1
Low degree melting under the Southwest Indian Ridge: The roles of mantle
temperature, conductive cooling and wet melting
C.J. Robinson1, M.J. Bickle1, T.A. Minshull2 & R.S. White3 and A R.L. Nichols4
1
Dept. of Earth Sciences, Downing St., Cambridge, CB2 3EQ, UK.
Phone Number: +44 (0)1223 333400, Fax: +44 (0)1223 333450
2
Now at School of Ocean and Earth Science, University of Southampton, Southampton
Oceanography Centre, European Way, Southampton, SO14 3HZ, UK.
Phone Number: +44 (0)23 80596569, Fax: +44 (023) 8059 3052
3
Bullard Laboratories, Madingley Rise, Madingley Rd., Cambridge, CB3 0EZ, UK.
Phone Number: +44 (0)1223 337191, Fax: +44(0)1223 360779
4
Dept. of Earth Sciences, Wills Memorial Building, Queen's Road, Clifton, Bristol, BS8
1RJ, UK. Phone Number: +44(0)117 9545400, Fax: +44(0)117 9253385
Correspondence to C.J. Robinson
Fax: +44 (0)1223 333450
Email: cjr17@esc.cam.ac.uk
mb72@esc.cam.ac.uk, rwhite@esc.cam.ac.uk, tmin@mail.soc.soton.ac.uk,
Alex.Nichols@bristol.ac.uk
Word count: 5473
2
Abstract
Both low mantle temperatures and conductive cooling have been suggested as the cause
of the atypically thin oceanic crust and the incompatible element enrichment characteristic of
very slow-spreading ridges. Here we present a model of melting under the Southwest Indian
Ridge, which takes into account mantle temperature, conductive cooling, source composition
and wet melting. The model parameters are constrained by oceanic crustal thickness, lava
chemistry and isotopic composition and water content. The results suggest that conductive
cooling to a depth of around 20 km, expected in areas with a full spreading rate of 15 mm/yr,
is necessary to generate the Southwest Indian Ridge lava chemistry, but not that from faster
spreading rate ridges at 23°N on the Mid Atlantic Ridge or 45°N on the Juan de Fuca Ridge.
The mantle potential temperatures of ~1280°C, estimated for the Southwest Indian Ridge
lavas are close to the global average of the upper mantle. Mantle water contents of 150-300
ppm can explain the observed melt water contents and allow sufficient melting at depth to
explain the observed heavy rare earth element depletions in the melts.
Introduction
Oceanic spreading ridges fed by a relatively uniform composition mantle and largely
unaffected by conductive cooling, have long been the environment of choice for investigating
melting processes in the mantle [e.g. 1-7]. The volume and chemistry of mantle melting
beneath spreading centres are consistent with the melting being near-adiabatic, and taking
place, albeit at variable rates, between ~ 60 km depth and the base of the oceanic crust. The
prime control on variations in oceanic crustal thickness and major element chemistry is the
potential temperature of the mantle in the melt region [5,8]. Very slow-spreading ridges such
as the Southwest Indian Ridge (SWIR), the Arctic Ridge and the Mid-Cayman Rise lie at the
thin crust, low degree of melt end of trends that relate crustal thickness to chemical indices [6]
(Fig. 1). Although Langmuir et al. [6] saw no correlation between the degree of melting and
the spreading rate, Reid and Jackson [9], Bown and White [10], Shen & Forsyth [11] and Niu
& Hekinian [12] calculate that conductive cooling from above may restrict the top of the
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melting zone and thereby reduce melting under ridges spreading at rates slower than 15-20
mm/yr. In these environments, the low amounts of melting may therefore result from slowspreading rates rather than atypically low mantle temperatures. If oceanic crustal thicknesses
are to be used as a measure of the temperature structure of the mantle [cf. 7], it is important to
identify examples where melting is reduced by other factors. Furthermore, in low melt
production environments, the amount and chemistry of melts may be sensitive to a number of
compositional factors in addition to mantle temperature. In this study of melting under the
SWIR, we use the seismic crustal thickness, major element, REE and water contents, as well
as the isotopic compositions of volcanic glasses, to constrain a simple physical model of the
melting which takes conductive cooling and wet melting into account. The results indicate
that it is the very slow-spreading rates rather than low mantle temperatures, which account for
the reduced magmatism at the SWIR.
