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Article
Volume 6, Number 1
16 February 2005
Q02005, doi:10.1029/2004GC000723
AN ELECTRONIC JOURNAL OF THE EARTH SCIENCES
Published by AGU and the Geochemical Society
ISSN: 1525-2027
A Cenozoic diffuse alkaline magmatic province (DAMP) in
the southwest Pacific without rift or plume origin
Carol A. Finn
U.S. Geological Survey, Denver Federal Center, MS 945, Denver, Colorado 80226, USA
R. Dietmar Müller
School of Geosciences and University of Sydney Institute of Marine Science, University of Sydney, Edgeworth David
Building F05, Sydney, New South Wales 2006, Australia
Kurt S. Panter
Department of Geology, Bowling Green State University, Bowling Green, Ohio 53503-0218, USA
([email protected])
[1] Common geological, geochemical, and geophysical characteristics of continental fragments of East
Gondwana and adjacent oceanic lithosphere define a long-lived, low-volume, diffuse alkaline magmatic
province (DAMP) encompassing the easternmost part of the Indo-Australian Plate, West Antarctica, and
the southwest portion of the Pacific Plate. A key to generating the Cenozoic magmatism is the combination
of metasomatized lithosphere underlain by mantle at only slightly elevated temperatures, in contrast to
large igneous provinces where mantle temperatures are presumed to be high. The SW Pacific DAMP
magmatism has been conjecturally linked to rifting, strike-slip faulting, mantle plumes, or hundreds of hot
spots, but all of these associations have flaws. We suggest instead that sudden detachment and sinking of
subducted slabs in the late Cretaceous induced Rayleigh-Taylor instabilities along the former Gondwana
margin that in turn triggered lateral and vertical flow of warm Pacific mantle. The interaction of the warm
mantle with metasomatized subcontinental lithosphere that characterizes much of the SW Pacific DAMP
concentrates magmatism along zones of weakness. The model may also provide a mechanism for warming
south Pacific mantle and resulting Cenozoic alkaline magmatism, where the oceanic areas are characterized
primarily, but not exclusively, by short-lived hot spot tracks not readily explained by conventional mantle
plume theory. This proposed south Pacific DAMP is much larger and longer-lived than previously
considered.
Components: 16,174 words, 10 figures, 1 table.
Keywords: geochemistry; geophysics; tectonics; alkaline magmatism; metasomatism.
Index Terms: 3040 Marine Geology and Geophysics: Plate tectonics (8150, 8155, 8157, 8158); 3099 Marine Geology and
Geophysics: General or miscellaneous; 1099 Geochemistry: General or miscellaneous.
Received 2 March 2004; Revised 5 November 2004; Accepted 22 November 2004; Published 16 February 2005.
Finn, C. A., R. D. Müller, and K. S. Panter (2005), A Cenozoic diffuse alkaline magmatic province (DAMP) in the southwest
Pacific without rift or plume origin, Geochem. Geophys. Geosyst., 6, Q02005, doi:10.1029/2004GC000723.
1. Introduction
[2] Cenozoic alkaline igneous rocks cover continental fragments of East Gondwana and adjacent
oceanic lithosphere in parts of Antarctica, eastern
Copyright 2005 by the American Geophysical Union
Australia, the Tasman Sea, New Zealand, and
the Antarctic plate extending east of the AustraliaAntarctic discordance, south of the PacificAntarctic Ridge and west of the Antarctic Peninsula
(Figure 1). This region is distinguished by dis1 of 26
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Figure 1. Topographic and bathymetric map of south Pacific [Sandwell and Smith, 1997]. The SW Pacific DAMP
study area is indicated by the thick black line. AP, Antarctic Peninsula; AT, Adare Trough; BI, Balleny Islands; BFZ,
Balleny Fracture Zone; CP, Campbell Plateau; CR, Chatham Rise; CS, Coral Sea; LHR, Lord Howe Rise; LR,
Louisville Ridge; LTK, Lau-Tonga-Kermadec trench; NFB, North Fiji Basin; NQ, Northern Queensland; MBL,
Marie Byrd Land; MI, Macquarie Island; P-DG, Peter I and De Gerlache Seamounts; RS, Ross Sea;
TAM, Transantarctic Mountains; TS, Tasman Sea; TZ, Tasmania; WA, West Antarctica; WV, Western Victoria.
tinctly low velocity upper mantle with variably
enriched geochemical signatures (Figure 2). The
igneous activity has been related to adiabatic
decompression melting due to rifting [Johnson,
1989; Tessensohn, 1994; Wörner, 1999] or strikeslip faulting [Rocchi et al., 2002a, 2003], large
mantle plumes [Behrendt, 1999; LeMasurier and
Landis, 1996], or numerous separate, small hot
spots [Gaina et al., 2000; Lanyon et al., 1993;
Sutherland, 1991]. As we will argue, all of these
models are flawed, whether the magmatism is
considered as separate or related events.
[3] Accumulating geological, geophysical and
geochemical data collected over the last several
years provide a foundation to revisit the issue of
the origin of the volcanism. Our approach toward
developing a general model for magmatism is to
synthesize a variety of the modern data sets,
which has not been previously done. In this
paper, we describe the geological, geochemical,
and geophysical characteristics of the crust and
mantle that we use to define a diffuse Cenozoic
alkaline magmatic province (DAMP). On the
basis of our synthesis and analysis of the limitations of existing models, we identify the key
combination of elements required to bring about
the regional alkaline magmatism in the largely
continental fragments of East Gondwana and
suggest an alternate model linked to late Cretaceous slab detachments to explain these characteristics. Finally, this model will be used to show
that the largely continental magmatism in the SW
Pacific DAMP may be part of a much broader
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Figure 2. Rayleigh wave 150s group velocity map (120 km depth) [Larson and Ekström, 2001]. Only long-lived
hot spots with long traces [Clouard and Bonneville, 2001; Gaina et al., 2000; Ritsema and Allen, 2003] underlain by
low-velocity perturbations in the upper mantle are shown. Abbreviations same as Figure 1 and AAD, AustralianAntarctic Discordance; LHR, Lord Howe Rise; RS, Ross Sea; TAM, Transantarctic Mountains, BH, Bellingshausen
Sea. The SW Pacific DAMP study area is indicated by the thick black line. This region also includes the expected low
velocities associated with the mid-ocean ridges. The white line locates the projection of the location of the postulated
130 Ma subducted plate in the mantle [Müller et al., 1993].
province encompassing much of the south
Pacific.
2. Characteristics of the Cenozoic
Alkaline Magmatism
2.1. Location of Alkaline Magmatism
[4] Cenozoic dominantly alkaline igneous rocks
cover large, but discontinuous portions of the
SW Pacific (Figure 1). The continental basement
to the Cenozoic alkaline magmatism formed as a
result of subduction processes and is composed of
Paleozoic-Mesozoic arc plutonic roots of magmatic
arcs, and accreted sedimentary and oceanic crust,
covered in part by Jurassic igneous rocks [e.g.,
Dalziel, 1992]. In East Australia, the exposed,
mainly mafic (alkaline and tholeiitic basalts),
igneous rocks extend 4400 km from offshore of
its northern coast south to Tasmania and Victoria
(Figure 1). The 100–300 km wide belt contains
scattered local volcanic centers whose estimated
thickness derived from gravity modeling averages
70 m, yielding a total upper crustal (<2 km)
volume of 0.02 – 0.07 10 6 km 3 (surface
area times estimated thickness) [Wellman and
McDougall, 1974]. In New Zealand, Cenozoic
intraplate volcanic rock is widely distributed along
1000 km of the coastlines in 3 distinct, 50–
100 km wide, mostly mafic provinces from the
northeastern part of the North Island, to Auckland,
and the southern portions of the South Island.
Compared to Australia, most of the predominantly
mafic volcanic centers are considerably smaller
with estimated volumes an order of magnitude
lower [Weaver et al., 1989]. The thinned continental
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crust of the Campbell Plateau region and seafloor
between Australia and Antarctica host scattered
alkaline volcanic islands [Johnson, 1989] and small
seamounts (Figure 1).
