Download Melting depths and mantle heterogeneity beneath

JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 104, NO. B2, PAGES 2817–2829, FEBRUARY 10, 1999
Melting depths and mantle heterogeneity beneath Hawaii
and the East Pacific Rise: Constraints from Na/Ti
and rare earth element ratios
Keith Putirka
Lawrence Livermore National Laboratory, Livermore, California.
Abstract. Mantle melting calculations are presented that place constraints on the
mineralogy of the basalt source region and partial melting depths for oceanic basalts.
Melting depths are obtained from pressure-sensitive mineral-melt partition coefficients for
Na, Ti, Hf, and the rare earth elements (REE). Melting depths are estimated by
comparing model aggregate melt compositions to natural basalts from Hawaii and the
East Pacific Rise (EPR). Variations in melting depths in a peridotite mantle are sufficient
to yield observed differences in Na/Ti, Lu/Hf, and Sm/Yb between Hawaii and the EPR.
Initial melting depths of 95–120 km are calculated for EPR basalts, while melting depths
of 200 – 400 km are calculated for Hawaii, indicating a mantle that is 3008C hotter at
Hawaii. Some isotope ratios at Hawaii are correlated with Na/Ti, indicating vertical
stratification to isotopic heterogeneity in the mantle; similar comparisons involving EPR
lavas support a layered mantle model. Abundances of Na, Ti, and REE indicate that
garnet pyroxenite and eclogite are unlikely source components at Hawaii and may be
unnecessary at the EPR. The result that some geochemical features of oceanic lavas
appear to require only minor variations in mantle mineral proportions (2% or less) may
have important implications regarding the efficiency of mantle mixing. Heterogeneity
required by isotopic studies might be accompanied by only subtle differences in bulk
composition, and material that is recycled at subduction zones might not persist as
mineralogically distinct mantle components.
1.
Introduction
Fundamental questions in the Earth sciences concern the
mineralogy of the basalt source region and the temperatures
and depths at which basalts form. It has been proposed that
isotopic and other compositional differences between basalts
reflect mineralogical heterogeneity [Langmuir et al., 1977; Hofmann and White, 1982; Zindler et al., 1984; Allegre and Turcotte,
1986] (and more recently, Hirschmann and Stolper [1996] and
Hauri [1996]). However, isotopic differences need not be accompanied by substantial differences in mineralogy (compare
whole rocks and mineral separates from Okano and Tatsumoto
[1996] and Xu et al. [1998]). Moreover, variations in the pressures (P) and temperatures (T) of partial melting have a
significant impact on basalt composition [Klein and Langmuir,
1987; Kinzler and Grove, 1992b; Kinzler, 1997]. The nature and
degree of mantle heterogeneity cannot be fully ascertained
until such effects are evaluated. To investigate the potential
effects of P and T on melt composition, melting models are
developed here, and the results are compared to tholeiites
from the East Pacific Rise (EPR) and tholeiitic and alkalic
lavas from Hawaii. The models test whether some of the geochemical variations at these localities can be explained by differences in the P-T conditions of melting. These calculations
also constrain minimum degrees of source region heterogeneity. The models use Na, Ti, Hf, and rare earth element (REE)
abundance’s since their mineral/melt partition coefficients are
Copyright 1999 by the American Geophysical Union.
Paper number 1998JB900048.
0148-0227/99/1998JB900048$09.00
sensitive to both P and mineralogy. Partition coefficients for
these elements have also recently been calibrated to high P and
T [Putirka, 1998a, b; Salters, 1996], making these elements
particularly useful for examining mineralogic heterogeneity
and potential P-dependent geochemical variations.
The nature and degree of heterogeneity impact upon issues
of mantle composition and the efficiency of mantle mixing
[Hofmann et al., 1986; Kellogg, 1992; Farber et al., 1994], as well
as the thermal structure of the mantle. It is thus essential to
differentiate between heterogeneity with respect to mineralogy, isotope ratios, or differences in major or minor element
composition. To examine issues of heterogeneity and the conditions of partial melting, alkalic and tholeiitic basalts from
Hawaii and tholeiites from the EPR are compared. Hawaii is
the type example for a mantle plume, while EPR basalts are
representative of mid-ocean ridges [Langmuir, 1988]. As Hawaii represents a mantle hot spot [Wilson, 1963; Morgan, 1972;
Sleep et al., 1988], melting depths should be deep compared to
the EPR, where melting occurs in response to passive upwelling [Oxburgh and Turcotte, 1968; Cawthorn, 1975; Ahern
and Turcotte, 1979; Spiegelman, 1996]. The Hawaiian island
chain is also underlain by thick lithosphere (70 –100 km [Bock,
1991; Woods and Okal, 1996]). Faced with such geologic differences, we might ask: Can differences in mantle temperature
or lithosphere thickness explain some of the geochemical differences between Hawaii and the EPR? This question is addressed by comparing natural basalts to calculated melt compositions derived from a homogenous depleted peridotite
source (see the appendix and Table 1). While the mantle is
known to be heterogeneous, such calculations constrain heterogeneity. Geochemical variations that are unexplained by such
2817
2818
Table 1.
PUTIRKA: MELTING DEPTHS AND MANTLE HETEROGENEITY
Model Compositions
Major
Oxides
Value
Mineral Modes
Value
SiO2
TiO2
Al2O3
FeO
MgO
CaO
Na2O
46.316
0.174
3.935
7.653
38.44
3.192
0.281
olivine
orthopyroxene
clinopyroxene
garnet
0.60
0.09
0.17
0.14
models provide minimum estimates on source region differences. To test for mineralogic heterogeneity, melting models
using eclogite and various garnet pyroxenite starting compositions are also developed.
Basalts from Hawaii and the EPR exhibit notable geochemical differences. Both tholeiites and mildly alkalic Hawaiian
basalts have lower Na/Ti ratios, due to higher Ti and lower Na
(Figure 1). Lower melt fractions at Hawaii could account for
increased Ti but not for Na if Na is also incompatible. This
conundrum might be explained if mineral-melt partition coefficients for Na increase with increased P [Frey et al., 1994;
Kinzler, 1997]. Hawaiian basalts also exhibit high Sm/Yb and
low Lu/Hf ratios compared to EPR lavas [Salters, 1996; Salters
and Hart, 1991, 1989], and greater melting depths at Hawaii
have been invoked to explain such differences [Salters, 1996].
Finally, there exist significant intershield variations in Na/Ti at
Hawaii; these intershield differences hold whether alkalic or
Figure 1. Weight percent Na2O and TiO2 are compared for
tholeiites from the East Pacific Rise (EPR) and from several
Hawaiian volcanoes. The Mauna Kea and Loihi groups also
include alkalic lavas. Most Hawaiian basalts have lower Na/Ti
due to both lower Na and higher Ti contents. Basalts from the
Koolau shield have intermediate Na/Ti ratios. Data are for
Loihi from Garcia et al. [1993, 1995]; Mauna Kea from Frey et
al. [1991], Yang et al. [1996], Rhodes [1996], and Rhodes and
Hart [1995]; Mauna Loa from Rhodes [1988], Yang et al. [1996],
and Rhodes [1996]; Kilauea from Garcia et al. [1992] and Chen
et al. [1996]; and Koolau from Frey et al. [1994]. East Pacific
Rise basalt compositions are from Langmuir [1988] (see also
Langmuir et al. [1986]).
tholeiitic rocks are examined (Figure 1). The melting models
test whether P-T conditions during melting can explain such
variations.
2.
Geochemical Data and Corrections
To examine geochemical variation due to mantle melting, it
is essential to minimize the effects of shallow level fractionation. As a test for shallow level fractionation, Na/Ti and
Sm/Yb ratios are compared to MgO (Figure 2). For basaltic
rocks, MgO is a useful indicator of fractionation, since olivine
is normally crystallized at shallow depths. While some scatter
exists, Na/Ti and Sm/Yb ratios are independent of MgO for
Figure 2. (a) Na/Ti cation fraction ratios are compared to
MgO (weight percent) for EPR and Hawaiian basalts. Note
that alkalic and tholeiitic lavas from Loihi and Mauna Kea
feature similar Na/Ti ratios and are distinct compared to
Koolau, Mauna Loa or EPR tholeiites. At Hawaii, Na/Ti ratios
are insensitive to MgO content. Na/Ti ratios at Hawaii are thus
unlikely affected by low-pressure fractionation and should reflect source composition and melting conditions. In contrast,
EPR basalts are affected by low-pressure fractionation. Only
EPR basalts with .8 wt % MgO were used to constrain EPR
melting depths; Na/Ti ratios for such EPR basalts are higher
than at Hawaii. (b) The concentration ratio Sm/Yb is compared to wt % MgO for the same samples as in Figure 2a.
Sm/Yb ratios are independent of MgO content and should thus
reflect melting conditions and mantle composition.
PUTIRKA: MELTING DEPTHS AND MANTLE HETEROGENEITY
Hawaiian tholeiites and alkalic lavas and thus unaffected by
shallow level processing (Figure 2). Since accumulation or
fractionation of olivine does not affect Na/Ti or Sm/Yb ratios,
Hawaiian basalts with 7–13 wt % MgO are examined. An
interesting interisland difference is apparent (Figure 2a): basalts from Mauna Kea, Loihi, and Kilauea have Na/Ti 5 2–2.5,
while basalts from Mauna Loa are somewhat higher at 2.5–3.0,
and Koolau basalts are higher still with Na/Ti, ranging from 3
to 4.