Previous geochemical studies of the SWIR have largely concentrated on anomalous areas
such as those affected by mantle plumes [e.g. 13]. Geophysical studies have mostly been
conducted in the triple junction regions, although bathymetric and gravity mapping is now
more widespread [14-16]. There have been few crustal thickness measurements made on
Southwest Indian ocean crust and prior to this combined study [17], none along the ridge axis.
Geology
Two areas were dredged along the axis of the SWIR and a third area was dredged ~10 Ma
off-axis during the RRS Discovery Cruise 208 (Fig. 2). The ‘66°E’ study area is between
65°41’E and 66°19’E on-axis, where 7 dredge hauls sample an area over 80km in length. The
‘57°E’ area is between 57° 28.7’ E and 57° 44.2’ E, where 5 dredge hauls sample an area over
26 km in length. The off-axis area (‘site 3’) is at 31°21.8’S, 57°15.3’E and was chosen in
order to represent the extrusive section of the rocks drilled from the Atlantis Platform [18].
There was only one successful dredge haul in the site 3 area.
Wide-angle seismic surveys were carried out coincident with the 66°E sampling area and
around the Atlantis Platform (Fig. 2) [19]. No seismic survey was carried out on-axis at the
57°E area, except well off-axis crossing the Atlantis Platform, on 11Ma crust [20]. The 57°E
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samples and this seismic line are not used in the modelling because there is little reason to
believe that the crustal thickness measurement is representative of the melt production on-axis
at the present time [21]. However, the 57°E data are used to provide constraints on
fractionation trends for the site 3 data. The crustal thickness measurements obtained from the
seismic profiles suggest that the crustal thickness associated with the site 3 sample site is 4.0
± 1.0 km [18]. The crustal thickness under the 66°E area averaged across three spreading
segments is 4.3 ± 1.0 km [16].
Analytical methods
We picked volcanic glasses that were visibly alteration-free. Major element
concentrations were determined using the Cambridge Cameca SX50 electron microprobe
using a combination of energy and wavelength dispersive spectrometry [22]. Analysis of the
standard glass USNM 111240/91, interspersed with the unknown samples, was always within
error of the published values [23]. Accuracy and precision are better than 3% for the major
elements including Na2O. The rare earth element analyses were obtained using a PlasmaQuad
PQ2+ at the NERC ICPMS facility [24]. Sample SNS-7 [25] was used as a reference
standard. Reproducibility and accuracy of the REE are better than ± 10%. The standard
deviation in each REE in the sets of samples averaged for the REE inversions below, averages
65% and the analytical uncertainty is small compared with between-sample heterogeneities.
Nd isotope data were determined by thermal ionization mass spectrometry using the T40
VG sector 54 multicollector mass spectrometer at the University of Cambridge and chemical
separation and mass-spectrometric techniques described by Ahmad et al. [26]. The
fractionation was normalized to 146Nd/144Nd = 0.7219. The precision in 143Nd/144Nd is better
than 0.000015 (2). Nd blanks were < 1 part in 104 of the sample. Nd values are calculated
with respect to 143Nd/144Ndchur = 0.512638.
Water contents of the glasses (Table 1) were measured by Fourier Transform Infrared
analysis at the University of Bristol by Alex Nichols. The method was as described in Dixon
et al. [27], using the same peak (3550 cm-1) and molar absorptivity coefficient (63 l/mol.cm),
except a Nicolet Model 800 FTIR spectrometer was used with a 100 micron diameter aperture
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and only 256 scans were performed. Repeat analyses of individual glasses indicate analytical
precisions of ± 0.02 wt%.
Fractionation
Ocean ridge basalt samples undergo variable amounts of low-pressure fractionation and a
number of methods have been used to normalise elemental data to allow for comparison with
melting models [e.g. 8]. We have calculated compositions adjusted to 8 wt% MgO, for
comparison with data sets compiled by Langmuir et al. [6] then to Mg# 70 (10-11wt% MgO)
for comparison with primary magma compositions, assuming that primary magmas are in
equilibrium with mantle olivines of Fo90. The adjustments were applied to the whole data
sets.
Glass compositions (Table 1) were adjusted to 8 wt% MgO, using linear regressions
through element against MgO plots (e.g. Fig 3). The site 3 data show very limited major
element dispersion, so were corrected using linear regressions through the 57°E data. The data
are insensitive to the correction because their measured MgO contents are ~7-7.5 wt% MgO.