[5] Cenozoic alkaline igneous rocks crop out at the
edge and offshore of West Antarctica (Marie Byrd
Land seamounts and Peter I Island) and parts of
East Antarctica (e.g., Balleny Islands), on islands
in the Ross Sea and in the Transantarctic Mountains (Figure 1). In contrast to much of the rest of
the region, 20 volcanoes have summit elevations
greater than 2500 m, but thousands of small cinder
cones and flows are also observed [LeMasurier
and Thomson, 1990]. In the mostly ice-covered
regions of West Antarctica, only a few volcanic
edifices have been imaged in sub-ice bedrock
topography, none of them large (<2 km diameter
and 500 m elevation) and only one is inferred to be
active [Blankenship et al., 1993]. High-amplitude
and frequency magnetic anomalies are interpreted
as evidence for sub-ice mafic igneous rocks that
cover 50% of West Antarctica [Behrendt, 1999;
Damaske et al., 1994; Ferraccioli et al., 2002;
Luyendyk et al., 2003; Maslanyj and Storey,
1990; Pederson et al., 1981]. Joint modeling of
magnetic and gravity data over central west Antarctica shows that, with the exception of the large
volcanoes, the exposed igneous rocks cannot be
more than 1000 m thick, and are probably less than
600 m, leading to total volume estimates of 0.5–
1 106 km3 for much of the province [Behrendt et
al., 1994; Finn et al., 2001].
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[Weaver et al., 1989]. Intermittent Cenozoic
volcanism in the Campbell Plateau region began
40 Ma (Chatham Islands) and extends to 1 Ma
(Antipodes Islands) [Weaver et al., 1989]. In Antarctica, magmatism started at least 48 Ma in
the Transantarctic Mountains – Ross Sea region
[Tonarini et al., 1997], 37 Ma in Marie Byrd
Land [Rocchi et al., 2002b] and continues today
[LeMasurier and Thomson, 1990] (Figure 1).
Because much of the region is covered by ice,
age data are sparse, but indicate pulses of activity in
the Upper Oligocene-lower Miocene in the western
Ross Sea and Marie Byrd Land regions and the
last 10 Myr, with 20 volcanoes manifesting
Holocene activity [LeMasurier and Thomson,
1990; Rocchi et al., 2002a].
[7] Age dating of seamounts on the Pacific plate
[Crawford et al., 1997; Duncan and McDougall,
1989; Lanyon et al., 1993], inferences from stratigraphy [Behrendt et al., 1987], bathymetry, and
magnetic anomalies [Cande et al., 2000] suggest
that the seamounts are younger than the underlying
oceanic crust and therefore intraplate in nature.
Volcanism in the Tasman Sea occurred during
and after its opening between 90 and 52 Ma [Gaina
et al., 2000; Lanyon et al., 1993; McDougall and
Duncan, 1988]. Notable Cenozoic alkaline features
include the Balleny Islands (<10 Ma), Peter I
Island (<12 Ma) [LeMasurier and Thomson,
1990] and the De Gerlache seamounts (20 –
21 Ma) (Figure 1) [Gohl et al., 1997].
2.3. Tectonic Setting
2.2. Timing of Magmatism
[6] The earliest alkaline magmatism in Australia
started about 70 Ma, but most has been emplaced
episodically between 55–15 Ma and 5 Ma to
Recent with young (<13,000 years) volcanic
centers limited to NE Queensland and western
Victoria (Figure 1) [Johnson, 1989, and references
therein]. Intermittent activity over tens of millions
of years characterizes much of the region; for
example, western Victoria has been active at various times over the last 60 Myr [Johnson, 1989].
In New Zealand, Mid to Late Cretaceous alkaline
igneous rocks are exposed in several places on the
South Island and offshore (e.g., Chatham Islands).
Magmatism increased around 30 Ma, but most
centers are younger than 15 Ma with Recent
activity on the North Island (Figure 1) [Hoke et
al., 2000; Weaver et al., 1989]. As with Australia,
several centers exhibit pulses of activity over 10–
30 Myr periods, notably on the South Island
[8] Measurements and models of modern and
paleo-stress fields from east Australia show that
the continent has been under minor compression
since the Eocene, with little variation in intensity
[Dyksterhuis, 2002; Hillis et al., 1999; Reynolds et
al., 2002]. The change at 6 Ma from predominantly strike-slip to compressional tectonics led to
the cessation of volcanism in the South Island, NZ
[Walcott, 1998].
[9] Little is known about the current stress field of
Antarctica. Plate motion studies suggest that the
continent is under compression [LithgowBertelloni and Guynn, 2004; Wuming et al.,
1992] and that the last major regional extensional
event was the Late Cretaceous break-up of Gondwana [Lawver and Gahagan, 1994]. GPS measurements collected in Marie Byrd Land, indicate
no significant motion between East and West
Antarctica [Donnellan, 2003]. The lack of diffuse
seismicity indicates that most of West Antarctica is
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no longer actively extending [Winberry and
Anandakrishnan, 2003]. In addition, gravity lows
and low-velocity layers that typically reflect extensional, young sedimentary basins are sparse
and small (<40 km wide) in West Antarctica
[Bell et al., 1998; Luyendyk et al., 2003;
Studinger et al., 2001; ten Brink et al., 1993].
The lower crust of much of eastern Australia and
central West Antarctica, is electrically resistive,
not conductive as observed in active rifts
[Tammemagi and Lilley, 1971; Wannamaker et
al., 1996]. Lack of geologic evidence for
large Cenozoic faults does not support largemagnitude, regional extension in much of the
region [Siddoway et al., 2003; Walcott, 1998].
However, marine magnetic anomalies in the
Adare Trough [Cande et al., 2000] and stratigraphy and fault orientations of drill core in the
western Ross Sea (Figure 1) identify contemporaneous Oligocene-mid-Miocene extension [Cape
Roberts Science Team, 1998, 1999, 2000]. Along
the Transantarctic Mountains front and offshore,
faults and fractures identify a stress regime compatible with regional Late Cenozoic dextral transtension [Rocchi et al., 2002a, 2003; Wilson, 1995]
and GPS measurements indicate that the western
Ross Sea is slowly extending today [Willis et al.,
2004]. In addition, Cenozoic 40Ar/39Ar ages for
pseudotachylyte suggest coseismic fault activity in
North Victoria Land [Di Vincenzo et al., 2004].
2.4. Heat Flow
[10] Averaged heat flow values in eastern Australia
over the Recent volcanic fields of Queensland
and Victoria, Tasmania, and New Zealand are
90 mW/m2, but near the 60 mW/m2 continental
average elsewhere [Cook et al., 1999; Cull, 1991;
Godfrey et al., 2001; Pandey et al., 1981; Sass and
Lachenbruch, 1979]. Sparse measurements in the
western Ross Sea region [Blackman et al., 1987;
Decker and Bucher, 1982; Della Vedova et al.,
1991; Kyle, 1990] range from 60–100 mW/m2; a
single measurement of 75 mW/m2 was obtained in
central West Antarctica [Gow et al., 1968].
[11] Thermobarometric analyses of crustal xenoliths from the western Ross Sea [Berg et al.,
1989] and eastern Australia [O’Reilly and Griffin,
1985] reveal elevated temperature gradients in
the crust (Figure 3). The SE Australia geotherm
crosses the McMurdo geotherm, giving somewhat lower temperatures in the crust and higher
temperatures in the upper mantle, but this may
be due to lack of data (dashed lines, Figure 3)
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rather than real differences [Berg et al., 1989].
Both geotherms have been explained by heat
transport by mafic magmatic intrusion at the
base of the crust [Berg et al., 1989; Cull et al.,
1991; O’Reilly and Griffin, 1985; Sass and
Lachenbruch, 1979].
2.5. Seismic Velocity Anomalies
[ 12 ] Inversion of Rayleigh and Love waves
referenced to models of the average velocity structure for the crust and mantle yielding 3-D shear
wave velocity perturbation models for the upper
[Bannister et al., 2000; Debayle and Kennett,
2000b; Larson and Ekström, 2001; Ritzwoller et
al., 2001; Shapiro and Ritzwoller, 2002; Simons et
al., 1999] and whole mantle [Ritsema et al., 1999]
provide information on mantle velocity structure
and, indirectly, temperature and chemical variations including volatile and melt content. Although
resolution of models is difficult to determine,
estimates range from horizontal resolution of
250 and vertical resolution of 50–100 km in
Australia and to 600 km horizontal and 50–
200 km vertical resolution for much of the
south Pacific upper mantle including Antarctica
[Ritzwoller et al., 2001]. Velocity perturbation
values for the upper mantle generally range from
±6% and ±1.5% in the lower mantle. The
resolution for the lower mantle is 1000 km
laterally and 100 – 200 km vertically at
>1000 km depth with the poorest resolution
(>250–300 km vertical resolution) in the transition
zone (500–1000 km depth) [Ritsema et al., 1999].
