavas, whereas samples erupted in March
1977 are hereafter referred to as 1977 ejecta.
METHODS
Samples for geochemical analysis were collected from
the least altered outcrops (locations shown in Fig. 2)
and further screened for the effects of alteration when
examined in thin section. Loss on ignition varies between
+2% and –1%, with the majority of samples falling
between +1% and –1%. Consequently, we believe that
the data presented here represent fresh magmatic compositions. At least 1 kg of each sample was crushed and
the freshest chips were retained for milling in a tungsten
carbide concentric ring mill. Major and trace element
concentrations were determined at the University of
Canterbury (RW sample series) and at CSIRO Floreat
Park Laboratories (TRW sample series) by wavelengthdispersive X-ray fluorescence spectrometry (XRF), using
Philips PW 1400 and 1480 automatic X-ray spectrometers, respectively. Estimated precision and interlaboratory uncertainty are within 1% and 5% for the
major and trace elements, respectively. Analytical procedure follows that of Norrish & Hutton (1969), and
further details on detection limits and calibrated analytical
uncertainty have been given by Weaver et al. (1990).
Volatile loss on ignition (LOI) was determined by weight
loss after fusing at 1000°C for 20 min. Whole-rock and
mineral mg-number are calculated by Mg2+/(Mg2+ +
Fe2+) with all Fe as Fe2+ (mg-numbers are therefore
minimum values), except for pyroxene where Fe2+ and
870
COLE et al.
MAGMA ORIGIN AND EVOLUTION OF WHITE ISLAND
Fig. 3. Schematic stratigraphic logs from the north and west faces of the Ngatoro cone, White Island Volcano, illustrating the volcanic
stratigraphy and correlation of lavas between the sections. Lava eruption unit number (E11, etc.) and names are given on the side of the column
and the star indicates lava flows sampled for this study (see Appendix). (See key for other symbols.) Based on T. Thordarson & B. F. Houghton
(unpublished data, 1999).
Fe3+ are calculated according to the method described
by Droop (1987).
Analyses of minerals in White Island samples and
groundmass glass in RW34 (Table 2, below) were
undertaken at the University of Otago with a JEOL
Superprobe JXA 8600 operating with an accelerating
voltage of 15 kV, a beam current of 20 nA and a 5–20 m
beam diameter. Data were reduced on-line using standard
871
JOURNAL OF PETROLOGY
VOLUME 41
NUMBER 6
JUNE 2000
Fig. 4. Schematic stratigraphic logs from the north and south crater walls of the Central Cone, White Island Volcano, illustrating the volcanic
stratigraphy and correlation of lavas between the sections. Lava eruption unit number (E1, etc.) and names are given on the side of the column
and the star indicates lava flows sampled for this study (see Appendix). Inset shows the probable stratigraphic position of the Club Rocks lava
in relation to the south crater wall succession. HTAEB, hydrothermally altered ejecta bed. (See key for other symbols.) Based on T. Thordarson
& B. F. Houghton (unpublished data, 1999).
ZAF correction procedures (Sweatman & Long, 1969;
modified by Y. Kawachi & M. Trinder, unpublished
data, 1993). The groundmass glass analyses in samples
TRW34 and TRW35 (Table 2) were carried out at
872
COLE et al.
MAGMA ORIGIN AND EVOLUTION OF WHITE ISLAND
University of Hawaii at Manoa on a Cameca
SX50 microprobe using the same instrumental settings
as given above and on-line ZAF correction procedure.
Rare earth elements (REE), U, Th, Ta and Hf were
determined by inductively coupled plasma mass spectrometry (ICP-MS) at Analabs, Perth, Western Australia.
Sr, Nd and Pb isotope ratios were analysed at the
Department of Geology, Royal Holloway College, University of London using a VG Isomass 354 five-collector
mass spectrometer. Conventional ion exchange techniques were used to separate Sr, Pb and Nd. Data
collection for Sr and Nd was multidynamic, whereas that
for Pb was static. Blanks are >0·5 ng Sr, 0·3 ng Nd
and 0·3 ng Pb and not significant in this study. The
international NBS SRM987 standard gave 0·710241 ±
22 (2), internal errors are >±0·00001 (2). 143Nd/
144
Nd was normalizeed to 146Nd/144Nd = 0·7219, Aldrich
standard 0·511410 ± 7 (2). Internal errors are
>0·00001 (2) (Thirlwall, 1991). Pb isotope ratios were
normalized for mass fractionation to international NBS
SRM981 standard by >0·13% per a.m.u.; internal errors
were better than 0·005% and reproducibility was estimated at better than 0·05%.
8a). Larger vesicles typically have uneven globular outlines as a result of growth by coalescence of vesicles.
The main phenocryst assemblage is plagioclase, clinopyroxene and orthopyroxene. Spinel (magnetite and
chromite; Table 1) microphenocrysts occur as a minor
phase in all lava formations, except the spatter or scoria
bombs of E1, where it is also absent as a groundmass
phase. Anhedral olivine phenocrysts occur in trace to
minor amounts in most lavas and are commonly jacketed
by aggregates of small orthopyroxenes with intergrowth
of vermicular titanomagnetite. Neither this study nor that
of Duncan (1970) identified the primary biotite, reported
from White Island lavas by Black (1970). Within individual lava eruption units the phenocrysts are present
as single crystals, single mineral aggregates or as heterogeneous (multi-mineral) crystal aggregates. All lavas contain broken and highly fractured phenocrysts as well as
variable amounts of very small (Ζ0·4 mm) jagged crystal
fragments, apparently chipped from larger phenocrysts.
Occasionally, vesicles are seen to penetrate and split a
phenocryst.
Plagioclase
PETROGRAPHY AND MINERALOGY
A total of 16 of the 22 eruption units on White Island
(see Appendix for sample numbers within each unit) have
been described in detail from 24 thin sections covering
all lava types recognized on the island. Modes of these
sections are estimated using standard charts. These are
given in Table 1 and summarized in Fig. 5. These show
a clear bimodality of phenocryst modes with peaks at
25–30 and 40–45 modal %. Representative electron
probe microanalyses of mineral phases are given in Table
2 and shown graphically in Figs 6 and 7. The groundmass
crystallinity ranges from hyaline to holocrystalline. Hyaline groundmass consists of brown transparent glass
with 0–5% plagioclase ± pyroxene microlites, whereas
hypohyaline groundmass is typically microcrystalline with
a high abundance of aligned lath- to needle-shaped
plagioclase and lesser amount of interstitial magnetite
and pyroxene. Hypocrystalline to holocrystalline groundmass is microcrystalline to fine grained, with a preponderance of plagioclase laths, usually in parallel
arrangement, variable amount of interstitial glass, and
small amount of interstitial magnetite and pyroxene.
Texture is typically pilotaxitic, and less frequently intergranular (Table 1). The crystalline interior of lavas is
non- to poorly vesicular, featuring stretched, irregular
vesicles. Scoria bombs from the 1977 ejecta (E1 bombs)
have vesicularity in excess of 45% and contain numerous
relatively small spherical to ovoid-shaped vesicles (Fig.
Plagioclase phenocrysts are a heterogeneous assemblage
of crystals of variable size [0·1–5·0 mm, usually with a
bimodal size-distribution peaking in the sub-millimetre
(P1) and 1·5–2·5 mm (P2) size range, respectively], modes
(9–25·8%; Table 1; Fig. 5b), shape (sub- to euhedral),
coherency, compositional zoning (An52–97; Table 2; Fig.
6), dissolution–resorption structures, inclusions and fractures. Chemically the greatest compositional range is in
the mafic andesite lavas of Troup Head, Club Rocks and
Ngatoro (Fig. 6), with a compositional break between
more and less calcic plagioclase. The dacites from Central
Cone have a smaller continuous compositional range.
Six plagioclase sub-populations can be identified (P1–P6),
based on shape and internal features. P1 and P2 phenocryst types are the most common and are subdivided
into subtypes (P1a, P1b, P2a, P2b, etc.) on internal
features. Types 3–6 are only present in minor amounts.
Type P1—clear, euhedral phenocrysts
873
These phenocrysts have sharp, well-defined outlines
and tabular to elongate shape (Fig. 8a) and are typical
of the 1977 ejecta and the dacites. They are clearly
juvenile phenocrysts that have crystallized from the
host magma. Modes vary from 2 to 14·4% (Table 1;
Fig. 5c) and the size is typically <1·5 mm (total range
0·1–4·0 mm). P1 phenocrysts usually occur as single
crystals or as homogeneous crystal aggregates with a
single large crystal surrounded by smaller ones. However, they are also present in heterogeneous crystal
aggregates, either as the main phenocryst phase or as
40·1
38·2
25·4
25·2
33·3
28·0
27·0
23·5
E8
874
E9
E10
E12
E14
E15
E16
E22
15·3
15·0
12·0
6·1
5·0
2·0
5·6
3·8
2·0
9·3
8·0
6·9
7·5
9·2
10·0
10·0
13·3
8·0
10·1
16·5
14·0
3·6
7·1
15·2
0
0
0
0
0·2
0
1·0
1·0
0·8
0·5
5·1
5·1
6·0
10·0
13·3
10·0
12·1
9·3
14·0
7·7
7·1
11·1
12·9
<1
<1
<1
<1
<1
<1
>1
>1
<0·01
3·0
3·0
1·1
0·2
0·1
0·1
0·1
0
0
>1
>1
2·5
2·4
<1
<1
0·5
3·1
3·0
3·0
2·2
3·0
1·0
3·1
3·0
1·7
1·7
2·5
3·5
0·5
2·0
1·0
2·0
5·6
3·0
5·1
3·1
7·0
8·6
8·7
5·1
4·7
4·5
8·6
76·5
73·0
72·0
66·7
74·8
74·6
61·8
59·9
79·4
76·1
58·6
56·5
85·0
76·3
crystalline
hypocrystalline
crystalline
hypohyaline
hyaline/
hypohyaline
hyaline/
hypohyaline
hyaline/
crystalline
crystalline
pilotaxitic
hyalopilitic
pilotaxitic/
pilotaxitic
hyaline/hyalopilitic
hyaline/hyalopilitic
hyaline/hyalopilitic
intergranular
intergranular
pilotaxitic/
hyalopilitic
pilotaxitic/
pilotaxitic
hypocrystalline
hyalopilitic/
crystalline
hyalopilitic
pilotaxitic/
pilotaxitic
intergranular/
hypohyaline/
hypocrystalline
crystalline
intergranular
pilotaxitic/
hyalopilitic
crystalline
pilotaxitic/
hypohyaline
hypohyaline
crystalline/
—
pl; mt, px
pl; dev. glass, mt, px
pl; mt, px, dev. glass
dev. glass; pl, mt
dev. glass; pl, mt, px
dev. glass; pl, mt
pl; mt, px, glass
pl; mt, px, dev. glass
pl; glass, mt, px
glass, pl; mt, px
pl, glass; mt, px
pl; mt, px, glass
pl; mt, px, glass
pl, glass; mt, px
glass; pl, px
np
ap
ap.?
