Petrology of ultramafic nodules from São Tomé Island, Cameroon Volcanic Line (oceanic sector)

Journal of African Earth Sciences 34 (2002) 231–246 www.elsevier.com/locate/jafrearsci

~o Tome Island, Petrology of ultramafic nodules from Sa Cameroon Volcanic Line (oceanic sector) R. Caldeira

a,*

, J.M. Munh a

b

a

b

Centro de Geologia do Instituto de Investigacßa~o, Cientıfica Tropical (IICT) e Centro de Geologia da Univ. Lisboa, Alameda D. Afonso Henriques 41-4° D, 1000-123 Lisbon, Portugal Departamento e Centro de Geologia da Faculdade de Ci^ encias da Universidade de Lisboa, Campo Grande, 1749-016 Lisbon, Portugal Received 4 July 2000; accepted 1 November 2001

Abstract Ultramafic xenoliths were found in recent alkali basalts from S~ ao Tome Island. These include spinel peridotites (lherzolites, harzburgites and dunites) and pyroxenites (orthopyroxenites and clinopyroxenites). Textures and mineral compositions indicate that pyroxenites originated from crystal/liquid separation processes operating on magmas similar to those giving rise to their present host rocks whereas spinel peridotite xenoliths had an accidental origin; Fo (>89) and Ni (>0.36 wt.%) contents in olivines, Mg# (91–95) of orthopyroxenes and low Ti in clinopyroxene (primary crystals: TiO2 < 0:06 wt.%) and in spinel (TiO2 < 0:1 wt.%) are within the range reported for abyssal peridotites, indicating S~ao Tome spinel peridotites represent refractory residues of melting. Nevertheless, the lack of correlation between mineral chemistry and modal composition suggests that spinel peridotite xenoliths are not simple residues and were affected by infiltration of fluid/melts within the mantle. The wide temperature range obtained for spinel peridotites (700 to >1150 °C) is compatible with a long period of pre-entrainment cooling supporting Fitton’s [Tectonophysics 94 (1983) 473] hypothesis that proposes oceanic lithosphere uprising in the Cameroon Volcanic Line prior to the initiation of the current thermal regime, related to S~ao Tome magmatism. The association of upper mantle (peridotite) xenoliths with igneous cumulates (pyroxenites) suggests that the spinel peridotite suite originated in the uppermost mantle above the S~ ao Tome magma storage zone(s), probably in a region of high strain rate, near the boundary between the mantle and the overlying oceanic crust. Ó 2002 Elsevier Science Ltd. All rights reserved. Keywords: Xenoliths; Ultramafic; S~ao Tome; Peridotite; Pyroxenite; Mantle

1. Introduction S~ ao Tome Island is located on the Gulf of Guinea, South Atlantic, at the oceanic sector of the 1600 km long Cameroon Volcanic Line (CVL; Fig. 1). This has been considered a privileged zone for the comparative study of oceanic and continental intraplate alkaline volcanism and consequently of sub-oceanic and sub-continental mantle (Fitton and Dunlop, 1985; Fitton, 1987; Halliday et al., 1988). The tectono-magmatic origin of this volcanic line has been widely discussed but is still subject of controversy (Fitton, 1980, 1983; Moreau et al., 1987; Halliday et al., 1988, 1990; Deruelle et al., 1991; Deruelle, 1994; Lee et al., 1994; Meyers et al., 1998). *

Corresponding author. E-mail addresses: [email protected] (R. Caldeira), [email protected] (J.M. Munha).

Xenoliths entrained in alkali basalts are an important source of information about lithospheric composition and thermal evolution in mantle regions associated with alkaline volcanism (Menzies, 1987). However, xenoliths studies on the CVL are scarce and confined to the continental sector mainly (Lee et al., 1996; Dautria et al., 1983); on the oceanic sector there are only brief references to peridotitic xenoliths from Bioko (Deruelle et al., 1991), Pagalu (Cornen and Maury, 1980) and S~ ao Tome (Fitton, 1980). S~ao Tome is predominantly composed of basaltic lavas less than 1 Ma old (Munha et al., in press); well preserved recent (<0.4 Ma; Munha et al., in press) pyroclastic cones (sometimes associated with scoriaceous lavas) are frequent in the northeast as well as in the south extreme of the island (Fig. 1). An older volcanic period (3–8 Ma) is exposed in the south–southeast region, with phonolitic plugs associated with tephritic-basaltic lava flows (phonolitic-basaltic complexes; Fig. 1).

0899-5362/02/$ - see front matter Ó 2002 Elsevier Science Ltd. All rights reserved. PII: S 0 8 9 9 - 5 3 6 2 ( 0 2 ) 0 0 0 2 2 - 2

232

R. Caldeira, J.M. Munha / Journal of African Earth Sciences 34 (2002) 231–246

Fig. 1. Geological sketch map of S~ao Tome island (after Munha et al., in press) with nodules locations (stars). Insert map is a generalised map of the Cameroon Line showing the main volcanic centres.

Except for one occurrence of peridotite xenoliths in older basalts of the Mizambu complex (
7 Ma), ultramafic nodules (as well as gabbroic and syenitic inclusions already described by Caldeira et al., 1999) have been found in recent (<1 Ma) alkali basaltic lavas from the northeast region of the S~ ao Tome island (Fig. 1). In this study we provide mineralogical and petrological data on S~ ao Tome’s ultramafic inclusions as a contri-

bution to a better knowledge of the oceanic lithosphere beneath the CVL. The nodules discussed in this paper are mainly peridotitic with subordinate pyroxenites. The host basalts are essentially porphyritic alkali olivine basalts to basanites (see Table 1), with olivine (xFo  81) and augitic to diopsidic clinopyroxene (Wo4746 En2940 Fs2412 ) phenocrysts in a matrix generally containing plagioclase

R. Caldeira, J.M. Munha / Journal of African Earth Sciences 34 (2002) 231–246

233

Table 1 Whole-rock compositions of xenolith bearing lavas from S. Tome island Ref

ST95/9

ST95/45

ST95/54

ST95/60

ST95/73

ST95/89

SiO2 TiO2 Al2 O3 Fe2 Ot3 MnO MgO CaO Na2 O K2 O P2 O5 LOI Total

43.88 3.18 12.81 12.62 0.18 10.27 10.68 3.83 1.37 1.17 0.03 100.02

43.63 3.19 12.80 12.85 0.18 10.06 10.76 3.83 1.36 1.17 0.18 100.01

43.02 3.16 12.81 12.46 0.17 9.83 10.87 3.63 1.39 1.17 1.48 100.00

46.11 2.78 12.58 13.02 0.17 10.78 9.84 3.06 0.89 0.72 0.04 99.99

44.29 3.01 12.74 12.12 0.17 11.35 9.74 4.22 0.50 1.13 0.75 100.02

44.48 2.95 13.39 13.26 0.18 10.14 9.65 3.02 1.28 0.77 0.82 99.94

Analyses carried out by XRF at the Crystallography and Mineralogy Center of IICT by M.O. Figueiredo and T. Silva.

(An4460 ), clinopyroxene (Wo48 En39 Fs13 ), Ti-magnetite (usp ¼ 0:77–0.43), some olivine (Fo74 ) and interstitial analcite.

2. Petrography Ultramafic nodules are small (0.5–2 cm in size) except for those collected at Mizamb u that can reach up to 5.5 cm. They vary from angular to rounded shapes. According to their modal composition the inclusions were classified as peridotites (lherzolites, harzburgites and dunites) or pyroxenites (orthopyroxenites and clinopyroxenites) following the IUGS recommendations (Streckeisen, 1973, 1976; Le Bas et al., 1986). Hand specimens of pyroxenites are dark-brown whereas those of peridotites tend to be yellowish-green to dark green in colour. 2.1. Peridotites S~ ao Tome peridotites are allotriomorphic granular to porphyroclastic (Harte, 1977; Pike and Shwarzman, 1977) and show deformational features. Most of them have a thin reaction rim with the host basalt resulting in replacement of orthopyroxene by intergrowth of olivine and clinopyroxene and substitution of Al, Cr-rich brown spinel by iron-oxides. The lherzolites and harzburgites have abundant olivine (45–90 vol%) and variable orthopyroxene (10–45 vol%). Clinopyroxene, always present in lherzolites (5– 25 vol%), is scarce or absent in the harzburgites (<5 vol%). Dunites are almost exclusively composed of olivine, although pyroxenes and spinels can be present in very small quantities (<3–5 vol%). Primary olivine occurs as large porphyroclasts (3–7 mm in size) commonly showing strain features such as undulose extinction or kink band structures. It also occurs as small recrystallised grains (0.05–1 mm) together with clinopyroxene and spinel, surrounding

primary crystals. Olivine neoblasts are mostly polygonal and show typical triple junctions. Orthopyroxene forms anhedral dark-brown crystals (2–5 mm in size) often with exsolution lamellae (occasionally bended) or blebs of clinopyroxene that disappear near crystal borders. Small rounded spinel inclusions and lamellar exsolutions are also common. Newly formed orthopyroxene exhibits round to polygonal forms and is free of exsolutions or inclusions. Primary clinopyroxene is colourless to green and usually smaller and less abundant than the coexisting olivine and orthopyroxene (average grain size is
2 mm). Clinopyroxene may exhibit undulose extinction and, rarely, contains exsolution of orthopyroxene. Occurring interstitialy between grains of olivine and orthopyroxene clinopyroxene forms small anhedral crystals (<2 mm), rich in spinel inclusions, probably resulting from migration from orthopyroxene. Brown to reddish brown chromiferous spinel occurs as intergranular amoeboid to vermicular grains, as very small grains in primary inclusions in silicates, as trails of small euhedral secondary inclusions in porphyroclasts or as lamellar exsolutions in pyroxenes. Sometimes it forms symplectitic intergrowth with clinopyroxene and/ or olivine constituting characteristic clusters, similar to those described by Mercier and Nicolas (1975). Rounded or tubular fluid inclusions in silicate minerals are ubiquitous features of S~ao Tome peridotites. Although hydrous minerals are absent, some peridotite samples (L225, D274, D448, D276) contain interstitial carbonates; (de-vitrified) glass can also be seen in these areas. 2.2. Pyroxenites Clinopyroxenites are mainly composed of clinopyroxene (65–97 vol%) with variable amounts of olivine (1–35 vol%) and minor spinel (<5 vol%). These nodules are rounded in shape and 1–3 cm in diameter. Most samples show granular igneous cumulus type textures,