In contrast to Hawaiian lavas, EPR basalts display Na/Ti
ratios that decrease sharply with decreasing MgO (Figure 2a);
this probably reflects plagioclase fractionation. Ratios of
Sm/Yb do not vary with MgO (Figure 2b) and are thus apparently unaffected by such fractionation. Experimental studies
indicate that plagioclase saturation for mid-ocean ridge basalts
(MORB) occurs near 8 –10 wt % MgO [Tormey et al., 1987;
Grove and Bryan, 1983]. The highest Na/Ti ratios (Figure 2a)
may thus approach primitive values. To minimize the effects of
fractionation on Na/Ti, only basalts with .8 wt % MgO will be
compared to model values for constraining partial melting
depths. Na/Ti ratios for such EPR basalts (4 –5) are significantly higher, and Sm/Yb ratios are lower, compared to Hawaiian tholeiites and alkalic lavas.
3.
Results
The melting model of this study follows the work of Langmuir et al. [1992] (see the appendix) and examines the compositions of aggregate melts, i.e., melts that are pooled from a
range of depths within a melting column. Thus each choice for
mantle potential temperature (T p ) corresponds to a unique
initial depth of partial melting, and the choice of lithosphere
thickness corresponds to a final depth of melting. In addition,
for a particular starting composition, when lithosphere thickness and T p are specified, there exists a unique maximum and
mean extent of partial melting, or F (melt fraction). Since only
two variables (T p , initial melting depth, lithosphere thickness,
and mean F) need be specified for a given source composition,
the following discussion will refer to differences in initial melting depths and lithosphere thickness; differences in mean F
and T p are implied.
Initial melting depths and lithosphere thickness both significantly affect Na/Ti and Sm/Yb ratios in aggregate melts. This
is true for both peridotite and eclogite or garnet pyroxenite
mantle sources. In general, partial melting of a garnet pyroxenite leads to aggregate melts that have higher Na/Ti and
Sm/Yb ratios, compared to aggregate melts from a peridotite
source. High Na/Ti for garnet pyroxenite partial melts result, in
part, from higher Na/Ti in the source (5, compared to 4 for
peridotite). More significant, though, is the greater amount of
garnet in the garnet pyroxenite (see the appendix), which increases the garnet pyroxenite bulk mineral/melt distribution
melt
sol
coefficient for Ti (5D sol/melt
5 C sol
Ti
Ti /C Ti , where C Ti is the
concentration of Ti in the total solid and C melt
is
the
concenTi
tration of Ti in the melt). High Sm/Yb ratios similarly reflect
the high modal proportion of garnet, which results in an increased D sol/melt
. For both peridotite and pyroxenite starting
Yb
compositions, Na/Ti ratios decrease and Sm/Yb ratios increase
in aggregate melts, as lithosphere thickness increases.
Mid-ocean ridge basalts were modeled using a 10-km lithosphere; at Hawaii, lithosphere thickness’ are less certain. For
example, convective thinning is presumed to reduce an initially
90 –100 km thick lithosphere to 50 km or less [Detrick and
2819
Crough, 1978; Liu and Chase, 1989]. However, heat flow [VonHerzen et al., 1989] and seismic studies [Woods and Okal, 1996]
are inconsistent with such thinning, and an average lithosphere
thickness of 100 km is indicated [Woods and Okal, 1996].
Lithosphere thickness, though, is unlikely to be constant along
the Hawaiian chain and may influence magma production
[Morgan et al., 1995; Wessel, 1993]. Since estimates of 75– 80 km
lithosphere have been obtained at Oahu [Bock, 1991], a 75–120
km range was explored.
When thick lithosphere is present, Na/Ti and Sm/Yb ratios
decrease as initial melting depth (or T p ) increases (Figure 3).
Na/Ti ratios are higher beneath thin lithosphere and decrease
as initial melting depth increases (Figure 3a). Na/Ti ratios
decrease with increased melting depth because the K cpx/melt
Na
melt
(Na partition coefficient for clinopyroxene/melt, 5 C cpx
Na /C Na ,
i
where C Na is the concentration of Na in the superscripted
phase) increases with increased P, while clinopyroxene and
garnet K min/melt
remain constant or decrease [Langmuir et al.,
Ti
1992; Blundy et al., 1995; Kinzler, 1997; Walter, 1998; Putirka,
1998a, b]. Most importantly, without the P-induced change in
K cpx/melt
, Ti and Na are not significantly fractionated from one
Na
another, and aggregate melt Na/Ti ratios are nearly invariant.
Low Na/Ti ratios thus indicate a contribution from melts produced at substantial depth.
The decrease in Sm/Yb beneath thick lithosphere is due to
the progressive removal of garnet from the residue during
partial melting. When melting begins, the amount of garnet is
highest, and so is D sol/melt
, (since K garnet/melt
is high). As meltYb
Yb
ing progresses, though, garnet is removed from the residue,
and D sol/melt
decreases. By increasing initial melting depths
Yb
(i.e., increasing mean and maximum F) the amount of garnet
in the total melting column is decreased, as is the average
D sol/melt
in the melting column. Aggregate melts at the surface
Yb
reflect this average D sol/melt
. When lithosphere thickness exYb
ceeds 75 km, all melting occurs in the garnet stability field
[Takahashi et al., 1993; Longhi, 1995; Kinzler, 1997; Walter,
1998], and the progressive consumption of garnet dominates
the Sm/Yb signal. For both Na/Ti and Sm/Yb the aggregate
melt curves flatten with depth because melt productivity is
reduced at increased P [Putirka, 1997]. The Na/Ti curves exhibit less flattening with depth, compared to Sm/Yb, because
K cpx/melt
increases steadily with depth throughout the melting
Na
column. In contrast, D sol/melt
, which depends upon the modal
Yb
abundance of garnet, does not vary greatly due to the small
amount of partial melt that is produced (and hence the small
amount of garnet that is consumed) at high P.
In the case of thin lithosphere, changes in Sm/Yb ratios are
less dramatic (Figure 3b) because melts produced deep in the
column (in the presence of garnet) are overwhelmed by the
comparatively large amounts of melt produced at shallow
depths (,75 km). With 10 km lithosphere, Sm/Yb first decreases then exhibits a rapid increase once the garnet stability
field is entered (75 km). The initial decrease in Sm/Yb, between initial melting depths of 50 and 75 km, is due to variations in mean F. At these depths D sol/melt
. D sol/melt
; in the
Yb
Sm
absence of garnet (,75 km), though, these D REE are very
small, and Sm and Yb are only fractionated from one another
when F is very low. As melting depths increase from 75 to 50
km, both maximum and mean F increase, and Sm/Yb ratios
approach initial (unfractionated) values. Sm/Yb ratios for aggregate melts increase sharply when initial melting depths
reach the garnet stability field (75 km), since K gar/liq
is very
Yb
high. Sm/Yb ratios are also sensitive to melting stoichiometry
2820
PUTIRKA: MELTING DEPTHS AND MANTLE HETEROGENEITY
(Figure 3c). As an illustration, peridotite 1 (melting coefficient
for garnet is 0.2) is compared to peridotite 2 (coefficient is 1.0).
These values range beyond observed values (0.7 [Walter,
1998]), but Figure 3c shows the magnitude by which melting
stoichiometry might influence Sm/Yb systematics at depth.
4.
Discussion
Variations in the P-T conditions of melting alone might
account for substantial geochemical variations in natural basalts (Figures 3–5). Such variations encompass the large-scale
differences between the EPR and Hawaii, as well as some
intershield variation at the Hawaiian islands. In particular, the
present model supports the Frey et al. [1994] and Kinzler [1997]
hypotheses regarding the importance of melting depth for reconciling variations in Na2O and TiO2 in oceanic basalts. These
results do not imply that the mantle is homogenous; heterogeneity appears to be required at both Hawaii and the EPR
[Langmuir and Hanson, 1980; Roden et al., 1994; Prinzhofer et
al., 1989; Hekinian et al., 1989; Reynolds et al., 1992]. However,
such heterogeneity might not encompass all major oxides, let
alone involve major differences in source region mineralogy.
Since the P-T conditions of melting can account for significant
major and minor element systematics, source region differences may be limited.
4.1.
Figure 3. (a) Na/Ti ratios (cation fraction) of aggregate
melts produced from peridotite and garnet pyroxenite sources
are compared at different P-T conditions. Lithosphere thickness indicates depth to the top of the melting column; “initial
melting depth” represents depth to the base of the melting
column and hence also reflects mantle potential temperature.
(b) Sm/Yb ratios for aggregate melt compositions are compared, and (c) a close-up view of the 10-km lithosphere calculations is given. In Figure 3c the “peridotite 1” model uses a
garnet melting coefficient of 0.2; the “peridotite 2” model uses
a garnet melting coefficient of 1. For a given lithosphere thickness a garnet pyroxenite source yields higher Na/Ti and Sm/Yb
ratios compared to a peridotite source.
