Download Lithospheric Removal as aTrigger for Flood

JOURNAL OF PETROLOGY
VOLUME 50
NUMBER 11
PAGES 2157^2186
2009
doi:10.1093/petrology/egp072
Lithospheric Removal as aTrigger for Flood
Basalt Magmatism in the Trans-Mexican
Volcanic Belt
LAURA MORI1*, ARTURO GO¤MEZ-TUENA2, PETER SCHAAF3,
STEVEN L. GOLDSTEIN4, OFELIA PE¤REZ-ARVIZU2 AND
GABRIELA SOLI¤S-PICHARDO1
INSTITUTO DE GEOLOGI¤A, UNIVERSIDAD NACIONAL AUTO¤NOMA DE ME¤XICO, 04510 MEXICO CITY, MEXICO
CENTRO DE GEOCIENCIAS, UNIVERSIDAD NACIONAL AUTO¤NOMA DE ME¤XICO, 76230 QUERE¤TARO, MEXICO
1
2
3
INSTITUTO DE GEOFI¤SICA, UNIVERSIDAD NACIONAL AUTO¤NOMA DE ME¤XICO, 04510 MEXICO CITY, MEXICO
4
LAMONT^DOHERTY EARTH OBSERVATORY AND DEPARTMENT OF EARTH AND ENVIRONMENTAL SCIENCES,
COLUMBIA UNIVERSITY, NEW YORK, NY 10964, USA
RECEIVED FEBRUARY 10, 2009; ACCEPTED OCTOBER 1, 2009
ADVANCE ACCESS PUBLICATION NOVEMBER 6, 2009
The voluminous succession of tholeiitic basalts, calc-alkaline andesites and minor high-K basalts that form the Late Miocene Altos de
Jalisco mafic province of the western Trans-Mexican Volcanic Belt
is interpreted as the magmatic manifestation of a lithospheric dripping event, which removed mantle lithosphere and lower crustal
lithologies beneath the study area. During this process, the release of
fluids from the foundering materials, coupled with mantle upwelling
around the sinking mass, promoted abundant melting of a spinel peridotite and the production of large volumes of tholeiitic magma with
low La/Yb and Gd/Yb ratios. Negative correlations of these ratios
with MgO contents, Nd isotopes and Rb/Nd ratios indicate that
the parental basalts subsequently experienced high-pressure fractional crystallization and contamination with a newly exposed felsic
continental crust, thus producing the more evolved calc-alkaline compositions. Stronger garnet signatures and marked enrichments in
highly incompatible elements in the high-K suite support derivation
from a garnet- and phlogopite-bearing pyroxenitic source, presumably
formed by reaction of mantle peridotites with hydrous silicic melts
derived from the foundering lithologies.This new petrogenetic model
for the Altos de Jalisco volcanic district suggests that the loss of
mafic lower crust during lithospheric dripping might be balanced by
production of abundant flood basalts within continents, and thus
*Corresponding author. Telephone: þ52 55 56224329, 206. Fax: þ52 55
56224317. E-mail: [email protected]
indicates that additional mechanisms may be required for the stabilization of andesitic crust on Earth.
KEY WORDS: continental flood basalts; high-K magmas; lithospheric
removal; mantle;Trans-Mexican Volcanic Belt
I N T RO D U C T I O N
The typical arc-like trace element characteristics of the
bulk continental crust have led to the general consensus
that convergent margins are the principal sites for continental growth (Brown & Rushmer, 2006, and references
therein). Nevertheless, magmatic fluxes at most modern
volcanic arcs are dominantly basaltic, whereas the bulk
continental crust is andesitic (Rudnick & Gao, 2003). If
the origin of the continental crust is ultimately related to
mantle melting processes (see, e.g. Hofmann, 1988), then
the chemical discrepancy between the continents and
mantle-derived magmas indicates that a significant mass
of mafic^ultramafic residues is periodically reintroduced
back into the mantle by either lithospheric delamination
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JOURNAL OF PETROLOGY
VOLUME 50
(i.e. mechanical removal) or dripping (i.e. foundering via
ductile gravitational instabilities) (Arndt & Goldstein,
1989; Kay & Kay, 1993; Jull & Kelemen, 2001; Kelemen
et al., 2003; Aeolus Lee et al., 2006; Go«g u«s & Pysklywec,
2008). In addition, recycling of lower crustal materials
back into the Earth’s interior by lithospheric foundering
seems to be a requirement for the formation of distinctive
mantle heterogeneities, such as the Enriched Mantle I
(EM I) component (characterized by low 206Pb/204Pb,
208
Pb/204Pb and Nd isotope ratios) recognized in the
source of some hotspot magmas (Tatsumi, 2000; Lustrino,
2005).
Lithospheric removal, however, may also represent an
indirect but important means for generating new continental crust, as it has the potential to enhance magmatism by
a combination of decompression and flux melting of the
upper mantle (Elkins-Tanton, 2007). Decompression melting can be induced by mantle upwelling around the sinking instability, whereas flux melting might be promoted
by the release of aqueous fluids from the lithospheric materials as they progressively founder, heat and dehydrate.
Partial melts of the downwelling material might also react
with mantle peridotite to form metasomatic pyroxenitic
lithologies that could be the source of some unusually
enriched mafic magmas (Elkins-Tanton & Grove, 2003).
In this sense, magmatic manifestations of lithospheric
removal have been documented in the Sierra Nevada
(Elkins-Tanton & Grove, 2003) and in the Andean arc
(Kay & Kay, 1993), and typically correspond to smallvolume volcanic episodes characterized by the eruption of
hydrous potassic mafic magmas. Nevertheless, numerical
models predict that this process could also create a variety
of magmatic compositions, and induce the production of
melt volumes up to the size of continental flood basalt provinces (Elkins-Tanton, 2005, 2007; Lustrino, 2005; ElkinsTanton et al., 2006).
In this contribution we report the results of a comprehensive geochemical study of the Late Miocene Altos de
Jalisco volcanic district, a widespread and voluminous
province of mafic plateau lavas emplaced in the western
sector of the Trans-Mexican Volcanic Belt (TMVB; Fig. 1).
This magmatic episode displays geological, volcanological
and compositional features (e.g. considerable volume
and areal distribution, dominant mafic character) that
make it a ‘volcanic anomaly’ compared with the more typical arc-like products of the TMVB; indeed, although they
have a more restricted volume, the Altos de Jalisco
magmas resemble those emplaced within continental
flood basalt provinces. The new geochemical data and the
geological characteristics of the Altos de Jalisco mafic
suites appear to be in agreement with those predicted by
numerical models of lithospheric dripping. These volcanic
successions therefore offer an excellent opportunity to
examine the effects of foundering of mafic and ultramafic
NUMBER 11
NOVEMBER 2009
Fig. 1. Tectonic map of the Mexican convergent margin, showing the
distribution of the subduction-related volcanic provinces of the Sierra
Madre del Sur (SMS), Sierra Madre Occidental (SMO) and TransMexican Volcanic Belt (TMVB). The Altos de Jalisco region is
highlighted with a darker grey color. Important cities are included as
reference: Guadalajara (Gdl), Quere¤taro (Qro) and Mexico City
(MC). The map also indicates the location of Site 487 of the Deep
Sea Drilling Project (DSDP 487), at which a representative section of
the subducted materials has been sampled (LaGatta, 2003). EPR,
East Pacific Rise; MAT, Middle American Trench.
lithologies on magma generation in an active convergent
margin.
G EOLO G IC A L F R A M E WOR K :
TH E MIOCENE TMV B
Although some controversies still exist about the origin of
the TMVB (see, e.g. Sheth et al., 2000; Verma, 2002), it is
generally agreed that this volcanic province is a continental magmatic arc related to the subduction of the Cocos
and Rivera oceanic plates beneath North America along
the Middle American Trench (MAT; Fig. 1; Demant, 1978).
The formation of the TMVB as a distinctive geological
entity dates back to the Early Miocene, when a major geodynamic reorganization involving North America and the
Pacific oceanic plates induced the progressive migration
and counterclockwise rotation of the Paleogene Sierra
Madre del Sur (SMS) and Oligocene Sierra Madre
Occidental (SMO) volcanic arcs (Fig. 1; Mammerickx &
Klitgord, 1982; Stock & Hodges, 1989; Ferrari et al., 1999;
Mora¤n-Zenteno et al., 1999).
The earliest magmatic products of the TMVB crop out
in the central and eastern sectors of the arc in close proximity to the modern volcanic front (Fig. 2a), and are represented by small trondhjemitic domes and typical arc-like
lavas with intermediate compositions (21^15 Ma;
Pasquare' et al., 1991; Capra et al., 1997; Ferrari et al., 2003;
Go¤mez-Tuena et al., 2008). Magmatism subsequently
migrated inland until reaching the northernmost limits of
the arc at 15^10 Ma, with the emplacement of a belt of
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MORI et al.
LITHOSPHERIC REMOVAL AND CRUST FORMATION
Fig. 2. Geological map of the Miocene TMVB, modified from Go¤mez-Tuena et al. (2007). (a) The Early^Middle Miocene activity of the
TMVB was restricted to the central and eastern sectors of the arc. The current location of the volcanic front and important cities (abbreviations
as in Fig. 1) are shown as reference. (b) Late Miocene mafic episode of the TMVB. The study area corresponds to the Altos de Jalisco region.
The field limited by a dotted line represents the inferred maximum areal extent of the mafic province of the western TMVB (Ferrari et al.,
2000). Published ages for the arc-like mafic lavas emplaced between longitudes 1058W and 1018W are almost the same within errors (11^8 Ma;
see text for references), whereas the mafic products emplaced east of longitude 1018W are progressively younger (see text for references), and
also display geochemical features similar to those of intraplate magmas. The diagram of age vs longitude shown in the inset is modified from
Ferrari (2004).
stratovolcanoes, domes and plutonic bodies with andesiti^
dacitic compositions and adakitic geochemical characteristics (Fig. 2a; Pe¤rez-Venzor et al., 1996; Valde¤z-Moreno et al.,
1998; Go¤mez-Tuena & Carrasco-Nu¤n‹ez, 2000; Go¤mezTuena et al., 2003; Verma & Carrasco-Nu¤n‹ez, 2003; Mori
et al., 2007).
Surprisingly, there is no evidence for the existence of an
Early^Middle Miocene arc in the western TMVB,
although subduction beneath the area has been a continuous process at least since the Late Cretaceous (Sdrolias &
Mu«ller, 2006; Go¤mez-Tuena et al., 2007). Indeed, SMOrelated silicic volcanism in this region ended at 22 Ma
(Ferrari et al., 2002), and was followed by an extended
hiatus in effusive activity that lasted for a period of 10
Myr (Fig. 2a). Volcanism subsequently resumed in the
Late Miocene (Fig. 2b), and was characterized by the generation of a widespread igneous province with different
compositional and volcanological characteristics from
those of the typical stratovolcanoes that were emplaced in
the central and eastern TMVB during the same period
(Mori et al., 2007). In fact, this magmatic episode might
be better described as a ‘small-scale continental flood
basalt event’, as it produced large volumes of fissural
mafic lavas (up to 3800 km3) and the formation of plateau structures distributed within an estimated area of up
to 15 500 km2 (Ferrari et al., 2000). Published ages for the
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JOURNAL OF PETROLOGY
VOLUME 50
mafic successions emplaced between longitudes 1058W and
1018W range from 11 to 8 Ma and do not display any
clear migration pattern (Fig. 2b; Nieto-Obrego¤n et al.,
1981; Moore et al., 1994; Ferrari et al., 2000; Rosas-Elguera
et al., 2003), indicating that the ‘flood basalts’ began erupting almost simultaneously over this vast territory, and
that this event occurred within a relatively short time
span, at a calculated eruption rate of 1·66 km3/ka
(Ferrari et al., 2000). Although these magmatic volumes
and eruption rates are unusually high for a normal arc setting, they are still several orders of magnitude smaller
than those of typical continental flood basalt provinces
(i.e. volumes up to 2 106 km3 and eruption rates as high
as 1000 km3/ka; Farmer, 2003).
A mafic episode of much smaller magnitude also
affected the Quere¤taro area in the central portion of the
arc at 8^6 Ma (Fig. 2b; Pasquare' et al., 1991; Valde¤zMoreno et al., 1998; Aguirre-D|¤ az & Lo¤pez-Mart|¤ nez,
2001), and the eastern TMVB between the Late Miocene
and Early Pliocene (7·5^3·5 Ma; Ferrari et al., 2005b;
Fig. 2b). It has been shown that the lava sequences
emplaced in the western and central TMVB display arclike geochemical characteristics (e.g. high Ba/Nb ratios;
Fig. 2b; Ferrari et al., 2000; Mori et al., 2007), whereas an
intraplate affinity (e.g. high Nb contents and low Ba/Nb
ratios) with minor contributions from the subducted slab
has been reported for the progressively younger mafic
products that crop out in the eastern sector of the volcanic
belt (Fig. 2b; Go¤mez-Tuena et al., 2003; Orozco-Esquivel
et al., 2007).
The changes in location and composition of volcanism
during the early geological history of the TMVB are still
not well understood, but they have usually been related to
modifications in the geometry of the subducted slab. In
particular, the northward migration of volcanism that
took place during the Middle Miocene in the central and
eastern TMVB, and the generation of an adakitic belt at a
distance of 500 km from the MAT, have been explained
by invoking a transition to a sub-horizontal subduction
geometry that favored heating and partial melting of the
oceanic crust (Go¤mez-Tuena et al., 2003; Mori et al., 2007).
Yet, there is currently no explanation for the lack of
Early^Middle Miocene volcanism in the western portion
of the arc. In contrast, the unusual massive ‘flooding’ of
continental basalts that marked the inception of magmatic
activity in the western TMVB has been related to extraordinary geological events, such as the arrival of a mantle
plume beneath the region (Moore et al., 1994; Ma¤rquez
et al., 1999) or the ascent of hotter asthenospheric material
through a detached slab (Ferrari, 2004).
G EOLO GY OF T H E ST U DY A R E A
The geological character, age and composition of the
oldest basement rocks below the TMVB are mostly
NUMBER 11
NOVEMBER 2009
unknown. Nevertheless, the Grenvillian Nd model ages of
Cretaceous coastal plutons (0·9^1·2 Ga; Schaaf et al.,
1995) and those of crustal xenoliths in Oligocene volcanic
rocks (1·4^1·6 Ga; El|¤ as-Herrera et al., 1998), as well as
the identification of Precambrian and Paleozoic continental signatures in zircons from Mesozoic sedimentary and
volcanic sequences (Centeno-Garc|¤ a et al., 2008; Martini
et al., 2009), indicate the possible existence of an ancient
continental basement beneath western and southwestern
Mexico.
