ABSTRACT. The petrogenesis of pyroxenite layers within the Beni Bousera peridotite massif is investigated by means of elemental and Nd-Sr-Pb-O-S isotope ...
0022-3530/93 $3.00
Geochemical Constraints on the Petrogenesis of Diamond Fades Pyroxenites from the Beni Bousera Peridotite Massif, North Morocco by D. G. PEARSON*, G. R. DAVIESf, AND P. H. NIXON Department of Earth Sciences, The University of Leeds, Leeds LS2 9JT, UK (Received 3 February 1992; revised typescript accepted 16 June 1992)
ABSTRACT The petrogenesis of pyroxenite layers within the Beni Bousera peridotite massif is investigated by means of elemental and Nd-Sr-Pb-O-S isotope analyses. The light rare earth element (LREE) depleted nature of many of the pyroxenites, their wide variation in composition, and lack of correlation between incompatible elements and fractionation indices preclude them from representing crystallized melts from a peridotitic source. The physical characteristics of the pyroxenites and their large (greater than a factor of 20) range in Ni rule out partial melting as the cause of their petrological and geochemical diversity. Major and compatible trace element geochemistry is consistent with formation of most of the pyroxenite suite via high-pressure crystal segregation in magma conduits intruding the peridotites. These magmas crystallized clinopyroxene, orthopyroxene, and garnet. The pressure of crystallization is constrained to be above ~ 4 5 kbar from the presence of graphitized diamonds in pyroxenite layers. Lack of correlation between fractionation indices and highly incompatible elements and the wide variation in incompatible element abundances suggest that the suite did not form from genetically related magmas. The presence of positive and negative Eu anomalies (Eu/Eu* = 054-2-0) in pyroxenites which crystallized at pressures much greater than the plagioclase stability field (~ 45 kbar) suggests that the parental magmas originated from precursors which formed in the crust. Oxygen isotope compositions of coexisting minerals in the pyroxenites indicate high-temperature equilibration but (51SO values vary from +4-9 to + 9-3%o, ruling out their derivation from the host peridotites or other normal mantle sources. The extreme O-isotope variation, together with 2 m thick contain graphitized diamonds (Slodkevich, 1983; Pearson et al, 1989) and have up to 15% graphite locally. Graphite octahedra contain faceted inclusions of GT and CPX, and constrain the depth of origin of pyroxenites to be in the diamond stability field (> 45 kbar). All the pyroxenites display exsolution and porphyroclastic to equant metamorphic textures, as a result of sub-solidus recrystallization and equilibration. Sub-solidus equilibration in GT-CPX cores records temperatures of ~ 1100 °C for most pyroxenites (Pearson et al, in press) at pressures of ~20 kbar based on the occurrence of Ca-Ts-rich CPX with corundum (Kornprobst et al, 1990). Peridotite/pyroxenite relationships Contact between pyroxenite layers and peridotites are generally sharp and nongradational, layer boundaries being parallel or sub-parallel to any foliation developed in the host peridotites. Layer boundaries are generally planar, possibly because of layer-parallel shearing during folding. However, small rounded apophyses are sometimes observed penetrating the surrounding peridotites; this feature indicates that the layers were formed by igneous processes. Deformation in the peridotites generally increases towards the contact with a pyroxenite layer. Olivines display a high density of subgrain boundaries and severe grain size reduction. Green chrome-diopsides are present at the extreme margins of some Alaugite pyroxenite layers, indicating peridotite-pyroxenite equilibration at high temperature.
PYROXENITE GEOCHEMISTRY Major elements One of the fundamental questions we wish to address is the relationship between the peridotites and the pyroxenites. Major and trace element analyses (Pearson, 1989, in prep.) indicate that the peridotites represent residues of variable amounts (5-35%) of partial melt extraction in the garnet and spinel stability fields. Some of the peridotites have also experienced FeO, Na2O, and light rare earth element (LREE) enrichment, possibly related to interaction with pyroxenite-related magmas. Very thin pyroxenite layers have lower SiO2 and A12O3 and high MgO, apparently as a result of equilibration with the host peridotites. Pyroxenite interaction with the peridotites is
130
D. G. PEARSON ET AL.
also indicated by an increase in mg-number and Cr content of silicate minerals in the pyroxenites at the pyroxenite-peridotite contacts. Similar effects have been described in detail from composite xenoliths (Irving, 1980) and peridotite massifs (Nielson et al, in press). Additionally, as garnet pyroxenites become progressively thinner the pyrope content of their garnet increases markedly. To avoid these effects, geochemical analyses were performed on pyroxenite samples taken from the centre of layers > 10 cm thick, unless otherwise stated. Major element compositions of some garnet pyroxenites (Table 1) superficially resemble those of picritic basalts, with MgO varying from 10-8 to 19-5%, but the pyroxenites have very low abundances of K2O and P 2 O 5 . Overall, the suite shows large variations in MgO, CaO, and A12O3; m^-number [100Mg/(Mg + Fe)] varies from 58 in the garnet clinopyroxenites to ~90 in the websterites and orthopyroxenites. The suite shows an iron enrichment trend and plots within thefieldsdefined by mafic and ultramafic cumulates from ophiolite massifs on an AFM plot (Fig. 2). Kornprobst et al. (1987, 1990) defined a separate group of pyroxenites (Group II) which had much higher Mg/Fe ratios than the main pyroxenite suite. Group II included corundum-bearing pyroxenites containing >21% A12O3. Partial melting of a peridotite source would produce a series of liquids andresidueswhose compositions are governed by the degree of melting. When plotted against a fractionation index such as m#-number or MgO (Fig. 3), the pyroxenite major element data show a considerable degree of scatter. If we assume that the Beni Bousera peridotites define a series of melting residues from a fertile source peridotite similar to the Ronda peridotite R717 analysed by Frey et al. (1985), we may define an 'extract line' between a residual peridotite, the fertile source (R717), and the calculated equilibrium melt. This extract line is valid only if there is no exhaustion of a phase during partial melting; this is the case in all the peridotites analysed in this study, which were four-phase assemblages that always contained CPX.