Estimates of Melting Depths and Temperatures
At the EPR, initial melting depths of 120 6 15 km reproduce
the approximate concentrations of Na, Ti, Sm, and Yb (Figure
4). These depths overlap with depth estimates implied by Salters [1996, Figure 2] and are consistent with depths of lowvelocity and electromagnetic anomalies at mid-ocean ridges
[MELT Seismic Team, 1998; Zhang and Tanimoto, 1993;
Anderson et al., 1992]. Mantle temperature estimates are sensitive to the temperature of the mantle solidus and are much
less robust than melting depth estimates (see the appendix). If
the mantle has a composition similar to KLB-1, a melting
depth of 110 km suggests a T p of 1831 K (15588C). While
mantle temperatures are not well constrained from geophysical
studies [Parsons and Sclater, 1977; Stein and Stein, 1992], T p
inferred using a KLB-1 composition overlaps with mantle temperatures of 1723 6 250 K (at 95 km depth), obtained from
global studies of heat flow and ocean bathymetry [Stein and
Stein, 1992]. It cannot be overemphasized, though, that temperature estimates are only as reliable as our knowledge of
mantle composition and solidus temperatures. Experimental
studies show a range of melting temperatures for different
peridotite compositions [Hirose and Kushiro, 1993]; actual temperatures could be lower by as much as 50 K, or more, if the
mantle is more fertile than KLB-1.
Much greater initial melting depths are implied for Hawaiian tholeiite and alkalic lavas. Abundances of Na, Ti, Hf, and
REE can be produced using the same depleted mantle composition as that for the EPR but with increased lithosphere
thickness and initial melting depths of 200 – 400 km. An undepleted mantle source would yield higher initial Na/Ti ratios
(Figure 5); to achieve the low Na/Ti ratios of Hawaii for an
undepleted source, increased initial melting depths (by ;50 –
100 km) would be required.
Estimates of melting depths at Hawaii exceed those obtained by Watson and McKenzie [1991]. There are numerous
reasons for these differences. First, it should be noted that the
Watson and McKenzie [1991] model does not successfully re-
PUTIRKA: MELTING DEPTHS AND MANTLE HETEROGENEITY
2821
in a melt productivity that is too high. High melt productivity
results in (1) elevated estimates of mean F and (2) lower
incompatible element concentrations for a given initial depth
of melting. The present model, with greater initial melting
depths and a more realistic DS f (;420 J/kg), successfully yields
Na, Ti, and REE concentrations for Hawaiian and EPR lavas
(Figure 4).
Since garnet is gradually consumed during melting, a garnetfree residue is produced at the top of the melting column at
Hawaii. If the melting coefficient for clinopyroxene is .0.5,
then clinopyroxene is also consumed by depths of 100 km,
leaving a harzburgite residue at shallow depths. This predicted
residue might resolve the apparent conflict that high-MgO
Kilauea tholeiites are harzburgite saturated at pressures ,3.5
GPa [Eggins, 1992], yet show a garnet trace-element signature
[Budahn and Schmitt, 1985; see also Frey et al., 1994]. Using the
solidus of KLB-1, a T p of 2130 K (18578C), is implied to
achieve melting at 300 km. Such temperatures represent a
mantle that is approximately 300 K hotter at Hawaii compared
to the EPR, which is consistent with temperature differences
inferred from tomographic [Romanowicz, 1994; Anderson et al.,
1992] and geodynamic studies [Sleep, 1990; Ribe and Christensen, 1994]. Calculated melting depths are also consistent
with depths for low seismic velocities and high seismic wave
attenuation [Bhattacharyya et al., 1996; Romanowicz, 1995;
Zhang and Tanimoto, 1993]. The temperature difference between Hawaii and the EPR is probably more robust compared
to absolute temperature estimates at Hawaii, which are subject
to the same uncertainties as noted above for the EPR.
4.2.
Figure 4. (a) Natural and model Sm concentrations (ppm)
and Sm/Yb ratios are compared. (b) Calculated and observed
concentrations for Na and Ti are compared. Hawaiian basalts
were modeled using a 100-km-thick lithosphere and an initial
melting depth of 300 km; EPR basalts were modeled using a
10-km lithosphere and an initial melting depth of 115 km;
these model calculations are indicated with the letters H (Hawaii) and E (EPR) (circled). Observed abundance’s of Na, Ti,
and some REE can be reproduced by varying the P-T melting
conditions of a mantle that has a depleted pyrolite bulk composition and a peridotite mineralogy.
produce Na2O and TiO2 contents at Hawaii. Their predicted
Na/Ti ratios [Watson and McKenzie, 1991, Table 2] are much
too high, indicating that their melting depth estimates (partition coefficients for Na) are too low. Watson and McKenzie
[1991] also use a final melting depth of 82 km and assume that
spinel peridotite is stable to depths apparently exceeding 100
km. The lithosphere beneath some parts of Hawaii is conceivably as thin as proposed by Watson and McKenzie [1991], but
spinel peridotite is probably not stable, even at depths as shallow as 82 km [Takahashi et al., 1993; Longhi, 1995; Kinzler,
1997; Walter, 1998]. Using a more realistic garnet lherzolite
mineralogy, such shallow melting depths yield Sm/Yb ratios
that are too high compared to observed values (Figure 5).
Finally, the entropy of fusion (DS f ) used by Watson and
McKenzie [1991] is too low, perhaps by a factor of 2, resulting
Mantle Heterogeneity
The degree and nature of heterogeneity in the basalt source
region can be constrained when the potential P-T effects on
partial melt compositions are evaluated. Geochemical variations that are unexplained by a homogenous peridotite mantle
model provide minimum estimates on source region differences. In some cases, such as at Hawaii, differences in P-T
conditions are probably not sufficient to explain all geochemical variations, and minimum differences in modal mineralogy
for source regions are calculated. In addition to peridotite,
eclogite and various garnet pyroxenite starting compositions
are also examined to resolve whether such sources are consistent with observed geochemical signals.
4.2.1. Heterogeneity beneath the EPR. Hirschmann and
Stolper [1996] have proposed that garnet pyroxenite exists in
the MORB source region. Their argument was in part based
upon the apparent discrepancy between the deep levels of
melting required to produce the garnet signature in MORB
and the relatively thin crust that is formed at oceanic spreading
centers. As noted by Putirka [1997], though, fractional melting
processes limit melt productivity and may be invoked to satisfy
both observed crustal thickness [White et al., 1992] and inferred
melting depths of .100 km [Salters, 1996]. Regarding geochemical variation, Na/Ti and Sm/Yb ratios at the EPR are
consistent with partial melting of a peridotite source. Some
mixing with a garnet pyroxenite source is allowed, since the
EPR samples plot between the peridotite and garnet pyroxenite curves (Figure 5b). However, only the major oxides carry
any significant leverage on mineral abundances. It is possible
that the observed geochemical differences, at least with respect
to Na and Ti, might be supported by minor mineralogic variations within a source that has an overall peridotite mineralogy. For example, a variation of 60.2 wt % Na in the EPR
2822
PUTIRKA: MELTING DEPTHS AND MANTLE HETEROGENEITY
source region could explain observed ranges of Na/Ti ratios in
the high MgO EPR lavas; such differences can be accommodated by very small (,2%; see below) changes in peridotite
mineral abundances.
As recognized by Hirschmann and Stolper [1996], partial
melts from neither garnet pyroxenite nor peridotite sources
exhibit sufficient variation to explain the full range of MORB
Lu/Hf and Sm/Nd ratios. If the lithosphere thickness is increased to 30 –50 km, peridotite partial melts can reproduce
the low Lu/Hf and Sm/Nd of the MORB array (Figure 5c).
Thicker lithosphere may not be surprising at such localities as
the Atlantic-Antarctic Discordance (AAD), since cooler mantle temperatures are indicated from comparisons of mid-ocean
ridge bathymetry and geochemistry [Klein and Langmuir, 1987;
Langmuir et al., 1992; Salters, 1996]. Neither the garnet pyroxenite nor the peridotite models, though, are capable of reproducing the high Lu/Hf and Sm/Nd ratios observed at Kolbeinsy. In contrast, Salters’ [1996] peridotite melting model
appears to account for all but the Kolbeinsy data by taking into
account variations in the timing of source region depletion.
4.2.2. Heterogeneity beneath Hawaii. At Hawaii, eclogite
has been hypothesized to exist in the source region, with eclogite partial melts contributing most greatly (up to 20%) to
some Koolau basalts [Hauri, 1996]. Hauri’s [1996] hypothesis
stems, in part, from the observation that Hawaiian lavas are
enriched in SiO2 compared to partial melts obtained from
experimental studies of lherzolites. However, several isotope
and trace element studies suggest that a recycled source at
Koolau might be minimal [Roden et al., 1994; Bennett et al.,
1996]. Furthermore, Frey et al. [1994] interpret interisland SiO2
systematics as indicating depth of melting differences, rather
than source region heterogeneity. Wagner and Grove [1998]
have also shown that basalts can obtain elevated SiO2 if orthopyroxene is assimilated at shallow depths. (Note: orthopyroxene assimilation does not affect Na/Ti or Sm/Yb. Using
data from Wagner and Grove [1998], assimilation of 20% or-
Figure 5. (opposite) (a) Aggregate melt compositions are
from a depleted peridotite source. High Na/Ti (cation fraction)
at Koolau may be explained by the presence of thinner lithosphere and cooler mantle temperatures. Intravolcano variation
probably reflects heterogeneity. Peridotite models use a depleted mantle composition (see the appendix and Table 1);
“bulk composition uncertainty” reflects the increase in Na/Ti
and Sm/Yb ratios that are obtained when an undepleted mantle source is input into the melting models (see the appendix).