ap
ap
ap
ap
ap?
np
ap
ap
np
np
np
—
ACC
NUMBER 6
16·7
12·0
12·1
25·8
23·0
11·3
15·1
5·1
4·1
>1
2·0
hyaline/
GRM constituents
VOLUME 41
TPM, total phenocryst mode; PL, total plagioclase phenocrysts; P1, Type 1 plagioclase phenocrysts; P2, Type 2 plagioclase; Po, Types 3–6 plagioclase phenocrysts;
PX, pyroxene; OL, olivine; OP, opaques; HCA, heterogeneous crystal aggregates; GRM, groundmass; ACC, accessory minerals; cpx/opx, clinopyroxene/orthopyroxene
ratio; dev, devitrified; mt, magnetite; ap, apatite; np, not present.
20·6
E7
23·9
E6
25·3
12·4
5·0
<0·01
58·0
hyaline/hyalopilitic
41·4
8·2
1·5
>1
2·8
—
E5
24·7
3·0
7·4
0
72·8
43·5
4·5
1·1
0·8
—
E4·5
9·0
9·6
0·95
0·1
15·0
3·6
20·5
6·2
E4
14·3
4·3
1·05
23·7
2·2
11·3
E3
14·4
—
—
20·7
—
42·0
E1bomb
—
27·2
E1block
9·6
GRM texture
Formation TPM (%) PL (%) P1 (%) P2 (%) Po (%) PX (%) cpx/opx OL (%) OP (%) HCA (%) GRM (%) GRM crystallinity
Table 1: White Island eruptive rocks: phenocryst modes of selected eruption units (on vesicle-free basis)
JOURNAL OF PETROLOGY
JUNE 2000
0·6
18·3
MnO
MgO
875
0·07
0·00
4·07
71·09
25·59
3·32
0·85
0·02
1·04
0·06
0·00
4·00
52·79
44·16
Mn
Mg
Ca
Na
Cation total
Enstatite
Ferrosilite
Wollastonite 3·05
1·50
0·01
0·53
0·00
0·01
Ti
Fe2+
1·90
0·06
1·98
0·04
Si
99·8
0·0
1·9
27·3
0·3
17·3
51·5
pcryst
OPX
Al
99·9
26·7
FeO
Total
0·3
TiO2
0·0
0·8
Al2O3
1·5
51·7
SiO2
CaO
0·2
gmass
Site:
Na2O
1·3
OPX
Mineral:
E3
Sample:
RW38
E3
RW38
E no.:
42·29
14·65
41·41
4·04
0·02
0·85
0·81
0·01
0·34
0·01
0·10
1·91
100·9
0·3
21·2
14·5
0·3
10·9
0·5
2·2
51·0
pcryst
CPX
RW38
E3
43·0
17·6
39·4
4·00
0·01
0·83
0·76
0·01
0·34
0·01
0·06
1·97
100·4
0·2
20·8
13·7
0·3
10·9
0·3
1·4
52·8
gmass
CPX
RW14
E5?
2·0
38·0
60·0
4·00
0·00
0·04
1·18
0·03
0·74
0·00
0·02
1·99
100·5
0·0
1·0
21·1
0·8
23·8
0·1
0·5
53·2
pcryst
OPX
RW14
E5?
39·5
17·2
43·3
3·99
0·02
0·76
0·83
0·01
0·33
0·01
0·08
1·96
99·1
0·2
18·8
14·8
0·3
10·5
0·5
1·9
52·1
pcryst
CPX
RW51
E10
40·1
9·4
50·5
3·98
0·01
0·76
0·95
0·01
0·18
0·01
0·09
1·98
99·3
0·2
19·3
17·5
0·2
5·8
0·2
2·0
54·1
pcryst
CPX
RW51
E10
3·8
30·1
66·0
3·99
0·00
0·07
1·29
0·01
0·59
0·01
0·04
1·98
100·5
0·0
1·9
23·6
0·4
19·2
0·3
0·9
54·2
pcryst
OPX
RW51
E10
3·7
19·1
77·2
3·99
0·00
0·07
1·51
0·01
0·37
0·00
0·06
1·98
100
0·0
1·9
28·4
0·2
12·5
0·1
1·3
55·6
pcryst
OPX
RW51
E10
42·9
12·4
44·7
4·02
0·03
0·84
0·87
0·01
0·24
0·01
0·07
1·95
99·8
0·4
21·1
15·8
0·2
7·8
0·3
1·7
52·5
pcryst
CPX
RW22
E18
4·0
25·8
70·2
4·00
0·00
0·08
1·38
0·01
0·51
0·01
0·04
1·97
99·5
0·0
2·0
25·3
0·4
16·6
0·2
1·0
54·0
pcryst
OPX
RW22
E18
37·7
18·3
44·0
4·00
0·02
0·72
0·84
0·01
0·35
0·01
0·08
1·96
99·4
0·2
18
15·1
0·3
11·2
0·3
1·9
52·4
gmass
CPX
RW30
E22
1·9
38
60·1
3·99
0·00
0·04
1·17
0·03
0·74
0·00
0·03
1·99
100·6
0·0
0·9
21·0
0·9
23·7
0·1
0·6
53·4
gmass
OPX
RW30
E22
2·6
27·2
70·1
3·99
0·00
0·05
1·35
0·01
0·52
0·01
0·07
1·98
100·2
0·0
1·3
25·0
0·3
17·3
0·2
1·6
54·5
gmass
OPX
RW30
E22
42·6
14·3
43·1
4·00
0·01
0·82
0·83
0·01
0·28
0·01
0·11
1·94
99·9
0·2
20·6
15·0
0·2
8·9
0·4
2·4
52·2
pcryst
CPX
RW23
?
Table 2: Electron microprobe analyses of minerals and glass from representative eruption units at White Island
4·1
26·3
69·6
3·99
0·00
0·08
1·35
0·01
0·51
0·01
0·05
1·98
99·2
0·0
2·0
24·7
0·4
16·6
0·3
1·1
54·1
gmass
OPX
RW23
?
SiO2
Fo
Total
NiO
Cr2O3
CaO
MgO
MnO
FeO
TiO2
Al2O3
E1
90·37
98·37
0·25
0·08
0·19
48·54
0·13
9·23
0·0
0·06
39·89
pheno
OL
91·61
99·53
0·32
0·06
0·18
49·93
0·21
8·15
0·0
0·03
40·65
microp
OL
P41600 P41600
E1
COLE et al.
MAGMA ORIGIN AND EVOLUTION OF WHITE ISLAND
33·2
0
17·3
2·1
PLAG
pcryst
49·9
30·3
0·1
1·6
0·3
15·2
2·5
0·5
100·3
Mineral:
Site:
SiO2
Al2O3
TiO2
FeO
MnO
MgO
CaO
Na2O
Total
E3
Sample:
876
0·01
3·46
0·75
0·02
0·24
0·09
2·98
0·89
0·11
Mg
Ca
Na
0·5
1·4
60·9
37·7
20·12
0·06
1·56
2·52
0·02
0·11
0·01
6·31
9·53
99·7
0·3
4·4
12·9
0·1
0·7
0·1
29·3
52·1
PLAG
0
88·4
11·6
20·00
0·00
0·46
3·53
0·03
0·08
0·00
7·34
8·56
100·1
0
1·3
17·9
0·1
0·5
0
33·8
46·5
gmass
E5?
0·6
69·7
29·7
19·94
0·02
1·16
2·72
0·03
0·10
0·01
6·52
9·37
100·7
0·1
3·3
14
0·1
0·7
0·1
30·6
51·8
gmass
PLAG
RW14
0·6
74·9
24·5
19·96
0·00
0·18
3·82
4·00
0·08
0·00
7·52
8·37
100
0·1
2·7
14·9
0·1
0·7
0
31·1
50·4
pcryst
PLAG
RW51
E10
0
95·5
4·5
19·94
0·00
0·39
3·58
0·03
0·09
0·00
7·18
8·66
99·2
0
0·5
19·1
0
0·5
0
34·2
44·9
pcryst
PLAG
RW22
E18
0
90·1
9·9
20·00
0·00
0·39
3·63
0·00
0·08
0·00
7·44
8·47
99·8
0
1·1
18·1
0·1
0·6
0
33
46·9
pcryst
PLAG
RW22
E18
0
91·1
8·9
19·86
0·04
1·29
2·47
0·00
0·12
0·01
6·24
9·68
99·1
0
1·0
18·5
0·1
0·6
0
33·3
45·6
gmass
PLAG
RW30
E22
0·6
78·4
21
20·01
0·00
0·36
3·70
0·03
0·09
0·00
7·32
8·51
100·1
0·1
2·4
16·2
0
0·6
0
31·7
49·1
pcryst
PLAG
RW30
E22
0
90·3
9·7
20·01
0·02
0·85
3·18
0·00
0·09
0·00
6·85
9·00
100·9
0
1·1
18·5
0
0·5
0
34·5
46·3
pcryst
PLAG
RW23
?
