petrogenesis of ultramafic rocks and associated chromitites in the nan uttaradit ophiolite, northern...

LITHOS 0

Lithos 35 (1995) 153-182 ELSEVIER

Petrogenesis of ultramafic rocks and associated chromitites in the Nan Uttaradit ophiolite, Northern Thailand

B. Orberger”,“, J.P. Lorandb, J. Girardeau”, J.C.C. Merciera*d, S. Pitragool” “Laboratoire de P&-ologie Physique, Universite’ Paris-7, lnstitut de Physique du Globe de Paris, 2 Place Jussieu,

75251 Paris Cedex 05, France / CNRS, URA N” 1093, France

“Laboratoire de Mine’ralogie, Mus&an National d’Histoire Naturelle, UnitP associke au CNRS N” 736, 61 rue de Buffon. 75005 Paris, France “Laboratoire de Pe’trologie Structurale; Universite Nantes, UFR Science et Tkchniyues, 2 rue de la HoussiniPre,

44072 Nantes Cedex 03, France “Present address: Lab. d’ktudes Physiques et Chimiques appliquies ci la Terre, PBle Sciences and Technologie. UniversitP de L.a Rochelle,

Avenue Marillac, 17042 Lu Rochelle CJdex 01, France ‘Department of Geological Sciences, Faculty of Science, Chiang Mai University, Chiang Mai 50002, Thailand

Received 4 March, 1993; revised and accepted 21 July, 1994

Abstract

The ultramafic sequence and associated chromitites of the Nan-Uttaradit ophiolite in the northeastern part of Thailand have been studied in the field and by applying petrography and geochemistry to whole rock samples and minerals. The ultramafic rocks comprise irregulary shaped bodies of dunite, harzburgite, orthopyroxene-rich lherzolite and orthopyroxene-rich harzburgite, clinopyroxene-rich dunite and intrusive clinopyroxenite-websterite bodies. Three types of chromitite were distinguished. Type I chromitite lenses and type II layers which are hosted in orthopyroxenite in the northern part and in dunite in the central part of the ophiolite. Type III chromitite forms lenses or layers in clinopyroxenites in the central and southern parts of the belt. According to the modal and chemical composition the peridotites and orthopyroxenites are strongly refractory. They originated during different stages of interaction between percolating melts and peridotite. The chromitites of types I and 11, which are very rich in Cr (up to 68 wt.% Cr,OX), crystallized from a boninitic parental magma under highly reducing conditions in the northern part and moderate oxygen fugacities (FMQ) in the central part of the ophiolite. The chromitite of type III which are characterized by the highest Fe3+ /(Fe”+ + Cr + Al)-ratios, and hosted in intrusive clinopyroxenite-websterite-rocks, cumulated from a CaO- rich transitional boninitic melt underflS conditions around FMQ.

1. Introduction

Ophiolites are important to our understanding of magma generation and magma extraction from the

mantle. Both the pillow lava and the upper mantle section of ophiolites indicate their derivation from different magma generations (Beccaluva and Serri, 1988; Ohnenstetter et al., 1990 and references therein). The oldest component of ophiolite sequences is usually a MORB-type tholeiite where plagioclase has crystal-

0024-4937/95/$09.50 0 1995 Elsevier Science B.V. All rights reserved

SSDIOO24-4937(94)00041-7

lized first. It is overlain by various highly magnesian

andesites and basalts where pyroxene has crystallized before plagioclase. These rocks show the typical high-

field strength element (HFSE) depletion of island-arc

magmatism. Some rocks, which have crystallized

orthopyroxene or clinopyroxene and highly chromiferous Cr-spine1 first, show boninitic affinities. Due to the presence of the rocks formed from such magmas, most authors agree that ophiolites were likely formed in mar-

154 B. Orberger et al. /Lithos 35 (1995) 353-182

ginal basins above subduction-zones rather than at mid-

ocean ridges (e.g. Elthon, 199 1; Pearce, 199 1) .

Intrusive ultramafic bodies cross-cutting the mantle

sequence and the mantle/crust boundary have been

recognized in several ophiolite complexes (e.g. Troo-

dos: Benn and Laurent, 1987; Oman: Ernewein et al.

1988; Benn et al., 1988; Appalachian ophiolites, Can-

ada: Laurent and Hebert, 1989). They are characterized

by the crystallization of clinopyroxene and/or ortho-

pyroxene before olivine and plagioclase. It has been

suggested that these bodies may represent cumulates

segregated from the low-Ti magmas that produced the

upper pillow lavas (e.g. Thy and Moore, 1988; Ohnen-

stetter et al., 1990; Thy and Xenophontos, 1991).

Orthopyroxenite and clinopyroxenite bodies of this

type also occur within in the presently studied lower

ophiolite sequence of the Nan Uttaradit (NU) belt in

Thailand. They host different types of chromitite bodies

which are unusual rich in chromium. The present paper

presents a detailed mineralogical and geochemical

study of both, the chromitites and their associated ultra-

mafic rocks. Like many ophiolites related to continent-

a CHINA CHINESE

CR AT ON

continent collisions, the NU ophiolite is strongly

dismembered and hydrothermally altered. The princi-

pal aim of the present study was therefore to identify the protoliths of the ultramafic rocks by using whole

rock and mineral major and trace element data. Chrom-

spine1 compositions in conjunction with silicate para-

genesis were used to constrain the compositions and

redox conditions of the different generations of melt

that percolate the NU ophiolite.

2. Geology and tectonic setting

The Nan-Uttaradit (NU) ophiolite belt, in the northeast of Thailand, is a 10 km large and 150 km long NE-

SW trending suture zone separating two major

continental cratons (Fig. 1 a) : the Shan-Thai craton to the West and the Indosinian-Chinese craton to the East

(Bunopas and Vella, 1978; Thanasuthipitak, 1978;

Sengor, 1979; Ridd, 1980; Huang, 1984; Barr and Mac- donald, 1987; Cooper et al. 1989; Hutchison, 1989;

Panjasawatwong and Crawford, submitted). It extends

17O30’N

Fig. 1. Geotectonic situation of the Nan Uttaradit ophiolite. a. general geotectonic framework of southeast Asia. b. geologic map of the Nan

Uttaradit area, northeastern Thailand (modified after Thanasithapak et al., 1978).

B. Orberger et al. /Lithos 35 (1995) 153-182 155

from east of Mae Charim in the north to about 20 km south of the Sirikit reservoir, about 30 km east of Uttar-

adit and has probably its prolongation near Prachin Buri near the Cambodian border (Fig. la).

Mafic and ultramafic ophiolitic rocks appear as tec-

tonized slices within sediments. These sediments con-

sist of sandstones, slates, shales and conglomerates of

Permo-Triassic and Carboniferous (Fig. 1 b) age in the

west, Siluro-Devonian in the southeast and Jurassic in

the eastern and north-eastern parts (Baum et al., 1970).

Extensive faulting and thrusting during emplacement

have dismembered the Nan-Uttaradit ophiolite and

obscured the relationships with the host sediments.

The geotectonic setting in which the NU ophiolite

originated is still debated. Helmcke ( 1985) questioned

whether it represents a remnant of Paleotethys seafloor

or whether it was formed to the east of Gondwana or

within a marginal sea which was situated off Paleoeu-

rasia. A continent-continent collision model has been

proposed by many authors for Thailand. Bunopas and

Vella (1978), Barr and Macdonald (1987) and Barr

et al. ( 1990) suggested a westward subduction beneath the Shan-Thai craton whereas an eastward subduction

under the Indosinian craton was proposed by Beckin- sale et al. (1979). A third model considers a pair of

subduction zones, one dipping to the west and the other

dipping to the east (Bunopas and Vella, 1978; Thana- suthipitak, 1978; Cooper et al., 1989; Panjasawatwong

and Crawford, submitted). The latter authors suggest

that a northeast-dipping subduction of oceanic crust,

initiated beneath the Indochina craton, later changed to southwest-dipping subduction beneath the Shan-Thai

craton. There is also a considerable uncertainty about the timing of the continent-continent collision. Esti-

mated ages range from pre-Late Permian (Helmcke,

1985)) Middle to Late Permian (Barr and Macdonald, 1987), Permo-Triassic (Thanasuthipitak, 1978; Coo-

per et al., 1989; Barr et al., 1990)) Triassic (Beckinsale et al., 1979) and Late-Triassic (Bunopas and Vella,

1978; Sengiir, 1979; Hutchison, 1989). The contrasting opinions on the genesis of the Nan-Uttaradit suture are

due mainly to the poor exposure of the ophiolite rocks and the effects of tectonic events during and after the

collision, i.e. extensional collapse of the overthickened crust during the late Triassic and Late-Triassic to Ceno-

zoic transcurrent movements (Cooper et al., 1989). However, whatever the chosen model is, the formation of the Nan-Uttaradit ophiolite above a subduction zone

is confirmed by geochemical studies of the volcanic rocks. These comprise oceanic-island basalts, back-arc basalts and andesites and island-arc basalts and andesites of Carboniferous to Permo-Triassic age

(Panjasawatwong and Crawford, submitted).

2.1. Lithology of the NlJ ophiolite

Ultramafic rocks crop out all over the belt, whereas mafic rocks, in particular the gabbroic sequence, are

mostly exposed only in the central part (CIG-company,

unpubl. data). East-west profiles across the belt show

that the ophiolite consists of tectonic slices thrusted to

the east over sedimentary rocks of uncertain age. Around Pak Nai, the nearly NE-SW-trending thrust

contacts and fault zones are often characterized by silic-

ified serpentinite lenses, while massive magnesite

occurs in the northern (Rae Nan) and central parts (Pak

Nai) of the belt.

Fig. 2 displays a schematic cross-section through the ophiolite. Although the main units are often dismem-

bered, primary magmatic and sedimentary contacts are

Fig. 2. Schematic synthetic section of the Nan-Uttaradit ophiolite.

1.56 B. Orberpr et al. /Lithos 35 (1995) 153-182

preserved locally. The ophiolite can be subdivided into

( 1) an ultramafic unit, consisting of peridotites and

variably amphibolitized pyroxenites, both containing

chromitite orebodies and (2) a mafic unit, comprising

more or less amphibolitized layered and isotropic gab-

bros hosting small intrusive granitic bodies and doler-

ites. The dolerites are massive and do not form a well

defined sheeted-dike complex. The mafic sequence is

crosscut by different types of dikes: coarse-grained orthopyroxenites, plagioclasite (anorthite-rich with

large pyroxene crystals) in the gabbroic rocks and, fine-

grained epidotite and late basalts of various width (5

50 cm) in the dolerites. Pillowed lavas with their sed-

imentary cover, comprising bedded red, green and black cherts, rarely rest upon the mafic cumulates.

Ultramafic rocks appear in the central part of the belt

as three tectonic slabs, which are more or less con- nected. Our samples were collected in the area of Mae Charim and Rae Nan in the north, Pak Nai in the central

part of the belt, and Prachin Buri in the south (Fig. la). In the central part, the westernmost slab that borders

the Permo-Carboniferous (?) sediments mainly com-

prises serpentinized harzburgites and minor pyroxe-

nites. The central slab consists of highly serpentinized

harzburgites, orthopyroxenites and dunites. The largest

slab which crops out south of Pak Nai consists of more

or less amphibolitized pyroxenite layers and sheets (up

to 300 m thick) within serpentinized harzburgite and

orthopyroxenite. The pyroxenites form centimetre- to

metre-sized sills which are generally separated from the host peridotites by thrust contacts hiding their intru-

sive relationships. An even larger clinopyroxenite

body, up to 400 m thick, is clearly intrusive within the mantle/crust boundary (Fig. 2). It is in contact with

peridotites to its base and with layered gabbro at its top. The latter likely marks the bottom of the oceanic crust.

Both the clinopyroxenite and the gabbros are partly

amphibolitized and foliated because of regional metamorphism related to the continental collision.

The ultramafic rocks also contain three main types of chromitites (Figs. 2, 3) :

-type I chromitites are lenses and hosted in peridotites and orthopyroxenites (Fig. 3a). They were found essentially in two parts of the ophiolite belt. A large chromitite lens of about 20 m length and 10 m width is located near Mae Charim. It is separated from the sur- rounding peridotite by an EW-trending shear zone. Several chromitite lenses of centimetric to metric size

occur in the central part (Pak Nai). Mining work has

obscured the relationships between these chromitite

lenses and their hosts. Nevertheless, some small chromitites lenses show a thin serpentinized dunitic envelope.