0 Graphite garnet pyroxenites •
Garnet pyroxenites Orthopyroxenites
Oceanic ^ Tholeilte
^
\o
Websterites
FIG. Z AFM diagram showing Fe enrichment in pyroxenites. Area enclosed by dashed lines is the field for ophiolitic cumulates from Coleman (1977); average compositions for oceanic tholeiites and gabbros are from Elthon (1979).
PETROGENESIS OF DIAMOND FACIES PYROXENITES
131
20
10
20
30
40
50
MgO wt% FIG. 3. Anhydrous A12O3 and CaO vs. MgO for pyroxenites (symbols as in Fig. 2) and peridotites (closed triangles). Solid line represents an 'extraction' line between a Ronda peridotite of similar composition to pyrolite as a source (R7I7 of Frey et ai, 1985), an average Beni Bousera peridotite as a residue, and the calculated equilibrium melL
Figure 3 depicts extract lines calculated between the assumed source (R717) and an average Beni Bousera peridotite (data from Pearson, 1989). Few of the pyroxenite analyses plot near this 'extract line' or the calculated equilibrium melt, indicating that they are unlikely to represent simply frozen melts of the peridotites. Trace elements Ni and Cr show positive correlations with m#-number (Fig. 4), although correlation within discrete lithological groups is poor. The overall range in Ni abundance is very large, both for the entire suite (50-1368 ppm) and within individual lithological groups (50-550 ppm in the garnet pyroxenites), indicating that partial melting alone cannot account for their genesis and that some sort of fractional crystallization processes must have occurred. The graphitic garnet pyroxenites (GGP) have very low Ni abundances, 50-200 ppm, considering their relatively high MgO contents. Ni and Cr abundances increase towards layer margins in
FIG. 4. Compatible elements vs. mg-number for pyroxenites. Arrows show the effect of addition of garnet, clinopyroxene, and orthopyroxene, and are estimated from considerations of partition coefficients and the instrumental neutron activation analysis (INAA) data of Bodinier et al. (1987). Symbols as in Fig. 2.
2000
o z
in
•v m
D O
Sc Zr
Ba V
Y
Rb Sr
Zn
Co Ni Cu
1000 63 522 57 88 0 4-2 23 27 233 46 21
48-76 083 11-26 1119 0-27 14-86 11-37 1-38 004 005 7029
SiO 2 TiO 2 AI 2 O 3 FeO MnO MgO CaO Na 2 O K2O P2O5 mg-no.
Cr
GPI9 GP
Sample Lithology Thickness
728 63 365 58 98 1 22-5 23 58 250 56 9
47-86 075 11-84 1117 027 15-55 11-98 111 004 004 71-26
GP20 GP 170
458 53 50 19 94 0 1 66 46 5 580 76 9
46-68 079 13-66 13-88 033 11-37 11-87 1-37 001 004 59-33
GP25 GGP Float
1978 45 686 103 19 0 44 25 7 236 81 37
46-32 043 15-73 5-91 021 2O30 1013 092 003 003 85-95
GP28 WEB 30
2134 51 778 93 47 0 48-5 17 25 264 63 30
49-38 073 1098 7-46 017 18-39 11-52 1-32 001 004 76-97
46-54 058 9-77 8-33 016 22-97 1053 1-09 001 003 83-08 1942 67 1012 43 63 0 53 15 6 164 35 20
GP33a WEB 70
GP30 WEB 30
2130 55 597 105 37 0 45 24 3 233 59 40
45-82 055 16-38 8-39 020 15-74 11-80 107 000 005 76-98
GP33b WEB 70
929 66 364 6 66 0 169 18 35 364 62 8
45-45 065 1312 11-82 022 13-30 13-86 1-52 002 004 66-70
GP37 GP 50
7-2 30 15 354 57 12
0
850 48 139 20 77
4802 084 12-68 11-33 032 14-51 1026 1-98 002 004 69-52
GP47 GGP Float
305 59 79 42 78 1 41 45 46 658 80 9
46-91 059 13-59 13-81 030 1107 1213 1-50 007 004 58-81
GP8I GGP 2-5
Major and trace element analyses of Beni Bousera pyroxenites by X-ray fluorescence
TABLE 1
1587 51 487 29 52 0 53 17 0 287 115 33
46-38 057 14-81 813 019 17-41 11-81 067 0-00 004 79-23
GP87T GP
1499 53 351 21 42 0 381 16 7 253 59 17
46-34 067 13-66 911 019 14-80 14-33 086 000 004 74-31
18
GP
GP87M
to
m
X m Z
o
73
•
n
D > O
•n
o
to
00
O O m Z m
SO
H
m
•o
Ba V Sc Zr
Cr Co Ni Cu Zn Rb Sr Y
P2O3 mg-no.