(b) Calculated aggregate melts are shown for eclogite and
garnet pyroxenite sources (“HS” is the average garnet pyroxenite from Hirschmann and Stolper [1996]; “massif” refers to
garnet pyroxenites from massif peridotites; see the appendix; K
is the Koolau and ML is the Mauna Loa 1 Mauna Kea 1 Loihi
fields, as in Figure 5a). Na/Ti and Sm/Yb ratios at Hawaii
overlap only with some massif pyroxenites, but the predicted
Na, Ti and REE abundances are too high compared to lava
compositions (see text). (c) Lu/Hf and Sm/Nd from MORB
[Salters, 1996] are compared to Hawaiian compositions (AAD,
Atlantic-Antarctic Discordance). Initial melting depths for the
10-km lithosphere peridotite and garnet pyroxenite models
range from 50 to 400 km, moving downward in the figure.
Hawaiian data are adequately described with a 100-km lithosphere and initial melting depths of 200 – 400 km (increasing
upward in the figure). Lu/Hf and Sm/Nd for other MORBs
span a range of values that cannot be covered by varying the
initial melting depths for eclogite or peridotite partial melting.
PUTIRKA: MELTING DEPTHS AND MANTLE HETEROGENEITY
thopyroxene decreases Na/Ti and Sm/Yb ratios by .015 and .05
respectively). Finally, Okano and Tatsumoto [1996] observe
that Hawaiian spinel lherzolite xenoliths exhibit a wide range
of isotope ratios, with some isotopes approaching Koolau-like
values. This observation shows that a wide range of isotope
ratios can be supported by a lherzolite mineralogy and that
isotope variations at Hawaii might not require an eclogite
source. It is thus unclear that eclogite is a necessary source
component at Hawaii.
Most of the interisland variation in Na/Ti and Sm/Yb at
Hawaii can be accounted for by plausible ranges in the P-T
conditions of melting for a peridotite mantle (Figure 5a). For
example, Na/Ti-Sm/Yb partial melting arrays overlap the
Koolau field as lithosphere thickness decreases to 75 km, and
initial melting depths decrease to 170 km. Owing to differences
in initial melting depth (and T p ), different values for mean F
are also obtained for the Hawaiian tholeiites (Figure 5a). For
comparison, estimates of F at Loihi are similar to, but slightly
higher than, those obtained by Garcia et al. [1995] and are
within the very wide range of values noted by Feigenson et al.
[1996] for Mauna Kea. As noted above, though, differences in
melting depth (in the absence of heterogeneity) are required to
fractionate Na from Ti; initial melting depth, rather than F, is
the critical variable for explaining the range of Na/Ti ratios at
Hawaii.
Undoubtedly, some intraisland variation is due to heterogeneity, especially since some Hawaiian lavas plot outside of the
partial melting grid (Figure 5). Limited variations in melting
depths and lithosphere thickness at Hawaii, though, are at least
plausible. Melting depths might vary depending upon the position of the shield volcano and the plume axis [Farnetani and
Richards, 1995], especially if shield positions are influenced by
the state of stress in the lithosphere [Jackson and Shaw, 1975].
Relative maximum melting depths for Koolau, Kilauea, Mauna
Loa, and Loihi are qualitatively consistent with melting depths
inferred from major oxide systematics [Frey et al., 1994]; the
model is furthermore consistent with trace element and isotopic observations [Bennett et al., 1996; Roden et al., 1994] that
indicate a peridotite source for Hawaiian lavas. Moreover,
Koolau lies north of the Molokai fracture zone [Mammerickx,
1989; Atwater, 1989], and seismic work [Bock, 1991] at Oahu
indicates a 75– 80 km lithosphere in this region.
If the Koolau lithosphere approaches 100 km, the mantle
source at Koolau must differ from that beneath Loihi and
Mauna Kea. Additionally, since lithosphere thickness probably
does not vary greatly at any single volcano, intraisland heterogeneity is also implied by the range of Na/Ti ratios displayed
at each of the Hawaiian shields (Figure 5a). The lowest Na/Ti
ratios at Mauna Kea are also not explained by the present
model. These low Na/Ti ratios can only be explained if (1) the
top of the melting column exceeds .150 km depth or (2) the
mantle source has higher clinopyroxene (by 2%) and lower
garnet (by 2%) compared to the mineralogy of Table 1. Option
2 seems more likely, but lower garnet results in lower Sm/Yb
ratios, perhaps indicating REE heterogeneity.
While peridotite partial melting may explain interisland variation at Hawaii, can garnet pyroxenite or eclogite sources also
yield observed Na, Ti, and REE abundances at Koolau? To
illustrate the potential range of such sources, several starting
compositions were explored, including eclogite, and garnet
pyroxenites from various sources (see the appendix). As is
evident from Figure 5b, most pyroxenites and eclogites are
unsuitable as source components, since Sm/Yb ratios are high
2823
compared to Hawaiian tholeiites and alkalic lavas (Figure 5).
Partial melts from some massif garnet pyroxenites exhibit
Na/Ti and Sm/Yb ratios that overlap with some Hawaiian tholeiite and alkalic basalts. However, Na, Ti, Hf, and REE abundances from such sources are much higher than observed at
Hawaii. Partial melts from pyroxenite sources yield Na abundances between 0.08 – 0.20 cation fraction, and Sm abundance’s of 6.2–12.5 ppm (compare to Figure 4) and are highest
for massif pyroxenites. Minor element abundances might be
diluted if mixed with melts from a peridotite source. However,
to account for Na/Ti ratios at Koolau, massif pyroxenite partial
melts would appear to dominate the Koolau aggregate melts
(up to 100%; Figure 5b). In addition, massif pyroxenites exhibit a wide range of garnet/pyroxene ratios; only when garnet
is ,20% are Sm/Yb ratios in the correct range. Finally, the thin
lithosphere and cooler melting temperatures required in the
peridotite model (above) for Koolau are not avoided by assuming a pyroxenite source. Thus, in contrast to the peridotite
model, pyroxenite and eclogite sources provide poor fits to Na,
Ti, and REE abundances at Hawaii. It would therefore appear
that such sources at Hawaii are not only unnecessary but perhaps also unlikely.
As is evident from Figure 5, variations in lithosphere thickness may be invoked to explain interisland variations at Hawaii. On the other hand, if the hypothesized lithosphere variations are incorrect, only small changes in the source region
mineralogy are required based on Na and Ti abundances. Mass
balance calculations using a mantle peridotite and highpressure mineral compositions [Walter, 1998; Putirka, 1998a, b]
underscore this conclusion. For example, if the difference between the lowest and highest Na2O values at Koolau and
Kilauea (about 1 wt %; Figure 1) are due only to heterogeneity
in the source region, mass balance implies that this difference
can be accommodated by a source region difference in clinopyroxene 1 garnet of ,2.0% (assuming mean F is 0.15, mean
D min/melt
is 0.06, and minerals have constant composition).
Na
Such source differences are reduced if any Na2O variation
results from the P-T conditions of melting, assimilation in the
shallow mantle, adjustments in mineral composition in the
source region, or shallow level fractional crystallization. It
should be noted that while these calculations are based only on
Na2O, this oxide carries more leverage on calculated mineral
proportions compared to other major oxides. This is because
Na occurs only in clinopyroxene to any great extent, and even
at high pressure it is never the dominant component in this
phase. Large changes in source Na2O concentrations therefore
translate into large changes in clinopyroxene modal abundances. Na2O is thus an ideal oxide for determining whether a
mantle source is comprised of peridotite or pyroxenite. While
an analysis of other elements, especially Ca or Al, might also
be useful (given suitable high-pressure calibrations of K d ), it is
tentatively concluded that small modal variations in a peridotite assemblage of minerals can probably accommodate observed geochemical variations at Hawaii.
Mantle heterogeneity is a flexible hypothesis and cannot be
excluded. If eclogite is not required to describe geochemical
variations in oceanic lavas, though, there are important implications for mixing in Earth’s mantle. It seems likely that small
amounts of recycled material contribute to the trace element
and isotopic signature of some oceanic lavas [Hofmann and
White, 1982; Zindler et al., 1984; Hauri et al., 1996]. Such material, though, might not survive the recycling process as a
distinct mineralogical component. Mantle mixing might be ef-
2824
PUTIRKA: MELTING DEPTHS AND MANTLE HETEROGENEITY
Figure 6. Hawaiian shield averages for isotopes and major oxides [Hauri, 1996] are compared to test for
depth-dependent isotopic heterogeneity at Hawaii. Na/Ti ratios are correlated with (a) 87Sr/86Sr, (b) 206Pb/
204
Pb, and (c) 187Os/188Os. (Note that Hauri [1996] gives two averages for Loihi: one for alkalics and one for
tholeiites; both are plotted. R 2 values are for Hauri’s [1996] shield averages.) The Loa trend volcanoes are
consistent with two-component mixing between an upper isotopically enriched mantle component and a
deeper more isotopically depleted/less enriched source. Kea trend volcanoes are not consistent with twocomponent mixing and may indicate a third component [Chen, 1987]. Interestingly, Loihi lavas [Garcia et al.,
1995; Lanphere, 1983; Frey and Clague, 1983] and Mauna Loa lavas [Rhodes and Hart, 1995] show local internal
correlations for 87Sr/86Sr-Na/Ti that are opposite to the intershield trend but directed toward MORB (Figure 7).
ficient at erasing such heterogeneity, or perhaps the melting
process is itself a homogenizing agent.
4.3.
