JOURNAL OF PETROLOGY
VOLUME 51
NUMBER 4
PAGES 823^846
2010
doi:10.1093/petrology/egq002
Missing Oligocene Crust of the Izu^Bonin Arc:
Consumed or Rejuvenated During Collision?
YOSHIHIKO TAMURA1, OSAMU ISHIZUKA2, KAN AOIKE3,
SHINICHI KAWATE4, HIROSHI KAWABATA1, QING CHANG1,
SATOSHI SAITO1, YOSHIYUKI TATSUMI1, MAKOTO ARIMA5,
MASAKI TAKAHASHI6, TATSUO KANAMARU6, SHUICHI KODAIRA1
AND RICHARD S. FISKE7
1
INSTITUTE FOR RESEARCH ON EARTH EVOLUTION (IFREE), JAPAN AGENCY FOR MARINE^EARTH SCIENCE AND
TECHNOLOGY (JAMSTEC), YOKOSUKA 237-0061, JAPAN
2
INSTITUTE OF GEOSCIENCE, GEOLOGICAL SURVEY OF JAPAN/AIST, TSUKUBA 305-8567, JAPAN
3
CENTER FOR DEEP EARTH EXPLORATION (CDEX), JAPAN AGENCY FOR MARINE^EARTH SCIENCE AND
TECHNOLOGY, YOKOHAMA 236-0001, JAPAN
4
MUSASHI HIGH SCHOOL, TOKYO 176-8535, JAPAN
5
GRADUATE SCHOOL OF ENVIRONMENT AND INFORMATION SCIENCES, YOKOHAMA NATIONAL UNIVERSITY,
YOKOHAMA 240-8501, JAPAN
6
INSTITUTE OF NATURAL SCIENCES, NIHON UNIVERSITY, TOKYO 156-8550, JAPAN
7
SMITHSONIAN INSTITUTION, NMNH MRC-119, WASHINGTON, DC 20013-7012, USA
RECEIVED JULY 6, 2009; ACCEPTED JANUARY 7, 2010
ADVANCE ACCESS PUBLICATION FEBRUARY 6, 2010
The 50 Myr old Izu^Bonin^Mariana (IBM) arc consists
mostly of Oligocene middle and lower crust that underlies the upper
crust; these units are in turn covered by Quaternary volcanic rocks.
Seismic imaging, forearc geology, Ocean Drilling Program drilling
and magnetic anomalies suggest that most IBM arc crust was created in Eocene^Oligocene times. However, remnants of this old
crust have never been found at the northern end of the arc, where it
is colliding with the Honshu arc (Izu collision zone).Two batholiths
in this collision zone (the Tanzawa tonalites and the Kofu Granitic
Complex) were emplaced during the Miocene (4^17 Ma). Major
elements, Zr/Y, rare earth element ratios and normalized abundance
patterns, and Sr^Nd isotopic data indicate that these plutonic
bodies are compositionally similar to the Oligocene IBM volcanic
rocks, and that they are dissimilar to the Miocene, Pliocene and
Quaternary IBM lavas and volcaniclastic rocks. We suggest that
the Miocene plutonic rocks in the Izu collision zone were derived
from partially melted Oligocene middle crust. A model is proposed
Corresponding author. Telephone: þ81-46-867-9761.
Fax: þ81-46-867-9625. E-mail: tamuray@jamstec.go.jp
in which IBM arc middle crust in the collision zone was partially
melted during the collision and then intruded into the overlying
upper crust of the Honshu and IBM arcs. This resulted in the complete loss of chronological information related to their original source.
WORDS: collision zone; granite; IBM arc; Oligocene;
remobilization; tonalite
KEY
I N T RO D U C T I O N
The Pacific plate began subducting beneath the Philippine
Sea plate about 50 Myr ago to produce the currently
active Izu^Bonin^Mariana (IBM) arc. This 50 Myr old
subduction system contains remnant arcs, such as the
Kyushu^Palau Ridge, West Mariana Ridge and Bonin
Ridge, and extinct spreading centers, such as the Shikoku,
ß The Author 2010. Published by Oxford University Press. All
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JOURNAL OF PETROLOGY
VOLUME 51
Parece Vela and West Philippine Basins, as well as the
active Mariana Trough back-arc basin (Stern et al., 2003;
Tatsumi & Stern, 2006) (Fig. 1). Here, we present a new
petrological and geological model based on recent geophysical observations of the seismic crustal structure in
the northern IBM arc, which we hereafter refer to as the
Izu^Bonin arc. Seismic imaging suggests that most of the
present Izu^Bonin arc crust was created in Eocene^
Oligocene times (Kodaira et al., 2008); these older rocks
are mostly buried and make up much of the infrastructure
(middle and lower crust) of the IBM arc. However, remnants of this older crust have never been found in the Izu
collision zone. The absence of Eocene^Oligocene Izu^
Bonin arc crust in this zone is unexpected, unexplained,
and is the focus of this study. Here we compare the major
element, trace element, rare earth element (REE) patterns
and Sr- and Nd-isotope compositions of Oligocene,
Miocene, Pliocene and Quaternary volcanic rocks and
the collision zone tonalites. We integrate these results with
recent geophysical images of the arc and conclude that
the Miocene plutonic rocks in the Izu collision zone were
derived from the Oligocene middle crust, which was partially melted, remobilized and rejuvenated during the collision. Melts and partially melted masses of the
remobilized middle crust rose buoyantly to form the
Miocene tonalite and granitic plutons that crop out today.
Similar processes may have operated in other collision
zones.
O L I G O C E N E M I D D L E C RU S T I N
T H E I Z U ^ B O N I N A RC
Recent studies have demonstrated that most of the submarine exposures in the IBM forearc are Oligocene to
Eocene in age (Ishizuka et al., 2006). Moreover, most of
these rocks have differentiated from mantle melts to varying degrees, implying that plutonic rocks from this era
probably make up most of the crust. The abundance of
Oligocene volcaniclastic deposits drilled in the forearc
[Ocean Drilling Program (ODP) sites 787, 792 and 793,
Fig. 1] also suggests the important role that Oligocene magmatism played in the evolution of the Izu^Bonin arc
(Hiscott & Gill, 1992; Gill et al., 1994).
Magnetic anomalies and crustal structure
Yamazaki & Yuasa (1998) recognized three conspicuous
north^south-trending rows of long-wavelength magnetic
anomalies along the Izu^Bonin arc, oriented slightly oblique to the present volcanic front (Fig. 1). The easternmost
anomalies correlate with frontal arc bathymetric highs,
such as the Omachi Seamount; the westernmost anomalies
coincide with the Kyushu^Palau Ridge (the remnant arc);
and the central anomalies lying near 1398E cross linear
arrays of Miocene volcanoes (Fig. 1). Yamazaki & Yuasa
NUMBER 4
APRIL 2010
attributed all three magnetic anomalies to loci of
Oligocene magmatic centers and suggested that the magnetic anomalies were caused by induced magnetization
associated with mafic Paleogene plutonic bodies constituting the middle to lower crust.
Full-crustal velocity profiles for the IBM arc (Suyehiro
et al., 1996) define continuous layers extending 200 km
across the arc (Figs 1 and 2b). Kodaira et al. (2007a,
2007b, 2008) conducted several active source wide-angle
seismic studies in the northern part of the Izu^Bonin arc;
these profiles extended for 1050 km along the volcanic
front and for 500 km along the rear-arc, 150 km west of
the volcanic front (Figs 1 and 2a). Kodaira et al. (2007a,
2007b) documented systematic crustal variations beneath
the volcanic front (wavelengths of 80^100 km), defined by
average crustal velocities and the thickness of felsic- to
intermediate-composition middle crust [P-wave velocity
(Vp) of 6·0^6·8 km/s] (Fig. 2a). Kodaira et al. (2008) also
showed undulating crustal thicknesses crossing Miocene
volcanic cross-chains of the Izu^Bonin rear-arc, 150 km
west of the present-day volcanic front (Figs 1 and 2).
Crustal thicknesses are as great as 25^30 km beneath the
rear-arc and have along-arc wavelengths of about 100 km.
Interestingly, the crust underlying the rear-arc thickens
and thins, but these variations are not related to the surface volcanoes, in contrast to the situation along the volcanic front (Kodaira et al., 2007a; Tamura et al., 2009).
Kodaira et al. (2008) suggested that most of the thick crust
in rear-arc areas was created in Eocene^Oligocene times,
before the Shikoku Basin began to open. Moreover, the
undulating pattern of total- and middle-crust thicknesses
and the variations of average seismic velocity, reflecting
the bulk composition of the crust, form patterns similar to
those along the present-day volcanic front. Three discrete
thick crustal segments (20^25 km thick) in the rear-arc,
and possible counterparts beneath the volcanic front
(Kodaira et al., 2008), are shown in Fig. 2a. The crust
beneath the volcanic front is thicker than that of the rear
arc, possibly because of Quaternary magmatism (Kodaira
et al., 2007a, 2007b; Tatsumi et al., 2008; Tamura et al.,
2009) and/or a greater magmatic production rate beneath
the Eocene^Oligocene volcanic front. Taylor (1992) interpreted the frontal arc highs (50 km east of the present
volcanic front) to be arc volcanoes of the Oligocene volcanic front. However, the similarities in crustal variation patterns beneath the volcanic front and the possible paleoarc
(rear-arc) crust suggest that many parts of the crust and
its undulating pattern beneath the volcanic front might
have been also created before the Miocene opening of the
Shikoku Basin. The seismic profile lies 20^50 km west of
the magnetic anomaly along the rear arc (Fig. 1), which
shows five distinct magnetic highs (see Kodaira et al.,
2008, fig. 8). Kodaira et al. (2008) found good correlation
between the seismic velocity image and the arrangement
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TAMURA et al.
