Article
Volume 13, Number 1
1 June 2012
Q0AM04, doi:10.1029/2011GC003974
ISSN: 1525-2027
Large volumes of rejuvenated volcanism in Samoa: Evidence
supporting a tectonic influence on late-stage volcanism
Jasper G. Konter
Department of Geological Sciences, University of Texas at El Paso, El Paso, Texas 79968, USA
([email protected])
Matthew G. Jackson
Department of Earth Sciences, Boston University, Boston, Massachusetts 02215, USA
[1] Hot spot volcanoes are commonly constructed in a characteristic sequence of stages. After the volumet-
rically dominant shield stage, a protracted period of quiescence ends with a final stage of activity: rejuvenated volcanism. The mechanism responsible for generating rejuvenated volcanism is not generally
agreed upon. New data obtained for samples 200 m down-section in a deeply incised canyon on Savai‘i
(Samoa) are unusually enriched isotopically and indicate a relatively voluminous rejuvenated stage compared to other intraplate volcanoes. Using a modified flexural model originally proposed for Hawai‘i, we
suggest that the location of Samoa near the Tonga Trench terminus causes plate flexure resulting in upward
flow of the shallow mantle driving partial melting. In particular, subduction-related plate bending in the
Samoan region may cause a larger flexural amplitude than generated by volcanic loading in Hawai‘i. The
larger amplitude may explain the larger volume of rejuvenated melt in Samoa, constrained by our new data.
Moreover, we argue that Sr-Nd-Pb-Os-He-Ne isotopes in Samoan rejuvenated lavas are all consistent with
sampling of a lithospheric component that is characterized by a metasomatic imprint from the Pacific Plate’s
earlier passage over the Rarotonga hot spot. Furthermore, temperature estimates for the melts suggest a drop
in temperature during the predicted shallower melting due to flexural uplift, compared to the conditions during shield volcanism. Thus, flexural bending and metasomatism of the Samoan lithosphere may have generated the voluminous and geochemically distinct Samoan rejuvenated lavas, implying the lithosphere may
play an important role during this stage in non- Hawaiian hot spots.
Components: 14,800 words, 10 figures.
Keywords: OIB; Samoa; hot spot; plate flexure; radiogenic isotopes; rejuvenated volcanism.
Index Terms: 1033 Geochemistry: Intra-plate processes (3615, 8415); 1040 Geochemistry: Radiogenic isotope
geochemistry; 8178 Tectonophysics: Tectonics and magmatism.
Received 22 November 2011; Revised 13 March 2012; Accepted 25 April 2012; Published 1 June 2012.
Konter, J. G., and M. G. Jackson (2012), Large volumes of rejuvenated volcanism in Samoa: Evidence supporting a tectonic
influence on late-stage volcanism, Geochem. Geophys. Geosyst., 13, Q0AM04, doi:10.1029/2011GC003974.
Theme: Studies of Seamount Trails: Implications for Geodynamic Mantle Flow Models
and the Geochemical Evolution of Primary Hot Spots
This paper is not subject to U.S. copyright.
Published in 2012 by the American Geophysical Union.
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1. Introduction
[2] Following the volumetrically dominant shield
stage volcanism, a period of quiescence of at least
0.5 Ma ends with eruptions of the rejuvenated
stage, the final stage of volcanism at hot spots [e.g.,
Clague and Dalrymple, 1987; Garcia et al., 2010].
In Hawai‘i, the archetype for hot spot volcanoes, the rejuvenated stage is erupted after a
0.5–2 million year long hiatus in volcanic activity,
and it is geochemically distinct from earlier, shieldstage volcanism [e.g., Garcia et al., 2010]. Temporal geochemical variations are clearly observed
in samples from vertical sections such as canyons
and drill cores [e.g., Stolper et al., 1996; DePaolo
et al., 2001a; Garcia et al., 2007, 2010], and have
helped constrain the comparatively small total
volumes for the rejuvenated volcanic stage. Despite
their small volumes, rejuvenated lavas have a surprisingly large impact on our understanding of the
origin of volcanism at hot spots.
[3] Hot spots are thought to be generated by
buoyantly upwelling mantle plumes that melt
beneath a moving plate, producing age-progressive
chains of volcanoes that extend away from a central
region of melting, or ‘hot spot’ [Morgan, 1971].
In Hawai‘i, volcanism consists of multiple stages
[e.g., Clague and Dalrymple, 1987] where the first
three stages of each volcano—pre-shield, shield,
and post-shield—are thought to be caused by the
volcano passing over and sampling different parts
of a mantle plume [e.g., Hauri, 1996; Kurz et al.,
1996; Lassiter et al., 1996; Frey and Rhodes,
1993; DePaolo et al., 2001b; Blichert-Toft et al.,
2003; Abouchami et al., 2005; Fekiacova et al.,
2007]. However, the fourth and final stage of
volcanism—rejuvenated volcanism—has always
posed a problem for the mantle plume origin of hot
spots, due to the preceding hiatus in activity and
the distinct composition of the lavas. A hiatus
between shield and rejuvenated volcanic stages
implies that the volcano is transported 35–140 km
(Pacific Plate motion of 70 km/Ma) away from
the location where the last shield-building eruptions occurred. Rejuvenated volcanism is difficult
to reconcile with a plume-origin for this final stage
of hot spot activity: the plume is far removed from
the locus of volcanism when rejuvenated volcanism
initiates. Therefore, other (non-plume) mechanisms, including plate flexure, have been proposed
to explain the origin of rejuvenated lavas [e.g., ten
Brink and Brocher, 1987; Bianco et al., 2005;
Konter et al., 2009]. Most models for rejuvenated
volcanism have been based on observations of
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Hawaiian volcanoes, but other ocean islands,
including Samoa, exhibit rejuvenated volcanism
and may provide insight into potential causes.
[4] Rejuvenated volcanism is a feature of volcanic evolution at a number of intraplate volcanic
chains, including Kerguelen, Canary Islands, Madeira,
Mauritius, Marquesas, Society Islands, and FieberlingGuadalupe Seamounts, but the Samoan hot spot
hosts some of the most extensive rejuvenated volcanism in the world (Figures 1 and 2) [Macdonald
et al., 1983; Wright and White, 1987; Woodhead,
1992; White and Duncan, 1996; Ielsch et al.,
1998; Paul et al., 2005; Workman et al., 2004;
Konter et al., 2009; Garcia et al., 2010; White,
2010]. The largest island in Samoa, Savai‘i (similar in size and age to the Hawaiian islands of
O‘ahu and Kaua‘i; Figure 2) is known for its substantial volume of rejuvenated (or post-erosional)
volcanism [Stearns, 1944; Kear and Wood, 1959;
Natland, 1980]. In Samoa, this stage is so extensive
that it has almost completely “resurfaced” the island
with a veneer of young lavas [Workman et al.,
2004]. The resurfacing is so complete that subearially exposed shield stage lavas were only
recently reported in the deeply incised canyon of the
Vanu river [Konter et al., 2010]. Hawkins and
Natland [1975] pointed out that the eruption of the
youngest stage stretches to the islands of Upolu
and Tutuila, but they are not as far along in their
volcanic construction, so the island of Savai‘i provides the best estimate of erupted volumes of rejuvenated volcanism. Moreover, the vertical exposure
afforded by the Vanu river canyon (Figure 3) provides a unique opportunity to estimate the volume
of erupted rejuvenated lavas in Samoa. The geochemical compositions of several lavas sampled
over the bottom 50 m of the canyon walls help
constrain the thickness of the rejuvenated lava
veneer on Savai‘i, since shield-stage lavas are geochemically distinct and should occur below the
lowest rejuvenated lavas. The thickness of this
veneer of rejuvenated volcanism in Samoa is
important for evaluating the proposed models for
the origin of this stage.
[5] The canyon of the Vanu river on Savai‘i is a
steep amphitheater-shaped canyon with a vertical
relief of 500 m. This canyon resembles the large
erosional canyons of Kohala Volcano (Hawai‘i),
which started forming potentially due to waterfall
erosion during the shield stage as evidenced by late
shield stage lavas flowing into the canyon [Lamb
et al., 2007]. In Samoa, the canyon west of the
Vanu river seems to have experienced significant
refilling by younger lavas based on the canyon
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Figure 1. Samoan Islands and seamounts, the northern Tonga Trench and Lau Basin. Vailulu‘u is the currently
active end of the volcanic chain, although rejuvenated volcanism has caused historic eruptions on Savai‘i. Insets on
the right show areas of shield (gray) and rejuvenated (white) lavas on islands. Inset on the left shows simplified
Cook-Austral hot spot tracks (blue [Wessel and Kroenke, 2008]), the Samoan track (red), and hot spot locations
(black stars).
morphology, but the eastern canyon is more steeply
incised and has alluvium exposed on the canyon
floor, resting on minor younger lava fill [Kear and
Wood, 1959]. Such canyons afford direct access to
the vertical stratigraphy of different volcanic stages.
One major difference with the canyons of Kohala is
that the mouth of the Samoan canyons lies near
400 m elevation, while those in Hawai‘i are at sea
level and the canyon walls form seacliffs. In the
most recent eruption, a major canyon was filled
with lava backing up against the offshore reef
[Anderson, 1910]. Given that the erosional profile
of the Vanu river flattens near 400 m, its canyon
may have been filled in a similar fashion. Since this
Figure 2. Mapped extent of rejuvenated (white) and shield (gray) lavas. (left) Savai‘i (Samoa) only exposes shield
stage lavas in the walls of the deeply incised canyon of the Vanu river. Vertical exposures of rejuvenated lavas series
are indicated with height. (middle) Kaua‘i is the Hawaiian island with the largest volume of rejuvenated lavas, covering nearly half the surface area (modified from Garcia et al. [2010]). (right) Erupted volumes on O‘ahu (modified from
Langenheim and Clague [1987]) are even smaller than on Kaua‘i.
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Figure 3. Canyon of the Vanu river, exposing 500 m of lava sequence. (a) Overview of the location of the canyon
with regard to the entire volcano (Shuttle Radar Topography Mission 90 m data; multibeam bathymetry collected during AVON and ALIA cruises reported on by, e.g., Workman et al. [2004] and Jackson et al. [2007]; gravity-derived
bathymetry from Smith and Sandwell [1997]). (b) Satellite imagery (Quickbird) draped over 90 m Shuttle Radar
Topography Mission data. (c) Picture of Sinaloa Falls (200 m) within the canyon, note people for scale.
is the deepest canyon on Savai‘i, combined with the
high eruption rate observed in Samoa [Anderson,
1910], it appears qualitatively that larger volumes
of lava are erupted in Samoa compared to Hawaii,
where major canyons erode down to sea level.
Despite a potential filling of the lower parts of the
Vanu river canyon, the 500 m nearly vertical
upper canyon wall provides constraints on the
thickness of the rejuvenated lavas on Savai‘i. We
show that this exposure suggests a much thicker
sequence of rejuvenated lavas that covers a much
larger portion (99%) of the volcano surface area
compared to the very small volumes of rejuvenated
volcanism that cover less than half of the surface of
a Hawaiian volcano [Garcia et al., 2010].
[6] We present new major element and Pb isotope
data that demonstrates a minimum thickness for the
rejuvenated lavas of 200 m. An upper bound on
this thickness—approximately 500 m, which is the
depth of the canyon—is provided by the presence
of shield-stage trachytes as cobbles in the Vanu
river that drains this canyon [Workman et al.,
2004]. We calculate a minimum thickness of
200 m, which we argue constitutes a volume of
erupted rejuvenated lava that is many times that
observed for rejuvenated volcanism in Hawai‘i
[Garcia et al., 2010]. The unusually large volumes
of rejuvenated lavas in Samoa greatly extend
the available set of observations for this volcanic
stage among hot spots globally. We use both
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geochemical compositions and melting temperature
estimates to critically evaluate proposed models
for the origin of rejuvenated volcanism in Samoa.
[7] The exceptional volumes and unusual geochemical compositions of Samoan rejuvenated volcanism
can help constrain its origins. We argue that the large
volumes of Samoan rejuvenated lavas are the result
of tectonic bending associated with Samoa’s location
near the Tonga trench. The tectonic regime near a
trench corner provides a mechanism for extensive
lithospheric flexure, and we use shallow upwelling
rates from modeling to calculate related partial
melting. We will show that the resulting melts may
have inherited a significant contribution from metasomatized lithosphere beneath Samoa that we argue
was enriched during the plate’s prior passage over
the Rarotonga hot spot (part of the Cook-Austral hot
spot chain) 15 million years ago. Therefore, the
model we present below suggests that the unique
tectonic regime permits melting of metasomatized
lithosphere and provides a common mechanism
linking the tectonic setting with the unusually large
volumes of isotopically enriched rejuvenated lavas in
Samoa.
