JOURNAL OF PETROLOGY
VOLUME 39
NUMBER 2
PAGES 277–295
1998
Mid-Ocean Ridge Melting: Constraints from
Lithospheric Xenoliths at Oahu, Hawaii
HUAI-JEN YANG1∗, GAUTAM SEN1 AND NOBUMICHI SHIMIZU2
1
DEPARTMENT OF GEOLOGY, FLORIDA INTERNATIONAL UNIVERSITY, MIAMI, FL 33199, USA
2
WOODS HOLE OCEANOGRAPHIC INSTITUTION, WOODS HOLE, MA 02543, USA
RECEIVED APRIL 29, 1997; REVISED TYPESCRIPT ACCEPTED SEPTEMBER 9, 1997
by melting a lherzolitic source. Instead, they reflect mixing between
two components: (1) the pooled melt derived from lherzolitic source
in a passive melting regime with an Fmax of 30% (~80% of the
mixture) and (2) melt derived from garnet pyroxenite by 40%
fractional melting (~20% of the mixture). This model produces
7·1 km of crust.
One plagioclase–spinel lherzolite and four spinel lherzolite xenoliths
from Oahu, Hawaii, contain clinopyroxene grains that show homogeneous rare earth element (REE) abundances and smooth REE
patterns with systematic depletion of light REE (LREE). These
five xenoliths are mid-ocean ridge basalt (MORB) magma-depleted
residues with compositions that were not modified by later metasomatism. Trace element systematics of these xenoliths were used to
investigate the melt production rate (dF/dP) within a 90-my-old
residual mantle column (RMC). Such rates were calculated as
ratios of the difference in extents of depletion (dF) to the difference
in equilibrium pressures (dP) between two xenolith samples. The
extents of melting were modeled from REE, Sr and Zr abundances
in clinopyroxene; and equilibrium pressures were inferred from the
two-pyroxene geobarometer of Mercier et al. (1984, Contributions
to Mineralogy and Petrology, 85, 391–403). Equilibrium
pressures range from 21 to 7 kbar and extents of melting vary from
2 to 8%. Together, these data constrain the maximum extent of
melting in the garnet lherzolite stability field to be <2% and a
dF/dP of 0·43%/kbar within the stability field of spinel lherzolite.
Uncertainties in the estimates of equilibrium pressure and extent
of depletion lead to a slightly broader range of dF/dP values
(0·26–0·78%/kbar). These values are significantly lower than
that of ~1·2%/kbar suggested by most previous studies. With the
best estimated mean dF/dP of 0·43%/kbar, only 3·9 km of crust
could have been generated by melting lherzolite in the pressure range
of 26–7 kbar. The thickness and composition of the crust that
overlies the 90 Ma Oahu RMC require a higher extent of melting.
Based on the REE abundances, most samples from the 90 Ma
crust and East Pacific Rise can be explained as pooled melts derived
from the lherzolitic source in a passive melting regime with a
maximum extent of melting (Fmax) of 30%. This model produces
a dF/dP of 4·4%/kbar in the pressure range of 7–2 kbar and
generates an additional 1·8 km of crust. In detail, the high [Sm/
Yb]DM ratios in most East Pacific Rise samples cannot be explained
Mid-ocean ridge basalts (MORBs) are generated by
adiabatic decompression melting of asthenospheric
mantle during upwelling as a consequence of passive
spreading of the overlying lithosphere. Primary vs nonprimary nature of erupted MORB was the main focus
of petrological studies conducted between the early 1960s
and mid-1980s. Some researchers (O’Hara, 1968; Elthon
& Scarfe, 1980, 1984; Jaques & Green, 1980; Stolper,
1980; Elthon, 1986) postulated that all MORBs are
derived by olivine fractionation from picritic parental
magmas generated at pressures of 15–30 kbar. Other
researchers (Green & Ringwood, 1967; Kushiro, 1973;
Presnall et al., 1979; Fujii & Bougault, 1983; Takahashi
& Kushiro, 1983; Presnall & Hoover, 1984, 1986) argued
that the compositions of the least fractionated MORBs are
primary because they resemble experimentally produced
melts at pressures of 8–10 kbar. All of these models
assumed that primary MORB magmas are generated by
batch melting. However, recent studies have presented
evidence that fractional melting must be important in
∗Corresponding author.
 Oxford University Press 1998
KEY WORDS:
Hawaii; MORB; xenoliths
INTRODUCTION
JOURNAL OF PETROLOGY
VOLUME 39
MORB genesis. Klein & Langmuir (1987) used a global
database of MORB compositions to show that fractionation-corrected MORB compositions are strongly
correlated with axial depths of mid-ocean ridges. They
suggested that MORB crust is produced from melts
pooled from small increments of melt generated over a
pressure range from ~20–25 kbar to the base of the
crust. Moreover, Johnson et al. (1990) showed that the
abundances of rare earth elements (REE) and high field
strength elements (HFSE) in clinopyroxene of abyssal
peridotites require near-fractional melting. Sobolev &
Shimizu (1993) described melt inclusions with extremely
depleted light REE (LREE) abundances that can only
be explained by near-fractional melting.
The following general model of ridge melting is accepted by most researchers. Melting occurs in response to
adiabatic decompression of the ascending asthenosphere.
Because the asthenospheric adiabat has a lower temperature gradient (~1–2°C/kbar) than that of the dry
mantle solidus (~12°C/kbar), an ascending mantle parcel
starts to melt when it crosses the solidus (Fig. 1a). Because
the vertical cross-section of a melting regime of a passive
spreading ridge is probably triangular (Fig. 1b), a mantle
parcel that rises directly beneath a spreading center
undergoes the largest extent of melting and forms the
top portion of the lithospheric mantle (Fig. 1b). Asthenosphere that rises away from a spreading ridge
undergoes less melting and accretes to the lower parts of
the lithosphere (Fig. 1b). Consequently, the lithosphere
that forms from this melting regime is a vertically stratified
column of melt-depleted residues in which the extent of
depletion increases from bottom to top (Fig. 1b), and
was called the ‘Residual Mantle Column (RMC)’ by
Plank & Langmuir (1992).
An important parameter in the dynamics of melt
production and the compositional variations of MORB
is the change of melt fraction with pressure (dF/dP) during
adiabatic decompression. We refer to this parameter as
the ‘melt production rate’. The melt production rate
is typically calculated on a thermodynamic basis with
constraints from experimental data. The values calculated
in such a way range from ~1 to 2%/kbar (McKenzie &
Bickle, 1988; Langmuir et al., 1992; Iwamori et al., 1995).
A melt production rate could also be obtained if one
could estimate the variation in the extent of melting in
an entire RMC. Because abyssal peridotites represent
only the uppermost part of the RMC and direct sampling
of a whole RMC by drilling the lithosphere is impossible,
the latter approach has not been attempted. However,
lithospheric xenoliths that represent much of a 90-myold RMC occur in the Honolulu volcanics on Oahu,
Hawaii (Sen, 1983, 1988). The lithosphere beneath Oahu
is inferred to be ~90 my old because the basalts recovered
from the Ocean Drilling Program (ODP) Site 843, 250
km to the west of Oahu, range from ~94 to 110 Ma
NUMBER 2
FEBRUARY 1998
(King et al., 1993). We determined the abundances of
REE, Ti, V, Cr, Sr, Zr, and Y of clinopyroxenes in a
suite of spinel and spinel–plagioclase lherzolite xenoliths
with the ion probe at Woods Hole Oceanographic Institution (Table 1). Detailed rim-to-rim (through core)
traverses of analyses were carried out for two clinopyroxene grains from each sample. Based on these data
and the mineral compositions reported by Sen (1988)
and Sen et al. (1993), we selected five of these xenoliths
to model the extent of melting as a function of depth
and derived the melt production rate for this 90-my-old
RMC. Using these melt production rates and crustal
compositions, we then calculated the crustal thickness.