-type I1 layered and banded chromitites are most

abundant. They are hosted in peridotites and orthopyroxenite all over of the ophiolite. The largest layers (30

cm thick) have been found at Rae Nan (Figs. 1 b, 3b).

Multiple thin chromitite bands, several centimeters in thickness crop out south of Pak Nai (Figs. 1 b, 3~). In

both cases, the chromitite orebodies grade into dissem-

inated chromite (50% Cr-spine], 50% silicate) in the

host peridotites. En-echelon chromitite veins crosscut

the hostrock (Fig. 3b). -type III chromitites are lenses of different sizes

within the lower part of the large intrusive clinopyrox-

enite body discovered in the Pak Nai area. They display sharp and locally secant contacts with the host pyrox-

enite (Fig. 3d). The largest bodies of metre-size are

folded and affected by the same metamorphic foliation as the host pyroxenite. The smallest discontinuous thin (cm-size) bands/layers are orientated parallel to the

pyroxenite foliation (Fig. 3d).

3. Analytical methods

Sixty rock samples have been investigated in pol-

ished thin sections by conventional microscopic tech- niques, using both transmitted and reflected light.

About fourty peridotites and sixteen pyroxenites have

been analysed for major and minor elements as well as for a number of trace elements by X-ray fluorescence

spectrometry. These samples were collected at dis- tances from massive magnesite or silicified serpentin- ites. A first set of major element analyses was collected

at the Institut fur Mineralogie und Lagerstattenlehre at the Technical University of Aachen, Germany, using

Phillips Typ PW 1400. A second set of analyses including some minor and trace elements (Cr, Ni, Co, SC, V, Zn, Sr, Ba, Ce, Nd, Zr, Nb) was performed at the School of Earth Science, University of Melbourne, Australia, using an ARL 8420 spectrometer. Major and minor elements as well as trace elements were analysed

using glass beads prepared according to the method described by Haukka and Thomas ( 1977). Represen- tative analyses of peridotites and pyroxenite subtypes


Fig. 3. Photographs of chromitites: a. type I chromitite pods with dunitic envelope (Pak Nai); b. type II layered chromitite in orthopyroxenite

(Rae Nan); c. type II chromitite bands in dunite (Pak Nai); d. type 111 chromitite lens in intrusive clinopyroxenite.

as well as detection limits of the analytical methods are

shown in Tables 1 and 2. Phase compositions were determined with an ARL-

electron microprobe at the Institut fiir Mineralogie und

Lagersttittenlehre, RWTH Aachen, Germany, using

acceleration voltage of 20 k, a current beam of 20 nA and a counting time of 20 s. Another set of data have been obtained with fully automated CAMEBAX-elec-

tron microprobes “Microbeam” at the BRGM (OrlCans, France) and CAMPARIS (Universitt Paris VI, France) at 15 kV and 40 nA with a counting time of 20 s. It was not possible to analyse all phase com-

positions in all the rock types because of the too strong hydrothermal alteration and metamorphism. Thus, a special attention has been paid to olivine, clinopyroxene and chromite in order to trace variations of mg ( 100 Mg/Mg + Fe in atoms) and Cr ( 100 Cr/Cr + Al in atoms) numbers. Representative analyses are reported in Table 3.

4. Petrography

4. I. Peridotites and orthopyroxenites

Modal compositions

Peridotites and orthopyroxenites collected from

shear zones (e.g. Pb 1, Pb 6) are strongly serpentinized

as demonstrated by the loss on ignition values (L.O.I.)

of whole-rock analyses (Table 1). Due to the high degree of serpentinization a metamorphic texture could

not be recognized. Spine1 is the less altered mineral. It

is slightly transformed into ferrite-chromite. Olivine is

replaced by hour-glass chrysotile-lizardite, sometimes coexisting with brucite in serpentinized dunites north

of Pak Nai. Orthopyroxene is bastitized. Fine-grained magnetite is abundant in the serpentine matrix or forms overgrowths on Cr-spinels. Even clinopyroxene are transformed into serpentine or Ca-amphibole in the Rae Nan and Prachin Buri samples. The degree of serpen-

1% B. Orberger et al. / Lithos 35 (I 995) 153-I 82

Table 1 Representative major and trace element compositions of peridotites and orthopyroxenites. Detection limits for minor elements (wt.%): TiO,: 0.02; CaO: 0.02; Na,O:

0.50); detection limits of trace elements (ppm): Cr, SC, V, Co, Ni: 3; Zn, Zr, Nb, 2; FelO, stands for total iron. Mg# = lW*Mg/(Mg+Fe) (using mol.%). Due

to the CaO loss in most of the pyroxene-rich lherzolites, original CaO contents were recalculated by assuming a constant Sc/Ca-ratio of 7 x 10m4. Original CaO/

Al,O, ratios (prior to CaO-loss) is reported as CaO/Al,O: Opxte: orthopyroxenite: Hzbg: harzburgite, Lherz: Iherzolite, LOI: loss on ignition

Sample no.

Opxte Opxte Opxte Lherz. Lherz. Lherz. Lherz. Hzbg Lhetz. Lherz. Lherz. PB6 MC 38 MC40 PB 1 PB 3 RNP RN64 BNT 133 RN 59 RN 63 MC51

skeletal opx skeletal opx skeletal opx skeletal opx skeletal opx skeletal opx skeletal opx skeletal opx

shearzone shearzone shearzone shearzone shearzone

SiO, (wt.%) 48.12

TiO, <0.02

Al?& 0.61

Fe& 8.63

MnO 0.11

MgD 29.86

CaO 0.05

Cr (ppm) 2711

SC 9

V 26

co 142

Zn 39

Ni 9585

LO1 (wt.%) 10.78

4-3.33

< 0.02

0.81

9.24

0.07

33.22

< 0.02

3312

44.08 41.02 40.08 40.67 40.47 41.03 40.2 39.63 40.86

<0.02 < 0.02 0.02 0.02 < 0.02 < 0.02 0.02 0.03 < 0.02 1.33 1.71 2.74 2.42 I .28 0.57 1.4 1.91 0.99

4.26 1.75 10.17 7.04 7.38 6.02 7.81 6.61 7.28

0.13 0.07 0.1 0.11 0.11 0.1 0.04 0.06 0.14

35.72 35.35 33.35 36.5 37.44 39.89 37.19 38.03 37.38

0.05 0.03 0.04 1.37 0.04 0.77 < 0.02 <0.02 I .53

12741 3155 3686 2311 2905 2958 2147 6079 2813

6 I4 18 II IO 8 9 5 12

15 65 80 43 33 24 43 33 48

123 102 111 94 107 98 103 104 97

57 67 76 35 41 36 48 31 32

3318 4973 6328 1992 2295 2360 228 1 3006 2150

12.27 12.74 12.38 11.07 12.59 10.03 12.44 12.25 12.41

54

142

41

5115

12.73

Total 99.43 100.32 99.52 99.54 99.93 99.72 99.89 98.99 99.67 99.65 101.13

Mg# 87.26 87.68 94.32 90.03 86.65 91.12 90.95 92.92 90.41 91.93 91.04

CaO/Al,O, I .4f 0.86* 0.4 I* I* 0.57 I* I .35 1.06* 0.27’ 1.53

Modal composition

Olivine 14

Onhopyroxene 80

Clinopyroxene 5

Spine1 I

34

61

4

22 45

75 44

3 IO

7 0.9

46 52 54 59 57.5 60 53

38 39 37.5 36 35.5 35 27

15 8 7.5 4 6 9 9

I 0.7 I I I 2 I

Sample no.

Hzbg Lherz. Hzbg Hzbg Hzbg Hzbg Hzbg Dunite Dunite Dunite Dunite

RN 65 BNT 138 MC55 MC53 MC47 BNT 136 BNT 135 BNT 103 BNT 134 MC 50 MC 49

skeletal opx skeletal opx neoblasts cpx

SiO,(wr%) 18.33 40.03 42.07 41.05 38.74 40.29 37.46 41.39 36.66 33.73 32.66

TiO, 0.02 <0.02 0.02 <0.02 < 0.02 io.02 <0.02 0.09 < 0.02 < 0.02 < 0.02

AI,0 I I .02 0.89 1.61 1.24 0.63 0.5 0.67 1.98 0.36 0.72 0.97

Fe201 7.97 7.7 I 8.59 7.94 8.58 7.93 8.04 13.67 8.59 8.57 9.6 I

MnO 0.1 I 0.11 0.12 0.12 0.13 0.12 0. I3 0.19 0.12 0.13 0.15

MgO 36.74 40.x5 39.57 40.98 37.61 39.69 40.55 31.43 43.08 40.03 39.57

CaO 0.02 0.55 1.1 1.27 0.02 0.21 0.35 6.54 0.21 0.17 0.22

Cr (ppm) 8352 3004 2790 2508 2614 1809 2586 264 1 2997 3176 2225

SC 5 14 9 10 4 5 4 35 8 6 6

V 18 33 43 39 8 3 21 I15 I9 I5 6

CO II9 102 107 105 I30 129 II3 125 126 115 125

Zn 34 32 52 45 43 24 41 63 42 37 44

Ni I503 2148 3511 2198 2171 2606 2280 645 2495 2495 I490

LO1 (wt.%) 13.73 8.3 5.24 6.28 13.46 10.02 I I.18 4.03 9.45 15.52 16.14

Total 99.06 99.03 99 99.41 99.72 99.25 98.94

Mg# 90.13 91.3 90.12 91.09 89.67 90.83 90.9

99.75 99.1 99.55 99.76

81.99 90.85 90.24 89.07

3.3s 0.58 0.24 0.22 CaO/AI,O:

Modal composition Olivine

Orthopyroxene

Clinopyroxene

Spine1

0.59* 2* 0.68 I .03 0.541 0.85’ 0.52

63 60 66 65 70 72 77 30.5 29 27 27 27 26 20

35 10 6 7 2 2 2 3 I 0.8 0.8 0.9 0.6 0.9

49 91

8

0

98

0

99

0

0

0.8

44

0.6

B. Orberger et al. / Lithos 3.5 (1995) 153-182 159

Table 2

Major and trace element composition of clinopyroxenites, websterites and layered gabbros in contact. Mg#= lOO*Mg/(Mg +Fe). am:

amphibolitised, s.am. slightly amphibolitised. BC: basis, close to chromitite, AC: away from chromitite, Ct gb: contact to layered gabbro; LOI:

loss on ignition. Detection limites not reported in table 1: K20: 0.10 (wt.%); Ba, Ce: 15 ppm, Sr, Rb, Y: 2 ppm

Sample no. BNS I?

am.

BNS 78 a

am.

BNS 82

s. am.

BNS 83 a

s. am.

BNT84

s. am.

BNT 105

Contact to gabbro

SiO, (wt.%) 52.30

TiO, 0.20

A&G 3.94

Pe,OJ 6.45

MnO 0.12

MgO 23.29

CaO 10.83

Na,O < 0.50

K,O <O.lO

P*Os <O.Ol

Cr (ppm) 1520

Ba 16

SC 81

Ce 17

V 147

co 50

Zn 14

Ni 1267

Ga 2

Zr 4

Y 4

Sr 32

Rb <2

Nb <2

Pb <4 -

LOI (wt.%) 2.04

Total (wt.%) 99.7 1 99.02 99.4 99.17 99.28 99.68

Mg# 80.00 78.00 83.00 79.00 83.00 70.00

51.32 50.87

0.23 0.1

4.52 1.6

6.36 4.88

0.15 0.12

20.48 20.49

13.19 19.06

< 0.50 < 0.50

<O.lO <O.lO

0.01 <O.Ol

1650 I909

18 25

76 58

< 15 < 15

202 113

49 51

10 17

937 204

4 <2

5

35

<2

2

5

2

<2

47

<2

<2

5

1.84

50.46 48.08

0.15 0.08

2.38 1.29

5.85 5.91

0.13 0.12

19.43 25.94

18.92 15.01

< 0.50 < 0.50

<O.lO <O.lO

<O.Ol <O.Ol

1122 3499

36 24

62 31

< 15 15

152 73

51 62

14 20

238 428

2 <2

4 4

<2 <2

31 17

<2 <2

<2 <2

<4 (4

I .28 2.31

54.12

0.61

8.29

5.55

0.11

1254

13.3

2.07

I.1 I

0.19

2436

299

28

15

119

32

42

151

94

12

341

15

IO

1.42

tinization decreases at some distance of the shear zones

resulting into lower L.O.I. values (Table 1). This is

especially the case of Pak Nai (BNT) and Mae Charim

(MC) samples in which the olivine, clinopyroxene

and/or orthopyroxene survive. Modal compositions of ultramafic rocks have been

estimated by plotting whole-rock major element com-

positions recalculated to anhydrous basis in a MgO vs. SiO, diagram, because it was not always possible to

identify primary minerals through their alteration prod-

ucts (Fig. 4a). Such a diagram, just for illustrative purposes, simulates the projection from spine1 onto the Streckeisen’s olivine-orthopyroxene+linopyroxene ternary, but it gives phase proportions in wt.% (Fig. 4b). In addition, modal compositions have been cal-

culated from whole-rock compositions by least square calculation assuming that olivine (Fo,), enstatite

(En,), pure diopside (Mg%) and Cr-spine1 (Cr/

Cr + Al = 0.5) were the primary phases of the peridotites. Scandium was used to estimate the “primary” CaO content because the CaO was modified by

hydrothermalism. Thus the initial amount of clinopyroxenes was estimated (see below). All the available chromium was allotted to spine1 so that the relative

proportions of this phase might be slightly overestimated. There is good agreement between modal compositions estimated from the MgO vs. SiOl diagrams and computed values. The results are given in Table 1.