The Altos de Jalisco volcanic province rests unconformably on two large geological complexes: the Jurassic^
Cretaceous Guerrero Terrane, a tectonostratigraphic
assemblage of island arc magmatic sequences and sedimentary units that constitutes the Mesozoic basement of
western Mexico (Centeno-Garc|¤ a et al., 1993); and a thick
pile of silicic ignimbrites and intermediate lava sequences
related to the Oligocene^Early Miocene activity of the
SMO (Ferrari et al., 2002). Taking into account that the
subaerial expression of the SMO may represent only a
small proportion of the overall volume of mainly mafic
intrusions that probably reside in the deepest portions of
the continental crust, the lowermost crust beneath the
western TMVB may be much younger than the oldest
exposed basement rocks (Bryan et al., 2008).
The Late Miocene mafic province
The vicinity of Guadalajara and the Altos de Jalisco region
host the most spectacular and voluminous manifestations
of the flood basalt event that affected the western TMVB
during the Late Miocene. Volcanic activity in this area
produced 2000 km3 of magmas, and generated basaltic
plateaux distributed over an area of 8000 km2 (Ferrari
et al., 2000). The emplacement of the mafic successions was
favored and controlled by pre-existing zones of crustal
weakness, which were reactivated in a transtensional fashion during the Late Miocene (Ferrari et al., 2000). Lava
flows were mainly extruded as fissure eruptions, or emitted
from small shield volcanoes that cap the lava successions
(Fig. 3). Published K^Ar ages (whole-rock and groundmass) for the lava sequences and the overlying shield volcanoes range from 11 to 8 Ma (Watkins et al., 1971; Damon
et al., 1979; Nieto-Obrego¤n et al., 1981; Nixon et al., 1987;
Castillo-Herna¤ndez & Romero-R|¤os, 1991; Moore et al.,
1994); nevertheless, most ages are concentrated between
10·3 and 9·5 Ma (Fig. 3), indicating that the major magmatic outburst within the study region occurred in less
than 1 Myr.
The most impressive exposures of the mafic sequences
can be observed along the walls of the R|¤o Santiago
canyon, NE of Guadalajara (Fig. 3), and consist of an
700 m thick, monotonous succession of lava flows, each
with a thickness of 2^10 m. The lack of erosional contacts,
paleosols or sedimentary beds between adjacent units confirms that this thick volcanic succession was emplaced
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MORI et al.
LITHOSPHERIC REMOVAL AND CRUST FORMATION
Fig. 3. Geological map of the Guadalajara area and the Altos de Jalisco region, modified from Ferrari et al. (2005a). Also shown are sample
locations and isotopic ages available from the literature (see text for references). The field photograph was taken from the Mirador Huentita¤n,
NE of Guadalajara: it shows a panoramic view of the mafic plateaux that flank the R|¤o Santiago.
within a short time span. These lava sequences flooded a
pre-existing depression cut into Early Miocene ash flows
belonging to the SMO province (Ferrari et al., 2000), and
are overlain by volcanic products of Pliocene to
Quaternary age (Fig. 3; Gilbert et al., 1985). Mafic
sequences with an approximate thickness of 220 m are
also vertically exposed along the R|¤o Verde valley, in the
heart of the Altos de Jalisco (Ferrari et al., 2000). In this
area, they unconformably overlie Early Miocene ignimbrites related to SMO activity (Fig. 3), and are covered by
a characteristic red soil, which represents the typical top
layer of many continental flood basalts (see, e.g. Ollier &
Sheth, 2008).
There are structural, compositional and textural differences between the lava units that build up the Altos de
Jalisco plateaux. Most lava flows appear dense and massive, but some volcanic units have slightly vesicular structures, in which the small vesicles are occasionally filled
with secondary mineralization of calcite or zeolite. The
volcanic products range in composition from basalt to
basaltic andesite with a minor proportion of andesitic
rocks and show a range of petrographic features depending
on their mafic or intermediate character. Lava flows are
commonly affected by surficial alteration, which overprints
the original dark grey color of the rocks with reddish to
greenish banding. In some cases, lava blocks display spheroidal structures produced by more intense weathering
processes.
We carried out an extensive sampling of the Altos de
Jalisco lava flows, collecting rocks with a variety of textures and mineral assemblages (Table 1) representative of
the different volcanic units. Petrographic and geochemical
analyses allowed the recognition of two rock suites within
the study area (Fig. 3), which are classified and described
in subsequent sections.
A N A LY T I C A L M E T H O D S
Major elements were determined by X-ray fluorescence
spectrometry using a Siemens SRS-3000 instrument at the
Laboratorio Universitario de Geoqu|¤ mica Isoto¤pica
(LUGIS) of the Universidad Nacional Auto¤noma de
Me¤xico (UNAM), using procedures of Lozano-Santa
Cruz & Bernal (2005). Trace element data were obtained
by inductively coupled plasma mass spectrometry using a
Thermo Series XII instrument at the Centro de
Geociencias (CGEO) of UNAM, following the sample
preparation and measurement procedures described by
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JOURNAL OF PETROLOGY
VOLUME 50
Mori et al. (2007). Reproducibility of the trace element data
is given by the average concentrations and standard deviations of multiple digestions of the US Geological Survey
rock standards AGV-2, BHVO-2, BCR-2, and the
Geological Survey of Japan JB-2 standard (Table 2).
Sr, Nd and Pb isotopic ratios were measured by thermal
ionization mass spectrometry (TIMS) at LUGIS using a
Finnigan MAT 262 system equipped with eight Faraday
cups. Additional isotopic data were obtained at the
Lamont^Doherty Earth Observatory (LDEO) of
Columbia University using a VG Sector 54-30 TIMS
system equipped with nine Faraday collectors. Sample
preparation and measurement procedures for isotopic
analyses have been described by Schaaf et al. (2005) for
LUGIS, and by Go¤mez-Tuena et al. (2003) for LDEO.
87
Sr/86Sr ratios obtained in both laboratories were normalized to 86Sr/88Sr ¼ 0·1194 and corrected to a NBS-987 standard ratio of 87Sr/86Sr ¼ 0·710230. 143Nd/144Nd ratios were
normalized to 146Nd/144Nd ¼ 0·72190 and corrected to a
La Jolla standard value of 143Nd/144Nd ¼ 0·511860. During
two separate analysis intervals at LUGIS, the measured
values of the NBS-987 standard were 87Sr/86Sr ¼
0·710178 0·000011 (2s, n ¼ 2), and 0·710287 0·000011
(2s, n ¼ 5). The measured 143Nd/144Nd ratio of the La
Jolla standard at LUGIS was 143Nd/144Nd ¼
0·511855 0·000005 (2s, n ¼ 6). During two separate analysis intervals at LDEO, the measured values of the NBS987 standard were 87Sr/86Sr ¼ 0·710245 0·000016 (2s,
n ¼ 4), and 0·710271 0·000014 (2s, n ¼ 6). The measured
143
Nd/144Nd ratio of the La Jolla standard at LDEO was
0·511836 0·000014 (2s, n ¼14). Pb isotope ratios obtained
in both laboratories were corrected to NBS-981 standard
values of 206Pb/204Pb ¼16·9356, 207Pb/204Pb ¼15·4861,
208
Pb/204Pb ¼ 36·7006 (Todt et al., 1996). During two separate analysis intervals at LUGIS, the measured Pb isotope
ratios of the NBS-981 standard were 206Pb/204Pb ¼16·904,
207
Pb/204Pb ¼15·444, 208Pb/204Pb ¼ 36·561 (2s of 0·04%,
0·06%, 0·08%, respectively; n ¼ 7); and 206Pb/204Pb ¼
16·891, 207Pb/204Pb ¼15·427, 208Pb/204Pb ¼ 36·510 (2s of
0·04%, 0·08%, 0·10%, respectively; n ¼10). Pb isotopic
ratios obtained at LDEO were corrected for mass fractionation using the LDEO 207Pb^204Pb double spike (see
Mori et al., 2007, for details). The measured Pb isotope
ratios of the NBS-981 standard at LDEO were
206
Pb/204Pb ¼16·9356, 207Pb/204Pb ¼15·4912, 208Pb/204Pb
¼ 36·7037 (2s of 188, 304, 317 ppm, respectively; n ¼13).
NUMBER 11
NOVEMBER 2009
Rock classification and petrography
The analysis of a large number of samples shows that there
are two rock suites with different petrographic and geochemical characteristics within the Altos de Jalisco mafic
province. Most rocks from the study area display a coherent compositional variation ranging from low-K tholeiitic
basalts to medium-K calc-alkaline andesites (Fig. 4a and
b): this group is subsequently referred as the Altos de
Jalisco suite (AJ suite), as it is widespread throughout the
volcanic district, forming the typical plateaux (Fig. 3). On
the other hand, a small set of trachybasalts and basaltic
trachyandesites characterized by strong potassium enrichment defines a high-K group (Fig. 4a and b); these rocks
are mainly concentrated in the eastern border of the Altos
de Jalisco region (Fig. 3).
Samples belonging to the AJ suite display different
petrographic features according to their degree of differentiation. The tholeiitic basalts are aphyric, and contain
fine-grained (51mm) plagioclase, olivine, clinopyroxene
and oxides that form intergranular textures with subophitic domains (Fig. 4c); most of the larger olivine crystals
show iddingsitization along their rims and fractures; the
smallest crystals tend to be completely altered. In contrast,
calc-alkaline rocks typically have porphyritic textures in
which variable amounts of fine- to medium-grained (up to
2·5 mm) plagioclase, olivine and clinopyroxene phenocrysts are surrounded by a microcrystalline groundmass
with the same paragenesis plus additional orthopyroxene
and Fe^Ti oxides (Fig. 4d). Plagioclase phenocrysts commonly show disequilibrium features such as concentric
zoning, sieve textures or rounded shapes caused by partial
resorption. Olivine phenocrysts show intense iddingsitization, and some crystals also have corroded shapes or
embayments; they are often organized in glomeroporphyritic aggregates that sometimes include clinopyroxene.
The high-K group comprises porphyritic rocks in which
olivine is the only phenocryst phase (Fig. 4e). In these samples, fine- to medium-grained (up to 2·5 mm) olivine
phenocrysts, occasionally forming glomeroporphyritic
aggregates, are embedded in a microcrystalline groundmass of acicular plagioclase, olivine, clinopyroxene and
oxides. Olivine phenocrysts commonly display secondary
alteration along rims and fractures and the smallest crystals are completely iddingsitized; some crystals have
rounded shapes or embayments.
Geochemistry
R E S U LT S
The main phenocryst assemblages and modal proportions
of selected rocks from the study area are given in Table 1.
Major and trace element abundances and Sr, Nd and Pb
isotopic compositions of the analyzed samples are reported
in Tables 2 and 3.
The geochemical characteristics of the various groups that
have been identified within the Late Miocene mafic province are illustrated in the major and trace element variation diagrams of Figs 5 and 6.
Samples from both series display negative correlations
between silica content and TiO2, CaO (not shown),
Fe2O3tot and MgO (Fig. 5a and b); on the other hand,
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MORI et al.
LITHOSPHERIC REMOVAL AND CRUST FORMATION
Table 1: Modal mineralogy of selected samples from the studied rock sequences
Sample
Ol
Pl
Cpx
Opx
Op
Gms
AJ suite
Jal-99-1
Jal-99-2
4·0
0·3
1·5
21·6
—
—
—
—
—
—
94·4
78·1
Jal-99-3
Jal-99-4
9·3
15·7
59·2
64·9
29·4
19·4
—
—
1·5
—
—
—
Jal-99-5
Jal-99-7
16·8
17·1
64·2
66·1
16·7
15·9
—
—
2·3
0·9
—
—
Jal-99-8
Jal-99-9
2·2
0·1
19·3
—
6·0
6·2
—
—
Jal-99-10
Jal-99-11
3·7
23·8
15·2
58·8
5·2
15·6
—
—
—
—
0·7
1·8
75·0
—
Jal-99-18
8·0
8·8
—
—
Jal-99-19
Jal-99-20
18·8
23·8
64·4
59·2
15·8
16·2
—
—
1·0
0·8
Jal-99-22
Jal-06-18
15·4
5·4
64·8
20·7
19·4
5·1
—
—
0·4
—
—
68·8
Jal-06-19
Jal-06-20
3·6
2·5
7·7
4·9
8·6
8·5
—
—
—
—
80·1
84·1
16·3
42·9
1·8
1·9
—
1·0
—
—
81·9
54·2
—
—
—
—
88·7
83·7
1·3
—
—
—
92·0
79·4
Jal-06-21
Jal-06-22
—
—
—
—
—
72·4
93·7
83·2
—
—
Jal-06-23
Jal-06-24
3·0
3·1
8·3
13·2
Jal-06-25
Jal-06-26
3·1
3·3
3·0
9·8
0·6
7·6
AJ-07-4
AJ-07-5
2·9
14·4
3·6
67·6
—
17·6
—
—
—
0·4
93·5
—
AJ-07-6
AJ-07-8
6·8
4·2
3·0
4·4
—
2·1
—
—
—
—
90·2
89·3
AJ-07-9
AJ-07-13
6·0
3·1
6·4
3·1
—
—
—
—
—
—
87·6
93·8
AJ-07-14
AJ-07-15
1·5
4·3
11·5
3·0
—
—
—
—
—
—
87·0
92·7
AJ-07-16
AJ-07-17
0·5
4·1
13·8
2·6
1·7
—
—
—
—
—
74·9
93·3
AJ-07-18
AJ-07-19
13·9
14·4
62·0
63·0
22·7
21·6
—
—
1·4
1·0
—
—
AJ-07-20
AJ-07-21
21·0
16·8
59·5
64·4
17·8
18
—
—
0·1
0·8
—
—
AJ-07-22
AJ-07-23
10·0
21·0
70·0
63·6
18·0
14·2
—
—
1·0
1·0
—
—
AJ-07-24
AJ-07-25
0·9
2·5
4·0
5·7
—
0·6
—
—
—
—
95·1
91·1
AJ-07-26
AJ-07-27
0·5
19·4
4·9
67·8
0·2
11·4
—
—
—
1·4
94·4
—
AJ-07-31
13·6
67·6
17·4
—
AJ-07-32
AJ-07-35
4·4
3·1
4·3
3·7
—
0·7
—
—
—
—
91·3
92·5
AJ-07-36
AJ-07-41
10·6
6·1
68·4
10·9
20·2
—
—
—
0·8
—
—
82·9
AJ-07-42
High-K group
6·0
30·2
—
—
—
63·7
Jal-06-11
Jal-06-14
6·9
8·5
—
—
—
—
—
—
—
—
93·0
91·3
Jal-06-15
Jal-06-16
8·5
7·0
—
—
—
—
—
—
—
—
91·4
92·9
AJ-07-1
AJ-07-2
5·7
6·5
—
—
—
—
—
—
—
—
94·3
93·4
AJ-07-3
AJ-07-10
13·2
8·8
—
—
—
—
—
—
—
—
86·8
91·2
1·6
—
Ol, olivine; Pl, plagioclase; Cpx, clinopyroxene; Opx, orthopyroxene; Op, opaque minerals; Gms,
groundmass. Porphyritic rocks: modal proportions of phenocrysts (40·3 mm) from 1200 points.
