The ophiolite of northern Oman

The Ophiolite of Northern Oman Frontispiece. The northern Oman mountains as imaged by the Multispectral Scanner (MSS)...

Author: S. J. Lippard | A. W. Shelton | Ian Graham Gass

104 downloads 1182 Views 56MB Size Report

This content was uploaded by our users and we assume good faith they have the permission to share this book. If you own the copyright to this book and it is wrongfully on our website, we offer a simple DMCA procedure to remove your content from our site. Start by pressing the button below!

Report copyright / DMCA form

DOWNLOAD PDF

The Ophiolite of Northern Oman

Frontispiece. The northern Oman mountains as imaged by the Multispectral Scanner (MSS) carried by the Landsat satellite, using near infrared light (band 7). Mosaic made from parts of two MSS scenes: path 170 row 44 and path 171 row 43. The ophiolite forms the dark terrain occupying most of ~he mountain-belt. The darkest, most rugged unit is the mantle sequence; crustal rocks are slightly paler. The palest units are sediments with autochthonous Mesozoic limestones to the west, and the Batinah coastal plain sediments to the east, of the ophiolite. For reference to ophiolite block names and geographic localities refer to the left-hand inset map on the 1:250,000 geological map that accompanies this memoir.

The Ophiolite of Northern Oman S.J. L I P P A R D , A.W. S H E L T O N & I.G. GASS Department of Earth Sciences, The Open University, Milton Keynes, MK7 6AA, UK

Memoir No 11

1986 Published for The Geological Society by Blackwell Scientific Publications Oxford London Edinburgh Boston PaloAlto Melbourne

Published for The Geological Society by Blackwell Scientific Publications Editorial offices: Osney Mead, Oxford, OX2 0EL 8 John Street, London, WC1N 2ES 23 Ainslie Place, Edinburgh, EH3 6AJ 52 Beacon Street, Boston, Massachusetts 02108~ USA 667 Lytton Avenue, Palo Alto, California 94301, USA 107 Barry Street, Carlton, Victoria 3053, Australia 9 1986 The Geological Society. Authorization to photocopy items for internal or personal use, or the internal or personal use of specific clients, is granted by The Geological Society for libraries and other users registered with the Copyright Clearance Center (CCC) Transactional Reporting Service, provided that a base fee of $02.00 per copy is paid directly to CCC, 21 Congress Street, Salem, MA 01970 USA, 0305-8719/84 $02.00 First published 1986 Printed and bound in Great Britain by William Clowes Limited, Beccles and London

DISTRIBUTORS

USA and Canada Blackwell Scientific Publications Inc PO Box 50009, Palo Alto, California 94303 Australia Blackwell Scientific Publications (Australia) Pty Ltd 107 Barry Street, Carlton, Victoria 3053

British Library Cataloguing in Publication Data Lippard, S.J. The ophiolite of Northern Oman. 1. Rocks, Igneous--Oman I. Title II. Shelton, A.W. Ill. Gass, I.G. Society of London 552'. 1'095353 QE461

IV. Geological

ISBN 0-632-01587-X Library of Congress Cataloging-in-Publication Data Lippard, S.J. The ophiolite of northern Oman. (Memoirs/Geological Society: no. 11) Bibliography: p. Includes index. 1. Geoiogy--Oman. 2. Ophiolites--Oman. 1. Shelton, A.W. II. Gass, I.G. (lan Graham) III. Title. IV. Series: Memoirs (Geological Society of London): no. 11. QE291.05L56 1 9 8 6 555.3'53 86-11737 ISBN (,L632-01587-X

This memoir is dedicated to the memory of the late D R I S M A E L EL BOUSHI

former geological adviser to the Sultan of Oman, without whose help and whole-hearted co-operation this research project would not have been possible.

Contents

Preface

ix

The Geological Background 1.1 Introduction 1 1.2 Regional Tectonic Setting 2 1.2.1 The Zagros 2 1.2.2 The Makran 3 1.2.3 The Gulf of Oman 4 1.3 Tethyan Ophiolites 4 1.4 History of Geological Investigations 7 1.5 Tectonostratigraphy of the Oman Mountains 9 1.5.1 The Basement Rocks of the Arabian platform 1.5.2 The Allochthonous Units 12 1.5.3 Late Cretaceous Nappe Emplacement 14 1.5.4 Neoautochthonous Sediments and Tertiary Structures 16

11

Evolution of the Oman Tethys

17 2.1 Introduction 17 2.2 Rifting and Continental Break-up 18 2.2.1 Permian Sediments 18 2.2.2 Triassic Sediments 20 2.2.3 The Haybi Volcanics; Triassic Volcanism at the Opening of the Oman Tethys 22 2.3 The Oman Passive Margin 31 2.3.1 Shelf and Slope Facies Sediments 31 2.3.2 Rise and Basin Facies Sediments; the Middle and Upper Hawasina 32 2.3.3 Late Mesozoic Alkaline Igneous Activity 34 2.4 Mid-late Cretaceous Syn-tectonic Events on the Continental Margin - a prelude to late Cretaceous Nappe Emplacement 36 2.4.1 The Aruma Group 36 2.4.2 The Hawasina Melange 36 2.5 The Oman Tethys - a Summary 37

The Semail Ophiolite

39

3.1 Introduction 39 3.1.1 Internal Subdivisions 40 3.1.2 External Relations of the Semail Nappe 41 3.2 The Mantle Sequence 43 3.2.1 Harzburgites 46 3.2.1.1 The Basal Lherzolite 48 3.2.1.2 Geothermometry and Geobarometry 51 3.2.2 Dunites and Chromitites 52 3.2.2.1 Chromite Mineralization 55 3.2.3 Mantle Dykes 55 3.2.4 Structure 59 3.2.5 Alteration 62 3.2.5.1 Alteration Products 62 3.2.5.2 Physical and Chemical Effects of Serpentinization 63 3.2.6 Summary 64 3.3 The Petrological Moho 65 3.4 The Layered Series 66 3.4.1 The Layered Series of the Semail Ophiolite 68 3.4.1.1 Structure 68 3.4.1.2 Petrology and Mineralogy 72 3.4.1.3 Geochemistry 76

3.4.2 Large Scale Cyclic Units, Crystallization Sequences and Cryptic Variations 76 3.4.3 Magma Chamber Processes 78 3.4.3.1 Origin of the Cyclic Units 78 3.4.3.2 Magma Chamber Size and Shape 78 3.4.4 Fractionation Trends and Primary Magma Compositions 79 3.5 The High-Level Intrusives 79 3.5.1 Field Relations 79 3.5.2 Petrology 81 3.5.3 Geochemistry 82 3.5.4 Petrogenesis 86 3.6 Late Intrusive Complexes 86 3.6.1 Large Gabbro-diorite-plagiogranite bodies 87 3.6.2 Peridotite-gabbro Complexes 90 3.6.3 Geochemistry and Petrogenesis 91 3.7 The Sheeted Dyke Complex 94 3.7.1 Field Relations and Structure 95 3.7.2 Dyke Trends 100 3.7.3 Petrography and Mineralogy 100 3.7.4 Geochemistry 101 3.8 The Extrusive Sequence 103 3.8.1 Volcanic Stratigraphy 103 3.8.2 Geochemistry 107 3.8.3 Tectonic Setting 113 3.8.4 Clinopyroxene Compositions 115 3.8.5 Rare-earth Elements 116 3.9 Petrogenesis 117 3.9.1 Convective Processes 120 3.9.2 Partial Melting Processes 120 3.9.3 Magma Fractionation 123 3.10 Ocean-floor Metamorphism 124 3.10.1 Metamorphic Facies: Nomenclature and Conditions 124 3.10.2 Ocean-floor Metamorphism in the Semail Ophiolite 124 3.10.3 Summary 127 3.10.4 Massive Sulphide Deposits 127 3.11 Metalliferous and Pelagic Sediments 128 3.11.1 Fauna and Age 128 3.11.2 Stratigraphy and Field Relations 128 3.11.3 Geochemistry of the Pelagic Sediments 132 3.12 Isotopic and Magnetic Studies 133 3.12.1 Isotopic Studies 133 3.12.2 Magnetic Studies 135

Ophiolite Detachment, Emplacement and Subsequent Deformation 140 4.1 Introduction 140 4.2 Detachment 140 4.2.1 The Metamorphic Sheet 140 4.2.2 Peridotite Banded Unit 148 4.3 Emplacement and Related Processes 149 4.3.1 Metamorphism 149 4.3.2 Igneous Activity Associated with Emplacement (Biotite Granites) 151 4.3.3 Structural Evolution 153

References

167

Preface

This memoir is based on studies by Open University (OU) and associated personnel between 1975-85 and represents a fullsome precis, correlation and evaluation of work presented in eleven Ph.D. theses and numerous scientific publications. In total, it represents some 48 man-years of effort. The names of persons associated with the research project at various times, together with their status and affiliation, are listed in the table below. Early in the reign of Sultan Quaboos the Oman Government began to welcome overseas scientists. Until then, only geologists of the PDO (Petroleum Development Oman) a part-Oman Government part-Shell International Company, had been allowed to work in the north of the Sultanate. Work by PDO scientists, particularly the classic study of the Oman mountains by K.W. Glennie and his coileagues (Glennie et al. 1974), although primarily concerned with sedimentary formations, provided a firm foundation for this ophiolite-orientated project. Our own association with the Oman began in 1975 when one of us (IGG), on a visit to King Abdulaziz University (KAU) in Jeddah, Saudi Arabia, discussed with Dr A.O. Nasseef, PresiOpen University

dent of that University and now Secretary General of the Muslim World League, the scientific desirability of investigating the Oman Ophiolite and its potential as a post-graduate training ground. Dr Nasseef contacted Dr El Boushi, then Geological Advisor to the Sultan of Oman, and as a result Gass & Dr Abdurrazzak Bakor (KAU), together with Drs J.D. Smewing and A.D. Lewis (OU Research Fellows) visited the Sultanate later that year. It took only a few days to realize that the Oman mountains are a magnificent ophiolite, similar in many respects to the then better known Troodos massif of Cyprus. All present were convinced that it was a research area of the highest potential. It was envisaged that the area would be used as a training ground for Saudi, Omani and British graduate students who would work under the supervision of Dr I.M. El Boushi, and academic staff from the Open and King Abdulaziz Universities. Geological maps on the scale of 1:100,000 would be produced under an informal agreement with no contractual obligations on any of the parties. A grant of s from the Royal Society of London paved the way for a detailed feasibility study of the ophiolite in 1976 81

82

83

84

85

86

ACADEMIC STAFF

I.G. Gass J.A. Pearce A.J. Fleet, R.M. Shackleton RESEARCH FELLOWS

J.D. Smewing (OU) K.O. Simonian (OU) A.D. Lewis (OU) S.J. Lippard (OU) C.R. Neary (BGS: ODA; chromites) D.I.J. Mallick (BGS; ODA; remote sensing) R.B. Evans (BGS, ODA; geophysics) RESEARCHSTUDENTS:(abbreviated thesis titles and date D.T. Aidiss (NERC) M.A. Brown (ODA) G.M. Graham (Shell) M.P. Searle (OU) T. Alabaster (ODP) A.W. Shelton (ODA) D.A. Rothery (ODA) P. Browning (NERC) G. Stranger (Private) S. Roberts (EEC) I.D. Bartholomew (NERC)

(field leader 1976-7~

II

(field leader 197%81)

)resentanon given m parenthesis) ] (Granitic rocks of ophiolites: 1978) (Chromite studies: 1983) (AIIochthonous sediments. Hawasina Windo'~v: 19811) (AIIochthonous ttavbi complex and metamorphic sole: 19811) - - (Volcanic and hvdrothermal processes: 1982) ~ (Gravity and palaeomagnetic studies: 19841 (Remote sensing: 1982) (Plutonic rocks in Rustaq area: 1982) ~ (Groundwatcr hydrogcochemistrv: 1986) (Chromite metallogenesis: 1986) - - (Mantle sequence structures: 1983)

ASSOCIATED SCIENTISTS

I.M. Elboushi (Oman Government) N.I. Christengen (Seattle) J. Malpas (Memorial) A.H.F. Robertson (Edinburgh) A.G. Smith (Cambridge) N.H. Woodcock (Cambridge) D. Elliot (John's Hopkins) P. Tippitt (Texas)

(Mineralization) I (Seismic and related studies) (Field studies) I I [ [ [(Sedimentology) (Structure) ] I ] --% [ (Sedimentology and structure) (Structure) I I ] I (Radiolarian dating) 9

Preface by J.D. Sinewing, A.D. Lewis, K.O. Simonian and D.T. Aldiss. This fully confirmed that high quality geological studies could be made in this well exposed and relatively undeformed terrain. When it became likely that no KAU personnel would be further involved, Dr El Boushi agreed that the OU group should proceed alone. Subsequent to the initial Royal Society grant, the Natural Environmental Research Council (NERC) provided, over the period 1976-81, a research grant of s as well as three research studentships. However, by far the largest support (s 1977-82) came from the British Government's Ministry of Overseas Development (now Overseas Development Administration: ODA). Through the good offices of Dr David Bleakley, then the Director of the Overseas Division of the British Institute of Geological Sciences (now BGS: British Geological Survey) and Professor E. Machens, the project was put to the OEDC (Organization for Economic Cooperation and Development) whose enthusiastic support resulted in the ODA grant. These monies funded four research students and three BGS specialists and allowed the production of four 1:100,000 coloured geological maps as well as the 1:250,000 geological and gravity maps that accompany this memoir. Another major contributor was the Open University which funded 15 man-years of Research Fellowship and one research student. Shell and the European Economic Community (EEC) each funded a research studentship whilst the American Oil Company (AMOCO) funded a Research Fellow for two years. We are most grateful to all these funding agencies for their support. Throughout the project we have received invaluable assistance from numerous groups and individuals. In Oman, we are particularly indebted to the Department of Petroleum & Minerals (DPM) whose Director, Mr Mohammed Kassim, sponsored the project and provided the necessary documentation for project personnel to work in the Sultanate. We also acknowledge with many thanks the provision, by the DPM, of a house in Sohar that acted as the project's base for many

years. We were particularly fortunate that between 1977-79, our peak period of field work, we had the company in Oman of a University of California - United States Geological Survey research team who were making a detailed study of the ophiolite along a 25 km wide N-S strip extending south from Muscat. We benefited greatly from discussions with our American colleagues as we did from the PDO geologists who gave freely of their time and expertise. But, perhaps most of all, we will remember the unfailing and uniquely generous hospitality of the Omanis whose mountains we invaded. On the more practical side, the RAF and SOAF (Sultans Own Air Force) provided invaluable assistance in air lifting two landrovers to the Oman, and bringing back to the UK several tons of rock specimens. At the OU we received the fullest support and cooperation from the Department of Earth Sciences" technical and secretarial staff. We are most grateful to John Holbrook, Ian Chaplin, Andy Tindle & John Watson who provided prompt and efficient curating, specimen preparation and analytical services, to Helen Boxall who prepared all the text figures in this memoir and to John Taylor for the 1:250,000 geological and gravity maps that accompany it and to Pare Owen and Carol Whale who provided a first class secretarial service that saw the manuscript through numerous stages. The principle objective of this memoir is to present the major findings of the project under one cover as a coherent whole. We are greatly indebted to our former colleagues for allowing us to use their data in this way and also for commenting on and correcting those parts of the manuscript based on their findings. We are particularly grateful to the following who checked and improved earlier versions of the manuscript, Julian Pearce, Robert Shackleton, John Sinewing, Chris Neary, Tony Alabaster, Dave Rothery, Paul Browning, Gordon Stanger, Steve Roberts, lain Bartholomew, Micky Brown, John Malpas, Aiastair Robertson, Alan Smith & Martin Menzies.