The NU peridotites comprise orthopyroxenites, harzburgites and dunites which normally form the man-

160

Table 3

B. Orberger et ~11. / Lithos 35 (1995) 153-182

Representative microprobe analyses of olivines. clino-and orthopyroxenes

Olivines Sample no Rock type

BNT 135 BH 142 BH 142 BNT 133 BNT 133 RNP BNT 138 BNT 138 BNT 103 BNT 134 b Hzbg Dunite Dunite Hzbg Hzbg Lherz Lherz Lherz Dunite chromite

sq. opx sq. opx sq. opx sq. opx sq. opx cpx inclusion

SiOIl (wt.%) 40.49 42.218 41.744 40.697 TiOz 0.029 0.024 0.01 I 0.00 Fe0 8.72 5.994 5.66 I 8.153 MnO 0.131 0.15 0.225 0.1 I5 MgO SO.085 S I .857 52.235 so.794 NiO 0.394 0.29 0.364 0.327 CaO 0.00s 0.00 0.00 0.00 Total 99.855 100.634 100.303 100.08 Cation proportions calculated on the basis of 4 oxygens Si 0.99 I Ti 0.001 Fe 0.179 Mn 0.003 Mg 1.827 Ni 0.008 Ca 0.000 Total 3.0083 mg# 0.910 Clinopyroxene Sample no. A81 Rock type webst.

1.010 I.002 0.99 I 0.989 0.993 0.998 0.996 0.000 0.000 0.000 0.000 0.000 0.000 0.000 0.120 0.190 0.166 0.161 0.195 0.178 0.173 0.003 0.003 0.002 0.002 0.001 0.002 0.002 I.869 1.809 I.843 I.849 I.816 1.819 1.823 0.006 0.006 0.006 0.009 0.002 0.006 0.01 I 0.000 0.000 0.000 0.000 0.000 0.000 0.000 3.01 3.01 3.009 I 3.01 I 3.0092 3.0023 3.0043 0.939 0.943 0.920 0.920 0.903 0.91 I 0.914

0.991 0.000 0.373 0.006 1.636 0.002 0.000 3.0092 0.8 I4

Orthopyroxene

0.999 0.000 0.047 0.000 I .924 0.012 0.001 2.983 1 0.976

A 81 cpx Al4 Al4 A S7 A 157 BC 7li BC7ll AP 87. AP 87 webst. webst webst webst. webst. webst. webst. webst. webst

Si02 52.83 54.14 54.19 S3.98 53.32 56.00 54.84 53.37 56.87 57.16 TiOz 0.07 0.11 0.16 0.26 0.04 0.07 0.08 0.28 0.02 0.03 Fe0 3.80 4.71 4.15 4.88 2.52 2.92 I .96 2.79 S.65 5.41 MnO 0.14 0.19 0.14 0.15 0.2 I 0.18 0.1 I 0.10 0.13 0.15 MgO 12.58 15.18 17.2 IS.75 17.95 17.53 17.62 17.13 34.13 33.97 NiO 0.01 0.00 0.03 0.01 0.02 0.02 0.01 0.01 0.17 0.12 CaO 2s .03 24.49 23.48 23.01 25.54 23.05 24.42 24.23 0.46 0.45 Al?% 4.61 I .04 I.19 2.24 0.18 0.37 0.69 I .77 2.64 2.63

Cr20? 0.06 0.1 I 0.09 0.1 I 0.05 0.05 0.37 0.21 0.48 0.49

Na,O 0.24 0.28 0.06 0.04 0.07 0.09 0.04 0.05 0.01 0.06

Total 99.38 100.26 100.69 100.43 99.9 100.28 100.13 99.94 100.56 100.42

Cation proportions calculated on the basis of 6 oxygens Si I.943 I.985 I.967 I.965 Al IV 0.057 0.015 0.028 0.035 Al VI 0.143 0.030 0.023 0.06 1 Ti 0.002 0.003 0.004 0.007 Cr 0.002 0.003 0.003 0.003 Fe’ + 0.117 0.144 0.126 0.149 Mn 0.004 0.006 0.004 0.005 Mg 0.690 0.830 0.930 0.854 Ni 0.000 0.000 0.001 0.000 Ca 0.987 0.962 0.913 0.898 Na 0.017 0.020 0.004 0.003 Sum cations 3.962 3.998 4.004 3.980 En 38.466 42.850 47.239 44.959 wo s5.019 49.696 46.365 47.224 Fs 14.485 14.817 11.925 14.812

40.5s 1 40.406 40.578 40.81 I 39.018 42.423 0.022 0.003 0.00 0.001 0.008 0.023 7.918 9.266 8.637 8.459 17.576 2.39 0.109 0.132 0.103 0.073 0.279 0.023

SO.892 49.504 49.632 so. 144 43.237 54.79 0.446 0.407 0.299 0.545 0.119 0.628 0.009 0.02 0.00 0.013 0.015 0.00

799.947 99.738 99.248 100.047 100.252 100.904

I.956 2.020 1.977 I.947 1.946 I.955 0.008 0.000 0.023 0.053 0.054 0.045 0.000 0.000 0.005 0.023 0.053 0.061 0.00 I 0.002 0.002 0.008 0.000 0.001 0.00 I 0.001 0.009 0.006 0.013 0.013

0.077 0.088 0.058 0.085 0.162 0.155 0.007 0.005 0.003 0.003 0.004 0.004

0.981 0.942 0.957 0.93 1 I .740 1.731 0.001 0.00 I 0.001 0.000 0.005 0.003 I .004 0.891 0.966 0.947 0.017 0.016 0.005 0.006 0.003 0.004 0.000 0.000

4.04 I 3.957 4.004 4.007 3.994 3.984 47S79 49.047 48.580 47.430 90.69 91.00 48.673 46.368 48.400 47.430 0.88 0.86

7.303 8.548 5.860 8.380 8.50 8.21


BNT 103 0

30) WI 40 50 60 -_

SiO, (wt.%)

Olivine

A B

OPX CPX

Fig. 4a. MgO vs. SiO, for NU peridotites and orthopyroxenites. b. Modal composition of peridotites and olthopyroxenites. Symbols: Open triangles = dunite, dot = harzburgite, full square = orthopyroxenite. open square = pyroxene rich Iherzolite.

tle sequence of suprasubduction-zone ophiolites

(Pearce et al., 1984; Elthon, 1991; Pearce, 1991). Two

harzburgites (MC 53, MC 55) could be classified as

cpx-poor lherzolites since they contain slightly more

than 5 wt.% cpx. Nevertheless, the amount of cpx is

probably overestimated, because it has been assumed that no CaO enters orthopyroxene. This is the reason

why these rocks will be treated as harzburgites. The NU peridotites distinguish from ophiolitic and

erogenic lherzolites described so far by an important

group of orthopyroxene-rich (29-44 wt.% opx, Table 1) peridotites in the Rae Nan, Pak Nai and Prachin Buri area. They all come from the immediate contact of type I and II chromitite lenses.

In fact, all intermediate terms between opx-rich per-

idotites and orthopyroxenites may exist (Table 1) . The orthopyroxene-rich, olivine-poor peridotites are super- imposed by a clinopyroxene-enrichment, resulting in

the case of the cpx-rich dunite from Pak Nai (BNT 103)) in a broadly modal composition of wehrlite (Fig.

4a, b). In the field, these rocks are found close to the large intrusive clinopyroxenite body. In thin section,

large subautomorph clinopyroxenes replace orthopy-

roxenes. Such a wehrlitic modal trend has commonly been described in the mantle sequence of ophiolites and

was ascribed to precipitation of cpx from magmas infil-

trated within formerly refractory rocks (Nicolas and

Prinzhoffer, 1983; Bodinier, 1988).

Late shearing has completely obliterated primary

textures of orthopyroxenites. However, even during intensive serpentinization of the peridotites their man-

tle-derived equigranular neoblastic texture is preserved. The recrystallization of olivine and

orthopyroxene leads to an increase in grain size and

common triple points resulting in a mosaic texture.

Olivine is always present as anhedral, equant crystals

devoid of any prefered orientation and/or kink band-

ing. In the harzburgites, it forms a polycrystalline

matrix or a network isolating pyroxene patches (Fig.

5a). The mutual grain boundaries between each olivine

single crystal are lobate while 120”-triple junctions, typical for olivine neoblasts, are uncommon. Olivine

crystals up to 1 cm in diameter have been observed in

dunites. They are free of silicate inclusions, but contain euhedral Cr-spine1 inclusions. Pegmatoid olivine has

already been recognized in dunites from ophiolites ( Augt, 1983). The orthopyroxene-rich harzburgites

from Pak Nai and Rae Nan distinguish from the other

NU peridotites by the occurrence of a second genera-

tion of much smaller olivine crystals. The latter are

arranged as chains or patches corroding orthopyroxene

(Fig. 5b).

In the harzburgites, orthopyroxene (opx) occurs as small, ( <2OO by 200 pm in maximum dimensions)

equant crystals disseminated in the olivine network (Fig. 5b). It forms much larger elongated crystals up

to 1 cm length in the Pak Nai and Rae Nan opx-rich peridotites (Fig. 5~). The orthopyroxene crystal shapes

are euhedral or anhedral. Occasionally they show corrosion gulfs and embayments filled with polycristalline olivine as well as numerous olivine droplets inclusions. In an ultimate stage, the opx crystals are truncated and isolated within the second generation of fine-grained olivine (Fig. 5~). Such textural relationships point to

162 B. Orberger et al. / Lirhos 35 (1995) 153-182


a replacement of orthopyroxene by olivine and have

been described by Quick ( 1979) and Reimaidi ( 1993).

The negative correlation between olivine and ortho-

pyroxene modal proportions in the opx-rich peridotites

confirms this observation (Table 1). Clinopyroxene

(cpx) is often transformed into amphibole which has recrystallized into euhedral crystals during the regional

metamorphism coeval to the emplacement of the

ophiolite (Fig. 5d). When unaffected by amphiboliti-

zation, like in the Pak Nai pyroxene-rich peridotites,

cpx shows all the characteristics of the so-called

“impregnation” cpx ascribed to magma percolation by

Nicolas and Prinzhoffer ( 1983; Fig. 5d). It forms large

(500 by 300 pm) undeformed crystals independent of

orthopyroxene and/or spinel. Magmatic twins are

locally observed, especially in the cpx-rich dunite

(BNT 103). There, patches of several tens of cpx crys-

tals particularly rich in green spine1 exsolutions occur

within the olivine matrix (Fig 5d).

Chrome-spine1 is accessory ( < 1 wt.%) except in

three samples (MC 40, RN 63 and 65; Table 1) which

were collected close to chromitites lenses and layers

(type I and II). The shapes of spine1 crystals vary

between each peridotite sub-type (Fig. 5e, f). Harz-

burgites are characterized by ameboidal anhedral Cr

spinels. In the Pak Nai pyroxene-rich peridotites, Cr-

spine1 is randomly distributed irrespective of orthopy-

roxene and displays various shapes ranging from

anhedral to perfectly euhedral. The anhedral spine1

combines rounded, embayed, spongy and atoll-like tex-

tures (Fig. 5e). The atoll-like spine1 grains enclose

silicates (mainly olivine) forming a central cavity

inside the spinel. Such textures have been interpreted

as the result of corrosion and recrystallization of Cr-

spinels faced with “exotic” percolating melts (Lebl-

ant, 1980; Lorand and Cottin, 1987). The euhedral

Cr-spinels were formed when equilibrium was reached with this melt (Fig. 5f). They contain negative minute

silicate inclusions consisting of both anhydrous and

hydrous silicates evenly distributed. Euhedral Cr-spinels are included in both, olivine and orthopyroxene

(Fig. 5f).