Aphyric rocks with intergranular textures: modal proportions of crystals from 1200 points.
2163
JOURNAL OF PETROLOGY
VOLUME 50
NUMBER 11
NOVEMBER 2009
Table 2: Major and trace element analyses of the studied rock suites
Suite:
AJ
AJ
AJ
AJ
AJ
AJ
AJ
AJ
AJ
AJ
AJ
AJ
Sample:
Jal-99-1
Jal-99-2
Jal-99-3
Jal-99-4
Jal-99-5
Jal-99-7
Jal-99-10
Jal-99-11
Jal-99-18
Jal-99-19
Jal-99-20
Jal-99-22
B
B
B
B
B
B
B
B
B
B
B
Rock type: BA
Long. W: 102831·961’ 102830·239’ 102826·204’ 102825·584’ 102814·531’ 102809·302’ 102838·766’ 102838·949’ 103825·195’ 103825·445’ 103825·532’ 103825·041’
Lat. N:
20831·877’ 20832·255’ 20833·937’ 20834·160’ 20834·308’ 20840·496’ 20835·950’ 20836·217’ 21800·255’ 21800·421’ 21802·352’ 21801·028’
Major elements (wt %)
SiO2
53·31
49·20
50·93
49·08
48·69
49·76
49·47
48·12
50·21
47·88
47·12
TiO2
1·25
1·21
1·29
1·29
1·59
1·46
1·62
1·28
1·42
1·25
1·30
48·25
1·27
Al2O3
16·85
19·03
17·43
17·56
17·86
16·65
16·91
16·93
17·37
17·51
16·35
16·56
Fe2O3tot
9·12
9·23
9·74
9·76
10·77
10·78
10·76
10·15
9·66
9·72
11·04
9·89
MnO
0·14
0·13
0·15
0·15
0·15
0·16
0·16
0·14
0·13
0·15
0·15
0·15
MgO
5·72
5·34
6·35
7·74
6·07
6·04
6·66
8·76
6·27
7·89
8·30
8·10
CaO
7·74
10·58
10·10
10·75
9·65
9·55
9·76
9·30
8·50
9·56
9·47
9·09
Na2O
3·43
3·27
3·18
3·15
3·43
3·12
3·34
2·86
3·29
2·81
2·70
2·75
K2O
1·37
0·91
0·77
0·24
0·65
1·18
0·59
0·41
0·98
0·37
0·47
0·75
P2O5
0·48
0·26
0·22
0·17
0·31
0·34
0·26
0·17
0·33
0·16
0·18
0·18
LOI
0·09
0·72
-0·06
0·20
0·85
0·41
0·02
2·29
1·57
3·07
3·82
3·85
Total
99·51
99·88
100·09
100·09
100·02
99·44
99·55
100·41
99·71
100·37
100·90
100·83
Mg-no.
59
57
57
59
60
65
57
67
60
65
64
66
Trace elements (ppm)
Sc
23·6
V
174
Cr
195
24·9
188
96·0
Co
30·1
34·1
Ni
77·4
75·0
Cu
46
39
Zn
89
74
Ga
19·1
Li
Be
290
183
37·9
32·3
180
179
166
297
100
213
186
250
59
42
77
87
18·5
16·3
18·9
10·8
7·4
7·0
1·6
1·3
0·75
546
230
25·7
144
350
25·0
110
41·7
42·7
104
150
50
44
69
86
83
70
19·3
19·1
19·6
7·7
7·3
8·3
1·2
1·2
1·8
51
72
68
83
68
76
72
17·0
19·3
17·4
15·0
15·7
7·6
8·6
18·3
16·9
7·9
10·0
1·1
0·90
1·3
414
482
387
406
134
158
201
42·2
47
10
34·4
45·8
130
40
30·3
151
6·1
348
26·2
114
73·5
42·6
139
11
27·3
35·0
128
16
26·1
27·0
192
66
30·7
30·4
203
54
515
34·7
185
62·8
Sr
24·2
214
66·3
2·2
35·3
231
76·7
159
33·0
225
37·7
17
43·0
28·4
199
41·1
30
Zr
33·8
216
36·4
Rb
Y
35·2
219
17
534
23·5
158
0·76
0·84
3·4
5·1
317
24·1
106
444
24·0
104
0·86
12
804
22·9
106
Nb
11·2
4·43
2·52
4·99
5·85
4·95
5·23
2·09
11·8
2·74
3·34
3·79
Sn
1·4
0·9
1·0
1·0
1·0
1·3
1·1
1·1
0·9
0·7
0·7
0·7
Sb
0·03
0·01
0·05
0·03
0·01
0·01
0·02
0·02
Cs
1·06
0·08
0·55
0·06
0·33
0·16
0·12
Ba
575
0·10
339
La
24·7
13·4
Ce
53·9
30·8
Pr
Nd
7·19
29·7
4·48
19·9
99
5·88
16·0
2·48
12·0
264
264
408
0·11
222
11·1
12·2
15·9
10·4
25·9
28·5
30·9
25·4
3·68
16·7
4·07
18·4
4·88
21·7
3·89
18·3
133
6·67
17·3
2·61
12·6
383
15·5
34·9
4·67
19·9
0·20
142
6·12
16·1
2·44
11·8
0·95
96
6·50
16·9
2·51
12·0
1·01
148
6·98
17·7
2·64
12·6
Sm
6·45
4·76
3·39
4·21
4·55
5·52
4·78
3·53
4·58
3·24
3·31
Eu
1·78
1·54
1·26
1·38
1·56
1·77
1·61
1·28
1·44
1·18
1·16
3·42
1·20
Gd
6·16
4·81
3·97
4·60
4·86
5·93
5·35
4·01
4·63
3·93
3·93
3·94
Tb
0·932
0·724
0·668
0·734
0·757
0·923
0·862
0·688
0·714
0·655
0·646
0·639
Dy
5·31
4·18
4·22
4·50
4·54
5·54
5·27
4·30
4·15
4·19
4·11
3·97
Ho
1·08
0·85
0·89
0·94
0·93
1·16
1·09
0·92
0·85
0·89
0·87
0·83
Er
2·99
2·33
2·48
2·59
2·53
3·21
2·98
2·64
2·32
2·49
2·41
2·28
Yb
2·79
2·08
2·36
2·48
2·35
2·93
2·72
2·58
2·15
2·37
2·30
2·14
Lu
0·416
0·314
0·364
0·368
0·349
0·444
0·403
0·406
0·321
0·354
0·341
0·320
Hf
4·81
3·50
2·55
3·12
3·11
5·11
3·56
2·77
3·36
2·46
2·38
2·52
Ta
0·61
0·28
0·17
0·31
0·37
0·31
0·35
0·14
0·71
0·39
0·21
0·24
Tl
0·021
0·023
0·030
0·025
0·020
0·030
0·021
0·047
0·020
0·019
0·019
0·022
Pb
7·3
3·3
1·1
4·8
3·1
3·4
2·4
1·9
4·3
2·0
1·7
2·0
Th
2·61
2·09
0·37
2·12
1·15
1·93
0·94
0·68
1·42
0·50
0·45
0·55
U
0·824
0·609
0·114
0·567
0·398
0·994
0·160
0·186
0·495
0·137
0·139
0·186
(continued)
2164
MORI et al.
LITHOSPHERIC REMOVAL AND CRUST FORMATION
Table 2: Continued
Suite:
AJ
AJ
AJ
AJ
AJ
AJ
AJ
AJ
AJ
AJ
AJ
AJ
Sample:
Jal-99-23
Jal-06-18
Jal-06-19
Jal-06-21
Jal-06-22
Jal-06-23
Jal-06-24
Jal-06-25
Jal-06-26
AJ-07-5
AJ-07-6
AJ-07-9
BA
BA
BA
TA
BA
BA
A
BA
B
A
BA
Rock type: BTA
Long. W: 103830·404’ 102833·058’ 102833·515’ 102834·564’ 102835·332’ 102835·342’ 102834·815’ 102832·656’ 102835·733’ 102829·900’ 102826·650’ 102827·860’
Lat. N:
20850·886’ 20844·487’ 20844·193’ 20844·740’ 20845·564’ 20845·476’ 20844·882’ 20841·209’ 20842·816’ 20855·360’ 20852·950’ 20851·050’
Major elements (wt %)
SiO2
50·58
53·42
53·06
52·58
57·65
53·72
53·19
56·37
53·76
48·72
55·65
TiO2
1·86
1·02
1·00
1·20
1·00
1·12
1·11
1·18
1·00
1·36
1·08
55·72
1·11
Al2O3
16·53
17·47
17·61
17·64
17·36
17·45
17·27
16·97
17·29
17·16
16·16
16·27
Fe2O3tot
9·76
8·48
8·38
8·61
7·22
8·59
8·55
7·35
8·28
11·18
7·53
7·69
MnO
0·14
0·13
0·13
0·12
0·13
0·12
0·13
0·11
0·11
0·16
0·11
0·16
MgO
4·94
5·36
5·03
5·13
2·81
5·32
5·04
4·19
5·28
5·96
5·69
5·43
CaO
7·53
8·95
8·70
8·71
5·74
8·54
8·51
7·32
8·96
9·97
7·01
7·15
Na2O
3·86
3·19
3·17
3·44
4·56
3·39
3·39
3·34
3·25
2·90
3·09
3·36
K2O
1·24
1·10
1·11
1·21
1·84
1·18
1·22
1·89
1·11
0·99
1·62
1·40
P2O5
0·64
0·29
0·29
0·40
0·57
0·32
0·32
0·47
0·28
0·28
0·33
0·33
LOI
2·91
0·08
0·78
0·21
0·74
0·31
0·48
0·83
0·22
0·52
0·77
0·32
Total
99·98
99·48
99·25
99·25
99·62
100·07
99·20
100·00
99·54
99·20
99·05
98·93
Mg-no.
54
60
58
58
48
60
55
64
62
59
58
57
Trace elements (ppm)
Sc
V
26·6
206
25·4
216
25·1
216
23·4
198
13·6
106
2·13
24·7
201
24·7
207
18·9
156
24·9
208
Cr
40·7
81·6
83·2
77·7
71·9
72·9
55·2
95·3
Co
27·3
30·1
28·7
29·9
13·6
29·3
29·4
22·0
27·1
Ni
29·0
43·8
43·2
55·8
0·9
42·7
42·5
50·4
39·6
Cu
17
51
54
34
9
28
30
30
Zn
93
79
78
87
91
81
84
86
Ga
20·1
20·3
20·3
20·4
21·0
20·1
20·3
Li
18·5
5·4
7·8
7·9
11·2
8·3
8·6
Be
1·6
1·2
1·2
1·5
2·1
1·3
1·3
33·3
18·0
18·4
249
143
147
214
137
149
43·1
26·0
31·0
115
108
114
49
61
35
35
76
83
77
78
20·3
20·3
18·4
19·2
19·2
5·0
7·4
7·9
7·2
8·8
1·8
1·2
1·2
1·3
1·5
Rb
25
16
16
18
32
19
20
30
15
26
26
22
Sr
684
1088
1087
788
715
769
765
735
1047
433
745
728
Y
Zr
32·7
126
26·5
126
22·3
122
25·6
190
31·7
255
24·9
161
27·0
164
26·5
239
21·5
130
26·4
177
23·3
172
30·1
169
Nb
8·57
4·62
4·52
7·76
10·6
6·67
6·71
10·3
4·37
4·31
7·60
7·84
Sn
1·1
0·8
0·7
1·0
1·4
0·9
1·0
1·2
0·7
1·1
0·9
0·8
Sb
0·32
0·05
0·06
0·06
0·09
0·07
0·05
0·07
0·05
0·04
0·03
Cs
1·11
0·37
0·34
0·41
0·48
0·29
0·46
0·45
0·33
Ba
553
423
458
498
685
418
456
693
440
0·09
354
0·49
536
0·38
702
La
21·8
20·7
19·8
21·3
30·0
19·1
20·4
28·7
20·1
10·3
21·3
27·6
Ce
48·8
38·3
40·1
46·7
65·8
39·3
41·0
60·7
40·6
25·2
43·3
50·3
Pr
Nd
6·83
29·3
5·57
23·3
5·78
24·2
6·49
27·2
8·64
34·5
5·73
24·3
6·09
25·7
8·22
33·2
6·01
25·1
3·77
17·4
5·90
24·2
7·15
29·3
Sm
6·63
4·85
5·00
5·95
7·23
5·27
5·68
6·80
5·24
4·55
5·10
6·04
Eu
2·01
1·46
1·50
1·68
1·93
1·54
1·62
1·80
1·54
1·47
1·53
1·78
Gd
6·22
4·64
4·64
5·43
6·47
5·05
5·42
6·01
4·78
4·84
4·74
5·70
Tb
0·914
0·669
0·666
0·803
0·963
0·748
0·813
0·864
0·688
0·753
0·698
0·867
Dy
5·08
3·96
3·87
4·70
5·62
4·47
4·84
4·93
3·95
4·72
3·99
4·92
Ho
0·99
0·82
0·76
0·93
1·12
0·89
0·97
0·94
0·77
0·95
0·79
0·98
Er
2·58
2·26
2·07
2·52
3·11
2·43
2·62
2·58
2·10
2·62
2·17
2·68
Yb
2·14
2·00
1·85
2·33
3·01
2·22
2·41
2·32
1·91
2·45
1·99
2·46
Lu
0·288
0·306
0·277
0·348
0·452
0·335
0·360
0·343
0·282
0·362
0·302
0·361
Hf
2·61
3·20
3·12
4·37
5·63
3·91
3·94
5·39
3·29
4·57
3·97
3·90
Ta
0·56
0·26
0·26
0·45
0·63
0·39
0·39
0·58
0·25
0·27
0·45
0·47
Tl
0·024
0·078
0·055
0·010
0·019
0·005
0·050
0·090
0·075
0·027
0·044
0·026
Pb
5·8
5·4
5·3
7·4
9·2
7·3
6·5
8·4
4·8
2·3
6·3
5·7
Th
2·12
2·23
2·24
2·65
3·44
1·96
1·97
2·70
2·17
0·95
2·42
2·34
U
0·810
0·639
0·627
0·79
1·04
0·565
0·622
0·820
0·646
0·314
0·752
0·731
(continued)
2165
JOURNAL OF PETROLOGY
VOLUME 50
NUMBER 11
NOVEMBER 2009
Table 2: Continued
Suite:
AJ
AJ
AJ
AJ
AJ
AJ
AJ
AJ
AJ
AJ
AJ
AJ
Sample:
AJ-07-13
AJ-07-14
AJ-07-16
AJ-07-17
AJ-07-21
AJ-07-24
AJ-07-26
AJ-07-29
AJ-07-30
AJ-07-31
AJ-07-35
AJ-07-36
BA
A
BA
B
BA
BA
B
B
B
B
BA
Rock type: A
Long. W: 102847·250’ 102847·550’ 102852·100’ 102850·490’ 102851·390’ 102847·250’ 102845·000’ 102843·315’ 102844·090’ 102849·070’ 103812·880’ 103812·880’
Lat. N:
20856·860’ 20855·770’ 20852·700’ 20853·520’ 20842·000’ 20839·075’ 20839·400’ 20837·850’ 20836·260’ 20833·620’ 20841·000’ 20841·000’
Major elements (wt %)
SiO2
57·45
53·95
57·24
52·08
49·96
54·53
53·54
49·94
49·01
50·83
48·69
TiO2
0·91
1·21
1·18
1·18
1·28
1·01
1·19
1·48
1·01
1·62
1·46
51·34
1·16
Al2O3
16·97
17·57
17·58
16·70
16·85
18·64
17·90
16·50
18·00
16·94
17·16
17·91
Fe2O3tot
7·05
8·53
7·68
9·61
9·75
7·84
8·35
10·82
9·61
10·62
10·26
9·33
MnO
0·10
0·11
0·12
0·14
0·15
0·12
0·12
0·16
0·14
0·16
0·12
0·14
MgO
4·77
4·09
2·58
5·67
6·44
3·82
4·48
6·77
6·94
5·86
6·05
5·50
CaO
7·25
7·85
6·44
9·11
9·97
7·58
8·26
10·71
9·96
8·53
9·40
9·07
Na2O
3·35
3·06
4·15
3·01
3·06
3·59
3·33
3·18
2·85
3·24
3·14
3·46
K2O
1·48
2·02
1·62
1·04
0·83
1·45
1·81
0·46
0·59
1·08
1·02
1·07
P2O5
0·31
0·41
0·38
0·31
0·22
0·29
0·45
0·20
0·15
0·37
0·31
0·28
LOI
Total
Mg-no.