Chapter 1 The Geological Background of which lies between 500 and 1500 m, is bare and rocky with little or no vegetative cover. Although some wadis have perennial streams, most are dry except during severe winter storms when they are subject to flash flooding. East of the central part of the mountains there is a narrow coastal plain (the Batinah), underlain by recent sands and gravels and coastal deposits; whereas inland to the west and south are the desert wastes of the interior of Arabia, known as the Empty Quarter or Rub al Khali. Most of the mountain area is in the Sultanate of Oman, but between Wadis Hatta and Dibba an eastward extension of the United Arab Emirates (U.A.E.) separates the Musandam

1.1 I n t r o d u c t i o n The Oman Mountains, also known as the Hajar or Eastern Hajar (Hajar al Gharbi) range, lie at the eastern extremity of the Arabian sub-continent and run in a broad arc parallel to the Gulf of Oman coastline (Fig. 1.1). They extend from Ras Musandam, facing the Straits of Hormuz in the north, to Ras al Hadd in the east, a distance of nearly 800 km. The mountains have an average width of about 75 km, reaching a maximum of 130 km in the central part, and rise to a height of about 3000 m on Jebel Akhdar. The rugged mountain terrain, most 1

56 ~E

.~

N

o~

I

I,-~4~~4~/~~ MTt~s

~

[

~.o,~ Qo; ~ ", MAKRAN

\

,,,

RLJus al Jibal (2087 m)

26 ~N I.

/

~-"--~--J/

~)_GULF

ARABIA

rv"

,>

p,,,.~., "

//

W/k -n---, i

jl -j

::s

/'-' ~i Khor

1500-2000 m

~.O

I

Iloo0-15o0m

1

J 500-1000 m

[

] Land b e l o w 500 m Main c o m m u n i c a t i o n routes t h r o u g h the m o u n t a i n s International b o u n d a r y Major t o w n s Main m o u n t a i n peaks

.....

ohar _X

'~~'q!~,

INDIAN OCEAN

Land over 2000 m

\

/Buraimi",,

.sz

"4

Fakkan ~ ~13 Fujairah ~

z~

/

F~

,f

g

~ .f ! / ,4,1

OF OMAN

.

.~,

--,

I

",...1

0~4%~4L

9

o

4~

r4 uvs

@ ,.,\\ t e /

Seeb P

I a

i

n

,,

i '--Jebel Ka~ ~,,(2960 , ~

, 5"

"x \_\ 50 km

el b ra

'!".! Y 58~ I

Fig. 1.1 Physiographic map of the Oman Mountains showing major towns and communication routes. I

%

Chapter 1 oceanic collision boundary along the Zagros and Makran fold and thrust belts. Northern Oman has greater geological similarities with the Zagros ranges of southern and central Iran than with the Makran region which lies directly across the Gulf of Oman (Fig. 1.3). The continuity of structure from Zagros to Oman was noted by several early workers on the Alpine system, including Suess (1909), who showed Oman as an outer loop of the Zagros system that connected back towards the Himalayas through the northeast-southwest trending mountain ranges of southeast Pakistan. Ricou (1971) noted the similarity in geological setting and the age of ophiolite nappes thrust over the northern margin of Arabia in a crescent-shaped belt from Cyprus and Turkey along the Zagros thrust zone to Oman which he called "le croissant ophiolitique peri-Arabe'.

Peninsula from the rest of Oman. The most important communication routes crossing the mountains are along Wadis Ham, Hatta and Jizi in the U.A.E. and northern Oman and through the Semail "Gap" southwest of Muscat (Fig. 1.1). The Oman Mountains are geologically distinct from the rest of the Arabian Peninsula and are part of the Alpine-Himalayan fold belt that extends from the western Mediterranean to the Far East. They were deformed during two major orogenic events in the late Cretaceous and the mid-Tertiary. The first of these was dominated by the emplacement of nappes of Mesozoic continental margin and ocean floor rocks from northeast to southwest onto the Arabian continental margin. The second was marked by folding on largely upright axial planes, although there was some local thrust reactivation. The present topographic eminence of the mountains is the result of postMiocene uplift. The geology of the mountains is dominated by the Semail Nappe, a huge sheet of mid-Cretaceous ophiolitic rocks which is thrust over allochthonous units of marine sediments and volcanic rocks representing the remains of a Tethyan continental margin that lay formerly to the northeast of the present mountain area. The whole allochthon is emplaced over the basement rocks of the Arabian shield; these are composed of an Upper Proterozoic Pan-African crystalline basement overlain by late Precambrian to Cretaceous continental to shallow marine sediments that are exposed in the cores of Tertiary anticlines along the axis of the range.

1.2.1 The Zagros The Zagros fault zone is some 1600 km long and extends from the Turkey-Iran border to just north of the Straits of Hormuz (Fig. 1.3). Within the presently active fault belt, there is a series of nappes of shallow to deep-water Mesozoic sediments, including radiolarites with large "exotics" or olistoliths of shallow-water limestones, and ophiolite masses which were thrust southeastward over the autochthonous rocks of the Arabian platform forming the folded Zagros ranges to the south (Stocklin 1968, 1974; Hallam 1976; Ricou 1976; Ricou et al. 1977; Stoneley 1980). Ricou (1971) recognized three nappe units in the Neyriz area above the autochthonous succession: the Pichegun Nappe of Triassic to mid Cretaceous shallow to deep water sediments; the Melange Nappe with exotic blocks of Permian-Triassic limestones, radiolarites, lavas, metamorphics and serpentinite, and the uppermost Neyriz Ophiolite Nappe. All these units are transgressed by Rudist limestones of Maastrichtian age in turn overlain by Palaeogene sediments that are cut by the main Zagros Fault. Evidence for shallow marine to continental sedimentation in the Maastrichtian to early Tertiary, immediately following nappe emplacement, led some

1.2 Regional Tectonic Setting The Oman Mountains lie on the northeastern edge of the Arabian Plate (Fig. 1.2) which is bounded to the south and southwest by the active spreading axes of the Gulf of Aden and Red Sea, to the east and west by transcurrent fault zones of the Owen Fracture Zone and the Jordan or Dead Sea Rift, and to the north by a complex continent-continent to continent-

;~

2~3~

-

-

-

-

-

40 ~

6}) ~

810~

100~,E

-

40 ~ N

t -. ~ :.::::.:i::: r......... " 5 . . :. ,.? j.:.: :..,., .: : ~f>, " " " .:.

-

....

??

....

~,

~t,

i,, ,,- O M A N

:ii

MOUNTAINS

9

"""" " ~ l:"b'C': ~,\... :."..:::::::"~-:::.2 ": k , Alpine-Himalayan Fold Belts k'~ A A ~,

Active Subduction Zones (Aegean & Makran) Quaternary volcanoes

of Aden <

Fig. 1.2. Tectonic map of the Alpine-Himalayan fold belt showing the position of the Arabian plate and the Oman Mountains. Major mountain ranges: 1. Betic Cordillera; 2. Pyrenees; 3. Atlas: 4. Alps: 5. Apennines: 6. Dinarides: 7. Carpathians: 8. Hellenides: 9. Pontides: 10. Taurides: 11. Caucasus; 12. Zagros; 13. Alborz" 14. Hindukush: 15. Karakorum: 16. Himalaya: 17. Kunlun" 18. Indo-Burma Ranges.

The Geological B a c k g r o u n d I 60 ~E

Central Iran

/

Helmand Block

,7% 9

.DZ

~

Qatar Arch

]

~ ~ 9 ,

9 "

~

GULF OF

9

,

"

0 MA N .-4... ~ .-.4,._ ~

/~%',~ \ 0

;'. :..f"

~

-\ \

.-~.-........-......--..-.-.-....~-.L ', (+% ...-- v--- . ~ ' 7 .-7-. ~.

~

~:L-- ~--- ~

, ~ ...~.-.vV .....v.-v.-.vV.'V.V :v. v. )...-7..

.".. .. .": "-,, @ c"sk .

:.v:.':-.-:-.v.v.v.v.v?.,7..7.-7.-7.'7-7.-7.7.'):.-.).v.v v ".~'x,,.'... "s= ' \

' " ' J ' / ? / 5 . . .5 s

?. s .'5/s s

?/?J/? s

/

5 J5 " J 5 / J / J - ' J ~ ,

?:: :: :: :: .': :.-):):7:7:7:::.-::v: :.. :.-7:-: .v .vl ':L-7:)'.%tx v..v .v .'. ..'-.-. .. .'. .. Rub ai Khali ..v .-.-..-.v .v .v ..-..- ..-...\

))])[[:):[i]':[i)]:])[]:)i[17[i])'i:):: :: ... :: : ..:.: : i[i.iiii'iil)il])]:i)iil)::i[':[[)]]:))?.l

.i :i.v7--i.'.'7.'7.'i-t 84 9 .'jj'j

iiiiiiiiiii!!iiiii!iiiiiiii!!

I'1~

/

Masirah Is. ~

I

/

Faults. lineaments

/ o~o~

/

/

~Gulf. of/Aden / -~! , / o , ~ o ~

/

0

Tertiary fold axes

.....

Limit of Tertiary Z a g r o s - O m a n Folding

d~-iZ~)

Inland depressions Southern limit of Makran folds

_.A._

Limit of Oman mountain thrust belt

c} s~'/

,%/ /

/

300kin I

.i.vT..y

:!iiiiii!i!iiiiii!!iiiiiiii!ii!iiii!iiiii ~ v.....v.i...~y

20ON --

~

57 J "?[/"

Zagros thrust zone Submarine ridges and scarps

/

/

/

/ '~,~ / /

,~

~ ' Sheba/Carlsberg ridge/transform system

CFZ DZ OFZ ZF ,o

Chaman fault zone Dibba zone Owen fracture zone Zendan fault zone Quaternary Volcanoes H o r m u z Salt domes

Fig. 1.3. Tectonic map of eastern Arabia and adjoining regions. workers (Takin 1972; Stocklin 1974; Berberian & King 1981) to suggest that continental collision occurred along the Zagros suture in the late Cretaceous, whereas others (Dewey et al. 1973; Sengor & Kidd 1979) argue that the collision of Arabia and Central Iran did not take place until the mid to late Tertiary. To the north of the Zagros fault zone, the central part of Iran consists of a belt of continental margin "Cordilleran-type" calc-alkaline plutonic and volcanic complexes with some remnants of older Precambrian and Palaeozoic basement (Berberian & King 1981). The magmatic activity spans most of the Mesozoic and Cenozoic (Alavi 1980, Berberian et al. 1982) suggesting subduction beneath the area throughout this period. Central Iran is just one of a number of small continen-

tal "microplates" that lie between the Eurasian and AfroArabian plates (Dewey et al. 1973). Palaeomagnetic studies (Krumsiek 1976) and geological considerations (Crawford 1972; Falcon 1974; Stocklin 1974) support the view that these continental masses formerly belonged to part of the southern supercontinent of Gondwanaland. Their separation apparently occurred in the Permo-Triassic (see Chapter II, Section 2.2 for an account of the break-up of the Oman continental margin), but details of the subsequent movements are unclear owing to their complex Mesozoic-Cenozoic histories. 1.2.2 The Makran

The Makran is the coastal region of southeast Iran which lies

4

Chapter

directly across the Gulf from Oman (Fig. 1.3). It is bounded to the north by the Jaz Murian depression at the southern end of the Lut block, which is mainly underlain by Precambrian basement. The inner, northern part of the Makran consists of a complex zone of ophiolites and pelagic sediments, ranging in age from Jurassic to Palaeocene and overlain unconformably by Eocene limestones (McCall & Kidd 1982). These rocks are faulted to the south against a ridge of folded Palaeozoic, Mesozoic and early Cenozoic platform-type sediments, that are in turn thrust southwards over Eocene to early Miocene flysch that underlies a 200 km wide belt in central and southern Makran. The northern edge of the flysch belt is marked by a narrow zone of Maastrichtian to Palaeocene ophiolitic melange and the flysch deposits themselves occur in southward younging thrust units (McCall & Kidd 1982). These rocks are unconformably overlain by flat-lying to mildly folded Miocene and Pliocene shallow marine sediments that are interpreted by McCall & Kidd (1982) as the deposits of Neogene forearc basins. The Makran and its southern continuation beneath the Gulf of Oman has been interpreted as a broad Cenozoic accretionary prism formed above a shallow, northwarddipping present-day subduction zone (Farhoudi & Karig 1977; Jacob & Quittmeyer 1979). The trace of the present day subduction zone can be followed offshore where flat-lying sediments of the floor of the Gulf of Oman are deformed into east-west trending folds (White & Ross 1979; White 1982). Four to five hundred kilometres north of the Makran coast there is a chain of Quaternary andesite volcanoes which is interpreted as a volcanic arc related to the subduction (Farhoudi & Karig 1977). Cordilleran-type magmatic activity in this region extends back to the late Cretaceous (Arthurton et al. 1982; Berberian et al. 1982) and it seems likely that subduction of the Gulf of Oman oceanic crust beneath the Makran has been more or less continuous since that time. 1.2.3 The Gulf of Oman

The floor of the Gulf of Oman lies at oceanic depths of 30003400 m in the deepest part off the northeast coast of Oman where the continental margin is particularly steep, reaching 3000 m depth only 30 km offshore. The floor of the Gulf is heavily blanketed by sediments, 4-6 km thick, but heat-flow and depth data (Hutchinson et al. 1981) indicate that it is oceanic and probably at least 70 Ma old. West of longitude 58~ the Gulf shallows and the outer parts of the Arabian and Iranian continental margins are in contact. White & Ross (1979) interpret this line, approximately along the Zendan Fault to the north, as marking the boundary between continental collision (Zagros zone) to the west and oceanic subduction (Makran zone) to the east. The variations along the collision zone of the Zagros-Oman-Makran system can be explained by irregularities in the shapes of the continental margins, particularly the "Oman Embayment" (Section 2.5), on the northern edge of the Arabian plate, which allowed the Gulf of Oman to remain an unconsumed oceanic area whilst Cenozoic continent-continent collision and underthrusting occurred along the Zagros thrust zone. The onland boundary between the Zagros and Makran zones is a complex north-south trending fault zone, known locally as the Zendan Fault (Shearman 1976). The fault shows evidence of both westward overthrusting and right-lateral strike-slip movement and some workers have projected it across the Gulf of Oman to link up with the NE trending Dibba Zone at the northern end of the Oman Mountains (Fig.