4.2. Intrusive pyroxenites

Samples within the intrusive pyroxenite body have

been collected close to the contact with the type III

chromitite lenses (BNS 71, 72, 77, 78), in the central

part of the body up to the contact with the layered

gabbro sequence (BNT 105). All of them have inter-

mediate modal composition between clinopyroxenite

and websterites (opx + 01 Q 20 vol.%). The textures

and mineralogy of the clinopyroxenites-websterites

vary according to their locations. The clinopyroxenites

close to type III chromitite are almost completely trans-

formed to Ca-amphibole and recrystallized into mosaic

texture. The observed relict clinopyroxenes are free of

spinel-exsolutions. Upsection (BNT 82, 83) up to the contact with the layered gabbros (BNT 105) the cli-

nopyroxenites are sometimes less than 10% amphibolitized and display typical magmatic textures (Fig. 5g).

These clinopyroxenites distinguish from those in the

vicinity of type III chromitite by numerous spinel-exsolutions within clinopyroxenes and the presence of iron-

titanium oxides which are interstitial to the cpx or

included in amphibole. The cpx is texturally similar to

those observed in the cpx-rich dunite (BNT 103) rich

in Cr-spine1 exsolution.

4.3. Chromitite ore bodies

The three types of chromitites recognized in the field can also be distinguished by their texture and miner-

alogy of the silicate matrix. Type I chromitite lenses are massive. Interstitial sil-

icates represent less than 20 vol.%. The nature of interstitial silicates and wall rocks to type I chromitite

orebodies vary on a regional scale. The most massive

Fig. 5. Microtextures of the ultramafic rocks and chromitites. a. harzburgite (MC 53); b. harzburgite with opx I crystals interstitial to partly

serpentinized olivine relicts (BNT 135); c. orthopyroxene-rich lherzolite displaying skeletal orthopyroxene II crystals corroded and truncated

by polycristalline olivine (BNT 138); d. large undeformed cpx crystals (“impregnation” clinopyroxene) partly transformed into amphibole in

the cpx-rich dunite (BNT 103); e. ameboidal anhedral Cr spinels in harzburgite showing incipient atoll like texture (MC 55); f. euhedral Cr

spinels in dunite (MC 49); g. intrusive clinopyroxenite showing mosaic-textured clinopyroxene (BNS 83 a); h. massive type I chromitite, the

interstitial silicates are Mg chlorites and serpentine (MC 1); i. type II layered chromitite composed of euhedral Cr spine1 crystals; the matrix

silicates are serpentines (RN 66); j. type III chromitite in contact with amphibolitized clinopyroxenite; the matrix silicates are chnopyroxene

or tremolite (BNS 7 I ).

164 B. Orberger et al. / Lithos 35 (1995) 153-182

ores from Mae Charim and Prachin Buri are mainly

hosted in orthopyroxenites. Their interstitial silicates

consist largely of orthopyroxene, now partly bastitized (Fig. 5h). Graphite flakes have been observed in sam-

ple MCI. The formation of Mg-rich chlorite and ser-

pentine suggests that the silicate matrix contained in

minor amount olivine prior to serpentinization. Chro-

mite is massive and free of deformation, except in mil-

limeter-thick narrow bands along shear planes caused

by the emplacement of the ophiolite. It exhibits incip-

ient replacement by ferritchromite along these fracture zones.

Unlike the Mae Charim and Prachin Buri orebodies,

Pak Nai type I chromitites have dunitic wall rocks and

exhibit nodular textures. Chromite nodules are closely

packed. To the contact of the interstitial silicates, they

display crystalline faces. Olivine droplets up to 100 pm

in diameter that are randomly distributed are the only solid inclusions observed in the type I chromitite (Fig.

5h). A great number of these inclusions are unserpentinized. Interstitial silicates may occupy up to 30-40%

of the total volume of the Pak Nai type I chromitite

orebodies. As shown by the abundance of serpentine

and Mg-chlorite, olivine was predominant in the sili-

cate matrix prior to serpentinization.

Type II chromitite layers and bands are made up

from undeformed delicate chromite crystals (about 300

pm), closely packed and generally showing well devel-

oped crystalline faces (Fig. 5i) with little signs of cor-

rosion features. Matrix silicates amount up to 50 vol.%.

As observed for type I chromitites, the Pak Nai type II

chromitites have dunitic wall rocks and thus olivine

dominated prior to serpentinization, coexisting with opx and minor amphibolitized cpx. In contrast, the Rae

Nan type II chromitite layers are hosted by orthopyrox-

enites and relict opx are predominant among the interstitial silicates. Solid inclusions in chromite are

numerous. They have all the characteristics of those described in chromitites (Talkington et al., 1984; Lor-

and and Ceuleneer, 1989; Bacuta et al., 1990). These are mainly negative crystals, but also rounded cavities with scallopped margins. Unlike the single phase olivine inclusions in type I chromitite, the solid inclusions in type II chromitite are polyphase. The commonest mineral assemblage is olivine + phlogopite, sometimes

containing minute base metal sulfides. Type III chromitite lenses distinguish from the type

II chromitites by the abundance of interstial clinopy-

roxene (70-80 vol.%), coexisting with minor olivine

and opx (Fig. 5j). Compared to the host clinopyroxenites, the cpx interstitial to the chromite crystals is free

of any Cr-spine1 and opx exsolution. A single cpx crys-

tal may enclose several tens of chromite crystals. Cli-

nopyroxene is partly amphibolitized due to the

emplacement metamorphism while opx and olivine are

partly bastitized and serpentinized, respectively. The

shapes of chromite crystals seem to have been modified

by the amphibolite metamorphism. When hosted in

cpx, chromite occurs as euhedral to anhedral crystals

showing slight embayments and corrosion features

(Fig. 5j) enclosing the same silicate assemblage as the

type II chromitite. When occurring in amphiboles,

chromite forms perfect octahedral translucent crystals and is free of any inclusion.

5. Whole-rock chemistry

5.1. Peridotites and orthopyroxenites

Major, minor and trace transition elements have been

plotted against A1203 which is both moderately incom-

patible in mantle melting processes and inert during hydrothermal alteration (Dungan, 1979) (Fig. 6). In

addition transition elements have been normalized to

SC1 (Jagoutz et al., 1979) and presented in spider-

diagrams in Fig. 7.

The majority of the NU peridotites have MgO similar

to those of mantle peridotites. High field strength ele-

ment concentrations are all close to (e.g. TiO,) or below (e.g. Zr, Nb) the detection limits of the XRF

procedure. Low contents of high-field strength ele-

ments and mantle-normalized plots of Fig. 7 reveal strong analogies with peridotites from supra-subduc-

tion zone ophiolites such as the Leka complex, in Nor- way (Furnes et al., 1992). Of course, the positive Cr

anomalies in these plots correspond to the chromite- rich samples (MC 40, RN 63, RN 65). Likewise, the abundance of cpx in wehrlite BNT 103 accounts for the high V and SC and low Ni contents compared to the other peridotites. In contrast, the Co and Ni concentration peaks observed in Prachin Buri and Mae Charim samples (up to 1 wt.% Ni; Figs. 6 and 7)) reflect incipient lateritization which might also have disturbed the

vanadium concentrations.

B. Orberger er al. / Lithos 35 (1995) 153-182 165

6

0 1 2 3

1

1 0 II I 2 3

cl

. 0

0

140 1

I20 0

100

80

60

JO

20

0 I 2 3 4

Al 2

0 3

(w1.?‘o)

Fig. 6. A-J. Major and trace element covariation diagrams of the peridotites and orthopyroxenites as a function of A&O3 contents. Whole rock

analyses have been recalculated to an anhydrous basis according to the formula: (element in oxide wt.%) / ( 100 - LOI) * 100. Same symbols as

in Fig. 4.

166 B. Orberger et al. / Lithos 35 (1995) 153-182

10.” 7 ’ 1 r 1 1 I 1 SC Ti V Cr MnFe Co Ni

10’

10”

10

10

I’eridolite Leka

SC Ti V Cr Mn Fe Co Ni

Fig. 7. Mantle normalized (SC 1) transition element patterns of the NU peridotites and orthopyroxenite NU ophiolite (a) and the Leka ophiolitic

peridotites (b) (Fumes et al., 1992); SC1 composition after Jagoutz et al. ( 1979).

Variations of MgO and trace transition elements V,

SC and Co in all of the peridotites-subtypes follow the

general trends defined from unserpentinized worldwide

erogenic and ophiolitic peridotites (Fig. 6a-i). MgO

and roughly Co concentrations increase with decreas-

ing A1203 contents, in agreement with their compatible

characters whereas the contents of the incompatible elements V and SC decrease (Fig. 6). The pyroxene-

rich peridotites are more or less fertile with A&O3 con-

centrations ranging from 0.6 to 3.12 wt.% (Table 1) . Dunites and harzburgites are strongly refractory and

are similar to those of other ophiolitic peridotites stud-

ied so far (e.g. Jaques and Chapell, 1980; Furnes et al.,

1992). The data points plot at the depleted apex of the

mantle trends. However, their Mg# are low (89-9 1)

compared to the other ophiolitic dunites and do not

vary between harzburgites and dunites. Mg numbers

decrease in the pyroxene-rich peridotites with the abun-

dance of cpx (90-92 vs. 82 in sample BNT 103)

whereas those of the orthopyroxenites show more var-

iable Mg#. The high Mg-number (94.3) of sample

MC 40 (Fig. 6b, Table I) is due to the high Cr content

of the sample (1.4 wt.%) which leads to a low Fe0

content (4.37 wt.%, Table 1). The CaO and TiO, contents totally contrasts with

those of reference mantle peridotites (Fig. 6c, d) . The

cpx-rich dunite BNT 103 has much higher CaO/Al,O, ratios than the mantle peridotite (3.3 vs. cO.87,

respectively; Table 1) . This indicates the addition of Al-poor cpx compositionally similar to those from the intrusive clinopyroxenites. Serpentinized dunites have

CaO/Al,O,-ratios of unserpentinized dunites (see

Bodinier, 1988). Harzburgites and pyroxene-rich peridotites in contrast are almost devoid of CaO in spite

of microscopic and modal evidence of cpx (Table 1; Fig. 6~). Here calcium was removed by hydrothermal

alterations which can be confirmed by looking at SC

which is incorporated into the cpx structure and inert to serpentinization (Wedepohl, 1968-1978). The samples with CaO/A1203 ratios close to zero have high SC/

CaO ratios, consistent with the presence of cpx prior to

alterations (Fig. 8). Using the SC vs. CaO diagram the

initial CaO content can be recalculated by assuming Sc/CaO ratios of 7 X lop4 in unaltered mantle peri-

dotites (Bodinier, 1988; Bodinier et al., 1988). The

results show all the pyroxene-rich Iherzolites, but three

1 A Dunlte

6.. n Orthopyraxenite 0 Px-rich peridotite

l Harzburgile

i 0 10 2 0 3 0 4 0

SC (PP”lI

Fig 8. CaO vs. Scandium diagram for peridotites and orthopyroxenites. Symbols as in Fig. 4.


0.00 0.02 0.04 0.06 0.08 0.10

Ti02 (wt.%) Fip. 9. Scandium vs. TiOz for peridotites and orthopyroxenite. Sym-

bols as in Fig. 4

0.6

A A

A

A

A4

A

AA

A AAA AA A A

0.7 0.8 0.9

Mg/Mgt Fe

F 0.6. BNT 105 A

5

6 0m4- A

r+ A

A 0.2. A

A A

A A A

A A4

0.01

3.6 0.7 0.8 0.9

MglMgtFe

Fig. 10a. CaO/A1,O1 vs. Mg#; b. TiO, vs. Mg# forclinopyroxenites

and websterites.

Rae Nan samples had initial CaO/A1203 higher than predicted by mantle melting trends ( 1 .O-2.0 vs. 0.87- 0.20; e.g. Frey et al., 1985). Such ratios provide strong

support to the hypothesis of a general cpx-enrichment

trend in the NU pyroxene-rich peridotites, culminating in the cpx-rich dunite BNT 103. It is suggested that

calcium was removed during hydrothermal alteration

of the oceanic crust before the emplacement of the

ophiolite (cf. Coleman and Keith, 1971) rather than

during amphibolitization related to its emplacement.