The calculated linear corrections correspond to assemblages of olivine + plagioclase +
clinopyroxene in the ratios 0.21:0.50:0.29 for the 57°E and site 3 data and to olivine-only for
the 66°E data. The same assemblages fit the trace element data well. The adjustment of
magma compositions to Mg# 70 is made by assuming that olivine alone crystallized until they
reached 8wt% MgO. The adjustment is made by calculating equilibrium olivine compositions
using the partition coefficients from Beattie et al. [28] and Beattie [29] and adding this to the
lava compositions in increments of 0.5% until Mg# 70 (10-11wt% MgO) is reached. The
assumption of olivine-only fractionation between Mg# 70 and 8 wt% MgO is believed to be
reasonable. Firstly, according to the observed mineralogy, the site 3 and 66°E samples are
saturated only with olivine (± spinel). Secondly, the calculated phase appearances (using
MELTS, a liquid line of descent program [30]) over a range of pressures [see 21], coincide
with this observation. Thirdly, the adjustments made to the data are quite small. Spinel
fractionation has no affect on the MgO contents.
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Additional data used in this paper are from 23°N on the Mid Atlantic Ridge [31] and
45°N on the Juan de Fuca ridge [32]. Analyses with MgO > 8 wt% and Mg# > 60 were
selected and adjusted for olivine fractionation as above.
Structure of the Melting Region
The structure of decompression melting regimes is determined by the potential
temperature of the underlying mantle, the extent of conductive cooling from the surface, the
mantle solidus and the variation of melt fraction with height in the melting region, itself a
function of the mechanism of two phase flow in the melt region (Fig. 4). Significant degrees
of melting are only encountered above the dry peridotite solidus, which probably varies little
in pressure-temperature space for the range of upper mantle compositions sampled by oceanic
spreading ridges. The uniform, segment-averaged thickness of most oceanic crust spreading
faster than 20 mm/yr attests to the relative uniformity of both mantle composition and mantle
potential temperature [e.g. 7]. Small degrees of melting fluxed by the minor amount of water
in the mantle must occur at depths below the dry solidus. Addition of these small-volume, wet
melts may be important for incompatible element concentrations in oceanic basalts, their
relative contributions being proportionally more important in regimes with lower degrees of
melting. Another important uncertainty about processes in the melting region is the extent to
which fluid-solid reactions high in the melt region change the chemistry of melts generated at
depth [33,34]. In slow-spreading regimes, if there is a thick zone of conductively-cooled
mantle above the melt region [e.g. 12], such melt-mantle reactions are likely to be enhanced.
The amount of melting as reflected by the thickness of the oceanic crust and the relative
enrichment of incompatible elements such as Na8 does not distinguish between control by
mantle temperature or conductive cooling. The main concerns of this paper are the extent to
which the chemistry of the lavas can be used to distinguish whether reduced melting at the
SWIR reflects conductive cooling or reduced mantle temperatures and how important small
degrees of wet melting below the dry solidus are for the lava chemistry.
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Chemistry of Mantle Melts at the SWIR
Major and trace element compositions of the SWIR lavas are consistent with their being
low degree melts. Na8 values of the site 3 lavas are around 3.4 wt% whereas the 66°E lavas
range between 3.7 and 4.1 wt% (Fig. 1). The lavas lie at the thin crust, high Na8 end of the
global correlations between Na8 and crustal thickness or Fe8 [6]. The 66°E lavas lie at
significantly higher Na8 for their seismic crustal thickness than the global trend. Rare earth
elements are 20 – 90 % more enriched in the site 3 lavas and by up to 120% in the 66°E lavas,
relative to average ‘normal’ MORB [35]. The more marked enrichment in light rare earth
elements (LREE) relative to heavy rare earths (HREE) is interpreted below to indicate some
melting within the stability field of garnet. Similar enrichments relative to N-MORB are also
seen amongst the other incompatible trace elements in the site 3 samples. However,
enrichments of up to a factor of 4 are seen in the incompatible trace elements from 66°E.
Values of Ca/Al are 80 to 90% lower than average ‘normal’ mid-ocean ridge basalt in the site
3 and 66°E lavas, respectively.