K2O
Na 2 O
CaO
MgO
86-3 9 14 247 63 15
0
1906 51 402 41 34
46-45 0-61 13-41 8-40 0-18 16-98 13-03 0-92 001 002 78-26
SiO 2 TiO 2 A1 2 O, FeO
MnO
GP87B GP
Sample Lithology Thickness
47 3
108
12
6
943 51 349 102 30 0 16-3
45-95 0-17 16-68 6-26 0-13 14-43 13-89 2-57 002 0O3 80-40
GP97 WEB 50
121 3 7 187 38 4
0
3832 39 963 58 43
51-93 0-10 7-24 4-22 0-12 20-75 13-73 2-25 003 003 89-87
WEB 10
GPlOla
645 56 219 82 53 0 43 13 14 256 58 5
45-62 0-39 16-21 9-50 019 12-06 14-79 1-22 000 002 69-33
GP200 GP 20
277 46 22
1234 66 567 70 80 2 38 10 12
ooo ooo
73-20
48-29 066 11-91 1O80 019 16-55 1039 1-21
GP204 GP 50
923 48 399 145 30 0 77 4 10 124 44 1
15-73 13-73 1 39 OOO 002 81-83
014
15-88 6-22
018
46-71
WEB 16
GP125
TABLE 1 (Continued)
3562 79 1368 64 63 0 1 2 2 123 19 1
021 000 OOO 90O0
210
5409 005 5-87 619 013 31-36
GP130 WEB 10
9775 78 1799 73 62 0 27 2 0 124 35 2
34-44 619 054 002 002 9O96
Oil 6O8 610 014
46-38
3568
53-28 044 407 7-54 018 32-43 1-75 O31 OOO OOO 88-45
GP 132(1) GP137 WEB OPX1TE 20 05
57
49 0 352
903 55 188 32 79 0 40
48-30 1-34 12-42 1097 028 14-35 1020 207 001 007 69-97
GP13S GP
35 7 504 69 15
769 58 156 43 95 0 26-8
12-67 11-38 1-82 001 005 64-35
028
12-70 12-46
085
47-80
250
GP139 GP
Z
o
•v m >
9
1-96 001 003 72-62
819 52 176 35 67 0 12 24 9 397 55 23
Cr Co Ni Cu Zn Rb Sr Y Ba V Sc Zr
27 3 379 57 8
827 50 128 27 81
47-63 0-83 13-04 11-87 0-38 14-97 9-54 1 70 0O1 003 69-19
48-87 0-71 12-38 10-34 O30 15-40
SiO 2 TiOj AljO 3 FcO MnO MgO CaO Na 2 O K2O P2O5 mg-no.
1001
GP143 GP
GP140 GP
Sample Lithology Thickness
861 53 138 24 80 0 2-54 30 0 377 58 7
47-48 079 13-38 12-23 039 1511 9-42 115 OOO O05 68-75
GP147 GGP 260
12
946 57 192 33 79
48-99 081 12-36 11-25 034 14-64 9-67 1-89 OOO 005 69-81
GP148a GGP
1767 61 713 69 47 3 65 16 4 159 33 27
47-07 048 13-66 8-20 017 19-70 9-40 1-27 O01 004 6O30
52-40 025 5-29 8-54 019 24-34 816 081 002 OOO 83-54 3234 66 715 12 59 0 11 6 17 404 45 3
GP183 GP 25
GP170 WEB Float
TABLE 1 (Continued)
3493 78 1289 49 66 0 1-95 0 0 121 18 0-
54-44 006 5-68 6-27 013 31-28 1-91 021 OOO 002 89-87
GP188 WEB 15
2276 43 408 38 36 0 63 15 10 204 50 18
48-81 060 11-28 6-90 017 17-73 14-03 118 019 002 82-07
GP194M GP 14
2451 49 543 12 32 1 51 18 0 242 59 27
47-55 083 1097 810 002 1911 12-45 093 001 003 8O78
GP196 GP 10
580 65 194 37 53 1 13 28 7 380 74 9
46O6 034 15O0 12-29 029 1402 1048 1-48 001 003 6702
GP207 GP 30
1026 65 478 88 65 2 20 14 0 234 53 16
47-41 045 13-74 1076 022 15-31 1088 1-22 000 003 71-71
GP231 GP 20
tfl
z
m
O X
73
•
p o
PETROGENESIS OF DIAMOND FACIES PYROXENITES
137
mono-lithological layers (e.g., GP87B and T vs. 87M; Table 1). This may reflect crystallization of the marginal phases from a less fractionated liquid, or the effects of sub-solidus chemical interaction with the peridotites, or a combination of both processes. Elements such as Sc, V, and Y show scattered negative correlations with m#-number which appear to be controlled by the modal abundance of primary phases (Fig. 4). For example, the most garnet- and CPX-rich rocks have the highest Sc contents, consistent with the high partition coefficient of Sc in these two minerals. Incompatible elements such as Zr, Sm, and Sr show no correlation with mg-number, a feature also observed by Suen & Frey (1987) for Ronda pyroxenites. Some of the Sr variability may be due to post-crystallization alteration by crustally derived fluids, as suspected by Polve & Allegre (1980) and Zindler et al. (1983). This is well illustrated in some cases by the disparity between whole-rock Sr concentrations, measured by isotope dilution, and whole-rock concentrations computed from calculated modal abundances and analysis of acid-washed mineral separates by isotope dilution and ion probe. One of the most striking examples is GP81, with a measured whole-rock Sr concentration of 42 ppm and a calculated whole-rock Sr concentration of only 2-46 ppm, indicating the presence of a considerable amount of leachable, grain boundary Sr (see below). This apparent secondary Sr mobility is not ubiquitous and Sr concentrations in most specimens do not appear to have been significantly affected. Calculated Nd concentrations are generally in agreement with measured values, e.g., GP81 measured whole-rock Nd is 0-42 ppm whereas the calculated value is 043 ppm. REE geochemistry Whole rocks The chemical diversity of the pyroxenites is further illustrated