OAHU XENOLITHS
Xenoliths collected from the Honolulu volcanic vents
on Oahu, Hawaii, include lherzolites, pyroxenites and
dunites (White, 1966; Jackson & Wright, 1970; Sen,
1983, 1988; Sen & Presnall, 1986). 187Os/186Os ratios
provide important constraints on the origin of these rocks
because they remain virtually unchanged during fluid
metasomatism, as a result of the low Os concentration
in the fluid phases. Lherzolites from the Pali vent display
187
Os/186Os ratios in the range of abyssal peridotites
(Lassiter & Hauri, 1996), indicating that these samples
represent parts of an old RMC. In contrast to Os isotopes,
the abundances of incompatible elements such as LREE,
Zr, Sr and Ti in xenoliths reflect reequilibration of the
xenoliths with their host magmas or with earlier melts
that percolated through the lithospheric mantle (Navon
& Stolper, 1987). Many lherzolite xenoliths from the
Pali and Kaau vents contain clinopyroxenes that show
concave-upward REE patterns (Sen et al., 1993): for
example, a clinopyroxene in spinel lherzolite sample
77PAII-10 shows concave-upward REE patterns along
with systematic increases in La and Ce from grain cores
to rims (Fig. 2). These patterns result from metasomatic
enrichment of LREE by melts (Navon & Stolper, 1987;
Sen et al., 1993). Sr, Nd and Pb isotopic data for Salt
Lake xenoliths plot outside the field of MORB but overlap
the field of Honolulu volcanics (Okano & Tatsumoto,
1996), suggesting such enrichment may have resulted
from interactions with Hawaiian alkalic melts.
In this study, we determined the abundances of REE
and other trace elements in clinopyroxene from Pali,
Kaau and Salt Lake xenoliths [locations given by Sen
(1988)]. Our goal was to find xenoliths that contain
clinopyroxene grains that exhibit systematic depletion
of LREE, which is characteristic of MORB-depleted
lithospheric mantle that has not been affected by metasomatism, to estimate a melt production rate for the 90my-old RMC.
278
YANG et al.
MID-OCEAN RIDGE MELTING
Fig. 1. (a) A schematic diagram illustrating adiabatic decompression melting. The asthenospheric adiabat has a lower temperature gradient
(~1–2°C/kbar) than that of the dry mantle solidus (~12°C/kbar). Therefore, the ascending mantle parcel starts melting when it crosses the
solidus. (b) A schematic diagram of a triangular melting regime beneath a mid-ocean ridge and the resulting residual mantle column (RMC).
The dashed lines are flow lines of the ascending mantle. The mantle parcel that rises in the center of melting regime undergoes the greatest
extent of melting and eventually accretes to form the upper part of the RMC. The mantle parcel rising at the edge of the triangular melting
regime undergoes smaller extents of melting and accretes to the lower part of the lithosphere. Therefore, the extent of depletion in the RMC
increases from the bottom to the top.
ANALYTICAL METHODS
Clinopyroxene grains in polished xenolith thin sections
were analyzed in situ for abundances of REE, Ti, V, Cr,
Sr, Y, and Zr with the Cameca IMS-3F ion microprobe
at Woods Hole Oceanographic Institution. The operating
conditions were the same as those reported by Johnson
et al. (1990), except that La was also measured using
a basaltic glass as a standard. In each sample, two
clinopyroxene grains were analyzed. A rim–core–rim
traverse of 6–17 spot analyses was carried out on each
grain. The distance between two adjacent spots was
~0·1–0·2 mm. The uncertainties are ±7–15% for LREE,
±5–7% for middle and heavy REE (HREE), and ±10%
for other trace elements.
279
Olivine grains were analyzed for Ca contents with the
ion probe using two standards for which Ca abundances
had been determined by isotope dilution. One to six
spots on each of two or more olivine grains from each
lherzolite xenolith were analyzed (Table 2). The standard
deviations for analyses in each olivine are listed in Table
2.
RESULTS
One plagioclase–spinel lherzolite (77PAII-9) and three
spinel lherzolite xenoliths (77PAII-1A-1, 77KAPS-8 and
77KAPS-26) contain clinopyroxene grains that are characterized by smooth REE patterns with systematic
JOURNAL OF PETROLOGY
VOLUME 39
NUMBER 2
FEBRUARY 1998
Table 1: Abundances (ppm) of REE, Ti, V, Cr, Sr, Y and Zr in clinopyroxenes from three Pali and two
Kaau xenoliths
C1 ch∗
La
Ce
Nd
Sm
Eu
Dy
Er
Yb
0·235
0·603
0·452
0·147
0·056
0·243
0·159
0·163
436
57
2660
7·8
0·378
4·55
2·73
2·28
1535
264
6524
0·3
17·3
4·7
0·195
0·221
2·43
1·96
1·39
1·37
1·52
1·28
1467
1109
262
236
6272
7338
0·5
0·5
15·6
11·1
4·0
1·6
0·289
0·303
1·77
2·17
1·27
1·54
1·46
1·61
1272
1136
244
233
6205
5939
0·6
0·7
13·3
11·0
3·1
2·4
0·305
0·267
2·64
1·90
1·89
1·32
1·65
1·27
1038
1402
225
259
5806
6384
0·4
0·3
10·7
14·8
2·6
3·4
0·258
0·310
3·19
3·00
2·03
1·73
1·59
1·37
1518
276
6769
0·6
19·0
4·6
1·19
0·565
2·52
1·84
1·32
2608
269
3925
24·6
15·3
21·2
1·29
0·636
2·31
1·59
1·56
1·33
0·709
2·61
1·56
1·37
1·57
0·638
2·79
1·46
1·68
2894
3067
2629
2838
2697
2712
2609
2740
270
272
252
275
253
266
251
259
4353
4814
4473
5944
4080
4942
3326
4337
24·0
26·8
22·0
25·8
25·2
24·9
25·4
25·4
14·7
15·8
14·6
17·0
14·9
14·8
15·8
15·2
20·8
21·0
17·0
18·4
18·4
19·2
22·0
20·2
2659
2268
253
240
3758
4395
24·3
18·6
15·2
12·5
21·2
16·6
2605
2713
260
256
4899
3948
22·4
24·7
13·3
15·3
17·6
18·4
77PAII-9 (a plagioclase–spinel lherzolite from Pali)