Titanium oxide contents are generally below the

detection limit of the XRF-analyses (0.02 wt.%; Table

1) for all of the rock types but the cpx rich dunite (BNT

103). Titanium is believed to be inert during hydro-

thermal alteration (Pearce and Norry, 1979)) thus its

low contents cannot be attributed to the serpentiniza-

tion and other metamorphic events. One has to note that the pyroxene-rich peridotites have the lowest

Ti02/A1203 of the NU peridotites and their Sc/TiO,

ratios is much higher than the typical mantle melting

arrays (Fig. 9). This observation suggests that cpx has precipitated in these rocks from a low-TiO, magma.

5.2. Intrusive clinopyroxenites

Clinopyroxenites display large variations in CaO/

A&O,-and Mg/Mg + Fe-ratios as well as TiO, and Cr

contents (Fig. lOa, b; Table 2). In a CaO/Al,O, vs.

Mg/Mg + Fe diagram, one group of samples is char-

acterized by high and constant Mg# (86-90). It must

be pointed out that all these samples have been col-

lected away from layered gabbros. The variation of

CaO/Al,O, is clearly related to the replacement of cpx

into Ca-amphibole (actinolite, tremolite; Fig. 10a).

The less amphibolitized samples (BNS 82, 83) have

typical cumulate compositional features and resemble the late intrusive wehrlitic bodies crosscutting the top

of the mantle sequence and the layered gabbroic cumu-

lates of the Semail ophiolite (Lachize, 1993).

Mg# decrease with increasing TiO, contents in the

samples collected at the immediate contact of the lay-

ered gabbros of the oceanic crust (Fig. 1 Ob). The clinopyroxenite sample BNT 105 is characterised by the

lowest Mg# and CaO/A1203 ratio and the highest TiO, content of all the NU clinopyroxenites analysed (Table 2, Fig. lob).

6. Phase chemistry

Oliuine has typical mantle forsterite (Fo) and NiO contents (Table 3; Fig. 11). Its Mg-number reflects

168 B. Orberger et cd. / Lithos 35 (1995) 153-182

0.015 Olivine

O.OlO-

0.000 7 0.8 0.9 1 .o

0.02

Mg/(Mg+Fe)

A Clinopgroxene

1 AA

AF

cp

AAA c A

L

0.00 0.0 0,l 02 O-3

Al IV +\‘I

1.0 C

A A

A cpx in cumulates

B

A A

a G 0

0.8 ’ I 0.0 0. I 0 .2 0.3

AI IV + VI

Fig. I 1, Chemical composition of olivine, clino- and orthopyroxenes. a. Ni vs. Mg/Mg + Fe in olivine from peridotites and chromitite; b, c. clinopyroxene composition plotted in the Ti vs. Al and MglMg + Fe vs. Al diagrams (Ti and Al as cation proportion). Clinopyroxenes in

cumulates: data from Augk 1983; Emewein et al., 1988; HBbert and Laurent, 1989.

B. Orberger el al. / Lirhos 35 (1995) 153-182 169

whole rock Mg#. The lowest Mg# (Fo,,) and NiO

contents correspond to the cpx-rich dunite (BNT 103).

The dunitic envelopes of small type I chromitite lenses contain more magnesian olivine (Fo,,) in agreement

with their whole rock Mg# (94). Similar Fo-rich oli-

vine were described in the dunitic envelopes of podi- form chromitites from the New Caledonia ophiolite

(Johan and AugC 1986). In that case, olivine is not

NiO enriched compared to that in the other NU peri-

dotites. In contrast, olivine enclosed in type I chromitite

lenses displays extremely high Fo~~,_~~) and NiO contents (0.63-0.76 wt.%) that are commonly reported

from ophiolitic massive chromitite (Talkington et al.,

1984 and references therein). As suggested by many authors (e.g. Irvine, 1967; Hatton and von Gruene-

waldt, 1985) and experimentally demonstrated by Leh-

mann ( 1983), the high Mg-number of the olivine inclusions can be explained by preferential partitition-

ing of MgO into olivine during subsolidus re-equili-

bration with the volumetrically dominant host chromite. Except those peridotites collected close to

chromitites and intrusive clinopyroxenites, the Mg-

number of olivine vary in a narrow range from Fog0

toFo,,, irrespective of modal compositions and refrac-

tory index of host peridotites (Table 3). It does not

increase significantly from the less depleted harzbur-

gite to the strongly refractory dunites.

(Cr = 0.95). All three types of chromitites display compositional variations on a regional scale, that are

clearly related to mineralogical variations in host wall rocks and chromite-to-silicatemass ratios in each single

orebody.

Chromites from type I lens-shaped chromitites (Fig. 12b) are among the most chromiferous chromitites

ever described in ophiolites (see Haggerty, 1991) . The

massive lenses of Rae Nan and Mae Charim hosted in

orthopyroxenites have the highest Cr203 contents (up

to 68.1 wt.%) and Cr/Cr + Al ratios (Cr# = 0.95) in

Fig 12b. These spinels are almost devoid of TiOz and

ferric iron (Table 4). To the exception of inclusions in

diamond and meteorite chromites, highly chromiferous

chromite low in TiO, characterizes extrusive boninites

(Fig. 12). Similar chromite composition are known

only from the Heazlewood River layered complex in

Tasmania, also of definitely boninitic affinity (Peck

and Keays, 1990) and in orthopyroxenite dykes cross-

cutting New Caledonia and Papua New Guinea ophiol-

ites (Jaques and Chapell, 1980; Leblanc, 1985).

Orthupyroxenes are enstatites (En: 90.7-9 1 .O, WO:

0.4-0.9 and Fs: 8.2-8.8; Table 3).

Clinopyroxenes are diopsides, ranging from En:

43.77, Wo: 49.2, Fs: 13.4 to En: 55.8 Wo: 48.4 Fs: 7.00

to (Table 3). Their Mg# reflect bulk rock compositions (Table 3). The most magnesian diopsides are

characterised by very low TiOz and A1203 contents. Na,O contents are mainly below the detection limit of

0.5 wt.%. These compositions resemble clinopyroxenes analysed in intrusive wehrlites and clinopyroxeni-

tes from Troodos, Oman and Thetford Appalachian ophiolites (AugC, 1983; Ernewein et al., 1988; Htbert and Laurent, 1989). This comparison provides strong

support to the interpretation of the NU intrusive clinopyroxenites as early cumulates from island arc magmas of broadly boninitic composition. Likewise,

crystallization of iron-titanium oxides may explain the subtle variation of TiOz.

Compositional variations of the type I chromitites

involve both Cr/Cr + Al (Cr#), Mg/Mg + Fe (Mg#)

and Fe3+/Ti ratios. The orthopyroxenite-hosted mas-

sive lenses (Mae Charim and Prachin Buri, type I) have the highest Mg#‘s (0.60-0.75) (Fig. 12b). The

Cr#-decrease between the two areas is balanced by

increasing Al and Ti contents at nearly constant Fe3+ contents (Table 4). The Pak Nai type I chromitites

associated with dunites are characterized by large var-

iations of both Cr# (0.954.7) and Mg# (0.75-0.28).

Cr# decrease in response to decreasing Cr,03 and

increasing Fe”+ and Ti contents. A1203 increases only slightly compared to the Mae Charim type I orebody.

These variations displace the data points toward the

field of island-arc tholeiitic magma spine1 field in Fig. 12e, f. The large variations of Mg# reflect the chro-

mite/olivine mass ratios. They must be attributed to subsolidus re-equilibration of the Fe-Mg partitioning

with their host olivine. A decrease in Mg# is neces-

sarily accompanied by coeval increases in Cr#‘s (Dick and Bullen, 1984), because Mg forces Al to enter the

spine1 structure (e.g. Sack, 1982). This is the reason for the large scatter of the data points corresponding to the Pak Nai type I chromitites in Fig. 12b.

Chromite compositions span all the range defined for On average, type II layered and banded chromitites alpine-type peridotite hosted chromitites (Fig. 12a) associated with orthopyroxenites have lower Cr# and even extend toward much higher Cr-ratios (0.63-0.79) than type I chromitites. Regional com-

170 B. Orberger et al. / Lirhos 35 (1995) 153-182

e 0.X

h t

o 0.6

2 3 0.4

1.0

0.8

Il.6

0.4

0.2 tes 0.0

[ ' chromite -- -_/

bonini

0.X 0.6 0.4

Mg/Mg+Fe

0.2 0.0 0.00

Mae Charim dunite hosted b

inites

0.65 ,--typelj 0.R 0.6 0.4 0.2

.

0.00 0.10

Fe3+/(Fe3++Cr+Al)

0.6 0.4 0.2

Mg/(Mg+Fe)

type III d 0.4 0.6

Cr/Cr+Al

0.X 1.0

0.16’ cpxte

0.75.

0.6s

(I.5.s

0.45

ll.J5

0.2s 1 08 0.6 0.4

Mg/(hlg+Fe)

0.2


positional variations in type II chromitites produce two trends in Fig. 12~. The opx-hosted Rae Nan type II layered chromitites have a restricted Mg# range (0.63- 0.70) and their data point also partly overlap the boninitic spine1 field (Fig. 12a). This is due to the Cr# decrease in response to both decreasing Cr,O, and increasing AlzO, at nearly constant Fe,03 and TiO, contents relative to type I (Fig. 12e, f, g). In contrast, like their type I nodular counterpart, the Pak Nai type II banded chromitites hosted in dunites have variable Mg# (0.70-0.35) at nearly constant Cr# (0.75-0.80). This opposite behavior of Mg# between opx- and olivine-hosted chromitites attests to the efficiency of Mg partitioning into olivine relative to opx at decreasing temperature (see Irvine, 1967; Fabribs, 1979). The decrease in Cr# relative to the type I chromitites is balanced by higher TiOz contents (up to 0.32 wt.%) while Al,O, increases less rapidly. For that reason, the data points of the Pak Nai type II chromitites protrudes into the island-arc magma spine1 field in Fig. 12e, f, similar to spinels from chromitite type I.

Unlike the type II chomitites spinels from the type III chromitite lenses in intrusive clinopyroxenites, have relatively constant Mg# (0.47-0.60) whereas their Cr# largely vary from 0.8 to 0.3. In diagrams involving TiOl and Fe203 (Fig. 12e, f, g), the type III chromitites clearly distinguish from the other NU chromitite orebodies by higher TiOz and Fe3+ /Fe3+ +A1 +Cr (up to 0.8 wt.% and 0.12, respectively; Fig. 12e, f). This displaces the data point outside the boninitic spine1 field but toward island-arc low-Ti tholeiitic field. The type III chromitites also show marked compositional similarities with disseminated chromite orebodies hosted in intrusive wehrlites and websterites from Oman and Bay of Islands ophiolite complexes (Fig. 12a).

7. Discussion

7. I. Origin of the peridotites: partial melting residues or mantle/melt reactions products?

Strongly refractory peridotites (harzburgites and dunites) found in oceanic or continental lithospheric mantle have usually been interpreted as resulting from high degrees of partial melting affecting a chemically (basalt undepleted) homogeneous mantle (Dick and Bullen, 1984; Dick et al., 1984; Frey et al., 1985; Bod- inier, 1988; Bodinier et al., 1988; Johnson and Dick, 1992). The strong hydrothermal alteration events that have affected the NU peridotites only imply the use of the most inert elements i.e. Mg, Al, Ni and SC (Fig. 6) to obtain some informations on the degrees of melting.

According to Kostopoulos ( 199 1) about 42% melting is required to dissolve cpx completely from a MORB-pyrolite mantle, whereas the formation of opx-free dunites would require even higher melting degrees ( > 60%). It is unlikely that such high melting degrees could be produced by a single-stage process. Jaques and Green ( 1980)) MC Kenzie ( 1984) and Ribe (1988) showed that melts are readily extracted by buoyancy-driven segregation from a deformable matrix. Two-stage or three-stage melting processes are often proposed in the literature to explain the formation of highly magnesian lavas (e.g. Duncan and Green, 1987). This would result in a much more pronounced incompatible-element depletions than that observed in the Nan Uttaradit dunites and opx-poor harzburgites. The protolith of the NU ophiolitic peridotite is unknown. Textural features of the cpx, as well as CaO/ A&O: > 1 and unusually high recalculated Sc/TiO, of the cpx rich lherzolites rule out the hypothesis that these rocks could represent the undepleted mantle source of the NU peridotites. The source has therefore been assumed to have the composition of the undepleted spine1 lherzolite R717 of Frey et al. ( 1985). Corre- sponding abundances for representative traces are cal-

Fig. 12. Cr-spine1 compositions of type I, II and III chromitite. a. Cr/Cr + Al vs. Mg/Mg +Fe summerizing the compositional field of the NU chromitites compared to massive and disseminated chromites from ophiolitic websterite (Newfoundland; Dick and Bullen, 1984; Haggerty, 1991) and chrome spine1 from boninites (Arai, 1992: Roeder and Reynolds, 1991); b-d. enlargement of (a) showing variations in Cr# and Mg# composition in the three chromitite types as a function of their hostrock; e. TiOz versus Fe3+ /Fe’+ + Cr + Al, compared to those in mid- ocean ridge (MORB), island arc basalts and boninites (Arai, 1992); f. enlargement of (e) showing the variations according to chromite types and their hostrocks. g. TiOz vs. Cr/Cr + Al are compared with spinels cristallized from MORB-, island-arc- and boninite-type magmas. Circle: type I dunite hosted, dot: type I orthopyroxene hosted, open triangle: type II dunite hosted, full triangle: type II orthopyroxenite hosted, open square: type III.