2. Background on the Samoan Hot Spot
2.1. Samoa as an Age Progressive
Volcanic Chain
[8] The Samoan volcanoes are located in a linear
chain 100 km northeast of the northern terminus
of the Tonga trench (Figure 1). While the eastern
seamount, Vailulu‘u, is considered the youngest
volcano [Hart et al., 2000; Konter et al., 2004;
Staudigel et al., 2006; Sims et al., 2008], the
western-most subaerial volcano, Savai‘i, has been
active since at least 5.29 Ma and was last active in
1905–1911 [e.g., Sapper, 1906; Anderson, 1910;
Natland and Turner, 1985; Koppers et al., 2008].
However, construction of the volcanic shields in
Samoa follows an age progression that is anchored
at the easternmost volcano, Vailulu‘u Seamount, and
continues 1750 km to west to Alexa Bank, which
has lavas dated at 23–24 Ma [e.g., Duncan, 1985;
Natland and Turner, 1985; Hart et al., 2004;
Koppers et al., 2008, 2011; Németh and Cronin,
2009; McDougall, 2010]. On the eastern young
(0–2 Ma) side of the chain, the islands and seamounts east of Tutuila Island form volcanic edifices
that constitute various parts of the shield-building
stage of Samoan volcanism. Moving west, the
islands of Tutuila, Upolu and Savai‘i each host
rejuvenated stage volcanism, while volcanoes even
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further to the west have only been sampled on their
lower submarine flanks such that no extensive set of
rejuvenated lavas are available. The subaerial rejuvenated lava cover increases from less than one
quarter cover on Tutuila, to nearly half cover on
Upolu, and a near-complete “resurfacing” of the
island of Savai‘i with rejuvenated lavas (Figures 1
and 2). Moreover, on Savai‘i the only shield exposure is present in the canyon of the Vanu river
[Konter et al., 2010], constraining the total erupted
volume of rejuvenated lavas.
2.2. Volcanic Stages in Samoa
[9] Volcanism in Samoa spans at least two phases of
activity: a shield stage that makes up the majority
of the volume of a Samoan volcano, which is followed by a period of quiescence that is interrupted by rejuvenated volcanism. Recent efforts
combining geochemical and geochronological data
[Workman et al., 2004; Jackson et al., 2007;
Koppers et al., 2008; Jackson et al., 2010a; Koppers
et al., 2011] have helped refine Samoan eruptive
history [Stearns, 1944; Kear and Wood, 1959;
Hawkins and Natland, 1975; Natland, 1980;
Natland and Turner, 1985]. Such an evolution
ending with a rejuvenated stage is a characteristic
aspect of many intraplate volcanoes, although
Samoa is unique in its occurrence of these stages
near a trench-transform intersection [Macdonald
et al., 1983; Wright and White, 1987; Woodhead,
1992; White and Duncan, 1996; Ielsch et al., 1998;
Paul et al., 2005; Workman et al., 2004; Konter
et al., 2009; Garcia et al., 2010; White, 2010].
Although shield and rejuvenated lavas are also
observed in the Hawaiian archetype volcanoes [e.g.,
Clague and Dalrymple, 1987], it is not clear if all
Hawaiian stages exist in Samoa, particularly the preshield and post-shield stages [Natland and Turner,
1985]. The shield and rejuvenated stages in Samoa
show a number of similarities and differences with
Hawaiian stages. In Savai‘i, the shield stage lasted
from 5.29 Ma [Koppers et al., 2008] until the
appearance of trachytes [Natland and Turner, 1985].
This occurred around 2 Ma based on an imprecise
total fusion age for a cobble originating in the deeply
incised canyon of the Vanu river [Workman et al.,
2004], and therefore the first appearance is not
well-constrained. Nonetheless, Samoan rejuvenated
volcanism followed the shield stage and began in
earnest by 386 Ka, continuing with substantial activity in the last 2000 years, and erupting as recently
as 1905–1911 [Workman et al., 2004; Németh
and Cronin, 2009]. The shield stage in Samoa consists of alkalic basalts through trachytes and
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Figure 4. Total alkali versus SiO2 diagram illustrating basanitic to picritic nature of samples from the canyon. Shield
lavas (blue) are overlap but are generally lower in total alkalis compared to rejuvenated lavas (yellow, orange). Data
sources: Workman et al. [2004], Jackson et al. [2007, 2009], and this work. Sample U-10 S of Workman et al. [2004]
may have been misidentified as rejuvenated instead of shield stage in the field (M. Regelous, personal communication,
2012), and it was therefore omitted from this figure.
phonolites [Natland, 1980]. The young, rejuvenated
lavas consist of alkalic basalt, basanites and nephelinites, with a larger fraction of samples with generally
lower SiO2 compared to shield stage lavas (Figure 4),
and in contrast with Hawai‘i, melilite is absent in
Samoan rejuvenated lavas [Natland and Turner,
1985]. Unlike Hawaiian shield stage lavas, which
are dominantly tholeiitic, nearly all Samoan lavas,
including shield lavas, are alkalic in nature, and tholeiites are rare in Samoa [e.g., Macdonald, 1968;
Natland, 1980; Workman et al., 2004; Hart et al.,
2004; Jackson et al., 2010a]. However, it is common for the shield-stage in oceanic hot spots to be
dominated by alkalic volcanism [e.g., Hoernle and
Schmincke, 1993; Maaløe, 1998; Geldmacher et al.,
2001; Paul et al., 2005; Konter et al., 2009].
[10] The isotopic and trace element composition
of Samoan shield and rejuvenated lavas are extreme
in the context of global oceanic volcanic rocks:
The isotopic composition of Samoan lavas is defined
by highly radiogenic Sr isotope ratios, relatively
unradiogenic Nd isotopes, and intermediate Pb
isotope compositions [Hofmann and White, 1982;
Natland and Turner, 1985; Wright and White,
1987; Workman et al., 2004; Jackson et al., 2007,
2010a]. The Sr isotope ratios of several submarine
lavas from Savai‘i are the highest documented for
oceanic intraplate volcanism, and these data therefore
define one extreme end-member composition (EM2)
[Zindler and Hart, 1986] in the total compositional
range of oceanic volcanic lavas [Jackson et al.,
2007]. By contrast, the rejuvenated lavas are distinct and trend away from the field defined by
Samoan shield lavas toward a component with
moderately elevated 208Pb/204Pb and relatively low
206
Pb/204Pb, moderately high 87Sr/86Sr and low
143
Nd/144Nd [Natland and Turner, 1985; Wright and
White, 1987; Workman et al., 2004]. Moreover, these
lavas also show high Ba/(Sm, Nb, Th) ratios compared to shield stage lavas (Figure 5) [Hart et al.,
2004; Workman et al., 2004; Jackson et al., 2010a].
Such compositions are different from rejuvenated
lavas in Hawai‘i, where the isotopic composition
trend toward the compositions of mid-ocean ridge
basalts, which exhibit low 87Sr/86Sr and 206Pb/204Pb
isotope ratios and high 143Nd/144Nd. While Hawaiian
rejuvenated lavas do have similar high Ba/(Sm, Nb,
Th) [Garcia et al., 2010], the distinct isotopic compositions of the rejuvenated lavas in both island
chains imply that different source materials are
involved in the generation of the melts. We will
examine these compositions in terms of the potential
processes and sources that may were involved in
generating the observed geochemical signatures in
Samoan rejuvenated lavas.
2.3. Rejuvenated Volcanism in Samoa
[11] Although the origin of the shield stage and its
compositional signature is typically assumed to be
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Figure 5. Trace element compositions (Ba/Th versus Zr/Hf) of Samoan lavas show shield and rejuvenated lavas are
distinct. Large overlap exists between the compositions of the rejuvenated lavas and Rarotonga compositions. Data
sources: Workman et al. [2004], Jackson et al. [2007, 2009], and GEOROC database. Sample U-10 S of Workman
et al. [2004] was omitted (see Figure 4).
generated by melting of plume components [e.g.,
Stille et al., 1983; Staudigel et al., 1984; Clague and
Dalrymple, 1987; Frey et al., 1990; Roden et al.,
1994b; Valbracht et al., 1996; Mukhopadhyay
et al., 2003; Frey et al., 2005; Abouchami et al.,
2005; Fekiacova et al., 2007; Tanaka et al., 2008;
Xu et al., 2007; Konter et al., 2009; Ren et al., 2009],
a variety of models exist for the generation of rejuvenated lavas. These models focus on processes that
generate the rejuvenated melt and the nature of the
mantle source material. In the case of Samoa, a
relationship has been proposed between the unusual
volumes of rejuvenated volcanism in Savai‘i, and the
island’s location closer to the northern terminus of
the Tonga trench than any other Samoan island.
Stresses associated with the bending and tearing of
the Pacific plate at the northern terminus of the Tonga
trench may cause a propagating crack, which may in
turn increase the volume of rejuvenated volcanism at
Savai‘i (inset in Figure 6) [Hawkins and Natland,
1975; Natland, 1980; Natland and Turner, 1985;
Hart et al., 2004].
[12] We will demonstrate a further role of the litho-
sphere in modifying volcanism at Savai‘i is
expressed in the distinct isotope geochemistry of its
rejuvenated lavas: Instead of trending toward a
depleted mantle component like Hawaiian rejuvenated lavas, Samoan rejuvenated lavas sample
an enriched EM1-like component that forms an
isotopically distinct group from the EM2
component found in Samoan shield lavas (EM1,
EM2: enriched mantle 1 and 2 [Zindler and Hart,
1986]). Below we argue that flexural forces from
the nearby trench cause shallow melting of an EM1
rich component that may be related to the Pacific
Plate’s passage over the EM1 hot spot of Rarotonga
(Figure 1) [e.g., Chauvel et al., 1997; Konter et al.,
2008; Jackson et al., 2010a].
3. Samples and Methods
3.1. Sample Description
[13] Samples were collected during a field cam-
paign in 2009 on the island of Savai‘i that focused
on sampling the shield stage of Savai‘i and the
oldest rejuvenated lavas, exposed in the canyon of
the Vanu river. This canyon was mapped as shield
stage volcanics (Fagaloa Series) by Kear and Wood
[1959], but the new data show that the amount of
exposed shield lavas was overestimated [Konter
et al., 2010]. Rejuvenated lavas that we recovered
are used here to examine the minimum total erupted
volumes of rejuvenated lavas at Savai‘i. Collection
sites are located within the upper canyon of the
Vanu river, encompassing the lower 50 m of a
200 m vertical waterfall section (Sinaloa Falls). The
samples we collected from the waterfall section are
all olivine-phyric basanites and picrites. Groundmass phases consist of olivine, plagioclase, spinel
and clinopyroxene. Olivine phenocrysts occur both
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Figure 6. Lithospheric deformation around the corner of the Tonga Trench from a 3D numerical model (modified
from Govers and Wortel [2005]). Color shows vertical velocity of the plate, with upward flow under Tutuila, Upolu
and Savai‘i. Blue and red symbols show strain rates are only high very close to the corner, Samoan bathymetry
was overlain (to scale) for comparison. Inset shows model of Natland [1980], with cracks extending further west than
predicted by strain in the 3D model.
as anhedral and euhedral crystals, and the overall
Mg number of the olivines spans a wide range that
indicates that only some of the olivine phenocrysts
are in equilibrium with the Mg number of their
whole rock composition (see Text S1 and Figure S1
in the auxiliary material).1 Rare small phenocrysts
of clinopyroxone or plagioclase are anhedral. These
observations are in agreement with the picritic
major element composition of the samples, suggesting accumulation of olivine and clinopyroxene.
The geochemical composition of these samples
suggests that they are rejuvenated-stage lavas.
1
Auxiliary materials are available in the HTML. doi:10.1029/
2011GC003974.
3.2. Analytical Techniques
[14] All samples were crushed avoiding contact
with any metal, and the freshest chips were picked
with a binocular microscope. Major element analyses were performed on unleached whole-rock
powders by X-ray fluorescence (XRF) at the
GeoAnalytical Lab at Washington State University.
The technique employed is described by Johnson
et al. [1999]. Isotopic analysis of Pb was performed at the University of Texas at El Paso, using
standard HBr-HCl-based chemistry [Hanan and
Schilling, 1989] to separate and purify Pb in a
class 100 environment. Mass spectrometric analysis
was carried out on the Nu Plasma high resolution
multicollector inductively coupled plasma mass
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spectrometer (HR MC-ICP-MS) at the University
of Texas at El Paso. To monitor machine performance on the MC-ICP-MS, standards were run
systematically after every two samples. The Pb
isotopic analyses were carried out using a combination of bracketing samples with SRM 981 Pb
(correcting to values of Todt et al. [1996]) and
monitoring by Tl doping (SRM 997 Tl and
205
Tl/203Tl = 2.3889 [White et al., 2000; Hanan et al.,
2004, 2008]) to correct for mass fractionation and
machine bias [White et al., 2000]. The standard SRM
981 Pb (20 ppb solution, N = 70) achieves a longterm (October 2010–August 2011) average of
206
Pb/204Pb = 16.943 3, 207Pb/204Pb = 15.500 3,
and 208Pb/204Pb = 36.726 8 (2s). The reported
uncertainties for each sample represent in-run precision. Total procedural blanks were less then 50 pg
for Pb. Rock standard BCR2 is well within error of
the average reported by Weiss et al. [2006], where
the average renormalized to Todt et al. [1996] SRM
981 values corresponds to: 206Pb/204Pb = 18.7451 11, 207Pb/204Pb = 15.6192 9, and 208Pb/204Pb =
38.6987 24. All compositional data are presented
in Table S1 in the auxiliary material.