cpx-1-1
0·029
0·141
0·728
1·46
cpx-1-3
0·011
0·066
0·578
0·993
cpx-1-4
0·015
0·128
0·641
0·789
cpx-1-5
0·015
0·112
0·485
0·709
cpx-1-6
0·012
0·102
0·673
0·566
cpx-1-7
0·021
0·110
0·658
0·658
cpx-1-8
0·016
0·104
0·407
0·539
cpx-1-9
0·009
0·075
0·627
1·02
cpx-2
0·011
0·135
0·796
1·23
77KAPS-8 (a spinel lherzolite from Kaau)
cpx-1-0
0·115
0·851
1·56
cpx-1-1
0·185
1·15
2·43
cpx-1-2
cpx-1-3
cpx-1-4
0·186
1·17
2·68
cpx-1-5
cpx-1-6
cpx-1-8
cpx-1-9
0·209
1·27
2·51
cpx-1-10
cpx-1-11
cpx-1-12
cpx-1-13
cpx-1-14
0·058
0·467
0·890
cpx-1-15
cpx-1-16
0·211
1·15
2·12
cpx-1-17
cpx-1-18
cpx-1-19
cpx-1-20
0·193
1·02
2·09
cpx-2-1
cpx-2-2
0·235
1·43
2·59
cpx-2-3
cpx-2-4
0·170
1·20
2·62
cpx-2-5
0·167
1·40
2·46
cpx-2-6
0·166
1·17
2·42
cpx-2-7
cpx-2-8
0·144
1·14
1·97
cpx-2-9
0·155
0·931
2·11
cpx-2-10
0·166
1·01
2·28
cpx-2-11
cpx-2-12
cpx-2-13
cpx-2-14
0·191
0·930
1·94
77PA-39 (a spinel lherzolite from
cpx-1-1
0·227
0·429
cpx-1-2
0·220
0·411
cpx-1-3
0·235
0·332
Pali)
1·90
1·75
1·67
Ti
V
Cr
Sr
Y
1·56
Zr
3·94
0·472
0·318
2·05
1·02
1·17
2691
2666
257
259
3935
3714
25·4
24·2
15·0
15·0
19·5
18·8
1·47
0·587
2·56
1·20
1·52
1·35
0·655
2·22
1·52
1·46
1·59
0·640
2·81
1·67
1·72
1·40
1·25
1·25
0·642
0·630
0·545
2·59
2·80
2·36
1·58
1·89
1·39
1·48
1·79
1·54
1·18
1·39
1·15
0·563
0·582
0·494
2·34
2·62
2·56
1·53
1·61
1·30
1·64
1·61
1·43
1·27
0·630
2·49
1·44
1·18
2694
2741
2709
2970
2654
2633
2727
2860
2746
2320
2400
2487
2430
2466
2741
2512
2430
2193
2705
261
267
259
262
265
307
268
279
294
240
248
258
276
264
279
276
292
246
278
4450
3960
3934
4366
4091
7041
4385
4683
7412
3552
4000
3338
4899
3656
4404
4679
6644
3711
4523
23·3
24·9
25·6
24·2
24·3
35·9
23·2
18·2
18·0
14·6
16·9
20·5
17·2
16·8
18·7
17·3
17·5
15·8
19·6
15·4
16·0
15·4
15·2
14·4
14·5
15·6
15·7
15·4
13·0
13·2
14·4
13·8
13·2
15·9
14·9
12·4
11·9
15·9
19·1
21·7
23·7
20·6
18·4
18·4
22·5
19·1
18·4
14·8
16·2
17·0
15·7
18·1
18·5
17·6
15·4
15·1
20·1
1·46
1·28
1·20
0·581
0·524
0·446
2·45
2·36
2·49
1·59
1·83
1·59
1·76
1·58
1·62
2449
2412
2421
286
278
319
3786
4248
7237
25·7
23·3
26·9
17·0
16·3
16·4
10·2
9·8
10·0
280
YANG et al.
La
Ce
77PA-39 (a spinel lherzolite from
cpx-1-4
0·430
0·436
cpx-1-5
0·224
0·456
cpx-1-6
0·156
0·428
cpx-1-7
0·161
0·418
cpx-1-8
0·223
0·502
cpx-1-9
0·446
0·503
cpx-1-10
0·360
0·368
cpx-1-11
0·122
0·413
cpx-1-12
0·092
0·312
cpx-1-13
0·151
0·427
cpx-1-14
0·138
0·376
cpx-1-15
0·106
0·440
cpx-1-16
0·270
0·364
cpx-1-17
0·070
0·490
cpx-2-1
0·041
0·390
cpx-2-2
0·066
0·459
cpx-2-3
0·067
0·465
cpx-2-4
0·072
0·311
cpx-2-5
0·064
0·343
cpx-2-6
0·085
0·408
MID-OCEAN RIDGE MELTING
Nd
Sm
Eu
Dy
Er
Yb
Ti
V
Cr
Sr
Y
Zr
Pali)
1·92
1·79
1·58
1·62
1·98
1·83
1·90
1·61
1·40
1·76
1·59
1·62
1·65
1·79
1·50
1·91
1·44
1·55
1·67
1·35
1·20
1·21
1·13
1·06
0·98
1·16
1·13
1·23
1·05
1·32
1·12
0·99
1·32
1·26
1·13
1·08
0·95
0·90
1·21
1·31
0·517
0·537
0·501
0·596
0·648
0·538
0·561
0·513
0·405
0·580
0·450
0·642
0·580
0·575
0·572
0·551
0·466
0·584
0·678
0·467
2·67
2·84
2·72
2·63
2·46
2·54
2·68
2·55
2·02
3·26
2·65
2·70
2·86
2·83
2·61
3·00
2·65
2·78
2·67
2·45
1·69
1·87
1·70
1·71
1·47
1·60
1·42
1·52
1·24
1·83
1·69
1·81
1·85
1·77
1·77
1·65
1·72
1·73
1·70
1·63
1·52
1·63
1·71
1·83
1·85
1·57
1·80
1·62
1·36
1·91
1·58
1·74
1·87
1·78
1·74
1·62
1·59
1·80
1·59
1·58
2330
2447
2412
2446
2391
2325
2401
2375
2446
2426
2486
2518
2761
2717
2700
2855
2866
2781
2745
2836
296
279
273
276
271
267
300
273
272
289
281
289
302
291
295
312
323
315
316
306
5924
4330
4322
4391
4201
4023
6334
4106
4102
5199
4237
4583
3968
4353
3995
4024
4497
3868
3567
3781
15·4
14·0
17·3
10·1
8·8
11·2
14·2
14·8
12·2
11·6
13·2
15·9
12·6
14·6
6·7
6·4
9·3
8·6
6·7
7·1
16·3
17·1
16·2
16·7
16·1
16·5
17·0
16·4
16·5
16·1
16·7
16·2
17·1
17·1
17·0
18·9
19·4
19·3
18·8
17·7
9·4
9·3
8·7
9·0
8·2
8·5
8·6
9·9
10·3
8·3
9·6
10·9
10·4
10·3
10·6
12·6
12·9
11·0
11·6
11·8
1975
2137
2118
2249
2175
1965
2010
2044
2111
1916
1874
2079
2088
2075
1954
252
264
260
271
260
253
258
261
333
248
248
260
260
260
260
3611
4460
3849
4499
3868
4288
3907
3778
3815
3316
3479
4670
3992
4005
3696
28·2
24·2
22·5
21·9
27·7
25·6
25·2
26·0
27·8
23·3
24·9
26·1
24·7
23·6
26·2
14·3
15·4
15·2
14·6
14·5
14·7
15·0
14·0
14·8
13·7
13·7
14·9
15·0
14·5
14·3
26·6
24·2
26·5
25·9
26·1
25·9
25·3
25·5
26·4
24·0
22·2
27·2
28·2
27·6
25·5
77KAPS-26† (a spinel lherzolite from Kaau)
cpx-1-1
cpx-1-2
cpx-1-3
cpx-1-4
cpx-1-6
cpx-1-7
cpx-2-1
cpx-2-2
cpx-2-3
cpx-2-4
cpx-2-5
cpx-2-6
0·567
1·78
2·75
1·47
cpx-2-7
0·654
1·93
2·80
1·50
cpx-2-8
0·716
2·15
3·07
1·59
cpx-2-9
0·675
2·15
3·04
1·47
0·688
0·696
0·724
0·692
2·81
2·91
3·05
2·89
1·74
1·78
1·83
1·84
2·02
2·11
2·06
2·03
77PAII-1A-1 (a spinel lherzolite from Pali)
cpx-1-1
0·026
0·219
1·10
cpx-1-2
0·014
0·193
0·99
cpx-1-3
0·027
0·223
0·97
cpx-1-4
0·017
0·195
0·91
cpx-1-5
0·021
0·221
1·22
cpx-1-6
0·029
0·206
1·14
cpx-2-1
0·077
0·449
1·69
cpx-2-2
0·049
0·335
1·48
cpx-2-3
0·037
0·344
1·49
cpx-2-4
0·039
0·322
1·51
cpx-2-5
0·037
0·308
1·36
cpx-2-6
0·023
0·253
1·22
cpx-2-7
0·036
0·314
1·48
0·440
0·382
0·387
0·422
0·447
0·502
0·584
0·536
0·526
0·494
0·507
0·468
0·549
1·74
1·49
1·56
1·41
1·71
1·74
2·43
2·18
2·25
2·03
1·99
1·65
2·14
0·90
0·87
0·84
0·79
0·97
0·93
1·51
1·15
1·20
1·20
1·17
0·97
1·22
1·39
1·20
1·22
1·20
1·46
1·43
1·93
1·72
1·71
1·59
1·65
1·35
1·65
0·767
0·644
0·702
0·621
0·767
0·734
1·13
0·931
0·921
0·889
0·845
0·720
0·983
∗Chondrite values from Anders & Grevesse (1989).