IZU^BONIN MISSING OLIGOCENE CRUST
Fig. 1. Bathymetric features of the eastern Philippine Sea, including the Izu^Bonin^Mariana (IBM) arc system. Old seafloor (135^180 Ma) of
the western Pacific plate subducts beneath the active IBM arc at the Izu^Bonin^Mariana trenches. Spreading centers are active in the
Mariana Trough (7^0 Ma) and relict in the Shikoku and Parece Vela Basins (30^15 Ma) and West Philippine Basin (50^35 Ma). The
Ogasawara Plateau, Amami Plateau, Daito and Oki-Daito ridges are Cretaceous^Eocene features. The Kyushu^Palau Ridge (KPR) marks
the rifted western edge of the initial IBM arc system (50^30 Ma), subsequently separated by back-arc spreading into the Shikoku and Parece
Vela Basins. The black dashed lines show the locations of the wide-angle seismic profiles across the arc (Suyehiro et al., 1996), along the
present-day volcanic front (Kodaira et al., 2007a, 2007b) and along the rear-arc 150 km west of the volcanic front (Kodaira et al., 2008) (see
Fig. 2). The lines of white circles show three conspicuous north^south rows of long-wavelength magnetic anomalies identified by Yamazaki &
Yuasa (1998). Numbered boxes: 1, Miocene plutons (Kofu Granitic Complex and Tanzawa tonalites); 2, Oligocene pluton (Komahashi^Daini
Seamount); 3, 4, Oligocene^Miocene turbidites recovered during ODP Legs 125 and 126 (ODP 787, 792 and 793); 5^7, Oligocene volcanic
rocks (Omachi Seamount, Saipan, Rota, Guam and Palau Islands).
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VOLUME 51
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Fig. 2. (a) Three-dimensional block diagrams, bounded by the dashed lines in Fig. 1, showing seismic profiles across the volcanic front and the
rear arc 150 km west of the volcanic front (after Kodaira et al., 2008). Numbered circles indicate sites drilled during ODP Legs 125 and 126,
which recovered Oligocene and Neogene turbidites. Abbreviations show basalt-dominant Quaternary volcanoes (Mi, Miyake; Ha, Hachijo;
Ao, Aogashima; Su, Sumisu; To, Torishima) on the volcanic front and the andesitic Oligocene volcano (Om, Omachi Seamount) east of the
front. The stars on the rear-arc profile indicate Miocene^Pliocene volcanoes. Three discrete thick crustal segments (20^25 km thick) in the
rear-arc and their possible counterparts below the volcanic front (Kodaira et al., 2008) are numbered 1^3. (b) Schematic across-arc profile
(P-wave velocity) of the Izu^Bonin arc along the dashed line in Fig. 1 after Suyehiro et al. (1996). The 6·0^6·3 km/s, 7·1^7·3 km/s and 7·8 km/s
layers correspond to parts of middle crust, lower crust and upper mantle, respectively.
of magnetic highs; the strong magnetic highs lie immediately east of the thick crustal segments. The undulating
crustal structure coincides with magnetic anomalies along
the Izu^Bonin arc, suggesting the possible existence and
volumetric importance of Oligocene crust.
Quaternary rhyolites as probes into the
middle crust
Felsic magmas can originate from two end-member processes: by fractional crystallization of basaltic or andesitic
parents, or by partial melting of basaltic, andesitic, or sedimentary sources. Several lines of evidence in the Izu^
Bonin arc argue against fractional crystallization (see
Tamura et al., 2009, for details). Quaternary rhyolites may
be useful probes into the Izu^Bonin middle crust, if they
formed by the melting of such crust. Three chemical varieties of Quaternary rhyolite have been recognized by
Tamura et al. (2009) along the Izu^Bonin arc front; these
are closely related to volcano type and crustal structure.
R1 rhyolites are erupted from basalt-dominant volcanoes,
R2 rhyolites are associated with rhyolite-dominant submarine volcanoes, and R3 rhyolites are associated with rift
eruptions. Based on major element variation, trace element
ratios, REE patterns and Sr^Nd^Pb isotope ratios,
Tamura et al. (2009) concluded that R2 rhyolites are
partial melts of Oligocene middle crust, whereas R1 rhyolites are partial melts of Quaternary middle crust. Basalt
volcanoes consume new middle crust to produce R1 rhyolite magma, whereas R2 rhyolite volcanoes consume old
Oligocene middle crust. Moreover, rhyolite volcanoes
have no mantle roots beneath the crust; that is, there is no
evidence for a partially melted mantle source producing
basalt magmas beneath the rhyolite volcanoes. Instead,
arc-parallel dikes propagating from the basalt volcanoes
may provide the heat to partially melt the Oligocene crust
(Tamura et al., 2009). Huppert & Sparks (1988) showed
that when basalt sills are emplaced into preheated
continental crust, or as envisaged here a silicic to intermediate middle crust, a voluminous overlying layer of convecting silicic magma can be produced. Repeated
intrusion of dikes or sills from basalt volcanoes could preheat the crust between them, resulting in formation of
silicic magmas by melting the ‘continental crust’ components (middle crust) in the Izu^Bonin arc. This preheating
could stimulate the production of large volumes of melt
compared with the amount of basalt injected (Huppert &
Sparks, 1988).
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IZU^BONIN MISSING OLIGOCENE CRUST
AG E D I L E M M A O F E X P O S E D
M I D D L E C RU S T I N T H E I Z U
COLLISION ZONE
Suyehiro et al. (1996) and Takahashi et al. (1998) identified
low-velocity crust in the northern Izu^Bonin arc, consisting of 6 km thick middle crust (6·1^6·3 km/s) at a depth of
7^12 km and a 2 km thick upper part of the lower crust
(6·8 km/s) (Fig. 2b). The low velocity of the middle crust is
consistent with that of granitic rocks (Christensen &
Mooney, 1995). Kitamura et al. (2003) compared laboratory
measurements of P-wave velocities of Tanzawa tonalites
and hornblende gabbros (at various pressures and 258C)
with those obtained from seismic studies of the Izu^Bonin
arc (Suyehiro et al., 1996; Takahashi et al., 1998). Kitamura
et al. (2003) concluded that the middle crust and the
upper part of the lower crust in the IBM arc might consist
of tonalite and hornblende gabbro, respectively
(Kitamura et al., 2003). Rocks having the velocity of the
lower crust, except its upper part, were not observed
among the Tanzawa plutonic rocks.
The simplified geology of the Izu collision zone is shown
in Fig. 3a (after Aoike, 2001). The collision resulted in the
northward convex configuration of the Median Tectonic
Line in Central Japan. Figure 3b, which enlarges the
dashed rectangle in Fig. 3a, shows the location of the
Kofu Granitic Complex (KGC), which intrudes the
Cretaceous^Paleogene Shimanto Belt, and the Tanzawa
tonalites, which intrude a post-15 Ma accretionary terrane
(after Saito et al., 2007a). The KGC is the largest plutonic
complex in the Izu collision zone. K^Ar dates from the
KGC range from 15·7 to 7·4 Ma (Kawano & Ueda, 1966;
Shibata et al., 1984; Saito & Kato, 1996; Saito et al., 1997).
Sensitive high-resolution ion microprobe (SHRIMP)
zircon U^Pb ages for the KGC range from 16·8 to 10·6
Ma, overlapping the onset of IBM^Honshu collision at
15 Ma (Saito et al., 2007a). K^Ar and 40Ar^39Ar radiometric ages reported from the Tanzawa tonalite complex
vary widely, from 416 Ma to 4 Ma; these also overlap the
collision of the Tanzawa block at 6 Ma (compilation by
Yamada & Tagami, 2008). Exposed granites and tonalites
in the Izu Collision Zone are thought to represent felsic
middle crust exhumed by tectonic uplift during the continuing arc^arc collision (Kawate & Arima, 1998; Saito
et al., 2007a).
Here, we face a paradox, however. Forearc geology,
ODP results, magnetic anomalies and the seismic structure
of the Izu^Bonin arc suggest that this crust was created
mostly in Eocene^Oligocene times, a concept supported
petrologically by the fact that crustal partial melts
(Quaternary R2 rhyolites) have chemical signatures similar to those of Oligocene crust (Tamura et al., 2009).
However, the collision zone tonalites and granites, which
are deemed to be exhumed middle crust of the Izu^Bonin
arc, are Miocene in age. What has happened to the
Eocene^Oligocene Izu^Bonin arc crust in the process of
colliding with the Honshu arc?
Tonalitic rocks dredged from the Komahashi^Daini
Seamount, northern Kyushu^Palau Ridge, have radiometric ages of 37^38 Ma, which could represent felsic
middle crust in Eocene^Oligocene times (Haraguchi
et al., 2003). Interestingly, their chemical compositions are
similar to those of the collision zone plutons, as discussed
below.