4. Results
[15] The major element compositions, presence of
anhedral phenocrysts and occurrence of groundmass
olivine indicate that the samples we report on here
are alkalic lavas that likely accumulated some olivine and clinopyroxene. The difference in major
element compositions between samples is small
(e.g., Figure 4), but they have different modal contents in phenocrysts and differences in phenocryst
grain sizes. The samples have Pb isotope compositions ranging from 206Pb/204Pb = 18.801–18.808,
207
Pb/204Pb = 15.591–15.599, and 208Pb/204Pb =
38.981–39.003 (Figure 7). This narrow range makes
it difficult to distinguish these samples from each
other based on their Pb isotope compositions.
Compared to other rejuvenated lavas in Samoa,
these samples are more picritic in composition,
showing the lowest SiO2 and total alkali content
found in rejuvenated lavas from Savai‘i, although
rejuvenated lavas with lower SiO2 are found on
Upolu (Figure 4). The isotopic compositions fall
within the characteristic range of rejuvenated lavas
on Savai‘i, and are distinct from Samoan shield
lavas (Figure 7). Therefore, our samples from the
Vanu river canyon enhance aereal cover of the
rejuvenated lavas (the only available data in Upolu
and Tutuila) with an estimate of minimum thickness. On Savai‘i the thickness is 200 m (the height
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of the waterfall), and below we use these constraints
with the geochemical signatures of these lavas to
explore models for the origin of rejuvenated
volcanism.
5. Discussion
5.1. Volume Constraints on Samoan
Rejuvenated Lavas
[16] The stages of volcanism on Savai‘i were
crudely mapped by Kear and Wood [1959] based
on erosional profiles, soil thickness and rock
descriptions. However, geochemical compositions
are needed to distinguish different stage lavas.
Therefore the volume of the rejuvenated lavas on
the mature volcano of Savai‘i has not been previously constrained. Owing to the presence of shieldstage cobbles in the Vanu river [Konter et al.,
2010], we infer that the boundary between shield
stage and rejuvenated lavas is exposed in the deeply
incised canyon that is drained by the river. This
information can be used to place constraints on the
thickness and total volume of the rejuvenated lavas.
We identified rejuvenated lavas 200 m downsection in the wall of a >500 m deep canyon on
Savai‘i, based on their rock types (picrite/basanite)
and Pb isotope compositions compared to shield
lavas (Table S1 in the auxiliary material and
Figures 2, 4, and 7). The samples were collected
from the lower 50 m of the 200 m exposure;
although we do not have a detailed sample record
of the entire 200 m section, it is unlikely that a
change occurred up-section (i.e., back to the shield
stage). Therefore, the discovery of rejuvenated
lavas at the base of the 200 m stratigraphy of the
waterfall indicates that 200 m is the minimum
thickness for the rejuvenated lavas. Additionally,
the presence of a shield-stage cobble originating in
the canyon indicates that the 500 m depth of the
canyon represents a maximum thickness for the
rejuvenated stage. Unfortunately, no systematically
collected well cores [Garcia et al., 2010] or scientific drill cores [Stolper et al., 1996; DePaolo et al.,
2001a; Garcia et al., 2007] are available to estimate
the paleotopography and the thickness of the rejuvenated lavas in detail. Therefore, while the paleotopography could have an important effect on the
thickness distribution of the rejuvenated lavas, it is
not well-constrained on Savai‘i. The only other
location where a minimum thickness can be estimated is a 70 m rejuvenated lava escarpment near
the northern coast line (Figure 2) [Workman et al.,
2004], suggesting the thickness of the rejuvenated
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Figure 7. Pb and Sr isotope results showing that the canyon lavas are compositionally the same as rejuvenated lavas
from Savai‘i and Upolu. Data sources as in Figure 4. Sample U-10 S of Workman et al. [2004] was omitted (see
Figure 4).
lavas remains great across the island. We therefore
assume that the subaerial part of Savai‘i volcano is
covered by a minimum 200 m thickness estimated
in the canyon, despite known offshore occurrence
of rejuvenated lavas [Jackson et al., 2007, 2010a].
The observed eruption volume of the 1905 eruption
lends credibility to this approach as well. The flows
filled a 10 mile canyon with a lava series up to
130 m thick [Anderson, 1910], and the rare occurrence of deeply eroded canyons compared to Kaua‘i,
suggests that rejuvenated volcanism on Savai‘i can
easily smooth out an eroded shield like Tutuila or
Upolu. Assuming a thickness of 200 m results in a
total volume estimate of rejuvenated lavas on Savai‘i
of 0.75 103 km3. This is 1.8% of the total
volume of Savai‘i, which at 42 103 km3 is similar to the total volume of similarly aged Hawaiian
volcanoes (Figure 2) [Robinson and Eakins, 2006].
This likely still represents a conservative estimate,
since it uses the minimum thickness of 200 m
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and it assumes that offshore rejuvenated eruptions
(reported in Jackson et al. [2010a]) are volumetrically unimportant. Alternatively, if the veneer is
200 m near the summit of Savai‘i and tapers off,
disappearing at the coastline (an underestimate since
a 70 m escarpment is observed near the north coast),
a minimum estimate of 0.091 103 km3 is obtained
(0.22% of the total volume of Savai‘i). Thus, even
the most conservative estimate of the Samoan rejuvenated lava volume is nearly double the 0.1% relative volume estimated for the largest rejuvenated
stage outpouring in Hawai‘i (0.058 103 km3 on
Kaua‘i [Garcia et al., 2010]), and this is qualitatively
consistent with the areal extent (Figure 2) of the
rejuvenated lavas on Kaua‘i (half the island) versus
Savai‘i (nearly the entire island). However, using our
preferred estimate for Samoa (1.8%) suggests that the
erupted volume of rejuvenated lavas on Savai‘i
is more than an order of magnitude larger than that
for Kaua‘i.
5.2. Does Samoa’s Tectonic Setting Explain
the Large Volumes of Rejuvenated Lavas?
[17] A combination of modeling and observations
near the Tonga Trench terminus can be shown to be
in agreement with the idea that Samoa’s unique
tectonic setting may have an impact on the mantle
melting that produces the Samoan Islands. The
Tonga Trench south of the Samoan Islands is part
of a tectonically complex zone consisting of subduction along the Tonga-Kermadec subduction
zone and back-arc spreading in the Lau Basin,
which together accommodate the very fast convergence of the Pacific and Australian Plates (26–
27 cm/a [Hart et al., 2004]). While the Pacific Plate
is subducted south of the terminus of the Tonga
Trench, the Pacific plate including the Samoan
Islands north of the terminus continues further west
along a seismically observed tear in the plate that is
expressed as the trace of the Vitiaz Lineament [e.g.,
Millen and Hamburger, 1998; Hart et al., 2004].
The Samoan Islands are located just 100 km north
of the intersection of the Tonga Trench with the
Vitiaz Lineament, and their proximity to this complex plate boundary requires an assessment of the
role of the nearby trench on proposed models for
the origin of rejuvenated lavas.
[18] Owing to plate flexure outboard of the trench,
which extends north of the northern terminus, the
Pacific Plate is upwarped in the Samoan region
[Levitt and Sandwell, 1995; Govers and Wortel,
2005]. The resulting upward mantle flow under
the elastic part of the plate may cause small
amounts of decompression melting, in a manner not
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unlike that proposed by Bianco et al. [2005] for
plate flexure due to volcanic loading in Hawai‘i.
Bathymetric and gravity data across the trench have
been used to model flexure of the Pacific Plate in the
Samoan region. The data from three profiles suggest
a flexural amplitude of 250–1000 m, resulting
from tectonic forces from the Tonga trench [Levitt
and Sandwell, 1995]. Govers and Wortel [2005]
assumed visco-elastic rheology and employed a
finite element model of the trench-transform intersection to evaluate the near-instantaneous response
of the lithosphere and underlying mantle to a high
density, subducting slab. The results suggest significant flexure-driven uplift (1000 m) of the
lithosphere near Savai‘i (Figure 6), and the shallow upward velocities (1–4 mm/year) associated
with bending extend into the asthenosphere. While
tectonic forces associated with the nearby Tonga
trench have been suggested to crack the plate
parallel to the Vitiaz Lineament [Natland, 1980],
we suggest that plate upwarping drives decompression melting in the Samoan region. Below, we
assess whether this melt-generating mechanism
is greater in magnitude, but qualitatively similar to,
the flexure-driven melting in Hawaiian rejuvenated lavas.
5.3. Rejuvenated Melt Production
Due to Flexure in Samoa and Hawai‘i
[19] Melt production due to the flexural response of
the Pacific Plate under the growing load of a new
Hawaiian volcano has been modeled by Bianco et al.
[2005] with a 3D finite element model that includes
an evaluation of melt production. A full 3D numerical treatment of the complicated tectonic region
around Samoa goes beyond the scope of this work.
However, we have calculated a 1D melt production
example (a parcel of mantle flowing along the top of
the asthenosphere) and compare the cases of Samoa
and Hawai‘i using this approach (Figure 8). Our
calculation is based on the work of Phipps Morgan
[2001]. We estimate melt production per unit time
(dF/dt in Ma1) due to decompression:
dF dF dP
¼
⋅
dt
dP dt
where F is the melt fraction, P is pressure and t is
time. We follow Bianco et al. [2005] to estimate
dP/dt, using the vertical (upwelling) velocity:
dP
¼ r g vz
dt
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Figure 8. Calculated 1D melt productivity assuming decompression due to flexural uplift. (top) Productivity per Ma
at a given location due to the vertical upwelling along the flexural profile as a function of distance from Vailulu‘u.
Upwelling velocities for Samoa are estimated from a cross section through the results of Govers and Wortel [2005]
(see Figure 6), while the Hawaiian profile is derived from the flexural profile shown in Bianco et al. [2005] and the
total uplift they suggest in 1 Ma. (bottom) Cumulative melt production that would be generated while mantle flows
along the flexural profile as a function of time since initial hot spot volcano construction (assuming average plate
motion of 70 km/Ma). Particularly, Savai‘i and Upolu have already experienced the largest amount of melt production,
while Tutuila is on the cusp, and the younger volcanoes currently are suggested to have very small melt production due
to limited flexure.
This formulation only considers lithostatic decompression, using the density (r), acceleration of
gravity (g), and upwelling velocity (vz). The
upwelling velocity used here is derived from a cross
section through the results of Govers and Wortel
[2005] (see Figure 6), which corresponds to the lithospheric uplift velocity, providing the maximum
velocity for the asthenosphere directly underneath.
Evidence that their model produces reasonable
results comes from the flexural amplitude (up to
1000 m) calculated by Levitt and Sandwell [1995]
for the Tonga Trench. Next, we use Phipps Morgan’s
[2001] formulation to estimate the second term,
dF/dP:
fraction increments starting from the current hot
spot (Vailulu‘u) following the chain is the cumulative
melt fraction that would have been produced through
time (bottom panel). For the Hawaiian model, a
scaled (100 m uplift in 1 Ma) flexural profile from
Bianco et al. [2005] and its cumulative melt fraction
predict that rejuvenated melt production in Hawai‘i
is nearly ten times smaller, in agreement with a
flexural origin for Samoan rejuvenated volcanism.