The important question is whether the chemical compositions of the lavas can be used to
determine whether the low degrees of melting reflects 1) cool mantle and thus melting that is
restricted to shallow depths or 2) conductive cooling with melting correspondingly taking
place at higher pressures. Langmuir at al [6] calculated incremental decompression melting
curves assuming both equilibrium and fractional melting and showed a systematic increase in
Fe contents with the depth of solidus intersection. As mantle rises above the solidus, the
increase in degree of melting offsets the decrease in pressure of melting such that the
calculated Fe content of the melts changes little (Fig. 1). Langmuir et al. [6] interpret the
global data trend on the Na8:Fe8 diagram (Fig.1) to result from the correlation between the
amount of melting ( 1/Na8) and the increased mean depth of melting ( Fe8) in response to
higher mantle temperatures. The site 3 (and 57°E) lavas plot on or just above the global trend
with mean Fe8 of ~8.7 wt% and Na8 of 3.4 wt% compared with average normal-MORB [35]
which has between 8.5 and 10 wt% Fe8 and 3.0 to 2.6 wt% Na8. The 66°E lavas plot at higher
Na8 and lower Fe8 (~3.9 and 7.5 wt% respectively). An interpretation of these data is that the
SWIR lavas are produced at shallower depths caused by cooler mantle temperatures. If so,
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melting must have continued up to near the base of the oceanic crust or melt thicknesses
would be substantially less than observed. However melting up to the base of the oceanic
crust is inconsistent with the thermal calculations, which suggest significant conductive
cooling. It is also inconsistent with the REE patterns, which show evidence for melting in the
garnet stability field. One potential problem with interpretation of major, compatible elements
such as Fe is that shallow melt-mantle reactions may modify melts during transport. Below,
we use the REE compositions and water contents of the lavas to constrain decompression
melting models which take mantle temperature, spreading-rate conductive cooling and wet
melting into account.
A Decompression Melting Model
Rare earth element concentrations are used in combination with lava water contents,
source enrichment (constrained by 143Nd/144Nd ratios) and seismic crustal thicknesses, to
constrain mantle potential temperature and mantle water content and to investigate whether
the inclusion of conductive cooling significantly improves the fit of the model.
We use a decompression melting model to calculate melt thickness, melt water content
and melt REE concentrations as a function of mantle potential temperature, mantle water
content and spreading-rate dependent conductive cooling. The model results are evaluated by
the misfit between the predicted and observed REE concentrations, calculated melt
thicknesses compared with ocean crustal thickness and the calculated and measured melt
water contents. The REE composition of the mantle source is calculated from the observed
lava Nd in terms of a mixture between ‘primitive’ (Nd = 0) and ‘depleted’ (Nd = 10) mantle
(using the REE enrichments and Nd-isotopic compositions of McKenzie & O’Nions [36]).
Nd values for the Mid-Atlantic Ridge and Juan de Fuca Ridge areas are 11.9 and 10.2
respectively [37,38], and values for the SWIR areas are given in Table 1. In all three areas,
Nd values are close to the depleted mantle end-member [36] and the maximum fraction of
primitive mantle is 25% assumed for the source of the 66°E lavas. A set of models have been
run with conditions of 1) no conductive cooling and 2) conductive cooling using a mantle
upwelling rate of 7.5 mm/yr (assuming a 15 mm/yr full spreading rate for the SWIR [39]) and
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of 13 mm/yr for the Mid-Atlantic Ridge samples. The wedge angle (Fig. 4a) is taken as
45°, but it should be noted that constraints on the wedge angle are poor [e.g. 40].
A forward modelling approach is taken and a graphical portrayal of the results is used to
compare the results of the REE modelling, predicted melt thickness and melt water content
with those observed. The model depends critically on assumptions of source REE enrichment,
mineral-melt partition coefficients, the very approximate treatment of wet melting and the
geometry of the upwelling flow (wedge angle). As a result, the main value of the results is to
illustrate their sensitivity to the main controls on the melt regime. These are the mantle
potential temperature, the significance of conductive cooling and the mantle water content.
The decompression melting model calculates the change of melt fraction with height. It
makes essentially the same assumptions as McKenzie and Bickle [5], that is, constant entropy
of melting and variation of degree of melting as a cubic function of temperature normalised to
the solidus and liquidus temperatures at the appropriate pressure (McKenzie and Bickle [5],
equation 21). The important differences are that the otherwise adiabatic geotherm is modified
by conductive cooling from the top and the solidus, and the degree of melting, are modified to
allow for the water content of the melt. Given the melt-fraction - depth profile and a mantle
source composition, REE concentrations are calculated using the method, mantle phase
equilibria and partition coefficients of McKenzie & O’Nions [36] as modified by White et al.
[7]. The melt-fraction calculation uses an expression for variation in degree of melting with
pressure and temperature based on equilibrium (batch) melting experiments whereas
calculation of the REE concentrations assumes pooled fractional melting; this is because good
data are not yet available for the variation of degree of fractional melting with pressure and
temperature. Iwamori et al. [41] use a simple approximation to fractional melting and
calculate that mantle potential temperatures need to be 30 to 50°C higher for fractional
melting to produce the same melt volumes as equilibrium melting.