E O C E N E ^ O L I G O C E N E I B M DATA
S O U RC E S
The oldest arc rocks in the IBM system, the so-called ‘proto-arc’ sequences, range in age from 48 to 44 Ma
(Eocene), are in places boninitic, and are exposed in the
IBM fore-arc (Ishizuka et al., 2006). Increasing evidence is
being found that most of the proto-arc sequences are tholeiitic basalts (M. K. Reagan et al., personal communication; O. Ishizuka et al., personal communication).
High-silica rhyolites also erupted during the proto-arc
stage (Reagan et al., 2008). Late Eocene to Oligocene
(44^25 Ma) volcanism in the IBM system is marked by
eruption of basaltic to rhyolitic lavas with trace element
characteristics typical of subduction-related lavas, whose
resulting sequences have been called the ‘first-arc’ by Gill
et al. (1994). Stern & Bloomer (1992) suggested that an
early rift-like setting produced boninite magmas, whose
eruption rates might have been exceptionally high
(120^180 km3/km of arc/Myr). However, Kodaira et al.
(2010) studied the crustal structure of the Bonin Ridge on
the Izu^Bonin fore-arc (Fig. 1) and found that Chichijima
island, which exposes boninites, is underlain by thin crust.
Hahajima island, on the other hand, which is part of the
first-arc sequence, exposes tholeiites and is underlain by
thick and mature crust (Kodaira et al., 2010). These workers
concluded that the first-arc (Eocene^Oligocene) processes
created the wide extent of the present IBM crust.
Eocene^Oligocene first-arc rocks make up the Alutom
Formation on Guam, the Hagman Formation on Saipan
and first-arc exposures on Rota (Reagan et al., 2008).
Significant exposures of first-arc rocks also occur on
the Palau Islands, which consist of the Aimeliik and
Ngeremlengui Formations (Hawkins & Ishizuka, 2010).
Omachi Seamount (298150 N, 1408450 E), a broad feature
located 20 km east of the Quaternary volcanic front, has
east^west and north^south dimensions of 31·5 and 59 km,
respectively, and is an old basement high characterized by
broad positive magnetic anomalies (Yamazaki et al., 1991;
Yamazaki & Yuasa, 1998). Omachi is composed mainly of
early Oligocene (32^34 Ma) andesite lava flows and volcanic breccias and early Miocene turbidites (Yuasa et al.,
1988, 1998, 1999; Nishimura, 1992) and is truncated to the
827
Fig. 3. (a) Geological map of the collision zone between the Honshu and Izu^Bonin arcs [simplified from Aoike (2001)]. The Sambagawa, Chichibu and Shimanto Belts, respectively, represent
the high P/T metamorphosed Jurassic accretionary prism, the Jurassic accretionary prism and the Cretaceous to Tertiary accretionary prism. Iso-depth contours of the Pacific plate (dashed
lines) and Philippine Sea plate (gray lines) slabs estimated by Nakajima et al. (2009). Collision between these arcs results in the northward convex structure of the Median Tectonic Line
(MTL) in Central Japan (Kanto Syntaxis) and several large reverse faults. The Kofu Granitic Complex (KGC) and Tanzawa tonalites lie within the dashed rectangle. These intrusive complexes
invaded the Cretaceous^Tertiary Shimanto Belt and post-15 Ma accretionary terrain, respectively. Black and white triangles show basalt-dominant volcanoes and rhyolite-dominant volcanoes,
respectively. Volcanoes: As, Asama; Hk, Hakone; Os, Oshima; Mi, Miyake; Ha, Hachijo; Ao, Aogashima; Su, Sumisu. The crust and mantle profile along A^A’ is shown in Fig. 9. (b) A more
detailed geological map of the collision zone, an enlargement of the dashed rectangle in (a), showing the location of the Kofu Granitic Complex (KGC) and the Tanzawa tonalites (after Saito
et al., 2007b). Collided Eocene^Oligocene basement of the Izu^Bonin arc might be exposed in places such as the Mineoka^Setogawa Ophiolite Complex (MSC) of the Shimanto Belt.
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TAMURA et al.
IZU^BONIN MISSING OLIGOCENE CRUST
west by a normal fault of the Quaternary rift system
(Fig. 1). Omachi andesites are underlain by serpentinite,
exposed along the base of a fault scarp 3500^3100 m
below sea level (Ueda et al., 2004). The turbidites cored
during ODP Leg 126 in the Izu^Bonin arc range in age
from 0·1 to 31 Ma (Hiscott & Gill, 1992; Gill et al., 1994).
The Oligocene turbidites, almost entirely volcanogenic in
origin, were recovered from three ODP sites (787, 792 and
793). Sites 792 and 787 are 60 and 95 km, respectively,
east of Aogashima Island, and Site 793, slightly to the
south, is about 75 km east of the volcanic front in the
middle of the upper slope depositional area (Figs 1 and 2).
The first mature IBM arc formed at 35 Ma. Evidence
for this is preserved along the Kyushu^Palau Ridge
(KPR), which was separated into three parts by rifting to
form the Shikoku and Parece Vela Basins. The KPR is a
remnant arc, about one-third of which formed in the
Oligocene (Yamazaki & Yuasa, 1998). The original locations of ODP sites, Bonin Islands, Saipan, Rota and
Guam are based on the motions suggested by Deschamps
& Lallemand (2003), which, after 35 Ma, closed all of the
basins (e.g. the Shikoku and Parece Vela Basins, Mariana
Trough and Bonin rifts). These relations are shown in
Fig. 4.
C OL L I S ION ZON E P LU TON IC
RO C K S C O M PA R E D W I T H I B M
VO L C A N I C RO C K S A N D
TURBIDITES
Table 1 summarizes the sources of analytical data for the
Oligocene and Miocene, Pliocene and Quaternary lavas
(including Oligocene, Miocene and Pliocene volcaniclastic
turbidites) and plutonic rocks of the Izu^Bonin^Mariana
arc. Table 2 shows additional new data for the Oligocene
Omachi Seamount and Rota determined by GSJ, AIST
and IFREE, JAMSTEC, respectively. Details of the analytical methods have been given by Ishizuka et al. (2009)
and Tamura et al. (2009). Glasses from Izu^Bonin tephras
(e.g. Bryant et al., 2003) are not included because
water-rich andesitic magmas cannot make glasses, and
thus the data are more biased in composition compared
with the lavas and turbidites (for further discussion, see
Tamura & Tatsumi, 2002). It is likely that the plutons
were the sites of some crystal^liquid segregation during
cooling, thus the stage 1 gabbro suite (Kawate & Arima,
1998), the gabbros and anorthosites (Takahashi et al.,
2004) of the Tanzawa plutonic bodies, and the ‘cumulates’
of the KGC (Saito et al., 2007a) are not included in this
Fig. 4. The first mature IBM arc at 35 Ma, which was active along
the original Kyushu^Palau Ridge (KPR) before the opening of the
Shikoku and Parece Vela Basins, and which could have been subducted by the Pacific plate along the KPR. The present KPR is the
remnant one-third of the Oligocene IBM arc, which has been rifted
into three parts (Yamazaki & Yuasa, 1998). The original Oligocene
locations of ODP sites, Bonin Islands, Saipan, Rota and Guam are
shown by open triangles, based the reconstructions of Deschamps &
Lallemand (2003). The Central Basin spreading center in the West
Philippine Basin, active until 33 Ma, is shown by a dashed line.
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Table 1: Sources of analytical data (major and trace elements and isotopic data) for Miocene, Pliocene and Quaternary
volcanic rocks, Eocene^Oligocene volcanic rocks, and plutonic rocks of the Izu^Bonin^Mariana arc
References
Miocene, Pliocene and Quaternary volcanic rocks
Miyake-jima
Yokoyama et al. (2003)
Hachijo-jima
Ishizuka et al. (2008)
Sumisu
Tamura et al. (2005)
Torishima
Tamura et al. (2007)
Other Quaternary volcanoes of the Izu–Bonin arc
Tamura & Tatsumi (2002) and references therein,
ODP Miocene–Pliocene turbidites and tephras
Hiscott & Gill (1992), Gill et al. (1994), Schmidt (2001),
Tamura et al. (2009) and references therein
Straub et al. (2004)
Eocene–Oligocene (first arc) volcanic rocks
ODP Oligocene lavas, turbidites and tephras
Hiscott & Gill (1992), Taylor et al. (1992), Gill et al. (1994),
Bonin ridge
Ishizuka et al. (2006)
Hahajima
Taylor & Nesbitt (1995)
Taylor & Nesbitt (1998), Schmidt (2001)
Guam, Rota and Saipan
Reagan et al. (2008), this study
Palau
Hawkins & Ishizuka (2009)
Omachi Seamount
This study
Plutons
Komahashi–Daini Seamount
Haraguchi et al. (2003)
Kofu Granitic Complex (KGC)
Saito et al. (2007a)
Tanzawa tonalites and syn-plutonic dikes
Kawate (1996), Kawate & Arima (1998), Takahashi et al. (2004)
study. The R2 and R3 types of Quaternary rhyolite of the
Izu^Bonin arc are omitted because they are partial melts
of Oligocene crust (R2) or rift-type rhyolites (R3)
(Tamura et al., 2009). The discussion in this paper refers to
analyses normalized to 100% on a volatile-free basis with
total iron calculated as FeO.