Tab

le 4

R

epre

sent

ativ

e m

icro

prob

e an

alys

es

of C

r-sp

inel

s fr

om

lens

-sha

ped

chro

miti

te

(typ

e I)

, la

yere

d ch

rom

itite

(t

ype

II)

and

xeno

litic

ch

rom

itite

(t

ype

III)

. n.

a. =

not

ana

lyse

d

Typ

e I

Typ

e II

T

ype

III

Pods

N

odul

ar

Stra

tifor

m

vein

s B

ande

d L

ense

s H

ostr

ock

Ort

hopy

roxe

nite

D

tmite

D

unite

O

rtho

pyro

xeni

te

Ort

hopy

roxe

nite

D

unite

cp

xeni

teew

ebst

erite

Are

a M

ae C

hari

m

Prac

hin

Bur

i Pa

k N

ai

Pak

Nai

R

ae N

an

Rae

Nan

Pa

k N

ai

Pak

Nai

WA

66

.74

67.3

8 60

.58

62.2

7 60

.77

62.6

4 68

.14

58.1

4 62

.03

54.2

5

AL

O,

4.80

5.

29

9.66

9.

2 6.

64

6.78

5.

6 6.

53

3.92

16

.44

FeO

, 15

.07

14.2

4 14

.59

14.1

9 24

.17

18.0

7 15

.43

27.8

2 27

.97

15.8

8

TiO

, 0.

09

0.09

0.

23

0.23

0.

09

0.15

0.

05

0.17

0.

14

0.08

MgD

13

.39

13.0

4 15

.23

15.0

2 8.

44

12.8

1 11

.29

7.19

6.

11

13.5

9

MnO

0.

22

0.24

0.

24

0.24

0.

47

0.32

0.

25

0.44

0.

56

0.12

NiO

0.

03

0.08

0.

19

0.21

0.

06

0.14

0.

08

0.06

0.

03

0.18

CaO

0.

00

0.00

n.

a.

n.a.

n.

a.

na.

0.05

n.

a.

ma.

0.

00

Na,

O

0.00

0.

00

n.a.

n.

a.

n.a.

n.

a.

na.

n.a.

n.

a.

0.00

K

,O

0.00

0.

01

“.a.

n.

a.

na.

n.a.

n.

a.

n.a.

n.

a.

0.00

V

*O,

“.a.

n.

a.

0.07

0.

06

na.

n.a.

n.

a.

0.14

0

n.a.

Z

nO

n.a.

n.

a.

0.01

0.

01

n.a.

na

. n.

a.

0.12

0.

12

n.a.

Si

02

0.04

0.

02

n.a.

n.

a.

n.a.

na

. n.

a.

na.

n.a.

0.

00

Tot

al

100.

38

100.

39

100.

80

100.

23

100.

64

100.

91

100.

89

100.

61

100.

72

100.

54

Cat

ion

prop

ortio

ns

calc

ulat

ed

on t

he b

asis

of

4 o

xyge

ns

Cr

1.73

4 1.

752

1.52

2 1.

525

1.61

6 1.

611

Al

0.18

6 0.

205

0.36

2 0.

397

0.26

3 0.

260

Fe’

+ 0.

074

0.03

7 0.

104

0.06

6 0.

116

0.12

2 T

i 0.

002

0.00

2 0.

005

0.00

6 0.

002

0.00

4 Fe

2 +

0.34

0 0.

355

0.23

7 0.

302

0.56

4 0.

370

Mg

0.65

6 0.

639

0.72

1 0.

692

0.42

3 0.

621

Mn

0.00

6 0.

007

0.00

6 0.

006

0.01

3 0.

009

Ni

0.00

1 0.

002

0.00

5 0.

006

0.00

2 0.

004

Ca

0.00

0 0.

000

0.00

0 0.

000

0.00

0 0.

000

Na

0.00

0 0.

000

o.oo

o 0.

000

0.00

0 0.

000

K

0.00

0 0.

000

0.00

0 0.

000

0.00

0 0.

000

V

0.00

1 0.

000

0.00

0 0.

000

0.00

0 0.

000

Zn

0.00

0 0.

000

0.00

2 0.

001

0.00

0 0.

000

Sum

Cat

. 3.

000

2.99

9 2.

962

3.00

0 2.

999

3.00

1

1.78

0 I.

569

1.70

5 1.

339

1.21

4 1.

296

I .34

4

0.22

0 0.

263

0.16

1 0.

605

0.71

9 0.

480

0.48

4

0.00

0 0.

156

0.12

8 0.

053

0.03

1 0.

205

0.15

3

0.00

0 0.

004

0.00

4 0.

002

0.00

3 0.

009

0.00

9

0.43

0 0.

62 1

0.

667

0.36

2 0.

307

0.38

6 0.

47 I

0.

560

0.36

6 0.

317

0.63

2 0.

703

0.60

7 0.

530

0.01

0 0.

013

0.01

6 0.

003

0.00

1 0.

008

0.00

6

0.00

0 0.

002

0.00

1 0.

004

0.00

0 0.

003

0.00

2

0.00

0 0.

000

0.00

0 0.

000

0.00

0 0.

002

0.00

0

O.O

QO

0.

000

0.00

0 0.

000

0.00

5 0.

002

0.00

0

0.00

0 0.

000

0.00

0 0.

000

0.00

0 O

.OQ

O

0.00

0

0.00

0 0.

004

0.00

0 0.

000

0.00

0 0.

000

0.00

0

0.00

0 0.

003

0.00

3 0.

000

0.01

8 0.

000

0.00

0

3.00

0 3.

001

3.00

2 3.

000

2.98

3 2.

996

2.99

9

Mg#

0.

658

0.64

3 0.

750

0.70

0 0.

429

0.62

7 0.

570

0.37

1 0.

322

0.63

6 C

r#

0.90

3 0.

895

0.81

0 0.

790

0.86

0 0.

861

0.89

0 0.

856

0.91

4 0.

689

Fe#

0.

037

0.01

9 0.

052

0.03

3 0.

058

0.06

1 0.

000

0.07

8 0.

064

0.02

6

0.69

6

0.62

8

0.01

6

0.53

0

0.73

5

0.10

3

50.2

3 52

.18

52.2

1

19.9

4 12

.97

12.6

0

13.2

1 22

.49

22.9

0

0.11

0.

38

0.39

15.4

2 12

.97

10.9

3

0.05

0.

31

0.22

0.00

0.

13

0.09

0.01

0.

07

0.00

0.08

0.

03

0.00

0.00

0.

00

0.00

0.00

0.

05

0.0

I 0.

00

n.a.

n.

a.

0.58

n.

a.

“.a.

99.6

3 99

.50

99.3

5

0.52

8

0.73

3

0.07

7

57.7

5 57

.08

53.3

0 55

.52

49.9

8 26

.67

11.1

0 11

.15

10.6

8 9.

99

9.93

30

.93

16.7

3 18

.44

25.3

3 23

.74

22.4

6 28

.38

0.32

0.

32

0.33

0.

32

0.33

0.

50

13.9

6 12

.91

10.5

0 9.

80

8.64

Il

.52

0.30

0.

33

0.33

0.

49

0.39

0.

36

ca

0.10

0.

10

0.10

0.

12

0.12

0.

09

0

0.00

0.

00

0.00

0.

00

3.90

0.

01

P 0.

00

0.00

0.

00

0.02

0.

01

0.02

oz

t 7

0.00

0.

00

0.00

0.

01

0.01

0.

00

z 0.

00

0.00

0.

00

na.

na.

n.a.

E

0.00

0.

00

0.00

n.

a.

na.

n.a.

G

0.00

0.

00

0.00

0.

01

3.88

0.

13

B

100.

26

100.

33

100.

57

100.

01

99.6

4 98

.63

zJ

CII

;:

1.45

5 1.

447

1.37

2 1.

448

1.29

4 0.

636

E

b

0.41

7 0.

422

0.41

0 0.

389

0.38

3 1.

101

s 0.

112

0.11

6 0.

203

0.14

8 0.

053

0.23

3

0.00

8 0.

008

0.00

8 0.

008

0.00

8 0.

011

z 0.

334

0.37

9 0.

487

0.50

7 0.

562

0.48

3

0.66

3 0.

617

0.50

9 0.

482

0.42

2 0.

518

0.00

8 0.

009

0.00

9 0.

014

0.01

1 0.

009

0.00

3 0.

003

0.00

3 0.

003

0.00

3 0.

002

0.00

0 0.

000

0.00

0 0.

000

0.13

7 0.

000

0.00

0 0.

000

o.cO

O

0.00

1 0.

001

0.00

1

0.00

0 0.

000

0.00

0 0.

000

0.00

0 0.

000

0.00

0 0.

000

0.00

0 0.

000

0.00

0 0.

000

0.00

0 0.

000

0.00

0 0.

000

0.12

7 0.

004

3.00

0 3.

001

3.00

1 3.

000

3.00

1 2.

998

0.66

5 0.

620

0.51

I

0.48

7 0.

429

0.51

8

0.77

7 0.

774

0.77

0 0.

788

0.77

2 0.

366

0.05

7 0.

058

0.10

2 0.

074

0.03

1 0.

118


culated for both batch and fractional melting according to equations given by e.g. Frey et al. ( 1985). D values have been chosen from Frey et al. (1985), i.e. D A,Z03 = 0.1, D,, = 0.32. The calculation suggests that at least 25% and about 35% melting should have been necessary to produce whole-rock chemistry analogous to the harzburgite MC 55 and the dunite MC 50, respectively. Similar degrees of melting have been estimated by Duncan and Green (1987) for strongly depleted harzburgites belonging to supra-subduction zone ophiolites which, as the Nan Uttaradit harzburgite, yet contain traces of high-Ca cpx. A comparison of the selected trace element contents in harzburgite and dunite with the calculated ones show that the derived melting degree is incompatible with refractory modal assemblages such as harzburgite and dunite. Melting accompanied by melt percolation can however produce such a depletion. Finally, the strongest argument against a simple melting model is the nearly constant Mg# between harzburgites and dunites. Mantle melting produces typical sets of positive correlations between Mg#‘s of mafic phases (olivine, opx, cpx), Cr#‘s of spine1 and modal proportions of olivine (see e.g. Dick and Bullen, 1984). Similar correlations have been experimentally reproduced (Jaques and Green, 1980) and plotted in Fig. 13. It is apparent that the Mg#‘s of the NU dunites are too low to have been produced solely by partial melting even if melting had occurred at temperatures close to the solidus of a spine1 lherzolite.

I”

94'

92.

90.

88-

86.

84.

82.

Mg/Mg+Fe (rocks)

40 50 60 Fo (;:.%,

80 90 100

Fig. 13. Whole rock Mg# in peridotites as a function of olivine modal proportions. Melting curves at 1200” and 1600°C after Jaques and Green ( 1980)) star: olivine composition.