The dry solidus and liquidus temperatures are those of McKenzie & Bickle [5]. The wet
solidus temperature is calculated from the dry solidus temperature using the approximation of
Davies & Bickle [42]
10
Tw  Ts  al X w
1
where Tw is the wet solidus temperature, Ts the dry solidus temperature and al is a
constant (642°C). Xw is the mole fraction of water in the melt as defined by Davies and Bickle
[42]. The maximum weight fraction of water in the melt is taken to be 0.25 at P  30kb and
below 30 kb this is assumed to decrease linearly to zero at P = 0 kb [42]. The weight fraction
of water in the melt is calculated using the batch melting equation,
W
M wm
X  D  DX
2
where D, the bulk distribution coefficient between water and the melt, is assumed to be
0.01 [43], X is the melt fraction by weight, Mwm is the weight fraction of water in the mantle
and W, the weight fraction of water in the melt. The melt fraction with water present is
assumed to vary with temperature as for dry-melting except with the solidus and liquidus
temperatures decreased by the temperature reduction for the wet solidus given by equation 1.
The geotherm is calculated first for upwelling of an 'equivalent solid' without melting, but
allowing for conductive cooling, followed by calculation of the temperature reduction
consequent on adiabatic melting. The 'equivalent solid' geotherm is given by the solution to
the steady-state one-dimensional advection – diffusion equation incorporating adiabatic
decompression and ignoring internal heat production.
 2T u T ugT


0
C p
z 2  z
3
where T is temperature, u is the mantle upwelling rate, z is depth,  is the thermal
expansion coefficient, g, the acceleration due to gravity, Cp is the specific heat capacity of the
solid mantle and  is the thermal diffusivity. A good approximation to the solution of equation
3 is,
T  Tp e





gz 
C p 
 Tp e
  az 
  


The first term is the adiabatic gradient and the second gives the modification resulting
from conductive cooling and is important only when the Peclet number (avz/) is small. Tp is
4
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the mantle potential temperature (in K), a is the tangent of the wedge angle  (the angle to the
horizontal made by the lithosphere at the spreading centre), which is the factor relating the
half spreading rate, v, to the mantle upwelling rate, u (Fig. 4a).
The melt temperature and fraction are calculated as a function of depth, using the method
of McKenzie & Bickle [5], by first assuming no volatile phases are present, and taking the
appropriate mantle potential temperature at a given depth from equation 4. The dry melt
temperature and melt fraction are then used to calculate the wet melt temperature and melt
fraction assuming conservation of energy and ignoring volume changes as in Davies and
Bickle [42]. Calculation of conductive cooling from the 'equivalent solid' geotherm results in
a small overestimate in temperature and melt fraction.
The model aims only to provide first order insight into the effect of adding water to the
mantle source and of varying the mantle upwelling rate (i.e. conductive cooling) in a simple
passive upwelling regime. The major simplifying assumptions are: 1) All melts are extracted
to form crust, 2) small (~ tens of °C) variations in mantle potential temperature do not affect
the position of the spinel – garnet transition, 3) the upwelling velocity of the mantle is
constant within the melting regime, 4) the mantle solidus is not affected by melting, 5) the
upwelling rate of the mantle is directly proportional to the half spreading rate, 6) the latent
heat of melting and al are constant, 7) the effect of CO2 may be ignored and 8) melting across
the wet and dry solidi to be continuous in space and time.
Results
The output from models run at varying mantle potential temperatures, mantle water
contents and upwelling velocities are displayed in Fig.5 as contours of melt thickness, average
melt water content and misfit, defined as the root-mean squared difference between observed
and calculated REE concentrations. REE data were averaged for 66°E, site 3, the Mid Atlantic
Ridge at 23°N [31] and the Juan de Fuca Ridge at 45°N [32]. The two faster spreading sites
on the Mid Atlantic Ridge and Juan de Fuca Ridge are included as controls (full spreading
rates of 26 and 57 mm/yr). The best fits to the REE patterns for various models are shown in
Figs. 6 and 7. Most of the REE scatter results from systematic changes in concentration of all
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the REEs between samples. Consequently, the slopes of the REE patterns are more tightly
constrained than indicated by the error bars calculated from the sample sets and misfits to
slopes within these error bounds are significant.