Major elements
Figure 5a shows variation diagrams of wt % SiO2 vs major
element oxides (wt %) and Mg-number [100 Mg/
(Mg þ Fe)]. Izu^Bonin plutons include the Tanzawa and
KGC, which are Miocene collision-zone tonalites and a
granitic complex, respectively, and the Oligocene
Komahashi^Daini tonalites. These three intrusive complexes have similar and overlapping variations in major
element abundances and are thus plotted together in
Fig. 5a; henceforth, they are referred to as ‘plutons’. The
three columns illustrate data for Miocene, Pliocene and
Quaternary volcanic rocks, Oligocene volcanic rocks and
plutons, respectively. Figure 5b shows statistical assessments of these three groups of rocks and comparisons
between them. Each point in the diagrams shows an average one standard deviation, which are obtained in the
seven ranges of wt % SiO2 (45^50, 50^55, 55^60, 60^65,
65^70, 70^75 and 75^80). The most voluminous intrusions
in the Tanzawa plutonic complex comprise rocks with
60 wt % SiO2 (Kawate & Arima, 1998); the ranges of
55^65 wt % SiO2 are highlighted in gray in Fig. 5b to
emphasize the comparisons.
Weight per cent TiO2 vs silica diagrams show that
Miocene, Pliocene and Quaternary volcanic rocks have
higher TiO2 contents compared with the plutons.
Specifically, many of the Quaternary basalts and andesites
(563 wt % SiO2) have 41·0 wt % TiO2, significantly
higher than those of the plutons (0·5^1·0 wt % TiO2) at
the same silica content. Quaternary and Neogene dacites
and rhyolites are also systematically high in TiO2 compared with the plutons. On the other hand, Oligocene volcanic rocks are low in TiO2 (0·5^1·0 wt %) and in this
respect are similar to the plutons.
Weight per cent Al2O3 vs silica diagrams also show the
similarity between the Oligocene volcanic rocks and the
plutons. Many of the Quaternary and Neogene basalts
and andesites are low in Al2O3 (516 wt %), and the dacites
and rhyolites also have systematically lower Al2O3 values
than those of the Oligocene volcanic rocks and the plutons.
830
TAMURA et al.
IZU^BONIN MISSING OLIGOCENE CRUST
Table 2: Representative major and trace element data for Oligocene volcanic rocks from Omachi Seamount and Rota
Locality:
Cruise no.:
Submersible:
Omachi Seamount
YK01-04
Leg1
Shinkai 6500
Sample no.:
Latitude (8N):
Longitude (8E):
Depth (m b.s.l.):
Rock type:
608R001
29·1518
140·7044
3426
lava flow
61·16
SiO2
0·58
TiO2
16·81
Al2O3
5·94
Fe2O3
MnO
0·10
MgO
1·97
CaO
5·12
3·59
Na2O
2·89
K2O
0·16
P2O5
Total
98·31
Trace elements (ppm) by XRF
Ba
Ni
Cu
Zn
Pb
Th
Rb
Sr
Y
Zr
Nb
Trace elements (ppm) by ICP-MS
Sc
V
87·1
Cr
0·56
Co
Ni
Cu
Rb
17·4
Sr
233
Y
21·2
Zr
42·2
Nb
2·45
Cs
0·173
Ba
123
La
8·94
Ce
23·3
Pr
2·82
Nd
13·6
Sm
3·22
Eu
1·10
Gd
3·24
Tb
0·537
Dy
3·26
Ho
0·667
Er
2·03
Tm
0·318
Yb
2·07
Lu
0·301
Hf
1·50
Ta
0·237
Tl
Pb
2·60
Th
0·915
U
0·338
608R002
29·1521
140·7058
3410
lava flow
65·56
0·54
15·49
4·94
0·12
1·89
5·98
3·17
0·75
0·14
98·58
608R003
29·1521
140·7058
3388
lava flow
62·25
0·53
16·39
5·79
0·13
2·55
5·01
3·37
2·00
0·13
98·15
128·3
4·5
10·3
53·5
5·6
1·7
6·3
324·3
27
121·1
2·3
148·2
8·2
15·7
59·3
3·6
0·9
21·1
284·9
25·1
120·9
2·2
102·7
7·88
126·0
4·18
6·33
296
27·6
69·2
3·32
0·097
124
10·0
22·5
2·88
13·0
3·08
1·02
3·51
0·581
3·72
0·814
2·68
0·450
3·18
0·548
1·83
0·175
3·63
1·27
0·492
608R004
29·1524
140·7063
3353
lava flow
61·97
0·59
16·02
6·60
0·23
2·19
6·05
3·37
0·82
0·13
97·99
608R006
29·1529
140·7078
3148
lava flow
58·77
0·68
17·92
6·75
0·14
3·04
6·54
3·39
1·30
0·17
98·70
608R007
29·1531
140·7099
3055
lava flow
56·54
0·72
18·60
7·41
0·14
2·93
6·69
3·37
1·82
0·18
98·40
608R008
29·1538
140·7099
2964
lava flow
57·75
0·58
17·51
5·93
0·13
3·30
7·18
3·60
0·95
0·08
97·02
128·8
6·4
13·8
75·8
3·7
0·4
8·6
318·8
24·2
114·6
2·7
117·9
9·3
19·3
62·8
4·3
1·7
13·5
320·3
23·3
110
2·1
122·5
5·84
104·5
1·76
131·3
1·55
174·0
17·72
19·6
269
29·9
83·4
3·55
0·113
144
11·7
26·3
3·31
14·4
3·36
1·02
3·61
0·637
3·99
0·879
2·81
0·464
3·24
0·553
2·08
0·229
12·3
271
25·7
35·2
2·92
0·181
156
9·45
21·8
2·77
12·4
2·99
0·940
3·25
0·548
3·56
0·776
2·44
0·403
2·65
0·442
1·14
0·196
7·73
322
24·5
78·5
4·13
0·051
138
9·09
21·5
2·78
13·2
3·40
1·32
3·75
0·612
3·59
0·800
2·37
0·359
2·25
0·344
2·30
0·275
24·3
271
30·6
93·3
3·93
0·153
135
12·0
28·7
3·77
16·8
4·11
1·31
4·41
0·749
4·73
0·976
2·91
0·443
2·83
0·462
2·46
0·262
15·4
284
20·7
89·9
2·26
0·359
148
8·04
16·6
2·05
8·96
2·13
0·746
2·41
0·444
2·89
0·628
1·97
0·313
2·08
0·353
2·15
0·162
3·94
1·07
0·634
2·30
0·866
0·442
4·07
1·22
0·426
3·64
1·07
0·572
4·04
1·08
0·374
(continued)
831
JOURNAL OF PETROLOGY
VOLUME 51
NUMBER 4
APRIL 2010
Table 2: Continued
Locality:
Omachi Seamount
Cruise no.:
Submersible:
YK01-04
Leg1
Shinkai 6500
Sample no.:
Latitude (8N):
Longitude (8E):
Depth (m b.s.l.):
Rock type:
609R018
29·1249
140·7059
3226
lava flow
SiO2
57·81
TiO2
0·67
17·36
Al2O3
Fe2O3
7·24
MnO
0·19
MgO
2·85
CaO
7·74
Na2O
3·26
K2O
0·92
P2O5
0·18
Total
98·24
Trace elements (ppm) by XRF
Ba
Ni
Cu
Zn
Pb
Th
Rb
Sr
Y
Zr
Nb
Trace elements (ppm) by ICP-MS
Sc
V
148·9
Cr
1·00
Co
Ni
Cu
Rb
13·0
Sr
308
Y
25·2
Zr
92·7
Nb
1·83
Cs
0·401
Ba
134
La
6·86
Ce
15·9
Pr
2·23
Nd
11·5
Sm
3·21
Eu
1·07
Gd
3·55
Tb
0·607
Dy
3·84
Ho
0·826
Er
2·51
Tm
0·416
Yb
2·71
Lu
0·416
Hf
2·51
Ta
0·167
Tl
Pb
3·52
Th
0·633
U
1·57
YK00-08
Leg 1
Shinkai 6500
609R020
29·1261
140·7086
3176
tuff breccia
53·75
0·87
17·73
8·63
0·19
3·86
8·86
2·86
0·87
0·16
97·79
610R001
29·137
140·7055
3373
tuff breccia
55·07
0·89
18·36
8·50
0·20
3·77
8·83
3·02
0·70
0·20
99·54
610R003
29·1366
140·7078
3270
lava flow
57·99
0·64
17·55
6·56
0·13
4·00
7·54
3·27
0·57
0·09
98·34
73·5
7·1
44·2
94·2
4·5
1·2
9·7
312
31·9
120·8
1·7
161·1
12·5
28
58·3
4·8
1
4·1
290·7
16·4
102·7
1·4
208·7
1·80
178·1
16·71
189·5
19·86
10·6
267
26·0
99·5
1·72
0·321
75·2
4·72
12·8
1·87
10·2
2·91
1·08
3·60
0·640
4·44
0·887
2·70
0·449
2·86
0·435
2·53
0·153
10·5
289
35·3
130
2·70
0·252
71·3
6·78
18·4
2·70
14·0
3·90
1·26
4·73
0·824
5·33
1·137
3·46
0·564
3·48
0·592
2·90
0·125
3·84
0·365
0·265
4·50
0·347
0·247
570R004
29·1779
140·7097
2889
lava flow
57·66
0·62
17·45
6·48
0·11
3·24
6·47
3·29
3·00
0·10
98·42
570R005-2
29·178
140·7104
2818
lava flow
59·12
0·66
17·50
6·74
0·13
2·79
6·53
3·42
1·91
0·11
98·91
181·4
12
45·4
43
3
1·3
23
245·8
19·4
101·5
1·3
162·2
7·3
34·2
49·1
3·4
0·5
21·3
261·1
18·5
94·4
1·4
4·10
251
15·5
81·7
1·44
0·093
148
5·46
11·2
1·44
7·13
1·93
0·77
2·31
0·396
2·55
0·577
1·64
0·275
1·71
0·256
2·37
0·169
19·4
189
16·3
71·4
1·31
0·123
150
5·88
12·8
1·71
7·60
1·86
0·635
2·25
0·376
2·43
0·533
1·62
0·264
1·75
0·283
1·90
0·140
22·3
248
19·6
83·6
1·61
0·181
163
7·27
15·3
1·95
8·76
2·29
0·790
2·70
0·457
2·85
0·624
1·91
0·309
2·05
0·348
2·28
0·167
3·61
1·037
0·549
1·75
0·843
0·481
2·99
0·935
0·537
(continued)
832
TAMURA et al.