0·62
0·95
0·70
0·32
0·68
0·33
0·02
0·94
1·02
0·00
1·60
0·47
100·26
99·75
99·65
99·16
99·19
99·20
99·45
101·15
99·28
99·27
99·20
99·72
53
44
58
61
53
56
63
56
58
58
61
59
Trace elements (ppm)
Sc
V
19·3
149
19·2
185
Cr
92·6
50·4
Co
22·4
23·4
Ni
72·9
27·5
Cu
41
32
Zn
73
Ga
19·1
Li
7·8
Be
1·5
18·0
152
31·8
201
162
147
18·1
176
22·1
200
10·2
23·7
34·0
36·6
25·1
25·9
58·4
89·9
37·2
29·5
54
58
71
43
57
88
76
77
75
82
20·1
21·3
18·8
18·4
20·5
5·7
11·3
6·8
8·6
5·5
1·5
5·58
27·7
196
17·2
5·13
1·7
1·3
1·1
1·3
48
83
81
79
100
85
82
21·1
18·5
19·0
20·2
19·4
19·2
7·1
7·8
12·8
9·5
8·5
15·4
1·8
Sr
620
831
812
552
449
871
1020
187
194
163
143
141
137
35·9
212
41·5
129
1·0
7·3
348
32·6
155
1·1
1·6
37·9
98·9
60
35
Zr
118
44
19
24·2
142
62
15
31·6
126
72
17
41·2
192
23·1
188
81·1
29
33·1
26·2
216
95·5
48
21·4
26·4
206
35·3
29
23·9
23·9
198
34·2
Rb
Y
32·1
217
114
1·2
33·1
79·7
1·3
16
19
21
19
582
533
548
626
20·7
107
34·5
180
24·1
144
22·1
134
Nb
7·56
6·76
8·50
8·03
4·81
5·45
7·05
3·61
5·08
9·90
11·1
6·68
Sn
0·9
0·9
0·8
0·8
0·9
0·8
1·0
1·0
0·7
1·1
1·0
0·9
Sb
0·04
0·04
0·04
0·03
0·04
0·01
0·03
0·02
0·06
0·02
0·06
Cs
0·69
0·73
0·26
0·31
0·33
0·42
0·75
4·04
0·38
3·74
Ba
499
631
519
372
323
520
725
La
21·4
22·7
22·6
21·1
12·8
19·6
32·2
Ce
43·0
49·3
47·6
36·6
28·0
39·9
62·7
Pr
Nd
5·87
24·0
6·72
27·7
6·60
27·0
5·37
22·7
3·93
17·8
5·40
22·5
8·68
35·1
0·05
166
8·10
21·2
3·13
15·2
349
454
345
0·84
428
10·7
18·7
14·3
13·3
24·7
41·3
32·5
30·0
3·48
15·9
5·77
25·1
4·37
19·3
4·12
18·3
Sm
5·10
5·84
6·07
5·10
4·42
4·80
7·25
4·11
3·82
6·00
4·56
4·31
Eu
1·39
1·70
1·69
1·56
1·45
1·44
2·00
1·40
1·28
1·77
1·46
1·42
Gd
4·77
5·09
5·84
5·58
4·73
4·60
6·44
4·73
3·93
6·19
4·63
4·34
Tb
0·716
0·712
0·859
0·861
0·765
0·686
0·926
0·794
0·612
0·970
0·727
0·675
Dy
4·13
3·93
5·21
5·40
4·68
4·00
5·12
4·97
3·68
5·85
4·31
3·95
Ho
0·82
0·76
1·06
1·16
0·97
0·80
1·03
1·03
0·74
1·18
0·86
0·79
Er
2·26
2·03
2·93
3·28
2·68
2·21
2·87
2·85
2·00
3·25
2·35
2·13
Yb
2·15
1·83
2·70
2·95
2·58
2·05
2·50
2·78
1·89
3·05
2·19
2·01
Lu
0·322
0·272
0·404
0·451
0·386
0·306
0·382
0·407
0·284
0·452
0·325
0·298
Hf
4·11
4·80
3·87
3·26
3·13
3·38
4·92
3·02
2·62
4·16
3·27
3·19
Ta
0·43
0·39
0·54
0·48
0·31
0·30
0·40
0·27
0·36
0·59
0·66
0·47
Tl
0·026
0·034
0·021
0·021
0·022
0·028
0·036
0·019
0·021
0·024
0·021
0·027
Pb
7·1
6·6
5·1
4·3
3·7
7·1
7·8
2·2
4·1
5·1
3·9
4·7
Th
2·35
2·70
2·33
1·63
2·60
1·99
4·19
0·47
0·97
1·57
1·33
1·46
U
0·757
1·07
0·735
0·525
0·684
0·605
1·30
0·164
0·378
0·517
0·477
‘0·561
(continued)
2166
MORI et al.
LITHOSPHERIC REMOVAL AND CRUST FORMATION
Table 2: Continued
Suite:
AJ
AJ
AJ (enr.)
AJ (enr.)
AJ (enr.)
AJ (enr.)
AJ (enr.)
AJ (enr.)
AJ (enr.)
AJ (enr.)
AJ (enr.)
AJ (enr.)
Sample:
AJ-07-41
AJ-07-42
Jal-99-8
Jal-99-9
Jal-06-20
AJ-07-4
AJ-07-8
AJ-07-15
AJ-07-18
AJ-07-19
AJ-07-20
AJ-07-22
BA
BA
BA
BA
BTA
BA
BA
B
B
B
BA
Rock type: BA
Long. W: 103817·690’ 103817·960’ 102831·774’ 102832·562’ 102834·952’ 102831·490’ 102826·240’ 102848·290’ 102850·800’ 102847·950’ 102849·850’ 102847·700’
Lat. N:
20844·000’ 20844·290’ 20843·140’ 20841·698’ 20843·831’ 20856·050’ 20850·600’ 20853·070’ 20850·760’ 20850·295’ 20847·420’ 20840·185’
Major elements (wt %)
SiO2
50·76
51·93
53·85
53·60
54·09
53·20
55·72
52·41
48·12
49·29
48·52
TiO2
1·10
0·91
1·02
1·03
0·98
1·31
1·09
1·18
1·42
1·32
1·29
51·67
1·11
Al2O3
17·79
18·60
17·18
17·44
17·18
18·16
16·35
16·52
17·11
17·14
17·55
17·89
Fe2O3tot
9·05
8·34
8·34
8·32
8·32
8·02
7·84
9·50
11·67
10·42
10·83
9·32
MnO
0·13
0·14
0·12
0·12
0·12
0·12
0·14
0·14
0·16
0·17
0·17
0·16
MgO
6·58
5·59
5·36
5·59
5·62
4·93
5·89
6·03
7·42
6·49
6·78
5·22
CaO
8·71
8·68
8·57
8·78
8·67
7·83
6·97
9·07
10·20
10·05
10·00
9·03
Na2O
3·16
3·28
3·37
3·37
3·16
3·42
3·25
3·02
3·05
2·95
2·94
3·11
K2O
0·88
0·91
1·03
0·99
1·12
1·89
1·51
1·13
0·36
0·94
0·41
1·02
P2O5
0·25
0·17
0·32
0·31
0·28
0·34
0·37
0·30
0·19
0·29
0·17
0·28
LOI
0·86
0·74
0·57
0·28
0·69
0·05
0·37
0·00
-0·22
0·24
0·44
0·58
Total
99·26
99·30
99·71
99·83
100·23
99·27
99·48
99·33
99·48
99·31
99·09
99·38
Mg-no.
63
61
60
61
59
64
60
60
59
59
57
61
Trace elements (ppm)
Sc
V
22·5
181
29·7
25·8
229
201
115
101
25·2
199
27·3
31·4
33·5
35·2
200
215
226
216
181
151
196
177
242
26·7
213
82·4
30·3
40·2
27·7
28·1
30·9
26·5
Ni
80·7
83·1
46·6
45·2
50·8
72·3
Cu
53
58
54
54
48
30
30
60
71
64
74
56
Zn
83
68
76
83
82
77
83
78
82
92
81
85
Ga
19·2
17·7
20·5
20·5
19·8
20·5
18·9
19·0
18·9
18·5
18·6
20·8
Li
13·3
8·1
7·5
7·0
7·2
7·8
8·3
6·8
6·2
8·7
10·4
6·9
0·79
1·3
2·1
1·6
1·4
1·5
1·4
0·91
2·3
1·0
Rb
15
Sr
627
5·5
543
Y
17·7
21·8
Zr
88
83
78·4
18·3
145
Co
0·91
109
16·1
173
Cr
Be
95·2
25·9
209
28·8
141
32·9
56·0
15
14
19
33
25
18
991
1046
1039
1110
705
577
38·5
138
68·6
131
72·8
119
39·6
208
42·8
176
58·2
142
48·0
118
4·7
355
37·5
123
42·0
87·2
23
443
53·0
156
47·0
148
4·4
396
59·1
135
93·6
36·5
65·2
1·1
10
805
175·0
139
Nb
3·32
3·51
4·94
4·74
4·58
6·01
8·37
7·97
3·58
5·54
2·96
4·61
Sn
0·5
0·5
1·0
0·6
0·8
0·7
0·8
0·8
0·7
0·9
0·8
0·8
Sb
0·03
0·01
0·01
0·06
0·01
0·06
0·03
Cs
2·10
0·27
0·29
0·99
0·26
0·43
0·31
0·08
0·18
Ba
La
Ce
Pr
Nd
337
8·55
19·5
2·74
12·6
0·07
181
8·23
18·1
2·66
12·5
484
499
479
721
707
375
0·03
144
412
0·04
234
33·5
78·7
34·9
41·5
32·1
40·6
12·4
25·6
20·0
47·5
44·7
44·4
51·9
48·8
36·4
19·3
37·4
24·0
16·9
11·4
62·0
48·3
13·2
11·6
8·89
36·5
Sm
3·10
3·16
7·28
Eu
1·07
1·14
2·16
Gd
3·25
3·54
7·21
Tb
0·507
0·567
1·03
Dy
3·11
3·57
5·77
2·27
12·2
1·92
10·8
3·27
11·6
1·69
10·2
9·36
39·3
7·45
30·4
9·48
37·8
3·24
15·5
7·93
32·8
4·73
21·0
7·46
6·22
8·49
3·86
8·20
5·25
2·48
1·77
2·04
1·39
2·21
1·70
7·13
6·39
9·03
4·78
8·72
6·49
0·942
0·931
1·36
0·762
1·41
1·04
5·05
5·61
8·35
4·71
8·75
6·62
0·12
517
72·9
34·1
12·2
51·1
10·9
3·40
15·4
2·22
14·3
Ho
0·63
0·75
1·20
2·16
2·21
1·02
1·20
1·72
1·01
1·79
1·44
3·37
Er
1·74
2·09
3·29
5·95
5·95
2·72
3·36
4·77
2·75
5·04
4·00
9·73
Yb
1·67
2·02
2·84
5·37
5·47
2·13
2·88
4·35
2·46
4·94
3·57
7·76
Lu
0·249
0·304
0·430
0·764
0·836
0·329
0·446
0·639
0·365
0·723
0·534
1·26
Hf
2·22
2·01
3·43
3·28
2·93
5·13
4·10
3·24
2·66
3·71
2·65
3·43
Ta
0·21
0·23
0·28
0·27
0·26
0·35
0·50
0·48
0·23
0·34
0·20
0·28
Tl
0·021
0·018
0·026
0·026
0·076
0·034
0·034
0·029
0·020
0·024
0·020
0·022
Pb
4·5
2·1
5·2
4·9
5·7
4·5
6·9
4·1
1·3
2·9
2·4
4·6
Th
1·16
1·01
2·10
2·13
2·25
2·93
2·52
1·65
0·32
1·27
1·62
1·75
U
0·460
0·315
0·628
0·613
0·626
1·05
0·755
0·534
0·119
0·365
0·210
0·456
(continued)
2167
JOURNAL OF PETROLOGY
VOLUME 50
NUMBER 11
NOVEMBER 2009
Table 2: Continued
Suite:
AJ (enr.)
AJ (enr.)
AJ (enr.)
AJ (enr.)