Increasing either the mantle potential temperature or the mantle water content increases
the depth at which melting starts and increasing the mantle upwelling rate decreases the depth
at which it stops. The calculated melt thickness is predominantly controlled by the mantle
potential temperature. The effects of spreading rate on melt thickness are significant at
spreading rates below 20 mm/yr, but melt water content makes relatively little difference. The
melt water content is a function of the melt thickness and is directly proportional to mantle
water content. REE patterns for dry-melting regimes that differ only in mantle potential
temperature are more or less parallel, with increased temperatures leading to dilution of the
elements. Changing the mantle water content alters the slope of the pattern by allowing
melting in the garnet stability field below the dry solidus. Plausible uncertainties in source
compositions calculated from Nd affect only the light rare earth elements. Varying the mantle
upwelling rate alone, for example from 7.5 mm/yr to 13 mm/yr, has roughly the same effect
on melt thickness as a 30°C increase in mantle potential temperature on the melt thickness.
When the mantle upwelling rate is set to 7.5 mm/yr, conductive cooling is found to shut off
melting at around 6 kb, that is ~ 18 km depth (Fig. 4b).
Melt thicknesses in normal or fast-spreading rate oceanic crust have been calculated from
REE analyses by McKenzie and O’Nions [36] and White et al. [7]. We discuss the two normal
spreading rate localities where chemical data including REE, water contents and seismic
estimates of crustal thickness are available; this enables us to investigate the significance of
small degrees of wet melting on REE patterns and compare the results with those from the
slow-spreading segments studied. The fits between model and observed REE concentrations
for the Mid Atlantic Ridge 23°N sample suite and the Juan de Fuca Ridge 45°N sample suite
are illustrated in Fig. 6. REEs from the Mid-Atlantic Ridge suite are best modelled with a
mantle potential temperature of 1295°C and a mantle water content of ~150 ppm for an
upwelling rate of 13 mm/yr which terminates melting at 12 km depth (r2=0.17). The predicted
melt water content of 0.15 wt% is similar to the observed value of 0.21 wt% [44] and the
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predicted melt thickness of 6.1 km is within the range of the observed thickness between 6
and 7 km [45]. Modelling the Mid-Atlantic Ridge with a faster upwelling rate (i.e. removing
conductive cooling) changes these parameters only slightly, mantle potential temperature
being reduced to 1270°C, the predicted melt thickness remains the same and mantle water
content increases to 200 ppm, which increases the predicted melt water content to 0.22 wt%
(r2=0.17). REE from the Juan de Fuca ridge segment are best modelled with a mantle
potential temperature of 1265°C and a mantle water content of 250 ppm (r2=0.18); this model
predicts a crustal thickness of ~ 6 km identical to the thickness inferred by Hooft & Detrick
[46] from an analysis of gravity data and melt water contents of ~0.3 wt%, higher than the
0.15 wt% observed [27]. Unfortunately, there is no direct seismic constraint in this area. The
discrepancy in water contents may be explained either by systematic problems with analysis
of water in the samples or errors in the source composition or fractionation processes which
change the enrichment of the LREE. Neither the water contents nor the Nd isotope values are
measured on the same samples as the REE, making this the most likely source of errors.
The modelled distribution of melt with depth for the MAR and Juan de Fuca suites shows
that, although most melting takes place above the dry solidus in the spinel peridotite stability
field, small degrees of melting occur deeper. This result is similar to determinations of melt
distribution by direct inversion of REE profiles [e.g. 7,36]. The significance of the small wet
melt fractions for incompatible trace element compositions is illustrated by the poor model
fits to the REE concentrations assuming a dry mantle in which the calculated REE patterns are
shallower than the measured patterns (Fig. 6).
The model and observed REE concentrations for the site 3 samples are matched most
closely (r2=0.25) with a mantle potential temperature of 1280 ± 5 °C and mantle water
contents between 150 and 200 ppm, given an upwelling rate of 7.5 mm/yr (Figs. 5 & 7). The
range of calculated melt water content from these models, of between 0.18 and 0.28 wt%,
encompasses the measured value of 0.19 wt%. The calculated melt thickness (3.9 ± 0.2 km) is
within error of the estimated crustal thickness (4.0 ± 1.0 [18]). The best match between model
and observed REE concentrations for site 66°E samples (r2=0.11) occurs at similar mantle
potential temperature (1280 ± 5 °C) but slightly higher mantle water contents (300 to 350
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ppm), given an upwelling rate of 7.5 mm/yr (Figs. 5 & 7). At these conditions, the range of
calculated melt water contents (0.4 to 0.5 wt%) is close to the measured average of 0.35 wt%
and the calculated melt thicknesses (4.7 to 5.2 km) are similar to the seismic estimate of 4.3 ±
1.0 km [17]. The higher mantle water content required by the 66°E lavas is consistent with
their higher water contents and steeper REE patterns, attributed to increased melting in the
garnet stability field.