Areas with few earthquakes and seismic stations
like Antarctica will have lower resolution.
[13] Consistently, the regions of alkaline magmatism are characterized by slow velocity anomalies,
such as the velocity perturbations exceeding 2%
from eastern Australia to New Zealand and West
Antarctica (Figure 2). The low velocities are generally restricted to a zone between 60 and 200 km
depth (Figure 4) [Bannister et al., 2000; Debayle
and Kennett, 2000b; Larson and Ekström, 2001;
Ritzwoller et al., 2001; Simons et al., 1999] except
beneath the Tasman Sea (Figure 5a) and South
Pacific Ocean (Figure 5b) where low-velocity
perturbations (>0.4%) extend to 670 km and
800 km depths, respectively [Montelli et al.,
2004; Ritsema et al., 1999]. However, the near
constant velocity variations between 300 and
800 km depths might be due to limitations of the
model resolution at this depth range such that the
actual depth is poorly determined [Ritsema and
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Figure 3. Pressure versus temperature diagram for McMurdo, Antarctica [Berg et al., 1989], SE Australia
[O’Reilly and Griffin, 1986] to stability fields of amphibole and phlogopite [Class and Goldstein, 1997; Green
and Falloon, 1998]. Also shown are water-saturated and water-undersaturated solidi and an adiabatic path for
asthenospheric mantle. Orange areas outline incipient melt zone; blue regions denote major melt regimes [Green
and Falloon, 1998].
Allen, 2003]. Although low seismic velocity
anomalies are due to a largely undifferentiable
combination of temperature and chemical variations, elevated temperatures in the asthenosphere
are their most common explanation in the region
[Bannister et al., 2000; Debayle and Kennett,
2000b; Ritzwoller et al., 2001; Simons et al.,
1999].
[14] The low-velocity zones terminate in the west
and south at the boundaries with thick, highvelocity cratons in Australia and Antarctica, and
the 130-Myr subducted Pacific slab [Gurnis et al.,
1998] that divides distinct Indian and Pacific
geochemical reservoirs [Klein et al., 1988; Pyle
et al., 1992] at the Australia-Antarctic Discordance
(Figures 2, 4, and 5). In the north and northeast, the
boundaries of the low-velocity zone coincide with
high-velocity perturbations (>0.9%) of the subducting Pacific plates beneath the Lau-TongaKermadec trenches (Figure 5a), and old oceanic
lithosphere (e.g., east of the Tonga-Kermadec
trench) (Figures 2, 4b, and 4c). High-velocity
Precambrian craton [Maslanyj and Storey, 1990]
terminates low-velocity mantle beneath Marie
Byrd Land, as does the continuation of the
130-Myr slab east of the Antarctic Peninsula
[Gurnis et al., 1998] that coincides with the boundary of Pacific and Atlantic mantle determined by
geochemical tracing [Pearce et al., 2001] and
seismic wave anisotropy studies [Helffrich et
al., 1999]. Beneath the region, lower mantle
high seismic velocities (>0.6%) image detached
Mesozoic subducted slabs (Figure 5) [Gurnis et
al., 2000].
[15] Magmatism is also generally absent from areas
characterized by laterally extensive (>100 km),
high-velocities (perturbations >2% [Ritzwoller
et al., 2001]) in the upper 80 km (Figure 4) do
not typically host Cenozoic volcanic rocks. Examples include areas of old (>100 Ma), thick (60–
80 km) lithosphere like the area offshore Marie
Byrd Land north of 70S and east of 230 and
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Figure 4
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east of the Campbell Plateau (compare Figure 1 to
Figures 4b and 4c.).
2.6. Geochemical Signature
[16] For most of the region, alkaline magmatism is
a result of low degrees of melting (1–3%) of a
source enriched in incompatible elements relative
to primitive upper mantle. Some mantle sources
have been metasomatically enriched. Indeed, it has
long been proposed that sources for primary silicaundersaturated alkaline magmas contain volatiles
(H2O ± CO2) [Green and Falloon, 1998] and that
metasomatic enrichment of the upper mantle may
be a necessary precursor to alkaline magmatism
[Best and Christiansen, 2001; Sun and Hanson,
1976; Wass and Rogers, 1980]. The volatiles are
stabilized in the hydrous phases of phlogopite and
pargasitic or kaersutitic amphibole, which can exist
at pressures >3GPa [Mengel and Green, 1986;
Wallace and Green, 1991]. Their presence in the
residual source region of alkaline magmas can
be assessed using minor and trace element data
[Beswick, 1976; Class and Goldstein, 1997;
Greenough, 1988; Späth et al., 2001] in particular
Rb, Ba and K, which are retained in hydrous
relative to anhydrous minerals during melting
[Dalpé and Baker, 1994; LaTourrette et al.,
1995]. If melting does not consume all of the
hydrous minerals then the liquid extracted from
the source will be low in Rb, Ba and K relative to
other elements (e.g., LREE, Th, U) that partition
more readily into the melt phase.
[17] Negative K anomalies on mantle-normalized
multielement plots (Figure 6) demonstrate the
retention of hydrous minerals in the source. While
low relative K contents of alkaline rocks may be an
artifact of low source concentrations, modeling of
trace element data from West Antarctic basalts
predict sources with K concentrations 1–3 primitive upper mantle values and bulk partition coefficients consistent with the presence of residual
hydrous potassic phases [Hart et al., 1997; Rocchi
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et al., 2002a]. This is also supported by melting
experiments [Orlando et al., 2000] and the occurrence of metasomatized, amphibole- and phlogopite-bearing upper mantle xenoliths in alkaline
rocks from West Antarctica, SE Australia and
Southern New Zealand [Gamble et al., 1988;
O’Reilly et al., 1989]. The origin of alkaline
magmas in New Zealand and Australia have also
been linked to sources with residual amphibole or
phlogopite [Gamble et al., 1986; Panter et al.,
2000; Zhang and O’Reilly, 1997] and provincewide models for a metasomatic source for alkaline
volcanism have been proposed [O’Reilly, 1987;
Panter et al., 2000; Sun and McDonough, 1989].
[18] Variations in Sr, Nd and Pb isotopes of SW
Pacific basalts have been explained by mixing of
HIMU mantle (high 238U/204Pb sources that produce high time-integrated 206Pb/204Pb signatures)
with depleted sources (e.g., sources for mid-ocean
ridge basalts, MORB) and enriched sources (high
87
Sr/86Sr and low 143Nd/144Nd) such as oceanic
mantle EM1 and EM2, and subcontinental lithospheric mantle (SCLM). The depleted mantle
source for volcanism in all but the northern-most
region exhibits Pacific MORB mantle isotopic
fingerprints [Klein et al., 1988; Pearce et al.,
2001; Pyle et al., 1992; Zhang et al., 1999].
Beneath North Queensland [Zhang et al., 1999]
and the Lau and North Fiji basins [Hickey-Vargas
et al., 1995], Indian Ocean MORB-type mantle has
partially displaced Pacific Ocean MORB-type
mantle over the past 10 Myr; the present boundary of the two distinct mantle domains is in the
Vanuatu-Fiji-Tonga region (near NFB, Figure 2)
[Crawford et al., 1995]. In West Antarctica, Hart et
al. [1997] and Panter et al. [2000] suggest that the
depleted source for volcanism is not representative
of MORB but is similar to sources for the Balleny
and Scott Islands, and possibly Macquarie Island
[Kamenetsky et al., 2000]. The Balleny source is
very close to the oceanic FOZO (‘‘focus zone’’)
mantle end-member as defined in 3-D isotope
Figure 4. Shear velocity anomalies from a global three-dimensional diffraction tomographic model created from
inversion of surface wave fundamental model phases and group velocity measurements [Shapiro and Ritzwoller,
2002]. (a) Model slice at 150 km depth showing location of profiles. (b) Section at 34S. Gray line represents the
thickness of the lithosphere derived from geoid and flexural models [Zhang et al., 1998]. Thick dashed black line
represents the maximum lithospheric thickness based on seismic anisotropy [Debayle and Kennett, 2000a] and
magnetotelluric [Simpson, 2002] data. White circles indicate earthquake locations. (c) Section at 50S. Dashed gray
line represents the lithospheric thickness derived from flexural models [Godfrey et al., 2001]. (d) Section at 80S.