1

1.3). Several authors, among them Gansser (1955) and Falcon (1967), have called this lineament the "Oman" or "Dibba Line". However, the Dibba zone is a complex part of the Oman Mountains thrust front dominated by low-angle thrusts (Lippard et al. 1982) and there is no evidence for major rightlateral strike-slip movement as depicted on some regional tectonic maps (e.g. Murris 1980). Another NE-SW trending major strike-slip fault has often been postulated along the "Masirah Line", parallel to the straight southeastern coast of Oman (Fig. 1.3). Here, Masirah Island and the headlands of Ras al Madrakah and Ras Jibsch are composed of late Mesozoic ophiolitic rocks overlain unconformably by Eocene limestones (Moseley 1969; Glennie et al. 1974). These outcrops are separated from the mainland by a NE trending fault zone occupied by ophiolitic melange (Moseley & Abbots 1979). Although originally considered a strike-faulted portion of the Semail ophiolite (Glennie et al. 1974), giving the Masirah fault zone a right-lateral displacement, it now appears more probable that the Masirah ophiolite represents part of the floor of the Owen Basin (Fig. 1.3) obducted during left-lateral movement along a complex boundary between the Arabian and Indian plates (A. C. Ries pers. c o m m . 1983).

L3 Tethyan Ophiolites The Semail Nappe of the Oman Mountains is just one member of the long chain of Tethyan ophiolites that extends from the western Mediterranean to the Far East (Fig. 1.4). The presence of these complexes along the length of the Alpine-Himalayan mountain belts has long been recognized as one of the keys to understanding Tethyan evolution since their regional development was first recognized by Steinmann (1927), Dubertret (1953) and Brunn (1960). These early workers interpreted the ophiolites as gigantic submarine eruptions in geosynclinal troughs that formed at the early stages of the development of the Tethyan basins. Trumpy (1960) suggested that they marked a transition from an extensional to a compressive regime in the late Jurassic and early Cretaceous. Subsequent studies in which individual complexes have been characterized and dated, have shown both a diachroneity and periodicity in ophiolite formation and emplacement along the belt (Table 1.1 and references therein), suggesting the opening and closing of several ocean basins in the Mesozoic (Smith 1971). Several authors, e.g. Nicolas & Jackson (1972), Mesorian et al. (1973), Rocci et al. (1975) and Abbate et al. (1976), have divided the ophiolites of the Mediterranean region into two groups: Jurassic complexes in the western and central area (Alps-Apennines-Carpathians-Dinarides-Hellenides) and Cretaceous complexes in the eastern Mediterranean (CyprusTurkey-Syria). Petrological and geochemical studies have shown that there are important differences along the Tethyan ophiolite belt. The Alpine and Apennine (including Corsica) ophiolites have MORB-type chemistry (Bickle & Pearce 1975; Ferrara et al. 1976; Beccaluva et al. 1977), dominantly lherzolitic ultramafics and are characterized in the upper levels by serpentinite breccias and ophicalcites (mixed carbonate-serpentinite rocks). They lack sheeted dyke complexes and have been interpreted as thin oceanic crust and underlying mantle possibly formed in a fracture-zone setting (Barrett & Spooner 1977). Lherzolite ultramafics are predominant in the western ophiolite belt in Yugoslavia and Albania whereas in the eastern belt in Yugo-

The Geological Background 35 ~

55 ~ E

BlacklSea-~~ 9b

9,s " v " -

~

~Cn~ [ aspia

~12a

,| ".9d 9 e h _ o l l a P

MediSea terranean

35 ~ N

16a

~' ~16b\

~

4'

Ophiolitic suture zones

25 ~

.q

Major fault zones

4

"~qbx'~ Omon

I F1

t16d

~7

#d~# Ophiolites ....

o

,J/' 14a 14b

,,

.

----

~

40 ~ N

10 E~/-X

12" \

,,t16

,ff (

\~._

20 ~

20 ~

40 ~

6

o

80~

Fig. 1.4. T e t h y a n ophiolites; d a t a on n u m b e r e d occurrences given in Table 1.1.

slavia and northern Greece they are largely harzburgites (Pamic & Majer 1977; Karamata et al. 1980). This dichotomy of the Jurassic ophiolites is continued into central Greece where the well-known Vourinos complex is harzburgitic (Moores 1969; Ross et al. 1980) but the Pindos and Othris ophiolites are lherzolitic (Menzies & Allen 1974). It has been suggested on the basis of trace element and isotopic data, that the Vourinos complex formed in an island-arc setting (Noiret et al. 1981; Beccaluva et al. 1984) and Pindos and Othris in an associated marginal or back-arc basin (Capredi et al. 1980), although any interpretation is complicated because of the presence of both island-arc and ocean floor volcanic rocks types in these complexes (J. A. Pearce pers. c o m m . 1983). Further east, in northern Turkey, the Anatolian and Pontide belts contain Jurassic and Cretaceous ophiolites as dismem-

bered fragments in melanges (Bergonnan 1975), whereas the southern Tauride ophiolites are larger and more intact masses of mid to late Cretaceous age (Juteau et al. 1973; Juteau 1980). They are mainly composed of tectonized "mantle sequence" harzburgites with well-developed layered cumulate sequences in some areas, e.g. Antalya (Juteau & Whitechurch 1980). Sheeted dykes and lavas are lacking but the ultramafics are cut by tholeiitic dyke swarms dated at 85 Ma (Juteau 1980). The Troodos massif on Cyprus was early recognized as a late Cretaceous (probably Campanian age) ophiolite (Gass 1968, 1980; Moores & Vine 1971). It contains a complete sequence from near the top of the mantle sequence to the pelagic sediments and has been the subject of numerous detailed investigations. It lies at the western extension of a 3500 km long chain of mid to late Cretaceous ophiolites that extends

Chapter 1 Table 1.1.

Country~Region

Individual complexes

Formaliotl age"

1. Alps

Montgenevre, Pre-Upper Jurassic Monviso, Haute-Ubaye, Overlain by UJ-KL Lanzo, Zermatt-Sass sediments

2. A p e n n i n e s Corsica

Inzecca, Bologna, Mt. Maggiorasca, Voltri area

3. Calabria 4. Carpathians 5. Dinarides 6. Hellenides

As Apennines J? Western belt-Zlatibor Eastern belt Vourinos, Pindos, Othris, Guevgueli

7. Pontides, Anatolides 8. Lesser Caucasus

LJ (160-185 Ma m)

L J? LJ ?

J-K?

Obductionh/ emplacement' age

Occur in Pennine nappes. Involved in late Mesozoic-Cenozoic Alpine deformation and metamorphism. High pressure metamorphism overprints sea-floor assemblages. Lherzolites. MORB chemistry of dykes and lavas I-5 Complex geology, associated with Post-KL (Overlain by olistostromes and melanges. conformable UJ-KL Serpentinite breccias and sediments ophicalcites. Highly-faulted oceanic ridge <'s. Lherzolites. MORB chemistry 7-~. Similar to Apennines Little known, probably an extension of the Alpine zone. Western belt lherzolites, eastern 160-182 Ma ~3. Pre-Upperbelt harzburgites 1112 Jurassic Vourinos well studied 14,t5 160-172 Ma 22 Pre-Upperharzburgites ~ and probably arcJurassic related volcanics ~vq~Othris and Pindos more lherzolites2~, possibly marginal basin. Numerous small complexes 23, mainly in melange zones. Some high pressure metamorphism. Poorly known 2~>. Most probably an extension of the Anatolian belt. Largely harzburgites 2~>. Layered 86-104 Ma (Metamorphics sequence well exposed by cut by Campanian (75 Ma) Antalya >. No sheeted dykes or dykes) lavas. Metamorphic soles and mantle sequences cut by tholeiitic dykes. Intensely studied 31-~4. Recent work 87-89 Ma 37 Pre-Campanian on iavas suggests a supraformation subduction zone setting~s36. 86-96 Ma z~ Post-Campanian Complete sequences exposed. Harzburgitic 3s-4" Island-arc tholeiite - Pre-Maastrichtian chemistry~. emplacement

"'Pre-Cenomanian'"

10. Cyprus

Vedi zone, SevanAkara (a) Marmoris (b) Mersin-Kersanti (c) Antalya (d) Posanti (e) Beyshin (f) Hoyran Troodos

11. NW SyriaSE Turkey

(a) Hatay (b) Baer-Bassit

KM?

12. Zagros

(a) Khoy (b) Kermanshah (c) Neyriz

KM?

89? Ma ~ Post CampanianPre-Maastrichtian emplacement

13. Oman mts

Semail (Samail)

Cenomanian (98-94 Ma)

14. SE Oman

(a) Masirah Island (b) Ras Madraka

K?

96-85 Ma Post Campanian Pre-Maastrichtian emplacement Pre-Eocene, probably late Cretaceous emplacement

15. Makran

(a) Makustan-Fanuj (b) Iranshah

KU?

Pre-Palaeocene emplacement

16. East & Central Iran

(a) Nair (b) Baft (c) Esfandagheh (d) Tchehel-Kureh Sabzevar

KU?

Pre-Palaeocene emplacement

K?

Pre-Palaeocene emplacement

9. Taurides

17. North Iran

KM?

KM-KU? (79 Ma? 33)

Remarks

Ophiolites strung out along Zagros fault zone (2000 km long) 43. Harzburgite at Neyriz cut by low K tholeiite dykes 4344. Island-arc tholeiite lavas in other complexes 45. Harzburgitic mantle sequence. Island-arc tholeiite chemistry (see Chap. IlI for details). Harzburgites, MORB? chemistry47. Related to NE-SW "transform' faulting along Masirah Line 4s4~. Probably a part of the Owen Basin obducted from the SE. Late Cretaceous back-arc marginal basin 455~ Harzburgites. MORBIAT chemistries. Mainly harzburgites 51. Small ocean basins around Lut block. MORB and IAT chemistries45. Harzburgite 5253. Small ocean basin between Lut and Asian block. IAT? Chemistry.

The

Country~Region

Individual Complexes

18. West Pakistan Afghanistan

Bela, Muslim-Bagh, KU? Zhob, Kabul-Waziristan

Geological

Formation age"

Background

7

Obductionb/ emplacement c age

Remarks

Harzburgite 54. Obducted to SE during collision of Afghan and Indian plates 555~. Harzburgitic Spontang and Post-Campanian- PreJungbwa ophiolites are klippen to Eocene emplacement the south of the suture zone -~7. Lherzolitic Xigaze in the Tsangpo suture has a thin crustal sequence 5s and MORB chemistry. Pre-Cretaceous emplacement Jurassic suture zone 59. Little known but probably Cretaceous Tethyan ophiolites. Harzburgitic s~.

Post-Maastrichtian - PreEocene emplacement

19. Himalayas Ladakh, Spontang, (Indus-Tsangpo) Jungbwa, Xigaze

K? (Albian fossils in Xigaze)

20. North Tibet Donquiac 21. Northeast India- Nagaland Burma

J?

K?

Notes

a "Formation age" refers to isotopic ages of igneous minerals (usually of doubtful validity except in cases of U-Pb Zircon ages) and the palaeontological age of interbedded or conformable sediments. b "Obduction ages" given in Ma in this column are isotopic ages of amphibolites in metamorphic soles taken to indicate the time of initial oceanic displacement. c "Emplacement age" is generally a minimum age of nappe emplacement bracketed by the age of the youngest rocks involved in the thrusting and the oldest cover sediments, usually continental or shallow marine beds. References

1. Bickle & Nisbet (1972) 2. Bickle & Pearce (1975) 3. Bodinieretal. (1981), 4. Lewis & Sinewing (1980) 5. Venturellietal. (1981) 6. Barrett & Spooner (1977) 7. Beccaluva et al. (1977) 8. Bessi & Piccardo (1971) 9. Ferraraetal. (1976) 10. Ohnenstetter etal. (1981) 11. Pamic & Majer (1977)

12. Karamata et al. (1980) 13. Lanphere et al. (1975) 14. Moores (1969) 15. Jackson etal. (1975) 16. Rossetal. (1980) 17. Capredi et al. (1980) 18. Noiret et al. (1981) 19. Beccaluvaetal. (1984) 20. Montignyetal. (1973) 21. Menzies & Allen (1974) 22. Spray & Roddick (1980) 23. Bergougnan (1975) 24. Adamia et al. (1981)

25. Knipper & Khain (1980) 26. Juteau et al. (198(I) 27. Juteau etal. (1977) 28. Juteau (1980) 29. Juteau & Whitechurch (1980) 30. Thuizatetal. (1981) 31. Gass (1968) 32. Moores & Vine (1971) 33. Desmet et al. (1978) 34. Desmons et al. (198{)) 35. Robinson etal. (1983) 36. Schmincke etal. (1983)

from the Hatay and Baer-Bassit complexes (Parrot 1973, 1974) in N W Syria and S. Turkey through eastern Turkey and into Iran along the Zagros thrust zone to the Semail Nappe in O m a n , that was named the "Croissant Ophiolitique PeriA r a b e " by Ricou (1971). As noted by Pearce (1980) and Desmons & Beccaluva (1983), most of the lavas in these complexes are geochemically akin to island-arc tholeiites. Although recent geochemical and istopic studies of Troodos lavas suggest an island-arc origin (Robinson et al. 1983), Pearce et al. (1984) propose that a supra-subduction setting is more appropriate for this complex in which there is no other evidence of an island-arc environment. Moores et al. (1984) have suggested that all the ophiolites of the "'croissant'" formed in a marginal basin-island arc complex setting similar to the modern A n d a m a n Sea. Masirah Island and other ophiolite fragments along the coast of SE O m a n appear to be distinctly different from the Semail ophiolite and may well be part of the floor of the Owen Basin emplaced by N E - S W faulting along the Masirah "line" in the late Cretaceous to Palaeocene (Moseley & Abbotts 1979; A. C. Ries & R. M. Shackleton pers. c o m m . ) . The ophiolites of eastern Iran, W. Pakistan and Afghanistan occur

37. Spray & Roddick ( 1981 ) 38. Parrot (1973) 39. Parrot (1977) 40. Tinkleretal. (1981) 41. Delaloye & Wagner (1984) 42. Ricou (1968, 1971) 43. Pamicetal. (1979) 44. Arvin (1982) 45. Desmons & Beccaluva (1983) 46. Adib (1978) 47. Abbotts (1981) 48. Moseley (1969)

49. Moseley & Abbotts (1979) 50. McCall & Kidd (1982) 51. Davoudzedeh (1972) 52. Alavi (1977) 53. Lensch et al. (1977) 54. Abbas & Ahmed (1979) 55. Asrarullaetal. (1979) 56. Moores et al. (1980) 57. Thakur (1981) 58. Nicolas et al. ( 1981) 59. Girardeau et al. (1984) 60. Ghose & Singh (1981)

in several belts and, where data are available, appear to have slightly younger obduction or emplacement ages (PalaeoceneEocene) than those of the "croissant". These complexes appear to be related to the opening and closing in the late Cretaceous-early Cenozoic of several small ocean basins, some of Red Sea-type, others of back-arc character (McCall & Kidd 1982; Desmons & Beccaluva 1983). Further east, along the Indus-Tsangpo suture zone in the Himalayas, there are ophiolites that were probably formed and emplaced in the mid to late Cretaceous (Nicolas et al. 1981). North of the suture in northern Tibet there is an earlier, probably Jurassic, suture zone containing ophiolites (Girardeau et al. 1984). Beyond the NE Himalayan syntaxis, the Nagaland ophiolites (Ghose & Singh 1981) of N E India and B u r m a continue the Tethyan ophiolite belt to its eastern extremity (Mitchell 1984).