Spatial Variability of Isotope Sources
To test for depth distributions of isotopic heterogeneity in
the Hawaiian mantle, compositional averages for the Hawaiian
shields [Hauri, 1996] were examined. Such shield averages
yield good correlation’s between Na/Ti and some isotope ratios
(Figure 6). Since, as noted above, Na can only be fractionated
from Ti at depth, it appears that melts produced at depth (low
Na/Ti ratios) tap mantle sources that are isotopically distinct
from melts produced at more shallow levels of the mantle
(higher Na/Ti). The ordering of the Loa trend volcanoes is the
same in each plot in Figure 6 and is thus consistent with mixing
between two, vertically stratified, mantle sources. This ordering for the Loa trend holds, even for the more poorly correlated Sr isotope ratios (Figure 6), and supports prior arguments for two-component mixing at Hawaii [Bennett et al.,
1996; Chen, 1987; Chen and Frey, 1985]. It is also possible,
though, that isotope ratios vary continuously but systematically
with depth in the mantle. So far as Na/Ti ratios represent
melting depth, Koolau lavas reflect partial melting of a shallow
and variably enriched/depleted [see Richardson et al., 1982]
mantle component (Figure 7). A two-component mixing
scheme is not consistent with Kea trend lava compositions,
which may indicate a less vertically stratified source [Staudigel
et al., 1984; Stille et al., 1983; Chen, 1987].
Interestingly, Sr isotopes for individual lavas at Loihi and
Mauna Loa show a local, or internal, correlation that is opposite to the intershield trend (Figure 6) but directed toward
MORB. The intershield trend would appear to indicate a lateral (along-strike) component as well as a vertical component
to local heterogeneity at the Hawaiian islands. This might be
an enriched shallow layer, perhaps at or near the base of the
Hawaiian lithosphere, that is most strongly sampled at Koolau.
In contrast, both the local trend at Hawaii and the differences
between isotope and Na/Ti ratios at Hawaii and the EPR
(Figure 7) indicate a large-scale vertical component to isotopic
heterogeneity in the Pacific upper mantle. The Hawaii-EPR
comparison supports the conventional layered mantle model,
or perhaps a “marble cake” mantle with a systematic depth
distribution of isotopically distinct blobs.
5.
Summary
A mantle melting model was developed to estimate partial
melting depths for oceanic lavas and to constrain mantle heterogeneity, using Na/Ti and REE ratios. Na/Ti ratios are particularly useful since Na and Ti are only strongly fractionated
from one another at high P. Hafnium and the REE’s are also
useful since their mineral/melt partition coefficients are sensitive to pressure and mineralogy [Salters, 1996; Kinzler, 1997;
Walter, 1998; Putirka, 1998a, b]. Several important results are
obtained from the models. First, the P-T conditions of melting
alone appear to be capable of explaining both local variation at
Hawaii, as well as some large-scale geochemical differences
between Hawaii (an ocean island) and the EPR (a mid-ocean
ridge spreading center). This implies that bulk compositional
differences between Hawaiian and EPR source regions might
be minimal; this is consistent with near-uniform, but nonprimitive, minor element ratios for oceanic basalts [Hofmann et al.,
1986] and high-pressure diffusion experiments which indicate
efficient homogenization in parts of the upper mantle [Farber
et al., 1994]. Mantle heterogeneity is also constrained; inferred
source region differences can be accommodated by small (2%)
variations in peridotite phase proportions. Na2O, TiO2, and
REE abundances at Hawaii also appear to be inconsistent with
a variety of potential eclogite or garnet pyroxenite sources. The
mantle is undoubtedly mineralogically heterogeneous, but this
and other work [Sims and Depaolo, 1997] reveal the need to
test models of heterogeneity against possible P-T dependent
variations in melt composition.
Abundances of Na, Ti, Hf, and the REE for EPR lavas are
best reproduced when melting begins at 120 6 15 km, consis-
PUTIRKA: MELTING DEPTHS AND MANTLE HETEROGENEITY
2825
Figure 7. Isotopes and Na/Ti ratios for Hawaiian lavas are compared to lavas from the EPR. Differences in
Na/Ti between EPR and Hawaii indicate that the shallow upper mantle has experienced greater extents of
long-term depletion of (a) Rb relative to Sr and (b) Nd relative to Sm. Isotopic field for EPR lavas (and
neighboring seamounts) are from Zindler et al. [1984]. Note that individual Loihi and Mauna Loa lavas appear
to trend toward MORB, perpendicular to the mean interisland trend. Symbols and data are as in Figure 6.
tent with recent seismic work [MELT Seismic Team, 1998] and
prior depth estimates [Salters, 1996]. In contrast, Hawaiian
tholeiites can be modeled using a 100-km-thick lithosphere and
initial melting depths of 200 – 400 km. This implies a temperature difference between Hawaii and the EPR of 3008C, close
to estimates inferred from geodynamic studies [Sleep, 1990;
Schilling, 1991] and mantle tomography [Romanowicz, 1994].
The compositional dependencies of peridotite solidus temperatures, though, need to be explored so that these depth estimates may be used to better resolve the thermal structure of
Earth’s upper mantle. The present model also indicates that
garnet is consumed during melting, while aggregate melts retain a garnet signature. This result may help to resolve the
apparent conflict between trace element abundances, which
indicate garnet in the source region [Budahn and Schmitt,
1985; Frey et al., 1994], and experimental studies which indicate
that some primitive Kilauea basalts are not garnet saturated at
P , 3.5 GPa [Eggins, 1992].
Appendix: Mantle Melting Model
Oceanic basalt compositions can be successfully explained
by polybaric partial melting models [Klein and Langmuir, 1987;
Kinzler and Grove, 1992b; Langmuir et al., 1992; Kinzler, 1997].
Aggregate melt compositions were thus calculated for batch
and fractional melting processes using the equations of Langmuir et al. [1992] and Plank and Langmuir [1992] and the
mineral-melt partitioning relationships of Putirka [1998a, b]
and Salters [1996]. The calculations employ a step-wise approach [Putirka, 1997]. In this method, degrees of partial melting (F) and aggregate melt compositions are calculated at
depth increments of 2 km. The residual mineralogy is adjusted
at each step using the estimated value for F and the stoichiometric coefficients of the pertinent melting equilibria (see below).
Recent experimental work indicates that at near-solidus conditions the clinopyroxene-melt partition coefficient for Ti may
decrease with increasing depth in the mantle [Kinzler, 1997;
Walter, 1998]. This appears not to be an intrinsic pressure
effect but instead results from the low Al content of near
solidus liquids at high pressure [Putirka, 1998b]. A decrease in
the partition coefficient for Ti amplifies the pressure dependence of Na/Ti and decreases estimates of initial melting
depths. Since the clinopyroxene-melt partition coefficients for
Ti observed by Kinzler [1997] and Walter [1998] are probably
most appropriate for modeling mantle melting, they were
adopted for this study.
A1.
Mantle Composition
One set of mantle mineral proportions used for peridotite
calculations (Table 1) are from Ita and Stixrude [1992, Figure
4]. The Ita and Stixrude [1992] estimates were obtained as a
best fit to seismic data for mineral elastic properties. Such
mineral modes are consistent with other similar studies of
upper mantle mineralogy [Weidner and Ito, 1987] and provide
a reasonable starting point for modeling an initial mantle mineralogy. Initial element concentrations for a peridotite mantle
are from Kinzler and Grove [1992b] and Hirschmann and Stolper [1996] for a depleted mantle and McDonough and Sun
[1995] for an undepleted source. Minerals compositions from
high-pressure experimental studies [Walter, 1998; Putirka,
1998a] were used to compare the major oxide [Kinzler and
Grove, 1992b] and mineral modes [Ita and Stixrude, 1992] of
Table 1. Mass balance calculations indicate overlap between
the major oxide composition and mineral proportions, but
these estimates are very sensitive to the mineral compositions
selected for mass balance. Using the Kinzler and Grove [1992b]
major element abundances, it is possible to obtain mineral
modes that have lower garnet (by 2%) and higher clinopyroxene (up to 6%) balanced by lower olivine; these mineralogys
were also explored for modeling purposes. A model mantle
with slightly higher clinopyroxene and lower garnet and olivine
abundances results in lower depth estimates for Hawaiian la-
2826
PUTIRKA: MELTING DEPTHS AND MANTLE HETEROGENEITY
vas. Such a mineralogy can also more easily explain the lowest
Na/Ti ratios observed at Mauna Kea, although lower-thanobserved Sm/Yb ratios are also obtained. The mineralogy of
Table 1 appears to be the most successful mineralogy for simultaneously explaining EPR and Hawaiian lava compositions. Model results are less sensitive to small changes in initial
Na/Ti ratio, and a range of depleted mantle starting compositions are successful at explaining major element variations in
oceanic basalts.
The peridotite bulk composition uncertainty (Figure 5) reflects the difference between depleted and undepleted sources
when these sources are input into the melting models. The
average garnet pyroxenite composition is from Hirschmann
and Stolper [1996], and their modal mineralogy is used. Eclogite compositions are from Griffin and Brueckner [1985], Griffiths and Cornichet [1985], Miller et al. [1988], Taylor and Neal
[1989], and Beard et al. [1992]; model calculations use a garnet
fraction of 0.25. Eclogite garnet contents vary widely (0.25–
0.50), and higher amounts of garnet yield partial melts with
higher Sm/Yb and Na/Ti ratios (Figure 5). Garnet pyroxenites
from massifs are from Loubet et al. [1982], Bodineir et al.