Thick gray line represents the lithospheric thickness derived from seismic data [Winberry and Anandakrishnan,
2003]. EWM, Ellsworth-Whitmore Mountains. White boxes outline source regions of melt derived from geochemical
modeling [Hart et al., 1997; Hoke et al., 2000; O’Reilly and Zhang, 1995; Panter et al., 2000]. Thin black outlines in
Figures 4b– 4d indicate features persistent in all models.
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Figure 5. Two 180-wide cross sections of shear velocity anomalies from model S20RTS [Ritsema et al., 1999].
The thick dashed line indicates the 670-km discontinuity. (a) Section from the Southeast Indian Ridge, across the
Tasman Sea, and Kermadec trench. (b) Section from the Indian Ridge, East Antarctic craton, South Pacific Ocean,
Pacific-Antarctic Ridge and central Pacific. (c) Model slice at 2850 km depth [Ritsema et al., 1999]. Circles indicate
locations of hot spots.
space (87Sr/86Sr-143Nd/144Nd-206Pb/204Pb) by Hart
et al. [1992] (Figure 7). This FOZO-like component is found in both continental and oceanic
basalts and appears to be a large-scale geochemical
feature in the uppermost mantle of the SW Pacific
DAMP. The enriched isotopic signatures (Figure 7)
are limited to continental basalts (including continental fragments such as the Lord Howe Rise and
Campbell Plateau [Weaver et al., 1994]) and have
been attributed to EM1+ EM2 [Hart et al., 1995,
1997] as well as ancient metasomatized SCLM
[Ewart et al., 1988; Foden et al., 2002; O’Reilly,
1987; Rocholl et al., 1995; Wörner, 1999].
[19] The HIMU signature is strongest in continental basalts from Marie Byrd Land and Southern
New Zealand, approaching 206Pb/204Pb values of
21 (Figures 7 and 8). Using ocean island basalts as
a proxy, HIMU sources in the SW Pacific have
been linked to mantle plumes [Behrendt, 1999;
LeMasurier and Landis, 1996]. It has also been
suggested that an upwelling plume(s) played a role
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intermittence, broad regional extent, low magma
production rates, metasomatized and enriched
sources for volcanism that coincide with unusual
upper mantle low-velocity zones and moderate
to high-velocity lower mantle. Key issues for
understanding the origin of magmatism include
determining the timing of metasomatism, locating
the metasomatized portion of the melt source (in the
SCLM, the asthenosphere, or both), and defining
the roles of lithospheric architecture (thickness and
faults) and stress.
Figure 6. Primitive mantle normalized [McDonough
and Sun, 1995] multielement diagram comparing mean
values of continental alkaline basalts from West
Antarctica [Hart et al., 1997; Panter et al., 1997b,
2000, 2003; Rocchi et al., 2002a; Rocholl et al., 1995],
Australia and Tasmania [Ewart et al., 1988; McBride et
al., 2001; McDonough et al., 1985], and southern New
Zealand [Baker et al., 1994; Panter et al., 1997a]. Also
shown are oceanic samples that represent average
HIMU end-member compositions [Hauri and Hart,
1993, 1997; Woodhead, 1996], MORB from the PacificAntarctic ocean ridge system [Ferguson and Klein,
1993], and three ‘‘near primitive’’ glasses from
Macquarie Island, an uplifted block of oceanic crust at
the Australia Pacific plate boundary south of New
Zealand [Kamenetsky et al., 2000]. Macquarie samples
and all other basalts used to calculate mean values have
MgO concentrations greater than 7 wt.%.
in the Mid to Late Cretaceous breakup of the protoPacific margin of Gondwana [Lanyon et al., 1993;
Storey et al., 1999; Weaver et al., 1994]. Variants
on the plume model call upon an ancient ‘‘fossilized’’ source that was emplaced and frozen to the
base of the SCLM prior to Gondwanaland breakup,
either in the Cretaceous or Jurassic [Panter et al.,
2000; Rocholl et al., 1995].
[20] The mixtures of distinctly different isotopic
end-members by small volume alkaline magmas
imply mantle heterogeneity on a fine scale
[Meibom et al., 2002]. The broad region over
which these magmas were erupted also implies that
this heterogeneity is a regional feature in the upper
mantle beneath the Pacific [Lassiter et al., 2003;
Staudigel et al., 1991; Workman et al., 2004].
3. Key Characteristics of the Cenozoic
SW Pacific Diffuse Alkaline Magmatism
[21] Striking features of SW Pacific Cenozoic
alkaline magmatism are its longevity (50 Myr),
[22] The cause of the metasomatic enrichment is
unclear. Several authors have proposed volatile
flux from plume derived melts [Panter et al.,
2000; Hart et al., 1997]. But another potential
source is metasomatic fluids (hydrous fluids and
volatile-rich silicate melts) derived from the prolonged subduction in the Paleozoic-Mesozoic
(500 – 100 Ma) along the Pacific margin of
Gondwana. It has been a long-held view that
the melting of subduction-related metasomatized
sources will yield high La, Ba, Rb/Nb-Ta ratios as
observed in island and continental arc magmas.
Alkaline rocks in the SW Pacific show OIB trace
element signatures with high relative abundances
of Nb and Ta, indicating the lack of the classic
subduction component (Figure 6). But if subduction-modified SCLM exists beneath much of the
SW Pacific, then why is it not being tapped by the
volcanism? One explanation would be that melting
does not occur within the lithosphere; however,
this is contrary to the geochemical and geophysical
evidence. A second explanation is that only SCLM
that has been altered by plume-derived melts and
fluids is being consumed, but this would require
spatially discrete metasomatic domains. A third
possibility is that the alkaline melts may be generated, in part, within a region of the lithosphere that
has been metasomatized by subduction but does
not contain what would be regarded as a typical arc
signature [e.g., Petrone et al., 2003].
[23] On the basis of the concept of chromatographic
separation of trace elements in mantle environments, Ionov and Hofmann [1995], Navon and
Stolper [1987], and Stein et al. [1997] have developed subduction-related metasomatic models to
explain the retention of Nb (and Ta) in hydrous
minerals (amphibole) within the lowermost portion
of the mantle wedge above a dehydrating slab.
Further development of the model helped explain
geochemical characteristics of the source regions
for alkaline rocks from the northern Arabian- Nubian Shield [Stein et al., 1997]. In their model,
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Figure 7
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fluid-rock interactions retain Nb while Pb and Rb
are lost relative to U, Th, Sr and REE during the
preservation of the wedge. The process leads to
the development of low Th-U/Nb and Rb/Sr
ratios and high U/Pb ratios compatible with
sources for alkaline magmas. Isolation of this type
of geochemical domain over long periods of time
may produce isotopic signatures that are similar
to sources for HIMU-type ocean island basalts
(high 206Pb/204Pb and low 87Sr/86Sr). Although a
detailed evaluation of subduction-related versus
plume-related metasomatism for the SW Pacific
is beyond the scope of this paper, it is of interest
to note that source regions in the SW Pacific
with the longest history of subduction (Marie
Byrd Land, Campbell Plateau, Chatham Rise,
and southern New Zealand) correspond to the
highest 206Pb/204Pb values (Figure 8). It has also
been proposed that the impact of a Mid Cretaceous plume head beneath this same region can
explain the position of Antarctic-New Zealand
rifting and the distribution of the HIMU source
[Storey et al., 1999; Weaver et al., 1994]. However, the model has been criticized because of
the brief time interval (10–15 Myr) between
the cessation of subduction-zone magmatism and
magmatism related to the inception of a mantle
plume [Dalziel et al., 2000; Mukasa and Dalziel,
2000], leaving subduction as the likely origin of
metasomatism and enrichment.
Figure 8. Plate reconstructions of Gondwana; colors
indicate age of the ocean floor [Müller et al., 1993].
Keys to abbreviations in Figure 1. (a) A 100 Ma
reconstruction overlain by locations and contours for
alkaline samples with measured 206Pb/204Pb ratios (data
from sources in Figure 7). (b) A 50 Ma reconstruction,
approximate timing of onset of alkaline magmatism.