1.4 History of Geological Investigations Politically secluded from the outside world, the O m a n Mountains have, until recently, remained relatively remote and isolated. In consequence, until the mid-1970s, geological

Chapter 1 studies were of a reconnaissance nature. Only during the past few years have more systematic and detailed investigations, of which this Memoir forms a part, been made. Early observations of the geology around Muscat were made by officers of the Indian Geological Survey, including Carter (1850), Blanford (1872) and Pilgrim (1908) who called all the pre-Tertiary rocks exposed in the mountains the "Oman Series". Lees (1928), who made surveys throughout most parts of the mountain area which were accessible at that time, divided the succession into seven major units (Table 1.2). He described the Semail Igneous Series as being composed largely of serpentinite, gabbro, diorite and lavas and was the first

person to appreciate their enormous extent. He proposed that the igneous rocks form a huge thrust sheet which he called the "Semail Nappe" and also recognized the allochthonous nature of the underlying Hawasina Series which, on seeing the highly deformed state of these sediments, he described as . . . "a yielding and incompetent member between two rigid units; the Musandam limestones below, and the massive igneous rocks of the Semail above" . . . (Lees 1928, p. 601). Lees also recognized the similarities between the Oman sequence and the "strongly folded shales, radiolarites and basic igneous rocks of the Neyriz area" in the southeastern Zagros Mountains of Iran.

Table 1.2. Historical development of the tectonostratigraphic classification of the Oman Mountains. ALLEMANN & PETERS (1972) (northern mountains)

LEES (1928) 7.

Maastrichtian to Eocene Limestones

6.

Semail Igneous Series

5.

("Permian Klippen") Hawasina Series

Simsima Fm. Qahlah Fm. "Neoautochthon"

Maastrichtian- Lower Tertiary cover seds. Batinah Complex

Semail Ophiolite Nappe

Semail Ophiolite Nappe

Semail Nappe

Metamorphic Series

Oman Melange ./ ("Oman Exotics") ./ ,-, Hawasina AIIochthonous Unit Sumeini Group (Parautochthonous)

Haybi Complex

Hawasina Complex

.,r

4. 3. 2.

Musandam Limestones Elphinstone Beds Permian Limestones

1.

Pre-Permian quartzites, limestones and phyllites

PRESENT WORK

GLENNIE et al. (1974)

Ruus al Jibal Unit (Parautochthonous)

Muti Fm ~ Aruma Gp. Hajar Supergroup (Autochthonous) Pre-Permian Basement -=Huqf, Haima Gps. Hatat Metamorphics* Jebel Ja'alan schists and gneisses (Precambrian)

Thrust contact

............... Unconformity

In contrast to the ideas put forward by Lees, later workers, notably Morton (1959), disputed the allochthonous nature of the Semail and Hawasina units and argued that they form an essentially conformable succession of Upper Cretaceous rocks above the Mesozoic limestones. Morton cited critical exposures in Wadi Mia'adin south of Jebel Akhdar where the contact between the Hawasina beds and the underlying midCretaceous platform carbonates is (in his words) "'sharp but concordant". He added that the Hawasina passes westward by a facies change into Upper Cretaceous marls of the foreland succession. Wilson (1969) recognized that the Hawasina sediments were disturbed by gravity slides, particularly near the top where huge masses of Permian and Triassic "Oman Exotics" (his term for those massive limestones formerly called the "Permian Klippen" by Lees) were interpreted as having slid into a deep water basin9 He also noted the presence of metamorphic rocks at the base of the Semail, but dismissed shearing at the contact as being too slight to have been caused by large thrust movements. All the authors who propounded the autochthonist viewpoint were strongly influenced by the then currently favoured geosynclinal concept (Stille 1924; Aubouin 1965) in which the "Steinmann Trinity" of serpentinite, pillow lava and charts was considered characteristic of the initial stages of geosynclinal evolution in which a deep water basin forms in a formerly stable, intracratonic area. In the late 1960s detailed biostratigraphic studies of the Hawasina sediments by Shell Oil Company Geologists (Haremboure & Horstink 1967; Glennie et al. 1973, 1974) showed ages ranging from Triassic to Cretaceous demonstrating that

J

Hawasina Assemblage ....-

Mainly as Glennie et al.

*Now considered late Cretaceous

the Hawasina must be separated at the base of the sequence by a major tectonic break from the underlying Upper Cretaceous marls and conglomerates of the Muti Formation, formerly assigned to the Hawasina by Morton (1959). Glennie et al. (op. cit.) reinterpreted the Hawasina as a stack of thrust slices of hemi-pelagic and pelagic sediments which they envisaged as having been deposited in an ocean basin that originally lay to the northeast of the present mountain area. The Permian and Triassic limestone Exotics, the highest member of Glennie et al.'s "Hawasina Allochthonous Unit", were found to have foundations of pillow lavas and interpreted as having formed on ocean islands within the "Hawasina Ocean". The same authors also recognized the "Oman Melange", between the Hawasina and the base of the Semail Nappe, as a mixture of sedimentary, volcanic and metamorphic rocks and serpentinite which was largely of tectonic o r i g i n . . . "a mechanical mixture that formed when the Oman Exotics and the Semail Nappe were emplaced onto the uppermost units of the Hawasina" 9. . (Glennie et al. 1974, p. 180). The marked shift of opinion back to the allochthonist ideas in the late 1960s and early 1970s was partly inspired by the new plate tectonic concept of ophiolites as fragments of oceanic lithosphere emplaced tectonically onland during plate collisions (Coleman 1971; Dewey & Bird 1971). B. M. Reinhardt, a member of the Shell team, was the first person to describe the Semail Nappe in modern terms as an ophiolite complex (Reinhardt 1969; Glennie et al. 1974). He recognized that the serpentinized peridotite, which forms a large part of the outcrop, is mainly harzburgitic with a relict high temperature

The Geological I

Background I

56 ~

58 ~ E

26ON -

MUSANDAM PENINSULA

] Recent deposits

DIBBA

~

Maastrichtian-Tertiary sediments

~

Semail Ophiolite nappe

~

Hawasina and Haybi allochthonous units and Batinah complex

~

Sumeini Group (Triassic to Cenomanian carbonates)

-W. HAM {~--~ _

-W. HATTA

SUM[

Hajar Supergroup (Middle Permian to Cenomanian carbonates)

_

~

Late Precambrian--Ordovician sediments

'~-~

Jebel Ja'alan schists and gneisses (Precambrian, 850 Ma)

'~,,,

Major anticlines

"~

High angle faults

JlZl 0

50kin

t

'~ SOHAR

,.J

AHIN WINDOW 24 ~

ASJ U D I'~::!~,..'.'~

\

-

" MUSCAT

t : : 7X:: t: ;:.i.i

~t[

::

"~>~.......~ ~:; ~...-. ,'~ "-~,:.. ~\ \~. \\

V.. ,... "<" %'.;. \.,.

RAS AL

~,~ :k\,, %.%

9'?k

~ ~ .& -..:2, f

~o ~

-..:&.%

r~ N:

f

: i

~,.<_..::...~. ..:...,:,,:~.

......................

,,~'.:.:.::::.)1 ",,, .:::,: :;,..%:::::./ ~

...%

HADD

.~__

/

JR'ALAN /~ . .

.

. : . ." . " . . " . 7

-'.'.'-5-,;

"~ " ~ " . . . : : : : : i

I

l

I

/

I

Fig. 1.5. Simplified Geological Map of the Oman Mountains (largely after Glennie et al. 1974 with additions from OU and USGS mapping 1975-80).

mineralogy and tectonite fabric. Reinhardt showed that the peridotites pass upwards through gabbros and diabase dykes into spilitic pillow lavas and he regarded this mafic suite as having formed at the axial rift of an oceanic ridge. An Upper Cretaceous age (Cenomanian to Coniacian) was established by the palaeontological dating of pelagic sediments within the lava sequence. The emplacement of the ophiolite nappe was regarded by Glennie et al. (1974) as the result of gravity sliding from an elevated mid-ocean ridge in the Campanian. A study of the northern part of the mountains by Allemann & Peters (1972) provided additional support for the ophiolite nappe concept and new information on the composition and ages, by radiometric dating, of igneous and metamorphic rocks beneath the ophiolite.

The most important contribution made by the Shell team was to provide a thorough understanding of the major geological relationships in the Oman Mountains which has formed a basis for all subsequent work. The two-volume Memoir (Glennie et al. 1974), published as a special report of the Royal Dutch Geological Society, was accompanied by two coloured geological maps of the whole mountain area at a scale of 1:500,000.

1.5 T e c t o n o s t r a t i g r a p h y o f the O m a n M o u n t a i n s Fig. 1.5 shows a simplified geological map of the Oman Mountains slightly modified after that of Glennie et al. (1974) by the

o

Chapter 1 AUTOCHTHON

ALLOCHTHON i~_Salahi

-

M1

Let/vole (U~) > *Batinah Sediment Sheets" i .:~__,:~qE:Exotic Barghah Fm(U'h -KL) ;

6O

-:_ =

~ m ~-I

San con'

Cen

[__-~:::1o:Fm oJ ~

I~

Central Mts

c

.

:i.! '.i

Musandam/

~

.~.

(Part Melange)

SuhaylahFmf "upniolitic beds" (Cenomanian- S a l a h i IInit ~, Santonian) i i "Extrusives" -I!

:~.~.~ Lasail Alley Unit i ~.;~L-~+~ Unit Geotimes Unit 'fHH!!~[!!I

>K F-~IQ-A - ~ / ~ ' ~ 1 N a p p e ~

Crop

BATINAH COMPLEX

~ ~: "Batinah Melange" (KU)

SIMSIMA Fm

Maa

.

Fm (LJ-KM) Sakhin Fm (UT~-L J)

,~i ~.

"Sheeted Dyke Complex"

9 ',' " "High Level Intrusives"

~ - ~ _ . ."~~

wa.s,~ G ~ : : ~

"Cumulate Sequence"

"Late Intrusive Complexes" occur w~lhin these units

~. *Crustal Sequence"

SEMAIL NAPPE

=-~:7~--:=--" Pet r olo gic a I Moho" Kahmah Gp ~

"Mantle

[-~-r Unit ,4 -~ : z

Sequence"

~: "Banded Unit" -'". '-- "Metamorphic Sheet" (KM-KU)

~.e

Ghalilah ~alilah Fm :~ ~:

uo

M ~

L

~

~-Fm

Huqf Gpl Kharus Frn ~ Late

Mi'aidin Fm <~ Hajir

- Cambrian?

l

[: ~ -[ Amdeh \ ~_--::_-:-~ :m \ } - : :- ~ L-M Ord ? [- = - :~ Halma Gp ' -:-'-;~

Mistal Fm

~F . . ~

Sil Carb

~ R a n n

JEBEL Fm

QAMAR (Or(I)

Zulla Fm (M-h-U-h )

J

SAIH HATAT JEBEL AKHDAR

) V U r ~ [ ~ , ~ : _ r ~ !AI "Hawasina Melange" ~ ~ v ! Sedimentary Melange /) Q, ii ii ~! (Olistostrome K M ? ) / / , ~ ' .: .. ~i Hawasina Seds - - - Exotic Lets (P, U~ ~ "- Nadan Fm (LJ) o::~::'.: - AI Aridh/Ibra Fm (U'~) Halfa/Haliw Em ('~-KL) :-~J:~ AI Ayn Fm (U'h-L J) 5~_-zS: I - - Wahrah Fm (LJ-KM) _Nayid Fm (KL-KM) Sidr Fm (UJ-KL) Guwayza Fm (U-h -L J) . Hamrat 9 Duru Gp.

nelsses schists granites ( 8 5 8 M a )

JEBEL JA'ALAN

~ ~.~

HAYBI COMPLEX

Serpentinite Melange (KU) Haybi Volcanics (M-U'~) 4- Exotic Let (P. U~ )

(Melange)

"] /

I

L Hawasina or -Hawasina Series

HAWASINA ASSEMBLAGE

~J

Qumayrah Fm (KU) ( Muti Fm) ~ Mayhah Fm ,LJ-KM) ~ Sumeini Gp. Maqam Fm (MTa -U-h) /Jebel Wasa Fm (Uh)

Fig. 1.6. Tectonostratigraphy of the Oman Mountains. Terminology largely from Glennie et al. (1974), except for Haybi Complex (Searle 1980), Semail Ophiolite (Smewing 1980, Alabaster et al. 1980, Fleet & Robertson 1981)) and Batinah Complex (Woodcock & Robertson 1982a) subdivisions. Time-scale based on Harland et al. (1982) (note scale change tit 1(10 Ma).

results of Open University (Smewing 1979; Lippard 1980; Lippard & Rothery 1981; Browning & Lippard 1982) and U.S.G.S. (Bailey 1981) mapping. The geology of the area may be divided into three major tectonostratigraphic units (Fig.