A different feature of the model that is consistent
with observations is that only the Samoan islands
with significant flexurally induced melt production
host rejuvenated lavas.
m
∂T
agT
∂z F
c
dF
1
mp
¼
∂T
dP S
rg TDS m
þ
cp
∂F P
estimate (particularly dTm/dF) are not well characterized, and therefore the calculated curves may have
to be scaled appropriately. In general only rough
estimates are available for dry peridotite [Phipps
Morgan, 2001], while even less is known for other
lithologies, including metasomatized peridotite. As a
result, for the comparison of Samoa to Hawai‘i equal
values for dF/dP are used at both localities. The
difference in upwelling velocity due to flexure
therefore controls the difference in melt generation
in this calculation. This is probably a reasonable
assumption, considering the observed larger volumes
of Samoan rejuvenated lavas relative to Hawai‘i
can be produced by melting to a larger degree,
and/or melting a larger source volume. The dTm/dF
and we use his estimates for its parameters
(∂Tm/∂z = 4 C/km for the solidus gradient with
depth; agT/cp = 0.4 C/km for the temperature gradient for adiabatic expansion; TDS/cp = 550 C for
the enthalpy of fusion; and ∂Tm/∂F = 250 C for the
solidus depletion gradient). The calculated productivity per 1 Ma at a single location per 1 Ma (dF/dt;
Figure 8) shows a nonlinear profile across the
flexural bulge (Figure 8). The sum of these melt
[20] One caveat is that the parameters for the dF/dP
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term can change the most significantly, especially
showing an increase due to very low degree melting
[Phipps Morgan, 2001], which logically predicts
lower melt production (dF/dP). Therefore, other
reasonable values for dTm/dF only amplify the difference between Samoa and Hawai‘i. Alternatively,
melting of a larger source volume requires a larger
volume of mantle near its solidus and efficient melt
segregation. However, in the flexural calculation,
the region with vertical upwelling velocities is
larger in Samoa (Figure 8), and therefore a flexural
origin would still be consistent with the available
data. Therefore, we suggest that flexural upwarping
of the Pacific Plate is indeed quantitatively consistent with more rejuvenated-stage melt production
in Samoa. Moreover, this process generates large
quantities of rejuvenated volcanism in Samoa that
can erupt along established conduits of the island
volcanoes as they are rafted past the northern terminus of the trench.
5.4. Can the Hawaiian Flexural Model Be
Modified to Explain Samoan Rejuvenation?
[21] The source for rejuvenated lavas is isotopi-
cally depleted in many intraplate volcanoes (e.g.,
Hawai‘i) relative to shield-stage lavas, and rejuvenated lavas tend to trend toward a depleted mid
ocean ridge basalt (MORB) mantle source composition [e.g., Lanphere and Dalrymple, 1980; Stille
et al., 1983; Roden et al., 1994a; Chen and Frey,
1985]. Such a component has been proposed to
exist as an independent component in the plume
source, or as a MORB-related component in the
lithosphere or the shallow depleted mantle [Chen
and Frey, 1985; Reiners and Nelson, 1998;
Lassiter et al., 2000; Yang et al., 2003; Frey et al.,
2005; Garcia et al., 2010]. No consensus has
emerged regarding the origins of rejuvenated volcanism, and there is still significant uncertainty
regarding the mantle source (plume or mantle lithosphere) of rejuvenated volcanism.
[22] Similarly, the mechanism responsible for gen-
erating rejuvenated lavas is still not well known,
and the eruptive hiatus that separates rejuvenated
volcanism from shield stage volcanism (e.g., 0.5–
2 Ma in Hawai‘i [Garcia et al., 2010]) is a major
characteristic that needs an explanation. This hiatus
is not a natural result of any model for the internal
structure of a mantle plume. Some models for
rejuvenated volcanism place the rejuvenated mantle
sources in the plume, but vary the source geometry. For example, suggested geometries for the
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rejuvenated mantle component include concentric
zoning of the plume, filament structures or small
blobs within the plume [e.g., DePaolo et al., 2001b;
Abouchami et al., 2005; Marske et al., 2007].
However, these models fail to explain why a new
phase of activity initiates after a period of protracted quiescence, when the mantle plume
is hundreds of km away. As a result, plumelithosphere interaction has developed as an alternative for explaining the initiation of rejuvenated
volcanism. This class of models is divided into three
broad categories: 1.) Conductive melting of the
lithosphere driven by lithospheric heating from the
plume [Gurriet, 1987], 2.) Decompression of plume
material by lateral spreading of the plume head
underneath the lithosphere [Ribe and Christensen,
1999; Sherrod et al., 2003; Paul et al., 2005],
or 3.) Decompression due to plate flexure around a
newly constructed volcanic shield [ten Brink and
Brocher, 1987; Bianco et al., 2005].
[23] The Samoan volcanoes can contribute to the
evaluation of these models, given their exceptional
volumes of rejuvenated lavas and their unusual geochemical composition (Figure 1) [Hawkins and
Natland, 1975; Wright and White, 1987; Workman
et al., 2004; Koppers et al., 2008]. It is unlikely that
conductive heating or plume spreading can generate
more melt in Samoa compared to Hawai‘i, since the
estimated buoyancy flux for the Samoan plume
source is 6 times smaller and the potential temperature is lower than that for Hawai‘i [Courtillot et al.,
2003]. Instead of plume spreading or conductive
heating, we consider the flexural model the most
appropriate for generating rejuvenated lavas in
Samoa. The proximity to the Tonga Trench may lead
to more substantial flexure-driven melting compared
to the model for Hawai‘i [Bianco et al., 2005].
[24] The large volumes of rejuvenated lavas in
Samoa (up to 10 times more than in Hawai‘i) provide
a unique opportunity to assess the potential role of
lithospheric flexure in rejuvenated volcanism. In a
plate flexure model as applied in Hawai‘i, the load of
a large volcanic edifice on an elastic plate can cause
the plate to upwarp at a distance of 200–300 km, and
the resulting decompression may generate melts [ten
Brink and Brocher, 1987; Bianco et al., 2005].
Unlike Hawai‘i, which is located far from any plate
boundary, Samoa is located next to the northern terminus of the Tonga Trench, and we suggest that
melting under subducting and bending lithosphere
causes greater flexure-driven melting due to the
greater flexural amplitude compared to Hawai‘i.
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The plate flexure in Samoa (from 250–1000 m
[Levitt and Sandwell, 1995; Govers and Wortel,
2005]) can create greater uplift and decompression melting than the magnitude of plate flexure
predicted for Hawai‘i (100 m [e.g., Bianco et al.,
2005]). This observation makes decompression
melting due to flexural uplift an attractive explanation for explaining the greater volumes of rejuvenated volcanism in Samoa. Therefore, compared
to Hawai‘i, larger volumes of rejuvenated melt are
predicted (and observed) in Samoa.
5.5. What is the Difference With Previous
Lithosphere-Based Models for Samoan
Rejuvenated Volcanism?
[25] Hawkins and Natland [1975] proposed that tec-
tonic stresses in the Samoan lithosphere played an
important role in generating the unusually large
volumes of rejuvenated lavas along nearly 300 km of
volcanic chain, from Tutuila to Savai‘i. Bending of
the lithosphere southward, toward the tearing and
subducting plate, was proposed to create tensional
faults that provide pathways for melt [Hawkins and
Natland, 1975; Natland, 1980; Natland and Turner,
1985] (inset in Figure 6). However, 3D numerical
(finite element) modeling of the lithospheric response
to the Pacific Plate tearing and subducting suggests
that plate bending generates low strain rates under the
Samoan chain (Figure 6) [Govers and Wortel, 2005].
Strain rates are only high well within 100 km of the
subduction-transform intersection at the northern
terminus of the Tonga Trench (i.e., 100 km south
of Savai‘i Island). In contrast to the localized high
strain rates, the Govers and Wortel [2005] model
does show that the Samoan lithosphere flexurally
bends up around the northern terminus of the Tonga
trench, with upward velocities extending into the top
80 km of the asthenosphere. In fact, the vertical
velocity field of Govers and Wortel [2005] correlates
very well with the location of rejuvenated lava
eruption with Savai‘i, Upolu, and Tutuila, which are
located near the maximum in vertical velocities
(Figures 6 and 8). Similarly, upward mantle flow
caused by plate flexure under the flexural bulge of
Hawai‘i has been modeled to generate rejuvenated
lavas [Bianco et al., 2005]. It is important to note that
the flexural amplitude of the Pacific Plate entering
the Tonga Trench is at up to 10 times as large as the
flexural amplitude in Hawai‘i that is only related to
flexural loading instead of bending into a subduction
zone [Levitt and Sandwell, 1995; Govers and Wortel,
2005]. Additionally, in the Hawaiian flexure model,
the region in which lithospheric uplift and upwelling
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occurs is <200 km across. By comparison, the
region affected by flexural uplift near the Tonga
Trench spans 300 km along the Samoan hot spot.
From these observations, we suggest that upward
mantle flow beneath Samoa caused by lithospheric bending generates enhanced volumes of melt
through decompression melting.
6. Plate-Flexure Induced Melts
of the Lithospere in Samoa:
Evidence From Isotopes
[26] In the following sections we assess whether tra-
ditional heavy radiogenic isotopes (Sr-Nd-Pb), noble
gas isotopes and Os isotopes for the Samoan rejuvenated lavas are all in agreement with melt generation
from lithospheric flexure. Since the predicted flexural
melts are all shallowly based (section 5.3), it is
important to consider the geologic history of the
Pacific Plate. Plate reconstructions show that the
region of the Pacific plate now occupied by Samoa
passed over the Rarotonga hot spot 15 million years
ago (inset in Figure 1) [e.g., Chauvel et al., 1997;
Wessel and Kroenke, 2008; Konter et al., 2008;
Jackson et al., 2010a]. Given the similar isotopic and
trace element composition to the Rarotonga lavas, we
will argue that the Samoan rejuvenated mantle source
may contain a Rarotonga hot spot signature hosted in
the shallow Pacific mantle instead of being hosted
entirely in the upwelling Samoan mantle. During the
shield stage, such a component would likely be overwhelmed by melting of the Samoan mantle source,
since the predicted amount of melting due to flexure
is very low during the shield stage (Figure 8 and
section 5.3). With limited decompression melting due
to flexure, conductive melting due to heat transfer
from hot plume material is another potential cause of
melting of the rejuvenated source material. However,
this has been modeled in Hawai‘i by Gurriet [1987],
and would cause a delay in melting of 2–10 Ma,
suggesting melt production from heat conduction
into the lithosphere during the shield stage should
be minimal as well. Moreover, heat conduction and
the delay in melting it causes [Gurriet, 1987] is
unlikely to be the cause for rejuvenated volcanism
by itself, since the likely hotter mantle in Hawai‘i
[Courtillot et al., 2003; Herzberg and Gazel, 2009] is
producing smaller volumes of rejuvenated melt than
Samoa. Therefore, the lithosphere-driven hypothesis
is our preferred model, and we will assess below
whether the geochemical predictions associated
with this model (e.g., Rarotonga influence) can be
confirmed.
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Figure 9. Pb-Sr isotope data show that Samoan, rejuvenated lavas trend toward lithosphere metasomatized by a
Rarotonga (EM1) type composition. D 8/4 [Hart, 1984] defines the deviation from the mantle array and allows for a
more detailed view of the Pb data and mixing relationships. Rarotonga hot spot (purple) from Chauvel et al. [1997].
Samoa data: Workman et al. [2004] and Jackson et al. [2007, 2010a]. A mixing model is shown that explains the rejuvenated lava compositions as a result of lithosphere metasomatism by a Rarotonga component (0.5%; green arrow),
followed by mixing of melts from the metasomatized lithosphere with a small amount (<2%; orange arrow) of Samoan
mantle melts. Data sources as in Figure 4; EPR field based on data from the PETDB database. Sample U-10 S of
Workman et al. [2004] was omitted (see Figure 4).
6.1. Evidence From the Isotopes of
Incompatible Trace Elements (Sr-Nd-PbHe-Ne) for Involvement of EM1
(Rarotonga) Metasomatized Lithosphere
in Samoan Rejuvenated Lavas
[27] The rejuvenated lavas are distinct from the
shield lavas in Samoa, with a rejuvenated composition that is moderately elevated in 208Pb/204Pb for
its low 206Pb/204Pb (high D8/4 [Hart, 1984]), and
intermediate in 87Sr/86Sr and 143Nd/144Nd compared to the shield (Figure 9). Although individual
samples are scattered, the rejuvenated field is
elongated away from the shield compositions, and
it is nearly parallel to the East Pacific Rise (EPR)
MORB field in Pb isotope space (Figure 9). The
rejuvenated lavas have Pb isotopic compositions
that overlap with Rarotonga lavas. However, this
overlap does not exist for Pb-Sr isotope compositions. Instead the rejuvenated lavas trend toward
EM1-like isotope compositions in Rarotonga. In
more detail, the scatter of the Samoan rejuvenated
lava compositional field implies that the source
likely contains at least three components. Since the
compositions of the Savai‘i rejuvenated lavas are
all contained between the Samoan shield lavas, the
EPR MORB, and the most EM1-rich sample from
Rarotonga, we explore whether there is evidence
for a mixture of a Rarotonga component and a
lithospheric component in the Samoan rejuvenated
lavas. In this model, the Samoan component may
represent either the admixture of the Samoan shield
stage lavas with the lithosphere (in a similar way
to the Rarotonga metasomatizing component), or
Samoan plume mantle that was not completely
depleted during the shield stage.