Dick (1977) was among the first to point out that Mg-ratios tend to decrease in the most refractory dunites from the Josephine ophiolite peridotites. Sim- ilar decreasing Fo contents irrespective of the concentration of the other strongly compatible elements (Ni, Cr) were also reported by Berger and Vannier ( 1984) and Bodinier (1988) in dunites coexisting with lherzolites. In both cases, the authors call for a percolation of previously depleted dunites by less refractory melts. Nevertheless, this interpretation is unable to explain the depletion event predating the percolation process. An alternative model was first suggested by Quick ( 198 1) and further refined by Kelemen ( 1990). Kelemen et al. ( 1990a, b, 1992) and Reimaidi et al. ( 1993) interprete the dunites starting from a harzburgite protolith per- meated by silicate melts out of equilibrium with the harzburgite and dissolving its pyroxene component. The concept of this model is that most basaltic liquids which are initially orthopyroxene-saturated at depths, become progressively opx undersaturated in the shal- lower upper mantle. This is due to a shrink of the stability field of orthopyroxene in the systemMg,SiO,- Si02-CaMgSi206 at decreasing pressure (see Stolper, 1980; Fig. 14). By moving upward through a mantle column, these melts will tend to dissolve the pyroxene component of the transected peridotites, mainly opx, leading to formation of olivine-rich rocks. In that model, the mg number of the melt is buffered by the wall-rock harzburgites through olivine-liquid equilibrium. In the case where the amounts of the magma infiltrated and extracted are nearly constant, the mg number of mafic phases and whole rock remain con-

stant (Kelemen et al., 1990a, b). The Kelemen et al.

( 1990a, b) model would therefore be the only one that can reconcile both the strongly refractory compositions

of the NU dunites and their relatively low and constant

Mg#. There are abundant microtextural criteria suggesting

that the NU harzburgites and dunites were soaked by

silicate melts at a time of their history. Recrystallization

of ameboidal spine1 into euhedral spine1 has been

observed both in natural samples percolated by hydrous

melts (e.g. Lorand and Cottin, 1987) and in supra solidus experimental deformations of mantle lherzolites in presence of a capillary silicate melt (Bussod and

Christie, 1991). The latter authors have emphasized that the capillary melt facilitated the rotation of olivine

crystals during deformations and also chemical

174 B. Orberger et al. / Lithos 3.5 (1995) 153-182

CaMgSizO6

/ MgzSiOd En

CaMgSirO6

/^\

SiOz

B /\

MgzSiOz En SiO2

Fig. 14. Pseudo-liquidus ternary phase diagrams in the system forsterite-diopside-silica showing the displacement of olivine-rtho- pyroxene and olivine/clinopyroxene as a function of pressure (A) and water pressure (B) (after Stolper, 1980 and Quick, 1981). The 1 atm cotectics of a primitive mid ocean ridge basalt were determined by Walker et al. ( 1979). The multiple saturation points at 10, 15 and 20 kb were determined by Stolper ( 1980) for basaltic melts in equilibrium with pyroxene and olivine. Orientations of the high pressure cotectics are estimated. B. Projection showing approximate location of 5 kb cotectics for a basaltic melt and liquid lines of descent for melt Ml that is fractionating olivine (solid line), reacting with plagioclase lherzolite wall rocks (dotted line) and reacting with harzburgite wall rocks (dash-dot line). M 1 is generated with harzburgite wall rocks (dash-dot). Ml is generated at 15 kb at the ol-opx-cpx multiple saturation point.

exchanges. In their experiments, the final texture of the peridotite is characterized by randomly orientated olivine crystals showing interlocked grain boundaries, with little evidence of 120°-triple junctions, as observed

in the NU harzburgite and dunites. Moreover, as emphasized by Kelemen ( 1990) and Bodinier et al. ( 1992)) subduction environments facilitate porous flow melt propagation within the mantle wedge above the Benioff zone, because of volatile release from the

subducted slab and inversion of the geothermal gradi-

ent. In this situation, volatiles immediately promote

first hydrous melting at depths and the hydrous melts,

encountering inverted geothermal gradients, can per-

colate large volume of mantle rocks. Since Kushiro

( 1969)) it is well known that increasing water pressure

shifts the ol-opx cotectic towards the SiOz apex of the

Fo-An-SiO, ternary so that dunites are expected to be

abundant at low P and high PHzO. To summarize, a

subduction-zone tectonic setting and the presence of

volatile now encapsulated as phlogopite inclusions in

chrome spine1 can explain generation of km-sized

refractory harzburgite and dunites bodies in the NU

ophiolite like in most ophiolitic complexes.

The origin of the opx-rich peridotites clearly

involved a silicate melt percolation multistage history.

No Post-Archean peridotites resulting from partial

melting have been reported to contain more than 32

wt.% opx (Kostopoulos, 1991; Kelemen et al., 1992),

even if the primitive mantle source is an orthopyroxene-

rich mantle such as that proposed by Sun and Mc-

Donough (1989). It can be suggested that the NU

orthopyroxene-rich peridotites were first enriched in

orthopyroxene because they are closely associated with

the oldest type I and II chromitite orebodies hosted

themselves in orthopyroxenites. Similar addition of an

orthopyroxenite component to a harzburgite residuum

through melt impregnation or mechanical mixing was

proposed by Edwards (1990) to explain orthopyrox-

enite formation in the Springer Hills peridotites from

the Bay of Islands ophiolitic complex. The mechanism

of the opx addition leading to the NU opx rich perido-

tites remains unclear. It might have involved precipi-

tation of opx from melts injected by hydrofracturing in

the wall rocks of type I and II chromitite orebodies.

Alternatively, percolation of a highly siliceous magma

(see below) within a peridotite may yield to the inverse

reaction

( 1)olivine + SiO,-rich melt I

+ opx + SiO,-poorer melt 2

discussed in detail by Kelemen et al., ( 1992). Textural and modal relationships between orthopy-

roxene and olivine in the opx-rich peridotites clearly

point out that, once formed, euhedral shaped OPX II reacted to form secondary, fine-grained olivine, suggesting a change in the composition of percolating magma, decreasing P,O, or increasing water pressure to

il. Orberger et al. / Lithos 35 (1995) 153-182 175

invert reaction ( 1) . This reaction has to be linked with the general process of dunite formation involving hydrous melts which was discussed before.

The late cpx-enrichment trend is clearly related to percolation of the CaO-rich melt generation which gave rise to the intrusive clinopyroxenite body. This is demonstrated by the sharp increase of cpx modal proportion at the immediate vicinity of the intrusive clinopyroxenite body, the textural similarities between the “impregnation’ ’ clinopyroxene in the peridotites and those in the clinopyroxenites and the high Sc/TiO, ratios of the cpx-rich peridotites. The instability of olivine and opx in the peridotites faced with the melt parent to the intrusive clinopyroxenite is consistent with the late appearance of these phases in the clinopyroxenite body. As noticed before, the formation of abundant clinopyroxene in the wehrlite BNT 103 resulted in a sharp decrease of Mg#‘s down to 82. According to the Kelemen ( 1990)‘s theoretical modelling, Mg#‘s drop when the quantity of magma extracted from a percolated peridotite becomes much smaller that the entering flux. This situation is likely to occur when the peridotites are much colder than the percolating melt, provoking abundant crystallization of this latter. The formation of cpx-rich peridotites would have therefore took place at a relatively late stage of the NU ophiolite history, when the peridotite-percolating melt system was frozen out.

7.2. Petrogenesis of chromitites

Type I and II chromitite as early cumulates from boninites

In addition to their compositions, type I chromitites hosted in orthopyroxenites have all characteristics of early crystalline segregates from olivine, or bronzite, boninites. The latter precipitates olivine and chromite first above 1300°C but olivine is rapidly out of equilibrium with the melt being remplaced by enstatite which crystallize at 1270°C (see for example Cameron et al., 1979; Howard and Stolper, 1981; Bloomer and Haw- kins, 1987; Peck and Keays, 1990; Thy and Xeno- phontos, 1991). A wealth of experimental work exists demonstrating that Cr-spine1 is a highly sensitive petrogenetic indicator as regards Cr,03, Al,03, FeO,, FeO, MgO contents of the melts from which they segregated (Hill and Roeder, 1974; Fisk and Bence, 1980; Maurel and Maurel, 1982; Murck and Campbell, 1986;

Allan et al., 1988; Roeder and Reynolds, 1991). Com- positions of melts that were in equilibrium with type I and II chromitites orebodies have been reconstructed in Table 6 using partition coefficient of Maurel and Maurel 1982; Allan et al., 1988 and Roeder and Reyn- olds, 1991. The calculation has been done from the Mae Charim chromitite lens that has both the highest Mg# and the highest Cr# and, thus likely segregated from the most primitive magma. In addition, CaO, Ti, V, SC and Cr that are pivotal to characterize boninites, have been estimated from Dppx-“q partition coefficients applied to orthopyroxenite sample PB6. The results show strong similarities with olivine boninites with regard to Cr/Cr+Al, TiOJV, TiOJSc, TiOJCaO and Ti02/A1,03 ratios as well as contents of each of the afore-mentionned elements. FeO/MgO ratios estimated from the Mae Charim chromitite are much higher than those of olivine boninites; this may result from slight modifications of FeO/MgO in the chromitites by subsolidus re-equilibrations with host silicates. Like- wise, Cr contents estimated from orthopyroxenites are notably underestimated, probably because orthopyroxene had crystallized from a melt impoverished in chromium by prior chromite crystallization.

The unusually high Cr,O, contents of some NU type I chromitite is near or above the stoichiometric 67.9 wt.% necessary to completely fill the octaedral sites of chromite. The study of meteorite and diamond chromite inclusions suggests that above this threshold, tetrae- drally coordinated divalent chromium should be present (Bunch and Olstein, 1975) requiring strongly reducing conditions, at least 3-4 log unit below FMQ (Roeder and Reynolds, 1991). Such highly reducing conditions are supported by Fe”+ /Fe3+ + Cr + Al ratios close to 0 and presence of graphite matrix and graphite inclusions in the Mae Charim chromitites. It is known from Schreiber and Haskins ( 1976), Murck and Campbell (1986) and Roeder and Reynolds ( 1991) experiments that the solubility of chromium in basaltic melts increases with decreasingfl, because of the greater solubility of Cr*+ relative to C? + . Con- versely, it is only C$’ that controls Cr203 contents of chromite. Partition coefficient DE!m”“‘“q. are raised by high silica and alkali contents which cause polymeri- zation of the melt and thus decrease the number of octaedral sites available for Cr) +, This is depicted by Irvine ( 1976)‘s phase diagram in the system Fo-An- Si02 showing that a melt saturated with respect to opx

176 B. Orberger et al. /Lithos 35 (I995) 153-182

coexists with a chromite with Cr# = 0.87. To summa-

rize, the peculiar chemistry of Mae Charim type I

chromitites reflects both the boninite composition of

their parent melt and the strongly reduced conditions

having controlled their crystallization.

The decreasingCr#‘s make type II chromitites likely

fractional crystallization products of the residual liquid

left after type I chromitite crystallized. Likewise, the

change from single phase olivine inclusion to phlogo-

pite-bearing polyphase inclusions points to an increas-

ing water activity in type II chromitites that may be

accounted for by massive precipitation of type I chrom-

itite along with anhydrous silicates from the parent

boninitic melt. Actually, mineralogical variations in

wall-rock and silicate matrix as well as in chromite

compositions clearly distinguish the Pak Nai chromitite

orebodies from the other NU type I and II chromitite

orebodies. The Fe3+ /Fe”+ + Cr + Al ratios (up to 0.1)

of the former reflect crystallization under higher oxy-

gen fugacity, close to the FMQ reference buffer curve

if we compare with experimental results of Fisk and

Bence ( 1980) and Roeder and Reynolds ( 1991) .

Moreover, the olivine-dominated silicate matrix of the

Pak Nai type I and II chromitites can be interpreted in

term of higher water activity in their parent melt, which

results in a larger number of hydrous silicate enclosed

within chromites. Increasing water activity drastically

increases the stability field of olivine at the expense of

opx in the system Fo-En-SiO, so that even boninitic

like highly siliceous melts can precipitate olivine

instead opx at high uH20 (Fig. 14). Since all of the Pak

Nai chromitite orebodies (types I, II and III) display

similar signs of high@* and high water activity, the

latter are most likely inherited from the mantle source

of the parent magmas. It is commonly assumed that

boninites are generated from strongly refractory mantle

metasomatised by Ba, Sr-rich hydrous fluids released

from the downgoing subducted slab (e.g. Hickey and

Frey, 1982). We can speculate that the melt parent to

the Pak Nai chromitites derives from a mantle source

metasomatized by oxidised fluids whereas the Mae

Charim and Rae Nan chromitites melts originated in a

source having incorporated carbon-rich, reduced fluids.

Unfortunately, the dismembered character of the NU

ophiolite precludes further discussion of these hypoth-

eses.

Type Ill chromitites and host clinopyroxenites: Early cumulates from transitional boninites.