1.6): 1. An autochthonous* succession of Middle Permian to Upper Cretaceous carbonates resting unconformably on late Precambrian to Lower Palaeozoic sediments, representing part of the stable Arabian platform. 2. Allochthonous* units, composed mainly of Mesozoic rocks including the Semail ophiolite nappe, that were emplaced as a series of thrust nappes in the late Cretaceous. 3. A neoautochthonous* sequence of Maastrichtian and Lower Tertiary sediments. The subdivision and nomenclature of the Ha jar Supergroup, Sumeini Group and Hawasina units shown on Fig. 1.6 follow that of Glennie et al. (1974). The term "Oman Melange", for

* Glennie et al. (1974) described the rocks of Saih Hatat and Jebel Akhdar as "autochthonous" but recent work (S. Hanna pers. comm. 1983) suggests that they may have been thrust into their present position. They also describe the Sumeini Group carbonates as "'parautochthonous" but here they are included in the allochthonous units. "Neoautochthonous", as first used by Glennie et al. (1974), refers to the post-allochthon cover rocks.

the rocks between the top of the Hawasina and the base of the Semail Nappe, has been modified and redefined as the "Haybi Complex" (Searle 1980). The internal subdivision of the Semail ophiolite follows Smewing (1980a), Alabaster et al, (1980) and Fleet & Robertson (1980). An allochthonous unit of melange and Mesozoic sediments overlying the Semail Nappe has been recognized as the "Batinah Complex" by Woodcock & Robertson (1982a). 1.5.1 The basement rocks of the Arabian platform The autochthonous rocks of the Arabian platform are exposed in the central part of the Oman Mountains in the Jebel Akhdar and Saih Hatat anticlinal windows and in the north on the Musandam Peninsula (Fig. 1.5). The Musandam succession consists entirely of Permian to mid-Cretaceous carbonates whereas in the Jebel Akhdar and Saih Hatat areas erosion has breached the cover sequence to expose older pre-Permian rocks. 1.5.1.1 P r e - P e r m i a n o f E a s t e r n A r a b i a

A Precambrian crystalline basement is exposed at Jebel Ja'alan (Fig. 1.7) at the eastern end of the Oman Mountains and is inferred to underlie the whole of the eastern Arabian Peninsula. The Jebel Ja'alan gneisses and schists are cut by

The

Geological

intrusive granites dated by the Rb/Sr method at 858 + 16 Ma (Glennie et al. 1974). In Dhofar province, in southern Oman, the basement is exposed along the coast near Salalah and continues eastward to form the Kuria Muria Islands (Fig. 1.7) (Beydoun 1966; Hawkins et al. 1981; Gass et al. in prep. These small occurrences of mainly granitic gneisses and amphibolites in eastern Arabia appear to be an extension of the PanAfrican (1100-450 Ma) basement of the Arabian Shield that covers a large area of the sub-continent in western Saudi Arabia (Greenwood et al. 1976; Gass 1977). In Oman the crystalline basement is overlain by up to 3500 m of late Precambrian to Palaeozoic continental to shallow water marine sediments, comprising the Huqf (late Precambrian to Cambrian), Haima (Ordovician-Devonian) and Haushi (Carboniferous-Lower Permian) Groups (Tschopp 1967; Gorin et al. 1982). These are exposed in the south near Salalah, along the Indian Ocean coast in the NE-SW trending Huqf "uplift" and in the Oman Mountains, and are known from oil well logs, to underlie most of central Oman. In interior Oman the succession is most complete and includes glacial deposits of late Carboniferous to early Permian age (Braackman et al. 1982), whereas, in the Oman Mountains Ordovician rocks are unconformably overlain by Middle Permian carbonates (Lovelock et al. 1981). At the base of the pre-Permian succession in Jebel Akhdar the clastic Mistal Formation contains granite clasts, derived from the Pan-African basement, as dropstones in glaciogenic tillites and siltstones (Tschopp 1967; Glennie et al. 1974). These are overlain by interbedded black limestones and dolomites (Hajar Formation), a sequence of thinly bedded

54~

~Jebe~ba

_.///-J

60~

\\ Y~" ~. 24~ \\ Muscat \ \ \g'\'-/-.loJebel f--.4/ ~-R~5 h~AkT (l.~.~. A-"---_ i N S0, ." Hatat INTERIOR (boreholeOMAN data)

I HAIMAGROUP HUQFGROUP

JebelJa'alan~ /

Horrnuzsail Huqfuplift ,

0 1

I

II

siltstones and shales (Mi'aidin Formation) and then further carbonates (Kharus Formation). The total succession is about 1500 m thick and is undated but Gorin et al. (1982) correlate the four formations with the succession in the Huqf type area which would indicate a late Precambrian-Cambrian age. The finer grained sediments have been deformed into weakly cleaved phyllites and the structure is of open NE-SW trending folds that are truncated by the base of unconformably overlying Middle Permian dolomites. In the Saih Hatat area, some 50 to 100 km to the east of Jebel Akhdar, the pre-Permian sedimentary sequence comprises the Hijam (dolomite) and Amdeh (quartzite) Formations (Kapp & Llewellyn 1965; Tschopp 1967; Glennie et al. 1974). Lovelock et al. (1980) discovered a well-preserved mid-Ordovician fauna in the Amdeh Formation which suggests correlation with the lower part of the Haima Group of the interior. The 3400 m thick Amdeh sequence consists of current-bedded and ripplemarked quartz sandstones and siltstones containing abundant trace fossils indicative of shallow marine deposition. These Lower Palaeozoic sediments are unconformably overlain by Middle Permian dolomites on the southern and western margins of the Saih Hatat window, but to the north they are in thrust contact with a problematic group of phyllites, mica schists and greenschists which were originally assigned by Glennie et al. (1974) to part of the pre-Permian basement. Glennie et al. (op. cit.) reported a K-At whole rock age of 327 + 16 Ma for a metabasaltic greenstone in support of a Hercynian metamorphic event in this area. There is evidence, however, that these rocks were affected by a high pressure transitional blueschist-greenschist facies metamorphism (Boudiet & Michard 1981; Lippard 1983) and metamorphic micas in glaucophane schists from As Sifar in NE Saih Hatat have yielded mid to late Cretaceous K-Ar ages of 100-80 Ma (Lippard, op. cit. ). Thus, it is now considered more likely that the metamorphism in the northern part of the Saih Hatat area was contemporaneous with late Cretaceous thrusting and nappe emplacement (the relation of this metamorphic event to subduction and continental collision will be discussed in Section 4.3.1). A small occurrence of pre-Permian rocks on Jebel Qamar in the Dibba zone in the northern mountains consists of 240 m of Cruziana-bearing Ordovician quartzites (Rann Formation) overlain by a condensed sequence of Siluro-Devonian? shales and limestones that are overlain unconformably by Permian carbonates (Hudson et al. 1954a). The whole succession which, including the Permian and Triassic carbonates, is over 1500 m thick, occurs as an exotic block or olistolith in a Cretaceous melange (Allemann & Peters 1972; Searle et al. 1983). 1.5.1.2 Permian to Upper Cretaceous S h e l f carbonates

I

O

Background

19 ~

200km j

~xI~, fKURIAMURIA 9 oo 9 ISLANDS ~.-.., LateaPrecambri n[ -...[Pal eozoicsedimaents l Pan-AfricanBasement I

Fig. 1.7. Precambrian and Lower Palaeozoic rocks of eastern Arabia.

The Middle Permian to Upper Cretaceous (Cenomanian) Hajar Supergroup (Glennie et al. 1974) of eastern Arabia consists of massive marine carbonates which are exposed in the Musandam Peninsula and in the Jebel Akhdar and Saih Hatat areas (Hudson & Chatton 1959; Hudson 1960; Allemann & Peters 1972: Glennie et al. 1974: Ricateau & Richd 1980). They consist of up to 3000 m of shallow water sediments that represent part of a broad carbonate platform subjected to occasional epeiric movements throughout the Mesozoic with, in general, more open marine conditions to the north and east towards the open Tethys ocean (Murris 1980). The lower part of the succession consists of Permian and Triassic sabkha-cycle dolomites deposited in restricted conditions. In the early

I2

Chapter 1

Jurassic, after a period of emergence and renewed marine transgression, more open shelf conditions prevailed. The Lower Jurassic limestone contains interbedded quartz sandstones which represent the only major influx of clastic, terrigeneous material into the area throughout this period. The Middle and Upper Jurassic consists of shallow water limestones, but in the latest Jurassic, Upper Tithonian to Valanginian, deeper water hemi-pelagic lime mudstones were deposited above an erosional unconformity. By the Albian, shallow water conditions had returned and the Albian-Cenomanian is represented by a succession of marls and rudistid limestones (Wasia Group) that marks the end of Mesozoic carbonate shelf deposition. 1.5.1.3 The Aruma Group

Tectonic instability, represented by the Turonian "WasiaA r u m a break", resulted in uplift and erosion followed by a rapid deepening of the shelf area. The Coniacian to Campanian A r u m a Group is represented in the Oman Mountains by limestone conglomerates and marls of the Muti Formation which disconformably overlie the C e n o m a n i a n carbonates. In the sub-surface, to the west and southwest of the mountains, the Muti grades westwards into the Fiqa shales, a sequence of S a n t o n i a n - C a m p a n i a n deep water marls and calcareous shales (Glennie et al. 1974). These are overlain by the Juweiza Formation, only found in boreholes west of the northern mountains. This formation consists of a possible thickness of over 3000 m of conglomerates, sands and marls of Campanian to ?Lower Maastrichtian age (Glennie et al. op. cit.). The Juweiza contains abundant Hawasina and Semail clasts and is interpreted by Glennie et al. (op. cit.) as a flysch deposited in a foredeep in front of the advancing nappes. 1.5.2 The Allochthonous Units

The O m a n Mountains allochthon comprises a complex sequence of thrust slices or nappes (Fig. 1.5). The lowest structural unit is composed of Mesozoic carbonates, called the Sumeini Group by Glennie et al. (1974). This unit apparently travelled only a short distance compared to the higher units, and was described by them as "parautochthonous". The remaining allochthonous units were divided by Glennie et al. (1974) into the "Hawasina Allochthonous Unit" and the "Semail Nappe", the two being separated by a discontinuous unit of melange called the " O m a n Melange". The allochthon is here subdivided into five units (Fig. 1.6): BASE

The The The The The

Surneini Group Hawasina Assemblage* Haybi Complex* Sernail (Ophiolite) Nappe* Batinah Complex*

TOP

1.5.2.1 Sumeini Group

The Sumeini Group has a total thickness of about 2000 m (Glennie et al. 1974) and is exposed as a n u m b e r of anticlinal structures along the western edge of the mountains and in the Hawasina Window (Fig. 1.5). The base is nowhere seen, but the Sumeini carbonates are overthrust by the basal units of the Hawasina Assemblage. The rocks are mainly poorly fossiliferous fine-grained limestones of Triassic to Cenomanian age,

but with shallow water dolomites and limestones, including a local reef facies, at the base of the sequence. There are intraformational conglomerates, horizons which were interpreted by Glennie et al. (1974) as slump horizons. Glennie et al. (op. cit.) considered that the Sumeini Group represents Mesozoic continental slope deposits laid down to the northeast of the Arabian continental platform (Section 2.3). It is locally overlain by the conglomeratic Qumayrah beds, recognized by Glennie et al. (1974) as an equivalent of the Muti Formation. 1.5.2.2 The Hawasina Assemblaget The Hawasina Assemblage+ is composed of Mesozoic hemipelagic and pelagic sediments, mostly of turbidite origin, which form an array of complexly deformed thrust sheets between the autochthonous and parautochthonous limestones below, and the Semail Nappe above. It is equivalent to the Hawasina Series* of Lees (1928) or the Hawasina Allochthonous Unit- of Glennie et al. (1974). Glennie et al. (1974, p. 30) defined the Hawasina as " . . . a complex association of folded and faulted lithological sequences, comprising quartz sand and carbonate turbidites, silicified limestones and radiolarian cherts, containing fossils of Triassic to Cenomanian age and blocks of white, fractured and partly recrystallized shallow marine limestones of Permian and Triassic age that are either associated with deep-water sediments or have a substrate of sheared basalt pillow lavas". The Hawasina crops out along the southern and western sides of the mountains, notably in the Hamrat Duru ranges, but there are a number of inliers in the central part of the mountains, including the "type-area" of the Hawasina Window. Glennie et al. (1974) showed from detailed biostratigraphic studies that the Hawasina sediments range in age from at least as old as Middle Triassic to mid-Cretaceous (Cenomanian). The groups and formations introduced by the Glennie team (Fig. 1.6) are not conventional stratigraphic units, but each represents a separate fault-bounded slice containing sediments deposited throughout much of this period. The main units found in the central part of the mountains, in the stacking order of the thrust sheets from base to top, are as follows: (i) The Harnrat Duru Group is the thickest and most extensively developed of the Hawasina units consisting of over 1000 m of Lower Triassic to Cretaceous (Cenomanian) sediments divided into the Zulla, Guwayza, Sidr and Nayid Formations that consist of turbiditic quartz sandstones, siltstones, shales and redeposited limestones with some cherts. They are believed to have been deposited northeast of the Sumeini Group at the base of the continental slope and on the continental rise (Glennie et al. 1974; G r a h a m 1980b).

*The terms "nappe', "assemblage" and "complex" are defined as follows: "Nappe" is "an allochthonous tectonic sheet which has moved along a thrust fault" (McClay 1980). There is no implication of largescale recumbent folding or inversion of strata in this definition. "Assemblage" is "an allochthonous structural unit composed of several thrust slices of the same or closely related rock units occurring at the same general structural level" (Williams 1975). "'Complex" is "~one or more thrust slices of largely unrelated rocks brought together by tectonic processes" (Williams 1975). tThe informal term "assemblage" (following the definition of Williams (1975) is here used as the collective term for the Hawasina. Names such as "Group" and ~ that have been used in the past are not appropriate as they do not fit the accepted definitions of these terms (ACSN 1961).