Pb/204Pb–208Pb/204Pb
array of rejuvenated lavas stretches from the
Samoan shield lavas to a composition between an
extreme Rarotonga composition and EPR MORB,
it appears that an initial mixture between Rarotonga
and MORB may be involved. Therefore, if a
Rarotonga component is truly present in the
Samoan source, it likely mixed with a Pacific
MORB component, prior to exposure to Samoan
[28] Since the elongated
206
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shield compositions. One way to create such a
Rarotonga-lithosphere mixture would be metasomatic overprinting of lithospheric compositions
with Rarotonga melts when the Pacific Plate passed
over the Rarotonga hot spot. Pilet et al. [2008]
recently showed that ocean island alkalic basalts
(including nephelinites and basanites) can be generated from melting of metasomatized mantle, and
we here extend this idea to Samoan rejuvenated
volcanism by using an “easy-to-melt by flexure”
component hosted in the lithosphere. Since our
best estimate for a non-contaminated lithospheric
composition is EPR MORB, a range of potential
lithospheric compositions needs to be assessed to
describe the potential metasomatic process. To a
first order, we can assume a composition for a
Rarotonga component based on the most extreme
EM1-type Rarotonga sample of Nakamura and
Tatsumoto [1988], and mix such a melt with
either the “depleted” or “enriched” depleted (solid)
mantle composition of Workman and Hart [2005].
These compositions require only about 0.5% of the
Rarotonga melt component to explain the Pb isotope composition of the rejuvenated lavas, which
plot halfway between EPR MORB and Rarotonga
(Figure 9), while a less extreme Rarotonga composition would require a slightly larger melt contribution. If we adopt this 0.5% mixture (composed of
Rarotonga plus DMM) as one (solid) end-member
and a Samoan shield melt composition (dredge 115
from Savai‘i [Jackson et al., 2007]) as the other
(melt) component, then only about 1–2% contribution from the Samoan shield source is required to
explain the overall isotopic compositions of the
Samoan rejuvenated lavas. Overall, these calculations are consistent with our hypothesis that Samoan
rejuvenated lavas contain a metasomatic Rarotonga
component.
[29] If the Samoan rejuvenated lavas are melts of
MORB mantle metasomatized by the Rarotonga
hot spot, they would also be expected to host noble
gas isotopic signatures associated with an EM1
metasomatic component. Based on He and Ar isotope systematics, Burnard et al. [1998] argued for
multiple pulses of metasomatism in Samoan xenoliths, confirming earlier reports of extensive metasomatic overprinting of the Sr-Nd-Pb isotopic
compositions of Samoan xenoliths from the lithospheric mantle [Hauri et al., 1993; Hauri and Hart,
1994]. Other clues about the noble gas isotopic
composition of shallow mantle beneath Samoa
come from analyses of noble gases in Samoan
peridotite mantle xenoliths hosted in Savai‘i rejuvenated lavas [Poreda and Farley, 1992]. In a plot
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of He versus Ne isotopes (Figure S2 in the auxiliary
material), the Samoan xenoliths trend from a component similar to MORB (but with slightly less
radiogenic He) toward a component similar to the
EM1 mantle found in Pitcairn lavas [Honda and
Woodhead, 2005]. Since Ne isotopes have not
been measured in Rarotonga lavas, we assume that
the He-Ne signature for EM1 in Rarotonga lavas is
similar to the EM1 lavas from Pitcairn, given that
Rarotonga forms a trend that partly overlaps with
Pitcairn in Pb-Sr-Nd isotope space [Woodhead and
McCulloch, 1989]. A subset of the Samoan mantle
xenoliths indeed trend toward the EM1 composition we might expect based on observed Pitcairn
compositions. The Samoan xenoliths erupted during the rejuvenated stage plot between MORB,
EM1 and EM2 (Samoan shield lavas [Jackson
et al., 2009]). Therefore, the He-Ne isotope
observations in the rejuvenated lavas, similar to
Sr-Nd-Pb isotope compositions, appear to be consistent with our proposed model for a metasomatic
Rarotonga component in Samoan lithosphere.
[30] Combining all the results, the isotopic signatures
of Sr, Nd, Pb, He and Ne define a unique geochemical composition for Samoan rejuvenated lavas that
can be explained by metasomatic fluids from the
Rarotonga hot spot. We have shown with a mixing
calculation that a combination of an EM1 Rarotonga
component, the MORB mantle, and the EM2 shield
component can explain the Samoan rejuvenated
lava composition. Our model assumes that the
EM1 component can be contributed by the geochemical signature of Rarotonga (e.g., Rarotonga:
206
Pb/204Pb = 18.256–18.975, 207Pb/204Pb = 15.481–
15.564, 208Pb/204Pb = 38.598–38.994, 87Sr/86Sr =
0.70412–0.70444, 143Nd/144Nd = 0.512629–0.512750
[Nakamura and Tatsumoto, 1988; Schiano et al.,
2001]). In our model, this metasomatic component
is sampled by melting due to lithospheric flexure,
where upward mantle flow and melting is concentrated near the bottom of the plate [Govers and
Wortel, 2005]. According to our model, all the
isotope systems described above would be affected
by metasomatism, and we therefore need to evaluate evidence for lithospheric contributions using an
isotope system not affected by metasomatism.
6.2. Evidence From Os Isotopes for Lithosphere
Involvement in the Genesis of Samoan
Rejuvenated Lavas
[31] In order to evaluate the potential contribution
from the Pacific lithosphere, we here discuss the
available Os isotope data for Samoa. Since the
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parent element Re is incompatible while Os is
compatible during partial melting of the mantle,
melts will have an increased Re/Os ratio while
mantle peridotites will have a greatly diminished
Re/Os ratio [Shirey and Walker, 1998; Bizimis
et al., 2007]. Consequently radiogenic ingrowth
of 187Os is slow in peridotites, yielding low
187
Os/188Os ratios over time. Modification of these
low ratios by metasomatic effects is expected to be
minimal, given the relatively high Os concentrations in peridotites compared to metasomatizing
melts or fluids [Chesley et al., 2004; Bizimis et al.,
2007]. Moreover, the relatively short amount of
time (15 Ma) since metasomatism took place
would have prohibited significant radiogenic 187Os
ingrowth, even if Re were added by the metasomatic fluids. Melts that interact with the high Os
content mantle are likely to inherit the low Os isotope ratios [Shirey and Walker, 1998]. Thus, the
low 187Os/188Os of the lithospheric mantle may be
preserved and it would likely be one of the best
ways to establish a lithospheric origin for the rejuvenated lavas.
[32] Indeed, low Os isotope ratios are observed in
the Samoan lithospheric mantle samples studied by
Hauri and Hart [1993] and Jackson et al. [2010b],
even though noble gases and Sr-Nd-Pb isotopes
indicate that they have been highly metasomatized [Hauri et al., 1993; Burnard et al., 1998].
We suggest that the Sr-Nd-Pb-He-Ne isotopes
reflect a metasomatized Rarotonga plume signature
(Figure 9) that is decoupled from Os isotopes that
may reflect a lithospheric signature, owing to the
incompatibility of Sr, Nd and Pb isotopes compared
to Os. In fact, Samoan rejuvenated lavas host
among the lowest 187Os/188Os ratios among ocean
island basalts (OIBs) globally, with ratios that
extend down to 0.123 [Hauri and Hart, 1993].
When compared to the Os isotope data on Samoan
shield lavas, none of the rejuvenated lavas from
Savai‘i (>30 ppt Os) have 187Os/188Os that overlap
with Samoan shield lavas [Hauri and Hart, 1993;
Jackson and Shirey, 2011] (Figure S3 in the
auxiliary material). Critically, however, the Savai‘i
rejuvenated lavas form a trend that extends to lower
187
Os/188Os than shield lavas (Figure S3 in the
auxiliary material), and we explore whether this can
be explained with our metasomatized lithospheredriven melting model for the rejuvenated stage.
[33] The lithospheric mantle in the Samoan region
has remarkably unradiogenic Os signatures that
overlap with the very low 187Os/188Os ratios measured in Samoan rejuvenated lavas. Jackson et al.
[2010b] reported 187Os/188Os on 12 peridotite
10.1029/2011GC003974
mantle xenoliths from Samoa, and half of the
xenoliths yielded 187Os/188Os ratios equal to or
lower than 0.123, the lowest 187Os/188Os ratio
reported in Os-rich (>30 ppt Os) Savai‘i rejuvenated
lavas (all < 0.127). Additionally, the same set of
peridotite xenoliths examined for 187Os/188Os were
previously shown by Hauri and Hart [1994] to have
pressures of <2.5 GPa, based on the Ca olivineclinopyroxene exchange barometer of Köhler and
Brey [1990], which places the xenoliths within the
mantle lithosphere beneath Savai‘i. The shallow
origin suggests that the xenoliths do not sample a
sub-lithospheric (or “plume”) source, but instead
represent the oceanic mantle lithosphere beneath
Samoa. Therefore, the lithospheric mantle is a suitable geochemical reservoir for the low Os isotope
ratios observed in the Samoan rejuvenated lavas.
[34] In Hawai‘i, Os-isotopic systematics were also
used to argue that rejuvenated lavas are melts derived
from the oceanic mantle lithosphere [Lassiter et al.,
2000]. In Hawai‘i the depleted Sr-Nd-Pb isotopes,
combined with a lack of correlation with Os isotopes, were used to argue against a plume source in
rejuvenated lavas. Lassiter et al. [2000] argue that a
plume origin for the Os isotope signature would
imply significant contributions to the Sr-Nd-Pb
isotopes as well, which would likely not generate
the depleted signature observed in Hawaiian rejuvenated lavas. Lassiter et al. [2000] also argue that
metasomatized lithosphere is unlikely to be a source,
since the expected relation of silica content with Os
isotope ratios is opposite of what would be expected
for this scenario. Finally, Lassiter et al. [2000] suggest that the Hawaiian rejuvenated stage Os isotope
compositions may be the result of melting of a
lithospheric pyroxenite component combined with
a peridotitic component. In their model, large
degree pyroxenite melts contribute to much of the
Os-isotopic signature, while the other radiogenic
isotope signatures are controlled by small degree
peridotite melts. This argument hinges on pyroxenites being formed at the mid ocean ridge at
100 Ma as trapped melts or their cumulates.
However, the Pacific Plate’s previous passage over
Rarotonga before it reached Samoa is vital here;
in contrast to the Pacific plate “upstream” of the
Samoan hot spot, no hot spot is known to have
occupied the region of the Pacific plate now occupied by Hawai‘i. In the case of Samoa, old midocean ridge-related pyroxenites are unlikely to be
present due to the low pyroxenite solidus and high
melt productivity [e.g., Hirschmann and Stolper,
1996], which would have caused melting and
removal of most of this component during passage
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over the Rarotonga hot spot. As a result, even
though a lithospheric component may be present
in Samoan rejuvenated lavas, its composition would
likely not be the same as that suggested by Lassiter
et al. [2000] for Hawai‘i. It is, therefore, not
surprising that rejuvenated melts sampling the
lithosphere beneath Hawai‘i exhibit a more depleted
Sr-Nd-Pb isotope signature than Samoan rejuvenated lavas.
[35] Despite the agreement with our predictions, the
data support our model but do not exclude other
interpretations. For example, we cannot exclude a
model where rejuvenated lavas are generated by contamination of plume melts with lithosphere containing
high Os concentration and low Os isotope ratios,
a model that would imply mostly sub-lithospheric
melting that is at odds with our lithosphere-driven
model. We consider this model less likely given the
eruption of the rejuvenated lavas starts several million
years after the initial construction of a volcano,
implying the volcano has been transported up to
hundreds of km away from the original mantle
(plume) source. Additionally, we consider such a
model to be unlikely because it does not explain why
shield stage lavas have more radiogenic 187Os/188Os
ratios even though they pass through the same lithosphere as rejuvenated melts.
7. Evidence for Lithospheric Involvement
From Lower Temperatures
[36] Another approach for evaluating the litho-
spheric contribution of rejuvenated volcanism
stems from P-T estimates using major element
compositions for primitive lavas. Therefore, we
assess whether thermobarometry for the melts can
be used to test for the model of involvement of
lithosphere previously affected by Rarotonga. We
might expect melting depths and temperatures to
shift from deeper plume-based melting to shallow
flexurally driven melting, and such a shift does
seem present in temperature estimates based on
lava major element compositions.
[37] Melting is expected to occur at shallower levels
for rejuvenated compared to shield lavas based on
the 3D numerical model of Govers and Wortel
[2005] that shows upwelling velocities concentrated in the lithosphere and decreasing into the top
80 km of the asthenosphere while the estimates of
Farnetani and Hofmann [2010] for plume melting
in Hawai‘i suggest depths down to 175 km.
Therefore, a shift in temperatures and pressures
might be expected between the shield and rejuvenated stages, if our model is correct. As a first step,
10.1029/2011GC003974
we calculate conditions assuming a peridotite source
[Herzberg and Asimow, 2008; Lee et al., 2009],
assuming that the low Os isotope signatures in the
Samoan lavas [Hauri and Hart, 1993; Jackson and
Shirey, 2011] signifies that pyroxenite is not likely
to make up a significant component of the mantle
source for Samoan shield and rejuvenated lavas.