HK
HK
HK
HK
HK
HK
HK
HK
Sample:
AJ-07-23
AJ-07-25
AJ-07-27
AJ-07-32
Jal-06-11
Jal-06-14
Jal-06-15
Jal-06-16
AJ-07-1
AJ-07-2
AJ-07-3
AJ-07-10
BA
BA
BA
BTA
B
BTA
B
BTA
BTA
TB
BA
Rock type: B
Long. W: 102847·720’ 102845·500’ 102843·610’ 102856·510’ 102806·328’ 102809·136’ 102813·232’ 102814·790’ 102830·790’ 102829·650’ 102829·500’ 102848·900’
Lat. N:
20839·490’ 20839·420’ 20839·560’ 20832·160’ 20839·271’ 20834·035’ 20837·860’ 20839·191’ 21801·050’ 21804·850’ 21804·100’ 20858·190’
Major elements (wt %)
SiO2
50·28
53·28
51·39
51·58
53·75
50·33
51·00
51·10
53·09
52·89
49·92
TiO2
1·22
1·16
1·20
1·37
1·29
1·29
1·34
1·25
1·37
1·39
1·47
51·99
1·25
Al2O3
17·27
17·24
17·20
17·51
14·17
16·13
14·42
15·32
16·37
16·30
14·31
15·77
Fe2O3tot
9·21
8·44
9·43
9·41
7·98
9·76
9·59
9·62
9·03
9·25
10·00
9·25
MnO
0·15
0·13
0·18
0·13
0·11
0·15
0·13
0·14
0·12
0·13
0·14
0·13
MgO
6·27
4·89
5·40
5·98
6·79
6·66
8·54
7·91
5·60
5·57
7·80
6·92
CaO
8·59
8·41
9·44
8·63
7·93
9·16
7·98
9·11
8·03
7·93
9·48
8·32
Na2O
3·26
3·27
3·13
3·58
2·25
2·48
2·43
2·60
2·95
2·84
2·78
2·66
K2O
1·58
1·67
1·13
1·08
4·23
2·26
2·98
2·22
2·69
2·72
2·50
2·29
P2O5
0·38
0·39
0·30
0·28
0·63
0·48
0·66
0·46
0·45
0·46
0·47
0·49
LOI
1·96
0·26
0·62
-0·23
0·85
1·27
0·91
0·26
0·20
0·19
1·14
0·34
Total
100·18
99·13
99·41
99·33
99·98
99·97
99·98
99·99
99·90
99·67
100·01
99·41
Mg-no.
61
57
57
60
66
61
67
66
59
58
65
64
Trace elements (ppm)
Sc
23·6
23·7
28·0
23·7
23·8
34·5
26·3
30·0
22·3
22·7
23·1
22·6
V
200
200
211
191
302
274
246
246
227
234
223
218
Cr
160
36·5
125
78·4
251
285
387
382
131
143
374
221
Co
37·0
29·7
37·7
34·8
32·7
37·4
42·7
39·2
29·7
30·2
39·7
33·1
Ni
135
40·4
42·4
80·0
181
116
299
155
78·6
78·8
196
108
Cu
52
63
46
49
70
55
54
49
57
57
63
46
Zn
114
81
85
80
86
81
86
80
89
90
89
89
Ga
20·4
20·3
19·6
19·4
21·0
18·6
17·9
18·3
21·3
21·2
19·3
19·9
Li
9·3
8·2
6·9
8·2
7·4
9·0
8·4
7·8
9·2
7·8
9·9
11·2
Be
3·4
2·0
1·8
1·4
3·1
2·4
3·0
2·3
2·0
2·0
1·7
1·8
Rb
29
32
17
18
115
62
107
73
60
62
55
73
Sr
916
843
741
608
731
462
548
471
648
646
631
814
Y
63·7
64·9
116·4
48·2
32·6
28·3
28·4
25·7
23·9
23·9
24·2
22·1
Zr
202
213
162
154
510
324
371
281
339
311
277
297
Nb
7·46
7·46
5·92
9·89
5·00
6·46
7·09
5·16
8·51
8·50
12·3
6·16
Sn
1·1
1·0
0·9
0·9
2·5
1·8
2·3
1·9
1·4
1·4
1·2
1·3
Sb
0·02
0·03
0·01
0·05
0·08
0·05
0·06
0·07
0·04
0·04
0·02
0·08
Cs
0·23
0·54
0·19
0·34
0·83
0·28
0·44
0·67
0·67
0·76
0·36
0·80
Ba
674
625
467
387
1134
629
880
662
770
747
901
698
La
37·9
61·3
43·9
26·6
38·7
22·9
23·6
16·9
29·2
29·4
33·6
27·2
Ce
52·6
59·5
46·7
34·2
79·2
52·4
56·1
40·3
64·4
64·5
74·3
61·0
Pr
10·5
14·0
9·26
7·07
11·9
7·72
7·74
5·68
8·59
8·74
9·83
8·30
Nd
42·3
55·4
38·8
30·5
50·8
33·5
34·1
25·2
35·3
35·8
41·4
34·9
Sm
9·96
11·8
8·62
6·83
10·6
7·63
8·25
6·18
7·28
7·46
8·45
7·25
Eu
2·56
2·63
2·36
2·16
2·64
2·08
2·27
1·81
2·02
2·02
2·30
2·01
Gd
10·3
11·5
10·9
7·55
8·27
6·64
7·37
5·84
6·06
6·19
6·76
5·97
Tb
1·61
1·69
1·57
1·17
1·07
0·949
1·02
0·854
0·849
0·856
0·931
0·821
Dy
9·94
9·75
9·69
7·02
5·43
5·36
5·54
4·94
4·54
4·60
4·71
4·29
Ho
2·09
1·92
2·23
1·47
0·99
1·03
1·01
0·94
0·85
0·85
0·87
0·78
Er
6·18
5·22
6·32
4·04
2·58
2·77
2·60
2·52
2·28
2·28
2·26
2·08
Yb
6·63
4·70
5·04
3·59
2·05
2·50
2·20
2·24
2·01
2·00
1·93
1·83
Lu
1·06
0·679
0·797
0·539
0·307
0·372
0·323
0·330
0·296
0·294
0·279
0·266
Hf
4·96
5·15
3·86
3·48
14·0
8·09
9·33
7·37
8·02
7·82
6·76
7·57
Ta
0·44
0·74
0·35
0·61
0·29
0·39
0·42
0·31
0·50
0·50
0·72
0·35
Tl
0·030
0·034
0·022
0·023
0·89
0·34
0·34
0·29
0·052
0·050
0·042
0·059
Pb
6·5
6·3
4·6
5·0
8·2
4·9
5·6
6·0
7·8
7·7
5·6
6·2
Th
3·06
3·35
2·32
1·86
6·17
3·99
3·55
3·13
4·78
4·79
4·60
3·84
U
1·10
1·12
0·685
0·645
3·65
2·19
2·24
1·70
2·05
2·09
1·65
1·76
(continued)
2168
MORI et al.
LITHOSPHERIC REMOVAL AND CRUST FORMATION
Table 2: Continued
Suite:
AGV-2
AGV-2
BCR-2
BCR-2
BHVO-2
BHVO-2
JB-2
JB-2
Sample:
Rock type:
n ¼ 52
1s
n ¼ 53
1s
n ¼ 55
1s
n ¼ 50
1s
Long. W:
Lat. N:
Major elements (wt %)
SiO2
TiO2
Al2O3
Fe2O3tot
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Total
Mg-no.
Trace elements (ppm)
Sc
13·6
0·6
33·3
0·5
V
Cr
116
16·7
3
0·4
415
16·2
7
0·4
317
280
6
4
Co
Ni
15·7
18·5
0·2
0·4
36·7
12·4
0·5
0·4
44·4
116
0·8
2
Cu
Zn
53
86
2
2
2
1
137
103
6
2
Ga
Li
20·9
10·9
0·3
0·2
22·3
9·3
0·2
0·1
Be
Rb
2·4
69
0·1
1
2·3
48
0·1
1
Sr
Y
658
20·1
4
0·3
335
37·2
6
0·5
397
26·7
5
0·3
Zr
Nb
233
14·9
4
0·3
186
13·1
3
0·3
171
19·4
4
0·6
48
0·61
1
0·08
25
129
31·2
21·7
4·7
1·17
9·3
0·4
0·2
0·1
0·04
0·3
54·2
0·5
577
25·2
7
0·6
36·1
13·9
0·7
0·5
222
103
16·6
8·2
0·31
6·5
178
24·2
8
2
0·2
0·1
0·02
0·2
2
0·3
Sn
Sb
2·0
0·54
0·2
0·07
2·2
0·35
0·1
0·07
1·8
0·13
0·1
0·02
0·6
0·29
0·1
0·04
Cs
Ba
1·16
1142
0·02
7
1·14
680
0·02
11
0·09
129
0·01
4
0·79
216
0·02
7
La
Ce
38·1
69·1
0·2
0·8
24·9
52·3
0·2
0·8
15·1
37·4
0·2
0·8
2·27
6·7
0·07
0·3
Pr
Nd
8·23
30·3
0·06
0·3
6·83
28·4
0·06
0·3
5·34
24·4
0·07
0·3
1·11
6·4
0·02
0·1
Sm
Eu
5·64
1·56
0·07
0·02
6·70
1·91
0·06
0·02
6·22
1·99
0·05
0·02
2·27
0·82
0·03
0·01
Gd
4·66
0·06
6·82
0·06
6·24
0·05
3·25
0·04
Tb
Dy
0·661
3·56
0·008
0·05
1·068
6·55
0·008
0·07
0·948
5·42
0·007
0·06
0·573
4·02
0·006
0·06
Ho
Er
0·69
1·83
0·01
0·02
1·32
3·65
0·02
0·04
1·01
2·53
0·02
0·03
0·88
2·52
0·02
0·04
Yb
Lu
1·66
0·252
0·02
0·003
3·43
0·508
0·04
0·005
2·03
0·283
0·02
0·004
2·56
0·391
0·03
0·004
Hf
Ta
5·03
0·91
0·05
0·02
4·74
0·83
0·05
0·01
4·25
1·25
0·05
0·02
1·43
0·040
0·02
0·005
Tl
Pb
0·26
13·4
0·04
0·2
0·25
10·4
0·04
0·4
0·023
1·8
0·003
0·1
0·038
5·3
0·004
0·1
Th
U
6·25
1·90
0·07
0·04
6·02
1·69
0·09
0·03
1·23
0·41
0·01
0·01
0·26
0·148
0·02
0·004
Reproducibility of trace element data is given by the average concentrations and standard deviations of multiple digestions
of the US Geological Survey rock standards AGV-2, BHVO-2 and BCR-2, and the Geological Survey of Japan JB-2
standard. AJ, AJ suite; AJ (enr.), REE-enriched samples of the AJ suite; HK, high-K group. Rock names are given
according to the total alkali vs silica diagram of Le Bas et al. (1986) (Fig. 4a). B, basalt; BA, basaltic andesite; A, andesite;
TB, trachybasalt; BTA, basaltic trachyandesite; TA, trachyandesite. Mg-number ¼ 100Mg/(Mg þ 0·85Fetot), molar.
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JOURNAL OF PETROLOGY
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Table 3: Sr, Nd and Pb isotopic compositions of selected rocks from the study area
Sample
87
Sr/86Sr
2s mean
143
Nd/144Nd
2s mean
eNd
206
Pb/204Pb
207
Pb/204Pb
208
Pb/204Pb
AJ suite
Jal-99-1
0·703844
6
0·512900
28
5·1
18·7009
15·5965
38·4696
Jal-99-2
0·703627
7
0·512945
12
6·0
18·6174
15·5640
38·2808
Jal-99-3
0·703605
7
0·512928
9
5·7
18·6700
15·5879
38·4244
Jal-99-4
0·703216
7
0·512984
14
6·7
18·6451
15·5570
38·2673
Jal-99-5y
0·703516
10
0·512837
6
3·9
18·680
15·594
38·368
Jal-99-7
0·703729
6
0·512896
9
5·0
18·6996
15·5776
38·3762
Jal-99-8
0·703356
6
0·512877
8
4·7
18·6401
15·5725
38·3421
Jal-99-9
0·703395
7
0·512867
7
4·5
18·6429
15·5774
38·3582
Jal-99-10
0·703388
6
0·512940
7
5·9
18·6938
15·5750
38·3729
Jal-99-11
0·703470
6
0·512984
9
6·7
18·6063
15·5640
38·2524
Jal-99-18y
0·703987
9
0·512789
7
2·9
18·730
15·590
38·457
Jal-06-22y
0·703677
10
0·512800
5
3·2
18·718
15·623
38·547
AJ-07-6y
0·703663
11
0·512798
6
3·1
18·703
15·595
38·464
AJ-07-16y
0·703356
10
0·512800
5
3·2
18·701
15·585
38·411
AJ-07-24y
0·703552
10
0·512813
5
3·4
18·675
15·598
38·437
18·576
15·539
38·107
High-K group
Jal-06-11y
0·703981
11
0·512983
5
6·7
Jal-06-15y
0·703934
9
0·512898
5
5·1
AJ-07-2y
0·703948
11
0·512931
5
5·7
18·638
15·575
38·301
AJ-07-3y
0·703892
11
0·512899
4
5·1
18·640
15·559
38·250
AJ-07-10y
0·703904
9
0·512904
5
5·2
18·636
15·558
38·243
Reported values are not age-corrected and are taken as initial. The 2s mean for single Sr and Nd measurements are
multiplied by 106. Reproducibility for Pb isotopes is given by the 2s mean of multiple measurements of NBS-981 standard
(see text for details). At LUGIS, Sr and Nd isotopes were measured by static multicollection, with each analysis consisting
of 60 isotopic ratios. At LDEO, Sr and Nd were measured by dynamic multicollection, with each analysis consisting of 120
isotopic ratios. At both laboratories, Pb isotopes were measured by static multicollection, with each analysis consisting of
100 isotope ratios.
Sample analyzed at LDEO.
ySample analyzed at LUGIS.
Al2O3 and Na2O abundances remain almost constant with
differentiation in both sequences, although Na2O shows a
weak positive correlation with SiO2 in the AJ suite
(Fig. 5c and d). Despite the analogous trends in most
major element variation diagrams, the high-K group has
similar MgO contents to those of the AJ tholeiitic samples
but with higher SiO2, and also displays lower Al2O3 and
Na2O and higher P2O5 (not shown) than the AJ suite at
similar SiO2.