234

R. Caldeira, J.M. Munha / Journal of African Earth Sciences 34 (2002) 231–246

but some pyroxenites also display moderately to strongly deformed textures. Undeformed nodules (C236, C241) have sub-euhedral to anhedral (
3 mm) zoned clinopyroxene crystals, containing fluid inclusions and exsolution lamellae of spinel. Olivine is rare (1–2 vol%) and oxide minerals are Ti-magnetites, forming euhedral inclusions or intergranular anhedral crystals. Deformed clinopyroxenite (C646) has a texture between allotriomorphic granular and protogranular (Mercier and Nicolas, 1975), where primary crystals of deformed clinopyroxene are accompanied by intergranular green Al-spinel. Olivine (
35 vol%)–clinopyroxenite sample (C279) has a typical porphyroclastic texture where relics of strained clinopyroxene porphyroclasts (showing irregular borders and internal recrystallisation and exsolution) are set in a mylonite-like matrix of small elongated grains of clinopyroxene and olivine. The orthopyroxenites are small in size (<2 cm) and have allotriomorphic granular textures. Orthopyroxene crystals (0.5–2 mm) are anhedral, display only slight strain effects (weak deformation lamellae and undulose extinction) and often contain fluid and spinel inclusions. Minor, intergranular, olivine, clinopyroxene and spinel are also be present.

3. Mineral chemistry Mineral compositions were studied from 23 ultramafic xenoliths and host basalts. The average mineral analyses for each sample are listed in Tables 2–5. The analyses were carried out on an automated JEOL JCXA 733Ó electron micropobe at the Center of Geology (University of Lisbon). Operating conditions were 15 kV acceleration voltage (18 kV for oxides), 25 nA probe current and an electronic beam of 5 lm of diameter. Standards were natural homogeneous minerals (Siwollastonite and olivine; Ca-wollastonite; Ti and Nakaersutite; Ti-ilmenite; Al-disthene; Fe and Mg-olivine; Mn-rhodonite; K-adularia) and synthetic minerals of Cr, V and Ni. The analytical precision evaluated from repeated analyses of standards is better than 2%. 3.1. Olivine The majority of primary olivine in peridotite has Fo contents ranging from 90 to 92 (Table 2; Fig. 2), contrasting with olivine in host basalts (
Fo80 ). Recrystallised grains show a wider range of Fo (84–92), with the lower values (similar to the outermost rims of some zoned primary crystals) resulting from reaction with the host alkali magma. NiO (for average values see Table 2) is within the range indicated by Sato (1977) for olivines of mantle origin. Ca in olivine olivine crystals from peridotites is low (CaO < 0:08 wt.%; Table 2) suggesting a high pressure equilibration environment (Simkin and

Smith, 1970; Kohler and Brey, 1990). Some grains exhibit a slight increase in CaO from core to rim which can be attributed either to partial re-equilibration at low pressures (Stormer, 1973) or to diffusive processes induced by heating during transport by the host magma (Takahashi, 1980). No evidence was found of systematic variation of olivine composition with peridotite textures as suggested by Xu et al. (1998) nor with modal abundance as referred by Frey and Green (1974) for residues of progressive partial melting. The olivines of the pyroxenites are less forsteritic (Fo < 88) revealing considerably differences between samples (Table 2). NiO contents in olivine from orthopyroxenite (
0.38 wt.%) is similar to those in peridotites and higher than in clinopyroxenites (NiOolivines ¼ 0:08– 0.21 wt.%). The lower CaO contents of orthopyroxenite olivines (Table 2) compared to those of clinopyroxenite olivines (similar to host basalts: CaOoliv  0:24 wt.%) probably reflect a difference in crystallisation pressures between pyroxenites. 3.2. Orthopyroxene Orthopyroxenes in peridotite are enstatites (Morimoto et al., 1988) with a compositional range of Wo03 En8891 Fs810 . In orthopyroxenite they have an average composition of Wo0:7 En87 Fs12 (see Table 3). The Mg# of orthopyroxenes correlates with the Fo of coexistent olivine and indicates that Mg=ðMg þ FeÞopx > Mg=ðMg þ FeÞol (e.g. Qi et al., 1995; Vaselli et al., 1995). Peridotite orthopyroxenes show a wide range of Al2 O3 which is a negative correlation with Cr# of coexisting spinel (Fig. 3) as usual for residual peridotites (Frey and Prinz, 1978). Orthopyroxene neoblasts have lower CaO, Cr2 O3 and Al2 O3 contents than primary crystals (e.g. L275; L232 in Table 3). The latter are often zoned with decreasing of CaO, Cr2 O3 and Al2 O3 contents from core to rim (sample H234 is shown as an example in Table 3). 3.3. Clinopyroxenes Clinopyroxene in peridotites are mostly diopsides (Wo4149 En5442 Fs29 ), with Mg# in the range of 87–97 (Table 4); their Cr2 O3 contents (0.50–1.75 wt.%) make them chromiferous diopsides (Cr a.f.u: > 0:01: Morimoto et al., 1988). Primary crystals generally have higher Mg# values than those of coexisting olivine and orthopyroxene, suggesting that Mg#cpx > Mg#opx P Fooliv , as it is typical for residual peridotites (Galer and O’Nions, 1989; Heinrich and Besh, 1992). Except for some recrystallised grains (Table 4: H456; D274) that reacted with the host liquid, TiO2 and Al2 O3 concentrations in clinopyroxenes from peridotites are low, being less than 0.06 and 4 wt.%, respectively (Fig. 4), with AlVI < 0:1 and AlIV < 0:1 (Fig. 5; Table 4). These compositional

Table 2 Micropobe analyses of olivine (averages of several analyses) Peridotites

Pyroxenites

Lherzolites L225

Harzburgites

L275

L232

P

R

L455

L235

L239

L272

H234

H226

Dunites H243

H456

H763

P

R

P

R

Orth.

Clinopyroxenites

D448

D276

D274

D553

D459

O242

C279

C464

C236

40.60 0.01 0.00 0.45 8.73 0.13 49.62 0.01 99.55

40.45 0.01 0.02 0.37 9.12 0.17 49.11 0.07 99.31

39.80 40.77 0.00 0.01 0.01 0.01 0.36 0.38 13.22 8.61 0.15 0.13 46.02 50.09 0.01 0.03 99.58 100.02

40.62 0.00 0.00 0.39 8.55 0.13 49.47 0.02 99.19

40.59 0.00 0.01 0.41 8.51 0.12 50.01 0.02 99.66

40.51 0.01 0.01 0.36 9.35 0.14 49.57 0.05 100.00

40.59 0.00 0.00 0.39 8.97 0.12 49.45 0.07 99.59

40.61 0.01 0.01 0.37 8.46 0.13 49.57 0.02 99.17

40.45 0.01 0.00 0.38 8.24 0.12 49.93 0.03 99.15

40.54 0.01 0.01 0.37 8.42 0.12 49.62 0.04 99.12

40.60 0.01 0.00 0.41 8.51 0.11 49.25 0.04 98.93

41.43 0.11 0.93 0.35 10.42 0.13 45.75 0.19 99.31

40.56 0.01 0.00 0.38 9.67 0.21 49.56 0.03 100.41

39.79 0.03 0.00 0.29 15.38 0.24 44.78 0.13 100.63

40.92 0.00 0.00 0.41 8.36 0.13 50.19 0.02 100.03

40.80 0.00 0.02 0.40 8.66 0.14 49.88 0.03 99.94

40.65 0.00 0.00 0.37 9.46 0.17 49.15 0.04 99.85

40.41 0.00 0.00 0.40 8.41 0.16 49.86 0.01 99.26

40.78 0.01 0.01 0.40 8.51 0.10 50.12 0.03 99.95

40.00 0.02 0.00 0.38 11.48 0.23 46.79 0.08 98.97

39.78 0.01 0.03 0.17 13.53 0.23 46.06 0.25 100.06

37.73 0.04 0.05 0.08 20.51 0.50 38.83 0.40 98.14

38.89 0.02 0.03 0.21 16.97 0.21 42.53 0.22 99.07

Fo%

91.02

90.57

86.12

91.16

91.28

90.43

90.76

91.26

91.53

91.31

91.16

88.74

90.14

83.84

91.46

91.12

90.25

91.36

91.30

87.90

85.83

77.14

81.71

91.21

P––porphyroclasts; R––recrystallised or when there are significant compositional differences in the same sample.

Table 3 Micropobe analyses of orthopyroxene (averages of several analyses) Peridotites

Pyrox.

Lherzolites L225

Harzburgites

L275

L232

L455

P

R

P

R

L235

L239

L272

Dunites

H234 Average

Pcore

Prim

Orthp.