[38] To test this idea, we use CaO (and Al2O3)
versus MgO to filter for primitive lavas that have
not experienced clinopyroxene fractionation (see
Text S1 and Figure S4 in the auxiliary material) and
exclude any picritic samples in order to obtain P-T
estimates for the melts using the technique of Lee
et al. [2009]. Unfortunately, the majority of the
estimated pressures are around 4 GPa, based on the
silica activity in the melt, and assuming orthopyroxene and olivine are present [Lee et al., 2009].
However, at these pressures orthopyroxene is not
stable, except in combination with olivine and melt
at a degree of melting over 25% [Walter, 1998;
Herzberg and O’Hara, 2002], which is likely not
relevant to hot spots such as Samoa (where the
degree of melting is thought to be relatively low).
As a result, we only consider the temperature estimates reliable for the samples. Each temperature
estimate is offset from its solidus along a melting
adiabat, reflecting the degree of melting that took
place, and implying that the solidus was actually
crossed at a higher temperature than calculated.
Therefore, it is not the average temperature, reflecting average degrees of melting, but the highest temperatures for each stage that provide insight into the
difference in conditions between shield and rejuvenated lavas. This argument assumes that the highest
temperatures reflect the lowest degree melts from a
peridotite source with similar solidus, therefore being
closest to the solidus intersection. This is probably a
reasonable assumption, since rejuvenated lavas are
generally expected to be low-degree melts [e.g.,
Clague and Dalrymple, 1987].
[39] When the results for the temperature calcula-
tion are examined, it is clear that the shield lavas
range to higher temperatures than the rejuvenated
lavas by 50 C (Figure 10). Such a shift to lower
temperature is at least qualitatively consistent with
our model, where plume-driven deep melting gives
way to shallower, flexure-based melting. We tested
our results against Herzberg and Asimow’s [2008]
temperature estimates, which filter for pyroxenite
source material, clinopyroxene fractionation or
accumulation, volatile-rich sources, and allows for
adjustment of Fe oxidation state. Only a subset of
samples meet all the filter criteria of the Herzberg
and Asimow [2008] model (only 2 rejuvenated
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KONTER AND JACKSON: SAMOAN REJUVENATED LAVAS AND FLEXURE
Figure 10. Temperature estimates for melting (based
on Lee et al. [2009]), shown as histograms for both
shield and rejuvenated stage Samoan lavas. Samples
were filtered to exclude lavas with clear signs of clinopyroxene fractionation, evolved lavas, and picritic lavas.
Maximum temperatures likely best represent the difference in temperature conditions between shield and
rejuvenated volcanism.
lavas). This model suggests a similar difference in
temperature between shield and rejuvenated lavas
(75 C) even when Fe2O3/TiO2 is varied, although
absolute temperatures are slightly different than the
method by Lee et al. [2009]. Therefore, these models seem to be in agreement with melting under
lower temperature conditions during the rejuvenated stage.
[40] However, there is an important caveat. Since
we argue that the rejuvenated lavas may have a
metasomatized component in their source, it is
possible that the volatile content of this source is
significantly different from the shield stage source.
As a result, the solidus for the rejuvenated stage
may be different from that of the shield stage. In
particular, the presence of metasomatic phases
such as amphibole or phlogopite can affect the
solidus while still producing the observed alkalic
melts [e.g., Pilet et al., 2008]. However, hydrous
phases were never observed in the metasomatized
Samoan xenoliths [Hauri et al., 1993]. Nonetheless,
this uncertainty implies that the temperature estimates do not provide strong constraints on our model,
but they do seem consistent with it, to first order.
8. Conclusion
[41] The isotopic composition of new samples from
200 m down-section in a canyon on Savai‘i
(Samoa) allow us, for the first time, to constrain the
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unusually large volumes of rejuvenated lavas that
are erupted in Samoa. Such volumes and their
unusual isotopic compositions lead us to suggest
that the location of Samoa near the northern Tonga
Trench terminus plays an important role in the origin of this stage of volcanism. We propose a model
where subduction related upwelling of the shallow
mantle due to flexure causes decompression melting
involving a metasomatized lithospheric component.
Evidence for involvement of such a lithospheric
component comes from the composition of both
rejuvenated lavas and lithospheric xenoliths. These
compositions are in agreement with melts that contain a lithospheric component that includes a signature from the plate’s prior passage over the
Rarotonga hot spot. Therefore, the two features that
may explain the difference in erupted rejuvenated
lavas between Hawai‘i and Samoa are: (1.) the
location and flexure near a trench, and (2.) the
potential for metasomatism of the lithosphere, suggesting the lithosphere plays an active role during
this phase of volcanism. Evidence from Sr-Nd-Pb
isotope compositions of the rejuvenated lavas suggests a mixture involving a metasomatized MORB
component that incorporated a Rarotonga-type
composition and a subsequent Samoan component.
Os isotope signatures of lavas and xenoliths also
suggest involvement of a lithospheric mantle component. Furthermore, He-Ne isotope compostions
argue for a metasomatic component corresponding
to an EM1 mantle source in the Samoan lithosphere.
Last, temperature estimates for Savai‘i Island suggest a shift to lower temperatures during rejuvenated
stage volcanism, which might be expected if the
origin for magmatism shifts from a plume source to
a shallow source dominated by melting just below
and within the lithosphere. Data for each of these
areas by itself does not provide conclusive evidence
for our model, however the preponderance of data
that may be explained with our model argues that it
is likely the simplest explanation that explains the
most data, although additional data would allow us
to evaluate our model for each of these areas while
ruling out alternative models that cannot currently
be excluded.
Acknowledgments
[42] We would like to thank Warren Jopling, the good people
of the Safua Hotel, and the chief of Sili village and his sons for
making the field component a success. We are thankful for the
reviews by Loÿc Vanderkluysen, Christoph Beier, as well as
anonymous reviews and comments on various versions of this
work, comments from Claude Herzberg and Jim Natland, and
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KONTER AND JACKSON: SAMOAN REJUVENATED LAVAS AND FLEXURE
the editorial efforts by Joel Baker. We acknowledge NSF support, through grant EAR-0946752.
References
Abouchami, W., A. W. Hofmann, S. J. G. Galer, F. A. Frey,
J. Eisele, and M. Feigenson (2005), Lead isotopes reveal bilateral asymmetry and vertical continuity in the Hawaiian mantle
plume, Nature, 434, 851–856, doi:10.1038/nature03402.
Anderson, T. (1910), The volcano of Matavanu in Savaii, Q. J.
Geol. Soc. London, 66, 621–639, doi:10.1144/GSL.JGS.1910.
066.01-04.30.
Bianco, T. A., G. Ito, J. M. Becker, and M. O. Garcia (2005),
Secondary Hawaiian volcanism formed by flexural arch
decompression, Geochem. Geophys. Geosyst., 6, Q08009,
doi:10.1029/2005GC000945.
Bizimis, M., M. Griselin, J. C. Lassiter, V. J. M. Salters, and
G. Sen (2007), Ancient recycled mantle lithosphere in the
Hawaiian plume: Osmium-Hafnium isotope evidence from
peridotite mantle xenoliths, Earth Planet. Sci. Lett., 257,
259–273, doi:10.1016/j.epsl.2007.02.036.
Blichert-Toft, J., D. Weis, C. Maerschalk, A. Agranier, and
F. Albarède (2003), Hawaiian hot spot dynamics as inferred
from the Hf and Pb isotope evolution of Mauna Kea volcano,
Geochem. Geophys. Geosyst., 4(2), 8704, doi:10.1029/
2002GC000340.
Burnard, P. G., K. A. Farley, and G. Turner (1998), Multiple
fluid pulses in a Samoan harzburgite, Chem. Geol., 147,
99–114, doi:10.1016/S0009-2541(97)00175-7.
Chauvel, C., W. McDonough, G. Guille, R. Maury, and
R. Duncan (1997), Contrasting old and young volcanism in
Rurutu Island, Austral chain, Chem. Geol., 139, 125–143,
doi:10.1016/S0009-2541(97)00029-6.
Chen, C. Y., and F. A. Frey (1985), Trace-element and isotopic
geochemistry of lavas from Haleakala Volcano, East Maui,
Hawaii: Implications for the origin of Hawaiian basalts, J. Geophys. Res., 90, 8743–8768, doi:10.1029/JB090iB10p08743.
Chesley, J., K. Righter, and J. Ruiz (2004), Large-scale mantle
metasomatism: A Re-Os perspective, Earth Planet. Sci. Lett.,
219, 49–60, doi:10.1016/S0012-821X(03)00698-8.
Clague, D. A., and G. B. Dalrymple (1987), The HawaiianEmperor volcanic chain, part I: Geological evolution, in Volcanism in Hawaii, edited by R. Decker et al., U.S. Geol. Surv
Prof. Pap., 1350(1), 5–54.
Courtillot, V., A. Davaille, J. Besse, and J. Stock (2003), Three
distinct types of hotspots in the Earth’s mantle, Earth Planet.
Sci. Lett., 205, 295–308, doi:10.1016/S0012-821X(02)01048-8.
DePaolo, D. J., E. Stolper, and D. M. Thomas (2001a), Deep
drilling into a Hawaiian volcano, Eos Trans. AGU, 82,
149–155.
DePaolo, D. J., J. G. Bryce, A. Dodson, D. L. Shuster, and
B. M. Kennedy (2001b), Isotopic evolution of Mauna Loa
and the chemical structure of the Hawaiian plume, Geochem.
Geophys. Geosyst., 2(7), 1044, doi:10.1029/2000GC000139.
Duncan, R. A. (1985), Radiometric ages from volcanic rocks
along the New Hebrides-Samoa lineament, in Geological
Investigations of the Northern Melanesian Borderland, Earth
Sci. Ser., vol. 3, edited by T. M. Brocker, pp. 67–76, Circum.
Pac. Counc. for Energy and Miner. Resour., Houston, Tex.
Farnetani, C. G., and A. W. Hofmann (2010), Dynamics and
internal structure of the Hawaiian plume, Earth Planet. Sci.
Lett., 295, 231–240, doi:10.1016/j.epsl.2010.04.005.
Fekiacova, Z., W. Abouchami, S. J. G. Galer, M. O. Garcia,
and A. W. Hofmann (2007), Origin and temporal evolution
10.1029/2011GC003974
of Koolau Volcano, Hawaii: Inferences from isotope data on
the Koolau Scientific Drilling Project (KSDP), the Honolulu
Volcanics and ODP Site 843, Earth Planet. Sci. Lett., 261,
65–83, doi:10.1016/j.epsl.2007.06.005.
Frey, F. A., and J. M. Rhodes (1993), Intershield geochemical
differences among Hawaiian volcanoes: Implications for
source compositions, melting process and magma ascent
paths, Philos. Trans. R. Soc. London, Ser. B, 342, 121–136,
doi:10.1098/rsta.1993.0009.
Frey, F. A., W. S. Wise, M. O. Garcia III, H. West, S.-T.
Kwon, and A. Kennedy (1990), Evolution of Mauna Kea
Volcano, Hawaii: Petrologic and geochemical constraints
on postshield volcanism, J. Geophys. Res., 95, 1271–1300,
doi:10.1029/JB095iB02p01271.
Frey, F. A., S. Huang, J. Blichert-Toft, M. Regelous, and
M. Boyet (2005), Origin of depleted components in basalt
related to the Hawaiian hot spot: Evidence from isotopic
and incompatible element ratios, Geochem. Geophys. Geosyst.,
6, Q02L07, doi:10.1029/2004GC000757.
Garcia, M. O., E. H. Haskins, E. M. Stolper, and M. Baker
(2007), Stratigraphy of the Hawai‘i Scientific Drilling Project core
(HSDP2): Anatomy of a Hawaiian shield volcano, Geochem.
Geophys. Geosyst., 8, Q02G20, doi:10.1029/2006GC001379.
Garcia, M. O., L. Swinnard, D. Weis, A. R. Greene, T. Tagami,
H. Sano, and C. E. Gandy (2010), Petrology, geochemistry
and geochronology of Kaua‘i lavas over 4.5 Myr: Implications for the origin of rejuvenated volcanism and the evolution of the Hawaiian plume, J. Petrol., 51, 1507–1540,
doi:10.1093/petrology/egq027.
Geldmacher, J., K. Hoernle, P. van den Bogaard, G. Zankl, and
D. Garbe-Schonberg (2001), Earlier history of the ≥70-Ma-old
Canary hotspot based on temporal and geochemical evolution
of the Selvagen Archipelago and neighboring seamounts in
the eastern North Atlantic, J. Volcanol. Geotherm. Res., 111,
55–87, doi:10.1016/S0377-0273(01)00220-7.
Govers, R., and M. J. R. Wortel (2005), Lithosphere tearing at
STEP faults: Response to edges of subduction zones, Earth
Planet. Sci. Lett., 236, 505–523, doi:10.1016/j.epsl.2005.