IZU^BONIN MISSING OLIGOCENE CRUST
Table 2: Continued
Locality:
Rota
Sample no.:
Latitude (8N):
Longitude (8E):
Sample point:
RT02_7
14·0727
145·1056
Talakhaya
RT03_11
14·0725
145·1055
Talakhaya
RT04_14
14·0852
145·1115
Mt. Sabana
RT08_26
14·0839
145·1124
Mt. Sabana
Rock type:
conglomerate
conglomerate
volcanic breccia
volcanic breccia
62·46
0·47
16·71
6·55
0·23
2·89
7·21
3·31
0·82
0·11
100·76
54·99
0·52
17·71
7·62
0·12
5·86
9·92
2·17
1·03
0·17
100·11
271·5
21·3
28·2
64
2·5
11·4
144·1
40·3
60·4
0·6
107·7
30·9
65·8
60·2
3·1
1·9
4·9
469·6
16·9
41·7
0·6
25·9
30·0
29·0
20·7
26·0
10·39
139
39·1
59·8
0·728
0·139
236
7·44
12·6
2·17
11·0
3·30
0·860
4·87
0·814
5·41
1·25
3·88
0·565
3·76
0·612
1·91
0·051
0·035
2·09
0·517
0·266
32·4
35·3
56·9
4·42
448
16·6
40·8
0·345
0·092
100
5·99
12·3
1·90
9·14
2·42
0·856
2·81
0·438
2·70
0·581
1·74
0·249
1·63
0·250
1·41
0·020
0·023
2·37
1·30
1·63
55·63
SiO2
0·52
TiO2
17·81
Al2O3
9·29
Fe2O3
MnO
0·15
MgO
5·08
CaO
9·00
2·45
Na2O
0·52
K2O
0·07
P2O5
Total
100·51
Trace elements (ppm) by XRF
Ba
86
Ni
41·6
Cu
108
Zn
82·5
Pb
2·9
Th
Rb
8·5
Sr
146·6
Y
22·3
Zr
36·2
Nb
0·4
Trace elements (ppm) by ICP-MS
Sc
33·9
V
Cr
Co
38·8
Ni
43·3
Cu
97·9
Rb
8·00
Sr
144
Y
23·0
Zr
37·2
Nb
0·544
Cs
0·329
Ba
81·6
La
3·47
Ce
6·96
Pr
1·44
Nd
7·51
Sm
2·60
Eu
0·914
Gd
3·49
Tb
0·632
Dy
4·33
Ho
0·950
Er
2·97
Tm
0·458
Yb
3·25
Lu
0·504
Hf
1·25
Ta
0·036
Tl
0·044
Pb
2·00
Th
0·343
U
0·206
RT16_51
14·0703
145·1118
Okgok Fall
Talakhaya
floated rock
RT16_56
14·0703
145·1118
Okgok Fall
Talakhaya
floated rock
RT16_57
14·0703
145·1118
Okgok Fall
Talakhaya
floated rock
53·94
0·53
17·41
8·14
0·13
6·48
10·44
2·07
0·72
0·13
99·99
61·28
0·77
16·58
8·22
0·14
2·23
6·10
3·92
1·16
0·13
100·54
62·48
0·61
17·36
6·95
0·13
2·09
6·56
3·85
0·69
0·14
100·87
52·04
0·41
16·49
9·58
0·17
7·90
11·89
1·70
0·28
0·05
100·50
82·9
34·3
87·8
66·8
145·5
5·7
74·7
84·7
4
1·4
15·9
157·4
24·5
84·4
1·1
186·1
9·1
39·8
74·4
3·3
0·7
10·3
167·7
32·2
74·3
0·9
41·8
136
83·6
71·9
6
142
10·2
24·1
0·6
35·2
22·5
23·3
41·2
33·5
38·3
77·4
4·08
391
12·1
35·7
0·305
0·063
74·7
3·66
8·66
1·29
6·29
1·77
0·650
2·05
0·334
2·09
0·449
1·36
0·196
1·30
0·202
1·25
0·017
0·021
2·02
0·993
0·517
28·6
7·11
71·7
14·64
155
24·7
86·7
1·08
0·466
139
5·05
12·2
1·81
8·93
2·79
0·905
3·65
0·635
4·20
0·928
2·85
0·429
2·92
0·459
2·68
0·071
0·068
3·17
0·774
0·439
21·3
7·96
37·1
9·58
166
31·8
74·7
0·923
0·275
170
5·41
10·7
1·71
8·57
2·65
0·863
3·87
0·650
4·37
1·01
3·12
0·454
3·03
0·485
2·33
0·061
0·046
2·85
0·673
0·355
47·3
127
78·2
5·47
143
10·7
24·5
0·334
0·248
35·9
1·56
3·77
0·58
3·05
1·06
0·425
1·49
0·268
1·83
0·405
1·25
0·184
1·27
0·196
0·803
0·020
0·017
1·19
0·209
0·123
1·1
4·2
381·2
11·5
35·6
(continued)
833
JOURNAL OF PETROLOGY
Table 2: Continued
Locality:
Rota
Sample no.:
Latitude (8N):
Longitude (8E):
Depth:
RT16_61
14·0703
145·1118
Okgok Fall
Talakhaya
floated rock
Rock type:
51·88
SiO2
TiO2
0·40
15·79
Al2O3
Fe2O3
9·41
MnO
0·16
MgO
9·87
CaO
11·20
Na2O
1·63
K2O
0·29
0·05
P2O5
Total
100·68
Trace elements (ppm) by XRF
Ba
50·7
Ni
133
Cu
87·1
Zn
67
Pb
Th
Rb
5·4
Sr
131·4
Y
10·4
Zr
22·8
Nb
0·5
Trace elements (ppm) by ICP-MS
Sc
41·8
V
Cr
Co
46·3
Ni
134
Cu
84·3
Rb
5·06
Sr
135
Y
10·7
Zr
23·7
Nb
0·318
Cs
0·170
Ba
41·9
La
1·52
Ce
3·67
Pr
0·58
Nd
2·95
Sm
1·01
Eu
0·412
Gd
1·47
Tb
0·262
Dy
1·78
Ho
0·397
Er
1·23
Tm
0·181
Yb
1·24
Lu
0·191
Hf
0·78
Ta
0·018
Tl
0·020
Pb
1·12
Th
0·202
U
0·115
RT16_62
14·0703
145·1118
Okgok Fall
Talakhaya
floated rock
63·45
0·68
15·93
7·60
0·13
1·80
5·97
3·78
1·00
0·17
100·51
144
6·8
94·3
77
2·7
1·5
14
150·1
24·5
88
1·1
22·1
21·9
7·60
87·9
13·43
148
24·2
90·6
1·04
0·386
134
4·99
12·1
1·79
8·68
2·70
0·825
3·50
0·611
4·01
0·890
2·75
0·408
2·82
0·440
2·76
0·070
0·092
3·18
0·803
0·434
VOLUME 51
NUMBER 4
APRIL 2010
Miocene, Pliocene and Quaternary volcanic rocks have
higher wt % FeO than the plutons and Oligocene volcanic rocks at the same SiO2 content. In particular, the
basalts and andesites can contain as much as 16 wt %
FeO, values not observed in the plutons and Oligocene
volcanic rocks. Like the variations in TiO2 content, dacites
and rhyolites in the Miocene, Pliocene and Quaternary
units show higher abundances than those of the plutons
and Oligocene volcanic rocks. These high FeO contents
result from a tholeiitic differentiation trend, as will be
shown subsequently based on Mg-numbers.