1·2
64·9
33·9
19·94
0·02
0·96
2·92
0·03
0·11
0·00
6·70
9·21
100·8
0·2
3·7
12·8
0
0·8
0·1
29·4
53·8
gmass
PLAG
RW23
?
6·13
Fe
Al
Cr
Total
CaO
NiO
MgO
MnO
0·077
0·242
0·681
98·51
0·05
0·18
10·14
0·19
17·76
Fe2O
FeO
0·08
51·54
12·24
0·29
microp
V 2O 3
Cr2O3
Al2O3
TiO2
P41600
E1
Sample:
E no.:
0·086
0·247
0·667
98·81
0·02
0·16
10·81
0·20
17·01
6·84
0·07
50·72
12·59
0·39
microp
E1
0·16
P 2 O5
99·90
2·38
K 2O
5·75
2·53
0·19
6·14
1·01
14·69
99·97
0
2·19
3·28
5·75
1·44
0·12
5·31
0·59
15·59
65·71
(n = 2)
(n = 30)
64·10
GLASS
GLASS
3·01
Total
E11
TRW34/35 RW34
Na2O
CaO
MgO
MnO
FeO
TiO2
Al2O3
SiO2
Number (n):
CHROM CHROM Analysis:
P41600
E1
Location data can be found in Fig. 2. Cation formulae for pyroxenes based on six oxygens and plagioclase on 32 oxygens. E no., eruption unit number; OPX,
orthopyroxene; CPX, clinopyroxene; OL, olivine; PLAG, plagioclase; CHROM, chromite; gmass, groundmass; pcryst, phenocryst; microp, microphenocryst.
2·8
Orthoclase
81·8
17·7
22·4
74·9
Albite
Anorthite
20·20
20·04
PLAG
pcryst
E5?
RW14
NUMBER 6
Cation total
K
0·00
0·02
Ti
Fe2+
E3
RW38
VOLUME 41
0·11
7·31
6·56
Al
8·53
9·16
Si
99·1
0·1
0
0·7
45·7
gmass
PLAG
RW38
E3
RW38
E no.:
Table 2: continued
JOURNAL OF PETROLOGY
JUNE 2000
COLE et al.
MAGMA ORIGIN AND EVOLUTION OF WHITE ISLAND
Fig. 5. Histograms showing the modal per cent distribution of phenocrysts in White Island lavas: (a) total phenocrysts; (b) plagioclase; (c)
P1—Type 1 plagioclase phenocrysts; (d) P2—Type 2 plagioclase; (e) pyroxene (ortho- + clinopyroxene); (f ) opaque minerals (spinel); (g) olivine.
single crystals along the edges of the cluster. In
homogeneous crystal aggregates, the interstices (voids)
between crystals are filled by either pristine, crystalfree, pale to dark brown glass or hypohyaline to
crystalline groundmass material.
877
Subtype P1a—inclusion-free phenocrysts commonly
with a normally zoned core and a thin normal or
reversely zoned rim.
Subtype P1b—have small (10–100 m), spherical or
ovoid to tabular primary brown glass inclusions in
JOURNAL OF PETROLOGY
VOLUME 41
Fig. 6. Plagioclase ternary diagrams for White Island plagioclase
phenocrysts. CR, Club Rocks; TH, Troup Head; VR, Volckner Rocks.
Analyses of 1977 ejecta from Shiraki et al. (1994).
the core of the crystal. The glass is free of crystals
and microlites.
Subtype P1c—usually large (>1 mm) phenocrysts, featuring irregular and convoluted, thin-banded oscillatory zoning extending from the centre to the rim.
Type P2—clear rim but resorbed core
These typically occur as sub- to euhedral, tabular crystals,
commonly cracked, with a wide inclusion-filled core
and clear rims of 50–300 m thickness, usually free of
NUMBER 6
JUNE 2000
inclusions (Fig. 8b), and are more typical of the andesites.
Modes vary from 2·2 to 16·5% (Table 1; Fig. 5d) and
the size is usually >1 mm (total range 0·3–5·0 mm).
The inclusions in the core consist of either crystalline
groundmass or more commonly a pristine brown glass
that usually contains dispersed plagioclase, pyroxene and
opaque (magnetite?) microlites, amounting to 5–20 modal
%. Some of the glass is dark brown (nearly opaque) and
mottled (i.e. hydrated). Occasionally there is a gradational
transition from microlite-poor glass to crystalline groundmass. Small (p100 m) ovoid to tabular primary inclusions also occur in the core of the phenocrysts, and
are commonly aligned along the core–rim boundary or
along compositional zones.
These relationships indicate that the large, irregular
inclusions between the core in P2 phenocrysts are a
dissolution phenomenon, caused by reaction of either a
‘cognate’ or an ‘accessory’ crystal with the host melt.
The relative proportion of large irregular inclusions to
plagioclase in the P2 phenocryst cores is highly variable,
ranging from <10% to 100% inclusions. The rims have
oscillatory zoning expressed as thin (<10 m) bands,
although rims with continuous normal zoning also occur.
Rims are commonly thicker on large phenocrysts and in
a few plagioclase aggregates; the crystals are now sealed
together by a continuous rim.
Fig. 7. Pyroxene phenocryst compositions for White Island lavas. CR, Club Rocks; TH, Troup Head; VR, Volckner Rocks. Classification
scheme of Morimoto (1988). Analyses of 1977 ejecta from Shiraki et al. (1994).
878
COLE et al.
MAGMA ORIGIN AND EVOLUTION OF WHITE ISLAND
Fig. 8. Photomicrographs of lavas from White Island. (a) ‘Clean’ plagioclase (Type 1) and pyroxene phenocrysts plus spherical- to ovoid-shaped
vesicles in hyaline groundmass, 1977 bomb (TRW-35); (b) ‘dirty’ plagioclase phenocryst (Type P2) from Ngatoro Cone (TRW-4); (c) resorbed
olivine surrounded by a corona composed largely of orthopyroxene and magnetite, SW slope of Central Cone (TRW-26); (d) plagioclase–pyroxene
heterogeneous crystal aggregate (HCA) from dyke, NW cliffs, Ngatoro Cone (TRW-23); (e) ‘clean’ dacite lava from Central Cone (TRW-28);
(f ) ‘dirty’ andesite lava from Troup Head (TRW-31). Bar scale on each photomicrograph represents 1 mm.
Type P2 can be subdivided into subgroups on the
extent of resorption:
Subtype P2a—incipient resorption (1–3 large dissolution
inclusions).
Subtype P2b—little resorption (>10–30% of the core
and extinction is patchy).
879
Subtype P2c—moderate resorption (>30–70% of the
core and extinction is patchy).
Subtype P2d—high resorption (>70% of the core and
‘primary’ extinction is largely obliterated).
Subtype P2e—extreme resorption (core is completely
dissolved).
JOURNAL OF PETROLOGY
Subtype
sions.
P2f—numerous
small
dissolution
VOLUME 41
These are typically <2 mm, anhedral to subhedral and
often with rounded outlines. They have no or only a
very thin and discontinuous rim and their core typically
contains between 40 and 70% of small rectangular or
thin elongated inclusions in parallel arrangement and in
alignment with a crystallographic axis. Inclusions are
filled by glass or crystalline groundmass material and
those at the outer margins of phenocrysts open into the
surrounding groundmass.
JUNE 2000
Olivine
inclu-
Type P3—sieve-textured phenocrysts
NUMBER 6
Olivine phenocrysts (0·1–2·0 mm) have been identified
in most lava formations and typically occur in trace
(<0·5%) or subordinate amounts (1–3%), except in E1
block samples (6·2%; Table 1; Fig. 5g), where they are
forsteritic (Fo88–92; Graham & Cole, 1991; Shiraki et al.,
1994). The olivines generally form anhedral or rare
euhedral crystals, many of which are jacketed by a thick
corona of equigranular orthopyroxene and vermicular
magnetite (Fig. 8c). Jacketed anhedral and partly resorbed
phenocrysts also occur within the heterogeneous crystal
aggregates in some lavas.
Type P4—phenocrysts with strained extinction patterns
Spinels (magnetite and chromite)
Type P5—phenocrysts partly replaced by sericite or clay
minerals
Type P6—phenocrysts with rounded or convoluted outlines
as a result of dissolution
Types 3–6 are also largely ‘cognate’ or ‘accessory’
phenocrysts, which are out of equilibrium with the host
melt.
These are present as small (0·1–0·5 mm) dispersed, anhedral to euhedral equant (cubic) microphenocrysts
(modes vary from 0 to 3·5%, although most commonly
in the range of 1·5–3% (Table 1; Fig. 5f ) or as free
crystals or interstitial phase in crystal aggregates. Spinels
also occur as inclusions in pyroxene and olivine. Most
are magnetite but rare small octahedra of chromite up
to 0·02 mm (Table 2) also occur (Shiraki et al., 1994).
Pyroxene (clinopyroxene + orthopyroxene)
All White Island lavas contain sub- to euhedral clinopyroxene and orthopyroxene phenocrysts (0·1–3·0 mm in
size, although most commonly between 0·5 and 1·5 mm),
typically in about equal proportions (modes vary from 5
to 20·5%, although most commonly in the range of
5–12%; Table 1; Fig. 5e). Pyroxene phenocrysts occur
as single mineral grains and in aggregates, and consist
of a weakly normally zoned core and a thin normally or
reversely zoned rim. A small, but significant, amount
of the single-mineral-grain phenocrysts is resorbed. All
clinopyroxenes are augites, whereas orthopyroxenes
range from En60Fs26Wo4 to En77Fs19Wo4 in prehistoric
lavas and from En73Fs22Wo4 to En68Fs29Wo4 in the 1977
lavas (Table 2; Fig. 7). Phenocrysts with complex oscillatory zoning and orthopyroxene jacketed by clinopyroxene or pigeonite occur in most lavas but in low
abundances.