(c) A 15 Ma reconstruction.
[ 24 ] Evidence for metasomatized sources for
97 Ma mafic rocks in southern New Zealand
[Baker et al., 1998] and 80 Ma volcanics
on the Campbell Plateau [Weaver et al., 1994;
Panter et al., 1997a], 440 Ma model ages
from metasomatized ultramafic xenoliths from
the Ross Sea region [McGibbon, 1991] and
Neoproterozoic Nd model ages from basalts from
Australia [Zhang and O’Reilly, 1997] and 500–
300 Ma ages from metasomatized xenoliths
[Griffin et al., 1988] from southeastern Australia
suggest that the enrichment had occurred by the
Mid-Cretaceous. No evidence exists for metasomatism concurrent with Cenozoic magmatism
Figure 7. Values of 206Pb/204Pb versus (a) 207Pb/204Pb, (b) 208Pb/204Pb, and (c) 87Sr/86Sr diagrams showing fields
representative of basalts from the southwest Pacific Ocean and continental areas. Data sources are the same as in
Figure 6 with additional samples from the Pacific-Antarctic Ridge [Vlastelic et al., 1999], Balleny and Scott Islands
[Hart et al., 1992, 1995; Hart and Kyle, 1994], Antarctic Peninsula [Hole et al., 1993], and Peter I Island [Hart et al.,
1995]. Also indicated are the approximate locations of possible source regions for magmatism: HIMU, FOZO (Focus
Zone [Hart et. al., 1992]), EM1, and EM2 [Workman et al., 2004]. Torlesse metasediments, South Island, New
Zealand [Graham et al., 1992] represent a potential mid-crustal contaminant. The Northern Hemisphere Reference
Line [Hart, 1984] and 4.5 Ga geochron are also shown for reference.
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[O’Reilly, 1987]. A rough estimate of the age
for sources in the SW Pacific can be made on
the basis of the correlation of data on 207Pb/
204
Pb– 206Pb/204Pb plots (Figure 7a). If we assume
a two-stage mantle evolution from bulk-earth, the
slope of the data arrays (±1s) yield pseudo-isochrons between 950 and 650 Ma. We believe the
model ages represent a mixture of sources, including metasomatized SCLM. A younger metasomatic
event may be responsible for higher U/Pb ratios
(either addition of U or depletion of Pb). If so, how
much time would be required to produce a HIMUlike signature? Again we use a simple two-stage
mantle evolution model. First, initial U/Pb and Pb
isotope ratios of a premetasomatized source were
calculated for any time in the past using a presentday 206Pb/204Pb ratio of 19.5. This value approximates that of a ubiquitous FOZO-like lower U/Pb
component in the SW Pacific mantle that is recognized by the convergence of data arrays in 3-D
isotope space [Hart et al., 1997; Panter et al.,
2000]. In Figure 7, the lower U/Pb source lies in a
region between the values for the Balleny and Scott
Islands, Macquarie Island and the FOZO mantle
reservoir. Second, the U/Pb ratio was adjusted to
evolve the lower U/Pb source to match what is
considered to be the HIMU end-member of SW
Pacific DAMP (206Pb/204Pb 21.0). Adjustments
to higher U/Pb ratios thus simulate fractionation
due to metasomatism. Our calculations suggest that
in-growth of HIMU-like values for Pb isotopes can
be obtained within 550 to 250 Myr, with U/Pb
ratios between 0.5 and 0.8 (or 238U/204Pb (m) 32–
53). The model U/Pb ratios are comparable with
values for metasomatized peridotites from SCLM
[Hawkesworth et al., 1986; Ionov and Hofmann,
1995; Lee et al., 1996]. While speculative the
calculations lend support, in conjunction with the
depth of low seismic velocities, evidence for residual hydrous phases and prior history of subduction, for a relatively young and shallow
metasomatic origin of the HIMU signature in the
SW Pacific.
[25] If a SCLM source modified by metasomatism
can account for the geochemical signatures in the
Gondwana fragments of SW Pacific, what is the
explanation for similar signatures in the oceanic
islands such as Balleny Islands, seamounts to the
north, Peter I Island and basalts that formed at
(Macquarie Island) [Kamenetsky et al., 2000] or
very near (e.g., seamounts east of Tasmania)
spreading centers (Figure 8c) [Crawford et al.,
1997; Lanyon et al., 1993]? Previous interpretations include contamination from the Balleny
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plume [Crawford et al., 1997; Lanyon et al.,
1993] whose origin is disputed [Gaina et al.,
2000] or a regional asthenospheric signature
[Kamenetsky et al., 2000]. Movement of the continental fragments during Gondwana breakup
could have delaminated part of weak SCLM and
smeared the geochemical signature throughout
much of the asthenosphere in the region. Also,
perhaps relatively close (500 km) continental
fragments contaminated the ridge magmas as oceanic crust was formed (Figures 8b and 8c). Therefore older (>10 Ma) regions such as the Tasman
Sea (adjacent to Australia and the Lord Howe, New
Zealand, and Campbell Plateau continental fragments), Macquarie Island and Ridge (near the
Campbell Plateau) and Balleny Islands, Peter I
Island and De Gerlache seamounts (near West
Antarctica) are contaminated, in contrast to oceanic
regions lacking enriched signatures which are far
from continental lithosphere (>500 km) such as
the Pacific-Antarctic Ridge (PAC, Figures 5 and 6;
crust <10 Ma, Figure 8c). A plate reconstruction
from the approximate onset of alkaline magmatism
at 50 Ma (Figure 8b) and at 15 Ma (Figure 8c)
illustrates the close proximity of enriched crustal
fragments and adjacent oceanic crust during much
of the magmatism. We conclude that the sources of
the magmas lie in SCLM below the continents, and
in asthenosphere contaminated by adjacent or
smeared SCLM beneath oceanic crust. The proposed melt and volatile sources, then, would dominate the low seismic velocity signature in the
SCLM (white boxes, Figures 3b–3d), rather than
temperature.
[26] Lithospheric thickness has implications for
locating sources of magmas. If the <80 km thickness of the high-velocity seismic lid for most of
SW Pacific approximates the lithosphere, the
melts, inferred from geochemical evidence to originate at 100–140 km depth, would lie in the
asthenosphere. Temperatures at the top of the
asthenosphere are often defined to be 1300C
[McKenzie and Bickle, 1988], which, for much of
SW Pacific, would occur at shallow levels (50–
80 km). In this case, the continental geotherms
would resemble the asthenosphere adiabat with
major melting and generation of tholeiites at
50 – 60 km (Figure 3). This is not observed
and argues against shallow asthenospheric melt
sources. Lithospheric thickness estimates of 100–
150 km based on elastic models and inferred
‘‘frozen-in’’ directions of seismic anisotropy from
Australia [Debayle and Kennett, 2000a; Simons et
al., 2003] and scattered broad-band seismometer
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data from West Antarctica [Winberry and Anandakrishnan, 2003], suggest also that part of the
observed low-velocity zones (<100 – 150 km
depth) (Figure 4) reflects melt/volatile sources
within the SCLM, not in the asthenosphere.
[27] Lithospheric architecture and stress influences
the localization of magmatism in a variety of
ways. Uncertainties related to thickness estimates
aside, the magmatism is restricted to continental
lithosphere no greater than 150 km thick and
oceanic lithosphere <80 km thick. The disconnected high-velocity anomalies of the continental
fragments (e.g., Lord Howe Rise (Figure 4b),
Tasman basin and Campbell Plateau (Figure 4c)
and West Antarctica (Figure 4d) contrast with the
more coherent anomalies associated with Mesozoic oceanic lithosphere (Figures 4b and 4c),
suggesting that the continental fragments have
been broken up and are therefore more susceptible
to melt incursion than more coherent oceanic
lithosphere of similar thickness. The coincidence
of volcanoes (Peter I Island, De Gerlache and
Marie Byrd Land seamounts, Figures 2 and 8c)
with major trans-lithospheric structures separating
thick, old (>50 Ma) from younger (<20 Ma)
oceanic lithosphere [Gohl et al., 1997] imply that
these structures are necessary to promote voluminous magmatism in thick lithosphere. In continental areas, most volcanoes are not localized by
large faults displaying significant Cenozoic motion [Siddoway et al., 2003; Wilson, 1995], with
the exception of large trans-lithospheric boundaries in the western Ross Sea and Marie Byrd
Land [Damaske et al., 1994; Kyle and Cole,
1974; LeMasurier and Rex, 1989; Luyendyk et
al., 2001].