†Only four REE analyses from 77KAPS-26 are reported, as others have slightly high Sm, probably because of analytical
error.
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Fig. 2. This example illustrates the level of detail at which we analyzed the xenolith clinopyroxenes. Chondrite-normalized (Anders & Grevesse,
1989) REE concentrations of a clinopyroxene in spinel lherzolite sample 77PAII-10 show concave-upward REE patterns (except one analysis in
the core) with systematic increases in La and Ce from core to rim. The insert shows the chondrite-normalized La contents in a rim–core–rim
traverse.
Table 2: Ion-probe analyzed Ca contents (ppm) in olivines from Pali and Kaau xenoliths
77PAII-9
ol-1
No. of analysis
Mean (ppm)
SD (%)
77KAPS-26
ol-2
ol-3
ol-1
Sample average
ol-6
ol-7
3
3
3
1
2
2
2
2
110
119
240
257
236
273
247
252
3
5
11
7
—
0
5
9
3
112
ol-1
SD (%)
ol-5
3
77PA-39
ol-2
ol-3
ol-10
1
252
—
251
77KAPS-8
Mean (ppm)
ol-3
107
Sample average
No. of analysis
ol-2
ol-4
ol-1
77PAII-1A-1
ol-2
ol-1
ol-3
ol-4
3
3
2
2
6
3
5
3
4
241
215
234
220
269
285
236
248
243
5
4
5
10
15
5
6
5
2
228
277
depletions in LREE (Fig. 3a), similar to patterns seen
in abyssal peridotites ( Johnson et al., 1990). A large
clinopyroxene porphyroclast in plagioclase–spinel lherzolite sample 77PAII-9 shows negative Eu anomalies
ol-5
2
233
6
240
only at its rim, which suggests that the rim had reequilibrated with plagioclase (Fig. 3a, Table 1). The
similarities between REE patterns of clinopyroxene grains
from abyssal peridotites and from these four xenoliths,
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MID-OCEAN RIDGE MELTING
as well as the homogeneities of REE, Ti, Zr, Sr and Y
abundances within and between clinopyroxene grains in
each xenolith sample (Fig. 3a) indicate that the compositions of these four xenoliths were not modified by
Hawaiian magmatism (or any other melt–wall rock interaction). A fourth spinel lherzolite xenolith sample, 77PA39, contains a clinopyroxene grain, for which some
analyses show slight enrichment of La relative to Ce
(Table 1). The abundances of other incompatible elements in this grain are homogeneous and REE patterns
are smooth with systematic depletions from Sm to Ce
(Table 1). Another analyzed clinopyroxene grain in this
sample has smooth REE patterns with systematic depletions from Sm to La (Fig. 3a).
The LREE, Zr, Sr, and Y contents of clinopyroxene
grains differ significantly among these five samples (Figs
3a and 4). For example, the [La]n values [subscript ‘n’
indicates normalized to the chondritic values (Anders &
Grevesse, 1989)] of clinopyroxene grains from 77KAPS26 are 60 times higher than those from 77PAII-9, implying that the former sample is less depleted than the
latter (Fig. 3a).
The Ca abundances of olivines in these five xenoliths
range from 112 to 277 ppm (Table 2). The standard
deviations in all but one olivine grain from 77PA-39 are
Ζ11%, mostly Ζ6% (Table 2).
DISCUSSION
Extents of depletion
Partition coefficients, mineral proportions in the source
and those entering the melt, as well as source and residue
compositions are needed to calculate the extent to which
each xenolith sample has been depleted. The MORB
source is inferred to be a residue after removing 2% melt
generated by non-modal batch melting in the garnet
lherzolite stability field from the primitive mantle composition of McDonough & Sun (1995) [see Hirschmann
& Stolper (1996) for mineral proportions in the source
and melt]. The 2% depletion value was chosen so that
Sm/Nd and Lu/Hf ratios in the residue (the MORB
source) can generate the 143Nd/144Nd and 177Hf/176Hf
ratios observed in present-day MORB with a 2 Ga model
age. To model melting of spinel lherzolite, garnet in this
depleted source was converted to spinel by the reaction
of Mg3Al2Si3O12 (gt) + Mg2SiO4 (ol) = MgAl2O4 (sp) +
2Mg2Si2O6 (opx). The REE contents of clinopyroxene
grains in the resultant depleted source resemble those
estimated by Johnson et al. (1990) from abundance ratios
of MORB. The mineral proportions calculated for this
depleted source (olivine:orthopyroxene:clinopyroxene:
spinel = 51·6:29·2:17·0:2·2) are similar to those Kinzler
& Grove (1992a) inferred from their experimentally produced phase compositions.
Using the source compositions (Table 3) and mineral
proportions described above, along with the experimentally determined melting stoichiometry for spinel
lherzolites (Kinzler & Grove, 1992b) and partition coefficients compiled from literature (Table 3), we calculated
REE, Sr, and Zr abundances of clinopyroxene in residues
produced by partial melting. We considered (1) batch
melting, (2) fractional melting, and (3) incremental melting. The equations are those given by Johnson et al.
(1990):
C
D io
C icpx
o,cpx=
o
Ci
D i +F(1−Pi)
D
(1)
for non-modal batch melting, and
C D
C icpx
PF
= 1− i o
C o,cpx
Di
i
A B
1
−1
Pi
(2)
for non-modal fractional melting, in which Cio,cpx and Cicpx
are the concentrations of element i in clinopyroxene in
the initial source and residue, respectively; Dio is the bulk
partition coefficient of element i in the initial source; Pi
is the sum of the partition coefficients of phases in the
proportions that they enter the melt; and F is the extent
of depletion. In model (3), each increment of melt was
produced by 1% batch melting. Because melt presumably
would segregate from its source after 1% batch melting,
resulting residue from each increment becomes the source
for the next 1% batch melt.
Batch melting clearly fails to produce the REE patterns
exhibited by the xenolith clinopyroxenes. For example,
22% melting of the modeled source generates lower
HREE and MREE abundances but higher LREE abundances than the most depleted xenolith, 77PAII-9 (Fig. 3b).
In this calculation, clinopyroxene completely disappears
from the residue after 22% melting, and yet 77PAII-9 has
several clinopyroxene grains. Thus, both the petrography
and REE abundances of this xenolith are inconsistent
with it being a residue of batch melting.