H226

H243

H456

H763

D276

D274

D553

D459

O242

SiO2 TiO2 Al2 O3 Cr2 O3 Fe2 O3 FeO MnO MgO CaO Na2 O K2 O Total

55.77 0.00 2.71 0.86 0.30 5.84 0.18 33.01 1.28 0.02 0.01 99.98

55.54 0.01 2.75 0.85 0.33 5.79 0.15 32.85 1.24 0.01 0.00 99.53

57.81 0.02 0.92 0.08 1.14 5.39 0.17 35.47 0.21 0.02 0.01 101.23

56.09 0.02 1.87 0.57 1.34 4.52 0.17 34.21 1.05 0.01 0.00 99.85

57.38 0.00 0.33 0.04 0.25 5.72 0.15 34.96 0.24 0.02 0.00 99.09

55.64 0.01 2.07 0.60 0.43 5.42 0.18 33.14 1.04 0.01 0.01 98.54

55.77 0.01 2.18 0.61 1.05 5.00 0.17 33.83 0.94 0.01 0.00 99.56

55.62 0.02 2.64 0.78 0.18 6.08 0.15 32.82 1.14 0.06 0.00 99.49

54.86 0.01 3.18 0.70 0.88 5.11 0.13 33.14 0.74 0.04 0.00 98.78

55.71 0.02 2.43 0.72 0.23 5.55 0.15 33.47 0.84 0.02 0.00 99.15

55.73 0.02 2.56 0.75 0.13 5.64 0.15 33.32 0.96 0.02 0.00 99.29

55.64 0.03 1.92 0.64 0.59 5.16 0.15 34.05 0.36 0.02 0.00 98.56

55.39 0.04 2.93 0.92 0.60 5.15 0.13 33.35 1.10 0.02 0.01 99.64

53.45 0.04 4.31 1.02 2.03 4.08 0.15 32.62 1.02 0.03 0.00 98.75

57.37 0.00 1.68 0.23 0.00 5.65 0.17 34.51 0.43 0.02 0.00 100.06

58.09 0.01 0.18 0.39 0.12 5.10 0.16 34.41 1.84 0.09 0.02 100.41

56.57 0.01 0.99 0.16 1.16 4.98 0.15 34.86 0.26 0.01 0.00 99.15

57.73 0.03 0.60 0.04 0.48 5.83 0.15 35.15 0.18 0.04 0.00 100.23

57.71 0.01 0.35 0.14 0.32 5.51 0.12 35.44 0.13 0.00 0.01 99.73

56.71 0.03 2.56 0.33 0.50 5.18 0.13 34.72 0.40 0.03 0.00 100.58

57.43 0.00 0.30 0.09 0.00 7.90 0.29 33.23 0.36 0.04 0.01 99.65

Wo En Fs

2.45 88.13 9.42

2.39 88.21 9.40

0.38 90.23 9.40

1.99 89.37 8.64

0.46 90.68 8.86

2.01 88.97 9.02

1.78 89.17 9.05

2.20 88.17 9.63

1.45 89.44 9.11

1.62 89.52 8.86

1.85 89.26 8.89

0.68 90.59 8.73

2.12 89.15 8.73

2.00 88.75 9.25

0.81 90.63 8.56

3.41 88.82 7.77

0.49 90.52 8.99

0.33 90.41 9.25

0.24 91.22 8.54

0.74 90.80 8.45

0.68 87.26 12.06

Mg#

90.97

91.00

92.15

93.09

91.60

91.59

92.36

90.59

92.05

91.49

91.33

92.17

91.27

90.77

91.59

92.32

92.59

91.48

91.98

92.27

88.24

235

P––porphyroclasts; R––recrystallised or when there are significant compositional differences in the same sample.

R. Caldeira, J.M. Munha / Journal of African Earth Sciences 34 (2002) 231–246

SiO2 TiO2 Al2 O3 NiO FeO MnO MgO CaO Total

236

Table 4 Micropobe analyses of clinopyroxene (averages of several analyses) Peridotites

Pyroxenites

Lherzolites L225

Harzburgites

L275

L232

L455

L235

L239

L272

H234

H226

Dunites H243

H456

H763

Exsol.

R

D448

Clinopyroxenites D274

D553

D459

C279

C464 Pcore

Prim

C236

C241

52.36 0.03 3.19 1.28 0.07 2.32 0.09 16.92 22.85 0.10 0.00 99.21

54.35 0.03 1.19 0.96 0.00 2.38 0.12 17.72 21.29 0.53 0.01 98.57

53.00 0.07 2.60 1.49 0.67 1.73 0.10 16.96 22.21 0.62 0.01 99.47

52.24 0.04 2.55 0.99 0.77 1.56 0.07 17.31 23.30 0.03 0.01 98.86

53.07 0.03 2.62 1.05 0.67 1.63 0.11 18.69 22.01 0.04 0.01 99.94

53.18 0.03 3.20 0.66 0.88 2.23 0.07 16.57 22.11 0.74 0.01 99.68

52.90 0.04 2.41 0.51 0.41 1.78 0.09 16.53 23.51 0.40 0.00 98.56

53.80 0.04 2.03 0.79 0.00 2.53 0.06 17.74 22.11 0.28 0.00 99.38

54.08 0.00 1.45 0.90 0.00 2.35 0.07 17.91 22.10 0.35 0.00 99.23

53.20 0.06 2.18 0.31 0.26 1.96 0.10 16.71 23.81 0.28 0.01 98.88

52.58 0.95 1.74 0.82 0.79 2.93 0.11 16.50 21.34 0.83 0.02 98.61

50.32 1.18 4.02 0.99 0.97 3.99 0.05 14.81 21.81 0.61 0.00 98.76

55.09 0.00 0.87 1.45 1.13 2.03 0.04 16.31 21.69 1.50 0.00 100.12

52.48 0.49 2.36 1.19 0.00 4.37 0.15 17.00 20.01 0.00 0.01 98.08

47.25 2.08 6.65 0.81 2.40 3.07 0.11 13.22 21.69 0.82 0.01 98.12

51.90 0.02 3.99 1.35 2.02 0.79 0.07 16.04 22.81 0.73 0.00 99.73

54.32 0.02 1.95 0.18 0.72 1.57 0.06 17.38 23.89 0.38 0.00 100.48

50.90 0.40 4.88 1.15 0.62 3.44 0.09 14.46 22.57 0.58 0.03 99.12

52.68 0.10 1.82 0.20 0.29 4.00 0.17 15.00 23.52 0.37 0.01 98.16

49.16 0.51 7.34 0.49 0.95 2.91 0.05 15.07 21.22 0.48 0.00 98.18

49.08 1.42 5.46 0.25 2.48 3.72 0.11 14.50 20.88 0.76 0.00 98.65

46.32 2.61 7.31 0.04 1.75 5.74 0.17 11.22 22.49 0.65 0.01 98.30

Wo En Fs

47.29 48.71 4.00

44.31 51.57 4.12

46.52 49.48 3.99

47.37 48.96 3.67

44.16 52.16 3.68

46.42 48.48 5.09

48.71 47.66 3.63

45.20 50.62 4.17

45.04 51.04 3.92

48.73 47.60 3.67

45.18 48.61 6.22

47.14 44.55 8.31

46.35 48.49 5.16

42.40 50.11 7.49

49.01 41.57 9.42

48.31 47.28 4.42

47.92 48.50 3.57

49.20 43.83 6.97

49.14 43.62 7.25

46.98 46.42 6.60

45.59 44.09 10.33

51.17 35.52 13.31

Mg#

92.87

93.06

94.70

95.20

95.30

93.00

94.30

92.65

93.22

93.82

90.93

86.86

93.48

87.38

88.48

97.33

95.17

88.21

86.99

90.24

87.51

77.69

Table 5 Micropobe analyses of spinel (averages of several analyses) Peridotites

Pyroxenites

Lherzolites

Harzburgites

Dunites D274

Orth.

Clinopyronenites

L225

L275

L232

L455

L235

L239

L272

H234

H226

H243

H456

H763

D448

D276

D459

D553

O242

C279

Al2 O3 TiO2 Cr2 O3 V2 O3 Fe2 O3 FeO MnO MgO Total

27.60 0.03 34.96 0.23 7.22 15.50 0.23 13.27 99.04

26.20 0.16 35.41 0.22 7.60 15.35 0.25 13.14 98.34

16.23 0.06 50.11 0.37 3.93 15.30 0.19 12.16 98.35

10.92 0.03 49.99 0.27 10.66 14.34 0.20 12.21 98.61

27.19 0.02 38.52 0.17 3.79 16.29 0.22 12.72 98.92

29.44 0.03 37.81 0.23 3.03 13.82 0.20 14.65 99.21

31.08 0.20 34.52 0.23 4.64 12.35 0.16 15.85 99.05

44.39 0.02 21.68 0.12 4.45 10.81 0.13 18.33 99.94

28.47 0.04 37.16 0.16 4.94 14.81 0.23 14.00 99.81

27.86 0.06 39.18 0.18 1.72 16.04 0.16 12.86 98.07

44.85 0.04 21.58 0.07 4.45 10.62 0.13 18.61 100.35

30.46 0.06 32.60 0.10 7.09 12.17 0.16 15.64 98.27

13.82 3.10 37.79 0.20 11.90 22.42 0.36 8.89 98.47

2.73 0.02 60.78 0.15 5.70 24.49 0.46 4.70 99.03

27.60 0.02 37.82 0.20 5.43 13.21 0.22 14.85 99.35

13.39 0.30 44.98 0.25 11.60 19.50 0.26 9.42 99.69

13.95 3.65 32.56 0.19 16.07 25.18 0.28 7.65 99.52

11.57 1.30 48.70 0.13 6.74 22.01 0.28 8.03 98.77

11.65 0.25 51.20 0.11 6.31 20.90 0.32 8.15 98.89

42.91 0.07 22.64 0.12 5.55 9.34 0.13 19.15 99.92

15.36 0.08 45.50 0.24 8.91 14.99 0.30 11.95 97.33

7.43 0.67 46.27 0.07 15.00 20.69 0.37 7.88 98.39

44.07 1.50 9.28 0.12 12.70 18.26 0.23 14.26 100.42

Mg# Cr# % Usp

60.42 42.14 –

60.15 43.50 –

58.63 64.21 –

60.27 65.42 –

58.19 46.60 –

65.38 44.71 –

69.58 40.49 –

75.13 23.55 –

62.76 44.09 –

58.84 47.58 –

75.73 69.62 23.32 38.46 – –

41.42 54.21 –

25.49 86.49 –

66.71 44.97 –

46.27 59.31 –

35.14 47.43 –

39.41 67.31 –

40.99 68.66 –

78.51 24.64 –

58.69 59.18 –

40.44 64.57 –

C464

C236

C241

2.40 21.37 1.43 0.37 24.73 43.13 0.60 4.70 98.72

55.93 6.486 3.51 3.59 0.44 1.117 13.91 19.25 3.52 0.077 0.29 0.14 0.08 0.132 0.24 0.35 9.05 61.048 38.34 29.29 14.05 23.021 36.40 41.62 0.18 0.903 0.64 0.53 17.81 5.726 4.45 4.71 101.06 98.510 97.78 99.48