[28] Although the magmatism has been linked
broadly to tensional stress fields, these are not
requirements as evidenced by the Australia volcanism which has occurred in mildly compressional stress fields for most of its history
[Dyksterhuis, 2005; Zhao and Müller, 2001].
Stronger compressional forces such as those associated with subduction and translation along the
Alpine fault, as started in New Zealand in the
Miocene inhibit volcanism [Hoke et al., 2000;
Walcott, 1998]. Extension in the western Ross Sea
coincides with heightened periods of magmatism
in the Oligocene [McIntosh, 2000], a relation that
continues today [Willis et al., 2004; Wilson,
2002]. Decompression melting due vertical flow
of asthenospheric from 100–150 km thick lithosphere in central West Antarctica and offshore to
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the thinner (75 km) lithosphere of Marie Byrd
Land may help explain extensive magmatism
there.
4. Possible Triggers for Cenozoic SW
Pacific Diffuse Alkaline Magmatism
[29] The triggers for magmatism include changing composition to lower melting temperature;
depressurization, usually by extension; and
increasing temperature. The melting regime
for mantle-derived basaltic magmas for the SW
Pacific alkaline rock types requires the presence
of volatiles such as carbon and hydrogen in the
melt phase, which reduces the solidus temperature (compare dehydration to dry solidus,
Figure 3). Magmas are derived from an ‘‘incipient melting’’ regime which lies at temperatures
below the dry solidus marking entry to the
‘‘major melting’’ regime [Green and Falloon,
1998] (Figure 3). Assuming low melting temperature SCLM underlies much of the region, we
evaluate various models for extension and heating triggers for volcanism.
4.1. Extension Triggers
[30] Can extension alone trigger incipient melting
in the region? A model for North Victoria Land
suggests that reactivation of preexisting translithospheric faults 43 Ma induced lithospheric
pull-apart and small-scale mantle convection at
the edge of the East Antarctic craton and triggered local decompression melting of mantle
enriched by veining associated with amagmatic
late Cretaceous rifting [Rocchi et al., 2002a].
However, as the lack of evidence for regional
extension indicates, this model would only account for magmatism in the western Ross Sea
region. On the basis of recent evidence from
drilling in the western Ross Sea [Cape Roberts
Science Team, 1998, 1999, 2000], rifting did not
occur amagmatically in the Cretaceous but contemporaneously with magmatism from Oligocenepresent. In addition, although mantle veining could
produce enriched trace element signatures (high U,
Th, Rb, Ba, and LREE/HREE ratios; Figure 6), it
would be difficult to explain the isotopic signatures
in particular, high 206Pb/204Pb ratios (>20.5), if the
enrichment was less than the 40 Ma initiation of
rifting. If metasomatism did occur before Gondwana breakup as suggested above, the rifting in
most of the SW Pacific should have triggered
magmatism. This is generally not observed, suggesting that extension alone in the Cenozoic would
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also not trigger magmatism and a heating event is
required.
4.2. Hot Spots, Hot Lines, and Mantle
Plumes
[31] Lack of evidence for significant regional extension during the Cenozoic, as well as the
enriched isotopic signatures of the alkaline rocks,
motivated models of increased temperatures caused
by regional mantle plumes, local hot spots and hot
lines with origins in both the upper and lower
mantle. Due to the recognition of their low magmatic volumes (<100,000 km3), and long duration
of activity (>50 Ma), Australia and New Zealand
have not been viewed as large volume flood basalt
provinces resulting from large, deep-seated mantle
plumes [Johnson, 1989]. Instead numerous small
diameter hot spots with both lower and upper
mantle origin, have been proposed to explain some
[Duncan and McDougall, 1989; Gaina et al.,
2000; Johnson, 1989; McDougall and Duncan,
1988; Sutherland, 1998; Wellman, 1983] if not all
[Sutherland, 1991, 1994] of the volcanism.
[32] Several volcanic chains do indeed fit the hot
spot reference frame and rates of movement of
Australia relative to Antarctica. Alignments of
tholeiitic, highly differentiated volcanic centers
contemporaneous with the alkalic volcanism
(<35 Ma) young progressively southward in eastern
Australia (colored triangles, Figure 1) and the
Tasman Sea, consistent with the separation rate of
the Australian and Antarctic plates, [Eggins et al.,
1991; McDougall and Duncan, 1988; McDougall
et al., 1981; Sutherland, 1991; Wellman, 1983;
Wellman and McDougall, 1974], and therefore
passage over the Bass hot spot [Gaina et al.,
2000] which has been imaged seismically (Figures
2 and 4a) [Montelli et al., 2004; Ritsema et al.,
2004]. However, other proposed hot spot tracks
such as in the Balleny Islands [Lanyon et al.,
1993], New Zealand, and offshore Antarctica
[Sutherland, 1991] do not fit hot spot models
(Table 1) [Gaina et al., 2000] and are not imaged
seismically (Figures 2, 4, and 5).
[33] In contrast to the rest of the region, West
Antarctica has been compared to flood basalt
provinces resulting from large, deep-seated mantle
plumes [Behrendt, 1999; LeMasurier, 1990]. Comparison of various attributes of flood basalt provinces often attributed to mantle plume activity
shows that West Antarctica and the SW Pacific
are not similar (Table 1). A key parameter
for comparison is the magma production rate.
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Although estimating magma production rates is
difficult due to erosion, underplating of unknown
amounts of material and paucity of age dates,
comparison with flood basalt provinces (similarly
computed) is revealing. In Australia and New
Zealand, averaging the estimated volume of surficial (upper 5 km) Cenozoic igneous rocks
over 30 –50 Myr yields a production rate of
.002 km3/Myr [Hoke et al., 2000; Wellman and
McDougall, 1974]. For the more voluminous
magmatism reported for Antarctica [Behrendt et
al., 1994; Finn et al., 2001], over its 35–50 Myr
history, a crude rate of 0.026 km3/Myr obtains. In
comparison, 600,000 km3 were erupted from
the Deccan traps in 5 Ma [Bhattacharji et al.,
1996] (magma production rate of 1 km3/yr);
1.3 106 km3 from the Siberian traps in 1 Ma
[Reichow et al., 2002] (magma production rate of
1.3 km3/yr); and from the relatively small Columbia
Plateau, 40,000 km3 in 1 Ma (magma production
rate of .2 km3/yr) [Swanson et al., 1975]. None of
these provinces are as areally extensive as SW
Pacific DAMP (or even West Antarctica alone).
[34] Assignment of other characteristics commonly
associated with mantle plumes to the SW Pacific is
also problematic (Table 1). In particular, with the
exception of local anomalies under the Tasman and
Ross Seas, low seismic velocities are generally
restricted to the upper 250 km of the mantle in
the region, in contrast to low-velocity zones under
East Africa [Nyblade et al., 2000], and the central
Pacific [Montelli et al., 2004; Ritsema and Allen,
2003] that extend to >400 km depth. The velocities
in the lower mantle beneath the region are relatively high (e.g., Figures 4a–4c), in contrast to
regions under the central Pacific (Figures 4a and
4c) and Africa [Montelli et al., 2004; Ritsema and
Allen, 2003].