by their chondritenormalized REE patterns, which show wide variations in REE abundance (Cen 0-05-3-3, Yb n 0-56-28-4) and highly variable LREE/HREE depletion. Many pyroxenites have lower LREE abundances than the peridotites (Fig. 5). Most of the garnetiferous pyroxenites are HREE enriched, with Ybn > 20; consequently, the parameter (Ce/Yb)n is not a realistic indicator of LREE depletion, and (Ce/Sm)n is used in preference. Whole-rock (Ce/Sm)n values vary from slightly LREE depleted to highly LREE depleted (0-91-0-016; Table 2). All GGP samples and GP139 are characterized by extreme LREE depletion. Many samples show significant Eu anomalies (Eu/Eu* 0-54-20) despite containing no primary plagioclase. Ion probe analyses of plagioclase neoblasts and exsolution 'blebs' in CPX from PHN5731 and 5734 N. Shimizu & G. Pearson, unpub. data) reveal very low Sr contents of ~ 2 ppm, similar to the porphyroclastic CPX and confirming the sub-solidus origin of the minor plagioclase found in these rocks. Kornprobst et al. (1987, 1990) proposed that pyroxenites showing LREE depletion, low HREE abundances, and positive Eu anomalies (his Group II) represent relatively unmodified oceanic lithosphere directly recycled back into the convecting mantle. Four such samples were encountered during this study. Garnet websterite PHN5739 and garnet clinopyroxenite GP37 have low HREE abundances (Ybn 4-5 and 3-4) although they contain abundant garnet PHN5739 also has a large positive Eu anomaly (Eu/Eu* = 1-6). However, neither PHN5739 or GP37 could be classified as Type II pyroxenites on the basis of their major element chemistry, as they are too enriched in Fe. Garnet websterites GP97 and GP125 also have similar REE patterns to Group II pyroxenites. Despite containing up to 40% original garnet, GP97 and GP125 have two of the lowest Yb abundances (2-9 and 2-5 times chondritic, respectively), low Y abundances compared with the other garnetiferous pyroxenites, and also contain significant positive Eu anomalies (Fig. 5).
0.01
0.1 - .
1 -.
-
0.1
0.5 -
1 —
5
10 —
0.1
1 —
La C» Pr Nd
La Cs Pr Nd
Websterites
Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
Garnet
•i
GP97 GP125
z
O
VI
70
m >
•a
o
PHN5739
Sm Eu Qd Tb Dy Ho Er Tm Yb Lu
GP188
GP28 GP33a PHN5732 GP236
Fio 5 Chondrite-normalized REE patterns for selected whole-rock pyroxenites by isotope dilution. Garnet websterites (open squares) are chemically equivalent to the Group II pyroxenites of Komprobst et al. (1990). Normalization factors are after Nakamura (1974). Shaded field is range for Bern Bousera peridotites (Menzies et aU 1977; Loubet & Allegre, 1982; Pearson, 1989).
X
O
o
GP139
•5
Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
Sm Eu Qd Tb Dy Ho Er Tm Yb Lu
X
o
La C* Pr Nd
La Ca Pr Nd
O O
•o
100 -a
0.01
100-9
139
PETROGENESIS OF DIAMOND FACIES PYROXENITES
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FlG. 9. Nd-Sr isotope variation in CPX from peridotites and pyroxenites compared with fields for MORB, OIB, and the Ronda peridotites [from Reisberg & Zindler (1987) and Reisberg et al. (1989)]. Data are initial values corrected to an emplacement age of 21 Ma. Shaded field represents mixing between a magma in equilibrium with pyroxenite GP30 and the surrounding orust (kinzigite; data from Table 4).
1 44
143
0.5135
0.5140
D. G. PEARSON ET AL.
152
0.5140 " 143 144
Nd
Nd 0.5130
0.5120 0.0
0.5
1.0
147
0.712
I
Srn/ -
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144
2.0
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0.708 0.245
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0.024 ^_
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|
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0.000
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0.010
87
0.015
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Rb/ 8 6 Sr
20.0
206 204
Pb
19.0
Pb
18.0 20.0
10.0 »»U/
204
Pb
FIG. 10. Sm-Nd, Rb-Sr, and U-Pb isochron diagrams for CPX from pyroxenites. Intersection of lines on Sm-Nd diagram defines the estimated Bulk Earth. Symbols as in Fig. 9.
significantly displaced from the NHRL on a plot of 2 0 7 Pb/ 2 0 4 Pb vs. 2 0 6 Pb/ 2 0 4 Pb but is below the NHRL in terms of 2 0 8 Pb/ 2 0 4 Pb vs. 2 0 6 Pb/ 2 0 4 Pb. 2 3 8 U/ 2 0 4 Pb (n) values of the CPXs are very variable (2-2-20-4) and are uncorrelated with either 2 0 6 Pb/ 2 0 4 Pb or 207 Pb/ 2 0 4 Pb.