03.022.
Gurriet, P. (1987), A thermal model for the origin of posterosional alkalic lava, Hawaii, Earth Planet. Sci. Lett., 82,
153–158, doi:10.1016/0012-821X(87)90115-4.
Hanan, B. B., and J.-G. Schilling (1989), Easter microplate evolution: Pb isotope evidence, J. Geophys. Res., 94, 7432–7448,
doi:10.1029/JB094iB06p07432.
Hanan, B. B., J. Blichert-Toft, D. Pyle, and D. Christie (2004),
Contrasting origins of the upper mantle MORB source
revealed by Hf and Pb isotopes from the Australian-Antarctic
Discordance, Nature, 432, 91–94, doi:10.1038/nature03026.
Hanan, B. B., J. W. Shervais, and S. K. Vetter (2008), Yellowstone plume-continental lithosphere interaction beneath the
Snake River Plain, Geology, 36, 51–54, doi:10.1130/
G23935A.1.
Hart, S. R. (1984), A large-scale isotope anomaly in the Southern Hemisphere mantle, Nature, 309, 753–757, doi:10.1038/
309753a0.
Hart, S. R., et al. (2000), Vailulu‘u undersea volcano: The
New Samoa, Geochem. Geophys. Geosyst., 1(12), 1056,
doi:10.1029/2000GC000108.
Hart, S. R., M. Coetzee, R. K. Workman, J. Blusztajn, K. T. M.
Johnson, J. Sinton, B. Steinberger, and J. W. Hawkins
(2004), Genesis of the Western Samoa seamount province:
Age, geochemical fingerprint and tectonics, Earth Planet.
Sci. Lett., 227, 37–56, doi:10.1016/j.epsl.2004.08.005.
20 of 23
Geochemistry
Geophysics
Geosystems
3
G
KONTER AND JACKSON: SAMOAN REJUVENATED LAVAS AND FLEXURE
Hauri, E. H. (1996), Major-element variability in the Hawaiian
mantle plume, Nature, 382, 415–419, doi:10.1038/
382415a0.
Hauri, E. H., and S. R. Hart (1993), Re-Os isotope systematics
of HIMU and EMII oceanic island basalts from the South
Pacific Ocean, Earth Planet. Sci. Lett., 114, 353–371,
doi:10.1016/0012-821X(93)90036-9.
Hauri, E. H., and S. R. Hart (1994), Constraints on melt migration from mantle plumes: A trace element study of peridotite
xenoliths from Savai‘i, Western Samoa, J. Geophys. Res., 99,
24,301–24,321, doi:10.1029/94JB01553.
Hauri, E. H., N. Shimizu, J. J. Dieu, and S. R. Hart (1993),
Evidence for hotspot-related carbonatite metasomatism in the
oceanic upper mantle, Nature, 365, 221–227.
Hawkins, J. W., and J. H. Natland (1975), Nephelinites and
basanites of the Samoan linear volcanic chain: Their possible
tectonic significance, Earth Planet. Sci. Lett., 24, 427–439,
doi:10.1016/0012-821X(75)90150-8.
Herzberg, C., and P. D. Asimow (2008), Petrology of some
oceanic island basalts: PRIMELT2.XLS software for primary
magma calculation, Geochem. Geophys. Geosyst., 9, Q09001,
doi:10.1029/2008GC002057.
Herzberg, C., and E. Gazel (2009), Petrological evidence for
secular cooling in mantle plumes, Nature, 458, 619–622,
doi:10.1038/nature07857.
Herzberg, C., and M. J. O’Hara (2002), Plume-associated
ultramafic magmas of Phanerozoic age, J. Petrol., 43,
1857–1883, doi:10.1093/petrology/43.10.1857.
Hirschmann, M. M., and R. A. Stolper (1996), A possible role
for garnet pyroxenite in the origin of the ‘garnet signature’ in
MORB, Contrib. Mineral. Petrol., 124, 185–208, doi:10.1007/
s004100050184.
Hoernle, K., and H.-U. Schmincke (1993), The petrology of
the Tholeiites through melilite nephelinites on Gran Canaria,
Canary Islands: Crystal fractionation, accumulation, and
depths of melting, J. Petrol., 34, 573–597.
Hofmann, A. W., and W. M. White (1982), Mantle plumes
from ancient oceanic crust, Earth Planet. Sci. Lett., 57,
421–436, doi:10.1016/0012-821X(82)90161-3.
Honda, M., and J. D. Woodhead (2005), A primordial solarneon enriched component in the 635 source of EM-I-type
ocean island basalts from the Pitcairn Seamounts. Polynesia,
Earth Planet. Sci. Lett., 236, 597–612, doi:10.1016/j.epsl.
2005.05.038.
Ielsch, G., M. Caroff, H. G. Barsczus, R. C. Maury, H. Guillou,
G. Guille, and J. Cotton (1998), Géochimie des basaltes de
l’île de Ua Huka (archipel des Marquises): Variation du taux
de fusion partielle et hétérogénéité de la source mantellique,
C. R. Acad. Sci., Ser. II Sci. Terre Planets, 326, 413–420.
Jackson, M. G., and S. Shirey (2011), Re-Os systematics in
Samoan shield lavas and the use of Os-isotopes in olivine
phenocrysts to determine primary magmatic compositions,
Earth Planet. Sci. Lett., 312, 91–101, doi:10.1016/j.epsl.
2011.09.046.
Jackson, M. G., S. R. Hart, A. A. P. Koppers, H. Staudigel,
J. Konter, J. Blusztajn, M. D. Kurz, and J. A. Russell
(2007), The return of subducted continental crust in Samoan
lavas, Nature, 448, 684–687, doi:10.1038/nature06048.
Jackson, M. G., M. D. Kurz, and S. R. Hart (2009), Helium and
neon isotopes in phenocrysts from Samoan lavas: Evidence
for heterogeneity in the terrestrial high 3He/4He mantle,
Earth Planet. Sci. Lett., 287, 519–528, doi:10.1016/j.
epsl.2009.08.039.
Jackson, M. G., S. R. Hart, J. G. Konter, A. A. P. Koppers,
H. Staudigel, M. D. Kurz, J. Blusztajn, and J. M. Sinton
10.1029/2011GC003974
(2010a), Samoan hot spot track on a “hot spot highway”:
Implications for mantle plumes and a deep Samoan mantle
source, Geochem. Geophys. Geosyst., 11, Q12009, doi:10.1029/
2010GC003232.
Jackson, M. G., E. Hauri, and S. Shirey (2010b), Re-Os reveals
ancient mantle components in xenoliths from the CookAustrals and Samoan hotspots, Geochim. Cosmochim. Acta,
74, suppl. 1, A451.
Johnson, D. M., P. R. Hooper, and R. M. Conrey (1999), XRF
Analysis of rocks and minerals for major and trace elements
on a single low dilution Li-tetraborate fused bead, Adv. X Ray
Anal., 41, 843–867.
Kear, D., and B. L. Wood (1959), The geology and hydrology
of Western Samoa, N. Z. Geol. Surv. Bull., 63, 1–90.
Köhler, T. P., and G. P. Brey (1990), Calcium exchange
between olivine and clinopyroxene calibrated as a geothermobarometer for natural peridotites from 2 to 60 kb with
applications, Geochim. Cosmochim. Acta, 54, 2375–2388,
doi:10.1016/0016-7037(90)90226-B.
Konter, J. G., H. Staudigel, S. R. Hart, and P. M. Shearer
(2004), Seafloor seismic monitoring of an active submarine
volcano: Local seismicity at Vailulu’u Seamount, Samoa,
Geochem. Geophys. Geosyst., 5, Q06007, doi:10.1029/
2004GC000702.
Konter, J. G., B. B. Hanan, J. Blichert-Toft, A. A. P. Koppers,
T. Plank, and H. Staudigel (2008), One hundred million
years of mantle geochemical history suggest the retiring of
mantle plumes is premature, Earth Planet. Sci. Lett., 275,
285–295, doi:10.1016/j.epsl.2008.08.023.
Konter, J. G., H. Staudigel, J. Blichert-Toft, B. B. Hanan,
M. Polvé, G. Davies, N. Shimizu, and P. Schiffman (2009),
Geochemical stages at Jasper Seamount and the origin of
intraplate volcanoes, Geochem. Geophys. Geosyst., 10,
Q02001, doi:10.1029/2008GC002236.
Konter, J. G., M. G. Jackson, and L. Storm (2010), Exceptional
volumes of rejuvenated volcanism in Samoa, Abstract U51A0017 presented at 2010 Fall Meeting, AGU, San Francisco,
Calif., 13–17 Dec.
Koppers, A. A. P., J. A. Russell, M. G. Jackson, J. Konter,
H. Staudigel, and S. R. Hart (2008), Samoa reinstated as a
primary hotspot trail, Geology, 36, 435–438, doi:10.1130/
G24630A.1.
Koppers, A. A. P., J. A. Russell, J. Roberts, M. G. Jackson,
J. G. Konter, D. J. Wright, H. Staudigel, and S. R. Hart
(2011), Age systematics of two young en echelon Samoan
volcanic trails, Geochem. Geophys. Geosyst., 12, Q07025,
doi:10.1029/2010GC003438.
Kurz, M. D., T. C. Kenna, J. C. Lassiter, and D. J. DePaolo
(1996), Helium isotopic evolution of Mauna Kea Volcano:
First results from the 1-km drill core, J. Geophys. Res.,
101, 11,781–11,791, doi:10.1029/95JB03345.
Lamb, M. P., A. D. Howard, W. E. Dietrich, and J. Taylor Perron
(2007), Formation of amphitheater-headed valleys by waterfall
erosion after large-scale slumping on Hawai‘i, Geol. Soc. Am.
Bull., 119, 805–822, doi:10.1130/B25986.1.
Langenheim, V. A. M., and D. A. Clague (1987), The
Hawaiian-Emperor volcanic chain, part I: Geological evolution, in Volcanism in Hawaii, edited by R. Decker et
al., U.S. Geol. Surv. Prof. Pap. 1350(1), 55–84.
Lanphere, M. A., and G. B. Dalrymple (1980), Age and strontium isotopic composition of the Honolulu volcanic series,
Oahu, Hawaii, Am. J. Sci., 280A, 736–751.
Lassiter, J. C., D. J. DePaolo, and M. Tatsumoto (1996), Isotopic evolution of Mauna Kea volcano: Results from the initial
21 of 23
Geochemistry
Geophysics
Geosystems
3
G
KONTER AND JACKSON: SAMOAN REJUVENATED LAVAS AND FLEXURE
phase of the Hawaiian Scientific Drilling Project, J. Geophys.
Res., 101, 11,769–11,780, doi:10.1029/96JB00181.
Lassiter, J. C., E. H. Hauri, P. W. Reiners, and M. O. Garcia
(2000), Generation of Hawaiian post-erosional lavas by melting of a mixed lherzolite/pyroxenite source, Earth Planet. Sci.
Lett., 178, 269–284, doi:10.1016/S0012-821X(00)00084-4.
Lee, C.-T., P. Luffi, T. Plank, H. Dalton, and W. P. Leeman
(2009), Constraints on the depths and temperatures of basaltic
magma generation on Earth and other terrestrial planets using
new thermobarometers for mafic magmas, Earth Planet. Sci.
Lett., 279, 20–33, doi:10.1016/j.epsl.2008.12.020.
Levitt, D., and D. Sandwell (1995), Lithospheric bending at
subduction zones based on depth soundings and satellite gravity, J. Geophys. Res., 100, 379–400, doi:10.1029/94JB02468.
Maaløe, S. (1998), Melt dynamics of a layered mantle plume
source, Contrib. Mineral. Petrol., 133, 83–95, doi:10.1007/
s004100050439.
Macdonald, G. A. (1968), A contribution to the petrology of
Tutui1a, American Samoa, Sond. Geol. Rundsch., 57, 821–837,
doi:10.1007/BF01845367.
Macdonald, G. A., A. T. Abbott, and F. L. Peterson (1983),
Volcanoes in the Sea: The Geology of Hawaii, Univ. of
Hawaii Press, Honolulu.
Marske, J. P., A. J. Pietruszka, D. Weis, M. O. Garcia, and
J. M. Rhodes (2007), Rapid passage of a small-scale heterogeneity through the melting regions of Kilauea and Mauna Loa
volcanoes, Hawaii, Earth Planet. Sci. Lett., 259, 34–50,
doi:10.1016/j.epsl.2007.04.026.
McDougall, I. (2010), Age of volcanism and its migration in
the Samoa Islands, Geol. Mag., 147(5), 705–717, doi:10.1017/
S0016756810000038.
Millen, D. W., and M. W. Hamburger (1998), Seismological
evidence for tearing of the Pacific plate at the northern termination of the Tonga subduction zone, Geology, 26, 659–662,
doi:10.1130/0091-7613(1998)026<0659:SEFTOT>2.3.CO;2.