[35] The origin of HIMU signatures is hotly debated [e.g., Anderson, 1995; Hoffmann, 1997]. HIMU
signatures have been attributed to recycling of
ancient (>200 Ma) oceanic crust within plumes
rising from the deep mantle [Hart et al., 1992;
Hoffmann, 1997] and, along with EM types, which
are attributed to recycling of ancient sediments,
considered to be diagnostic of lower mantle plumes
because of the inferred long periods of isolation
required to generate the isotopic signatures [Hart,
1984]. Another model suggests that enriched layers
formed by subduction recycling of sedimentary
rocks and oceanic lithosphere metasomatized by
hydrothermal activity at ridges resides in the upper
200 km of the mantle and provides the isotopic
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Table 1. Conventional Characteristics of Deep-Seated Mantle Plume/Shallow Hot Spot Versus SW Pacific DAMP
Flood Basalt/OIB Provinces
Dominantly tholeiitic volcanism
HIMU (lead isotope signatures >20.5)
[Hart, 1984]
Atmospheric normalized He3/He4 >12 Ra
[Lupton, 1983]
Dome (stationary plate), uplift
[Davies, 1988]
Linear age progression (moving plate)
[Morgan, 1971]
High heat flow (>75 mW/m2)
High magma production rates
(>1 km3/yr, neglecting underplating)
[Bhattacharji et al., 1996; Reichow et al., 2002]
Short-lived volcanism (1 – 5 Myr)
[Richards et al., 1989]
High (5 – 8%) degrees of melting
[Hart et al., 1992]
>1450C upper mantle temperatures
[McKenzie and Bickle, 1988]
Vertically extensive (>400 km)
low-velocity zone
Geoid high [Davies, 1988]
Dimensions 1000 1000 km2
[Davies, 1988]
signatures [Anderson, 1995]. Alternate models include mantle heterogeneities at a variety of scales
[Meibom and Anderson, 2004] or preservation
of enrichment in SCLM following mantle
metasomatism [Hawkesworth et al., 1986; O’Reilly
and Zhang, 1995]. The debate indicates that the
HIMU signatures are not sufficiently diagnostic to
uniquely define mantle plumes beneath a region.
4.3. Slab Detachment Model
[36] Unusual magmatism in modern subduction
[Levin et al., 2002; Wortel and Spakman, 2000]
and collision [Kosarev et al., 1999; Seber et al.,
1996] zones has been linked to detachment of
subducting slabs. Numerical models of mantle
convection suggest that slabs deflected horizontally
in the mantle transition zone are gravitationally
unstable [Christensen, 1997] and capable of triggering dramatic [Solheim and Peltier, 1994;
Tackley et al., 1993] or moderate [Davies, 1995;
SW Pacific DAMP
Dominantly alkaline [Johnson, 1989]
Only in Marie Byrd Land [Hart et al., 1997;
Panter et al., 2000] and S. New Zealand
[Panter et al., 1997a].
<8.5 Ra in NZ [Hoke et al., 2000] and
4 – 7 in western Ross Sea [Nardini et al., 2003].
Marie Byrd Land dome only [LeMasurier and Landis, 1996]
and existence disputed [Luyendyk et al., 2001].
Timing of uplift of eastern Highlands, Australia
not coincident with passage of hot spot [Johnson, 1989].
Only Bass, Louisville, and Tasman Sea hot
spots produces tholeiitic magmatism in eastern
Australia and offshore that fit in plate tectonic frame
of reference [Gaina et al., 2000].
90 – 120 mW/m2 [Blackman et al., 1987; Cull, 1982;
Della Vedova et al., 1991; Hoke et al., 2000;
Pandey et al., 1981; Purss and Cull, 2001].
Low magma production rates (averaged over 50 Myr)
<0.002 – 0.03 km3/yr) [Finn et al., 2001;
Hoke et al., 2000; Johnson, 1989;
Wellman and McDougall, 1974].
Long-lived (50 – 70 Myr)
[Johnson, 1989; Tonarini et al., 1997]
Low (1 – 3%) degrees of melting
[Hart et al., 1995; O’Reilly and Griffin, 1985]
1100 – 1300C upper mantle temperatures
[Hart et al., 1997; O’Reilly, 1987; Panter et al., 2000]
Low velocity zone generally restricted to
upper 200 km of in the upper mantle
[Debayle and Kennett, 2000b; Ritzwoller et al., 2001].
Geoid low for much of the region; high
due to subducting slabs in north
Minimum of 2000 7,000 km2
Zhong and Gurnis, 1994] episodes of mixing and
whole mantle flow in a primarily 2-layer system. Detachment of subducted slabs also generates vertical and lateral viscous upper mantle
flow resulting in magmatism [Pysklywec et al.,
2003]. Lateral flow of mantle resulting from
slab detachment [Kosarev et al., 1999] related
to continental collision in east Asia has been
linked to widespread Cenozoic alkaline volcanism [Flower et al., 1998] similar to that in the
SW Pacific.
[37] Mantle convection and subsidence models for
Australia suggest that eastward migration of the
Gondwana continent over the subducting slab may
have sheared it, causing detachment 130–90 Ma
[Gurnis et al., 1998]. Subduction continued off the
Antarctic portion of Gondwana until 100 Ma,
when the Phoenix plate may have been captured by
the north-moving Pacific plate, initiating separation
of New Zealand, and other continental fragments
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from Marie Byrd Land [Luyendyk, 1995]. Synthetic
3-dimensional density models of the mantle based
on plate convergence and a simple simulation of
whole mantle flow [Lithgow-Bertelloni and
Richards, 1998] as well as recent mantle convection
models [Steinberger et al., 2004] suggest that the
detached Pacific Gondwana slab was present in the
upper mantle for much of the region until 75–
65 Ma, when it sank into the lower mantle,
where pieces remains today (e.g., Figures 5b and
5c). This detachment could have occurred episodically as in the Mediterranean region [Wortel and
Spakman, 2000] with initial detachment off Australia, eventually migrating to West Antarctica
20–30 Myr later.
[38] A synopsis of the development of conditions
that led to alkaline magmatism in the SW Pacific
(Figure 9) suggests that subduction-related processes contribute to every phase. During the
Paleozoic-Mesozoic, east Gondwana lithosphere
formed largely by magmatism and accretion accompanied by volatile flux from the subducting
plate, and/or Jurassic plume activity that metasomatized the SCLM (Figure 9a). This same activity
imprinted at least part of the enriched geochemical
signature in the SCLM.
Figure 9. Cartoon depicting proposed model for
Cenozoic alkaline magmatism in SW Pacific DAMP
based on subduction history models [Lithgow-Bertelloni
and Richards, 1998] with mantle flow velocity vectors
generalized (and meant to be schematic, not quantitative) from a model of the North Fiji Basin [Pysklywec et
al., 2003]. (a) Paleozoic-Mesozoic subduction along the
Gondwana margin (Figure 8a) formed the lithosphere
and provided enrichment of the SCLM in the region.
Additional fluids may have been introduced during
postulated Jurassic mantle plume activity. (b) Late
Cretaceous detachment of the subducting slab along the
Gondwana margin. (c). Cenozoic alkaline magmatism
related to slab detachment and sinking into the lower
mantle (see Figure 5 for high-velocity anomalies that
have been related to detached slabs) that resulted in flow
of warm Pacific mantle into the SCLM, catalyzing
melting and producing the observed seismic lowvelocity zone (pink box) and geochemical signatures.
[39] Slab detachment in the late Cretaceous could
have induced a change in mantle flow. The migration paths of the mantle (arrows, Figures 9b and 9c)
mimic those from figures of geodynamic models of
the North Fiji Basin [Pysklywec et al., 2003] which
are used to guide discussion of the relation of
mantle flow to magmatism, but here are only
schematic and do not follow the geodynamic
models exactly. Mantle flow during the initial
tearing of the slab in Late Cretaceous [LithgowBertelloni and Richards, 1998] would be moderate
according to the geodynamic models for North Fiji
(Figure 9b). The time gap between the proposed
slab detachment 90 Ma (Figure 9b) and initiation
of magmatism (55 – 60 Ma in Australia)
(Figure 9c) can be explained by the P-T diagram
(Figure 3). The present-day SE Australia geotherm
would decay to the conductive Phanerozoic geotherm in 40–50 My [O’Reilly et al., 1997; Sass
and Lachenbruch, 1979]. This also represents the
maximum amount of time to conductively heat
from the Phanerozoic to present-day geotherm. If
convective processes occur, as evidenced by magmatism and underplating in the region, the heating
time would be considerably shorter, in the extreme
case of convective processes alone, nearly instantaneously [Lachenbruch and Morgan, 1990]. As
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finn et al.: alkaline magmatic province
the onset of the magmatism was 55 Ma, the
processes to generate magmatism would have been
set in motion 55–95 Ma, coincident with the slab
detachments proposed by plate reconstruction
models (Figure 9b). As indicated by the plate
history models [Lithgow-Bertelloni and Richards,
1998], the Pacific slab would have sunk completely
into the lower mantle by 65 Ma (Figure 9c).
Mantle flow is supported by stage poles derived
from plate motion rotations, which show west and
southwesterly lateral flow of Pacific mantle beneath
the region between about 80 and 40 Ma [Gaina et
al., 2000].