58.18 16.25 10.64 5.02 – 0.59

69.31 30.718 17.91 16.79 3.68 0.114 0.69 0.42 – – 0.33 0.52

R. Caldeira, J.M. Munha / Journal of African Earth Sciences 34 (2002) 231–246

SiO2 TiO2 Al2 O3 Cr2 O3 Fe2 O3 FeO MnO MgO CaO Na2 O K2 O Total

R. Caldeira, J.M. Munha / Journal of African Earth Sciences 34 (2002) 231–246

237

clinopyroxenes plot in the same alkaline field as their host basalt clinopyroxenes. 3.4. Spinel

Fig. 2. Histogram of average Fo contents in olivine of the ultramafic xenoliths from S~ ao Tome.

features are in agreement with the values found in abyssal peridotites (Arai and Fujii, 1979; Niida, 1997) and in clinopyroxenes of residual mantle origin as defined by Wass (1979) and Jagoutz et al. (1979). Contrasting with peridotite clinopyroxenes, those from clinopyroxenites have higher contents of aluminia (Al2 O3 > 4 wt.%; Table 4). All clinopyroxenes from the ultramafic nodules have AlVI =AlIV > 0:25 reflecting conditions of high pressure crystallisation (Aoki and Shiba, 1973; Wass, 1979). In Fig. 6 it is evident that peridotite and strongly deformed clinopyroxenite clinopyroxenes plot in the non-alkaline field while undeformed clinopyroxenite

Fig. 3. Diagram showing the Al2 O3 opx (wt.%) vs. Cr# spinel relationships in the peridotitic xenoliths from S~ao Tome. Data points are average mineral analyses for each sample. Black symbols correspond to samples with recrystallised opx exclusively.

Spinels show large and characteristic compositional variations in peridotites, orthopyroxenites and clinopyroxenites (Table 5). Peridotite and orthopyroxenite spinels are (Mg,Fe)Al2 O4 –(Mg,Fe)Cr2 O4 solid solutions corresponding to chromiferous spinels (Al > Cr, Mg > Fe2þ ); magnesiochromites (Cr > Al, Mg > Fe2þ ) and chromites (Cr > Al, Fe2þ > Mg: Deer et al., 1963; Sigurdsson, 1977). They fall within the range found by Haggerty (1976, 1979) which is typical of ultramafic xenoliths in basalts and kimberlites and of alpine type peridotites. The large and reciprocal variation of Al2 O3 and Cr2 O3 (Table 5) is well displayed by the Cr# ¼ 100Cr=ðCr þ Al þ Fe3þ ) ratio which is an inverse correlation with Mg# (Fig. 7). In this diagram most spinels from peridotites are in the range defined for abyssal peridotites (Dick and Fisher, 1984; Dick and Bullen,

Fig. 4. Mg# versus Cr2 O3 , TiO2 and Al2 O3 (wt.%) plot of average clinopyroxene compositions from S~ao Tome ultramafic xenoliths and host rocks; Squares––harzburgites, diamonds––lherzolites, triangles––dunites,  cross––clinopyroxenites, þ cross––host basalts; c––center; r––rim. Black symbols as in Fig. 3.

238

R. Caldeira, J.M. Munha / Journal of African Earth Sciences 34 (2002) 231–246

Fig. 5. AlVI vs. AlIV (afu) plot of average clinopyroxene from ultramafic xenoliths and host basalt. Field for refractory mantle clinopyroxenes (RMC) after Jagoutz et al. (1979). Symbols as in Fig. 4. Black filled symbols correspond to samples with recrystallised cpx exclusively.

1984; Niida, 1997); however some samples show high Cr# (e.g. H763) falling in the alpine type peridotites field (Dick and Bullen, 1984). Others have Mg# contents too low for the corresponding Cr# suggesting late-stage reequilibration (Dick and Bullen, 1984). Sometimes, even in the same sample, there is a significant Cr# variation (e.g. samples L275 and H456) indicating that some spinels are out of equilibrium with the coexisting minerals (Henry and Medaris, 1980). Orthopyroxenite spinels are richer in TiO2 (
0.67 wt.%) than the majority of the peridotite spinels whose titanium contents are generally very low (TiO2 < 0:1 wt.%) and well below the upper limit defined for abyssal peridotite spinels (0.8 wt.%: e.g. Arai and Fujii, 1979; Hamlyn and Bonatti, 1980; Dick, 1989). Small grains

Fig. 6. SiO2 vs. Al2 O3 (wt.%) plot of average clinopyroxene compositions from S~ ao Tome ultramafic xenoliths and host rocks. Fields after Le Bas (1962). Symbols as in Figs. 4 and 5. Peridotites clinopyroxenes cleary have a different nature from the host basalts and most clinopyroxenites. Sample C464 (c––center; r––rim) is probably a xenocrystal that was partially recrystallised by action of the host magma (see text for explanation).

near the contact with the host basalt (or in recrystallisation areas that reacted with the host melt), reveal enrichment in TiO2 (up to 3.7 wt.%: Table 5) that are accompanied by higher Fe2 O3 values. Deformed clinopyroxenites have aluminous green to dark green spinels that are basically solid solutions between spinel s.s. (MgAl2 O4 ) and hercynite (FeAl2 O4 ), with Mg2þ :Fe2þ between 1 and 3 characteristic of pleonast (Deer et al., 1963). These spinel compositions (Table 5) fall within the range referred to spinels of type II inclusions by Frey and Prinz (1978). In undeformed clinopyroxenites (C241, C236) spinel belongs to the magnetite-ulvospinel series (Deer et al., 1963). Most spinels are high TiO2 Ti-magnetites (Table 5), similar to those of the host basalts (TiO2 ¼ 10:54–24.69 wt.%), but a few are Al, Mg-rich magnetites (Al2 O3  6:5 wt.%, MgO  6 wt.%), reflecting the characteristic incorporation of spinel s.s in magnetites crystallising from primitive alkaline basaltic magmas.

4. Geothermobarometry 4.1. Equilibration temperatures In this study we have applied several methods (see Table 6) to estimate equilibration temperatures for the ultramafic xenoliths found in S~ao Tome island. The selected nodule samples show relatively homogeneous mineral compositions and a good positive correlation between Mg# of coexisting silicate phases suggesting

Fig. 7. Plot of Cr# vs. Mg# of average spinel values from peridotite xenoliths from S~ao Tome.

R. Caldeira, J.M. Munha / Journal of African Earth Sciences 34 (2002) 231–246

239

Table 6 Geotermometric estimates for the ultramafic xenoliths of S. Tome (temperatures in °C) Sample

Lc.

Methods Cpx–opx

Opx-ca

Ol–Esp

W&B

WEL

B&M

KRET

B&K1

B&K2

P R E P R E P R P B P B R P B E P R E

1024 – 1079 1011 1194 1207 979 1112 1066 1049 1252 – – 1084 933 1093 992 – 938

912 – 982 891 1106 1133 851 1013 965 932 1163 – – 988 812 999 861 – 809

829 – 929 795 1128 1147 801 1002 906 875 1163 – – 956 700 993 764 – 703

862 – 959 918 1186 1186 854 1085 967 945 1159 – – 1029 1118 1054 889 – 847

867 – 957 799 1164 1160 770 1040 914 876 1150 – – 990 705 1005 792 – 692

1097 – 1164 1121 721 1084 1191 749 1110 961 1145 811 – 1125 1037 1024 884 831 1107

733 769 – 855 879 – 940 913 734 721 745 746 934 900 886 – 823 887 –

702 751 – 898 844 – 923 877 695 676 717 718 1008 963 942 – 936 1041 –

713 753 – 839 890 – 998 962 694 680 698 704 928 889 873 – 770 836 –

Harzburgites H234 P R E H226 P R H243 P B E H456 P H763 P R

– 1211 924 880 1223 1007 – 980 1062 – 980

– 1141 800 743 1146 870 – 829 952 – 854

– 1143 699 613 1145 749 – 707 897 – 789

– 1178 991 826 1182 891 – 840 1035 – 996

– 1170 903 570 1170 785 – 715 958 – 818

940 920 1090 1071 1039 892 842 1253 823 1208 1194

755 780 – 712 715 809 858 – 872 692 814

729 751 – 666 671 942 979 – 921 519 684

725 755 – 661 665 755 807 – 859 679 828

P R P R P R P R P

– – – – – 958 – 982 –

– – – – – 835 – 858 –

– – – – – 771 – 836 –

– – – – – 809 – 894 –

– – – – – 851 – 807 –

– – 761 761 – 702 – 659 820

838 842 704 756 749 703 867 825 931

842 844 599 641 634 574 828 761 1093

824 829 718 776 747 706 891 843 895

Ortopyroxenite O242 P

–

–

–

–

–

794

853

766

910

Clinopyroxenites C279 P R C464 P C236 P

– – – –

– – – –

– – – –

– – – –

– – – –

– – – –

603 689 1494 887

680 852 2142 1014

630 743 1216 1089

Lherzolites L225

L275

L232 L455 L235

L239

L272

Dunites D448 D276 D274 D553 D459

FAB

ROED

O&W

W&B––Wood and Banno (1973); WEL––Wells (1977); B&M––Bertrand and Mercier (1985); KRET––Kretz (1982); B&K1 and B&K2––Brey and Kohler (1990); FAB––Fabries (1979); ROED––Roeder et al. (1979); O&W––O’Neill and Wall (1987). P––primary crystals, B––crystal rims, R––recrystallised crystals, E––exsolution of clpx in opx. See text for explanation.

that equilibrium was attained––a pre-condition for a successful calculation of their formation temperatures.