The Geological Background (ii) The Wahrah Formation consists of some 200 m of Jurassic to Cretaceous limestone turbidites and cherts that are generally finer grained and more siliceous than the Hamrat Duru sediments of equivalent age and were probably deposited further from the shelf edge (Glennie et al. 1974; Graham 1980b). (iii) The A l A y n Formation is a dominantly quartz-sandstone unit about 350 m thick containing some fine grained limestones, shales and cherts. A Middle Triassic to Lower Jurassic age indicates correlation with the Zulla and lower Guwayza Formations of the Hamrat Duru Group. The high tectonic position of the AI Ayn suggests that some of the quartz sands were derived from local highs within the basin itself (Glennie et al. 1974). (iv) The Halfa and Haliw Formations are thin sequences of Triassic to Lower Cretaceous cherts and fine grained silicified limestones. These are rarely more than a few tens of metres thick and are highly deformed. They contain interbedded basaltic lavas or, more commonly, sills which in some cases were emplaced soon after sediment deposition and have been radiometrically dated as mid-Jurassic to mid-Cretaceous (Lippard & Rex 1982). These distal sediments were interpreted by Glennie et al. (1974) as having been laid down on oceanic basement in the deepest part of the Hawasina basin. (v) The A l Aridh and Ibra Formations consist of up to 125 m of coarse conglomerates and breccias, composed largely of fragments of shallow water limestones with occasional clasts of basalt. The rudites may be interbedded with redeposited Halobia-bearing limestones and cherts of Triassic age. They were considered by Glennie et al. (1974) to be deep water deposits formed close to the sites of Exotic limestone deposition. In the Hawasina Window the Hawasina Assemblage consists of lower thrust sheets composed of relatively proximal turbidite facies sediments (Hamrat Duru, Wahrah and A1 Ayn Formations) with a combined thickness of 1500-2000 m tectonically overlain by higher thrust slices composed largely of distal turbidites and pelagic sediments (Halfa and Haliw formations) associated with redeposited limestone breccias (AI Aridh Formation) and some large exotic limestone masses (Graham 1980a,b). Tectonic thickening (by factors of 2.5-3.0 in the lower units and more in the upper units) of each thrust sheet is the result of secondary thrusting or reverse faulting and associated folding that occurred during late Cretaceous nappe emplacement. The folds in the Hawasina sediments are usually asymmetrical tight to isoclinal, semi-recumbent or reclined structures and the thrusts are usually bedding-parallel with small displacements on the short, overturned limbs of folds. Folding of competent sandstone and limestone layers is by flexural slip followed by flattening and often accompanied by boudinage. Incompetent shales are deformed by homogeneous shear and in most cases they have developed a secondary cleavage. In the thin, centimetre-bedded cherts, buckling and flattening have resulted in the formation of complex kink folds.

I3

Searle et al. (1980)), Exotic limestones, and Hawasina-type sediments contained in a fine grained pelitic matrix (the Hawasina Melange) and a thrust slice of sub-ophiolite metamorphic rocks and serpentinite (Basal Serpentinite) derived from the base of the overlying ophiolite. It is now recognized that the melanges in this unit have diverse origins and include both sedimentary deposits, formed before nappe emplacement, and syn-tectonic melanges (Graham 1980b; Searle & Malpas 1980; Lippard et al. 1983). Some are composite with tectonic effects overprinting earlier sedimentary fabrics. The Haybi Complex as a whole is highly deformed, with numerous thrust slices and imbrications, so that one or more of its components may be repeated or missing. The stacking order of the thrust sheets may be locally reversed or jumbled although, in general, the Basal Serpentinite is the highest unit. In places, the Haybi complex is missing or reduced to a thin zone of serpentinite, elsewhere it is over 1000 m thick. (i) The Exotic Limestones (Wilson 1969; Glennie et al. 1974; Searle & Graham 1982) are large allochthonous masses of massive shallow-water carbonates of Middle to Upper Permian and Upper Triassic ages that occur near the base of the Semail Nappe or, in some cases, locally overlie it. They form masses of all shapes and sizes ranging from metre-sized blocks up to huge masses of Upper Triassic limestones; for example, Jebel Khawr which has a volume of c. 600 km 3 (Glennie et al. 1974). Some of the limestones are associated with pelagic sediments or pillow basalts (see below) and they have generally been interpreted as carbonate build-ups on ocean islands distant from the continental margin (Glennie et al. 1974), although the Permian examples may have formed on a continental platform or shelf area prior to Triassic rifting (Graham 1980b) (Section 2.2). On Jebel Kawr, there are Jurassic cherts conformably overlying the Upper Triassic limestones. (ii) Haybi Volcanics (Searle et al. 1980): These lavas and tufts, including tholeiitic, transitional and alkaline basalts with some trachytes, are up to 700 m thick and have yielded K-Ar late Triassic radiometric ages of c. 220 Ma on separated biotites (Searle et al. 1980; Lippard & Rex 1982). They contain interbedded Upper Triassic shallow water Exotic limestones, redeposited limestone/volcanic conglomerates, cherts and radiolarites. (iii) Hawasina Melange (Graham 1980b): This consists of mega-breccias and conglomerates containing large blocks ("olistoliths") of Exotic limestones, Haybi volcanic rocks and Hawasina sediments. Where there is a sedimentary matrix of red shale or sheared mudstone these deposits are interpreted as olistostromes (Graham 1980b). Radiolarian faunas indicate a probable mid-Cretaceous age. (iv) Basal Serpentinite This is a highly sheared serpentinite melange containing tectonic blocks of the sub-ophiolite metamorphic rocks as well as Haybi volcanics, Exotic limestones and Hawasina sediments and in some places forms much of the upper part of the Haybi complex. The serpentinite was derived from the base of the Semail ophiolite during nappe emplacement (Searle & Malpas 1980).

1.5.2.3 The Haybi Complex

The term Haybi Complex was introduced by Searle (1980) and Searle & Malpas (1980) for the allochthonous rocks that are tectonically sandwiched between relatively well-organized thrust sheets of Hawasina sediments below and the Semail Nappe above (Fig. 1.5). This unit is in part equivalent to the Oman Melange of Glennie et al. (1974) and includes thrust slices and olistoliths of volcanic rocks (the Haybi Volcanics of

1.5.2.4 The Semail Nappe

This is the largest and most extensive unit of the allochthon with a present day outcrop area of almost 20,000 km 2. As a result of syn- and post-emplacement faulting and folding, the ophiolite nappe has been broken into a number of more or less intact structural blocks. The largest of these expose a classic "Penrose" ophiolite stratigraphy from a basal metamorphic

I4

Chapter 1

sheet (0-500 m thick) tectonically overlain by ultramafic tectonites ("Mantle Sequence") (8-12 km) in turn overlain by layered peridotites and gabbros (0.5-4 km), high-level plutonic rocks (10-500 m), a sheeted dyke complex (1-1.5 km) and, at the top of the succession, extrusive lavas and interbedded pelagic sediments (0.5-2 km) (Reinhardt 1969; Glennie et al. 1974; Coleman 1977, 1981; Sinewing 1980a). The whole represents a complete section up to 20 km thick through midCretaceous oceanic lithosphere. Geophysical data (Shelton 1984) show that the maximum structural thickness of the nappe is between 5 and 10 km suggesting that the basal thrust cuts across the internal stratigraphic units. This is seen at the edges of the blocks where there is evidence of marked thinning of the nappe towards the inland (leading) and coastal (trailing) edges of the nappe als0 in cross-strike fault zones such as Wadi Jizi and Wadi Ahin. Along much of the leading edge, the nappe is affected by imbricate thrust faulting giving rise to repetition of the stratigraphy; in many other places it is extended by normal faulting. Many of the faults in the upper part of the ophiolite can be demonstrated to be oceanic structures even though they may have been reactivated during and after emplacement (Smewing et al. 1977). The base of the nappe is marked by a zone of mylonitized and highly sheared peridotites, known as the "Banded Unit" (Searle 1980), the structures of which clearly overprint an earlier mantle tectonite fabric (Boudier & Coleman 1981) and which appear to have formed during amphibolite facies metamorphism in the metamorphic sheet at an early stage of emplacement of the nappe (Searle & Malpas 1980). Radiometric dating of the metamorphic rocks yields mid to late Cretaceous K-Ar and Rb-Sr mineral ages ranging from 95 to 70 Ma (Allemann & Peters 1972; Lanphere 1981). During the later stages of nappe emplacement and post-emplacement the lower ultramafic part of the ophiolite has been extensively serpentinized. Apart from these effects, the Semail Nappe represents a more or less complete and largely unaltered section of mid-Cretaceous oceanic upper mantle and crust. A detailed description of the stratigraphy, ocean-floor structure and igneous and metamorphic history of the Semail ophiolite is the main theme of this volume. 1.5.2.5 The Batinah Complex

The Batinah Complex (Woodcock & Robertson 1982a) comprises several allochthonous units which overlie the Semail Nappe along the eastern edge of the mountains. It consists of a lower melange, the "Batinah Melange", composed of blocks of the Semail ophiolite, Exotic limestones, Haybi volcanics, Hawasina sediment, metamorphic rocks and serpentinite, which locally has a depositional contact with the top of the ophiolite. The melange is tectonically overlain by several thrust sheets of "Hawasina-type" sediments (Glennie et al., 1974; Graham 1980a) that have been subdivided by Woodcock & Robertson (1982b) into three units: a lower thrust slice of Triassic to Lower Cretaceous cherts and fine grained limestones (Barghah Formation); a middle unit of Upper Triassic limestone Exotics with some volcanic rocks and serpentinite, and several upper thrust sheets of Upper Triassic to Upper Cretaceous turbidite sandstones and limestones, shales and cherts (Sakhin and Salahi Formations). These nappes, which are the highest tectonic unit of the allochthon, are collectively referred to by Woodcock & Robertson (1982a) as the "Batinah Sediment Sheets".

[ I 9 I

METAMORPHIC AGES ~ Amphibolites ti B ueschists __ I I G r e e n s c h i s t s

MAASTRICHTIAN -70

CAMPANIAN

-80

CONIACIAN TURONIAN

-90

CENOMANIAN ALBIAN Sedimentary Formations

/ J

100

Igneou~ events

Fig. 1.8. Timing of late Cretaceous events in the Oman Mountains. 1.5.3 Late Cretaceous nappe emplacement

The Oman Mountains were subjected to a major deformation in the late Cretaceous (Turonian-Campanian) (Fig. 1.8) with the emplacement of far-travelled thrust nappes from northeast to southwest onto the Arabian continental margin (Lees 1928; Glennie et al. 1973, 1974). The large translation distances involved, several hundred kilometres, lead to the superposition of rocks of similar age, but markedly contrasted type and facies. The earliest deformation apparently began with intraoceanic thrusting which resulted in the detachment of a slice, up to 20 km thick, of mid-Cretaceous oceanic lithosphere (the Semail Nappe) along a low-angle thrust plane. As the hot, relatively newly formed oceanic lithosphere was emplaced upon the upper parts of the overthrust oceanic rocks, these were converted into the high-grade amphibolite facies metamorphic rocks (dated at 96-85 Ma, Allemann & Peters 1972; Lanphere 1981). As thrusting progressed southwestwards towards the Arabian foreland, thrust slices of Mesozoic ocean floor and continental rise sediments, comprising the Haybi Complex and Hawasina Assemblage, were incorporated into the nappe pile and, close to the base of the Semail Nappe, were locally metamorphosed to give the greenschist facies rocks dated at 85-70 Ma (Lanphere 1981). The Hawasina nappes were emplaced over the contemporaneous Mesozoic slope and shelf facies carbonates which, in places, formed parautochthonous thrust slices at the base of the nappe pile. In the north Saih Hatat area near Muscat, parts of the lower allochthonous units and the underlying autochthon were subjected to a high pressure blueschist to greenschist facies metamorphism which has been dated at 100-80 Ma (Lippard 1983). During the last stages of nappe emplacement in the Campanian, the Semaii Nappe overrode all the underlying units. This late movement was facilitated by serpentinization of the basal ultramafic part of the nappe and was accompanied by the formation of tectonic melanges. At about the same time the ophiolite nappe broke up into a number of tectonic blocks along structural lines of weakness, including several crossstrike fault zones, and melanges were emplaced between the blocks and onto the surface of the nappe (Batinah Melange). The last emplacement event was the westward thrusting of several sheets of Mesozoic Hawasina-type sediments (Batinah Sediment Sheets) on top of the ophiolite and its melange blanket (Woodcock & Robertson 1982b). Nappe emplacement onto the Arabian foreland was preceded by a deepening of the Arabian shelf margin in the mid to late Cretaceous with the deposition of deep-water sediments

The Geological Background (Muti and Fiqa formations) disconformably above the Cenomanian shelf and slope limestones following subsidence of the continental margin (Glennie et al., 1974). Deep water sedimentary melanges (olistostromes), found in the upper part of the Hawasina and Haybi Complexes, probably developed at an earlier time in a more distal setting. Syn-emplacement

flysch deposits of Campanian age (Juweiza Formation) were deposited in a fore-deep in front of the advancing nappes. Post-emplacement folding, uplift and erosion of the nappes was complete by the Maastrichtian when a transgressive series of clastic sediments and shallow water limestones were deposited across the mountain area. Late Cretaceous nappe emplacement in the Oman Mountains has several features in common with other "forelandtype" fold and thrust belts (Dahlstrom 1970; Elliott & Johnson

1980) in that thrusting developed progressively from an external (oceanic in this case) area towards an internal continental region. As a result the farthest travelled nappe (the Semail Nappe) occurs highest in the sequence and the lowest thrust slices (Sumeini unit) were the last to be detached and move. The incorporation of successively lower thrust slices gives rise to the characteristic "piggy-back" style of thrusting (Elliott 1976). It should, however, be noted that there are exceptions to these generalizations probably caused by the irregular nature of the surface over which the nappes were emplaced and the lenticularity of many of the thrust slices. For example, the Batinah Sheets, composed of continental margin sediments, are locally emplaced onto the Semail Nappe. There are several other examples of late thrust movements being 'out of sequence' causing lower units to occur above what are usually

~3~

I

I

~

56 ~

_

~5

RasMusandam

~z~

26ON

I: :.jl

-

Outcrops of Maastrichtian-Tertiary sediments

$

Tertiary fold axes

I

Major faults (tick on downthrow side) -..I-. Possible basement faults ormonoclines

\

Jebel Hafit

~"~)

.~

k~Sohar

\~\

x\

24 ~

o

Jebel Awaynah

, ~ ~

,

t.

~

Ibri \~

/

.

<.l " ~ ) ' ' ' ' "

9

..t

(

~ ~ ._"(~

~

~

q'> Qahlah

~.

V. .,~ ~ " . - \ \ ~ ..~,,,~)

Sur

Rasal Hadd

I / I

Fig. 1.9. Tertiary outcrops and structures.