Two types of phenocrysts can again be recognized—
clear ‘juvenile’ crystals and ‘cognate’ or ‘accessory’ crystals, which are cracked in a similar way to the P2-type
plagioclases, and contain 10–500 m, spherical to ovoid
glass inclusions. Some pyroxene phenocrysts also contained small inclusions of magnetite (± chromite) and
often have iddingsite in cracks and along crystal surfaces
providing evidence of deuteric alteration.
Heterogeneous crystal aggregates (a cluster
or aggregate of crystals consisting of two
or more mineral phases)
Heterogeneous crystal aggregates (HCA) occur in all
White Island lavas, although their modes (1–12%) and
size (1–7 mm) vary considerably (Table 1). HCA consist
of the same mineral phases as make up the single-crystal
phenocryst population (i.e. plagioclase, clinopyroxene,
orthopyroxene ± spinels) and occur as loose- to closepacked aggregates of sub- to euhedral, 0·4–3 mm crystals,
with a variable internal arrangement of mineral phases
(Fig. 8d). The loose-packed aggregates have numerous
angular intercrystal voids that are occupied by either
crystalline groundmass material or pristine brown glass
containing a few dispersed plagioclase and pyroxene
microlites. Sometimes anhedral olivines jacketed with
thick equigranular coronas of orthopyroxene are included
in the plagioclase–clinopyroxene–orthopyroxene HCA.
Aggregates consisting of plagioclase–orthopyroxene and
clinopyroxene–orthopyroxene also occur, but are rare.
A conspicuous group of HCA that occurs in trace
amounts in many lavas consists of equigranular orthopyroxenes (<0·5 mm diameter) with intergrowths of vermicular titanomagnetite. These aggregates are identical
to the coronas found around many partly resorbed olivine
phenocrysts and are undoubtedly of similar origin.
880
COLE et al.
MAGMA ORIGIN AND EVOLUTION OF WHITE ISLAND
‘Clean’ and ‘dirty’ lavas
On the basis of mineralogy and texture, the prehistoric
andesites and dacites can be divided into those that have
predominantly clear euhedral ( juvenile) phenocrysts,
which are referred to subsequently in the paper as ‘clean’
lavas (Fig. 8e), and those that contain dominantly resorbed, sieved or strained phenocrysts, which are referred
to as ‘dirty’ lavas (Fig. 8f ). The 1977 bombs and blocks
and the high-silica andesites and dacites from Central
and Ngatoro cones are predominantly ‘clean’, whereas
most of the other intermediate andesites, for example
from Club Rocks, Troup Head and Ngatoro, are predominantly ‘dirty’.
Enclaves and xenoliths
Microdiorite enclaves, ranging in size from a few millimetres up to >200 mm across (Fig. 8f ), are common in
lavas from the Troup Head and Pinnacle Head suites
(<1 vol. %). A true assessment of the abundance of
enclaves is, however, difficult to ascertain because of the
ubiquitous presence of a red–brown alteration rind on
all lava and breccia surfaces except along the coast.
Porphyritic lava xenoliths are common throughout White
Island lavas, as are gneissose crustal xenoliths or fragments of quartzite. The latter probably represent recrystallized sandstone and quartz veins from the
underlying basement.
extrusion. Systematic analysis of lavas through an established stratigraphic sequence (Fig. 10) suggests a series
of cycles, each culminating with eruption of widespread
high-silica andesite–dacite lava flows. Normalized incompatible trace element plots for White Island andesites
yield typical subduction zone patterns (Fig. 11) with
high large ion lithophile elements (LILE), low high field
strength elements (HFSE) and negative Nb and P anomalies.
A single scoria bomb (Pakihikura scoria; E2) was analysed from the sequence of pyroclastic deposits of North
Bench (erupted <2 ka; T. Thordarson & B. F. Houghton;
unpublished 14C date, 1999) and has significantly lower
SiO2 (56·3 wt %) than the youngest effusive lava of Central
Cone (E3; 62·84–63·48 wt %). The 1977 eruptives (E1)
are generally higher in MgO and are of two compositions:
bombs have >58 wt % SiO2 and 7·5 wt % MgO, and
blocks have >56 wt % SiO2 and 9·3 wt % MgO (Fig.
9; Cole & Graham, 1987). Both types are high in Cr
and Ni (Table 3). Juvenile samples from magmatic eruptions in 1991–1992 (Wood & Browne, 1996) have slightly
lower concentrations of compatible elements and higher
incompatible elements than the 1997 eruptives.
Most incompatible major and trace elements (e.g. K2O,
Ba, Zr) show a generally concordant oxide or element
distribution trend against SiO2, with Ba significantly
higher in the dacites (Fig. 9). Compatible elements (e.g.
MgO, TiO2, Ni) define more than one trend (Fig. 9),
with high MgO and Ni trends defined by those lavas
containing olivine, whereas the high TiO2 trend appears
to relate to those samples having titaniferous augite as a
phenocryst phase.
GEOCHEMISTRY
Major and trace elements
Representative major, trace and rare earth element data
for each formation on White Island are presented in
Table 3. A full list of 128 analyses, together with averages
for each eruption unit, is available from J.W.C. or
the Journal of Petrology web site (at http://www.
petrology.oupjournals.org) All samples from White
Island are medium-K calc-alkaline lavas ranging from
basaltic andesite to dacite (Fig. 9). mg-number ranges
from 42 to 72.
For the prehistoric lavas, the composition of lavas from
Ngatoro (E11–E21) and Central Cone (E3, E4, E6, E7,
E10) overlap on Harker variation diagrams (Fig. 9), and
Club Rocks (E5?) could be part of either cone, but there
is good stratigraphic evidence that they are part of Central
Cone. There are major and trace element differences
between lavas of Central Cone and Troup Head (Fig.
9), but again there is good evidence that Troup Head
lavas (E8, E9) are stratigraphically part of Central Cone.
Volckner Rocks (E22) are most similar chemically to
lavas of Troup Head and Club Rocks, but are a separate
Rare earth elements
Representative REE data for White Island are listed in
Table 3 and a chondrite-normalized REE plot is shown
in Fig. 12. (La/Yb)n ratios vary within a narrow range
(2·2–4·4) and increase with increasing SiO2 (Fig. 13a),
but prehistoric lavas do not show a correlation between
(La/Yb)n and mg-number (Fig. 13b). Although the correlation with SiO2 is suggestive of crystal fractionation,
the absence of a similar correlation with mg-number
may imply a more complex relationship. Samples show
variable Eu anomalies (Fig. 12), reflecting either extraction or accumulation of plagioclase during fractionation, or its retention or preferential separation in
the source region.
Isotopes
881
New isotope ratio data for prehistoric White Island lavas
are given in Table 4. 87Sr/86Sr isotope ratios range
JOURNAL OF PETROLOGY
VOLUME 41
NUMBER 6
JUNE 2000
Table 3: Representative chemical analyses for individual eruption units at White Island
E no.:
E1
E1
E2
E3
E4
E5?
E5
E6
E7
E8
Location:
1977–1982
1977–1982
Central
Central
Central
Central
Central
Central
Central
Troup Head Troup Head Central
Cone
Cone
Cone
Type:
1977 block 1977 bomb Pakihikura
Cone
Cone
Cone
Cone
Mt
SW lava
Club Rocks Otaketake PtOtaketake PtNorth
Gisborne
P41600
TRW-35
SiO2
55·94
58·05
TiO2
0·63
0·61
Al2O3
13·26
13·80
TRW-7
E10
Cone
Wilson Bay Wilson Bay Shark Bay
Bench
scoria fall
Sample:
E9
lava 3
lava 2
TRW-2
TRW-61
TRW-12
basal lava
RW-40
TRW-28
RW-14
RW-2
RW-4
TRW-47
56·59
63·28
60·39
56·97
60·21
61·86
62·10
59·03
57·21
0·65
0·79
0·68
0·59
0·50
0·72
0·72
0·60
0·57
0·61
15·76
14·71
15·41
17·12
15·21
14·52
14·63
15·95
16·40
14·65
58·67
FeO
7·33
7·66
8·01
6·24
7·27
8·03
6·66
6·44
6·20
7·75
8·13
7·50
MnO
0·14
0·14
0·15
0·11
0·13
0·15
0·13
0·11
0·11
0·15
0·13
0·14
MgO
10·14
7·80
5·98
3·35
4·05
3·97
4·44
3·73
3·57
4·41
3·82
6·44
CaO
8·96
8·26
8·47
5·50
6·27
7·90
7·01
5·95
5·79
7·57
7·89
8·47
Na2O
2·39
2·38
2·30
3·22
2·61
2·62
2·82
3·10
2·79
2·60
2·57
2·52
K 2O
1·15
1·36
1·15
2·22
1·46
1·07
1·42
2·09
2·21
1·32
1·06
1·37
P 2O5
0·07
0·08
0·08
0·13
0·09
0·08
0·07
0·11
0·10
0·08
0·08
0·08
LOI
0·00
-0·16
1·38
0·10
1·94
0·16
1·12
0·33
1·88
0·26
0·67
0·03
S
n.a.
Total
100·77
0·01
0·04
99·99
100·56
n.a.
Ba
481
485
388
792
538
473
512
695
701
508
486
Ce
10
12
27
25
32
5
25
24
28
11
16
20
Cl
n.a.
680
1590
n.a.
2080
n.a.
3520
1080
1520
n.a.
n.a.
970
Cr
497
352
151
86
78
27
140
102
89
37
23
243
Co
n.a.
46
46
n.a.
48
n.a.
39
39
29
n.a.
n.a.
35
Cu
90
76
73
n.a.
93
n.a.