Figure 5 shows that some Oligocene volcanic rocks are
higher in Na2O and K2O at 50^60 wt % SiO2 than the
plutons and Quaternary and Neogene volcanic rocks.
However, the Oligocene volcanic rocks and plutons are
similar, in that both have 3^5 wt % Na2O at 50^60 wt %
SiO2. Moreover, the strong positive trend of the Miocene,
Pliocene and Quaternary volcanic rocks between Na2O
and SiO2 cannot be observed in the plutons and
Oligocene volcanic rocks, where the trends are more
broad and relatively flat from 50 to 70 wt % SiO2
(Fig. 5b). Low-K Miocene, Pliocene and Quaternary volcanic rocks from the Izu^Bonin volcanic front are different
from the plutons and Oligocene rocks, both of which have
a wider range and higher values of K2O at the same SiO2
content. We can see two K^Si trends in plutons 465 wt %
SiO2. Such trends are different from those in both the
Miocene, Pliocene and Quaternary volcanic rocks and
the Oligocene volcanic rocks. It is likely that some of the
plutons have experienced crystal^liquid segregation
during cooling, and some silicic rocks might represent
such liquids. Thus, the range of 55^65 wt % SiO2 might
be the most appropriate for comparisons between the
three groups.
The Mg-numbers [Mg-number ¼100Mg/(Mg þ Fe)]
of the Miocene, Pliocene and Quaternary volcanic rocks
reflect their high FeO contents. Many basalts and andesites have Mg-number of 540, but some andesites have
values 440, corresponding respectively to the tholeiitic
and calc-alkaline trends defined by Miyashiro (1974). Thus
both tholeiitic and calc-alkaline rocks exist in the
Miocene, Pliocene and Quaternary of the Izu^Bonin arc,
as they do in the NE Japan arc (e.g. Miyashiro, 1974;
Tatsumi & Kogiso, 2003). On the other hand, most plutons
and Oligocene volcanic rocks have high Mg-numbers and
could be classified as calc-alkaline rocks by the definition
of Miyashiro (1974). Thus, again, the plutons and
Oligocene rocks are similar in terms of high Mg-number.
Zr, Y and Zr/Y
Figure 6a shows the variation of wt % SiO2 vs Zr, Y and
Zr/Y. Importantly, the Zr contents of all IBM volcanic
rocks and plutons are low (5200 ppm), and thus
zircon-saturation temperatures are 58508C (Watson &
Harrison, 1983). Figure 6b shows statistical assessments of
b.s.l., below sea level.
834
TAMURA et al.
IZU^BONIN MISSING OLIGOCENE CRUST
Fig. 5. (a) Variation diagrams of wt % SiO2 vs major element oxides (wt %) and Mg-number [100Mg/(Mg þ Fe)]. The three columns illustrate Miocene, Pliocene and Quaternary volcanic rocks, Oligocene volcanic rocks and plutons, respectively. Data sources and sample locations
are shown in Table 1 and Figs 1^4. (b) Statistical assessments of these three groups of rocks in (a) and comparisons between them. Each point
in the diagrams shows an average one standard deviation, calculated for seven ranges of wt % SiO2 (45^50, 50^55, 55^60, 60^65, 65^70,
70^75 and 75^80). The ranges for 55^65 wt % SiO2 are highlighted in gray to emphasize the comparisons.
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Fig. 5. Continued.
these three groups of rocks and comparisons between
them. Zr/Y values are petrogenetically significant because
(1) they are easily measured, (2) Zr and Y are not transported in fluids or affected by moderate alteration, and
(3) high ratios indicate a relatively fertile source, whereas
low values indicate a depleted source. In the Izu^Bonin
arc, Zr/Y and REE patterns changed to depleted ones in
the Miocene, Pliocene and Quaternary (Gill et al., 1994)
836
TAMURA et al.
IZU^BONIN MISSING OLIGOCENE CRUST
concurrent with the formation of the Shikoku back-arc
basin between 25 and 15 Ma.
As can be seen in Fig. 6, the Zr/Y values of the Miocene,
Pliocene and Quaternary volcanic rocks are low (1^3), in
contrast to the more variable and high (mostly 2^4)
values of the Oligocene volcanic rocks. Within the
Oligocene volcanic rocks, Oligocene turbidites in the
northern part of the Izu^Bonin arc are more variable
(1^8) in Zr/Y, compared with lavas of similar age to the
south. Omachi Seamount, the Oligocene Mariana arc
(Guam, Rota and Saipan) and the Oligocene Palau rocks
have Zr/Y values in the range of 2^6·5,1^4 and 1^5, respectively. Most Oligocene volcanic rocks have Zr/Y values of
2^4, and thus 10^50% partial melts of such protoliths
could have Zr/Y values similar to those of R2 rhyolites
(4^5) (see Tamura et al., 2009, fig. 19).
Zr/Y values of the plutons define a broad trend that
increases with silica content (Fig. 6). Tanzawa tonalites,
KGC and Komahashi^Daini tonalites each show wide
variation in Zr/Y, ranging from 52 to 47, similar to the
high values of the Oligocene volcanic rocks. Plutons containing 469 wt % SiO2 are also highly variable in Zr/Y,
but all of the three plutons overlap the fields of R2 rhyolites, suggesting a close relationship with the Oligocene volcanoes. However, as noted above, it is likely that the
plutons have experienced some crystal^liquid segregation
during cooling, which may explain the systematic increase
in Zr/Y with silica content. Moreover, some high-silica
KGC samples have unusually low Zr/Y, which could be
related to late-stage fractionation of zircon (Saito et al.,
2007a). Thus, again, the range of 55^65 wt % SiO2 may
be the most appropriate for comparison between the three
groups (Fig. 6b).
We emphasize here that both the Miocene plutons
(Tanzawa and KGC) and the Oligocene Pluton
(Komahashi^Daini) have higher Zr/Y values than those
of the Miocene, Pliocene and Quaternary volcanic rocks
and are more similar in Zr/Y ratio, if not systematics, to
the Oligocene volcanic rocks in the range 55^65 wt %
SiO2 (Fig. 6b).
REE patterns and ratios
Based on a study of Izu^Bonin turbidites recovered during
ODP Leg 126, Gill et al. (1994) concluded that as the
Shikoku back-arc basin formed between 25 and 15 Ma
REE patterns changed from the flat to enriched characteristics of the Oligocene turbidites to the depleted ones of
the Miocene, Pliocene and Quaternary. Figure 7 shows
the contrasting REE patterns between the Quaternary
lavas, Miocene^Pliocene turbidites, Oligocene lavas and
turbidites and the plutons. Quaternary Izu^Bonin basalts,
andesites and R1 rhyolites (Fig. 7a) are depleted in light
REE (LREE) compared with the middle REE and heavy
REE. Importantly, the LREE-depleted patterns of the R1
rhyolites are subparallel to those of Quaternary Izu
arc-front basalts and andesites (Fig. 7a). Miocene^
Pliocene (4^15 Ma) Izu turbidites also have
LREE-depleted patterns that are indistinguishable from
those of the Quaternary lavas (Fig. 7b). Oligocene rocks
have depleted to LREE-enriched patterns (Fig. 7c and d)
and the plutons have patterns similar to those of the
Oligocene rocks ((Fig. 7e and f).
Figure 7g and h compares the range of REE-abundance
patterns of the plutons with those of the Miocene,
Pliocene and Quaternary volcanic rocks and the
Oligocene volcanic rocks, respectively. It can be seen that
the field of Miocene, Pliocene and Quaternary volcanic
rocks does not match that of the plutons (Fig. 7g), whereas
the field of Oligocene volcanic rocks is mostly included
within the pluton field (Fig. 7h). Figure 7i shows the variation of La/Sm vs Dy/Yb of Miocene, Pliocene and
Quaternary volcanic rocks, Oligocene volcanic rocks and
plutons. Statistical assessments and comparisons are
shown in Fig. 7j. Both the Oligocene volcanic rocks and
the plutons have much larger standard deviations of La/
Sm than the Miocene, Pliocene and Quaternary volcanic
rocks, but their averages are similar to (2) but higher
than the latter (51). Figure 7k shows Ce/Yb vs wt % SiO2
for the three subgroups, and Fig. 7l presents their statistical
evaluation. The Ce/Yb ratios exhibit large standard deviations for the Oligocene volcanic rocks and plutons, but
values are systematically higher than those of the
Miocene, Pliocene and Quaternary volcanic rocks at the
same SiO2 content. The average Ce/Yb ratios in the range
55^65 wt % SiO2 are 5·4, 6·2 and 2·8, respectively.
87
Sr/86Sr and
143
Nd/144Nd ratios
Figure 8a shows along-arc variations of 87Sr/86Sr and
143
Nd/144Nd for the collision zone plutons (KGC and
Tanzawa tonalites), Miocene, Pliocene and Quaternary
frontal-arc volcanoes and rear-arc volcanoes of northern
Izu and the Mariana Trough (see Isse et al., 2009, for references). Additional data for Miocene^Pliocene volcanic
glasses from ODP Site 782 are from Straub et al. (2004).