"

/

!

x6

Chapter 1

tectonically higher ones. The last movement on the Semail Thrust appears to have been relatively "late" so that it cuts across imbricate structures in the underlying Haybi and Hawasina complexes (Searle 1980). Emplacement of the nappes was preceded by a rapid collapse of the continental margin which allowed it to accommodate over 10 km thickness of allochthonous rocks. As emplacement proceeded, the Aruma basin foredeep developed and migrated in front of the advancing nappe pile locally accumulating up to 3-4 km of deep water clastic sediments. Rapid uplift (isostatic rebound?) of the area occurred during the last stage of emplacement. 1.5.4 Neoautochthonous sediments and tertiary structures

Maastrichtian, Palaeocene and Eocene sediments, mainly limestones with a total thickness up to 600 m, unconformably overlie the late Cretaceous allochthonous units of the Oman Mountains (Lees 1928; Morton 1959; Glennie et al. 1974) along the eastern and western flanks of the mountains (Fig. 1.9). There are two facies: fluviatile conglomerates, sands and gypsiferous marls (Qahlah Formation) and fossiliferous marine limestones (Simsima Formation). The coarse grained clastics are largely restricted to the eastern side of the mountains; whereas, by contrast, on the western side, the limestones rest directly on older rocks. In places subaerial exposure and lateritic weathering of the ophiolite occurred before Maastrichtian sedimentation took place (Coleman 1981). The Maastrichtian-Lower Tertiary sediments were deformed during the mid-Tertiary, post-Middle Eocene. The tectonism was particularly severe in the north where, along the west side of the Musandam Peninsula, the Mesozoic limestones of the Ruus al Jibal massif were displaced several

kilometres westward along the Hagab thrust (Hudson et al. 1954b; Allemann & Peters 1972). Along the line of the thrust the limestones are locally overturned and thrust over Eocene shales (Searle et al. 1983). In general, the Tertiary fold axes trend parallel to the mountains from N-S in the north, to NW-SE and E-W further south although there are local variations and changes in strike. It has been suggested that the Tertiary folding is the result of reactivation of basement faults and local uplift, perhaps partly due to salt doming (Tschopp 1967; Glennie et al. 1974). A regional origin is however, suggested by their widespread development and the consistency of trends and directions of the fold axes (Searle et al. 1983) (Fig. 1.9). The post-Miocene history of the area is one of repeated uplift of the mountains and subsidence of the flanks with thick sediment deposition. Glennie et al. (1974) report over 4000 m of Neogene sediments in an offshore borehole on the Oman continental margin where extensive slumping of the thick semi-consolidated sediments has been identified on seismic profiles (White & Ross 1979). Cycles of repeated uplift are recorded by a series of gravel-capped terraces on the flanks of the mountains. Elevated terraces up to 300 m above sea level are found on Jebel Nakhl and on the northern edge of Saih Hatat just south of Muscat, but are as yet undated. In the major mountain wadis cemented gravels are "perched" up to 50 m above the present valley floors. Raised beaches and marine terraces characterize most of the coastal area from Dibba southwards (Lees (1928) reports "sub-recent" shells west of Shinas at 375 m above sea level), but to the north the Musandam Peninsula has a drowned, ria-type coastline and has been sinking in recent times (Vita-Finzi 1973).

Chapter 2 Evolution of the Oman Tethys 2.1 Introduction

Gondwanaland supercontinent, to produce a narrow ocean basin by the end of the Triassic, followed by the Jurassic to mid-Cretaceous development of a northeast-facing passivetype continental margin. The history of the Oman margin, as recorded by the sediments that overlie the pre-Permian basement, begins after a marine transgression in the Middle Permian that followed Palaeozoic (post-Ordovician) folding, uplift and erosion. The Triassic marks the first appearance of a deep water basin in which the earliest hemi-pelagic and pelagic Hawasina sediments were deposited. Volcanicity in the late Triassic (Haybi Volcanics) shows a trend from early alkaline through transitional to later tholeiitic types, reflecting the further development of the ocean basin. At the end of the Triassic in the present area of the Oman Mountains there was a broad continental shelf bordered to the northeast by a deep-water (Hawasina) basin and beyond this an outer margin "high" composed of shallow-water limestones (now represented by the "Oman Exotics") which formed as carbonate platforms developed on a basement of the late Triassic volcanics. In the mid-Jurassic the deposition of largely quartz-rich terrigenous turbidites in the Hawasina basin, which had begun in the mid-Triassic, was succeeded by calciturbidite-chert deposition. The rift-related volcanicity ceased at the end of the Triassic and the outer margin carbonate platforms subsided and were blanketed by pelagic sediments. In the late Jurassic-early Cretaceous there were relatively deep-water, sediment-starved conditions across the whole continental margin from the shelf to the ocean basin. By the mid-Cretaceous (Cenomanian) continued subsidence had resulted in the deposition of over 3000 m of shelf (Hajar Supergroup), 2000 m of slope (Sumeini Group) and at least 1000 m of continental rise (proximal Hawasina facies) sediments, whilst only a few tens of metres of pelagic limestones and cherts (distal Hawasina sediments) had been deposited in an abyssal oceanic environment between the late Triassic and mid-Cretaceous. Contemporaneous magmatism is represented by alkali dolerite and wehrlite sills (dated at 16093 Ma) that intrude the Hawasina sediments. In the mid-Cretaceous tectonic instability of the continental margin was reflected by uplift and erosion (the Turonian "Wasia-Aruma break") followed by subsidence and relatively deep-water sedimentation in the Coniacian to Campanian (Aruma Group). In the outer part of the continental margin, sedimentary melanges (olistostromes) formed in the midCretaceous as a response to rapid differential uplift and subsidence probably marking the beginning of subduction. These events heralded the emplacement of the continental margin sediments and volcanics as a series of thrust nappes (Hawasina Assemblage and Haybi complex). The Semail Nappe represents a slice of oceanic lithosphere that was generated at a midCretaceous spreading axis to the NE of the Oman margin and which was probably formed above a northward-dipping subduction zone in a marginal basin tectonic setting (see Chapter III). The ophiolite nappe was detached by intra-oceanic thrusting in the Turonian (ca. 90 Ma) and was then emplaced (obducted) southwestwards over the Hawasina and Haybi allochthons onto the continental margin in the Santonian to Campanian (87-72 Ma). Late in the emplacement sequence tectonic melanges formed beneath the Semail Nappe and were

The "Oman Tethys" was part of an ocean area that extended from the Mediterranean to eastern Asia between the ancient supercontinents of Gondwanaland (Africa-Arabia-India) to the south and Eurasia to the north (Fig. 2.1). The ocean appears to have formed in the early Mesozoic by the rifting off and northward migration of several small continental fragments or "microplates"; including Anatolia, Central Iran, Afghanistan and Central Tibet, away from the remainder of Gondwanaland (Dewey et al. 1973; Stoneley 1974; Boulin 1981). This seaway has been called the "Southern Tethys" or "Neo-Tethys" to distinguish it from the older "Palaeo-Tethys'" that lay to the north of the microplates and which was largely destroyed by pre-Liassic plate collision (Takin 1972; Sengor 1985; Teleki 1981). On a local scale, Glennie et al. (1974) referred to the Mesozoic ocean to the north of Oman as the Hawasina Ocean; here it is called the Oman Tethys. The evolution of the Oman Tethys throughout the late Palaeozoic and Mesozoic, up until its destruction by ophiolite obduction and nappe emplacement onto the Arabian platform in the late Cretaceous, is here discussed in terms of the rifting of the northern margin of the Arabian continent, then part of a

30~N ~ . ~ ~

Y

~

============================================ ~':,:,iiiiiiii!iii!!!!!iiiiiiiii ! !i ! : : Ap \4:q::::::::::: i i i i i ~_.

//

/ / / / / / / 7"--,._

An

~'--... >-

"~::::::::::: ........ : :
"-S, %

//~ 30os

#"~ /

"

~

xx

o.. X

x

x::

\:

~'. \\X \x

/

/ANTARCTICA / / / / / . "/'~'2~/AU / / ~ / /sTR/ /ALl/A~' ~ , ~]

[

Eurasia ~

Gondwanatand ~

Neo-Tethys ~

Palaeotethys

[Tethyan Microcontinents; Ap Apulia, An Anatolia, Cl Central Iran, Af Afghanistan, T Tibet

Fig. 2.1. Probable Late Triassic palaeogeography of Tethys (after Boulin 1981).

I'/

I8

Chapter 2

included in imbricate and cross-cutting fault zones and locally emplaced onto the ophiolite surface. The highest, and last emplaced, units of the allochthon are several thrust sheets of Mesozoic Hawasina-type sediments (Batinah Complex). The allochthonous units, which have a combined tectonic thickness in excess of 20 kin, were uplifted and eroded before the deposition of a transgressive series of Maastrichtian to early Cenozoic clastics and shallow water limestones. Mid-Tertiary (early Miocene?) folding and faulting produced the present mountain belt which has since undergone uplift and erosion. Thick Neogene molasse deposits flank the mountains and occur on the margin and floor of the Gulf of Oman. Geological and geophysical evidence indicates that the Gulf of Oman is oceanic and may be a remnant of the late Mesozoic Tethys. Since late Cretaceous-early Cenozoic time it has been subducted northwards beneath the Makran active margin. In this chapter the history of the Oman Tethys is discussed in three stages:

Michard 1983) and seems to have been unrelated to the later rifting events. Permian rocks are of restricted occurrence in the Oman Mountains and are represented by the lower part of the autochthonous shelf carbonate succession (the Saiq Formation of the Hajar Supergroup (Glennie et al. 1974) and the allochthonous Permian limestone Exotics (Lees 1928; Wilson 1969; Glennie et al. 1974) (Fig. 2.2). The autochthonous succession consists of Middle Permian limestones and dolomites followed by Upper Permian dark dolomitic mudstones which are typical of the Middle and Upper Permian of eastern Arabia that is generally of restricted shallow marine to intertidal sabkha facies with more open water conditions to the north and east (Murris 1980). The Permian Exotics are shallow water marine limestones containing abundant Middle to Upper Permian faunas (Glenhie et al. 1974) (Table 2.1). They occur as allochthonous masses, as separate tectonic slices or as blocks in melanges, and are to be found in the upper part of the Hawasina allochthon or in the Haybi and Batinah complexes. Lees (1928) called all the exotic limestones the "Permian Klippen", but later work by Wilson (1969), who first called them the "Oman Exotics", and Glennie et al. (1974) proved that most of the large allochthonous limestone masses in the Oman Mountains are of Upper Triassic age. The Permian examples are generally small in size, from a few metres up to 2 km across, are few in number and none has a stratigraphic thickness greater than about 200 m (Table 2.1). In general, the Permian Exotic limestones are of more open marine facies than the contemporaneous shelf carbonates of the Saiq Formation (Glennie et al. 1974; Graham 1980a). Three lithofacies have been recognized by Graham (1980b): a reef facies of vuggy to massive limestones consisting mainly of coral-algal boundstone with some bryozoa and stromatoporoids in which void space is filled with shell fragments, including gastropod, brachiopod and echinoderm debris (this is the only true reef facies found in the Oman Exotics); a f o r e - r e e f facies of massive breccias of reef limestone with boulder-sized clasts

(1) Permo-Triassic rifting and continental break-up (2) Jurassic to mid-Cretaceous passive margin development (3) Mid to late Cretaceous syn- and post-tectonic events

2.2 Rifting and Continental Break-up This section describes the history of rifting and break-up of the northern margin of the Arabian platform to form an ocean basin by the end of the Triassic. The sequence of events is recorded by (1) Permian sediments, (2) Triassic sediments and (3) the late Triassic Haybi Volcanics of the central and northern Oman Mountains. 2.2.1 Permian sediments

Mid-Palaeozoic folding and uplift in Oman occurred mainly on NE-SW trending axes (Tschopp 1967" Glennie et al. 1974:

SHELF (Hajar Supergp.) shallow water carbonates

Upper m o o3

SLOPE (Sumeini Group) shallowmoderately deep water carbonates

BASIN (Hawasina series)

Proximal Hamrat Duru AI Ayn Grouo formation (see Fig 2 3 'or details)

Y/ K/J

t--

'

=

At Aridh formation

'q

I

Upper Triassic exotics

"

I

Distal

9

Haybi volcanics

Middle

E: )PARAUTOCHTHON Lower

' Z'//

t-2 -2 -3'

ALLOCHTHON

i

Permian exotics - ~ Dolomites Z <

Upper

~

n" ~~

Middle

~

Limestones

Conglomerates 1 Sandstones

I

Shaly limestone

[ - - 7 1 Cherts

,[

Shales

[ ~

jL

Volcan,cs

AUTOCHTHON

Fig. 2.2. Permian and Triassic sediments and facies in the Oman Mountains.

Halfa/Haliw formations

Evolution of the Oman Tethys

19

Table 2.1. Comparisons of Permian and Triassic Exotics.

PERMIAN

TRIASSIC

Age range Size

Middle to Upper Permian Metre-sized blocks. Rarely as large masses up to 200 m by 2 km 2 (Jebel Yanqui).

Stratigraphic relations

Not preserved. Jebel Qamar is not a typical example (see text).

Lithofacies

Reefal. Coral-algal boundstones. Fore-reef.

Upper Triassic (Carnian-Norian) Metre to kilometre-sized blocks. Several large masses in excess of 300 m by 5 km 2. Jebel Kawr (900 m by 600 km 2) is exceptional. Often underlain by Haybi volcanics, occasionally by chert-limestone sequences. Overlain by volcanics or Jurassic cherts. "Back-ree]~'. Megalodontid wackstones and mudstones with occasional pelletoid, oolitic and fragmental grainstones and packstones. "Fore-reef". AI Aridh facies. Megabreccias of Exotic limestone clasts in a lime mudstone matrix. Interbedded with redeposited grainstone turbidites and thin-bedded limestones containing

Clast-supported calcirudites; angular blocks of reef limestone in a lime mudstone to grainstone matrix. Back-reef. Thin-bedded mudstones, packstones and grainstones.

Fauna

Rugose corals, stromatoporoids, bryozoa, gastropods, echinoderms. Fusulinids: Verbeckia,

Neoschwagerina, Parafusulina. Tubiphytes. Foraminifera; Hernigordiopsis.

Diagenesis

Environment of deposition

Palaeogeographic setting

Abundant evidence of early diagenesis and cavity formation. Encrustation of early cement by algae. Interbanding of early cement and skeletal debris. Cavities infilled by fibrous calcite, followed by deposition of vadose silt and dolomite. Shallow water, high-energy reef. Periodic emergence. Fore- and back-reef facies deposited in lower energy conditions. No evidence of deep water environments. Uncertain. Deposited on local "'highs" on an intracontinental shelf'?

in a lime mudstone matrix and interbedded with finer grained calcirudites and calcarenites, and a rare back-reef facies of massive to metre-bedded fusulinid bearing grainstones and packstones that resemble limestones more typical of the Upper Triassic Exotics (Table 2.1). On the Jebel Q a m a r Exotic in the Dibba zone there is a continuous sequence of Permian and Triassic carbonates unconformably overlying the Palaeozoic basement (Hudson et al. 1954a). Glennie et al. (1974) describe the Jebel Q a m a r limestones as a shelf-edge or slope facies, but they are of typical platform types and were probably once part of a continuous shelf area. Nowhere else in the Oman Mountains has an in situ substrate to a Permian Exotic been found, for although some are apparently underlain by volcanic rocks (Glennie et al. 1974), the contacts are everywhere tectonic or the limestones occur as blocks in melanges. Thus, although Glennie et al. (op. tit.) suggested that they formed as carbonate cappings to oceanic seamounts, there are no associated volcanic rocks or deep water sediments of Permian age to support this hypothesis. G r a h a m (1980a) regarded the Permian Exotics as formed on highs within a rifted, intracontinental basin; however, it is equally likely that they formed on a once-continuous shelf area, part of the Arabian platform, that was broken up by subsequent Triassic rifting.