153
PETROGENESIS OF DIAMOND FACIES PYROXENITES 41.0
Marine Seds
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15.4 17.5
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FIG. 11. Pb isotope plots of CPX in pyroxenites. Also plotted is the field for oceanic sediments [data from White & Dupre (1985) and Ben Othman el a\. (1989)]. Dashed lines join leachates (crosses) with residues (symbols as in Fig. 10). Solid line indicates mixing between a parental pyroxenite magma with an isotopic composition of GP30 and Pb content of 1 ppm, with hemi-pelagic sediment V34-47 from Ben Othman el al. (1989). Numbers by tick marks indicate percentage of sediment in mix. NHRL is the Northern Hemisphere Reference Line defined by Hart (1984).
Oxygen isotopes CPXs from the Beni Bousera peridotites have relatively homogeneous O isotope compositions (45 kbar. Wide variation in incompatible element abundances suggests that the pyroxenite 'suite' did not crystallize from cogenetic melts. The presence of Eu anomalies and their heterogeneous oxygen isotope compositions indicate that the parental melts were derived from crustal precursors, probably hydrothermally altered oceanic crust. Pb and Nd isotope systematics require that some pyroxenites have incorporated a component with high time-integrated Th/Pb, and U/Pb, and low Sm/Nd in their source. This component is interpreted to be hemi-pelagic sediment, probably incorporated during subduction of oceanic crust. 5 km) seawater hydrothermal alteration at mid ocean ridges. J. Geophys. Res. 86, 2737-55. 1986. Non-equilibrium, metasomatic " O / 1 6 O effects in upper mantle mineral assemblages. Contr. Miner. Petrol. 93, 124-35. Hamelin, B., & Allcgre, C. J., 1988. Lead isotope study of orogenic lherzolite massifs. Earth Planet Sci. Lett. 91, 112-31. Hart, S. R., 1984. A large-scale isotope anomaly in the Southern Hemisphere mantle. Nature 309, 753-7. Hatton, C. J, 1978. The geochemistry and origin of xenoliths from the Roberts Victor Mine. Unpublished Ph.D. Thesis, University of Capetown, South Africa, 179 pp. Henderson, P., 1981 Inorganic Geochemistry. London: Pergamon Press, p 91. Hooker, P. J, O'Nions, R. K , & Pankhurst, R. J, 1975. Determination of rare earth elements in U.S.G.S. standard rocks by mixed solvent ion exchange and mass spectrometric isotope dilution. Chem. Geol. 16, 189-96. Irving, A. J., 1980. Petrology and geochemistry of composite ultramafic xenoliths in alkalic basalts and implications for magmatic processes in the mantle. Am. J. Sci. 280a, 389-426. Ito, E., White, W. M., & Gopel, C , 1987. The O, Sr, Nd and Pb isotope geochemistry of MORB. Chem. Geol. 62, 177-89. Jagoutz, E-, 1988. Nd and Sr systematics in an edogite xcnolith from Tanzania: evidence for frozen mineral equilibria in the continental lithosphere. Geochim. Cosmochim. Acta 52, 1285-93. Dawson, J. B_, Hoernes, S., Spettel, B., & Wanke, H , 1984. Anorthositic oceanic crust in the Archaean Earth. Lunar and Planetary Institute, Houston, TX. LPI Tech. Rep. 85-01, 40-41. Johnson, K. T. M , Dick, H. J. B., & Frey, F. A, 1990. Melting in the oceanic upper mantle: an ion microprobe study of diopsidcs in abyssal peridotites. / . Geophys. Res. 95, 2661-78. Kellog, L. H, & Turcotte, D. L_, 1986. Homogenisation of the mantle by convective mixing and diffusion. Earth Planet. Sci. Lett. 81, 371-8. KimbalL K. L-, & Gerlach, D. C , 1986. Sr isotopic constraints on hydrothermal alteration of ultramafic rocks in two oceanic fracture zones from the South Atlantic Ocean. Ibid. 78, 177-88.