The fractional melting model fits the data better than
does batch melting (Fig. 3c). Indeed, 77PAII-9, the most
depleted xenolith, can be modeled as a residue that lost
6% MORB melt. However, a model of 1% incremental
batch melting yields the best fit to the REE, Sr and Zr
abundances in clinopyroxene grains from all five analyzed
lherzolites (Figs 3d and 4). This result is similar to that
of Johnson et al. (1990), who showed that the REE
contents of clinopyroxene grains from abyssal peridotites
require a near-fractional melting process. Results of 1%
incremental melting model indicate that the most depleted plagioclase–spinel lherzolite (77PAII-9) underwent
8% melting and the least depleted spinel lherzolite
(77KAPS-26) lost only 2% of melt (Figs 3d and 4).
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Fig. 3. (a) Chondrite-normalized (Anders & Grevesse, 1989) REE concentrations of clinopyroxene in five xenoliths from Oahu, Hawaii. For
each sample, cross-grain traverse analyses were made for two clinopyroxene grains. Except for 77PA-39, all analyses in the other four samples
show smooth LREE-depleted patterns which are similar to those in abyssal peridotites ( Johnson et al., 1990). Analyses for 77PA-39 show similar
patterns, except for some that have slightly enriched La (>0·13 ppm) which are not shown for clarity. (b–d) Comparisons between analyses in
(a) and modeled residues from melting a depleted spinel lherzolite with 2 Ga Nd and Hf model ages. One representative analysis (approximately
the mean) is shown for each sample. (b) Batch melting produces too low HREE and MREE and too high LREE for 77PAII-9. (c) Six percent
fractional melting produces slightly higher MREE and lower LREE than 77PAII-9. (d) Incremental melting with 1% melting steps provides the
best fit to the data. The error bar indicates the uncertainty of [La]n for 8% melting (see text for discussion).
Therefore, the upper limit of melting that could have
occurred within the garnet lherzolite field at higher
pressures is ~2%.
F estimates depend on D cpx/melt
LREE,Sr,Zr and clinopyroxene
melt
fraction that enters melt (X cpx
). This is because cpx/
melt partition coefficients for LREE, Sr and Zr are
2–4 orders greater than those for the other phases in
spinel lherzolites (Table 3). Phase equilibrium exmelt
increases with pressure
periments have shown that X cpx
(Walter et al., 1995; Kinzler, 1997). In our calculations
of spinel lherzolite melting we used stoichiometric
coefficients that are appropriate at 15 kbar (Kinzler &
Grove, 1992b; Walter et al., 1995; Kinzler, 1997). If
we use the 10 kbar [inferred from Walter et al. (1995)
and Kinzler (1997)] and 20 kbar coefficients (Walter
et al., 1995), the [La]n value for 8% incremental melting
only changes from 0·032 (at 15 kbar) to 0·037 and
0·023, respectively (Fig. 3d). The uncertainty in an F
estimate is dominated by D cpx/melt
LREE,Sr,Zr. The combination
of 0·04 (approximately the lowest in our
of a D cpx/melt
La
compilation) and the melting stoichiometry at 20 kbar
results in the lowest F estimate of ~6% for 77PAII-9
and 1% for 77KAPS-26. The highest estimated F
values for 77PAII-9 and 77KAPS-26 are 9% and 2%,
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MID-OCEAN RIDGE MELTING
Fig. 4. Comparison between the Zr and Sr contents in clinopyroxene from Pali xenoliths and those calculated by the 1% incremental melting
model (continuous line). Each tick mark on the continuous line indicates 1% melting. Each square indicates the median value of analyses in one
sample. The bars show the ranges of analyses in one sample. The inflection to higher Zr at low Sr content results from spinel–plagioclase
transformation.
Table 3: Depleted mantle source composition (ppm) and partition coefficients for modeling calculations
Element
Source∗
ol/melt†
opx/melt†
cpx/melt†
sp/melt†
La
1·40
0·00004
0·0028
0·065
0·00002
0·009
0·0348
Ce
4·63
0·00006
0·0032
0·09
0·00003
0·035
0·0278
Nd
5·21
0·00031
0·0041
0·19
0·00004
0·19
0·0179
Sm
2·05
0·00065
0·0058
0·30
0·00009
0·67
0·0132
Eu
0·831
0·00075
0·0078
0·40
0·00014
1·15
0·3
Dy
3·83
0·0027
0·012
0·52
0·00021
2·6
0·0112
Er
2·36
0·011
0·022
0·55
0·00033
3·35
0·0116
Yb
2·20
0·027
0·039
0·58
0·00046
5·94
0·016
Lu
0·314
0·030
0·053
0·53
0·0008
5·92
1·20
Hf
gt/melt†
plag/melt†
0·0029
0·023
0·31
0·05
0·30
Zr
39·4
0·0010
0·021
0·16
0·06
0·50
0·0092
Sr
38·1
0·00015
0·075
0·12
0·00003
0·0065
2·5
∗Clinopyroxene composition in the depleted source.
†Partition coefficients are the median values from the following database: olivine (ol): Beattie (1993, 1994), Kennedy et
al. (1993); orthopyroxene (opx): Beattie (1993), Kennedy et al. (1993); clinopyroxene (cpx): Fujimaki et al. (1984), Irving &
Frey (1984), Hart & Dunn (1993), Hack et al. (1994), Hauri et al. (1994), Skulski et al. (1994); spinel (sp): Stosch (1982),
Horn et al. (1994); garnet (gt): Shimizu & Kushiro (1975), Irving & Frey (1978), Fujimaki et al. (1984), Beattie (1994), Hauri
et al. (1994); and plagioclase (plag): Fujimaki et al. (1984).
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respectively, when a D cpx/melt
of 0·09 (approximately the
La
highest in our compilation) and the melting stoichiometry at 10 kbar are used.
Xenolith equilibrium pressures
Sensitive geobarometers are lacking for spinel lherzolites
(Sen & Jones, 1989). The Al2O3 content of orthopyroxene,
which is a good geobarometer for garnet and plagioclase
lherzolites, is virtually insensitive to pressure in the spinel
lherzolite stability field (Gasparik, 1984; Sen, 1985). Ca
exchange between olivine and clinopyroxene has been
suggested as a geobarometer for spinel lherzolites (Finnerty & Boyd, 1978; Adams & Bishop, 1982, 1986;
Ko¨hler & Brey, 1990). In an effort to use the proposed
Ca exchange barometer (Ko¨hler & Brey, 1990), we used
the ion probe to analyze Ca in olivine. However, we
were unable to obtain reasonable equilibrium pressures
for Hawaiian xenoliths using this geobarometer along
with the geothermometer of Brey & Ko¨hler (1990). For
example, we obtained a pressure estimate of 35 kbar for
plagioclase–spinel lherzolite sample 77PAII-9 and 42
kbar for spinel lherzolite sample 77KAPS-26. These
estimates are clearly incorrect, because experimental
studies have shown that (a) plagioclase cannot be stable
in a lherzolite above ~10 kbar and (b) garnet is likely to
be stable at 42 kbar [see Sen (1988) for references]. The
experimental data (Adams & Bishop, 1986; Ko¨hler &
Brey, 1990) show that the Ca content of olivine is low
and less sensitive to pressure at temperatures Ζ1000°C
(Fig. 5). The ion-probe-determined Ca abundances of
olivine from our samples (Table 2) are lower than those
in the experiments at 8–30 kbar (Adams & Bishop, 1986;
Ko¨hler & Brey, 1990) (Fig. 5). Apparently, the barometer
of Ko¨hler & Brey (1990) cannot be extrapolated to low
temperatures (<1000°C) at which Ca contents in olivines
are low (<300 ppm; Fig. 5).