[40] As the plate detaches and sinks into the lower
mantle (Figure 5c) (Late Cretaceous) [LithgowBertelloni and Richards, 1998; Steinberger et al.,
2004], relatively warm, Pacific mantle would flow
under the metasomatized SCLM (Figure 9c). The
Pacific mantle is considered to be warm enough to
generate melts, an observation supported by plate
history models, the low seismic velocity anomalies
in central Pacific mantle (Figure 5) and numerous
seamounts [Anderson, 1994]. Lack of cooling by
subduction for millions of years is one explanation
for the warm temperatures [Anderson, 1995]. Several lines of evidence suggest that the central
Pacific superswell, a broad region of uplift dynamically supported by buoyant, low-viscosity lowvelocity upper mantle material [McNutt, 1998],
was warmed in the late Cretaceous. Recent analysis
of bathymetry data suggest that warm mantle with
a temperature gradient of 0.014C/km flows from
the superswell to the East Pacific Rise and that the
onset of the swell was 98 Ma [Hillier and Watts,
2004], roughly synchronous with the proposed slab
detachments. In addition, thermal models constrained by seismic surface wave disperson data
suggests that the central Pacific lithosphere
was reheated between ages of 70 and 100 Ma
predominantly at depths between 70 and 150 km
[Ritzwoller et al., 2004].
[41] On the basis of the above discussion, the effect
of removing the slab and migration of Pacific
mantle beneath the Gondwana pieces could raise
temperatures sufficiently to generate alkaline melts
as indicated by the P-T diagrams (Figure 3). If the
late Cretaceous geotherm for Gondwana is similar
to that estimated for Phanerozoic Australia (green
line, Figure 7), and pressure did not decrease (that
is, little to no extension) in the Cenozoic, a
temperature increase of 100C (indicated by red
arrow, Figure 3) is required to match the current SE
Australia geotherm (orange line, Figure 3) and
10.1029/2004GC000723
perhaps much of the rest of the SW Pacific. If
the temperature of mantle material at 150 –
200 km beneath >50 Ma oceanic lithosphere is
1300 –1400C [Shapiro and Ritzwoller, 2002;
Ritzwoller et al., 2004], flow of this material
beneath the 100 – 150 km thick continents
(Figure 9c) could partially melt both the metasomatized SCLM mantle (due to its low melting
temperature) and rising asthenospheric mantle
(though decompression melting). Subsequent magmatism could produce the observed isotopic signatures interpreted to indicate mixing of different
reservoirs (SCLM and asthenosphere) that is a
combination of the MORB and FOZO-like components, as discussed previously. Extension superimposed on this system, such as in the western
Ross Sea during the Oligocene and currently, could
account for the higher temperature gradients [Berg
et al., 1989] (blue line, Figure 3) and more voluminous magmatism there than elsewhere in the SW
Pacific. Volcanism should persist until the lowmelting point metasomatized layer is depleted or
subduction is renewed.
5. A Regional Model for a Cenozoic
South Pacific Diffuse Alkaline
Magmatic Province
[42] Catastrophic slab detachments in the late Cretaceous would mostly likely induce mantle flow in
a broad region and therefore may explain Cenozoic
magmatism not only in the continental pieces
described here, but over much of the south Pacific,
including the superswell region, which contains
many scattered, short-lived volcanic chains with
linear age progression that cannot easily be
explained by conventional plume theory (e.g.,
Austral Islands; Figure 2) [McNutt and Bonneville,
2000; McNutt et al., 1997].
[43] In order to investigate the potential link between slab detachments and mantle flow leading to
warming in the south Pacific, we reconstruct mantle density from 100 Ma to the present on the basis
of a time-dependent global mantle flow model
[Steinberger et al., 2004]. The models are generated by integrating a current mantle density field
backward in time with global plate motions model
superimposed as boundary conditions. The integration is accomplished by reversing the sign of the
density anomaly which effectively reverses the
convection back through time. The initial density
structure in the flow model was derived from a
shear wave tomography model [Becker and Boschi,
18 of 26
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a)
b)
c)
finn et al.: alkaline magmatic province
10.1029/2004GC000723
2002]. The velocity anomalies were converted to
density anomalies with a conversion factor of 0.25
below 220 km. In our model, current mantle
density anomalies are advected back to 100 Ma,
not restricted to the Tertiary as in previous models
[Steinberger et al., 2004]. At 100 Ma, high-density
mantle reflecting subducting slabs partially underlie the east Gondwana margin (Figure 10a). By
50 Ma (Figure 10b), the slabs have sunk into the
lower mantle. Between 50 Ma (Figure 10b) and the
present (Figure 10c), a superswell-size negative
mantle density anomaly, presumably reflecting
warm temperatures and perhaps volatile content,
forms in the southwest Pacific, suggesting a link
between slab detachment in the lower mantle
and upwelling. Three-dimensional modeling of
upper-mantle anelastic structure [Ekstrom and
Dziewonski, 1998] of the prominent low-velocity
anomaly beneath the Pacific superswell (Figures
4a, 4c, 5b, and 5c) also suggest that thermal
upwelling from the lower mantle carry enough
energy across the transition zone to create coherent
upwelling flow in the upper mantle [Romanowicz
and Gung, 2002] as observed in our mantle density
models (Figure 10). This upwelling could then
supply heat and horizontal flow to the low-viscosity asthenospheric channel, thereby feeding hot
spots in the superswell [Romanowicz and Gung,
2002].
[44] Therefore we suggest that the continental
alkaline magmatism described here may lie at the
southwestern edge of a Cenozoic Pacific diffuse
alkaline magmatic province (DAMP) largely
defined by broad, discontinuous regions of relatively low volume alkaline basalts erupted intermittently since 55–30 Ma in 80–150 km thick
lithosphere. The extent of the DAMP most likely
covers the south Pacific region associated with
unusual mantle low-velocity zones (e.g., most of
the region of Figure 2 and the low-velocity region
of Figure 5c). The thermal flow model for the
superswell bathymetry suggests that warm temperature mantle extends at least to the East Pacific
Rise [Hillier and Watts, 2004] as do our models
Figure 10. Mantle density anomaly reconstruction for
500 km depth (in units of 1/1000 kg/m3) at (a) 100 Ma,
(b) 50 Ma, and (c) the present based on the timedependent global mantle flow model from Steinberger et
al. [2004]. High-density slabs lie beneath the upper
mantle in parts of the SW Pacific DAMP at 100 Ma but
have sunk into the lower mantle by 50 Ma. Lighter,
presumably warmer mantle is developed during this
time period.
19 of 26
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finn et al.: alkaline magmatic province
(Figure 10). However, definition of the entire
province is beyond the scope of this paper.
[45] Metasomatized SCLM (60–150 km depth,
Figures 3b – 3d) partially sources the melts in
continental regions and produces part of the observed regional low-velocity signature as well as
enriched and HIMU-like geochemical signatures.
The low-velocity signature outside of SCLM probably reflects warm temperatures in the asthenosphere and, geochemically, is the source of
depleted (MORB) and FOZO-like components.
On the whole, the isotopic heterogeneity is similar
to that observed in basalts from the Central Pacific
(e.g., Austral-Cook islands [Schiano et al., 2001])
and mixing of Gondwana SCLM with Pacific
asthenosphere is consistent with the diversity of
sources that are required to explain the isotopic
arrays in Figure 7. All of these lines of evidence
support the notion that mantle flow induced by slab
detachments explains the SW Pacific DAMP and
perhaps short-lived volcanism along ephemeral hot
spot tracks in the entire south Pacific as well.
Acknowledgments
[46] Special thanks go to Mike Ritzwoller and Nikolai
Shapiro for access to tomographic models and overall responsiveness to questions. We thank Bernhard Steinberger for
providing the mantle density files for Figure 10. Carmen
Gaina and Eric Anderson assisted with the figures. Carolina
Lithgow-Bertelloni provided a mantle density model for
85 Ma. Very constructive and comprehensive reviews and
comments by Gerhard Wörner, Sergio Rocchi, an anonymous
reviewer, and editors William White and Mary Reid greatly
improved this paper. C.A.F.’s work was funded by National
Science Foundation grants OPP-9319877 and OPP-9618568
and the U.S. Geological Survey Mineral Resource Program.
K.S.P. funding is from NSF grants OPP-9419686 and OPP0003702.
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