[1987], and Stosch and Lugmair [1990]; a garnet fraction of 0.2
was used for model calculations. As with eclogites, pyroxenites
feature a wide range of garnet fractions (0 –50%), and increased garnet yields increased Sm/Yb and Na/Ti ratios for
partial melts (Figure 5).
A2. The Mantle Liquidus and Solidus and Mantle
Adiabat
The mantle melting model uses a mantle adiabat of 1.2
K/GPa and liquidus and solidus curves from Zhang and Herzberg [1994]. Liquidus and solidus pressures were converted to
depth using the preliminary reference Earth model [Anderson,
1989].
The spinel-garnet transition was placed at 2.3 GPa [Gudfinnsson and Presnall, 1996; Kinzler, 1997]; a cusp on the solidus
at the spinel-garnet phase transition was not modeled. While a
change in slope must occur at the spinel-garnet transition, no
clear constraints on the magnitude of this change exist. A cusp
is not clearly detectable in the data of Takahashi et al. [1993]
for KLB-1, nor in the CaO-MgO-Al2O3-SiO2 (CMAS) system
[Gudfinnsson and Presnall, 1996], which may indicate that the
transition is not sharp. Any error introduced by ignoring small
cusps is also insignificant compared to error on the solidus
temperature.
Langmuir et al. [1992] showed that compositional differences
between batch and fractional melting may not be very great
when polybaric aggregate melts are compared. The present
calculations support this notion. The results of fractional and
batch melting models are similar, though batch melting results
compare somewhat better to the geochemical data (fractional
melting results in greater degrees of depletion for elements
with low partition coefficients), in agreement with Feigenson et
al. [1996]. Such differences, though, do not greatly impact P-T
estimates (compare the fractional melting curve in Figure 5a).
Additionally, estimates of depth are more robust than T since
K cpx/melt
is much more sensitive to P than T [Putirka, 1998b],
Na
and the K min/melt
for Ti, Hf, and REEs are expressed only as a
d
function of P [Salters, 1996]. Reported T estimates rely on
experimental measurements of solidus temperatures. Recently, solidus temperatures have been measured at high pressure for two similar peridotite nodules KR003 [Walter, 1998]
and KLB-1 [Zhang and Herzberg, 1994]. Unfortunately, other
peridotite solidi have not been measured with the same temperature precision (Walter [1998] and Zhang and Herzberg
[1994] account for thermal gradients in their experiments; see
Zhang and Herzberg [1994] for comparisons), and compositional effects on the mantle solidus cannot yet be quantified.
The data presented by Hirose and Kushiro [1993], though,
indicate that T solidus differences of up to 50 K are possible,
depending on mantle composition.
A3.
Melting Stoichiometry
Melting stoichiometry affects the estimation of mineral proportions throughout the melting column and thus impacts
upon D sol/melt
. Stoichiometric coefficients to depths of 0 –75
i
km are from J. Longhi [see Salters, 1996], Longhi [1995], and
Kinzler and Grove [1992a]. These coefficients (experimentally
determined at P , 3.5 GPa) might not extrapolate to high P.
With increased pressure, clinopyroxene and garnet survive to
greater extents of melting [Wei et al., 1990; Takahashi et al.,
1993]. In work by Takahashi et al [1993], clinopyroxene disappears at 38% melting at 3 GPa, while at 7 GPa clinopyroxene
survives to .70% melting. Less clinopyroxene must be consumed for each increment of melt produced at 7 GPa, and
melting coefficients should be reduced. At depths .75 km the
models (Figures 3–5) use Walter’s [1998] coefficients, although
lower coefficients were explored.
The amount of orthopyroxene on the mantle solidus is expected to change with depth since solid solution between orthopyroxene and clinopyroxene increases with increased temperature [Lindsley, 1983; Longhi and Bertka, 1996]. While
orthopyroxene appears in the products side of the melting
equilibrium at 3.0 –3.5 GPa [Longhi, 1995], orthopyroxene may
disappear from the melting equilibria at high pressure [Zhang
and Herzberg, 1994; Takahashi et al., 1993]. However, the precise P at which orthopyroxene disappears is unclear [see Canil,
1992; Walter, 1998; Bertka and Holloway, 1993]. Orthopyroxene
was treated as being absent from the melting interval at P .
6.5 GPa. Since orthopyroxene exerts little leverage on Na/Ti
or REE ratios, small errors regarding the placement of orthopyroxene in the melting interval do not affect model calculations.
Acknowledgments. I thank Charlie Langmuir for many helpful discussions concerning this project and for introducing to me some of the
problems concerning Na and Ti abundances in oceanic basalts. I also
greatly appreciate the reviews of Michael Garcia, Vincent Salters, and
Kenneth Cameron, which resulted in a much improved version of this
manuscript. Finally, I thank David Walker, Rosamond Kinzler, and
John Longhi for helpful comments and reviews of earlier versions of
this manuscript and Rick Ryerson for his support. UCRL-JC-131482
References
Ahern, J. L., and D. L. Turcotte, Magma migration beneath an ocean
ridge, Earth Planet. Sci. Lett., 45, 115–122, 1979.
Allegre, C. J., and D. L. Turcotte, Implications of a 2-component
marble cake mantle, Nature, 323, 123–127, 1986.
Anderson, D. L., Theory of the Earth, 366 pp., Blackwell Sci., Malden,
Mass., 1989.
Anderson, D. L., T. Tanimoto, and Y. Zhang, Plate tectonics and
hotspots: The third dimension, Science, 256, 1645–1651, 1992.
Atwater, T., Plate tectonic history of the northeast Pacific and western
North America, in The Geology of North America, vol. N, The Eastern
Pacific Ocean and Hawaii, edited by E. L. Winterer, D. M. Hussong,
and R. W. Decker, pp. 21–72, Geol. Soc. of Am., Boulder, Colo.,
1989.
Beard, B. L., L. G. Medaris, C. M. Johnson, H. Brueckner, and Z.
PUTIRKA: MELTING DEPTHS AND MANTLE HETEROGENEITY
Misar, Petrogenesis of Variscan high-temperature Group A
eclogites from the Moldanubian Zone of the Bohemian Massif,
Czechoslovakia, Contrib. Mineral. Petrol., 111, 468 – 483, 1992.
Bennett, V. C., T. M. Esat, and M. D. Norman, Two mantle-plume
components in Hawaiian picrites inferred from correlated Os-Pb
isotopes, Nature, 381, 221–223, 1996.
Bertka, C. M., and J. R. Holloway, Pigeonite at solidus temperatures:
Implications for partial melting, J. Geophys. Res., 98, 19,755–19,766,
1993.
Bhattacharyya, J., G. Masters, and P. Shearer, Global lateral variations
of shear wave attenuation in the upper mantle, J. Geophys. Res., 101,
22,273–22,289, 1996.
Blundy, J. D., T. J. Falloon, B. J. Wood, and J. A. Dalton, Sodium
partitioning between clinopyroxene and silicate melts, J. Geophys.
Res., 100, 15,501–15,515, 1995.
Bock, G., Long-period S to P converted waves and the onset of partial
melting beneath Oahu, Hawaii, Geophys. Res. Lett., 18, 869 – 872,
1991.
Bodineir, J. L., M. Guiraud, J. Fabries, J. Dostal, and C. Dupuy,
Petrogenesis of layered pyroxenites from the Lherz, Freychinede
and Prades ultramafic bodies (Ariege, Grench Pyrenees), Geochim.
Cosmochim. Acta, 51, 279 –290, 1987.
Budahn, J. R., and R. A. Schmitt, Petrogenetic modeling of Hawaiian
tholeiitic basalts: A geochemical approach, Geochim. Cosmochim.
Acta, 49, 67– 87, 1985.
Canil, D., Orthopyroxene stability along the peridotite solidus and the
origin of cratonic lithosphere beneath southern Africa, Earth Planet.
Sci. Lett., 111, 83–95, 1992.
Cawthorn, R. G., Degrees of melting in mantle diapirs and the origin
of ultrabasic liquids, Earth Planet. Sci. Lett., 27, 113–120, 1975.
Chen, C. Y., Lead isotopic constraints on the origin of Hawaiian
basalts, Nature, 327, 49 –52, 1987.
Chen, C. Y., and F. A. Frey, 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, 1985.
Chen, C. Y., F. A. Frey, J. M. Rhodes, and R. M. Easton, Temporal
geochemical evolution of Kilauea Volcano: Comparison of Hilina
and Puna Basalt, in Earth Processes: Reading the Isotopic Code,
Geophys. Monogr. Ser., vol. 95, edited by S. Hart and A. Basu, pp.
161–181, AGU, Washington, D. C., 1996.
Detrick, R. S., and T. S. Crough, Island subsidence, hot spots, and
lithospheric thinning, J. Geophys. Res., 83, 1236 –1244, 1978.
Eggins, S. M., Petrogenesis of Hawaiian tholeiites, 1, Phase equilibria
constraints, Contrib. Mineral. Petrol., 110, 387–397, 1992.
Farber, D. L., Q. Williams, and F. J. Ryerson, Diffusion in Mg2SiO4
polymorphs and chemical heterogeneity in the mantle transition
zone, Nature, 371, 693– 695, 1994.
Farnetani, D. G., and M. A. Richards, Thermal entrainment and melting in mantle plumes, Earth Planet. Sci. Lett., 136, 251–267, 1995.