47
63
69
n.a.
n.a.
58
Ga
13
14
14
15
17
15
16
15
14
8
13
14
La
4
17
6
18
10
8
4
14
17
11
5
7
Ni
158
115
52
36
25
20
34
39
39
20
19
69
99·67
0·01
100·32
n.a.
98·66
0·03
0·00
0·04
99·62
98·95
100·14
n.a.
99·71
n.a.
98·52
0·00
100·48
473
Nb
2
4
0
4
1
0
0
1
2
0
0
0
Pb
9
15
19
9
15
1
13
20
19
10
37
13
Rb
34
42
33
71
50
35
43
70
72
43
35
45
Sr
169
165
182
171
175
213
183
180
182
195
206
179
V
209
200
218
161
201
199
139
186
177
200
189
193
Y
21
20
21
26
24
19
16
23
25
22
15
22
Zn
67
63
63
54
76
67
60
54
57
56
60
62
Zr
88
98
82
142
102
95
90
134
134
103
96
93
Th
2
2
4
8
n.a.
5
6
7
7
6
4
5
Nd
8
10
10
16
n.a.
11
10
14
16
13
11
10
Sc
29
n.a.
25
22
26
n.a.
26
n.a.
17
n.a.
La
5·2
n.a.
n.a.
13·6
n.a.
Ce
11·0
n.a.
n.a.
29·3
n.a.
n.a.
n.a.
7·98
17·3
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
8·19
17·4
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
Pr
1·2
n.a.
n.a.
n.a.
n.a.
n.a.
Nd
5·1
n.a.
n.a.
13·5
n.a.
8·6
n.a.
n.a.
n.a.
8·6
n.a.
n.a.
Sm
1·4
n.a.
n.a.
3·3
n.a.
2·3
n.a.
n.a.
n.a.
2·2
n.a.
n.a.
Eu
0·48
n.a.
n.a.
1·05
n.a.
0·8
n.a.
n.a.
n.a.
0·83
n.a.
n.a.
Gd
1·8
n.a.
n.a.
3·5
n.a.
2·5
n.a.
n.a.
n.a.
2·5
n.a.
n.a.
Tb
0·3
n.a.
n.a.
0·57
n.a.
0·42
n.a.
n.a.
n.a.
0·41
n.a.
n.a.
Dy
2·2
n.a.
n.a.
3·7
n.a.
2·8
n.a.
n.a.
n.a.
2·8
n.a.
n.a.
n.a.
n.a.
Ho
0·46
n.a.
n.a.
n.a.
n.a.
n.a.
Er
1·3
n.a.
n.a.
n.a.
2·3
n.a.
n.a.
n.a.
1·9
n.a.
n.a.
n.a.
n.a.
1·8
n.a.
n.a.
Yb
1·3
n.a.
n.a.
2·4
n.a.
1·8
n.a.
n.a.
n.a.
1·8
n.a.
n.a.
Lu
n.a.
n.a.
n.a.
0·4
n.a.
0·3
n.a.
n.a.
n.a.
0·3
n.a.
n.a.
Hf
n.a.
n.a.
n.a.
4·3
n.a.
2·3
n.a.
n.a.
n.a.
2·3
n.a.
n.a.
Ta
n.a.
n.a.
n.a.
0·6
n.a.
0·7
n.a.
n.a.
n.a.
0·7
n.a.
n.a.
Th
n.a.
n.a.
n.a.
6·86
n.a.
3·18
n.a.
n.a.
n.a.
3·55
n.a.
n.a.
U
n.a.
n.a.
n.a.
1·62
n.a.
0·84
n.a.
n.a.
n.a.
0·84
n.a.
n.a.
882
COLE et al.
MAGMA ORIGIN AND EVOLUTION OF WHITE ISLAND
Table 3: continued
E no.:
E11
E12
E14
E15
E16
E17
E18
E19
E20
E22
?
?
Location:
Ngatoro
Ngatoro
Ngatoro
Ngatoro
Ngatoro
Ngatoro
Ngatoro
Ngatoro
Ngatoro
Volckner
Central
Ngatoro
Cone
Cone
Cone
Cone
Cone
Cone
Cone
Cone
Cone
Type:
Summit
Te
Te
SW side
SW side
NW Cliff
NW Cliff
NW CliffNW Cliff
Matawiwi
Matawiwi
basal lava
lava 3
lava I
Lava 2
TRW-17
TRW-16
RW-19
RW-22
RW-42
Sample:
RW-34
2
I
TRW-5
TRW-63
Rocks
Cone?
Cone?
south end
NW Cliffs
NW Cliffs
dyke
dyke
RW-20
RW-30
RW-23
RW-26
61·81
SiO2
57·79
61·37
60·18
61·03
62·80
59·10
55·21
61·21
59·40
60·41
63·21
TiO2
0·72
0·55
0·52
0·49
0·52
0·60
0·63
0·56
0·63
0·58
0·69
0·67
Al2O3
16·36
14·62
14·51
14·53
15·03
15·40
15·95
14·76
15·11
16·89
14·96
14·92
FeO
8·20
6·48
6·43
5·18
5·81
7·50
8·70
6·78
7·52
6·99
6·42
6·81
MnO
0·16
0·12
0·13
0·08
0·09
0·20
0·15
0·12
0·10
0·13
0·10
0·13
MgO
3·67
4·94
5·70
3·30
3·95
5·20
6·38
4·45
4·93
3·50
3·47
4·54
CaO
7·14
6·85
6·95
4·65
5·84
7·70
8·69
6·78
6·89
7·65
4·86
6·41
Na2O
2·59
2·78
2·83
2·30
2·81
2·50
2·13
2·68
2·50
2·77
2·98
2·76
K 2O
1·36
1·60
1·49
1·67
1·70
1·40
0·87
1·58
1·44
1·40
1·71
1·75
P 2O5
0·10
0·07
0·06
0·09
0·08
0·10
0·09
0·08
0·09
0·08
0·12
0·11
LOI
0·86
0·28
1·36
5·84
1·46
0·10
0·04
0·27
0·91
−0·13
0·48
-0·48
S
Total
Ba
n.a.
98·95
483
0·00
0·09
1·44
0·10
99·67
100·25
100·60
100·18
568
553
534
574
n.a.
n.a.
99·80
516
98·84
334
n.a.
99·26
591
n.a.
99·52
518
n.a.
100·26
577
n.a.
98·98
630
n.a.
99·42
621
Ce
31
15
26
31
28
26
17
25
30
23
40
17
Cl
n.a.
2310
2920
830
610
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
Cr
44
153
187
134
117
96
116
119
81
15
38
111
Co
n.a.
35
29
53
42
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
Cu
n.a.
33
51
71
42
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
Ga
15
14
15
16
17
16
15
16
16
21
12
15
La
11
10
9
8
11
10
8
14
11
10
15
13
Ni
19
46
57
101
43
30
37
51
33
9
18
33
Nb
2
5
3
0
1
0
0
3
n.a.
0
4
4
Pb
7
14
15
16
13
3
0
19
n.a.
0
4
58
Rb
39
47
45
59
51
35
20
49
39
42
54
35
Sr
210
171
171
145
185
202
206
178
195
217
185
188
V
233
166
141
126
148
207
247
167
188
174
182
168
Y
23
19
19
20
24
22
16
24
20
20
23
27
Zn
76
57
52
78
76
54
68
69
47
59
48
48
Zr
117
103
106
102
115
108
89
125
115
110
160
141
Th
11
6
3
n.a.
n.a.
12
8
4
4
8
4
3
Nd
17
12
15
n.a.
n.a.
10
14
11
17
13
11
14
Sc
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
La
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
Ce
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
20
n.a.
8·76
18·5
n.a.
7·04
15·6
n.a.
9·45
12·5
n.a.
27·3
n.a.
Pr
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
Nd
n.a.
n.a.
n.a.
n.a.
n.a.
8·6
8·2
n.a.
n.a.
9·3
n.a.
12·8
n.a.
n.a.
Sm
n.a.
n.a.
n.a.
n.a.
n.a.
2·3
2·2
n.a.
n.a.
2·3
3·2
n.a.
Eu
n.a.
n.a.
n.a.
n.a.
n.a.
0·83
0·79
n.a.
n.a.
0·84
1·0
n.a.
Gd
n.a.
n.a.
n.a.
n.a.
n.a.
2·5
2·4
n.a.
n.a.
2·4
3·2
n.a.
Tb
n.a.
n.a.
n.a.
n.a.
n.a.
0·43
0·39
n.a.
n.a.
0·39
0·51
n.a.
Dy
n.a.
n.a.
n.a.
n.a.
n.a.
2·7
2·5
n.a.
n.a.
2·6
3·3
n.a.
Ho
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
Er
n.a.
n.a.
n.a.
n.a.
n.a.
1·8
1·6
n.a.
n.a.
1·7
2·0
n.a.
Yb
n.a.
n.a.
n.a.
n.a.
n.a.
1·8
1·6
n.a.
n.a.
1·7
1·9
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
Lu
n.a.
n.a.
n.a.
n.a.
n.a.
0·3
0·2
n.a.
n.a.
0·3
0·3
n.a.
Hf
n.a.
n.a.
n.a.
n.a.
n.a.
2·7
1·7
n.a.
n.a.
2·5
4·0
n.a.
Ta
n.a.
n.a.
n.a.
n.a.
n.a.
0·6
0·7
n.a.
n.a.
0·6
0·5
n.a.
Th
n.a.
n.a.
n.a.
n.a.
n.a.
3·97
2·5
n.a.
n.a.
4·21
6·08
n.a.
U
n.a.
n.a.
n.a.
n.a.
n.a.
0·92
0·56
n.a.
n.a.
0·99
1·44
n.a.
Loss on ignition (LOI) at 1000°C and Fe as total FeO. REE elements and Hf, Ta, U and Th determined by ICP-MS. Samples
are located in Fig. 2. E no., eruption unit number; n.a., element not analysed.