Data for collision zone plutons (KGC and Tanzawa tonalites) are from Saito et al. (2007a) and Kawate (1996),
respectively. Collision zone plutons are exposed at the
northern end of the Izu^Bonin arc, but their isotopic compositions are different in two important ways from the
lavas of the Miocene, Pliocene and Quaternary northern
Izu^Bonin arc. First, the northern Izu^Bonin arc should
mostly be genetically related to the collision zone plutons.
Most Izu rear-arc volcanoes are Miocene (17^3 Ma) in
age (Ishizuka et al., 2003). However, they are medium-K
volcanic rocks with low 87Sr/86Sr (0·703), which are different from those of the plutons. Frontal volcanoes of the
northern Izu^Bonin arc are systematically lower in
87
Sr/86Sr and higher in 143Nd/144Nd compared with the
plutons. Second, the southern Izu^Bonin arc and
Mariana Northern Seamount Province (NSP) have the
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Fig. 6. (a) Variation diagrams of wt % SiO2 vs Zr, Y and Zr/Y. The first, second and third columns plot Miocene, Pliocene and Quaternary
volcanic rocks, Oligocene volcanic rocks and plutons, respectively. (b) Statistical assessments of these three groups of rocks and comparisons
between them. Each point in the diagrams shows an average one standard deviation, which are calculated for the seven ranges of wt %
SiO2 (45^50, 50^55, 55^60, 60^65, 65^70, 70^75 and 75^80). As in Fig. 5, the range for 55^65 wt % SiO2 is highlighted in gray to emphasize
the comparisons.
highest 87Sr/86Sr values and the lowest 143Nd/144Nd in the
IBM arc; many of them are high-K and shoshonitic rocks
(Sun & Stern, 2001). Moreover, many of them are much
lower in 143Nd/144Nd than the plutons. Geographically,
the plutons are far removed from the Mariana Central
Island Province (CIP) and Southern Seamount Province
(SSP), and in addition, the latter are low in 87Sr/86Sr and
are thus different from the plutons.
Figure 8b shows along-arc variations of 87Sr/86Sr and
143
Nd/144Nd for the collision zone plutons and Eocene^
Oligocene lavas (Taylor et al., 1992; Taylor & Nesbitt, 1995,
1998; Schmidt, 2001; Haraguchi et al., 2003; Ishizuka et al.,
2006; Reagan et al., 2008). There appears to be only limited
and no systematic along-arc variation of 87Sr/86Sr and
143
Nd/144Nd in Eocene^Oligocene times (Fig. 8b), which
is also suggested from a study of the Kyushu^Palau ridge
(O. Ishizuka, personal communication). Lavas of the
Eocene^Oligocene IBM have systematically higher
87
Sr/86Sr and lower 143Nd/144Nd than those of the
Miocene, Pliocene and Quaternary northern Izu^Bonin
lavas; however, these are similar to those of the collision
zone plutons.
DISCUSSION
We have shown that major elements, Zr/Y, REE ratios and
patterns, and 87Sr/86Sr and 143Nd/144Nd compositions indicate that the KGC and Tanzawa tonalites (and their associated syn-plutonic dikes) have strong similarity to the
Eocene^Oligocene IBM volcanic rocks and the Oligocene
Komahashi^Daini plutonic complex, and that they are dissimilar to the Miocene, Pliocene and Quaternary Izu^
Bonin magmas. How might this be explained?
Return of the Oligocene middle crust?
The chemical differences between the Oligocene and
Miocene^Quaternary Izu^Bonin magmas have been
related to the formation of the Shikoku back-arc basin
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TAMURA et al.
IZU^BONIN MISSING OLIGOCENE CRUST
Fig. 7. REE patterns of, (a) Izu^Bonin Quaternary arc-front basalts, andesites and R1 rhyolites, (b) Miocene^Pliocene (4^15 Ma) Izu turbidites, (c) Oligocene Omachi Seamount and Oligocene Palau, (d) Oligocene Izu turbidites and Oligocene Mariana arc (Guam, Rota and
Saipan), (e) Kofu Granitic Complex (KGC), Tanzawa tonalites and Tanzawa syn-plutonic dikes, (f) Oligocene Komahashi^Daini tonalites.
Combined normalized REE abundance patterns are illustrated in (g) comparison of plutons with Miocene, Pliocene and Quaternary volcanic
rocks and (h) comparison of plutons with Oligocene volcanic rocks, respectively. (i) La/Sm vs Dy/Yb ratios of Miocene, Pliocene and
Quaternary volcanic rocks, Oligocene volcanic rocks and plutons. (j) Statistical assessment and comparison of La/Sm vs Dy/Yb. (k) Ce/Yb vs
wt % SiO2 of the three groups and (l) their statistical assessments. Each point in (j) and (l) shows an average one standard deviation for
the three ranges of wt % SiO2 (45^55, 55^65, and 65^75). The range for 55^65 wt % SiO2 in (l) is highlighted in gray to emphasize the
comparisons.
(Gill et al., 1994), and thus they most probably reflect chemical changes within the mantle wedge, which transformed
the mantle source of the parental mafic magmas from
enriched to depleted.
Straub (2003) suggested that ultra-depleted subarc
mantle, which produced boninites, existed during arc initiation and was then gradually replaced by Indian
mid-ocean ridge basalt (MORB) mantle during the
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Fig. 7. Continued.
Eocene to late Oligocene. Straub (2003) also suggested that
this replacement, driven by mantle convection, was a gradual process. This gradual replacement readily accounts for
the wide range of trace element and isotopic ratios in the
Eocene^Oligocene arc lavas and is best explained by
source mixing (e.g. Hickey-Vargas & Reagan, 1987;
Straub, 2003).
The geochemical characteristics and wide range of trace
element and isotope ratios of the KGC and Tanzawa tonalites are similar to those of the Eocene^Oligocene arc
lavas. A large body of geochronological data (K^Ar and
40
Ar^39Ar radiometric ages and SHRIMP zircon U^Pb
ages), however, shows that the two large plutonic complexes in the Izu collision zone were emplaced in
840
TAMURA et al.
IZU^BONIN MISSING OLIGOCENE CRUST
Fig. 8. (a) Along-arc variations in 87Sr/86Sr and 143Nd/144Nd for collision zone plutons (KGC and Tanzawa tonalites), Miocene, Pliocene and
Quaternary frontal-arc volcanoes (open circles) and the rear-arc volcanoes of northern Izu and the Mariana Trough (solid triangles) (see Isse
et al., 2009, for references). Additional data for Miocene^Pliocene volcanic glasses from ODP Site 782 are from Straub et al. (2004). Data for collision zone plutons (KGC and Tanzawa tonalites) from Saito et al. (2007a) and Kawate (1996), respectively. Horizontal lines show averages of
frontal volcanoes of the northern Izu^Bonin arc and Mariana CIP and SSP, respectively. (b) Along-arc variations of 87Sr/86Sr and
143
Nd/144Nd of collision zone plutons and Eocene^Oligocene lavas (Taylor et al., 1992; Taylor & Nesbitt, 1995, 1998; Schmidt, 2001; Haraguchi
et al., 2003; Ishizuka et al., 2006; Reagan et al., 2008). Again, horizontal lines show averages of Miocene, Pliocene and Quaternary frontal volcanoes of northern Izu^Bonin arc and Mariana CIP and SSP, respectively. KDS; Komahashi^Daini Seamount.
Miocene times (4^17 Ma) [for a summary, see Saito et al.
(2007a) and Yamada & Tagami (2008)]. One possibility is
that the KGC and Tanzawa tonalites were produced by
partial melting of Oligocene crust, or that they represent
remobilized Oligocene magma bodies. Shukuno et al.
(2006) and Tamura et al. (2009) described melting experiments using the Tanzawa tonalite that showed the following: at 9008C and 3 kbar, Tanzawa tonalite with 62·3%
SiO2 becomes a mixture of 20% of rhyolitic melt (76%
SiO2) and 50% plagioclase, 20% pyroxene and a few
per cent of Fe^Ti oxides and quartz. Thus, it is possible
that partial remelting and remobilization of these plutonic
bodies could result in the complete loss of chronological
information relating to their source. Such information was
originally recorded in hornblende, biotite, zircon and the
whole-rocks, and would record post-cooling chronological
information following remelting and remobilization.
Whole-rock Tanzawa tonalites contain 1wt % H2O
(Shukuno et al., 2006), and their rhyolitic partial melts
would contain 5^8 wt % H2O. Thus the hydrous experiments of Watson & Harrison (1983) could explain
the behavior of zircon in the remobilized bodies.
The zirconium contents in Oligocene^Quaternary Izu^
Bonin volcanic rocks and plutons are low (5200 ppm)
(Fig. 6), and thus zircon-saturated temperatures are
58508C (Watson & Harrison, 1983). Saito et al. (2007b)
showed peak metamorphic conditions of 3 kbar and
7808C for contact aureoles in the KGC. Thus it is reasonable to assume that a granitic body would have a temperature of 48508C before intruding the host-rocks. For such
low-Zr high-temperature I-type granitoids, the likely dissolution of most of the recycled zircon before attainment of
saturation would result in the complete loss of chronological information related to their source (Watson &
Harrison, 1983). Many plutonic rocks and sedimentary
rocks show a remarkable memory of tens of million years
preserved in zircon ages. Zircons are clearly stable when
they are separated from the magma from which they crystallized; however, they cannot survive when they are surrounded by low-Zr partial melts such as those of the IBM
rhyolites.