Feigenson, M. D., L. C. Patino, and M. J. Carr, Constraints on partial
melting imposed by rare earth element variations in Mauna Kea
basalts, J. Geophys. Res., 101, 11,815–11,829, 1996.
Frey, F. A., and D. A. Clague, Geochemistry of diverse basalt types
from Loihi Seamount, Hawaii: Petrogenetic implications, Earth
Planet. Sci. Lett., 66, 337–355, 1983.
Frey, F. A., M. O. Garcia, W. S. Wise, A. Kennedy, P. Gurriet, and F.
Albarede, The evolution of Mauna Kea Volcano, Hawaii: Petrogenesis of tholeiitic and alkalic basalts, J. Geophys. Res., 96, 14,347–
14,375, 1991.
Frey, F. A., M. O. Garcia, and M. F. Roden, Geochemical characteristics of Koolau Volcano: Implications of inter-shield geochemical
differences among Hawaiian volcanoes, Geochim. Cosmochim. Acta,
58, 1441–1462, 1994.
Garcia, M. O., J. M. Rhodes, E. W. Wolfe, G. E. Ulrich, and R. A. Ho,
Petrology of lavas from episodes 2– 47 of the Pu’u O’o eruption of
Kilauea Volcano, Hawaii: Evaluation of magmatic processes, Bull.
Volcanol., 55, 1–16, 1992.
Garcia, M. O., B. A. Jorgenson, J. J. Mahoney, E. Ito, and A. J. Irving,
An evaluation of temporal geochemical evolution of Loihi summit
lavas: Results from Alvin submersible dives, J. Geophys. Res., 98,
537–550, 1993.
Garcia, M. O., D. J. P. Foss, H. B. West, and J. J. Mahoney, Geochemical and isotopic evolution of Loihi volcano, Hawaii, J. Petrol.,
36, 1647–1674, 1995.
Griffin, W. L., and H. K. Brueckner, REE, Rb-Sr and Sm-Nd studies
of Norwegian eclogites, Chem. Geol., 52, 249 –271, 1985.
2827
Griffiths, J. B., and J. Cornichet, Origin of eclogites from south Brittany, France: A Sm-Nd isotopic study, Chem. Geol., 52, 185–201,
1985.
Grove, T. L., and W. B. Bryan, Fractionation of pyroxene-phyric
MORB at low pressure: An experimental study, Contrib. Mineral.
Petrol., 84, 293–309, 1983.
Gudfinnsson, G. H., and D. C. Presnall, Melting relations of model
lherzolite in the system CaO-MgO-Al2O3-SiO2 at 2.4 –3.4 GPa and
the generation of komatiites, J. Geophys. Res., 101, 27,701–27,709,
1996.
Hauri, E. H., Major-element variability in the Hawaiian mantle plume,
Nature, 382, 415– 419, 1996.
Hauri, E. H., J. C. Lassiter, and D. J. Depaolo, Osmium isotope
systematics of drilled lavas from Mauna Loa, Hawaii, J. Geophys.
Res., 101, 11,793–11,806, 1996.
Hekinian, R., G. Thompson, and D. Bideau, Axial and off-axial heterogeneity of basaltic rocks from the East Pacific Rise at 128359N–
128519N and 118269N–118309N, J. Geophys. Res., 94, 17,437–17,463,
1989.
Hirose, K., and I. Kushiro, Partial melting of dry peridotites at high
pressures: Determination of compositions of melts segregated from
peridotite using aggregates of diamond, Earth Planet. Sci. Lett., 114,
477– 490, 1993.
Hirschmann, M. M., and E. M. Stolper, A possible role for garnet
pyroxenite in the origin of the “garnet signature” in MORB, Contrib.
Mineral. Petrol., 124, 185–208, 1996.
Hofmann, A. W., and W. M. White, Mantle plumes from ancient
oceanic-crust, Earth Planet. Sci. Lett., 57, 421– 436, 1982.
Hofmann, A. W., K. P. Jochum, M. Seufert, and W. M. White, Nb and
Pb in oceanic basalts: New constraints on mantle evolution, Earth
Planet. Sci. Lett., 79, 33– 45, 1986.
Ita, J., and L. Stixrude, Petrology, elasticity, and composition of the
mantle transition zone, J. Geophys. Res., 97, 6849 – 6866, 1992.
Jackson, E. D., and H. R. Shaw, Stress fields in central portions of the
Pacific plate: Delineated in time by linear volcanic chains, J. Geophys. Res., 80, 1861–1874, 1975.
Kellogg, L. H., Mixing in the mantle, Annu. Rev. Earth Planet. Sci., 20,
365–388, 1992.
Kinzler, R., Melting of mantle peridotite at pressures approaching the
spinel to garnet transition: Application to mid-ocean ridge basalt
petrogenesis, J. Geophys. Res., 102, 853– 874, 1997.
Kinzler, R. J., and T. L. Grove, Primary magmas of mid-ocean ridge
basalts, 1, Experiments and methods, J. Geophys. Res., 97, 6885–
6906, 1992a.
Kinzler, R. J., and T. L. Grove, Primary magmas of mid-ocean ridge
basalts, 2, Applications, J. Geophys. Res., 97, 6907– 6926, 1992b.
Klein, E. M., and C. H. Langmuir, Global correlations of ocean ridge
basalt chemistry with axial depth and crustal thickness, J. Geophys.
Res., 92, 8089 – 8115, 1987.
Langmuir, C. H., Petrology data base, vols. II and III, in East Pacific
Rise Data Synthesis Final Report, edited by S. Tingh, pp. 183–278,
Joint Oceanogr. Inst. Inc., Washington, D. C., 1988.
Langmuir, C. H., and G. N. Hanson, An evaluation of major-element
heterogeneity in the mantle sources of basalts, Philos. Trans. R. Soc.
London Ser. A, 297, 383– 407, 1980.
Langmuir, C. H., J. F. Bender, A. E. Bence, and G. N. Hanson,
Petrogenesis of basalts from the FAMOUS area: Mid-Atlantic
Ridge, Earth Planet. Sci. Lett., 36, 133–156, 1977.
Langmuir, C. H., J. F. Bender, and R. Batiza, Petrological and tectonic
segmentation of the East Pacific Rise, 58309–148309N, Nature, 322,
422– 429, 1986.
Langmuir, C. H., E. M. Klein, and T. Plank, Petrological systematics of
mid-ocean ridge basalts: Constraints on melt generation beneath
ocean ridges, in Mantle Flow and Melt Generation at Mid-Ocean
Ridges, Geophys. Monogr. Ser., vol. 71, edited by J. Phipps Morgan,
D. K. Blackman, and J. M. Sinton, pp. 183–280, AGU, Washington,
D. C., 1992.
Lanphere, M., 87Sr/86Sr ratios for basalt from Loihi Seamount, Hawaii, Earth Planet. Sci. Lett., 66, 380 –387, 1983.
Lindsley, D. H., Pyroxene thermometry, Am. Mineral., 68, 477– 493,
1983.
Liu, M., and C. G. Chase, Evolution of hotspot swells: Numerical
solutions, J. Geophys. Res., 94, 5571–5584, 1989.
Longhi, J., Liquidus equilibria of some primary lunar and terrestrial
melts in the garnet stability field, Geochim. Cosmochim. Acta, 59,
2375–2386, 1995.
2828
PUTIRKA: MELTING DEPTHS AND MANTLE HETEROGENEITY
Longhi, J., and C. M. Bertka, Graphical analysis of pigeonite-augite
liquidus equilibria, Am. Mineral., 81, 685– 695, 1996.
Loubet, M., and C. J. Allegre, Trace elements in orogenic lherzolites
reveal the complex history of the upper mantle, Nature, 298, 809 –
814, 1982.
Mammerickx, J., Large-scale undersea features of the northeast Pacific, in The Geology of North America, vol. N, The Eastern Pacific
Ocean and Hawaii, edited by E. L. Winterer, D. M. Hussong, and
R. W. Decker, pp. 5–13, Geol. Soc. of Am., Boulder, Colo., 1989.
McDonough, W. F., and S. Sun, The composition of the Earth, Chem.
Geol., 120, 223–253, 1995.
MELT Seismic Team, Imaging the deep seismic structure beneath a
mid-ocean ridge: The MELT experiment, Science, 280, 1215–1218,
1998.
Miller, C., H. G. Stosch, and S. Hoernes, Geochemistry and origin of
eclogites from the type locality Koralpe and Saulpe, eastern Alps,
Austria, Chem. Geol., 67, 103–118, 1988.
Morgan, J. P., W. J. Morgan, and E. Price, Hotspot melting generates
both hotspot volcanism and a hotspot swell?, J. Geophys. Res., 100,
8045– 8062, 1995.
Morgan, J. W., Plate motions and deep mantle convection, Mem. Geol.
Soc. Am., 132, 7–22, 1972.
Okano, O., and M. Tatsumoto, Petrogensis of ultramafic xenoliths
from Hawaii inferred from Sr, Nd and Pb isotopes, in Earth Processes: Reading the Isotopic Code, Geophys. Monogr. Ser., vol. 95,
edited by S. Hart and A. Basu, pp. 135–147, AGU, Washington,
D. C., 1996.
Oxburgh, E. R., and D. L. Turcotte, Mid-ocean ridges and geotherm
distribution during mantle convection, J. Geophys. Res., 73, 2643–
2661, 1968.
Parsons, B., and J. G. Sclater, An analysis of the variation of ocean
floor bathymetry and heat flow with age, J. Geophys. Res., 82, 803–
827, 1977.