Rocks of the AJ suite show negative correlations between
compatible trace elements (i.e. Ni; Fig. 6a) and SiO2,
although the most evolved samples have variable Ni contents. Incompatible elements such as Rb, Th and La generally increase with the degree of differentiation in this
group (Fig. 6b and c), whereas relatively less incompatible
elements such as Zr, Hf, Gd, Yb and Y remain almost constant with increasing silica content (Fig. 6d). Within the
AJ volcanic suite, the tholeiitic samples display the lowest
abundances of large ion lithophile elements (LILE) and
light rare earth elements (LREE), as well as the lowest
concentrations of high field strength elements (HFSE).
Rocks of the high-K series do not display clear correlations
between trace elements and silica (Fig. 6a^c).
Nevertheless, they are notable for their higher Ni, V and
Cr abundances, their marked enrichments in highly
incompatible elements such as Rb, Th and U, and for displaying higher Zr and Hf concentrations than the AJ suite
at similar silica contents.
Some samples of the AJ suite display prominent negative
Ce anomalies in mid-ocean ridge basalt (MORB)-normalized trace element diagrams and unusual Y and REE
enrichments (Figs 6c, d and 7b). Interestingly, the major
and other trace elements in these rocks do not show atypical concentrations relative to the other samples (they all
2170
MORI et al.
LITHOSPHERIC REMOVAL AND CRUST FORMATION
Fig. 4. (a) The total alkalis vs SiO2 diagram (Le Bas et al.,1986; alkaline^subalkaline division from Irvine & Baragar, 1971) and (b) the K2O vs
SiO2 discrimination diagram (Le Maitre et al., 1989; Rickwood, 1989) allow the recognition of two main magmatic groups within the study
area. Most rocks range in composition from low-K basalt to medium-K andesite and define the AJ suite, whereas a few samples with potassium
enrichments define the high-K group. Abundances of oxides are normalized to 100% volatile-free. Within the AJ suite, tholeiitic basalts are
aphyric and display intergranular-subophitic textures (c), whereas the calc-alkaline intermediate rocks are mainly porphyritic (d). (e) High-K
samples also display porphyritic textures, but olivine is the only phenocryst phase in these rocks. Microphotographs were taken under planepolarized light. Cpx, clinopyroxene; Ol, olivine; Pl, plagioclase.
plot within the range of ‘normal’AJ rocks), neither do they
show any petrographic differences. The origin of the anomalous REE abundances will be discussed below.
Both rock sequences display trace element patterns that
are typical of arc magmas (Fig. 7), such as enrichments in
LILE and Pb with respect to the HFSE, and fractionated
REE patterns showing higher LREE contents relative to
the heavy REE (HREE). Within the AJ suite, REE fractionation is relatively low in the most mafic samples, and progressively increases in the calc-alkaline rocks at almost
constant Yb contents (Fig. 7a). The high-K group displays
higher abundances of LILE and Th than the AJ suite at
similar HFSE contents, and also extends to higher La/Yb
and Gd/Yb ratios (Fig. 7c).
Relationships between the Sr, Nd and Pb isotope compositions of the studied rock suites (not age-corrected)
and potential sources are shown in Fig. 8. The Sr and Nd
isotopic compositions of the AJ suite display an almost
vertical trend (Fig. 8a) bracketed between a ‘mantle-like’
end-member with lower 143Nd/144Nd and slightly higher
87
Sr/86Sr ratios than the East Pacific Rise mid-ocean ridge
basalts (EPR-MORB; based on data from the PetDB database, Lehnert et al., 2000); and an upper crustal component
that might either be represented by the local basement
(SMO; Albrecht & Goldstein, 2000; Verma & CarrascoNu¤n‹ez, 2003) or by MAT sediments from Site 487 of the
Deep Sea Drilling Project (DSDP 487; LaGatta, 2003;
Fig. 1). The Pb isotope compositions of the AJ suite display
a discernibly positive linear array that also plots between
the compositions of a ‘mantle-like’ component and an
enriched upper crustal end-member (Fig. 8b). Because the
most depleted isotope ratios are observed in the most primitive samples of the AJ suite (as will be shown below), we
will consider them as the closest approximation to the isotopic composition of the mantle wedge.
High-K rocks and the most mafic samples of the AJ suite
exhibit similarly radiogenic Nd isotopic compositions
(143Nd/144Nd 0·5130); however, the former have Sr isotope
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Fig. 5. Major element variation diagrams for the studied rock suites. Abundances of oxides are normalized to 100% volatile-free. (a) Fe2O3tot vs
SiO2; (b) MgO vs SiO2; (c) Al2O3 vs SiO2; (d) Na2O vs SiO2.
compositions that are shifted towards slightly higher
values, as well as lower Pb isotope ratios. The high-K
group also shows smaller overall variations in Sr, Nd and
Pb isotope compositions than the AJ suite.
DISCUSSION
Origin of the anomalous REE abundances
and Ce anomalies in the AJ enriched
samples
The fractionation of Ce from the other REE is inferred to
be a consequence of changes in the valence state of this element from trivalent to tetravalent under highly oxidizing
conditions, whereas the other REE remain trivalent
(Cotten et al.,1995). Negative Ce anomalies have been identified in the volcanic products of many convergent margins
worldwide, and their origin has been often related to the
involvement of a subducted sediment component in the
petrogenesis of these magmas (Elliott et al., 1997; Class
et al., 2000); in particular, the contribution of pelagic sediments that inherit their typical negative Ce anomalies
from seawater (Elderfield & Greaves, 1982; Plank &
Langmuir, 1998) is considered to be responsible for these
peculiar chemical features.
Nevertheless, it has also been documented that negative
Ce anomalies accompanied by Y and REE enrichment in
volcanic rocks might be caused by severe alteration processes (Cotten et al., 1995; Patino et al., 2003). In particular,
in settings where the development of spheroidal weathering is advanced, acidic solutions can leach the trivalent
REE and Y out of the more extensively altered material
(i.e. the external shell) and transport them to the less
altered core, thus producing an unusual enrichment of
these elements in the internal portion of the rock; this
enrichment does not include Ce4þ, as it is not as easily
mobilized as the trivalent REE (Patino et al., 2003).
Interestingly, this peculiar type of weathering has no
major effects on other incompatible trace elements that
are typically mobilized during hydrothermal alteration,
and it does not modify the isotopic composition of the
altered rocks. Moreover, Yand REE enrichments in ‘spheroidally weathered’ rocks do not appear to be systematically
2172
MORI et al.
LITHOSPHERIC REMOVAL AND CRUST FORMATION
Fig. 6. Trace element variation diagrams for the studied rock suites. SiO2 abundances are normalized to 100% volatile-free. (a) Ni vs SiO2; (b)
Rb vs SiO2; (c) La vs SiO2; (d) Y vs SiO2.
correlated with loss of ignition (LOI) values (Cotten et al.,
1995).
In the Altos de Jalisco, mafic rocks in the field are occasionally affected by spheroidal weathering. Comparison of
the REE-enriched lavas from the study area with the
other samples with normal Y and REE abundances supports the idea that element^element variations in the
former are controlled by weathering-induced mobility
rather than magmatic processes; indeed, Y and REE
enrichments (such as La; Fig. 9a) are completely
decoupled from Ce, which is not mobile during weathering
(Cotten et al., 1995; Class & le Roex, 2008), whereas these
elements should form a well-defined positive correlation if
their abundances were governed by partial melting or fractional crystallization (as can be observed for the other
samples of the mafic province), as a result of their similar
partition coefficients. Moreover, the samples from the
Altos de Jalisco region with negative Ce anomalies lack
additional trace element signatures such as high U/Nb
ratios that might reflect important sediment contributions
(Class et al., 2000; Fig. 9b), providing further evidence
that the special features of these rocks reflect weathering.
Based on these considerations, we do not take into
account the Yand REE data of these peculiarly weathered
samples for the discussion on the petrogenesis of the AJ
suite, as they do not represent primary magmatic features.
Evidence for high-pressure fractional
crystallization and crustal contamination
in the Altos de Jalisco
The coherent arrays displayed by samples of the AJ suite in
major and trace element variation diagrams can be interpreted as fractional crystallization trends from parental,
mantle-derived, tholeiitic basalts (Figs 5 and 6).
Interestingly, these rocks also display a negative correlation in a plot of MgO content vs Gd/Yb (and La/Yb)
ratios (Fig. 10a), indicating that the differentiation process
involved the crystallization of mineral phases that were
capable of modifying the REE budget of the residual
liquids (Mu«ntener et al., 2001). A typical low-pressure
assemblage including olivine, plagioclase and oxides could
not be responsible for the formation of these trends
(Fig. 10a and b), as these minerals are not major repository
phases for the REE; at the same time, although amphibole
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JOURNAL OF PETROLOGY
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NOVEMBER 2009
Fig. 7. Normal (N)-MORB normalized trace element patterns of the studied rock suites (normalization values after Sun & McDonough, 1989).
(a) AJ suite. (b) REE-enriched samples of the AJ suite; REE patterns shown in the inset are chondrite-normalized (McDonough & Sun,
1995). (c) High-K group.
crystallization could in principle generate high LREE/
HREE ratios in the residual liquids (Castillo et al., 1999),
it would also produce a decrease in the Gd/Yb ratios,
because this mineral preferentially retains Gd over Yb
(Bottazzi et al., 1999). On the other hand, fractionated
LREE and HREE patterns are reliable indicators of
garnet and clinopyroxene crystallization (MacPherson
et al., 2006). Because garnet crystallizes from hydrous
2174
MORI et al.
LITHOSPHERIC REMOVAL AND CRUST FORMATION
Fig. 8. Isotope variation diagrams for the studied rock suites and potential end-members. (a) Nd vs Sr isotopes; (b) 207Pb/204Pb vs 206Pb/204Pb.
The mantle wedge beneath the western TMVB is assumed to have isotopic compositions similar to those of the most depleted samples of the
AJ suite. Also shown are the data fields of EPR-MORB (Lehnert et al., 2000), mafic portions of the Guerrero Terrane (GT; Lapierre et al.,
1992), SMO (Albrecht & Goldstein, 2000; Verma & Carrasco-Nu¤n‹ez, 2003), and the average isotopic compositions of subducted sediments
from DSDP 487 (LaGatta, 2003).
Fig. 9. Variations of (a) La vs Ce and (b) U/Nb vs Ce anomaly show that the unusual Yand REE enrichments coupled with Ce anomalies in
some samples of the AJ suite could not be produced by sediment melt contributions or magmatic processes such as partial melting (PM) or fractional crystallization (FC). They more probably reflect weathering processes. Data field of pelagic sediments from LaGatta (2003).
Ce anomaly ¼ CeN/[(2/3 LaN) þ (1/3 NdN)].
basalts at pressures greater than 1·2 GPa (Mu«ntener et al.,
2001), we conclude that the geochemical trends described
by the AJ suite in Fig. 10a and b reflect high-pressure fractionation of clinopyroxene and garnet from a hydrous parental tholeiitic basalt. Notably, increasing garnet
signatures in the AJ suite closely correlate with decreasing
Nd isotope ratios, indicating that garnet and clinopyroxene crystallization from mantle-derived magmas occurred
simultaneously with the assimilation of isotopically
enriched materials (Fig. 10c).
To constrain the nature of the crustal component
involved in the petrogenesis of the AJ group, the samples
and possible contaminants have been plotted on a diagram
of 143Nd/144Nd vs Rb/Nd ratios (Fig. 10d). Also shown for
comparison are the data fields of the volcanic products
that were emplaced during Miocene times in the central
TMVB (Quere¤taro area), and that suffered contamination
at different crustal levels (Mori et al., 2007); in particular,
the Middle Miocene adakitic rocks assimilated upper
crustal lithologies (Fig. 10d), whereas the geochemical
variations of the Late Miocene mafic sequences
reflect deep fractional crystallization (Fig. 10a) and contamination with lower crustal materials with low
143
Nd/144Nd, Rb/Nb and Rb/Nd ratios (Mori et al., 2007;
Fig. 10d).
Assimilation of the mafic Mesozoic sequences of the
Guerrero Terrane, characterized by overall high Nd isotope ratios (Lapierre et al., 1992), could not be responsible
for the geochemical variations observed in the AJ suite
(Fig. 10c and d). At the same time, the participation of
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Fig. 10. (a) The negative correlation between MgO (abundances normalized to 100% volatile-free) and Gd/Yb ratio is indicative of highpressure fractional crystallization (high-P FC) of garnet and pyroxene within the AJ suite, and within the Late Miocene mafic sequences of
the central TMVB (i.e. the Quere¤taro Volcanic Succession: QVS; Mori et al., 2007). As observed for EPR-MORB (Lehnert et al., 2000), lowpressure fractional crystallization (low-P FC) would not affect the Gd/Yb ratios of residual liquids. The lack of correlation between MgO and
Gd/Yb in the high-K group indicates that REE variations in this suite were governed by mantle processes rather than fractional crystallization.
(b) Variations of La/Yb and Gd/Yb ratios are useful to identify the mantle sources involved in the generation of the high-K sequence and of
the most primitive magmas of the AJ suite: calculated batch melting models for different mantle sources (see Table 4 for details) indicate that
high-K magmas could be derived from variable extents of melting of a garnet (grt) pyroxenite source, whereas the most primitive samples of
the AJ suite could be derived from 10% melting of a spinel (sp) peridotite source. Variations of Nd isotopes with (c) Gd/Yb and (d) Rb/Nd
ratios show that high-P FC in the AJ suite occurred simultaneously with assimilation of crustal materials. Whereas the mafic rocks of the QVS
experienced contamination with lower crustal lithologies with low Rb/Nd ratios, the AJ suite assimilated a more felsic component with higher
Rb/Nd ratios, similar to the crust that was assimilated by the Middle Miocene adakites of the central TMVB (i.e. the Palo Hue¤rfano^La
Joya^Zamorano Volcanic Complex: PH-LJ-Z; Mori et al., 2007). Higher Gd/Yb and Rb/Nd ratios of the high-K rocks compared with those of
the most primitive samples of the AJ suite at similar Nd isotope compositions indicate higher proportions of residual garnet during mantle melting, and preferential melting of a mineral phase enriched in incompatible elements. Also shown are the data fields of Mexican lower crustal
xenoliths (LC; Schaaf et al., 1994), mafic portions of the Guerrero Terrane (GT; Lapierre et al., 1992), SMO (Albrecht & Goldstein, 2000; Verma
& Carrasco-Nu¤n‹ez, 2003), and the average composition of subducted sediments from DSDP 487 (LaGatta, 2003).
lower crustal rocks with typically low Rb/Nd ratios
(Rudnick & Gao, 2003) can be reasonably ruled out, as
samples of the AJ sequence with the highest Nd isotope
ratios display low Rb/Nd ratios that are almost identical
to those of documented Mexican lower crustal xenoliths
(Schaaf et al., 1994). Contamination with lower continental
crust depleted in Rb would produce a nearly vertical
trend in which lower Nd isotope ratios would be accompanied by almost constant Rb/Nd. In contrast, samples of
the AJ group show a scattered but discernible negative
correlation, extending towards a more evolved endmember with low 143Nd/144Nd and higher Rb/Nd, which
is more similar to the upper crustal component that was
assimilated by the Middle Miocene adakites of the central
TMVB (Fig. 10d). The nature of the contaminant material
is difficult to define, but it could be represented either by
the continental sediments located at the base of the
Guerrero Terrane (Centeno-Garc|¤ a et al., 2008) or by the
tonalitic^granodioritic plutons that form the intrusive
counterparts of subduction-related volcanic activity, and
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MORI et al.