A geobarometer based on Ca exchange between
orthopyroxene and clinopyroxene was proposed for pressures ranging from 5 to 100 kbar by Mercier et al.
(1984). At P = 5–30 kbar and T = 900–1300°C this
geobarometer predicts 80% of the pressures of the experiments they used to calibrate the barometer within
±5 kbar (Fig. 6). Larger deviations occur at 25–30 kbar
(Fig. 6).
Several high-pressure pyroxene equilibration experiments have been carried out (e.g. Sen & Jones, 1989;
Brey et al., 1990) since the publication of the Mercier et
al. (1984) geobarometer. Sen & Jones (1989; see also Fig.
6) noted that this barometer predicts the pressures of
their spinel lherzolite subsolidus equilibrium experiments
within ±3 kbar. The experiments of Brey et al. (1990)
at 15–30 kbar can also be reproduced within ±5 kbar,
except for two experiments at 10 kbar (Fig. 6). Therefore,
NUMBER 2
FEBRUARY 1998
we consider the Mercier et al. (1984) geobarometer to be
appropriate for our purpose. This geobarometer yields
a pressure estimate of 6·8 kbar for plagioclase–spinel
lherzolite sample 77PAII-9, which is identical to the
estimate obtained from Gasparik’s [Al2O3]opx geobarometer (Gasparik, 1987; Sen, 1988). The Mercier et
al. (1984) barometer gives equilibration pressures between
12 and 21 kbar for the other three spinel lherzolites. A
spinel lherzolite sample from the Salt Lake Crater SL13, which displays clinopyroxene REE patterns similar
to those in garnet-bearing assemblages, yielded a pressure
estimate of 26 kbar. This is similar to that obtained from
the Al2O3 content in its orthopyroxene [25–27 kbar from
fig. 8 of Gasparik (1987)]. Spinel lherzolite sample 77PA39 yielded a pressure estimate of 58 kbar. As indicated
in an earlier section, a clinopyroxene grain from this
sample shows some La enrichment. Ca abundances in
an olivine grain show a larger standard deviation (15%)
than do values for other samples (Table 2). The unreasonably high pressure estimate probably reflects the
metasomatic enrichment that affected this xenolith.
In summary, several lines of evidence support our
contention that the pressure estimates we obtained are
meaningful. These are (1) a reasonable reproduction of
pressures of experimental runs using the Mercier et al.
(1984) barometer, (2) consistency between pressure estimates and phase equilibria of lherzolite mineral assemblages, and (3) a positive correlation between pressure
estimates and LREE contents in clinopyroxene [i.e. Ce
vs pressure (Fig. 7)], which is consistent with the expectation that the extent of depletion increases from the
bottom to the top of an RMC.
Melt production rate (dF/dP) in the 90 Ma
Oahu RMC
Melt production rates were calculated as ratios of the
difference in extents of depletion (dF ) to the difference in
equilibrium pressures (dP ) between two xenolith samples.
Based on the incremental melting model, the difference
in extent of melting between the most depleted (77PAII9) and the least depleted xenoliths (77KAPS-26) is 6%.
The difference in the equilibrium pressure between these
two samples is 14 kbar. These values yield a melt production rate (dF/dP ) of ~0·42%/kbar for pressures between 21 and 7 kbar. This estimated dF/dP is
independent of the source composition. The error associated with this estimate arises from the uncertainties
in the estimates of extent of depletion and equilibrium
pressure. As discussed in the preceding section, the largest
and smallest differences in extent of depletion between
the most and least depleted samples are 7% and 5%,
respectively. Because the pressure estimate for the plagioclase–spinel lherzolite (77PAII-9) was calculated using
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YANG et al.
MID-OCEAN RIDGE MELTING
Fig. 5. Experimental pressure vs Ca content in olivine. Filled symbols are for Ko¨hler & Brey (1990), and open symbols are for Adams & Bishop
(1986). Triangles, circles, squares and inverted triangles are for temperatures at 1300, 1200, 1100, and 1000°C, respectively. Ca content in
olivine becomes less sensitive to pressure at low temperature. The shaded area indicates the Ca contents of olivines from five Hawaiian xenoliths
(four spinel lherzolites and a plagioclase–spinel lherzolite) in this study.
Fig. 6. Experimental run pressures (Mori & Green, 1975; Lindsley & Dixon, 1976; Perkins & Newton, 1980; Sen & Jones, 1989; Brey et al.,
1990) versus those calculated with the geobarometer of Mercier et al. (1984). The continuous line indicates the prefect fit to experimental
pressure. The dashed lines indicate region wherein the data are fitted to ±3 kbar. The dotted lines indicate region wherein the data are fitted
to ±5 kbar.
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Fig. 7. Pressure calculated from the equation of Mercier et al. (1984) vs Ce content in clinopyroxene normalized to chondritic value of Anders
& Grevesse (1989). The vertical bar on each datum point indicates the range of [Ce]n in that sample. The insert shows one more spinel lherzolite
which contains clinopyroxene with REE patterns in equilibrium with garnet. The positive correlation between pressure and [Ce]n shows that
deeper samples are less depleted, which agrees with the concept of residual mantle column.
two different geobarometers and it is consistent with the
pressure range of the spinel to plagioclase transition (7–10
kbar) (Takahashi & Kushiro, 1983), the error on this
pressure estimate is probably insignificant. If we consider
a rather liberal error of ±5 kbar for 77KAPS-26 and
the uncertainty in the estimate of extent of depletion,
the melt production rate in spinel lherzolite stability field
of this 90-my-old RMC was in the range of 0·26–0·78%/
kbar (Fig. 8). We emphasize that our best estimate is
0·42%/kbar.
Comparison between melt production rate
in the Oahu RMC and previous studies
Before this study, melt production rate beneath midocean ridges was estimated using thermodynamic criteria
with constraints from experimentally determined solidus
and liquidus of mantle lherzolite (McKenzie & Bickle,
1988; Langmuir et al., 1992; Iwamori et al., 1995). Based
on enthalpy conservation during melting, Langmuir et
al. (1992) obtained a melt production rate of ~1·2%/
kbar (Fig. 8). They suggested that the rate decreases
with decompression during continual upwelling of the
asthenosphere in the melting regime. McKenzie & Bickle
(1988) concluded that decompression melting of a mantle
with a potential temperature of 1280°C can produce
the compositions of pooled MORB magmas. Because a
potential temperature of 1280°C intercepts the solidus
at ~15 kbar, a mean melt production rate of 2·5%/kbar
(Fig. 8) will generate a 7 km thick crust, if melting
continues to the bottom of the crust and obeys the
equation of Forsyth (1993):
Crustal Thickness=[dF/dP·(Zo2−Zf2)]/1·8
(3)
where Z o is the depth where melting begins, Z f is the
depth where melting ceases and, Z = 0 at the base of
crust. McKenzie & Bickle (1988) also claimed that the
melt production rate decreases at low pressures, which
means that dF/dP at high pressures should be greater
than the mean dF/dP of 2·5%/kbar. On the basis of
experimentally determined solidus and liquidus of the
lherzolite KLB-1 (Takahashi, 1986; Takahashi et al.,
1993) and conservation of mass and entropy during
melting, Iwamori et al. (1995) suggested that the melt
production rate does not vary systematically with depth.