PETROGENESIS OF DIAMOND FACIES PYROXENITES
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Kornprobst, J., 1969. Le massif ultrabasique des Beni Bouchers (Rif Interne, Maroc). Contr. Miner. Petrol. 23, 283-322. 1970. Les peridotites et les pyroxenolites du massif ultrabasique des Beni Bouchera: une etude experimentale entre 1100 et 1550 °C sous 15 a 30 kb de pression seche. Ibid. 29, 209-309. 1974. Contribution a l'etude petrographique et structurale de la zone interne du Rif (Maroc Septentrional). Notes Serv. G'eol. Maroc. T 251, 256 pp. Piboule, M., Roden, M., & Tabit, A, 1990. Corundum-bearing garnet clinopyroxenites at Beni Bousera (Morocco): original plagiodase-rich gabbros recrystallised at depth within the mantle? J. Petrology 31,717-45. Tabit, A-, 1987. Diversite des clinopyroxenites a grenat associees aux massifs ultramafiques orogeniques: eclogitcs, ariegites, griquaites et grospydites; une discussion. Bull. Soc. Geol. Fr. 2(8), 345-51. Kyser, T. K., 1990. Stable isotopes in the continental lithospheric mantle. In: Menzies, M. A. (ed.) Continental Mantle. Oxford: Qarendon Press, 127-56. CNeiL J. R., & Cannichael, I. S. EL, 1986. Reply to "Possible non-equilibrium oxygen isotope effects in mantle nodules, an alternative to the Kyser-CNeil-Carmichael 1 8 O/ 1 6 O geothermometer". Contr. Miner. Petrol. 93, 120-3. Lesher, C. E , 1990. Decoupling of chemical and isotopic exchange during magma mixing. Nature 344, 235-7. Loomis, T. P., 1975. Tertiary mantle diapirism, orogeny and plate tectonics east of the Strait of Gibraltar. Am. J. Set. 275, 1-30. Lorand, J. P., 1985. The behavior of the upper sulphide component during the incipient alteration of'Alpine'-type peridotites as illustrated by the Beni Bousera (N. Morocco) and Ronda (S. Spain) ultramafic bodies. Tschermaks Miner. Petrogr. Mitt. 34, 183-209. Loubet, M , & Allegre, C. J., 1979. Trace element studies in the Alpine type peridotite of Beni Bouchera (Morocco). Geochem. J. 13, 69-75. 1982. Trace elements in orogenic lherzolites reveal the complex history of the upper mantle. Nature 298, 809-11. Luck, J. M , & Allegre, C. J., 1990. Osmium isotopes and mantle processes. Terra Abstr. 2, 10. MacGregor, I. D., & Manton, W. I., 1986. The Roberts Victor eclogites: ancient oceanic crust J. Geophys. Res. 91, 14063-79. Machado, N., Brooks, C , & Hart, S. R., 1986. Determination of initial 87 Sr/" 6 Sr and '* 3 Nd/ 1 M Nd in primary minerals from mafic and ultramafic rocks: experimental procedure and implication for the isotopic characteristics of the Archean mantle under the Abitibi greenstone belt, Canada. Geochim. Cosmochim. Ada 50, 2335-^8. Martin, C. E., 1991. Osmium isotopic characteristics of mantle-derived rocks. Ibid. 55, 1421-34. McCulloch, M. T., Gregory, R. T., Wasserburg, G. J , & Taylor, H. P , 1981. A neodymium, strontium and oxygen isotopic study of the Cretaceous Semail ophiolite and implications for the petrogenesis and seawaterhydrothermal alteration of oceanic crust Earth Planet. Sci. Lett. 46, 201-11. McKay, G. A., 1989. Partitioning of rare earth elements between major silicate minerals and basaltic melts. In: Lipin, B. R., & McKay, G. A. (eds.) Geochemistry and Mineralogy of Rare Earth Elements. Miner. Soc. Am. Rev. Miner. 21, 45-77. McKenzie, D., & Bickle, M. J., 1988. The volume and composition of melt generated by extension of the lithosphere. J. Petrology 29, 625-79. Meijer, A., Kwon, T. K., & Tilton, G. R_, 1990. U-Th-Pb partitioning behavior during partial melting in the upper mantle: implications for the origin of High Mu components and the "Pb paradox". J. Geophys. Res. 95,433-48. Menzies, M. A., 1984. Chemical and isotopic heterogeneities in orogenic and ophiolitic peridotites. In: Gass, L G-, Lippard, S. J., & Shelton, A. W. (eds.) Ophiolites and Oceanic Lithosphere. London: Black well, 231-40. 1990. Archaean, Proterozoic and Phanerozoic lithospheres. In: Menzies, M. A. (ed.) Continental Mantle. Oxford: Clarendon Press, 66-86. Blanchard, D_, & Korotev, R., 1977. Rare earth element geochemistry of fused ophiolitic and alpine Iherzolites II. Beni Bouchera, Ronda and Lanzo. Contr. Miner. Petrol. 64, 53-74. Murthy, V. R, 1980. Enriched mantle: Nd and Sr isotopes in diopsides from kimberfa'te nodules. Nature 283,634-6. Michard, A., & Albarede, F , 1985. Hydrotbermal uranium uptake at ridge crests. Ibid. 317, 244-6. Mukasa, S. B., Shervais, J. W., Wilshire, H. Gn & Nielson, J. E ? 1991. Intrinsic Nd, Pb and Sr isotopic heterogeneities exhibited by the Lherz alpine peridotite massif, French Pyrenees. J. Petrology, Special Lherzolites Issue, 117-34. Nakamura, N , 1974. Determination of REE, Ba, Fe, Mg, Na and K in carbonaceous and ordinary chondrites. Geochim. Cosmochim. Ada 38, 757-75. Neal, C. R., Taylor, L. A^ Davidson, J. P., Holden, P , Halliday, A. N , Nixon, P. H., Paces, J. B., a a y t o n , R. N^ & Mayeda, T. K , 1990. Eclogites with oceanic crustal and mantle signatures from the Bellsbank kimberlite. South Africa, part 2: Sr, Nd and O isotope geochemistry. Earth Planet. Set Lett. 99, 362-79. Nicholls, I. A., & Harris, K. L , 1980. Experimental rare earth partition coefficients for garnet, clinopyroxene and amphibole coexisting with andesitic and basaltic liquids. Geochim. Cosmochim. Ada 44, 287-308. Nielsen, J. E , Shervais, J. W., Wilshire, H. G., & Mukasa, S. B , in press. Petrology and geochemistry of the Lherz Alpine peridotite. J. Petrology. Nixon, P. H., Davies, G. R., Slodkevich, V. V, & Bergman, S. C , 1986. Graphite pseudomorphs after diamond in the edogite peridotite massif of Beni Bousera, Morocco, and a review of anomalous diamond occurrences. Fourth Int. KimberlUe Conf. Perth, Extended Abstr. Geol. Soc. Aust. No. 16, 412-14.