LITHOSPHERIC REMOVAL AND CRUST FORMATION
that have been generated throughout the geological history
of the Mexican convergent margin at least since the Late
Cretaceous (Schaaf et al., 1995).
In any case, the overall geochemical evidence indicates
that crustal assimilation and fractional crystallization in
the AJ suite must have occurred at least at a depth of
40 km (corresponding to 1·2 GPa), because otherwise
garnet would have been an unlikely liquidus phase
(Mu«ntener et al., 2001). Interestingly, this depth corresponds to the present-day crustal thickness of the Altos de
Jalisco region (Urrutia-Fucugauchi & Flores-Ruiz, 1996),
suggesting that the felsic lithologies with high Rb/Nd
ratios might have been in direct contact with the upper
mantle, and thus exposed to basaltic intrusion, at least
since the Late Miocene. This is in contrast to the thicker
and Rb-depleted lower continental crust beneath the central TMVB (45^50 km; Urrutia-Fucugauchi & FloresRuiz, 1996) that contaminated the slightly younger Late
Miocene mafic sequences emplaced in the Quere¤taro area
(Mori et al., 2007).
In summary, the Late Miocene mafic magmas emplaced
in the Altos de Jalisco and in the Quere¤taro areas both
experienced deep fractional crystallization and contamination at the base of the continental crust. Nevertheless,
they assimilated different contaminants, represented by
felsic, Rb-enriched lithologies and mafic, Rb-depleted
materials, respectively. As both regions presumably shared
a common crustal architecture, and were similarly affected
by the large tectono-thermal event that generated the
silicic SMO province, it appears that a considerable
volume of lower crustal mafic^ultramafic lithologies was
missing below western Mexico during the emplacement of
the AJ suite.
Different mantle sources and metasomatic
agents for the AJ suite and the high-K
group
The higher K2O and Ni abundances and lower Al2O3 and
Na2O at similar SiO2 contents (Figs 4b, 5c, d and 6a)
show that the high-K rocks from the study area do not
derive from fractional crystallization of a more mafic
magma belonging to the AJ suite. Also, the lack of correlation between Gd/Yb ratios and MgO contents in the
high-K group indicates that the REE budget of these
rocks was not controlled by deep crystal fractionation
(Fig. 10a). At the same time, the REE variations within
this suite could not have been generated by subduction
components alone, because these elements are largely
insoluble in aqueous fluids (McCulloch & Gamble, 1991;
Pearce & Parkinson, 1993), whereas melt contributions
from DSDP 487 LREE-enriched sediments (LaGatta,
2003) would have induced a marked depletion in Nd isotopic compositions, which is not observed in the high-K
group. These considerations indicate that the REE budget
of the high-K suite should ultimately reflect distinct
mantle sources and processes.
Variations of La/Yb and Gd/Yb in the studied rock
suites might provide insight into the extent of partial melting, as well as into the mineral assemblages of the mantle
sources (Johnson, 1994). Modeled trajectories for spinel
peridotite, garnet peridotite and garnet pyroxenite melting
(see Table 4 for details) are illustrated in Fig. 10b. The modeling results show that the most mafic compositions of the
AJ suite, characterized by the lowest Gd/Yb and La/Yb
ratios, were probably generated by 10% partial melting
of a peridotitic source within the spinel stability field
(Fig. 10b). In contrast, the large ranges in Gd/Yb and La/
Yb displayed by the high-K group are far outside those
predicted by melting of a spinel or garnet peridotite
source, and probably reflect variable degrees of melting
(8^25%) of a pyroxenitic mantle, containing significant
proportions of modal garnet (Fig. 10b). A derivation from
garnet pyroxenite rather than peridotitic lithology for the
high-K suite is further supported by their higher Ni abundances and lower Al2O3 contents relative to those of the
AJ suite at similar MgO (Figs 5c and 6a; Kogiso et al.,
2004).
Diagrams in Fig. 10c and d provide additional evidence
for the involvement of mantle sources with different mineral assemblages in the genesis of the studied sequences.
The high-K group displays similar 143Nd/144Nd ratios to
those of the most primitive rocks of the AJ suite, suggesting
that the mantle wedge beneath the study region had a
rather
homogeneous
Nd
isotopic
composition.
Nevertheless, the generally higher Gd/Yb, Rb/Nd and
Rb/Sr ratios of the high-K rocks (Fig. 10c and d) might
indicate preferential melting of phlogopite as the principal
carrier for highly incompatible elements within the garnet
pyroxenite source (Schmidt et al., 1999). These considerations are in good agreement with a variety of petrological
studies that attribute the origin of high-K magmas to partial melting of an extensively metasomatized pyroxenitic
mantle source containing garnet, apatite and K-rich
hydrous phases such as phlogopite in its mineral assemblage (Lange & Carmichael, 1991; Foley, 1992; Carmichael
et al., 1996; Luhr, 1997; Elkins-Tanton & Grove, 2003;
Schiano et al., 2004). In particular, the pyroxenitic source
of potassic magmas could derive from the interaction
between mantle peridotites and incompatible elementenriched hydrous silicic melts, through metasomatic reactions that dissolve primary mantle minerals (such as
olivine) to form phlogopite and pyroxene (Sekine & Wyllie,
1982; Wyllie & Sekine, 1982; Schneider & Eggler, 1986).
The geochemical differences observed between the studied sequences support the idea that different hydrous components and metasomatic agents were involved in their
petrogenesis. Indeed, the relatively low LILE/HFSE ratios
and the almost flat REE patterns that characterize the
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Table 4: Source compositions and partition coefficients used in the models
Peridotite1
Pyroxenite2
Kdol3
Kdcpx3
Kdopx3
Kdsp3
Kdgrt3
Dsp
La
0·64
2·03
0·00005
0·053
0·0005
0·0006
0·001
0·009
0·010
0·028
Gd
0·92
1·66
0·00099
0·37
0·016
0·0006
0·80
0·068
0·099
0·400
Yb
0·68
1·29
0·017
0·432
0·047
0·0045
4·18
0·096
0·228
1·284
4
per/melt
Dgrt
5
per/melt
Dgrt
6
pyr/melt
1
Peridotitic mantle composition is an average EPR-MORB (Donnelly, 2002) inverted at 10% batch melting of a spinel
peridotite (53% Ol, 17% Cpx, 28% Opx, 2% Sp).
mantle is sample DMP425v from Liu et al. (2005).
Partition coefficients for single minerals from Hart & Dunn (1993), Kelemen et al. (1993), Johnson (1994) and Salters &
Longhi (1999).
4
Bulk solid–melt partition coefficients for a spinel peridotite assuming a residual mantle mineralogy of 53% Ol, 17% Cpx,
28% Opx, 2% Sp.
5
Bulk solid–melt partition coefficients for a garnet peridotite assuming a residual mantle mineralogy of 54% Ol, 19% Cpx,
24% Opx, 3% Grt.
6
Bulk solid–melt partition coefficients for a garnet pyroxenite assuming a residual mantle mineralogy of 53% Cpx, 22%
Opx, 25% Grt.
2
Pyroxenitic
3
most primitive magmas of the AJ suite (Fig. 7a) suggest
that a spinel peridotitic source was fluxed by aqueous
fluids, which produced a moderate enrichment in mobile
elements without modifying the HFSE and REE budgets
of the primitive mantle melts. On the other hand, the
inferred mineralogical features of the mantle source of
high-K magmas, their prominent enrichment in LILE
and other highly incompatible elements, and their fractionated REE patterns indicate that the metasomatic agent
involved in the petrogenesis of this suite was probably a
hydrous silicic melt enriched in K2O and other LILE,
which reacted with mantle peridotites to form phlogopiteand garnet-bearing pyroxenitic lithologies.
If the hydrous components involved in the genesis of
both suites were derived from the same source (whose
nature will be discussed below), then the difference in Sr
isotope ratios between the AJ tholeiites and high-K rocks
at similar 143Nd/144Nd ratios (Fig. 8) may indicate that the
metasomatic agents (aqueous fluids and silicate melts,
respectively) reacted in different proportions with the peridotitic mantle (i.e. different fluid/mantle and melt/mantle
ratios). In particular, the lower Sr isotope ratios of the AJ
tholeiites probably indicate that the aqueous fluids induced
a smaller change to the original isotopic composition of
the peridotitic mantle (i.e. small fluid/mantle ratio),
whereas higher 87Sr/86Sr ratios in the high-K samples suggest a stronger influence of the potassium-rich silicic melts
on the isotopic composition of the pyroxenitic veins (as a
result of a higher metasomatic melt-Sr/mantle-Sr ratio).
Alternatively, the more radiogenic Sr isotope compositions
of the high-K magmas might indicate that their pyroxenitic source already existed beneath the Altos de Jalisco
region, and that it was created during older metasomatic
processes. Because metasomatic, phlogopite-bearing veins
and the high-K suite display higher Rb/Sr ratios than typical peridotitic mantle (Schmidt et al., 1999), at least 130
Myr of isotopic ageing are necessary to shift the Sr isotopic
composition of the pyroxenites to the slightly higher
values observed in high-K magmas, if we assume an initial
87
Sr/86Sr ratio analogous to those of the AJ tholeiites, and
a maximum Rb/Sr ratio of 0·13 (average value for the
high-K group).
A P E T RO G E N E T I C ^ T E C T O N I C
M O D E L F O R T H E A LT O S D E
J A L I S C O VO L C A N I C P RO V I N C E
Previous models
The Late Miocene episode of flood basalt eruptions that
marked the beginning of magmatic activity in the western
TMVB represents an extraordinary event, displaying geological, volcanological and compositional features that are
rarely observed in other continental arcs; indeed, these
lavas clearly differ from the more typical arc-like volcanic
products that were emplaced during the same period in
the central and eastern sectors of the TMVB. In this
sense, it does not seem surprising that the models that
have been proposed to explain this volcanic episode
depart from the usual magmatic scenario of a convergent
margin.
Moore et al. (1994) and Ma¤rquez et al. (1999) related the
Late Miocene mafic province of the western sector of
the arc to the presence of a mantle plume beneath the
Guadalajara region. According to those workers, melting
of an asthenospheric plume might account for the large
volumes of mafic lavas and for their relatively weak subduction signatures. In contrast, Ferrari (2004) interpreted
the mafic volcanic successions emplaced along the arc
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MORI et al.
LITHOSPHERIC REMOVAL AND CRUST FORMATION
between 11 and 3·5 Ma, including the Altos de Jalisco, as
the surface expression of a slab tear that developed at the
mouth of the Gulf of California in the Late Miocene, and
gradually propagated eastward. In this model, the tear in
the slab allowed the ascent of deeper and hotter asthenospheric material, which produced a considerable increase
of temperature at the base of the mantle wedge and
induced its partial melting. The compositional variability
of the mafic successions emplaced along the arc is in turn
attributed to the existence of a heterogeneous mantle with
distinct histories and compositions beneath the TMVB.
The mafic lavas located west of Mexico City, which show
arc-like geochemical signatures, would derive from a
mantle source that has experienced subduction-related
metasomatism and depletion by melt extraction since the
Cretaceous, whereas the intraplate-like alkaline magmas
in the eastern TMVB would derive from a less depleted
mantle, which had not been modified by subduction
agents nor has undergone significant melting for the last
250 Myr (Ferrari, 2004).
The mantle plume hypothesis proposed by Moore et al.
(1994) and Ma¤rquez et al. (1999) for the Altos de Jalisco volcanic province has many analogies with the models that
have been traditionally invoked to explain the origin of
continental flood basalts: upwelling of a hot mantle plume
that favors the formation of a thick mantle melting
column, which in turn provides the conditions for the
generation of enormous volumes of basaltic magmas
(Morgan, 1971; McKenzie & Bickle, 1988). Nevertheless, it
has been documented that decompression melting of asthenosphere within an ascending mantle plume essentially
produces tholeiitic magmas that display geochemical affinities with ocean island basalts (OIB; Farmer, 2003).
These primary compositions are commonly observed in
the volcanic products of many continental flood basalt provinces (Farmer, 2003), but are very different from those of
the Altos de Jalisco district. In particular, the low titanium
contents and relatively high LILE/HFSE ratios of the
most mafic tholeiites of the AJ suite are difficult to explain
by partial melting of an upwelling enriched asthenospheric
mantle alone.
On the other hand, the Altos de Jalisco sequences display
compositional similarities to the volcanic products of
some continental flood basalt provinces such as the
Parana¤, Brazil, which show tholeiitic affinities and arclike geochemical features (i.e. high LILE/HFSE ratios).
These magmas have been interpreted as partial melts of a
hydrous lithospheric source heated conductively by an
upwelling mantle plume (Gallagher & Hawkesworth,
1992; Turner et al., 1996; Peate, 1997). Interestingly, the
slab detachment model proposed by Ferrari (2004)
predicts an analogous scenario, according to which the
Late Miocene mafic province of the western TMVB
would derive from melting of a previously metasomatized
and relatively hydrous mantle wedge induced by a temperature increase at its base; nevertheless, in this case
conductive heating of the wedge would be produced by
the passive ascent of hot asthenosphere through an opening slab window, rather than by a mantle plume
sensu stricto.
When melting of a wet peridotite is induced by a temperature increase from below, the wettest portions of the
mantle should be expected to melt first (at relatively
lower temperatures), and they should also melt to larger
extents than the drier counterparts. Therefore, the melting
degree should be proportional to the water content of the
mantle wedge. This is analogous to the well-documented
mechanism for magma generation at subduction zones, in
which the extent of peridotite melting increases with
increasing supply of slab-derived hydrous components to
the mantle wedge (Stolper & Newman, 1994).