Their model showed that a parcel of mantle with a
potential temperature of ~1350°C starts melting at P
~20 kbar and can generate 7 km of oceanic crust. This
implies a mean melting rate of 1·3%/kbar (Fig. 8). Using
a constant melt production rate of 0·9%/kbar, Kinzler
& Grove (1992a) showed that the variations in major
elements of MORB can be modeled by ~6–18% of
melting from the depleted source of Hart & Zindler
(1986). In this model, melting begins at 25–12 kbar and
ceases at 4 kbar. Shen & Forsyth (1995) used a melt
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YANG et al.
MID-OCEAN RIDGE MELTING
Fig. 8. Melt production rate (%/kbar) vs depth (km). (See text for the derivation of melt production rates for each model.) The thick continuous
lines are the best estimates from the present study. The thick dashed lines indicate the possible range of dF/dP when uncertainties of F and P
are considered. Our modeled rates are similar to the model of Asimow et al. (1997), but are very different from others. The curve of Asimow et
al. (1997) is only a schematic representation. The reader should refer to their paper for the exact results from MELTS.
production rate of 1·2%/kbar to model major and trace
element variations in MORB. They suggested that melting begins at 20–23 kbar (~60–70 km) and ceases at
different levels.
Asimow et al. (1997) modeled reversible adiabatic
upwelling by balancing entropy in a hypothetical twocomponent system. They argued that the melt production
rate increases at low pressures. Using the MELTS algorithm to model a nine-component system with an adiabat
that intersects the solidus at 22 kbar, Asimow et al. (1997)
obtained results similar to the two-component system,
namely, a low melt production rate of ~0·2%/kbar from
22 to 10 kbar followed by an increase from 0·2%/kbar
at 10 kbar to 3%/kbar at 5 kbar (Fig. 8).
In sharp contrast to all these studies, our data provide
direct constraints on dF/dP based on major (required
for barometry) and trace element (required for calculating
extent of depletion) systematics in a 90-my-old RMC.
Our estimated range of dF/dP (0·26–0·78%/kbar) is
substantially lower than the values of ~1%/kbar (or
higher) that have been used in most geochemical modeling studies (e.g. Kinzler & Grove, 1992b; Langmuir et
al., 1992; Shen & Forsyth, 1995; Kinzler, 1997). Our best
estimate resembles the low values obtained by Asimow et
al. (1997).
To investigate whether the melt production rate varies
much within the spinel stability field, we used spinel
lherzolite 77KAPS-8, which we estimate equilibrated at
19·7 kbar and lost 3% melt. By comparing this sample
with the least depleted sample (77KAPS-26), we estimate
that melt production rate at pressures of 20–21 kbar
is 0·71%/kbar. Comparing 77KAPS-8 with the most
depleted sample (77PAII-9) yields a melt production rate
of 0·39%/kbar. The decrease of melt production rate
towards the top of the RMC is consistent with inferences
from the thermodynamic models (Langmuir et al., 1992),
but we must add the caveat that the Mercier et al. (1984)
geobarometer cannot resolve a pressure difference of 1
kbar.
Crustal thickness: inferences from RMC
and crustal compositions
Crustal thickness can be inferred from the melt production rate and depth (pressure) interval of melting
using equation (3). If the melt production rate of 0·42%/
kbar prevails at pressures >21 kbar, then melting must
start at P ~26 kbar to produce the 2% depletion at 21
kbar estimated for sample 77KAPS-26. If melting did
occur over the pressure range of 26–7 kbar with a
melt production rate of 0·42%/kbar, we infer a crustal
thickness of 3·9 km. With the same logic, the models
with melt production rates of 0·78 and 0·26%/kbar
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Fig. 9. Comparison between REE abundances in calculated melt pooled from a passive melting regime (dashed lines without symbols) and
those in the EPR (Allan et al., 1989; Batiza & Niu, 1992; Niu et al., 1996) (shaded) and 90 Ma crust drilled from DSDP and ODP Sites. Four
calculated melt compositions derived from melting lherzolite with F max of 8%, 22%, 30% and 35% are shown. In our calculation, first 2%
melting occurs in the garnet stability field, followed by 6% melting in the spinel stability field, then melting continues in the plagioclase stability
field to reach F max of 22%, 30% and 35%. Plagioclase and clinopyroxene are exhausted at ~14% and ~23% melting, respectively. The inflection
of LREE in most 90 Ma crust may be due to alteration.
require melting to start at 19 and 34 kbar, respectively.
If melting ceases at 7 kbar, the former model would
generate 3·4 km of crust and the latter 4·3 km of crust.
These estimates of crustal thickness are much less than
the 6–7 km thickness estimated from the seismic data
from the vicinity of Hawaiian Islands (Ten Brink &
Brocher, 1987). To generate a 6–7 km crust, melting
must have continued at pressures <7 kbar.
Crustal compositions can add constraints on crustal
thickness. Unfortunately, only a few altered crustal
samples were drilled from the periphery of Oahu. Abundances of trace elements in four of seven samples cored
from Site 843 of the Ocean Drilling Program (ODP)
(King et al., 1993) show large positive Eu and Sr anomalies. Based on the positive correlation between Eu∗/
Eu and Zr/Nd ratios, King et al. (1993) proposed that
these plagioclase signatures did not result from plagioclase
accumulation; instead, they reflect a source distinct from
N-MORB. Three other samples reported by King et al.
(1993) and one sample from Site 164 of the Deep Sea
Drilling Project (DSDP) (Bass et al., 1973) have REE
abundances that fall within the values for samples from
the East Pacific Rise (EPR) (Fig. 9). Two samples from
ODP Site 169 show abundances of HREE lower than
those of the EPR (Bass et al., 1973). The samples with
REE abundances similar to those of the EPR are compared with the calculated pooled melt compositions (in
the next section) to estimate the extent of melting.
REE abundances in pooled melt were calculated by
integrating incremental melts in equilibrium with residues
from a passive melting regime using the method outlined
by Langmuir et al. (1992). Pooled melts from melting at
pressures of 26–7 kbar (maximum extent of melting
F max = 8%; first 2% melting in the garnet stability field,
then 3–8% melting in the spinel stability field) display
LREE abundances that are too high for these melts to
be parental magmas of the lavas that formed 90-Ma
crust (Fig. 9). More depleted melt increments derived
from plagioclase lherzolite are required. This result is
consistent with the inference from crustal thickness calculation with our estimates of dF/dP. If F max = 30%,
the REE abundances in pooled melts can be parental to
those of EPR lavas (Fig. 9). The parental magma for two
samples from DSDP Site 169 requires an F max >35%.
However, these two samples are unusual in their low
HREE abundances relative to EPR lavas. With an F max
of 30%, the melt production rate is 4·4%/kbar in the
plagioclase lherzolite field [(30%–8%)/(7 kbar–2 kbar)].
This estimate exceeds the highest value of ~3%/kbar
obtained by Asimow et al. (1997). With such a melt
production rate, melting at 7–2 kbar produces ~1·8 km
of crust. Therefore, melting of a lherzolitic source can
generate a total of ~5·7 (3·9 + 1·8) km of crust. This
estimate is close to the seismic data that suggest a 6–7
km thick crust (Ten Brink & Brocher, 1987). The model
requires that a lherzolitic source underwent ~30% melting (Fig. 9), which would leave harzburgitic residues, a
feature consistent with widespread occurrence of harzburgite at the top of the mantle sections of ophiolite
sequences (i.e. Elthon, 1979).
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MID-OCEAN RIDGE MELTING
Garnet signatures in crustal composition:
from garnet lherzolite or garnet pyroxenite?