170
D. G. P E A R S O N ET
AL.
O'Hara, M. J., & Yoder, H. S., Jr., 1967. Formation and fractionation of basic magmas at high pressure. Scott. J. Ceo/. 3, 67-117. Patriat, P., Segoufin, J., Schlich, R., Goslin, J., Auende, J. M., Benzart, P., Bonnin, J., & Olivet, J. L, 1981 Les mouvements relatifs de l'lnde, de 1'Afrique et de l'Eurasie. Bull. Soc. Giol. FT. (7), XXIV (2i 363-73. Pearson, D. G, 1989. The petrogenesis of pyroxenites containing octahedral graphite and associated mafic ultramafic rocks of the Beni Bousera peridotite massif, North Morocco. Unpublished Ph.D. Thesis, University of Leeds, 412 pp. Davies, G. R., & Nixon, P. H., in press. Orogenic ultramafic rocks of ultrahigh-pressure (diamond facies) origin. In: Coleman, R. G., & Wang, X. (eds.) Vltrahigh Pressure Metamorphism. Cambridge: Cambridge University Press. Greenwood, P. B., & Mattey, D. P., 1991a. Oxygen isotope evidence for the origin of pyroxenites in the Beni Bousera peridotite massif, North Morocco: derivation from subducted oceanic lithosphere. Earth Planet. Sci. Lett. 102, 289-301. Mattey, D. P., 1991c. A carbon isotope study of diamond facies pyroxenites and associated rocks from the Beni Bousera peridotite massif, North Morocco. J. Petrology, Special Lherzolites Issue, 175-89. Milledge, H. J., 1989. Graphitized diamonds from a peridotite massif in Morocco and implications for anomalous diamond occurrences. Nature 338, 60-2. Shirey, S. B, Carlson, R. W., Boyd, F. R., Nixon, P. H., Pokhilenko, N. P., & Brown, L , 1991ft. Rhenium-osmium isotope systematics in southern African and Siberian peridotite xenoliths and the evolution of subcontinental lithospheric mantle. Extended Abstracts, 5th International Klmberlite Conference, Brazil, 329-33. Polve, M , 1983. Les isotopes du Nd et du Sr dans les lherzolites orogeniques: contribution a la determination de la structure et de la dynamique du manteau superieur. Unpublished Ph.D. Thesis, University of Paris 7. Allegre, C. J., 1980. Orogenic lherzolite complexes studied by 8 7 Rb- 8 7 Sr. A clue to understanding mantle convection processes? Earth Planet. Sci. Lett. 51, 71-81. Prinzhofer, A^ Lewin, E., & AUegre, C. J., 1989. Stochastic melting of the marble cake mantle: evidence from local study of the East Pacific Rise at 12°50TM. Ibid. 92, 189-206. Reisberg, L. C, AUegre, C. J., & Luck, J. M , 1991. The Re-Os systematics of the Ronda Ultramafic Complex of southern Spain. Ibid. 105, 196-213. Zindler, A., 1987. Extreme isotopic variations in the upper mantle: evidence from Ronda. Ibid. 81, 29-45. Jagoutz, E., 1989. Sr and Nd isotopic compositions of garnet and spinel bearing peridotites in the Ronda Ultramafic Complex. Ibid. 96, 161-80. Reuber, I., Mkhard, A~, Chalcouan, A., Juteau, T., & Jermoumi, B., 1982. Structure and emplacement of the Alpine type peridotites from Beni Bousera, Rif Morocco: a polyphase tectonic interpretation. Tectonophys. 82, 231-51. Roden, M. F., Irving, A. J., & Murthy, V. R., 1988. Isotopic and trace element composition of the upper mantle beneath a young continental rift: results from Kilbourne Hole, New Mexico. Geochim. Cosmochim. Ada 52, 461-73. Sacks, I. S., & JCincaid, C , 1990. On slab melting in subduction zones. EOS, Trans. Am. Geophys. Union 71, 1714. Sakai, H., Des Marias, D. J., Ueda, A., & Moore, J. G., 1984. Concentrations and isotope ratios of carbon, nitrogen and sulphur in ocean floor basalts. Geochim. Cosmochim. Ada 48, 2433-41. Schiflmann, P., Evarts, R. C , Williams, A. E., & Pickthorn, W. J., 1991. Hydrothermal metamorphism in oceanic crust from the Coast Range Ophiolite of California: fluid-rock interaction in a rifted island arc. In: Peters, T. J., Nicolas, A., & Coleman, R. G. (eds.) Ophiolite Genesis and Evolution of the Oceanic Lithosphere, Dordrecht: Kluwer, 399-426. Shimizu, H., 1980. Experimental study on rare earth element partitioning in minerals formed at 20 and 30 kb for basaltic systems. Geocnem. J. 14, 185-201 Sleep, N. H., 1988. Trapping of melt by veins and dikes. J. Geophys. Res. 93, 10255-71 Slodkevich, V. V., 1983. Graphite paramorphs after diamond. Int. Geol. Rev. 