To test if the compositional features of the Altos de
Jalisco sequences are compatible with this melting mechanism, we have plotted Na2O vs Ba/Nb (Fig. 11). Na2O contents are good proxies for the extent of mantle melting in
arc magmas if melting takes place at similar depth intervals, and if there is no evidence of slab melting (Plank &
Langmuir, 1988), whereas the Ba/Nb ratio can be taken as
an indicator of the amount of fluids in the mantle source
region, or as a rough proxy for water content (Cervantes
& Wallace, 2003). The negative trend displayed by samples
of the high-K group is consistent with a derivation of
these magmas from variable extents of melting of a waterrich, metasomatized mantle source. In contrast, the tholeiitic and calc-alkaline rocks of the AJ suite do not display
any clear correlation between water contents and the
degree of melting. These observations support the interpretation that the geochemical variations in this volcanic
sequence were mainly governed by deep fractionation and
crustal contamination of a parental tholeiitic basalt rather
than variable extents of mantle melting.
Derivation of the mafic successions of the Altos de Jalisco
from melting of a hydrous mantle wedge induced by the
ascent of hotter asthenosphere through a slab window
(Ferrari, 2004) appears to be inconsistent with numerical
models (Turner et al., 1996). Indeed, these models predict
that conductive heating of a wet mantle would be an extremely slow process, which would protract the duration of
magmatism for 10^15 Myr at very low eruption rates
(Turner et al., 1996). Moreover, magma production within
the wet mantle wedge would be eventually replaced by
more extensive melting of the upwelling asthenosphere,
resulting in the voluminous eruption of OIB-like magmas,
which would be stratigraphically superimposed on the wet
mantle melts throughout their areal distribution (Turner
et al., 1996). None of these features is observed within the
study area. In fact, the major magmatic outburst in the
region took place in less than 1 Myr, without the presence
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Fig. 11. The lack of negative correlation between Na2O contents
(abundances normalized to 100% volatile-free) and Ba/Nb ratios indicates that the geochemical variations within the AJ suite were not governed by mantle melting processes. In contrast, the high-K group
displays a negative correlation, consistent with a derivation of these
magmas from different extents of melting of a water-rich source.
of intraplate-like volcanic products in the entire succession.
Magmas with intraplate-like geochemical characteristics
were erupted during Pliocene^Quaternary times to the
west, in the vicinity of Guadalajara (Fig. 3; Gilbert et al.,
1985), but they are volumetrically insignificant when compared with the Altos de Jalisco mafic sequences.
A new model: lithospheric removal
beneath the Altos de Jalisco region
Although a mantle plume or a slab detachment event
cannot satisfactorily explain the overall features of the
Late Miocene episode of the TMVB, we agree with previous researchers on the idea that the genesis of this mafic
province was not exclusively related to the normal subduction process (i.e. to partial melting of the mantle wedge
fluxed by a hydrous component derived from the subducting slab), especially if its unusual geological, volcanological
and compositional features are compared with those of
the adakitic belt that was still forming at 11^10 Ma in
the central and eastern TMVB, and whose origin was
clearly related to the continuing subduction process (Mori
et al., 2007).
An extraordinary influx of mantle melts beneath the
western portion of the TMVB seems to be a necessary
requirement to account for the eruption of large volumes
of mafic magmas in such a short period of time.
Therefore, based on the geological, volcanological and
geochemical evidence of the studied rock suites, as well as
on recent numerical simulations (Elkins-Tanton, 2005), we
propose an alternative model that relates the genesis of
the Late Miocene mafic province of the Altos de Jalisco to
a process of lithospheric dripping (involving mantle
NUMBER 11
NOVEMBER 2009
lithosphere and lowermost crustal lithologies) beneath the
region.
Starting from the Late Cretaceous, abundant magmatic
activity related to the subduction of the Farallon oceanic
plate beneath North America contributed to the construction of the SMO, which is the largest silicic volcanic province on the planet. Despite the dispute regarding its origin
by fractional crystallization of mantle-derived magmas
(Cameron & Hanson, 1982; Wark, 1991; Smith et al., 1996)
or crustal anatexis (Ruiz et al., 1988), it seems inevitable
that enormous volumes of underplated basalts should
have contributed either by heat and/or mass transfer to its
formation (see, e.g. Ferrari et al., 2007). The continuous
accretion of mantle-derived magmas at the base of the
crust must have produced a progressive increase in lithospheric thickness, but it should also have generated a
dense lower lithosphere, as a result of crystallization and
accumulation of pyroxene and garnet from the hydrous
basaltic melts at depth (Mu«ntener et al., 2001), and of
high-pressure metamorphic reactions that probably transformed the underplated arc magmas into garnet amphibolites or even eclogites (Atherton & Petford, 1993).
Therefore, a relatively young and thick mafic^ultramafic
lithospheric root closely related to SMO activity should
have existed beneath the western portion of the TMVB at
least until Early Miocene times.
Between 22 and 11 Ma, there was a volcanic hiatus in
western Mexico, although subduction of the Farallon^
Cocos system was still continuing at that time, and was
capable of inducing arc magmatism in the central and
eastern sectors of the TMVB. The reasons for this 10
Myr hiatus are unknown, but they may be related to the
thickened nature of the continental lithosphere at that
time. Indeed, an overthickened crust beneath the region
might have inhibited the ascent and eruption of
subduction-derived magmas, forcing them to stall and
crystallize at progressively greater depths. Given a constant magmatic influx imparted by continuous hydrous
melting of the mantle wedge, the mafic^ultramafic lithospheric root probably thickened with time, until it reached
a critical density that could not be further sustained by
buoyancy over the underlying mantle, and that ultimately
led to the removal of the lower lithosphere through ductile
gravitational instabilitites (i.e. lithospheric dripping;
Elkins-Tanton, 2007). This interpretation is in agreement
with previous studies, which document that lower crustal
lithologies such as mafic^ultramafic cumulates and metamorphosed underplated basalts can produce gravitational
instabilities beneath volcanic arcs (Jull & Kelemen, 2001).
According to these studies, the high Moho temperatures
at subduction zones (47008C) would allow gravitational
instabilities to develop on a timescale of 10 Myr, further
supporting our model (Jull & Kelemen, 2001). Removal
of lower crustal materials via lithospheric dripping is also
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MORI et al.
LITHOSPHERIC REMOVAL AND CRUST FORMATION
consistent with the observation that the continental
crust underlying the study area is thinner than in the
adjacent regions (40 km beneath the Altos de Jalisco vs
45^50 km beneath the central TMVB; UrrutiaFucugauchi & Flores-Ruiz, 1996), although the influence of
the SMO activity was longer and more voluminous.
The compositional features of the Late Miocene mafic
province of the Altos de Jalisco are also entirely consistent
with the predicted manifestations of continental magmatism induced by the removal of lower lithosphere. Indeed,
it has been documented that lithospheric downwelling
may promote abundant magma production, and ultimately generate voluminous eruptions that can even
match the sizes of continental flood basalt provinces
(Elkins-Tanton et al., 2006; Elkins-Tanton, 2007). Largevolume magma production within this scenario is caused
by a combination of decompression and flux melting of
the mantle: voluminous mantle upwelling around the foundering instability results in decompression melting; at the
same time, mantle melting is enhanced by the continuous
injection of hydrous components that may be released
from the underlying subducting plate and/or be derived
from dewatering of the sinking materials (Elkins-Tanton,
2005, 2007).
In conclusion, downwelling of mantle lithosphere and
lower crustal lithologies beneath the Altos de Jalisco at the
end of the Miocene enhanced abundant hydrous and
decompression melting of the mantle wedge, and produced
the most primitive tholeiitic compositions of the AJ suite
(Fig. 12). During ascent, these magmas experienced highpressure pyroxene and garnet fractionation concomitant
with the assimilation of the newly exposed and relatively
more felsic crustal materials (Fig. 12).
The lithospheric dripping model can also account for
the generation of the high-K set of samples distributed
along the eastern boundary of the Altos de Jalisco volcanic
province. As explained above, high-K magmas are probably the product of partial melting of garnet pyroxenitic
lithologies that formed by reaction of mantle wedge peridotites with hydrous silicic melts enriched in K2O and
other highly incompatible elements such as U and Rb. A
slab-derived origin for these metasomatizing magmas is
unlikely, as melting of the subducted tholeiitic MORB
crust during the early evolutionary stages of the TMVB
has been shown to produce trondhjemitic compositions,
characterized by prominent Na2O enrichment over K2O
(Go¤mez-Tuena et al., 2008). On the other hand, it has been
documented that melting of arc-related lithologies such as
underplated amphibolitic metabasalts, typically enriched
in incompatible elements, might produce K2O-rich
hydrous silicic magmas at pressures of 1·5^2 GPa (Sen &
Dunn, 1994; Pe-Piper et al., 2009). When melting takes
place at these pressures, amphibole and plagioclase in the
source will react out to form eclogitic residual mineral
assemblages (Sen & Dunn, 1994), thus producing magmas
with strong LREE enrichment and HREE depletion, but
with no negative Eu anomalies. Based on these considerations, we propose that preferential melting of foundering
amphibolite-facies basaltic materials at the edges of the
sinking instability (which could have occurred at least at
2 GPa), induced by a stronger exposure to the flux of
heat from the upwelling asthenosphere, was responsible
for the generation of the hydrous silicic magmas that contributed to the formation of K-rich pyroxenitic veins
within the mantle wedge (Fig. 12). Even if we cannot disregard the possibility that the source material of the high-K
magmas may have already existed in the form of ‘inherited’ metasomatic veins beneath the TMVB, the peculiar
location of the high-K rocks along the fringes of the Altos
de Jalisco seems more consistent with an origin of their
pyroxenitic source by reaction of mantle peridotites with
magmas derived by preferential anatexis along the borders
of the foundering mass (Fig. 12).
CONC LUSIONS A N D
I M P L I C AT I O N S
The Late Miocene volcanic district of the Altos de Jalisco is
a large (8000 km2) and voluminous (2000 km3) province of mafic plateaux with a minor proportion of high-K
rocks, that erupted in less than 2 Myr in the western
TMVB. The emplacement of the Altos de Jalisco volcanic
successions has been previously related to unusual events
such as the arrival of a mantle plume beneath the region
(Moore et al., 1994; Ma¤rquez et al., 1999) or a slab detachment process (Ferrari, 2004). However, none of these
processes provide a consistent explanation for the geochemical data. In contrast, the new data presented here
support the idea that the Altos de Jalisco mafic province
might represent the surface manifestation of a lithospheric
dripping event, which occurred beneath the western
TMVB as a consequence of a long-lasting period of magmatic thickening and densification of the lower lithosphere
in a continental arc setting. Within this context, the release
of fluids from the dehydrating foundering materials,
coupled with mantle upwelling around the sinking mass,
induced abundant flux and decompression melting of the
mantle, thus producing large volumes of tholeiitic magma.
These melts subsequently interacted with a newly exposed
and relatively more felsic crust, and evolved to more differentiated calc-alkaline compositions through high-pressure
crystallization and crustal contamination. On the other
hand, preferential melting at the borders of the foundering
mafic lithologies, induced by a stronger exposure to hotter
upwelling mantle, generated hydrous silicic magmas that
reacted with mantle peridotites to form the garnet- and
phlogopite-bearing pyroxenitic source of the high-K rocks.
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NUMBER 11
NOVEMBER 2009
Fig. 12. Schematic diagram showing the proposed petrogenetic^tectonic model for the Late Miocene Altos de Jalisco mafic province.
Lithospheric foundering beneath the study region enhanced magma generation by a combination of hydrous and decompression melting (dehydration of the sinking materials and mantle upwelling around the foundering instability), thus producing the most primitive tholeiitic compositions of the AJ suite. During their ascent, these magmas experienced high-pressure fractional crystallization concomitant with the
assimilation of the newly exposed felsic crustal materials, generating the more differentiated products of the AJ series. Melting at the edges of
the foundering lower crustal lithologies, induced by a stronger exposure to hotter upwelling asthenosphere, produced hydrous silicic magmas
that reacted with mantle peridotites to form K-rich pyroxenitic veins. Preferential melting (and subsequent pyroxenite formation) along the borders of the sinking materials was responsible for the distribution of the high-K group along the boundaries of the Altos de Jalisco region. The
representation of lithospheric dripping is taken from Elkins-Tanton (2007).
Our interpretation of the Altos de Jalisco mafic district is
in agreement with numerical modeling results (ElkinsTanton, 2005) that predict the production of large volumes
of magma as a result of foundering of lower lithosphere.
Although direct or indirect contributions of upwelling
mantle plumes have usually been invoked for the formation of large igneous provinces (Farmer, 2003; but see, e.g.
Sheth, 2005, for a contrasting point of view), the model
provided here offers an alternative explanation for the generation of smaller provinces of continental flood basalts in
volcanic arcs or intraplate settings.
The geochemical features of the Altos de Jalisco volcanic
province of the TMVB also support the idea that the
removal of lower lithosphere can be a powerful means for
element recycling on a global scale (Aeolus Lee et al.,
2006; Elkins-Tanton, 2007). Indeed, we have shown that
dehydration and melting of the foundering instability not
only can impart ‘arc-like’ trace element signatures to the
newly formed volcanic products, but may also provide a
counterbalance to the loss of mafic lower crust by triggering abundant basaltic flooding at continents. Although it
is generally considered that the removal of lower lithosphere exerts a strong influence on driving the bulk crust
towards intermediate compositions, our petrogenetic
model suggests that additional mechanisms are required
for stabilizing andesitic continents on Earth.
AC K N O W L E D G E M E N T S
L.M. thanks D. J. Mora¤n-Zenteno and F. Ortega-Gutie¤rrez
for fruitful discussions on the petrogenesis of the Altos de
Jalisco volcanic province. We thank J. T. Va¤squez-Ram|¤ rez
(CGEO) for preparing the petrographic thin sections,
and M. Albara¤n-Murillo (CGEO) for sample crushing
and powdering. Invaluable help was provided by R.
Lozano-Santa Cruz (LUGIS) during major element
determinations. Our sincere thanks go to J. J. MoralesContreras, M. S. Herna¤ndez-Bernal and T. Herna¤ndezTrevin‹o (LUGIS) for help during isotopic analyses.
2182
MORI et al.
LITHOSPHERIC REMOVAL AND CRUST FORMATION
Constructive reviews by L. J. Elkins-Tanton, P. M. Holm,
P. T. Leat and H. C. Sheth led to significant improvements
in the manuscript. Editorial handling by Professor J.
Gamble is also highly appreciated.
FUNDING
This work was supported by Consejo Nacional de Ciencia
y Tecnolog|¤ a [39785 grant to A.G.-T.]; Programa de Apoyo
a Proyectos de Investigacio¤n e Innovacio¤n Tecnolo¤gicaUniversidad Nacional Auto¤noma de Me¤xico [IN103907
grant to A.G.-T.]; National Science Foundation [EAR 9614782 grant to S.L.G.]; and Consejo Te¤nico de la
Investigacio¤n Cient|¤ fica-Universidad Nacional Auto¤noma
de Me¤xico [postdoctoral fellowship to L.M.].
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