883
JOURNAL OF PETROLOGY
VOLUME 41
NUMBER 6
JUNE 2000
Fig. 9. Selected major and trace element Harker diagrams for White Island.
between 0·70522 and 0·70559 and 143Nd/144Nd isotope
ratios range between 0·512717 and 0·512782 (Fig. 14).
87
Sr/86Sr isotope ratios for White Island lavas lie within
the range of most Ruapehu andesites (Graham et al.,
1995) but are lower than those from Motuhora (Burt et
al., 1996), whereas 143Nd/144Nd isotope ratios tend to be
slightly less radiogenic. Neither isotope system shows
any systematic relationship with SiO2, unlike lavas from
Ruapehu, which show a clear assimilation–fractional
crystallization (AFC) trend (Fig. 15). The 1977 ejecta
have lower 87Sr/86Sr ratios and higher 143Nd/144Nd ratios.
Pb isotope ratios also show limited variation (Fig. 16).
The limited range in Sr and Nd isotopic ratios for
prehistoric White Island lavas indicates that all these
lavas are derived from isotopically similar source regions.
Graham & Cole (1991) identified a slight positive correlation between 87Sr/86Sr and SiO2 for White Island
lavas; however, the larger dataset used in this study does
not support this conclusion. Simple correlations between
isotope ratios and whole-rock chemistry or stratigraphic
succession are not evident. Although the isotopic data
indicate that prehistoric lavas on White Island are remarkably uniform, they differ significantly from the 1977
ejecta (Figs 14 and 16).
884
COLE et al.
MAGMA ORIGIN AND EVOLUTION OF WHITE ISLAND
Fig. 10. Variation in chemistry with stratigraphy on White Island. The E number column on left refers to eruption unit numbers (see Figs 3
and 4, and Appendix), with E1 being the youngest and E21 the oldest. E22 represents Volckner Rocks and its relative age is unknown. These
data represent all the flows exposed at the surface; most have several geochemical analyses taken from a number of points along the flow and
the point displayed represents an average value. Continuous line separates the Central and Ngatoro cone successions; broken lines indicate
breaks between inferred magma evolution cycles. Time elapsed between extrusion of youngest Central Cone units E1–E4 is >1000 years (T.
Thordarson & B. F. Houghton, unpublished data, 1999), but the repose period for the remainder of the Central Cone and Ngatoro Cone lavas
is unknown.
Fig. 11. Representative major and trace element spidergrams for White
Island. MORB normalization values from Sun & McDonough (1989).
Samples from White Island have typical ‘arc’ patterns, and are enriched
in LILE and depleted in HFSE. Analyses by XRF except for U, Ta
and Th, which were analysed by ICP-MS.
DISCUSSION
White Island magma system
The volcanic succession on White Island shows a clear
cyclic pattern of volcanism with major episodes of lava
extrusion, separated by a period of phreatomagmatic or
strombolian explosive eruption such as that producing
Fig. 12. Rare earth element plots for White Island samples. Data for
P41600 from Graham & Cole (1991). Chondrite normalization values
from Sun & McDonough (1989). Analyses by ICP-MS.
the 1976–1992 ejecta. The small volume and limited
dispersal of the deposits of the latter sequences means
that they are poorly preserved in the stratigraphic record.
The length of each cycle is unknown, but new 14C dates
(T. Thordarson & B. F. Houghton; unpublished data,
1999) suggest that the major episodes of lava extrusion
are separated by time intervals of 0·5–1·5 ky.
885
JOURNAL OF PETROLOGY
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JUNE 2000
multiple populations of plagioclase phenocrysts. These
features suggest that magma mixing is a common process
during the evolution of many of the White Island magmas
and, moreover, that mixing events probably occur many
times during the growth of a single phenocryst. Plagioclase, pyroxene and many HCA indicate two phenocryst populations (see modes in Fig. 5a). P1-type
plagioclase phenocrysts are in equilibrium with the melt
and may be regarded as ‘juvenile’. Types P2–6 have
been added to the melt, probably by interaction with
extremely crystal-rich magmas (perhaps a crystal-mush)
within the main magma chamber–plumbing–storage system of the White Island volcano. These ‘dirty’ lavas (see
Fig. 8f ) always have a higher total phenocryst abundance
(see Table 1).
The HCA usually contain both P1 and P2–6 type
plagioclase, but the presence of vesicular, lava groundmass in voids within the loose-packed HCA indicates
that they were incorporated into the erupted magma as
porous loosely bound clusters of crystals that with time
became infiltrated by the surrounding melt. This relationship suggests that they are also derived from the
magma chamber. The mineralogy of the HCA indicates
that most are initially derived from the magma plumbing
system, although some may be directly from the mantle.
The presence and absence of spinel, along with highly
variable order of crystallization, suggests that they originated from a range of depth-intervals within the White
Island plumbing system.
Significance of Cr/Ni ratios
Fig. 13. (a) Plot of (La/Yb) vs SiO2 content. (b) Plot of (La/Yb) vs mgnumber. Motuhora data from Burt et al. (1996); TVZ and Kermadec
data from Gamble et al. (1993b, 1995).
The 1977 eruptive products have highly magnesian
olivine phenocrysts (Fo90–92), and high Cr and Ni, suggesting equilibration in the mantle (Shiraki et al., 1994),
but pyroxene (mg-number 70–83) and plagioclase (An60–75)
compositions are less primitive and suggest at least two
stages of crystallization. Some crystallization may occur
in a magma chamber within the crust (2–7 km? depth),
but geophysical data (Sherburn & Scott, 1988; Houghton
& Nairn, 1989), and localized inflation around the active
crater (Clark & Otway, 1989) all indicate that the 1977
eruptions took place from very shallow levels (<2 km)
within the feeder system, and further crystallization may
have occurred at this level.
Prehistoric lavas on the other hand have variable
phenocryst core–rim compositional variations and
There are significant variations in both Cr and Ni values
at White Island (Fig. 17). Highest Cr values (497 ppm)
occur in the 1977 eruptives (E1) and lowest levels (24
ppm) in the mixed lavas of Troup Head (E9; Table 3).
Ni ranges from 158 ppm in E1 to 16 ppm in E9. This
variation implies strong fractionation. However, there is
little variation in comparable SiO2 or Zr values (Fig.
17a) within the most mafic samples (E1 has 55·9–58·04
wt % SiO2 and 88–89 ppm Zr; E9 has 57·9 wt % SiO2
and 84 ppm Zr). Other formations (e.g. E18: 55·85 wt
% SiO2 and 91 ppm Zr) have similar values. This suggests
source heterogeneity and variable fertility of the source,
as suggested by Gamble et al. (1995) for the southern
Kermadec region and the Ngatoro Basin, and indicates
that magma recharge into the magma chamber may
come from at least two sources in the mantle. The plot
of Cr/Ni vs CaO (Fig. 17b) also suggests two lines of
derivation before fractionation decreases both Cr/Ni
ratios and CaO in the high-silica andesites and dacites.
Origin of White Island high-Mg andesites
886
Slight chemical and petrographic variations between
1977 blocks (largely solidified) and bombs (largely fluid)
10·1
9·0
MgO
CaO
887
481
Zr
Ba
Pb/204Pb
Pb/204Pb
Pb/204Pb
206
207
208
—
—
0·512800∗
6·0
—
—
—
—
—
466
85
51
145
180
36
62·5
1·3
8·4
∗Isotope data from Graham & Cole (1991).
Nd/144Nd
143
—
88
Ni
0·70511∗
158
Cr
Sr/86Sr
497
Sr
87
34
169
Rb
73·7
mg-no.
1·2
13·3
Al2O3
15·6
0·7
38·71±3
15·62±1
18·83±1
0·512742±5
0·70541±1
792
142
36
86
171
71
59·1
2·2
5·5
3·4
14·7
0·8
63·3
Cone
56·3
Central
scoria
0·6
K 2O
E3
RW40
Pakihikura
55·9
Locality:
E2
RW64
SiO2
1977 block
Sample:
TiO2
E1
P41600
E no.:
38·78±3
15·65±1
18·83±1
0·512756±5
0·70528±2
473
95
20
27
213
35
56·9
1·1
7·9
4·0
17·1
0·6
57·0
Club Rocks
RW14
E5?
38·71±3
15·63±1
18·82±1
0·512749±5
0·70538±1
508
103
20
37
195
43
59·8
1·3
7·6
4·4
16·0
0·6
59·0
Troup Head
RW2
E8
—
—
—
—
0·70536±1
515
104
21
71
191
56
60·2
1·3
7·6
4·3
15·7
0·6
59·1
Troup Head
RW3
E8
38·77±3
15·64±1
18·84±1
0·512776±5
0·70527±1
516
108
30
96
202
35
64·6
1·4
7·7
5·2
15·4
0·6
59·1
Cone
Ngatoro
RW19
E17
—
—
—
0·512756±7
0·70542±1
334
89
37
116
206
20
66
0·9
8·7
6·4
16·0
0·6
55·2
Cone
Ngatoro
RW22
E18
—
—
—
—
0·70547±1
577
110
9
15
217
42
56·9
1·4
7·7
3·5
16·9
0·6
60·4
Rocks
Volckner
RW30
E22
38·77±3
15·64±1
18·84±1
0·512732±6
0·70548±1
630
160
18
38
185
54
59
1·7
4·9
3·5
15·0
0·7
63·2
Cone
Central
RW23
?
Table 4: New isotope analyses from White Island (isotopic data by mass spectrometry on a VG54E), with representative major and trace
element data for the same samples
COLE et al.