25(5), 497-514. Smyth, J. R., Caporuscio, F. A^ & McCormick, T., 1989. Mantle eclogites: evidence of igneous fractionation in the mantle. Earth Planet. Sci. Lett. 93, 133-41. Stoltz, A. J., & Davies, G. R, 1988. Chemical and isotopic evidence from spinel lherzolite xenoliths for episodic metasomatism of the upper mantle beneath southeast Australia. J. Petrology, Special Lithosphere Issue, 303-30. Stosch, H. C & Lugmair, G. W., 1986. Trace element and Sr and Nd isotope geochemistry of peridotite xenoliths from the Eifel (West Germany) and their bearing on the evolution of the subcontinental lithosphere. Earth Planet. Sci. Lett. 80, 281-98. Suen, C. J-, & Frey, F. A., 1987. Origins of the mafic and ultramafic rocks in the Ronda peridotite. Ibid. 85,183-202. Thirl wall, M. F., 1981 A triple filament method for rapid and precise analysis of rare earth elements by isotope dilution. Chen Geol. 35, 155-66. Veizer, J_, Compston, W., Clauer, N., & Schidlowski, MM 1983. 8 7 Sr/"Sr in late Proterozoic carbonates; evidence for a "mantle" event at ~900 Ma ago. Geochim. Cosmochim. Acta 47, 295-301 Vocke, R. D., Hanson, G. N., & Grunenfelder, M , 1987. Rare earth element mobility in the Roffina Gneiss, Switzerland. Contr. Miner. Petrol 95, 145-54.
PETROGENESIS OF DIAMOND FACIES PYROXENITES
171
Walker, R.J., Carlson, R. W., Shirey, S. B, & Boyd, F. R^ 1989. Os, Sr, Nd and Pb isotope systematics of southern African peridotites xenoliths: implications for the chemical evolution of subcontinental mantle. Geochim. Cosmochim. Acta 54, 1583-95. White, W. M., & Dupre, B., 1986. Sediment subduction and magma genesis in the Lesser Antilles: isotopic and trace element constraints. J. Geophys. Res. 91(86), 5927-41. Wilshire, H. G., & Shervais, J. W., 1975. Al-augite and Cr-diopside ultramafic xenoliths in basalt host rocks from the western United States. Phys. Chem. Earth 9, 257-72. Zeck, H. P., Monie, P., Villa, I. M., & Hansen, B. T., 1991 Very high rates of cooling and uplift in the Alpine belt of the Betic Cordilleras, southern Spain. Geology 20, 79-81 Zindler, A., & Hart, S. R., 1986. Chemical geodynamics. Ann. Rev. Earth Planet. Sci. 14, 493-571. Jagoutz, E., 1988. Mantle cryptology. Geochim. Cosmochim. Acta 52, 319-33. Staudigel, H., Hart, S. R., Endres, R., & Goldstein, S., 1983. Nd and Sr isotopic study of a mafic layer from Ronda ultramafic complex. Nature 304, 226-30.
APPENDIX: ANALYTICAL TECHNIQUES Geochemical analysis of the coarse-grained pyroxenites necessitated large sample sizes, generally > 2 kg, which were collected from areas showing the least mineralogical heterogeneity. Mineral separates were obtained, after cone and quartering of the initial coarse crush, by magnetic separation (GP147 garnet required the use of Clerici's solutionX and further purification by hand picking under alcohol, until pure glassy splits were obtained. Each separate was washed ultrasonically in sub-boiled quartz-distilled (SBQD) water followed by 30 min in 6 M HC1 at 20 °C, washed four times in SBQD water, 60 min in 6 M HC1 at 80 °C, 30 min in 40% HF at 20 °C, and another 30 min in 6 M HC1 at 20 °C. Separates were then washed in SBQD water, then distilled methanol, and dried under filtered air. Treatment of minerals for oxygen isotope analysis has been described by Pearson et al. (1991a). Major elements and non-REE trace elements were determined by X-rayfluorescence(XRF) at the University of Leeds; major elements were analysed from fused discs and trace elements from powder briquettes. All Zr data were duplicated at extended counting times, and data are reported as the mean of the two runs. Reproducibility (2 S.D.) based on separate preparations of representative pyroxenite compositions were: SiO2 0-7%, A12O3 0-6%, Fe2O31-5%, MgO 1%, CaO 0-5%, Na2O 8%, TiO2 4%, MnO 2%, K2O and P 2 O 5 30%, Rb, Sr, and Zr ~ 10% at the sub-10-ppm level but 5% between 10 and 40 ppm, Sc 4%, V 1%, Cr 10%, Ni 3%, Cu 3%, Zn 5%, Y 2%. Precisions (2 S.D.) for major elements are all 1% or better, except for Ti and Mn (2%), and Na (~ 5%) at the low abundances present. K and P are precise only to ~50% at the