If conductive cooling is ignored in the modelling of both the site 3 and 66°E data, the low
degree of melting, indicated by thin oceanic crust and relative REE enrichment, is achieved
by low mantle potential temperatures (~1230°C) and the HREE depletion is achieved by
increasing the estimated mantle water contents. However there is no combination of these
parameters which provides good fits to the slope of the REE data and the calculated misfits to
the REE data for the best fitting models are significantly increased (Fig. 7). Models run
assuming dry-melting give very poor fits to all the REE data (Figs. 6 & 7).
Discussion
The melt models for both normal and slow-spreading ridges show that incorporation of
the simple model for wet melting is capable of explaining the apparent garnet control on REE
concentrations without the necessity for dry-melting in the garnet peridotite field with its
implication of high mantle potential temperatures [e.g. 47]. The modelling of REE, melt
thickness and to a lesser extent melt water content show that conductive cooling as a
consequence of slow-spreading with near normal mantle potential temperatures (1280°C for a
batch melting model) provides a better explanation for the thin crust on the SWIR than
reduced mantle potential temperatures. This is in contradiction to the conclusion from the
relationship between Na8 and Fe8 in the lavas (Fig. 1) where the low Fe8 values imply shallow
solidus intersection depths [6] and low mantle temperatures (~1230 to 1250°C, using the
McKenzie & Bickle [5] mantle solidus). It is possible that the compatible element
concentrations are modified by mantle-melt reactions in the shallow mantle.
The first order conclusions are robust, despite the numerous simplifying assumptions in
the modelling. Perhaps the most critical of these is the assumption of batch melting to
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calculate the variation in degree of melting with depth. If fractional melting predominates and
the rate of melting decreases substantially as the degree of melting increases, then both the
modelling of melting in this paper and the interpretation of melt chemistry from Na8 to Fe8
relationships may need to be modified (note that the fractional melting models of Langmuir et
al. [6], are based on a simple variation of melting rate with degree of melt).
Another potential complication is that some or all of the garnet signature in the REEs may
be derived from melting garnet pyroxenite in the mantle [48]. If garnet pyroxenite is present
in the mantle, then some of the garnet signature in the HREEs may be generated without
requiring water in the mantle; this would lead to a reduction in the required mantle water
content and therefore in the melt water contents predicted by the model. It is also possible that
clinopyroxene in the mantle may contribute to the so-called garnet signature in the REEs.
New partition coefficients for the rare earth elements in near-solidus clinopyroxene [49]
suggest that the heavy rare earth elements may be compatible in clinopyroxene during
melting; this could reduce the requirement for garnet in the mantle source which, in turn,
would reduce the mantle water content and predicted melt water contents. A reduction in
predicted melt water content would improve the fits between the REE modelling, melt
thickness and melt water contents for the Juan de Fuca Ridge and perhaps a little for the 66°E
and site 3 data.
There are number of potential uncertainties in the estimation and modelling of melt
thicknesses. Melt retention in the mantle may lead to crustal thickness being an underestimate
of the melt thickness [e.g. 7]. Serpentinized mantle included in the seismic crustal thickness
measurement will lead to overestimation of melt thickness [e.g. 50-51]. The fractionation
correction applied to the data is also critical and uncertainties in the phase relationships of
fractionating assemblages, olivine partition coefficients, mantle oxidation state and Fe/Mg
ratios probably propagate through to at least ± 10% variation in melt thickness. The close
correspondence found between the melt thickness and crustal thickness suggests that none of
these processes is very important in the areas studied.
16
Conclusions
The measured crustal thicknesses, the rare earth element concentrations and water
contents of the SWIR lavas are fit better by the melting model if conductive cooling is used in
the modelling (Figs. 5, 6 & 7). When conductive cooling is included in the modelling, the
inferred mantle temperatures are similar to those calculated for average thickness normal
spreading rate oceanic crust such as that produced at the Mid-Atlantic Ridge at 23°N or Juan
de Fuca Ridge at 45°N. Fractionation corrected water contents of between 0.20 to 0.35wt% in
the SWIR lavas are consistent with between 150 and 300 ppm of water in the mantle and
melting of garnet-peridotite below the dry solidus.