Plank, T., and C. H. Langmuir, Effects of the melting regime on the
composition of the oceanic crust, J. Geophys. Res., 97, 19,749 –
19,770, 1992.
Prinzhofer, A., E. Lewin, and C. J. Allegre, Stochastic melting of the
marble-cake mantle: Evidence from local study of the EPR at 12850
N, Earth Planet. Sci. Lett., 92, 189 –206, 1989.
Putirka, K., Melt productivity during fractional melting and the apparent conflict between inferred melting depths and crustal thickness,
Eos Trans. AGU, 78(46), Fall Meet Suppl., F837–F838, 1997.
Putirka, K., Garnet 1 liquid equilibrium, Contrib. Mineral. Petrol., 131,
273–288, 1998a.
Putirka, K., Clinopyroxene 1 liquid equilibrium to 100 kbar and 2450
K, Contrib. Mineral. Petrol., in press, 1998b.
Reynolds, J. R., C. H. Langmuir, J. F. Bender, K. A. Kastens, and
W. B. F. Ryan, Spatial and temporal variability in the geochemistry
of basalts from the East Pacific Rise, Nature, 359, 493– 499, 1992.
Rhodes, J. M., Geochemistry of the 1984 Mauna Loa eruption: Implications for magma storage and supply, J. Geophys. Res., 93, 4453–
4466, 1988.
Rhodes, J. M., Geochemical stratigraphy of lava flows sampled by the
Hawaii scientific drilling project, J. Geophys. Res., 101, 11,729 –
11,746, 1996.
Rhodes, J. M., and S. R. Hart, Episodic trace element and isotopic
variations in historical Mauna Loa lavas: Implications for magma
and plume dynamics, in Mauna Loa Revealed, Geophys. Monogr. Ser.,
vol. 92, edited by J. M. Rhodes and J. P. Lockwood, pp. 263–288,
AGU, Washington, D. C., 1995.
Ribe, N. M., and U. R. Christensen, Three-dimensional modeling of
plume-lithosphere interaction, J. Geophys. Res., 99, 669 – 682, 1994.
Richardson, S. H., A. J. Erlank, A. R. Duncan, and D. L. Reid,
Correlated Nd, Sr and Pb isotope variation in Walvis Ridge basalts
and implications for the evolution of their mantle source, Earth
Planet. Sci. Lett., 59, 327–342, 1982.
Roden, M. F., T. Trull, S. R. Hart, and F. A. Frey, New He, Nd, Pb,
and Sr constraints on the constitution of the Hawaiian plume: Results from Koolau Volcano, Oahu, Hawaii, USA, Geochim. Cosmochim. Acta, 58, 1431–1440, 1994.
Romanowicz, B., Anelastic tomography: A new perspective on upper
mantle thermal structure, Earth Planet. Sci. Lett., 128, 113–121, 1994.
Romanowicz, B., A global tomographic model of shear attenuation in
the upper mantle, J. Geophys. Res., 100, 12,375–12,394, 1995.
Salters, V. J. M., The generation of mid-ocean ridge basalts from the
Hf and Nd isotope perspective, Earth Planet. Sci. Lett., 141, 109 –123,
1996.
Salters, V. J. M., and S. R. Hart, The Hf-paradox and the role of garnet
in the source of mid-ocean ridge basalts, Nature, 342, 420 – 422, 1989.
Salters, V. J. M., and S. R. Hart, The mantle sources of ocean ridges,
islands and arcs: The Hf-isotope connection, Earth Planet. Sci. Lett.,
104, 364 –380, 1991.
Schilling, J. G., Fluxes and excess temperatures of mantle plumes
inferred from their interaction with migrating mid-ocean ridges,
Nature, 352, 397– 403, 1991.
Sims, K. W. W., and D. J. Depaolo, Inferences about mantle magma
sources from incompatible element concentration ratios in oceanic
basalts, Geochim. Cosmochim. Acta, 61, 765–784, 1997.
Sleep, N. H., Hotspots and mantle plumes: Some phenomenology, J.
Geophys. Res., 95, 6715– 6736, 1990.
Sleep, N. H., M. A. Richards, and M. A. Hager, Onset of mantle
plumes in the presence of preexisting convection, J. Geophys. Res.,
93, 7672–7689, 1988.
Spiegelman, M., Geochemical consequences of melt transport in 2-D:
The sensitivity of trace-elements to mantle dynamics, Earth Planet.
Sci. Lett., 139, 115–132, 1996.
Staudigel, H., A. Zindler, S. R. Hart, T. Leslie, C. Y. Chen, and D.
Clague, The isotope 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, 1984.
Stein, C. A., and S. Stein, A model for the global variation in oceanic
depth and heat flow with lithospheric age, Nature, 359, 123–129,
1992.
Stille, P., D. M. Unruh, and M. Tatsumoto, Pb, Sr, Nd and Hf isotopic
evidence of multiple sources for Oahu, Hawaii basalts, Nature, 304,
25–29, 1983.
Stosch, H. G., and G. W. Lugmair, Geochemistry and evolution of
MORB-type eclogites from the Munchberg Massif, southern Germany, Earth Planet. Sci. Lett., 99, 230 –249, 1990.
Takahashi, E., T. Shimazaki, Y. Tsuzaki, and H. Yoshida, Melting
study of a peridotite KLB-1 to 6.5 GPa, and the origin of basaltic
magmas, Philos. Trans. R. Soc. London Ser. A, 342, 105–120, 1993.
Taylor, L. A., and C. R. Neal, Eclogites with oceanic crustal and
mantle signatures from the Bellsbank Kimberlite, South Africa, 1,
Mineralogy, petrography, and whole rock chemistry, J. Geol., 97,
551–567, 1989.
Tormey, D. R., T. L. Grove, and W. B. Bryan, Experimental petrology
of normal MORB near the Kane Fracture Zone: 228–258N, MidAtlantic Ridge, Contrib. Mineral. Petrol., 96, 121–139, 1987.
VonHerzen, R. P., M. J. Cordery, R. S. Detrick, and C. Fang, Heat
flow and the thermal origin of hot spot swells: The Hawaiian swell
revisited, J. Geophys. Res., 94, 13,783–13,799, 1989.
Wagner, T. P., and T. L. Grove, Melt/Harzburgite reaction in the
petrogenesis of tholeiitic magma from Kilauea Volcano, Hawaii,
Contrib. Mineral. Petrol., 131, 1–12, 1998.
Walter, M. J., Melting of garnet peridotite and the origin of komatiite
and depleted lithosphere, J. Petrol., 39, 29 – 60, 1998.
Watson, S., and D. McKenzie, Melt generation by plumes: A study of
Hawaiian volcanism, J. Petrol., 32, 501–537, 1991.
Wei, K., R. G. Tronnes, and C. M. Scarfe, Phase relations of aluminum-undepleted and aluminum-depleted komatiites at pressures of
4 –12 GPa, J. Geophys. Res., 95, 15,817–15,827, 1990.
Weidner, D. J., and E. Ito, Mineral physics constraints on a uniform
mantle composition, in High Pressure Research in Mineral Physics,
Geophys. Monogr. Ser., vol. 39, edited by M. H. Manghnani and Y.
Syono, pp. 439 – 446, AGU, Washington, D. C., 1987.
Wessel, P., Observational constraints on models of the Hawaiian hot
spot swell, J. Geophys. Res., 98, 16,095–16,104, 1993.
White, R. S., D. McKenzie, and K. Onions, Oceanic crustal thickness
from seismic measurements and rare earth element inversions, J.
Geophys. Res., 97, 19,683–19,715, 1992.
Wilson, J. T., A possible origin of the Hawaiian islands, Can. J. Phys.,
41, 863– 870, 1963.
Woods, M. T., and E. A. Okal, Rayleigh-wave dispersion along the
Hawaiian Swell: A test of lithosphere thinning by thermal rejuvenation at a hotspot, Geophys. J. Int., 125, 325–339, 1996.
Xu, Y., M. A. Menzies, P. Vroon, J.-C. Mercier, and C. Lin, Texturetemperature-geochemistry relationships in the upper mantle as revealed from spinel xenoliths from Wangqing, NE China, J. Petrol.,
39, 469 – 493, 1998.
Yang, H. J., F. A. Frey, J. M. Rhodes, and M. O. Garcia, Evolution of
PUTIRKA: MELTING DEPTHS AND MANTLE HETEROGENEITY
Mauna Kea volcano: Inferences from lava compositions recovered
in the Hawaii scientific drilling project, J. Geophys. Res., 101, 11,747–
11,767, 1996.
Zhang, J., and C. Herzberg, Melting experiments on anhydrous peridotite KLB-1 from 5.0 to 22.5 GPa, J. Geophys. Res., 99, 17,729 –
17,742, 1994.
Zhang, Y.-S., and T. Tanimoto, High-resolution global upper mantle
structure and plate tectonics, J. Geophys. Res., 98, 9793–9824, 1993.
Zindler, A., H. Staudigel, and R. Batiza, Isotope and trace element
geochemistry of young Pacific seamounts: Implications for the scale
2829
of upper mantle heterogeneity, Earth Planet. Sci. Lett., 70, 175–195,
1984.
K. Putirka, Lawrence Livermore National Laboratory, L-202, 7000
E. Ave., Livermore, CA 94550. ([email protected])
(Received February 9, 1998; revised September 30, 1998;
accepted October 7, 1998.)
2830