However, the obtained results must be analysed bearing in mind that, for some lithotypes, there are characteristic

240

R. Caldeira, J.M. Munha / Journal of African Earth Sciences 34 (2002) 231–246

disequilibrium features which can be attributed to longterm temperature variations. Thus, whenever possible, separate temperatures were calculated from core (P) and rim (B) areas of porphyroclasts, recrystallised neoblasts (R), and from exsolved phases (E). The results (Table 6) show that the geothermometers based on the same mineral pairs correlate well, whereas there is no correlation between thermometers based on different mineral assemblages. This absence of correlation is to be expected, reflecting different exchange closure temperatures for the analysed minerals (Herzberg and Chapman, 1976; Fabries, 1979). Wells (1977) geothermometer has the important advantage of showing no correlation with calculated Fe3þ from microprobe data (McGuire, 1988) and it has been considered as one of the most reliable geothermometric methods (e.g. Fujii and Scarfe, 1982; Borshon and Clague, 1988; Heinrich and Besh, 1992); therefore, it will be given preference when concerning the two-pyroxene geothermometers. Olivine–spinel (ol–sp) thermometers gave similar results for each peridotite sample but different results for pyroxenites; we will give preference to the O’Neill and Wall method (1987) because it is the only one that ponders Ti contents in spinels. Two-pyroxene (Wells, 1977) temperature estimates for peridotite porphyroclast cores indicate variable conditions of equilibrium (temperatures from core ðPÞ ¼ 750–1150 °C; Table 6), This is within the range found for spinel lherzolites on the continental sector of the CVL (Lee et al., 1996). However, the common occurrence of clinopyroxene exsolution lamellae in orthopyroxenes indicates that these estimates should not be interpreted as indicating original formation temperatures which may have been much higher (Borshon and Clague, 1988). Significantly, two-pyroxene temperatures calculated for neoblast (R: Table 6) are systematically higher (by up to 400 °C) than those derived from primary peridotite assemblages. Olivine–spinel peridotite temperatures range from 650 to 950 °C (Table 6), being systematically lower (by about 100–200 °C) than twopyroxene temperatures; higher two-pyroxene temperatures reflect lower cationic diffusion rate in pyroxenes relative to spinel–olivine FeMg1 exchange kinetics, which remained effective towards the later cooling stages (Fabries, 1979). Thus, the above features indicate that peridotites underwent complex thermal evolution, corresponding to a long-term period of pre-entrainment cooling (during pyroxene unmixing and deformation) prior to initiation of neoblast recrystallisation, which certainly involved significant re-heating related to magmatic activity in S~ ao Tome island. Temperature estimates for pyroxenites were restricted to olivine–spinel geothermometers. Calculated O’Neill and Wall (1987) olivine–spinel temperature values are 630–750 °C for strongly deformed clinopyroxenite, 910 °C for orthopyroxenite, and 1100–1200 °C for moder-

ately deformed to undeformed clinopyroxenites. All these values are significantly lower than the segregation temperatures (1340–1380 °C) of the host magmas, as inferred by applying Albarede’s model (1992) to S~ ao Tome basalt compositions (Table 1). 4.2. Pressure estimates Although there are some studies, mainly based on the Ca distribution of olivine and clinopyroxene (e.g. Simkin and Smith, 1970; Adams and Bishop, 1986; Kohler and Brey, 1990), there are no reliable geobarometers for spinel peridotites. However, we can indirectly presume the equilibrium pressure from pertinent experimental stability phase relationships. The absence of garnet and plagioclase constrains the equilibrium pressure of the studied xenoliths to lie within the range of the spinel stability field, between about 10 and 20 kb (Obata, 1976; Herzberg, 1978; O’Neill, 1981; Gasparik, 1984), which is equivalent to depths of 30–60 km (assuming lithostatic pressure). These same values can also be assumed for pyroxenite nodules. The equilibration pressure between a magma and a given mineral assemblage, containing olivine and orthopyroxene, can be evaluated if the compositions of these phases are known at the equilibrium temperature (Carmichael et al., 1977; Ghiorso et al., 1983; Ghiorso, 1987; Ghiorso and Carmichael, 1987). This method was applied to the orthopyroxenite (the only pyroxenite containing olivine þ orthopyroxene) assuming that it represents a cumulate from S~ao Tome magmas; thus, representative chemical compositions of S~ao Tome basalts (within the range of 44–46 SiO2 wt.% range; Table 1) and the orthopyroxenite mineral chemistry (Tables 2 and 3), together with appropriate thermochemical data and solution models (Wood and Banno, 1973; Nichols, 1976; Carmichael et al., 1977; O’Neill and Wall, 1987), enable us to estimate the equilibration pressure between the orthopyroxenite xenolith and the S~ao Tome magmas to have been 12–14 kb, at a crystallisation temperature of about 1000–1100 °C. This pressure is equivalent to a depth of about 39–43 km. The same method, and Albarede’s empirical equations (1992), were also applied to evaluate segregation pressures of S~ao Tome basaltic magmas from their peridotitic sources. At 1300–1400 °C (see above) the results obtained by the two methods are similar suggesting segregation depths between 62 and 75 km (22–24 kb), next to the transition between the stability fields of spinel and garnet. 4.3. Oxygen fugacity To determine oxygen fugacities for the ultramafic nodules of S~ao Tome, we employed the methods of O’Neill and Wall (1987), Mattioli and Wood (1988),

R. Caldeira, J.M. Munha / Journal of African Earth Sciences 34 (2002) 231–246

Nell and Wood (1991), and Ballhaus et al. (1991) all ðolivineÞ based on the reaction 6Fe2 SiO4 þ O2 ¼ ðorthopyroxeneÞ ðspinelÞ 3Fe2 Si2 O6 þ Fe3 O4 . For these calculations we adopted the geothermometric data obtained by the Wells method (1977) and the known compositions of the mineral association of olivine–orthopyroxene-spinel (Tables 2,3 and 5). Except for the O’Neill and Wall geobarometer(1987), all methods indicate a majority of f O2 values above that of the FMQ buffer (Fig. 8). The D logðf OFMQ Þ vs. Cr#ðspinelÞ diagrams (Fig. 9) show that 2 oxygen fugacity values calculated from the peridotite xenoliths of S~ ao Tome do not correlate with the level of depletion though there’s a slight positive correlation for low to moderate degrees of partial melting (Cr# < 60); the residues with Cr#ðspinelÞ > 60 show decreasing oxygen fugacities. Overall, the relatively high f O2 values are in accordance with an oxidised mantle. This is frequently attributed to metasomatic processes during influx of carbonatitic fluid (Bryndzia et al., 1989; Mattioli et al., 1989; Wood et al., 1990). The data are similar to those of enriched MORB and OIB basalts related to mantle plumes (Ballhaus et al., 1990). The large range of calculated f O2 is also typical for mantle heterogeneity.

241

5. Discussion An accidental origin of the peridotitic xenoliths from S~ao Tome seems to be obvious since they exhibit very different petrographic and chemical features when compared with their host alkali basalts. Fo (>89) and NiO (>0.36 wt.%) contents in olivines, Mg# (91–95) values of orthopyroxenes (which correlate with the olivine Fo) and very low Ti contents in (primary) clinopyroxene crystals (TiO2 < 0:06 wt.%) and spinels (TiO2 < 0:1 wt.%) are all in the range of type I xenoliths (Frey and Prinz, 1978), the Cr-diopside series (Wilshire and Shervais, 1975), or abyssal peridotites (Arai and Fujii, 1979; Dick and Fisher, 1984) indicating that S~ ao Tome peridotite xenoliths are refractory mantle residues. The compositional range of pyroxene and spinel (Al2 O3ðopxÞ , Cr#ðspinelÞ ; Fig. 3) found in S~ao Tome peridotite xenoliths reflects variable melting degrees of peridotite (Dick and Fisher, 1984; Michael and Bonatti, 1985) implying an heterogeneous mantle, compositionally zoned from a primitive to depleted mantle. Accordingly, most fertile peridotites (L272, H243, D459) have orthopyroxene richer in Al and spinel poorer in Cr

Fig. 8. Oxygen fugacities relative to FMQ buffer for the ultramafic xenoliths of S~ao Tome. Spinel activity calibrations as indicated in the histograms (see text).