A fundamental question is whether the collision of the
IBM and Honshu arcs could rejuvenate and remobilize
the Oligocene middle crust.
841
Fig. 9. Schematic cross-section of the Izu^Bonin arc and the Honshu arc along the A^A’ profile in Fig. 3, modified from Aoike (2001). The Kofu Granitic Complex (KGC) and the Tanzawa
tonalites were emplaced during the Miocene within the zone of collision, delamination and accretion between the two arcs. Crustal structure of the Izu^Bonin arc after Kodaira et al. (2007a).
Most parts of the middle crust of the Izu^Bonin arc were produced in Eocene^Oligocene times (Kodaira et al, 2008). The middle crust in the collision zone was dragged to mantle depths
(40^50 km) and temperatures (900^10008C). The resulting partial melting resulted in remobilization and delamination of the middle crust from the lower crust of the Philippine Sea plate.
Possible magma source mantle [mantle diapir or hot finger (Tamura et al., 2002)] after Tamura et al. (2009). Closed and open triangles show basalt-dominant volcanoes and rhyolite-dominant
volcanoes, respectively. Rhyolite volcanoes in the Izu^Bonin arc have no mantle roots, but dikes are inferred to travel laterally from the basalt volcanoes (as they did for 30 km during the
2000 eruption of Miyakejima (Nishimura et al., 2001; Geshi et al., 2002), providing heat to produce rhyolite magma having the characteristics of Oligocene middle crust (Tamura et al., 2009).
The Tonoki^Aikawa Tectonic Line (TATL) and Kannawa Fault (KF) are thought to be the 15 Ma and present-day plate boundaries, respectively. (Note the delamination of upper and middle
crust from the Izu^Bonin arc plate and their accretion to the Honshu arc in the collision zone.) NMTL, Niigata^Matsumoto Tectonic Line (informal name); MTL, Median Tectonic Line;
BTL, Butsuzo Tectonic Line; TATL, Tonoki^Aikawa Tectonic Line; KF, Kannawa Fault.
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IZU^BONIN MISSING OLIGOCENE CRUST
Collision, delamination and accretion of
the Oligocene crust
Figure 9 shows the A^A0 profile from Fig. 3, across the Izu^
Bonin arc, the collision zone and the Honshu arc (based
on Aoike, 2001; Kodaira et al., 2007a; Tamura et al., 2009).
Most of the Izu^Bonin middle crust was produced in
Eocene^Oligocene times (Kodaira et al., 2008); the IBM
and Honshu arcs began to collide at 15 Ma (e.g. Soh
et al., 1998). The Eocene^Oligocene middle crust was probably carried to mantle depths in the collision zone, where
high temperatures (49008C) could easily have been
attained beneath the doubled thickness of crust (Fig. 9).
Subduction zones are generally cold because of the subduction of cold slabs of oceanic lithosphere. However, this collision zone is hot because of the subduction of newly
produced Izu^Bonin crust, many parts of which have
been heated by Quaternary basalts from mantle diapirs
or hot fingers (Tamura et al., 2009). The down-dragged preheated middle crust would partially melt, but the mafic
lower crust would not, resulting in delamination and separation of the middle crust from the lower crust (Fig. 9).
The remobilized middle crust could then delaminate and
rise buoyantly; the earlier parts intruded the Shimanto
Belt to form the KGC, whereas the later parts invaded the
Miocene Izu^Bonin sequences (Tanzawa Group) to form
the Tanzawa tonalites.
The aseismic Philippine Sea plate has been detected in
the Izu^Bonin^Honshu collision zone, and its configuration and depth are consistent with this model. This plate
is subducting to depths of 130^140 km without evidence of
a tear or other gap, even beneath areas NW of the Izu collision zone (Nakajima et al., 2009). Arai et al. (2009) studied
the crustal structure in the collision zone east of profile
A^A0 in Fig. 3. Their profile shows the juxtaposition and
delamination of the Izu^Bonin upper crust with that of
the Honshu arc at depths of 10^20 km. It is not clear how
deep the middle crust of the Izu^Bonin arc subducts
beneath the Honshu arc before being remobilized to the
surface. A seismic velocity study of the Tanzawa tonalites
(Kitamura et al., 2003) suggests that most of the lower
crust is missing beneath the Tanzawa tonalites, consistent
with delamination of the Izu^Bonin middle and lower
crusts below the Honshu arc. Confirmation of our interpretations might be provided by drilling into the
Oligocene middle crust of the Izu^Bonin arc, as now proposed to the IODP.
The final test awaits the development of a new method of
geochronology, one that is not affected by remelting or
remobilization at depth. Given that all geochronometers
involve decay of radioactive elements in given rocks or
minerals, such a geochronometer is unlikely to be developed if magmatism always follows segregation of liquids
from residues. However, we suggest here that some remobilization of andesitic middle crust could result in tonalitic
or granitic rocks with compositions similar to those of the
middle crust. We might be able to identify such a group of
rocks little affected by melt segregation and to produce isochrons by using such rocks. The resulting isochron ages
might be much older than the estimated ages from single
rocks, and they could represent the original age of the
source rocks.
Where is the Eocene^Oligocene arc upper
crust in the collision zone between the
Izu^Bonin and Honshu arcs?
If it is true that most IBM crust is of Eocene^Oligocene
age, then it follows that accreted supracrustal sequences
in
the
Izu
collision
zone
should
include
volcano-sedimentary units of that age. One such possibility
is the Mineoka ophiolite, which is located in the eastern
extension of the Cretaceous^Paleogene Shimanto Belt
(Fig. 3a) and is overlain, unconformably, by fore-arc sedimentary rocks. The ophiolite crops out in the southern
part of the Boso Peninsula, near the Izu^Bonin arc collision system. It includes pelagic to hemi-pelagic sedimentary rocks, tholeiitic pillow basalts and dolerites,
alkali-basaltic sheet flows, calc-alkaline dioritic to gabbroic
rocks and serpentinized peridotites.
The tholeiitic basalts in the ophiolite have variable trace
element compositions, ranging from mid-ocean ridge
basalt to island-arc basalt, whereas the alkali basalts have
a within-plate affinity (Hirano et al., 2003). 40Ar/39Ar and
K^Ar dates yield ages of 49 13 Ma for the tholeiite (Fe^
Ti basalts), 19·62 0·90 Ma for alkali-basalts, and c. 25, 35
and 40 Ma for the plutonic rocks (two diorites and
gabbro) (Hirano et al., 2003). Eocene (37·0 0·6 Ma)
alkali basalts are also found on the Miura Peninsula
(Taniguchi & Ogawa, 1990).
The Setogawa ophiolite complex, which is an accretionary complex, occurs in the Tertiary Setogawa Terrace of
the Shimanto Belt, west of the Izu^Honshu arc collision
zone (Fig. 3b) (Arai, 1991; Ishiwatari, 1991; Shiraki et al.,
2005). This complex contains serpentinite, picrite, gabbro
to trondhjemite, tholeiitic to alkali basalt, and high-MgO
and high-SiO2 rocks (high-magnesian andesites). The
occurrence of boninite-like rocks implies that part of the
Setogawa ophiolite formed in a supra-subduction zone setting from a depleted mantle source (Shiraki et al., 2005).
Collided Eocene^Oligocene basement of the Izu^Bonin
arc might be exposed in places in the Shimanto Belt, as in
the Mineoka^Setogawa Complex, and the present study
should therefore stimulate future petrological studies of
this interesting and complex area.
CONC LUSIONS
We have shown that major elements, Zr/Y, REE ratios and
patterns, and 87Sr/86Sr and 143Nd/144Nd compositions indicate that the Miocene KGC and Tanzawa tonalites (and
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JOURNAL OF PETROLOGY
VOLUME 51
their syn-plutonic dikes) are more akin to the Eocene^
Oligocene IBM volcanic rocks and are dissimilar to the
Miocene, Pliocene and Quaternary Izu^Bonin magmas.
These Miocene intrusive complexes within the collision
zone are, therefore, interpreted to be remobilized middle
arc crust, most of which was produced in Eocene^
Oligocene times. During the collision, the down-dragged
preheated middle crust of the IBM was partially melted,
but the mafic lower crust was not, resulting in delamination and separation of the middle crust from the lower
crust. The remobilized middle crust then rose buoyantly;
earlier parts intruded the Shimanto Belt to form the
KGC, and later parts invaded the Miocene Izu^Bonin
sequences (Tanzawa Group) to form the Tanzawa tonalites.
Partial melting during collision erased the Eocene^
Oligocene age of this remobilized middle crust and
explains the Miocene age of these intrusive complexes
seen today.
AC K N O W L E D G E M E N T S
We thank Professor Robert J. Stern and two anonymous
reviewers for their thorough and constructive reviews. We
appreciate the encouragement and editorial help of the
editor Professor John Gamble.
F U N DI NG
This work was supported in part by the JSPS Grant-in-Aid
for Scientific Research (B) (17340165 and 20340122) and
Grant-in-Aid for Creative Scientific Research (19GS0211).
Many of the data were obtained from samples collected
during JAMSTEC cruises and the JAMSTEC GANSEKI
database (http://www.jamstec.go.jp/ganseki/index.html).
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