Although the associated type III chromitites do not have typical boninitic spine1 compositions, the intru-

sive clinopyroxenites have mineral compositions and

whole-rock chemistry, especially their high SiO?/

A&O3 and CaO/Al,O, ratios, their Mg# coupled with

low A&O3 and TiO, that point to a boninitic affinity in

the relatively broad sense of Bloomer and Hawkins

( 1987). However, the NU intrusive clinopyroxenites

have cumulate textural and chemical features. It is thus necessary to estimate the compositions of coexisting

liquids prior to any comparison with extrusive bonini-

tes. The calculations have been done from both the Pak

Nai amphibole-free monomineralic clinopyroxenite samples BNS 82 and 83 and the type III chromitite.

This latter was chosen because it shows the highest

Mg# and was therefore less subject to re-equilibration

with their host clinopyroxenites. Selected values of par-

tition coefficients and calculated liquids compositions are listed in Table 5. The agreement between CaO-rich

transitional boninite compositions and ours is fairly good, especially for inter-element ratios that are discri-

minant for boninitic compositions (e.g. TiO,/V; TiO,/

Zr; TiOJCaO and Ti02/A1,03 ratios) as well as TiO,

and A120, contents. CaO is slightly overestimated in

our calculation, probably because of the poor knowl- edge of the DCaO partition coefficient. In addition to

their boninitic affinities, our calculated liquids show

strong similarities with some compositions of low-Ti basalts forming the upper pillow lavas of ophiolitic

complexes (e.g. Troodos; Table 6). This similarity

strongly supports to the hypothesis that intrusive cli-

nopyroxenite-wehrlite bodies crosscutting the top of the mantle sequence and the gabbros in ophiolites are

likely high-pressure cumulates of the upper pillow

lavas. Type III chromitites display the highest Fe”+ /

Fe3+ + Cr + Al ratios, like the type I and II Pak Nai

chromitite, indicating crystallization under rather high ~0, (around FMQ). Extrusive boninitic magmas are rich in CaO and A&O3 and thus precipitate abundant clinopyroxene and generally contain plagioclase phe- nocrysts (transitional boninites, boninitic dacites; Bloomer and Hawkins, 1987). The absence of plagioclase in the Pak Nai intrusive clinopyroxenite indicates that, like type I and II chromitites from the same area, their crystallization was controlled by high water activ-


ity. Addition of water to the system forsterite-anorthite-530, + chromium has been experimentally proven to stabilize chrome spine1 and clinopyroxene, instead of olivine + plagioclase (see Nicholson and Mathez, 199 1 and references therein). Since cpx has a higher Cr/Al ratio than the coexisting melt, continuous fractionation of clinopyroxene tends to decrease Cr/Al in the silicate melt and consequently also in chromite (Roeder and Reynolds, 199 1) . The strong decrease in Cr/Al ratios of Pak Nai type III chromitites is consistent with massive fractionation of cpx in the large clinopyroxenite intrusive body. In the Rae Nan type III chromite, however, high Cr/Al ratios are due to the late appearance of clinopyroxene in the crystallization order.

As suggested by type III chromitite compositions and demonstrated by calculated melt compositions, the intrusive clinopyroxenites have crystallized from less refractory melts than the type I and II orebodies. There is no structural or mineralogical argument suggesting

that they could be cumulates separated by a continuous fractional crystallization process from a single magma. Orthopyroxene and olivine appear late in the intrusive clinopyroxenites, whereas they are early cumulus phases in type I and II chromitite orebodies. It was not possible to obtain melt compositions in equilibrium with type III +chromitite by fractionation of olivine+chromite from the parental liquid to type I chromitites by quantitative modelling. However, it appears that type III orebodies formed from multiple injections of transitional boninitic melts having different mantle sources as parental liquids to type I and II. Theoretical modelling (e.g. Kelemen et al., 1990) suggests that the intrusive wehrlitic-clinopyroxenite suites could be formed from ascending melt propagating slowly by porous flow and dissolving the clinopyroxene component of a lherzolitic matrix. According to field relationships between the different generations of chromitites and the chronology of events discussed previously, orthopyroxene-enrichment seems to predate

Table 5

Calculated composition of melts in equilibrium with intrusive clinopyroxenites

DCpx/liq. BNS 82 BNS83 type III chromitite Troodos UPL Transitional Boninites

(Ohnenstetter Mariana Forearc

et al., 1990) (Bloomer and Hawkins, 1987)

(wt.%) SiOz I’ 51.82 51.11 47.1-50.4 55.22-58.12

TiO, 0.38-0.47’~’ 0.21-0.26 0.30-0.32 0.2-0.3“ O.lW.28 0.38-0.40

A&O? 0.17’= 6.7-9.6 IO-14 1 1.6-12.0h 7.8-11.0 14.45-14.35

Fe0 11.4 9.78 8.5 6.07-5.5 1 WO 13.5 9.05 17.5’ 8.7-13.4 7.76-6.99

cao I .5-J 6-’ 12.1-12.9 11.9 8-10.2 8.44-7.03

NazO 0.08’ 1.87 3.75 0.4-3.54 2.6-3.03

wm Ni

Cr

V

SC

Zr

4.58 45 55 134-93

3.8’ 510 460 40881465 313-222

0.8* 143 192 193-247 160-143

I .33-l .6*= 37-46 40-49 na na

o.123 <33 <33 IO-12 60-8 1

FeO/MgO 0.25-’ 0.84 1.08 0.5-l .o 0.79

Cr/Cr+Al+Fej+ 0.00&u3.007 o.oo55Wl.019 0.0070

Ti/Zr 45-56 69-84 95-198 45-35

Ti/V 10-9.5 11.8-14 8-5.9 16.8-19.8

TiOz/A120, 0.021-0.038 0.021-0.032 0.0145~.0035 0.026.0.0278 TiO,/CaO 0.0162-0.021 0.027 0.035-0.016 0.045-0.057

i ‘Bender et al. ( 1978); * J.L. Bodinier et al. (1987); ’ Hart and Dunn ( 1993); ’ Cawthom and Biggar (1993) @?m’“‘iq estimated from the

highly siliceous and magnesian basalt CC 17 1 atjO* = FMQ and 1283°C; 5 Thy and Xenophontos ( 199 1) ; 6 Maurel and Maurel ( 1982) ; ’ Allan

et al. (1988); * Bloomerand Hawkins (1987); 9Roederand Reynolds (1991) Dc~cr’“‘““Wme” assumingft), = FMQ at 130°C; na = not analyzed.

178

Table 6

B. Orberger et al. /Lithos 35 (1995) 153-182

Calculated melt compositions in equilibrium with Mae Charim type I chromitites compared to boninitic compositions

from Pb6 from Mae Charim olivine andesitic concentration range

chromitites boninites boninites of type B boninites

Bloomer Bloomer (Crawford and Cameron, 1985)

Hawkins (1987) Hawkins ( 1987)

wt.% TiOz

AN, Fe0

MgQ CaO

ppm Cr

V

SC

FeO/MgO Cr/Cr+ Al

Ti/V

TilSc

Ti02/A1203

TiO,/CaO

0.11' 0.09

0.247’ 4

7h 450

0.35’ 83

0.5’ 20

0.26” 0.87

7.7

32

0.013

0.025

0.07-0.15* 0.160.22

7.51 10.0-I 1.2 <7# 7.0-9.0

> g4 13-18

4.3-7.5

1700’ 1055-1300

120-180

nd

< 0.925 0.3-0.6 0.63 0.67-1.15

0.025(0.01)7 0.02 0.011 0.025-0.0047

6.3-13.0 9 7.3-19

0.014-0.022

0.16 0.17-0.26 0.13-0.32

11.6 11.6-13.34 6.613.7

7.0 6.0-7.5 8.0-l 1.4

4.97 4.0-8.0 4.3-5.6

722 338-548

126 95-164

0.014 0.013

0.032 0.02-0.06

100-1345

12.5-196

24-36

0.73-1.6

0.037-0.0026

5.4-9.4

27-54

0.02

0.02-0.05

/ ‘J.L. Bodinier, pers. commun. to J.P. Lorand; ’ Cawthom and Biggar (1993) @pm ‘e-r@ estimated from the highly siliceous and magnesian

basalt GC 171 atjO,=FMQ and 1283’C; ’ Maurel and Maurel (1982); 4 Allan et al. (1988); ’ Thy and Xenophontos (1991); ‘Bloomer and

Hawkins ( 1987); ’ Roeder and Reynolds ( 199 1); assuming logpI = FMQ-3-4 log units; nd: not detected

clinopyroxene formation in the peridotites. This situa-

tion would be at odds with observations made in the

upper pillow lavas of other ophiolitic complexes where

lavas precipitating opx before cpx are always younger

than those precipitating early cpx (Flower and Levine,

1987; Ohnenstetter et al., 1990; Thy and Xenophontos,

1991).

It is worth noting that, like in most ophiolitic com-

plex studied so far, the Pak Nai intrusive clinopyrox-

enites crosscut the boundary of mantle and layered

cumulates. This suggests that boninitic-magmas may

pond at the base of the crust in small magma chambers

(Benn and Laurent, 1987). BCdard (1993) recently

proposed that the intrusion of hot ( 1300°C), water-rich siliceous magmas into the gabbroic cumulates may yield in part to its assimilations, thus to a TiOz and

Na,O addition. High water pressure in the Ca-rich boninitic magmas would have yielded to preferential dissolution of the plagioclase component from the

gabbros. This is the most likely explanation for the MgO and Cr-impoverishement and Al,O, and TiO,- enrichment trends observed in the clinopyroxenites at

the immediate contact with the gabbros (e.g. sample

BNT 10.5).

8. Conclusions

Petrographic and geochemical data show that the

Nan-Uttaradit ultramafic sequence and its various

chromitite deposits represent the mantle part of an

ophiolite. In spite of the poor quality of the exposures and the complex history of this ophiolite, including

amphibolite-facies metamorphism, hydrothermal alteration, tectonic dismembering and late-stage supergene

weathering, most characteristics of supra-subduction zone ophiolites are preserved (see Pearce et al., 1984; Elthon, 1991; Pearce, 1991).

The protolith of the mantle section is strongly refractory, being mainly composed of harzburgites and dunites. The peridotites were formed by multistage reactions with percolating melts rather than by repeated events of partial melting. Three main reactions have been recognized ( 1) opx enrichments from true oli-


vine-boninite, (2) destabilization of opx into olivine

and (3) late-stage crystallization of HFSE depleted

clinopyroxene from the transitional boninitic melt par-

ent to the intrusive clinopyroxenites.

On the basis of their silicate matrix and their chem-

ical compositions, the chromitites have two different

parental magma compositions, although both enter the

category of high Mg, low-Ti siliceous basaltic that

characterize island-arc settings. The oldest one, parent

to type I and II chromitites, was a boninite as demon-

strated by the common association between type I-II

chromite unusually rich in chromium and early crys-

tallizing orthopyroxene + olivine. The second one was

a CaO-rich transitional boninite.

Both generations of magmas provide indirect evi-

dence of fluids released from the subducted slab. Type

I chromitites has crystallized under highly reducing conditions (FMQ-3-4 log units), probably in relations

to C-H-rich fluids, as demonstrated by numerous

graphite occurrences in the Mae Charim lens. The three generations of Pak Nai chromitites have differentiated

under conditions of higher oxygen fugacity ( = FMQ) and water activity because of phlogopite enclosed in

chromitites, abundant crystallization of oli-

vine + clinopyroxene and the absence of primary plagioclase even in intrusive clinopyroxenites.

Acknowledgement

This work has been financially supported by the

Deutsche Forschungsgemeinschaft (DFG) , the National Science Foundation of Thailand. (NRCT)

and the European Economic Community-( project SC*-CT 91-006). We are grateful to P. Wathanakul

who initiated the project, to Mr. A. Fancy and H. Gair

(Hancig Ltd., Bangkok) for the hospitility and the per- mission to work in their exploration area and to Hancig geologists (S. Jeenawut, S. Pramaphan, K. Pattamak-

eaw, U. Chaisomboon, G. Angkaew, A. Wongyai, S. Premmanee) for their help during field work. Our thanks are extended to the Department of Chiang Mai

University for sample preparation, to G Hamm (URA CNRS 736) for preparing polished thin sections, to G. Friedrich and the Institut fur Mineralogie und Lager- stattenlehre RWTH, to R.R. Keays and the School of Earth Science at the Melbourne University, Australia for XRF- and microprobe analyses. The largest part of

the work has been done when Beate Orberger was guest

researcher at the Laboratoire de Petrologic Physique,

UA CNRS 1093. This paper was improved through discussions with J. Marcoux, Y. Panjasawatwong and

critical reviews by J.L. Bodinier and an anonymous

reviewer.

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