242

R. Caldeira, J.M. Munha / Journal of African Earth Sciences 34 (2002) 231–246

Fig. 9. Plot of D logðf O2 Þ vs. Cr# of spinel for the different methods used in the calculations. Symbols as in Fig. 3; symbols correspond to recrystallised paragenesis.

than the depleted samples (H763, D274, D276, D553). However, according to this criteria the more or less depleted character does not always conform to olivine Fo values nor to peridotite types (except for dunites) as it would be expected for residues of progressive partial melting. Moreover, the lack of correlation between silicate Mg# values and modal compositions (e.g. Frey and Green, 1974; Frey and Prinz, 1978; Dick et al., 1984) suggests that spinel peridotites are not simple residues. This apparent decoupling between mineral chemistry and modal composition and much of the complexity illustrated by the mineral chemistry of S~ ao Tome peridotitic nodules could be best explained by mixing of a depleted residual component with other enriched components (Frey and Green, 1974; Michael, 1988), inherent to mantle metasomatic processes (Frey et al., 1985; Menzies et al., 1985). In S~ ao Tome’s peridotitic xenoliths, the absence of metasomatic hydrous minerals precludes the possibility of modal metasomatism (Kempton, 1987; Menzies et al., 1987). However, the presence of fluid inclusions and carbonates might be an evidence for metasomatic fluid activity. Additional evidence of possible interaction with carbonatitic fluids may be inferred from low Al2 O3 (<4 wt.%) and high CaO (21–24 wt.%) concentrations in clinopyroxenes (see Rudnick et al., 1993) and from the relatively high f O2 obtained for S~ ao Tome peridotites (see Ballhaus et al., 1990). This is supported by the higher f O2 values calculated in dunites (which also have higher carbonate contents) and is in accordance with the possible preferred circulation of carbonate melts in olivine rich rocks (Hunter and McKenzie, 1989).

––cumulates. Black filled

The wide span of calculated equilibrium temperatures (650 to >1150 °C) obtained for spinel peridotites is compatible with long-term cooling, prior to their caption by the ascending magma and high temperature recrystallisation of neoblast assemblages. This preentrainment cooling may also be inferred from reaction textures in coexisting MORB-type gabbroic enclaves (Caldeira et al., 1999) and is well expressed in Ca, Cr and Al concentrations decreasing outwards in orthopyroxene. This is a typical feature of sub-solidus reequilibrium (Sinton, 1977; Preß, 1986). The common occurrence of reaction rims in S~ao Tome nodules also suggests substantial cooling before incorporation in their host magmas. Cooling and intense deformation, as observed in S~ao Tome peridotite xenoliths, are often related to diapiric uplift of hotter mantle into colder uppermost mantle, and subsequent thermal-mechanical re-equilibration (Witt and Seck, 1987; Nicolas et al., 1987; Heinrich and Besh, 1992). Furthermore, the occurrence of characteristic orthopyroxene-clinopyroxenespinel clusters in S~ao Tome peridotitic xenoliths suggests derivation by decompressional melting of garnet þ olivine (Nicolas et al., 1987) during mantle uprising processes. Gravimetric and seismic reflection data obtained by Meyers and Rosendahl (1991) and Meyers et al. (1998) also suggest asthenospheric upwelling in the oceanic sector of the CVL causing crustal and lithospheric mantle uprise. Dautria et al. (1983) concluded that peridotitic xenoliths on the continental sector of the CVL (Adamawa) have granularities and rheological profiles implying the existence of a diapiric margin subjected to high tensions. In S~ao Tome we

R. Caldeira, J.M. Munha / Journal of African Earth Sciences 34 (2002) 231–246

propose an influx of upwelling asthenosphere causing thermal erosion and thinning that would favour lithosphere uprising in the CVL and consequent cooling, in a process similar to that proposed by Xu et al. (1998) for NE China. Mineral chemistry of pyroxenite nodules indicates that they cannot have the same origin as the peridotite xenoliths. Pyroxenite minerals are richer in Fe and Ti (olivines Fo < 88; clinopyroxenes TiO2ðcpxÞ > 0:4 wt.%; spinels TiO2ðspinelÞ ¼ 0:5–2.0 wt.%), their composition being in the range defined by Frey and Prinz (1978) for type II inclusions and by Wilshire and Shervais (1975) for inclusions of the augite-Al series which have been interpreted as mantle crystallisation products of basaltic magmas (Wilshire and Shervais, 1975; Frey and Prinz, 1978). However, the clinopyroxenitic nodules (C279 and C464) record features that are difficult to reconcile with their classification as simple crystal cumulates from the alkaline host magmas. Clinopyroxene cores in sample C464 are chemically similar to peridotite clinopyroxenes; on the other hand rims in this sample are like the host basalt pyroxenes (Figs. 4 and 6), where olivine is relatively poor in Fo (¼77) and aluminous spinel shows relatively high Cr and Ti (TiO2 ¼ 0:31–0.58 wt.%; Cr2 O3 ¼ 2–5 wt.%). These chemical characteristics, the textural aspects, and the reduced dimension of the nodule suggest that it corresponds to a (mega)xenocrystal that was fractured and partially recrystallised by action of the host magma. The strongly deformed clinopyroxenite (C279) contains olivine similar to the olivines of the host basalt, especially in CaO, NiO and Fo. Spinels have too much Ti (TiO2 ¼ 0:40–1.59 wt.%) to be considered as mantle material (be it residual or fertile) or as MORB type (Dick and Bullen, 1984). However, the clinopyroxene chemistry (Table 4) shows that the magmatic affinity is not clear. These features, the deformed texture revealing intense deformation and the existence of a reaction rim with the host, suggest that this clinopyroxenite is a fragment of a strained crystalline cumulate, resulting from an earlier magmatic episode unrelated to the host basalts. In light of the proposed diapiric model, its strongly deformed nature indicates that it could be located in a diapir flank (Heinrich and Besh, 1992; Xu et al., 1998).

6. Conclusions On the basis of petrographic, mineral chemical, and geothermometric studies of the ultramafic nodules collected from alkaline lavas of S~ ao Tome island we conclude that peridotite type I xenoliths represent a random sampling of the depleted oceanic upper mantle beneath the island, through which the basaltic magma has passed on its way to the surface. Mineralogical and textural

243

features of S~ao Tome peridotitic xenoliths record a process of mantle uprising (as suggested by Fitton, 1983 for the CVL) before the initiation of the current thermal regime related to magmatic activity in the island. Mantle uprise may be related to installation of the St. Helen plume, which has been connected with the opening of the Atlantic ocean and the CVL magmatic activity (e.g. Halliday et al., 1988; Guiraud et al., 1992; Lee et al., 1994). Metasomatic enrichments by infiltration of small melt fractions or fluids could have been produced by decompression melting caused either by earlier asthenospheric upwelling, or be related with recent magmatism associated with the S~ao Tome island. The association of upper mantle peridotite xenoliths with igneous cumulates (pyroxenites) suggests that the peridotite suite originated in the uppermost mantle, above the magma storage zone(s) (
40 km deep beneath the S~ao Tome island) that were fed from deeper sources in the mantle (62–75 km).

Acknowledgements Field work was financed by the ICP (Instituto de Cooperacß~ao Portuguesa). We thank the government of the Republic of S. Tome and Prince for logistic support. We also wish to acknowledge the Tropical Sciences Research Institute (IICT) and the Center of Geology from the Faculty of Sciences of the Lisbon University for material and laboratory support. For critical review and fruitful comments the authors want to thank F. Melcher and an anonymous reviewer.

References Adams, G.E., Bishop, F.C., 1986. The olivine–clinopyroxene geobarometer: experimental results in the CaO–FeO–MgO–SiO2 system. Contributions to Mineralogy and Petrology 94, 230–237. Albarede, F., 1992. How deep do common basaltic magmas form and differentiate? Journal of Geophysical Research 9 (7), 10997– 11009. Aoki, K., Shiba, I., 1973. Pyroxenes from lherzolite inclusions of Itinomegata, Japan. Lithos 6, 41–51. Arai, S., Fujii, T., 1979. Petrology of ultramafic rocks from site 395. In: Initial Reports of the Deep Sea Drilling Project, 45. US Government Printing office, Washington, pp. 587–594. Ballhaus, C., Berry, R.F., Green, D.H., 1990. Oxygen fugacity controls in the Earth’s upper mantle. Nature 348, 437–440. Ballhaus, C., Berry, R.F., Green, D.H., 1991. High pressure experimental calibration of the olivine–orthopyroxene–spinel oxygen geobarometer: implcations for the oxidation state of the upper mantle. Contributions to Mineralogy and Petrology 107, 27–40. Bertrand, P., Mercier, J.C., 1985. The mutual solubility of coexisting ortho- and clinopyroxene: toward an absolute geothermometer for natural system. Earth and Planetary Science Letters 76, 109– 122. Borshon, W.A., Clague, D.A., 1988. Origin of ultramafic xenoliths containing exsolved pyroxenes from Hualalai volcano, Hawaii. Contributions to Mineralogy and Petrology 100, 139–155.