MAGMA ORIGIN AND EVOLUTION OF WHITE ISLAND
JOURNAL OF PETROLOGY
VOLUME 41
NUMBER 6
JUNE 2000
Fig. 14. 143Nd/144Nd vs 87Sr/86Sr. Data for 1977 eruptives are from
Graham & Cole (1991); Motuhora data from Burt et al. (1996), Ruapehu
data from I. J. Graham (unpublished data, 1995). Inset: TVZ and
Kermadec basalt data from Gamble et al. (1993b, 1995); TVZ rhyolite
data from McCulloch et al. (1994) and Sutton et al. (1995), and data
for Waipapa and Torlesse metasediments from Graham et al. (1992)
and McCulloch et al. (1994).
Fig. 16. Pb isotope variation diagram for White Island lavas. There
is minimal overlap with published TVZ analyses although some TVZ
basalts are also displaced from the mixing line between Kakuki basalt
and Torlesse metasediment. Sources for published data as for Fig. 14.
Fig. 15. (a) 87Sr/86Sr vs SiO2. The ‘flat’ trend of White Island prehistoric
lavas when compared with the ‘AFC’ style trend of Ruapehu should
be noted. (b) 143Nd/144Nd vs SiO2. The lack of overlap with the adjacent
Motuhora volcano and the limited variation in the data should be
noted. Sources for published data as for Fig. 14.
led Clark & Cole (1986) to suggest that the 1977 nearsurface magma chamber comprised a molten core surrounded by a semi-rigid carapace. Eruption of the molten
core generated lava bombs and fragmented the carapaceforming blocks. Clark & Cole (1986) further suggested
that the lava bombs, which are slightly less mafic than
the blocks, were derived from the carapace magmas,
after fractionation of forsteritic olivine. Shiraki et al. (1994)
considered the 1977 eruption products were formed by
mixing of high-Mg basaltic andesite and dacite, with the
dacite supplying both plagioclase and low-Mg pyroxene
phenocrysts, but there is little evidence for such extensive
mixing.
Any model for the origin of these high-Mg andesites
must account for:
(1) high LILE/LREE ratios (Fig. 18);
(2) primitive character, e.g. mantle-like olivine composition, high MgO, Ni and Cr;
(3) isotopic compositions, indicative of crustal contamination despite (2);
(4) generally low HFSE (Fig. 11), characteristic of a
depleted source (Woodhead et al., 1993).
The model must also consider experimental studies
(e.g. Tatsumi, 1982), which indicate that high-Mg andesite cannot be generated from basalt by either fractional
crystallization or partial melting, although Kelemen
(1995) suggested the possibility of partial melting of a
888
COLE et al.
MAGMA ORIGIN AND EVOLUTION OF WHITE ISLAND
Fig. 18. La vs Ba for White Island. White Island lavas fall along the
same trend as those from Bay of Plenty andesites. TVZ andesite field
includes Tongariro, Ruapehu, Rotokawau and Taupo caldera complex.
Data from Graham & Hackett (1987), Browne et al. (1992), Gamble et
al. (1993b), Sutton (1995), Burt et al. (1996), Hobden (1997), Cole et al.
(1998) and I. J. Graham (unpublished data, 1995).
Fig. 17. (a) Cr vs Zr for White Island data. (b) Cr/Ni vs CaO,
illustrating the possible end-members involved in magma evolution at
White Island Volcano. Symbols for samples are the same as in Fig. 9.
depleted, but subsequently enriched (metasomatized),
source for the high-Mg andesites. These high-Mg magmas
can, however, be generated at the slab–mantle interface
as small volume partial melts combined, which ascend
rapidly toward the surface.
Graham et al. (1992) suggested that the 1977 ejecta of
White Island have boninitic affinities, whereas Shiraki et
al. (1994) compared the same White Island lavas with
high-Mg andesites of the Setouchi volcanic belt in SW
Japan. The 1977 ejecta are not boninitic in terms of the
classification of Crawford et al. (1988), but are similar to
the Setouchi suite (Tatsumi, 1982). Both are LREE
enriched with flat HREE in contrast to the V-shaped
REE pattern of boninites, they have low HFSE/LILE
ratios and occur in a continental setting. Tatsumi (1982)
showed that high-Mg magmas of Setouchi had equilibrated with mantle olivine plus two-pyroxene assemblages at 1050–1150°C and 10–15 kbar pressure
under water-rich conditions. Experiments by Wolf &
Wyllie (1994) and Sisson & Grove (1993) have yielded
melts similar to the Setouchi high-Mg magmas at high
f O2 and PH2O. In contrast, experimental studies of highAl basalt with MgO >8% have shown that such magmas
can be produced by small-scale anhydrous partial melting
of plagioclase or spinel lherzolite at 10 kbar (e.g. Falloon
& Green, 1987; Bartels et al., 1991). These lines of
evidence suggest that high-Al basalts equilibrate at relatively low pressure and not at the slab–mantle interface,
whereas high-Mg magmas are generated at deeper levels
in the melt column.
High-Mg andesites are frequently considered to be the
result of either high heat flow in back-arc ocean basins
or the subduction of young oceanic crust (e.g. Kelemen,
1995). The subducting slab in the Taupo arc is, however,
Cretaceous in age (Mortimer & Parkinson, 1996) and
therefore relatively cold compared with younger crust
(Peacock, 1990; Gamble et al., 1996). The overlying
mantle wedge is also likely to be cooler.
Modelled P–T profiles of the mantle wedge also support
a lower temperature than suggested by petrological data
(350–500°C at 60 km depth). Thermal modelling by
Davies & Stevenson (1992) showed that maximum temperatures of 1100°C are attained at 90 km depth, but
Shiraki et al. (1994) have suggested that the 1977 magmas
attained temperatures of >1200°C. One possible mechanism to account for this conflict between petrologic
and modelled P–T profiles was proposed by Davies &
Stevenson (1992). They suggested that small degree partial melts, generated in and just above the slab, rise into
the overlying mantle through stress-induced fractures.
During ascent, the melt undergoes isenthalpic decompression heating (up to 230°C above the geotherm),
889
JOURNAL OF PETROLOGY
VOLUME 41
leading to further melting of the mantle, which reacts
with the ascending magma at <60 km depth. The composition of this second melt is controlled by the water
content of the initial melt. Reaction occurs between the
two melts during which the initial melt is consumed.
This occurs in a zone where experimental data indicate
that high-Al basalts are generated. High-Mg magmas
must, therefore, rise relatively rapidly through the mantle
or they would react in the upper-mantle wedge to form
high-Al basalt.
Our preferred model for the Mg-rich andesites at
White Island is that they were initially generated as highMg magmas by hydrous melting of mantle, metasomatized by fluids from the dehydrating slab at or near
the slab–mantle wedge interface, and then rise rapidly
through the mantle wedge and lower crust to relatively
shallow magma chambers (2–7 km? depth), where limited
mixing and crustal contamination occur before eruption.
The juvenile samples collected from phreatomagmatic
eruptions in 1991–1992 (Wood & Browne, 1996) have
slightly lower concentrations of compatible elements and
higher levels of incompatible elements than material
erupted on 24 March 1977 (Graham & Cole, 1991; this
study). This is consistent with additional fractionation of
clinopyroxene ± olivine in the same magma chamber
that produced the 1977 ejecta.
Origin of prehistoric andesite and dacite
Central Cone lavas were modelled by Graham & Cole
(1991) as AFC derivatives from a parental magma comparable with 1977 ejecta and contaminated by Torlesse
metasediments. Those workers also modelled lavas from
Ngatoro cone which, they suggested, were the result of
plagioclase–olivine and orthopyroxene–augite–magnetite
(POAM) fractionation from a parental low-Al basaltic
magma. Given the mineralogical evidence, the variation
in chemistry within each volcanic unit, the degree of
chemical overlap between units and the stratigraphic
control on sample collection, we suggest that new magmas
entered the magma column frequently throughout the
eruptive history of White Island, as implied by stratigraphic variation shown in Fig. 10, and similar to the
process recently proposed for Ruapehu lavas by Gamble
et al. (1999).
The prehistoric andesites and dacites at White Island
are likely to be derived from similar sources in the mantle
wedge to the high-Mg magmas, but with melts spending
longer in the crustal magma chamber where fractionation
and/or crustal contamination occurred. It seems likely
that small batches of magma rose intermittently to highlevel chambers, where further mixing and crustal contamination may have occurred, obscuring many details
of the earlier processes.
NUMBER 6
JUNE 2000
Crustal contamination
The role of sediment subduction, as proposed by Gamble
et al. (1996), is hard to evaluate at White Island. Ba/La
ratios are generally higher than those of onshore TVZ
andesites (Fig. 18), which is consistent with sediment
input and/or a high slab fluid flux into the mantle wedge.
Differences between isotope values for 1977 ejecta and
prehistoric lavas, but limited range within each group
(Figs 14 and 16), tend, however, to argue against it. Also
the apparent break in the compositional range for Ba vs
SiO2 between andesite and dacite (Fig. 9) makes it more
likely that Ba variation is a result of processes occurring
during the later stages of evolution of the magma than
initial partial melting.
Contamination of mantle-derived magmas by overlying
crust is easier to constrain. There are four possible
contaminants in the crust under TVZ: (1) Torlesse metasediments; (2) Waipapa metasediments; (3) rhyolite–
ignimbrite volcanics; (4) an isotopically enriched lower
crust, identified through granulite xenoliths entrained in
Ruapehu lavas (Graham et al., 1990). The distribution
of Torlesse and Waipapa metasediments beneath TVZ
is unclear, but Waipapa metasediments crop out west of
TVZ and Torlesse metasediments form the axial ranges
to the east. Offshore, Waipapa-like metasediments have
been dredged from the Colville Knolls, some 160 km
NNW of White Island (Gamble et al., 1993a), and confirm
the existence of continental crust to the north of White
Island. Beetham & Watters (1985) placed the boundary
between Waipapa and Torlesse to the east of the Tongariro Volcanic Centre, but metasedimentary xenoliths
from Ruapehu and Tongariro have unequivocal Torlesse
affinities (Graham, 1987). Furthermore, Mortimer (1995)
and Burt et al. (1996) concluded that the Torlesse–
Waipapa boundary lies within TVZ, but the boundary
is unlikely to be a simple linear suture and Torlesse and
Waipapa metasediments are most likely to be interleaved
tectonically beneath TVZ.