Morgan, W. J. (1971), Convection plumes in the lower mantle,
Nature, 230, 42–43, doi:10.1038/230042a0.
Mukhopadhyay, S., J. C. Lassiter, K. A. Farley, and S. W.
Bogue (2003), Geochemistry of Kauai shield-stage lavas:
Implications for the chemical evolution of the Hawaiian
plume, Geochem. Geophys. Geosyst., 4(1), 1009, doi:10.1029/
2002GC000342.
Nakamura, Y., and M. Tatsumoto (1988), Pb, Nd and Sr isotopic
evidence for a multicomponent source for rocks of CookAustral Islands and heterogeneities of mantle plumes,
Geochim. Cosmochim. Acta, 52, 2909–2924, doi:10.1016/
0016-7037(88)90157-3.
Natland, J. H. (1980), The progression of volcanism in the
Samoan linear volcanic chain, Am. J. Sci., 280A, 709–735.
Natland, J. H., and D. L. Turner (1985), Age progression and
petrological development of Samoan Shield volcanoes: Evidence from K-Ar ages, lava compositions and mineral studies,
in Geological Investigations of the Northern Melanesian
Borderland, Earth Sci. Ser., vol 3, edited by T. M. Brocker,
pp. 139–172, Circum-Pac. Counc. for Energy and Miner.
Resour., Houston, Tex.
Németh, K., and S. J. Cronin (2009), Volcanic structures and oral
traditions of volcanism of Western Samoa (SW Pacific) and
their implications for hazard education, J. Volcanol. Geotherm.
Res., 186, 223–237, doi:10.1016/j.jvolgeores.2009.06.010.
Paul, D., W. M. White, and J. Blichert-Toft (2005), Geochemistry of Mauritius and the origin of rejuvenescent volcanism
on oceanic island volcanoes, Geochem. Geophys. Geosyst.,
6, Q06007, doi:10.1029/2004GC000883.
10.1029/2011GC003974
Phipps Morgan, J. (2001), Thermodynamics of pressure release
melting of a veined plum pudding mantle, Geochem. Geophys.
Geosyst., 2(4), 1001, doi:10.1029/2000GC000049.
Pilet, S., M. B. Baker, and E. M. Stolper (2008), Metasomatized lithosphere and the origin of alkaline lavas, Science,
320, 916–919.
Poreda, R. J., and K. A. Farley (1992), Rare gases in Samoan
xenoliths, Earth Planet. Sci. Lett., 113(1–2), 129–144,
doi:10.1016/0012-821X(92)90215-H.
Reiners, P. W., and B. K. Nelson (1998), Temporal-compositionalisotopic trends in rejuvenated-stage magmas of Kauai,
Hawaii, and implications for mantle melting processes,
Geochim. Cosmochim. Acta, 62, 2347–2368, doi:10.1016/
S0016-7037(98)00141-0.
Ren, Z.-Y., T. Hanyu, T. Miyazaki, Q. Chang, T. Takahashi,
Y. Hireahara, A. R. L. Nichols, and Y. Tatsumi (2009), Geochemical differences of the Hawaiian shield lavas: Implications for melting processes in the heterogeneous Hawaiian
plume, J. Petrol., 50, 1553–1573, doi:10.1093/petrology/
egp041.
Ribe, N. M., and U. R. Christensen (1999), The dynamical origin of Hawaiian volcanism, Earth Planet. Sci. Lett., 171,
517–531, doi:10.1016/S0012-821X(99)00179-X.
Robinson, J. E., and B. W. Eakins (2006), Calculated volumes
of individual shield volcanoes at the young end of the Hawaiian ridge, J. Volcanol. Geotherm. Res., 151, 309–317,
doi:10.1016/j.jvolgeores.2005.07.033.
Roden, M. F., F. A. Frey, and D. A. Clague (1994a), Geochemistry of tholeiitic and alkalic lavas from the Koolau range,
Oaho, Hawaii: Implications for Hawaiian volcanism,
Geochim. Cosmochim. Acta, 58, 1431–1440, doi:10.1016/
0016-7037(94)90547-9.
Roden, M. F., T. Trull, S. R. Hart, and F. A. Frey (1994b), New
He, Nd, Pb, and Sr isotopic constraints on the constitution of
the Hawaiian plume: Results from Koolau Volcano, Oahu,
Hawaii, USA, Geochim. Cosmochim. Acta, 58, 1431–1440,
doi:10.1016/0016-7037(94)90547-9.
Sapper, K. (1906), Der Matavanu-Ausbruch auf Savai’i 1905/
06. Nach Aufzeichnungen von Pater Mennel, Mittelungen
von Dr B. Funk und gedruckten Berichten dargestellt,
Z. Gesellschaft Erdkunde Berlin, 41, 686–709.
Schiano, P., K. W. Burton, B. Dupré, J.-L. Birck, J. Guille, and
C.-J. Allègre (2001), Correlated Os-Pb-Nd-Sr isotopes in the
Austral-Cook chain basalts: The nature of mantle components in plume sources, Earth Planet. Sci. Lett., 186, 527–537,
doi:10.1016/S0012-821X(01)00265-5.
Sherrod, D. R., Y. Nishimitsu, and T. Tagami (2003), New
K-Ar ages and the geologic evidence against rejuvenated
stage volcanism at Haleakala, East Maui, a postshieldstage volcano of the Hawaiian island chain, Geol. Soc. Am.
Bull., 115, 683–694, doi:10.1130/0016-7606(2003)115<0683:
NKAATG>2.0.CO;2.
Shirey, S., and R. J. Walker (1998), The Re-Os isotope system
in cosmochemistry and high-temperature geochemistry,
Annu. Rev. Earth Planet. Sci., 26, 423–500, doi:10.1146/
annurev.earth.26.1.423.
Sims, K. W. W., S. R. Hart, M. K. Reagan, J. Blusztajn,
H. Staudigel, R. A. Sohn, G. D. Layne, L. A. Ball, and
J. Andrews (2008), 238U-230Th-226Ra-210Pb-210Po, 232Th-228Ra,
and 235U-231Pa constraints on the ages and petrogenesis of
Vailulu’u and Malumalu Lavas, Samoa, Geochem. Geophys.
Geosyst., 9, Q04003, doi:10.1029/2007GC001651.
Smith, W. H. F., and D. T. Sandwell (1997), Global sea floor
topography from satellite altimetry and ship depth soundings,
Science, 277, 1956–1962, doi:10.1126/science.277.5334.1956.
22 of 23
Geochemistry
Geophysics
Geosystems
3
G
KONTER AND JACKSON: SAMOAN REJUVENATED LAVAS AND FLEXURE
Staudigel, H., A. Zindler, S. R. Hart, T. Leslie, C.-Y. Chen, and
D. Clague (1984), The isotopic systematics of a juvenile
intraplate volcano: Pb, Nd, and Sr isotope ratios of basalts
from Loihi Seamount, Hawaii, Earth Planet. Sci. Lett., 69,
13–29, doi:10.1016/0012-821X(84)90071-2.
Staudigel, H., et al. (2006), Vailulu’u Seamount, Samoa: Life
and death on an active submarine volcano, Proc. Natl. Acad.
Sci. U. S. A., 103, 6448–6453, doi:10.1073/pnas.0600830103.
Stearns, H. T. (1944), Geology of the Samoan Islands, Geol.
Soc. Am. Bull., 55, 1279–1332.
Stille, P., D. M. Unruh, and M. Tatsumoto (1983), Pb, Sr, Nd
and Hf isotopic evidence of multiple sources for Oahu,
Hawaii basalts, Nature, 304, 25–29, doi:10.1038/304025a0.
Stolper, E. M., D. J. DePaolo, and D. M. Thomas (1996), The
Hawaii Scientific Drilling Project: Introduction to the special
section, J. Geophys. Res., 101, 11,593–11,598, doi:10.1029/
96JB00332.
Tanaka, R., A. Makishima, and E. Nakamura (2008), Hawaiian
double volcanic chain triggered by an episodic involvement
of recycled material: Constraints from temporal Sr-Nd-HfPb isotopic trend of the Loa-type volcanoes, Earth Planet.
Sci. Lett., 265, 450–465, doi:10.1016/j.epsl.2007.10.035.
ten Brink, U. S., and T. M. Brocher (1987), Multichannel seismic evidence for a subcrustal intrusive complex under Oahu
and a model for Hawaiian volcanism, J. Geophys. Res., 92,
13,687–13,707, doi:10.1029/JB092iB13p13687.
Todt, W., R. Cliff, A. Hanser, and A. W. Hofmann (1996),
Evaluation of a 202Pb-205Pb double spike for high precision lead isotope analysis, in Earth Processes: Reading the
isotopic code, Geophys. Monogr. Ser., vol. 95, edited by
A. Basu and S. R. Hart, pp. 429–437, AGU, Washington,
D. C., doi:10.1029/GM095p0429.
Valbracht, P. J., H. Staudigel, M. Honda, I. McDougall, and
G. R. Davies (1996), Isotopic tracing of volcanic source
regions from Hawaii: Decoupling of gaseous from lithophile
magma components, Earth Planet. Sci. Lett., 144, 185–198,
doi:10.1016/0012-821X(96)00126-4.
Walter, M. J. (1998), Melting of garnet peridotite and the origin of komatiite and depleted lithosphere, J. Petrol., 39,
29–60, doi:10.1093/petroj/39.1.29.
Weiss, D., et al. (2006), High-precision isotopic characterization of USGS reference materials by TIMS and MC-ICPMS, Geochem. Geophys. Geosyst., 7, Q08006, doi:10.1029/
2006GC001283.
Wessel, P., and L. W. Kroenke (2008), Pacific absolute plate
motion since 145 Ma: An assessment of the fixed hot spot
hypothesis, J. Geophys. Res., 113, B06101, doi:10.1029/
2007JB005499.
White, W. M. (2010), Oceanic island basalts and mantle
plumes: The geochemical perspective, Annu. Rev. Earth
10.1029/2011GC003974
Planet. Sci., 38, 133–160, doi:10.1146/annurev-earth040809-152450.
White, W. M., and R. A. Duncan (1996), Geochemistry and
geochronology of the Society Islands: New evidence for deep
mantle recycling, in Earth Processes: Reading the Isotopic
Code, Geophys. Monogr. Ser., vol. 95, edited by A. Basu
and S. R. Hart, pp. 183–206, AGU, Washington, D. C.,
doi:10.1029/GM095p0183.
White, W. M., F. Albarède, and P. Telouk (2000), Highprecision analysis of Pb isotope ratios by multi-collector
ICP-MS, Chem. Geol., 167, 257–270, doi:10.1016/S00092541(99)00182-5.
Woodhead, J. D. (1992), Temporal geochemical evolution in
oceanic intra-plate volcanics: A case study from the Marquesas (French Polynesia) and comparison with other hotspots,
Contrib. Mineral. Petrol., 111, 458–467, doi:10.1007/
BF00320901.
Woodhead, J. D., and M. T. McCulloch (1989), Ancient seafloor signals in Pitcairn Island lavas and evidence for large
amplitude, small length-scale mantle heterogeneities, Earth
Planet. Sci. Lett., 94, 257–273, doi:10.1016/0012-821X(89)
90145-3.
Workman, R. K., and S. R. Hart (2005), Major and trace element composition of the depleted MORB mantle (DMM),
Earth Planet. Sci. Lett., 231, 53–72, doi:10.1016/j.epsl.
2004.12.005.
Workman, R. K., S. R. Hart, M. Jackson, M. Regelous, K. A.
Farley, J. Blusztajn, M. Kurz, and H. Staudigel (2004),
Recycled metasomatized lithosphere as the origin of the
Enriched Mantle II (EM2) end-member: Evidence from the
Samoan Volcanic Chain, Geochem. Geophys. Geosyst., 5,
Q04008, doi:10.1029/2003GC000623.
Wright, E., and W. M. White (1987), The origin of Samoa:
New evidence from Sr, Nd, and Pb isotopes, Earth Planet.
Sci. Lett., 81, 151–162, doi:10.1016/0012-821X(87)90152-X.
Xu, G., F. A. Frey, D. A. Clague, W. Abouchami, J. BlichertToft, B. Cousens, and M. Weisler (2007), Geochemical characteristics of West Molokai shield- and postshield-stage
lavas: Constraints on Hawaiian plume models, Geochem.
Geophys. Geosyst., 8, Q08G21, doi:10.1029/2006GC001554.
Yang, H.-J., F. A. Frey, and D. A. Clague (2003), Constraints
on the source components of lavas forming the Hawaiian
North Arch and Honolulu Volcanics, J. Petrol., 44, 603–627,
doi:10.1093/petrology/44.4.603.
Zindler, A., and S. Hart (1986), Chemical geodynamics, Annu.
Rev. Earth Planet. Sci., 14, 493–571, doi:10.1146/annurev.
ea.14.050186.002425.
23 of 23