Acknowledgements:
Thanks to Chris Richardson helped with equations and Peter Jacobsen who helped with
the programming. Useful reviews by Julian Pearce and Yaoling Niu improved the manuscript.
Cambridge research into the structure of the ocean crust and mantle melting work is supported
by NERC (GR3/8838 and a PhD studentship for CJR). TAM was supported by a Royal
Society University Research Fellowship. Earth Sciences contribution number 6335.
17
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Fig. 1a. Na8 versus crustal thickness. Shaded area is after Langmuir et al.[6] and solid squares are
points with seismically determined crustal thickness. Open triangles are data from the Juan de
Fuca Ridge (JdF) and Mid Atlantic Ridge (MAR) discussed in this paper (see below). Labelled
points with error bars are data from this study. Note that the crustal thickness measurement
used with the 57°E on-axis chemical data was measured on 11Ma crust. Fig. 1b. Na8 versus Fe8
(wt%) after Langmuir et al. [6] showing fields for data from the Mid-Atlantic Ridge near 26°N,
Mid Atlantic Ridge 23°N (MAR), Juan de Fuca Ridge 45°N (JdF), the Reykjanes Peninsula
and data from this study. The solid line approximates the global trend where lavas with higher
Fe8 and lower Na8 are the product of higher degrees of melting at higher mean pressures due to
higher mantle temperatures.
Fig. 2 Summary map showing the Southwest Indian Ridge (solid line labelled SWIR) between ~
55°E and ~ 70°E with the study areas highlighted. Also marked are the main fracture zones, the
Atlantis Platform and the extent of SWIR crust (dotted lines). CIR is the Central Indian Ridge
and SEIR , the Southeast Indian Ridge.
Fig. 3 Element-MgO plots for glasses analysed in this study. 2 s.d. analytical error bars shown on
each plot. Crosses are site 3 data and open squares are data from 66°E. The 57°E data (filled
diamonds) are presented to show the liquid line of descent used to correct the site 3 data.
Fig. 4 Schematic diagrams to show the effect of conductive cooling and of varying the mantle
potential temperature on the melting regime. Fig. 4a shows mantle flow through a melt region
where significant melting starts as mantle upwells through dry solidus and melting continues to
base of oceanic crust unless conductive cooling terminates melting below oceanic crust. Small
amounts of wet-melting will occur below ‘dry melt region’. The base of the wet melt region is
effectively defined as the level at which melt fractions are too small to separate from the solid
mantle. u, the upwelling velocity is a function of the wedge angle,  and the half spreading
rate, v. The wedge angle describes the geometry and therefore velocity of the upwelling column
of mantle. Fig 4b shows location of the dry solidus, solid adiabat, upwelling solid mantle with
conductive cooling appropriate to spreading at 15 mm/yr (full rate) and the modification of
these geotherms by melting. Geotherms terminate where melting stops (melt fraction starts
decreasing). Note how conductive cooling causes slight decrease in degree of melting and
terminates melting well below base of oceanic crust.
Fig. 5 Contoured plots showing the magnitude of missfits (root-mean squared difference between
observed and calculated REE concentrations) as a function of mantle potential temperature and
mantle water content. Shallow contours show calculated melt water content and steep contours
are calculated melt thickness. Solid circles show the measured mean melt water content and
oceanic crustal thickness at the relevant sample site (see text).
Fig. 6 REE contents of lavas from the Mid-Atlantic Ridge at 23°N and Juan de Fuca ridge at 45°N,
normalized to an appropriate mantle composition as determined by Nd values. Points are
means and error bars are 2 s.e. about those means for each locality. Curves are best fits for
models of adiabatic melting assuming either 1) melting is only above the dry solidus with no
21
conductive cooling, 2) wet-melting but no conductive cooling or 3) wet-melting and conductive
cooling, at the appropriate spreading rate. See discussion in text.
Fig. 7 REE contents, normalized to an appropriate mantle composition as determined by Nd values,
for site 3 and 66°E on the Southwest Indian Ridge. See caption to Fig. 6 for details and
discussion in text. See discussion in text.
Table 1 Uncorrected data used for the modelling. * Fe2O3 calculated from FeOT using the FeO/Fe2O3
ratio measured in whole rock samples using wet-chemistry. The mean water contents and Nd
values were used in the modelling. See text for analytical details.