Neither the Waipapa metasediments nor the rhyolite–
ignimbrite volcanics are thought to be major contaminants of andesitic magmas in TVZ. Both have similar
isotopic compositions to White Island lavas, so large
volumes of contaminant would have to be incorporated
into the melt to generate the observed isotopic ratios,
and this is not supported by major and trace element
chemistry. Interstitial melt compositions in granulite
xenoliths from Ruapehu are comparable with the modelled crustal contaminant proposed for White Island lavas
by Graham & Cole (1991), but no granulite xenoliths
have been observed in White Island lavas, although
crustal gneisses have been found. Results from isotopic
modelling of TVZ basalts and andesites suggest that
Torlesse metasediments underlie at least part of the TVZ
(e.g. Graham et al., 1992; Gamble et al., 1993b; Burt et
890
COLE et al.
MAGMA ORIGIN AND EVOLUTION OF WHITE ISLAND
Fig. 19. Schematic representation of magma evolution at White Island. (a) Steady-state activity during periods of quiescence, creating phreatic
(fumarolic) eruptions. (b) Activity when system is recharged with high-Mg magma creating phreatomagmatic–strombolian eruptions like those
of 1977–1992. (c) Activity when magma recharge is at a decreasing rate and ‘dirty’ crystal-rich magma rises to the surface to form andesite
lavas. (d) Activity when magma recharge is steady state, but ‘clean’ magma rises to the surface to form high-silica andesite–dacite lavas.
al., 1996) and the White Island isotopic data plot between
TVZ basalt and Torlesse metasediment fields (Figs 14
and 16).
Pb isotope data from both historic and prehistoric lavas
from White Island volcano are close to the AFC trend
(Fig. 16) of Graham et al. (1992). This line is calculated
using Kakuki basalt, considered by Gamble et al. (1990)
to be the least contaminated of the TVZ basalts, and
average Torlesse metasediment as the contaminant. Although the trend of White Island data is parallel to that
of the calculated line, the data are displaced to higher
207
Pb/204Pb. This may be because the Kakuki basalt has
undergone significant crustal contamination and does
not have mid-ocean ridge basalt (MORB)-like isotopic
ratios, but it is more likely that White Island isotopic
characteristics are derived from the characteristics of
the magma sources (i.e. with a contribution from slab,
metasomatized asthenospheric mantle and lithospheric
mantle).
Evolution of White Island magma system
Chemical variation with time on White Island (see Fig.
10) suggests there are systematic changes in magma
composition, which relate directly to the situation in the
sub-volcanic plumbing system: how it responds to changes
in the long-term flux of mafic magma into the system
and how easy it is for the melt fraction to migrate up
through the conduit system at any one time (Fig. 19).
During periods of quiescence, there may be little magma
recharge into the magma chamber and only occasional
slugs or pockets of magma migrate through the conduit
system to accumulate beneath a viscous plug in the vent,
capped by crater-fill debris (Fig. 19a). This accounts
for the ‘normal background’ phreatic activity currently
experienced by White Island. If magma recharge increases, slugs of high-Mg magma pass fairly rapidly
through the magma chamber and rise directly to the
surface to form the high-Mg basaltic andesite ejecta, like
that erupted in 1977–1992 (Fig. 19b). Much of the
891
JOURNAL OF PETROLOGY
VOLUME 41
increased magma recharge will, however, remain in the
magma chamber, where it will interact with the preexisting crystal mush. In time, volatile-charged andesite–
dacite melt will be driven off by filter-pressing (e.g. Sisson
& Bacon, 1999) or by its own buoyancy (Fig. 19c). This
will form the petrographically heterogeneous ‘dirty’ lavas
like those of Troup Head and Club Rocks. As further
time passes the relative proportion of crystal mush decreases and interstitial liquid increases until a new critical
point is reached and the melt fraction will again move
upwards to erupt as a new major episode of high-silica
andesite–dacite lava extrusion (Fig. 19d), forming flows
such as those of Central Cone.
The proportion of mafic component in the system is
gradually increasing with time, and so is the heat flux.
Eventually the system may reach a point when it is
relatively easy for the high silica andesite–dacite melts to
migrate rapidly to the surface. Removal of this magma
will enhance the pressure gradient within the plumbing
system and eventually allow a new injection of parental
magma from the mantle wedge. The sequence is thus
cyclic.
NUMBER 6
JUNE 2000
and comments on this project, and the three referees—
John Wolff, John Gamble and Jon Davidson—for their
help in substantially improving the paper. Transport to
and from White Island was provided by John Baker in
the Ma Cherie, Robert Fleming of Vulcan Helicopters
and No. 3 squadron RNZAF. Yosuke Kawachi is thanked
for assistance in running the microprobe in Otago University, and Pieter Vroon for help with isotopic analysis
at Royal Holloway College, University of London. We
also thank Gregg Kurras for access to side-scan sonar
data and for providing chemical analysis of some of the
submarine lava samples. Funding for this research was
provided by the Foundation for Research, Science and
Technology through PGSF Contracts UOC 314 and
608, and by IGNS Post-doctoral Research Fellowship to
T. Thordarson.
CONCLUSIONS
(1) There is a clear pattern of volcanism over time on
White Island, with major episodes of lava extrusion each
probably separated by periods of phreatomagmatic or
strombolian volcanism.
(2) The White Island magmas probably all originate
from enriched (metasomatized) mantle at or just above
the slab–mantle wedge interface; the fluids come from
the dehydrating slab. There is, however, good evidence
for at least two separate source locations. Partial melts
rise to a magma chamber within the crust where mixing
and limited crustal contamination occurs.
(3) High-Mg magmas rise rapidly from the main
magma chamber as slugs or pockets of magma, which
geophysical data suggest are <2 km beneath the active
vents.
(4) Prehistoric andesite and dacite lavas of White Island
are also the products of eruptions from the same magma
chamber, but have experienced different processes. The
‘dirty’ lavas with inherited phenocrysts result from interaction of the magma recharge with pre-existing crystal
mush in the magma chamber. The ‘clean’ high-silica
andesites–dacites are a product of filter-pressing of a
more fractionated volatile-rich melt fraction.
(5) Such multi-magma plumbing systems probably
occur under most andesite–dacite arc volcanoes.
ACKNOWLEDGEMENTS
We thank John Gamble, Bruce Houghton, Dick Price,
David Shelley and Tod Waight for the exchange of ideas
892
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COLE et al.
MAGMA ORIGIN AND EVOLUTION OF WHITE ISLAND
APPENDIX
Appendix: White Island eruptive rocks: eruption unit (E) number, location and sample numbers
E no.
Location
Samples (this study)
E1
1977 block in Main Crater
26592, P41600 (from Cole & Graham, 1987)
E1
1977 bomb in Main Crater
TRW34, TRW35
E2
Pakihikura scoria fall deposit
TRW7, RW64
E3
Mt Gisborne lava
TRW24
E3
SW-slope lava III
TRW25, TRW27, TRW41, RW38, RW39, RW40
E3
Ngatoro islet lava (unnamed sea stack)
TRW6, RW27
E3
Ngatoro islet lava (submerged part)
TRW59, RW60
E3
Central Cone North Rim lava
TRW13, TRW14, TRW18, TRW19, TRW20, TRW21
E4
SW-slope lava II (welded air-fall phase)
TRW28, TRW29
E4
SW-slope lava II (lava phase)
TRW26
E5?
Club Rocks lava
TRW42, TRW43, RW13, RW14, RW15
E5
Otaketake Point lava II
TRW2, RW63
E5
SW-slope lava I (welded air-fall deposit)
TRW39, TRW40
E6
Otaketake Point lava I
TRW1, TRW3, RW62, RW66
E6
Otaketake Point lava I (submerged part)
TRW55, TRW61, TRW62
E6
Central Cone–Ngatoro Saddle lava
TRW15
E7
Crater Bay lava II
TRW33, RW52, RW53, RW54, RW56
E7
North Bench lava
TRW8, TRW9, TRW10, TRW11, TRW12
E7
North Point lava
TRW22, TRW23, RW16, RW17, RW18
E7
North Bench lava (submerged part)
TRW65
E8
Crater Bay lava I
TRW37, RW55
E8
Upper Troup Head lava
TRW32, RW9, RW10, RW11, RW57
E8
Upper Troup Head lava (submerged part)
TRW54
E8
Pinnacle Head lava
TRW31, RW1, RW2, RW3
E9
Lower Troup Head lava
TRW36, RW4, RW5, RW12
E10
Shark Bay lava
TRW30, RW50, RW51
E10
Shark Bay lava (submerged part)
TRW46, TRW47, TRW50
E11
Ngatoro Summit lava = Ngatoro Te Matawiwi
lava III?
RW34, RW35
E12
Ngatoro Te Matawiwi lava II
TRW4, TRW5, RW65, RW61
E13
Ngatoro North Cliff lava V = Ngatoro West Cliff
lava VII
Not sampled
E14
Ngatoro South Shore lava
TRW38
E14
Ngatoro Te Matawiwi lava I
RW28
E14
Ngatoro Te Matawiwi lava I (submerged part)
TRW63
E15
Ngatoro North Cliff lava III
TRW17
E16
Ngatoro North Cliff lava II
TRW16
E17
Ngatoro West Cliff lava IV
RW19
E18
Ngatoro West Cliff lava III
RW21, RW22
E19
Ngatoro North Cliff lava I
RW42, RW43, RW44
E20
Ngatoro West Cliff lava II
RW20
E21
Ngatoro West Cliff lava I
RW24A, RW24B, RW25
E22
Volckner Rocks lava
RW29, RW30, RW31, TRW44
895