244

R. Caldeira, J.M. Munha / Journal of African Earth Sciences 34 (2002) 231–246

Brey, G.P., Kohler, T., 1990. Geothermobarometry in four phase lherzolites II. New thermobarometers, and practical assessment of existing thermobarometers. Journal of Petrology 31/6, 1353–1378. Bryndzia, L.T., Wood, B.J., Dick, H.J.B., 1989. The oxidation state of the Earth’s sub-oceanic mantle from oxygen thermobarometry of abyssal spinel peridotites. Nature 341, 526–527. Caldeira, R., Munha, J.M., Afonso, R.S., 1999. Petrologia dos n odulos gabr oicos em lavas basalticas da ilha de S. Tome, Arquipelago de S. Tome e Principe’’ (English abstract). Actas do II Congresso Iberico de Geoquimica/XI Semana de Geoquimica, Lisboa, Portugal, pp. 69–72. Carmichael, I.S.E., Nicholls, J., Spera, F.J., Wood, B.J., Nelson, S.A., 1977. High-temperature properties of silicate liquids: applications to the equilibration and ascent of silicate magmas. Philosophical Translations Royal Society of London A 286, 373–431. Cornen, G., Maury, R.C., 1980. Petrology of the volcanic island of Annobon, Gulf of Guinea. Marine Geology 36, 253–267. Dautria, J.M., Girod, M., Rahaman, O., 1983. The upper mantle beneath eastern Nigeria: inferences from ultramafic xenoliths in Jos and Biu volcanics. Journal of African Earth Sciences 1 (3/4), 331– 338. Deer, W.A., Howie, R.A., Zussman, J., 1963. Rock Forming Minerals, 5: Non-Silicates. Longmans, Green and Co, London, p. 371. Deruelle, B., 1994. Le magmatisme de la Ligne du Cameroun: Le bilan positif des deux dernieres decennies de cooperation franco-camerounaise. Pangea, Conference Specialisee, pp. 30–36. Deruelle, B., Moreau, C., Nkoumbou, C., Kambou, R., Lissom, J., Njonfang, E., Ghogomu, R.T., Nono, A., 1991. The Cameroon Line: a review. In: Kampuru, A.B., Lubala, R.T. (Eds.), Magmatism in Extensional Structural Settings. The Phanerozoic African plate. Springer, Berlin, pp. 274–327. Dick, H.J.B., 1989. Abyssal peridotites, very slow spreading ridges and ocean ridge magmatism. In: Saunders, A.D., Norry, M.J. (Eds.), Magmatism in the Ocean Basins. Geological Society Special Publication 42, 71–105. Dick, H.J.B., Bullen, T., 1984. Chromian spinel as a petrogenetic indicator in abyssal and alpine-type peridotites and spatialy associated lavas. Contributions to Mineralogy and Petrology 86, 54–76. Dick, H.J.B., Fisher, R.L., 1984. Mineralogical studies of the residues of mantle melting: abyssal and alpine-type peridotites. In: Kornprobst, J. (Ed.), Kimberlites II: The Mantle and Crust-mantle Relationships. Elsevier, Amsterdam, pp. 295–308. Dick, H.J.B., Fisher, R.L., Bryan, W.B., 1984. Mineralogic variability of the uppermost mantle along mid-ocean ridges. Earth and Planetary Science Letters 69, 88–106. Fabries, J., 1979. Spinel–Olivine geothermometry in peridotites from ultramafic complexes. Contributions to Mineralogy and Petrology 69, 329–336. Fitton, J.G., 1980. The Benue through and Cameroun Line––a migrating rift system in West Africa. Earth and Planetary Science Letters 51, 132–138. Fitton, J.G., 1983. Active versus continental rifting: evidence from the West African rift system. Tectonophysics 94, 473–481. Fitton, J.G., 1987. The Cameroon Line, West Africa: a comparison between oceanic and continental alkaline volcanism. In: Fitton, J.G., Upton, B.G.J. (Eds.), Alkaline Igneous Rocks. Geological Society Special Publication 30, 273–291. Fitton, J.G., Dunlop, H.M., 1985. The Cameroon Line, West Africa and its bearing on the origin of oceanic and continental alkali basalt. Earth and Planetary Science Letters 72, 23–38. Frey, F.A., Green, D.H., 1974. The mineralogy, geochemistry and origin of lherzolite inclusions in Victorian basanites. Geochimica Cosmochimica Acta 38, 1023–1059. Frey, F.A., Prinz, M., 1978. Ultramafic inclusions from San Carlos, Arizona: petrological and geochemical data bearing on their petrogenesis. Earth and Planetary Science Letters 38, 129–176.

Frey, F.A., Suen, C.J., Stockman, H.W., 1985. The Ronda high temperature peridotite: geochemistry and petrogenesis. Geochimica Cosmochimica Acta 49, 2469–2491. Fujii, T., Scarfe, C.M., 1982. Petrology of ultramafic nodules from West Kettle River, near Kelowna, southern British Columbia. Contributions to Mineralogy and Petrology 80, 297–306. Galer, S.J.G., O’Nions, R.K., 1989. Chemical and isotopic studies of ultramafic inclusions from the San Carlos volcanic field, Arizona: a bearing on their petrogenesis. Journal of Petrology 30, 1033–1064. Gasparik, T., 1984. Two-pyroxene thermobarometry with new experimental data in the system CaO–MgO–Al2 O3 –SiO2 . Contributions to Mineralogy and Petrology 87, 87–97. Ghiorso, M.S., 1987. Modeling magmatic systems: thermodynamic relations. In: Carmichael, I.S.E., Eugster, H.P. (Eds.), Thermodynamic Modeling of Geological Materials: Minerals, Fluids and Melts. Reviews in Mineralogy 17, 443–465. Ghiorso, M.S., Carmichael, I.S.E., 1987. Modeling magmatic systems: petrologic applications. In: Carmichael, I.S.E., Eugster, H.P. (Eds.), Thermodynamic Modeling of Geological Materials: Minerals, Fluids and Melts. Reviews in Mineralogy 17, 467–499. Ghiorso, M.S., Carmichael, I.S.E., Rivers, M.L., Sack, R.O., 1983. The Gibbs free energy of mixing of natural silicate liquids: an expanded regular solution approximation for the calculation of magmatic intensive variables. Contributions to Mineralogy and Petrology 88, 260–268. Guiraud, R., Binks, R.M., Fairhead, J.D., Wilson, M., 1992. Chronology and geodynamic setting of Cretaceous-Cenozoic rifting in West and Central Africa. Tectonophysics 213, 227–234. Haggerty, S.E., 1976. Opaque mineral oxides in terrestrial igneous rocks. In: Rumble III, R. (Ed.), Oxide Minerals. Reviews in Mineralogy 3, 101–300. Haggerty, S.E., 1979. Spinels in high pressure regimes. In: Boyd, Fr., Meyer, H.O.A. (Eds.), The Mantle Sample: Inclusions in Kimberlites and other Volcanics. Proceedings of the 2nd Kimberlite conference, pp. 183–196. Halliday, A.N., Dickin, A.P., Fallick, A.E.E, Fitton, J.G., 1988. Mantle dynamics: A Nd, Sr, Pb and O isotopic study of the Cameroon Line volcanic chain. Journal of Petrology 29, 181–211. Halliday, A.N., Davidson, J.P., Holden, P., DeWolf, C.D., Lee, D.-C., Fitton, J.G., 1990. Trace-element fractionation in plumes and the origin of HIMU mantle beneath the Cameroon Line. Nature 347, 523–528. Hamlyn, P.R., Bonatti, E., 1980. Petrology of mantle-derived ultramafics from the Owen fracture zone, northwest indian ocean: implications for the nature of the oceanic upper mantle. Earth and Planetary Science Letters 48, 65–79. Harte, B., 1977. Rock nomenclature with particular relation to deformation and recrystallisation textures in olivine-bearing xenoliths. Journal of Geology 85, 279–288. Heinrich, W., Besh, T., 1992. Thermal history of the upper mantle beneath a young back-arc extensional zone: ultramafic xenoliths from San Luis Potosı, Central Mexico. Contributions to Mineralogy and Petrology 111, 126–142. Henry, D.J., Medaris, L.G., 1980. Application of pyroxene and olivine–spinel geothermometers to spinel peridotites in southwestern Oregon. American Journal Science 280-A, 211–231. Herzberg, C.T., 1978. Pyroxene geothermometry and geobarometry: experimental and thermodynamic evaluation of some subsolidus phase relations involving pyroxenes in the system CaO–MgO– Al2 O3 –SiO2 . Geochimica Cosmochimica Acta 42, 945–957. Herzberg, C.T., Chapman, N.A., 1976. Clinopyroxene geothermometry of spinel-lherzolites. American Mineralogist 61, 626–637. Hunter, R.H., McKenzie, D., 1989. The equilibrium geometry of carbonate melts in rocks of mantle composition. Earth and Planetary Science Letters 92, 347–356. Jagoutz, E., Palme, H., Baddenhausen, H., Blum, K., Cendales, M., Dreibus, G., Spettel, B., Lorenz, V., Wanke, H., 1979. The