Three compositional characteristics of MORB suggest
the presence of residual garnet in their source. These
are: (1) U–Th disequilibrium that shows evidence for
high U/Th ratios in the source; (2) high 176Hf/177Hf
ratios, which imply high Lu/Hf ratios in the source; and
(3) [Sm/Yb]DM ratios of 1·4–1·5. Sources that could
contain residual garnet are garnet lherzolite (Salters,
1996) and garnet pyroxenite (Alle`gre & Turcotte, 1986;
Prinzhofer et al., 1989; Chabaux & Alle`gre, 1994; Hirschmann & Stolper, 1996). The U–Th disequilibrium signature cannot be preserved in 90-my-old crust because
of the short half-life of 230Th. The 90-my-old crustal
samples have not been analyzed for Hf isotopic ratios.
Thus, we discuss the issue of garnet lherzolite vs garnet
pyroxenite sources for MORB by comparing the [Sm/
Yb]DM ratios in the crust and in calculated pooled melts
derived from lherzolite.
The composition of pooled melt derived from spinel
and plagioclase lherzolite in a passive melting regime
with an F max of 30% has an [Sm/Yb]DM of 1·16, whereas
that of a pooled melt with an F max of 30% that first
underwent 2% melting in the garnet stability field is 1·23.
The [Sm/Yb]DM in 90-my-old crust ranges from 1·18 to
1·33. Therefore, it is unlikely that melt from garnet
pyroxenite has made a significant contribution to these
samples of 90 Ma oceanic crust. However, many of the
EPR samples exhibit [Sm/Yb]DM ratios that range from
1·35 to 1·45 (Allan et al., 1989; Batiza & Niu, 1992; Niu
et al., 1996). These high values cannot be explained by
melting of a lherzolitic source if the results from the 90my-old RMC are valid. Melting of garnet pyroxenite
could thus be a viable model.
To have a first-order estimate of the proportion of
melt that would be derived from garnet pyroxenite to
produce EPR samples with [Sm/Yb]DM of 1·35–1·45, we
calculated Sm and Yb abundances in melts using the
average SOC (subducted oceanic crust) pyroxenite composition compiled by Hirschmann & Stolper (1996) for
the source. Following Hirschmann & Stolper (1996),
we modeled the melt compositions by modal fractional
melting, as little is known about pyroxenite melting
relationships. The mineral proportions were assumed to
be 20% garnet, 50% clinopyroxene and 30% orthopyroxene. Because garnet pyroxenites have lower solidus
temperatures than lherzolites, they begin to melt at
greater depths and undergo a greater extent of melting
than do the former (Hirschmann & Stolper, 1996). We
obtained melt compositions for 40%, 60%, and 80%
melting. The EPR samples with [Sm/Yb]DM of 1·35–1·45
can be produced by mixing ~20% of the melt generated
by melting SOC pyroxenite (40% melting) with ~80%
of pooled melt derived from a lherzolitic source in a
passive melting regime (case 1) (Fig. 10). The proportion
of melt from SOC pyroxenite increases to 40–50%, if
60% melting of SOC pyroxenite is used (case 2) (Fig.
10). Eighty percent melting of SOC pyroxenite produces
a melt with [Sm/Yb]DM of ~1·2, which cannot explain
the EPR samples with [Sm/Yb]DM of 1·35–1·45 (Fig. 10).
The crustal thickness calculated based on case 1 is ~7·1
km, whereas that based on case 2 is ~9·5 km. Therefore,
it appears that case 1 is a more reasonable model.
Although the details of pyroxenite melting may require
future experimental verification, our calculations do show
that involving melt from a garnet pyroxenite source may
help to explain the high Sm/Yb ratios of some zero-age
EPR MORB.
SUMMARY OF CONCLUSIONS AND A
MODEL
(1) We estimate a melt production rate for Oahu RMC
xenoliths in the spinel lherzolite stability field of 0·42%/
kbar, which is significantly lower than the values used in
most previous studies.
(2) Only ~2% melting could have occurred at pressures
>21 kbar if MORB has a depleted source. This limits
melting within the garnet lherzolite stability field to the
same amount, <2%.
(3) Although the melt production rate in the plagioclase
lherzolite field has not been determined from the xenoliths, the REE abundances in the 90 Ma crust and zeroage EPR lavas require a melt production rate of 4·4%/
kbar in the plagioclase lherzolite field.
(4) The high Sm/Yb ratio in the EPR crust cannot be
explained by melt derived from a lherzolitic source alone
because little melting could have occurred in the garnet
lherzolite field. Addition of a melt component derived
from garnet pyroxenite can account for the high Sm/
Yb ratios in EPR crust.
Based on these conclusions, Fig. 11 offers a scenario for
the development of this 90 Ma RMC. The asthenosphere
consisted of lherzolite with veins or pockets of pyroxenite
(Sen, 1983; Alle`gre & Turcotte, 1986). Because of the
lower melting temperature of pyroxenite (Hirschmann &
Stolper, 1996), they began to melt at greater depths than
the lherzolites did. The mantle parcel that rose at the
edge of the passive melting regime underwent smaller
extents of melting and accreted to form the lower part
of the lithosphere, thus preserving some pyroxenites. In
contrast, the mantle parcel that rose in the central part
of the melting regime underwent more extensive melting,
exhausting more pyroxenites, and eventually formed the
upper part of the RMC. Lherzolite began to melt at P
~26 kbar and underwent a melting rate of 0·42%/kbar
from 26 to 7 kbar, generating 3·9 km of crust. Melting
continued in the plagioclase lherzolite stability field,
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Fig. 10. Mixing between pooled melt derived from melting lherzolite in a passive melting regime and melts from garnet pyroxenite in the
[Sm]DM–[Sm/Yb]DM space. DM indicates normalized to the depleted mantle source (Table 3). Three mixing curves are shown. They are mixing
between pooled melt with F max = 30% and melts derived from 40%, 60% and 80% melting of the average garnet pyroxenite (average subducted
oceanic crust of Hirschmann & Stolper (1996)]. Each symbol on the mixing curves indicates 10% increment of a mixing end-member. The
shaded field are EPR samples (Allan et al., 1989; Batiza & Niu, 1992) after addition of 25% olivine.
Fig. 11. Our model: a schematic diagram of a passive melting regime in which lherzolitic mantle with veins or pockets of garnet pyroxenites
melts. The dashed lines indicate the flow paths of ascending mantle parcels. Because of their low melting point, garnet pyroxenites (ellipses filled
with small dots) start melting deeper than lherzolites do. As a result of the small proportions in the mantle and large extents of melting, only a
small amount of pyroxenites remains in the spinel stability field. Also shown to the right is the resulting residual mantle column.
292
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MID-OCEAN RIDGE MELTING
where it melted at a dF/dP of 4·4%/kbar and reached
an F max of 30%. Melting in the plagioclase lherzolite field
probably produced ~1·8 km of crust. Therefore, the total
melting of a lherzolitic source generated a ~5·7 km thick
crust. The high Sm/Yb ratios of many EPR MORBs
require an additional melt contribution from a garnet
pyroxenite source.
ACKNOWLEDGEMENTS
We benefited from discussions with Dean Presnall, Jim
Natland, Jackie Dixon, Rosemary Hickey-Vargas, Kevin
Johnson, Paul Asimow, Glenn Gaetani, John Lassiter
and Alberto Saal. We particularly thank Dean Presnall
and Fred Frey for their constructive comments on an
early version of manuscript. Thorough reviews by Kevin
Johnson, Rodey Batiza and Tom Wright, and editorial
effort by Sorena Svea Sorensen are gratefully acknowledged. H.-J. Yang thanks Ken Koga and Debra
Hassler for their assistance with ion-probe analyses. This
research was supported by NSF Grant OCE-9520409 to
G. Sen.
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