R. Caldeira, J.M. Munha / Journal of African Earth Sciences 34 (2002) 231–246 abundances of major, minor, and trace elements in the earth’s mantle as derived from primitive ultramafic nodules. Proceedings of the 10th Lunar Planetary Science Conference, pp. 2031–2050. Kempton, P.D., 1987. Mineralogic and geochemical evidence for different styles of metasomatism in spinel lherzolite xenoliths: enriched mantle source regions of basalts? In: Menzies, M.A., HawKesworth, C.J. (Eds.), Mantle Metasomatism, pp. 45–90. Kohler, T.P., Brey, G.P., 1990. Calcium exchange between olivine and clinopyroxene calibrated as a geothermobarometer for natural peridotites from 2 to 60 kb with applications. Geochimica Cosmochimica Acta 54, 2375–2388. Kretz, R., 1982. Transfer and exchange equilibria in a portion of the pyroxene quadrilateral as deduced from natural and experimental data. Geochimica Cosmochimica Acta 46, 411–421. Le Bas, M.J., 1962. Role of aluminium in igneous clinopyroxenes with relation to their parentage. American Journal of Science 260, 267– 288. Le Bas, M.J., Le Maitre, R.W., Streckeisen, A., Zanettin, B., 1986. A chemical classification of volcanic rocks based on the total alkalisilica diagram. Journal of Petrology 27, 745–750. Lee, D.-C., Halliday, A.N., Fitton, J.G, Poli, G., 1994. Isotopic variations with distance and time in the volcanic islands of the Cameroon Line: evidence for a mantle plume origin. Earth and Planetary Science Letters 123, 119–138. Lee, D.-C., Halliday, A.N., Davies, G.R., Essene, E.J., Fitton, J.G., Temdjim, R., 1996. Melt enrichment of shallow depleted mantle: a detailed petrological, trace element and isotopic study of mantle derived xenoliths and megacrysts from the Cameroon Line. Journal of Petrology 17 (2), 415–441. Mattioli, G.S., Wood, B.J., 1988. Magnetite activities across the MgAl2 O4 –Fe2 O3 spinel join, with application to thermobarometric estimates of upper mantle oxygen fugacity. Contributions to Mineralogy and Petrology 98, 148–162. Mattioli, G.S., Baker, M.B., Rutter, M.J., Stolper, E., 1989. Upper mantle oxygen fugacity and its relationship to metasomatism. Journal of Geology 97, 521–536. McGuire, A.V., 1988. Petrology of mantle xenoliths from Harrat Al Kishb: the mantle beneath Western Saudi Arabia. Journal of Petrology 29, 73–92. Menzies, M., 1987. Alkaline rocks and their inclusions: a window on the Earth’s interior. In: Fitton, J.G., Upton, B.G.J. (Eds.), Alkaline Igneous Rocks. Geological Society Special Publication 30, 15–27. Menzies, M.A., Kempton, P.D., Dungan, M.A., 1985. Interaction of continental lithosphere and asthenospheric melts bellow the Geronimo volcanic field, Arizona, USA. Journal of Petrology 26, 663– 693. Menzies, M., Rodgers, N., Tindle, A., Hawkesworth, C., 1987. Metasomatic and enrichment processes in lithospheric peridotites, an effect of the astenosphere-litosphere interaction. In: Menzies, M., Hawkesworth, C.J. (Eds.), Mantle Metasomatism. Academic Press, London, pp. 313–361. Mercier, J.C.C., Nicolas, A., 1975. Textures and fabrics of uppermantle peridotites as illustrated by xenoliths from basalts. Journal of Petrology 16 (2), 454–487. Meyers, J.B., Rosendahl, B.R., 1991. Seismic reflection character of the Cameroon Volcanic Line: evidence for uplifted oceanic crust. Geology 19, 1072–1076. Meyers, J.B., Rosendahl, B.R., Harrison, C.G.A., Ding, Z.-D., 1998. Deep-imaging seismic and gravity results from the offshore Cameroon Volcanic Line, and speculation of African hotlines. Tectonophysics 284, 31–63. Michael, P.J., 1988. The concentration, behavior and storage of H2 O in the suboceanic upper mantle: implications for mantle metasomatism. Geochimica Cosmochimica Acta 52, 555–566. Michael, P.J., Bonatti, E., 1985. Peridotite composition from the North Atlantic: regional and tectonic variations and implications for partial melting. Earth and Planetary Science Letters 73, 91–104.

245

Moreau, C., Regnoult, J.-M., Deruelle, B., Robineau, B., 1987. A new tectonic model for the Cameroon Line, Central Africa. Tectonophysics 139, 317–334. Morimoto, N., Fabries, J., Ferguson, A.K., Ginzburg, I.V., Ross, M., Seifert, F.A., Zussman, J., Aoki, K., Gottardi, G., 1988. Nomenclature of pyroxenes. American Mineralogist 73, 1123–1133. Munha, J., Afonso, R., Caldeira, R., Mata, J., in press. Estudo geol ogico preliminar da regi~ao nordeste da ilha de S. Tome(folha n°2, Ana Chaves). Garcia de Orta. Serie de Geologia. Nell, J., Wood, B.J., 1991. High-temperature electrical measurements and thermodynamic properties of Fe3 O4 –FeCr2 O4 –MgCr2 O4 – FeAl2 O4 spinels. American Mineralogist 76, 405–426. Nichols, J., 1976. The activity of components in natural silicate melts. In: Fraser, D.G. (Ed.), Thermodynamics in Geology. D. Reidel, Dordrecht, pp. 327–348. Nicolas, A., Lucazeau, F., Bayer, R., 1987. Peridotite xenoliths in Massif Central basalts, France: textural and geophysical evidence for asthenospheric diapirism. In: Nixon, P.H. (Ed.), Mantle Xenoliths. John Wiley and Sons Ltd, New York, pp. 563– 574. Niida, K., 1997. Mineralogy of MRK peridotites: replacement through magma channeling examined from Hole 920D, MARK area. In: Karson, J.A., Cannat, M., Miller, D.J., Elthon, D. (Eds.), Proccedings. ODP, Scientific Results, 153, College Station, TX (Ocean Drilling Program), pp. 265–275. O’Neill, H.St.C., 1981. The transition between spinel lherzolite and garnet lherzolite, and its use as geobarometer. Contributions to Mineralogy and Petrology 77, 251–255. O’Neill, H.St.C., Wall, V.J., 1987. The olivine–orthopyroxene-spinel oxygen geobarometer, the nickel precipitation curve, and the oxygen fugacity of the Earth’s upper mantle. Journal of Petrology 28 (6), 1169–1191. Obata, M., 1976. The solubility of Al2 O3 in orthopyroxenes in spinel and plagioclase peridotites and spinel pyroxenite. American Mineralogist 61, 804–816. Pike, J.E.N., Shwarzman, E.C., 1977. Classification of textures in ultramafic xenoliths. Journal of Geology 85, 49–61. Preß, S., Witt, G., Seck, H.A., Eonov, D., Kovalenko, V.I., 1986. Spinel peridotite xenoliths from the Tariat Depression, Mongolia. I: Major element chemistry and mineralogy of a primitive mantle xenolithe suite. Geochimica Cosmochimica Acta 50, 2587–2589. Qi, Q., Taylor, L.A., Zhou, X., 1995. Petrology and geochemistry of mantle peridotite xenoliths from SE China. Journal of Petrology 36 (1), 55–79. Roeder, L.P., Campbell, I.H., Jamieson, H.E., 1979. A re-evaluation of the olivine–spinel geothermometer. Contributions to Mineralogy and Petrology 68, 325–334. Rudnick, R.L., McDonough, W.F., Chappel, B.W., 1993. Carbonatite metassomatism in the northern Tanzanian mantle: petrographic and geochemical characteristics. Earth and Planetary Science Letters 114, 463–475. Sato, H., 1977. Nickel content of basaltic magmas: identification of primary magmas and a measure of the degree of olivine fractionation. Lithos 10, 113–120. Sigurdsson, H., 1977. Spinels in Leg 37 basalts and peridotites: phase chemistry and zoning. In: Initial Reports of the Deep Sea Drilling Project, 37. US Government Printing Office, Washington, pp. 883– 891. Sinton, J.M., 1977. Equilibration history of the basal Alpine-type peridotite, Red Mountain, New Zealand. Journal of Petrology 18, 216–246. Simkin, T., Smith, J.V., 1970. Minor element distribution in olivine. Journal of Geology 78, 304–325. Stormer Jr., J.C., 1973. Calcium zoning in olivine and its relationship to silica activity and pressure. Geochimica Cosmochimica Acta 42, 1829–1844.

246

R. Caldeira, J.M. Munha / Journal of African Earth Sciences 34 (2002) 231–246

Streckeisen, A., 1973. Classification and nomenclature of plutonic rocks: recommendations. Neues Jahrbuch Fur Mineralogie-Monatshefte Jg., 149–164. Streckeisen, A., 1976. To each plutonic rock its proper name. Earth Science Review 12, 1–33. Takahashi, E., 1980. Thermal history of lherzolite xenoliths. I: Petrology of lherzolite xenoliths from the Ichinomegata crater, Oga Peninsula, Northeast Japan. Geochimica Cosmochimica Acta 44, 1643–1658. Vaselli, O., Downes, H., Thirwall, M., Dobosi, G., Coradossi, N., Seghedi, I., Szakacs, A., Vannucci, R., 1995. Ultramafic xenoliths in plio-pleistocene alkali basalts from the eastern Transylvanian basin: depleted mantle enriched by vein metassomatism. Journal of Petrology 36 (3), 23–53. Wass, S., 1979. Multiple origins of clinopyronenes in alkali basaltic rocks. Lithos 12, 115–132. Wells, P.R.A., 1977. Pyroxene thermometry in simple and complex systems. Contributions to Mineralogy and Petrology 62, 129– 139.

Wilshire, H.G., Shervais, J.W., 1975. Al-augite and Cr-diopside ultramafic xenoliths in basaltic rocks from Western United States. In: Ahrens, L.H., Press, F., Runcorn, S.K., Urey, H.C. (Eds.), Physics and Chemistry of the Earth, vol. 9. Pergamon, Oxford, pp. 257–272. Witt, G., Seck, H.A., 1987. Temperature history of sheared mantle xenoliths from the West Eifel, West Germany: evidence for mantle diapirism beneath the Rhenish massif. Journal of Petrology 28, 475–493. Wood, B.J., Banno, S., 1973. Garnet-orthopyroxene and orthopyroxene-clinopyroxene relationships in simple and complex systems. Contributions to Mineralogy and Petrology 42, 109–124. Wood, B.J., Bryndzia, L.T., Johnson, K.E., 1990. Mantle oxidation state and its relationship to tectonic environment and fluid speciation. Science 248, 337–345. Xu, Y., Menzies, M.A., Vroon, P., Mercier, J.-C., Lin, C., 1998. Texture-temperature-geochemistry relationships in the upper mantle as revealed from spinel peridotite xenoliths from Wangqing, NE China. Journal of Petrology 39, 469–493.