Halobia. Megalodontid lamellibranchs, plus other thickshelled forms such as Perna and Opisoma. Foraminifera; lnvoluta. Brachiopods; Spiriferina greisbachi, S. alkivaga. Ammonites; Jovites, Juvavites, Leiophyllites, Siberites ? Corals, bryozoa, crinoids, gastropods, echinoderms. Early diagenetic effects largely obscured by recrystallization and dolomitization. Several generations of cementation with early dog-tooth and late drusy spar.

Shallow water, high to low energy. AI Aridh facies represents deposition in moderate to deep water. Carbonate platforms (lOOs km e in area). Largely built on a volcanic substrate distant from the continental margin. Steep margins passing rapidly into deep water. "'Bahamian type".

2.2.2 Triassic sediments Triassic sediments in the Oman Mountains are more diverse and widespread than those of Permian age and include, as well as massive shallow water carbonates of both shelf and Exotic facies, deep water turbidites and pelagic sediments of the lower Hawasina series (Fig. 2.2). Associated with the outer margin Triassic sediments are considerable volumes of contemporaneous voicanics, known as the Haybi Volcanics (Searle et al. 1980). The shallow water carbonates of the autochthonous shelf succession that conformably overlie the Permian are the massive dolomites of the Triassic Mahil Formation (Glennie et al. 1974). These represent part of a continuously subsiding shallow marine basin that extended across most of the Arabian platform at this time (Murris 1980). At the base of the parautochthonous Sumeini Group, Middle to Upper Triassic shallow water dolomites, marls and a local reef facies represent deposition in a slope area to the northeast of the present mountain area. Further to the north and east during the Triassic a deep water basin, or series of basins, developed in which coarse to fine grained turbiditic sediments, represented by the oldest units of the Hawasina, accumulated. Still further from the shelf edge a line of ridges or platforms, composed of Upper

20

Chapter 2

Triassic shallow water limestones overlying Triassic volcanics. formed outer margin highs around and beyond which deep water pelagic sediments were deposited. In this section the basin facies sediments are described. They are divided into two groups; (i) proximal turbidite facies; and (ii) deep water pelagics and the Upper Triassic Exotics.

Base of Guwayza limestones Disturbed shales with sandstone lenses Medium to coarse sandstones with shale partings

N

Thin bedded fine grained sandstones with shale partings Abundant trace fossils, including Palaeodictyon. Medium to coarse sandstones with conglomeratic bases.

2.2.2.1 The early Hawasina; Triassic deep-water sedimentation The oldest sediments of the Hawasina series comprise sandstones, shales, cherts and limestones of the Zulla Formation, at the base of the Hamrat Duru Group (Fig. 2.3) (Glennie et al. 1974; Graham 1980a). The upper part of the Zulla Formation has been dated as Upper Triassic by the presence of the foraminifera Involutina sp. and Diplotremina sp. (Glennie et al. 1974). The age of the lower beds is, however, uncertain and the base is not exposed, being truncated by a thrust at the base of the Hawasina allochthon. The sediments show rapid lateral and vertical variations, although the general sequence is of interbedded shales, siltstones and fine grained sandstones at the base, a middle coarse grained sandstone unit and an upper series of interbedded cherts and limestones (Fig. 2.3). The overlying Guwayza Formation consists of generally coarser grained and more massive sandstones with interbedded shales up to 250 m thick which range in age from Upper Triassic to Lower Jurassic (Glennie et al. 1974). The AI Ayn Formation, an equivalent of the lower Guwayza and perhaps part of the Zulla Formation, forms a separate, tectonically higher thrust slice. It is up to 400 m thick and is composed of siltstones, shales and fine grained sandstones at the base, a middle unit of thin-bedded limestones and an upper unit of more massive, coarse grained sandstones with shale partings (Fig. 2.3). The Triassic to Lower Jurassic quartz sandstones of the lower Hawasina units are typical turbidite deposits with welldeveloped Bouma (T,,_e) cycles and abundant sole markings. In complete units, a graded conglomeratic base passes up into parallel and cross-bedded sands with laminated siltstones and shales at the top. The sands are nearly all poorly sorted quartz arenites and are often bimodal with an abundant fine sand to silt fraction. The clasts are predominantly angular to subrounded polycrystalline quartz with a minor depleted heavy mineral suite of tourmaline, zircon and opaques. Rare lithic fragments, including phyllite and mica schist clasts are presumably derived from pre-Permian basement rocks (Graham 1980). The dominant mature quartz sand component was probably largely derived from the recycling of Lower Palaeozoic sandstones such as the Rann and Amdeh quartzites. Cross-bedding and bottom structures in the sandstones show predominant transport directions from the west or southwest, particularly for the Guwayza and AI Ayn formations (Fig. 2.3). The coarse grain size and immature nature of the sediments, particularly in the upper part of the AI Ayn Formation, suggested to Glennie et al. (1974) some local derivation of the clastic material from intra-basinal highs as well as from the continental platform to the west and southwest but it could be that the sands by-passed the shelf along submarine canyons. Quartz sands are confined to the Lower Jurassic of the platform succession, so it is unlikely that this area was an important source of the Triassic sands. 2.2.2.2 Outer Margin Triassic sediments

Upper Triassic limestone exotics, limestone breccias interbedded with redeposited Halobia limestones and fine grained

Redeposited oolitic and pelletoidal packstones and grainstones with lime mudstones containing Halobia. Thin-bedded radioiarian cherts with interbedded siltstones and lime mudstones N

Z O Fine sandstones with shale partings. Abundant bottom structures.

o

<[ d d

Sheared shales and thin bedded fine grained limestones.

N

Chlorite-sericite schists and quartz phyllites Thrust contact Top not seen

Massive metre to decimetre-bedded sandstones. Sandstones and shales with abundant sole markings and trace fossits.

0 z

>..

~ T - -

Interbedded dark redeposited limestones and shales

<

lOOm -J

0

L-

Thrust

Fig. 2.3. Representive sections (from Graham 1980b) of Triassic Hawasina sediments in the Hawasina Window. Palaeocurrent directions represented by rose diagrams (data from Graham 1980b and Giennie et al. 1974).

pelagic limestones and cherts comprise the Triassic distal facies sediments of the Hawasina (Fig. 2.2). The Upper Triassic Exotics are massive to well bedded and often highly fossiliferous rocks (Glennie et al. 1974). They range in size from metre-size blocks in melanges to the huge Jebel Kawr massif which covers an area of 600 km 2 and has a stratigraphic thickness of 900 m. Several other large examples, such as Jebels Ghawil and Ajran, on the west side of the mountains, Jebel Misht, north of Jebel Kawr, and Sohar Peak

E v o l u t i o n o f the O m a n Tethys I 56~

I 58~

26~N _

1/

Jebel Qamar (North and South P-T~ )

~

Semail Nappe

~

Haybi and Batinah complexes with Exotic Limestones

/:.

Autochthonous basement

t II "

"

[

P

Permian Exotics

T~

Tnass~c Exotics

Jebel Ghawil peaks(~

Jebel Mahada~tpk t

Jebel Abiad (~) . Jebel al Hut (~)

Jebel A j r a n - Hawrat al Asal

24:--

Jebel Nakhl (~ Jebel Yanqui (P) ~

\ / ~ ~X_

Hawasina W i n d o w / ~

(

Jebel Misht Jebel M i s f : . ~ JebeJ Kawr (T~ 50 Km - - ]

A/ Jebel Bahlah Jebel

i

Hawasina n

a

p -

Jebel Misfah Jebel Kawr,, \ p e ~ ~ _

Jebel Akhdar _ _

B ~

Jebel al Hur // sea lever

Fig. 2.4. The Oman Exotics- distribution and tectonic setting.

and Jebel al Hawr on the edge of the Batinah Plain (Fig. 2.4), form dramatically steep-sided and often isolated mountain peaks. G r a h a m (1980b) recognized several lithofacies in the Upper Triassic Exotics; most common are "back-reef" environments, but no true reef facies has been identified (Table 2.1). The most common rock type is a wackestone containing giant Megalodontid bivalves, often in high concentrations of both whole and broken shells. These thick-shelled organisms are characteristic of the Upper Triassic (Norian) shallow marine carbonate facies throughout the Tethys and are considered to be indicators of warm, shallow water, lagoonal or inter-tidal, and somewhat restricted marine conditions. The Megalodont limestones in the O m a n Exotics are locally interbedded with packstones and grainstones, including oolitic and pelletoidal types, that contain a more diverse fauna of small bivalves, gastropods, echinoderms, algae, calcisponges and rare corals. These probably represent more open water, higher energy conditions. Most of the Upper Triassic Exotics have stratigraphic thicknesses of 200-300 m; Jebel Kawr is exceptional with over 900 m of shallow-water carbonates. Calcirudites of late Triassic age, recognized as belonging to the A1 Aridh and Ibra formations by Glennie et al. (1974), consist of large angular blocks of shallow water Exotic-type limestones, up to 50 m across, and smaller rounded blocks of limestone and occasional lava set in a lime-mudstone matrix. These calcirudites occur as massive lenticular bodies, up to 150 m thick, and may be interbedded with cherts and redeposited fine grained Halobia-bearing limestones that contain small angular clasts of limestone, basalt and, rarely, quartz grains.

2I

These deposits are interpreted as a marginal facies of the Exotic limestones that formed on the margins of the carbonate platforms in relatively deep water. In most cases the original contacts of the Exotic limestones have been obscured either as a result of their redistribution as blocks in sedimentary melanges or by tectonic dismemberment during late Cretaceous thrusting; however, there are some cases where the primary stratigraphic relations are preserved. On Jebels Ghawil and Ajran and on several Exotics in the Hawasina window, notably Jebel Abiad, the limestones are conformably under and overlain by volcanic rocks, mostly submarine pillow lavas and tufts, belonging to the upper part of the Haybi Volcanics (Section 2.2.3). The volcanic rock/ limestone contacts are usually sharp and clearly primary, although in some cases they are marked by a few metres of sheared lava containing blocks of limestone. In some places there are small detached blocks of limestones within the volcanics which are undeformed and were clearly entrained within the lavas during eruptions. More rarely, as on Jebel Hamali in the Hawasina window, the massive limestones overlie fine-grained cherts and pelagic limestones and no volcanic rocks are present. On Jebel Kawr, the Upper Triassic limestones are conformably overlain by 30 m of Jurassic bedded cherts belonging to the Nadan Formation (Glennie et al. 1974). The Upper Triassic Exotics are shallow water limestones most like "Bahamian-type" carbonate platforms with a predominance of "back-reef" facies. They appear to have formed as carbonate banks built largely on a volcanic substrate isolated from the O m a n continental margin (Fig. 2.5). Glennie et al. (1974) and Searle & Graham (1982) describe them as forming the cappings to volcanic ocean islands. However, the presence of clastic quartz grains in the limestones themselves or in the marginal rudite deposits (Graham 1980b) suggests that some were formed on or close to a continental baserhent. At the end of the Triassic, shallow water deposition on the "exotic" carbonate platforms off the O m a n margin ceased. It is probable that the drowning of the platforms was produced by tectonic subsidence and/or rising sea-level (Vail et al. 1977). Massive shallow water Permo-Triassic limestones, similar in many respects to the O m a n Exotics and in which the Norian Megalodontid-facies is particularly characteristic, occur throughout the Tethys belt in the Alps (Bossellini & Rossi 1974), the eastern Mediterranean (Blumental 1963; Robertson & Woodcock 1979) and the Zagros mountains of southern Iran (Stocklin 1974; Hallam 1976). In many cases, as in Oman, the limestones occur as allochthonous masses associated with deep-water sediments and volcanic rocks. They are interpreted as having formed by the break-up of "external" carbonate platforms on rifted continental margins (Bernoulli & Jenkyns 1974) or as the cappings to ocean islands (Glennie et al. 1974; Hallam 1976; Searle & G r a h a m 1982). Pelagic sediments of Upper Triassic age occur in the lower parts of the Halfa and Haliw formations (Glennie et al. 1974) (Fig. 2.2). They consist of fine-grained redeposited limestones and cherts that were deposited in areas remote from clastic sediment sources. They occur as thin thrust slices or as blocks in melanges both tectonically below and above the Exotics. Where they are best developed they are up to 30 m thick and may be deposited on a basement of pillow lavas. The Upper Triassic cherts are identified by the presence of such Radiolaria as Veghicyclia sp., Gorgansium sp., Temphium sp., Sarla sp., Conosphaera sp., Capruchospaera sp. and Canoptium sp. (E. A. Pessagno, pers. comm.).

Chapter 2

22

UPPER TRIASSIC EXOTIC JEBEL QAMAR TYPE EXOTIC 1~ I " " - _ ~ P

" " I Masswe to well-bedded limestones and dolomJles

U~

Lower Palaeozoic basement

[ Alablan Platform

r

shelf slope facies carbonates -~.

I Megal hmestones OdOntid

[ : PERMIAN EXOTIC

-~~~-~-~_.=__-

li!ii

= : .

] JI I I I ~ _

Bas,n fac,~,s s(?d,m~,nts t Hawaslna)

Contmenlal C r u s t -

Ivlasslvr ~,rt~ff

[ [.

Ptllow basalts

\ \

~

See F~g 27 for

~ ~ ""/" ~ - - - _ . . _

sea lewwL el

Triassic volcanlcs t Hayb,) ~ -~

TransitJonalCrtJsl

Fig. 2.5. Schematic reconstruction of the palaeotectonic setting of the Oman Exotics on the Oman passive margin.

2.2.3 The Haybi Volcanics; Triassic volcanism at the opening of the Oman Tethys Glennie et al. (1974) showed the presence of lavas and some minor intrusive rocks associated with the Oman Exotics in the upper part of the Hawasina allochthon. They recognized that some of these rocks are petrographically quite distinct from the extrusive rocks of the overlying Semail ophiolite and included alkaline types, such as titanaugite-bearing "porphyrites" and "picrites", as well as altered spilitic pillow lavas, trachytes and rhyolites. They suggested that some of these volcanic rocks may be Permian or Triassic in age and perhaps related to the early rifting stage of the Hawasina Ocean. More specifically, the relations with the limestone Exotics were used as evidence that these formed as volcanic islands within an ocean basin. Searle et al. (1980), who called the thick sequences of volcanic rocks in the Haybi complex the Haybi Volcanics, described a wide range of lava types, from alkaline to tholeiitic in composition, and reported mid to late Triassic isotopic ages on the alkaline lavas. They demonstrated a sequence from alkaline rocks through transitional basalts to tholeiites and related this sequence to the development of some present day ocean islands. The Haybi Volcanics are here distinguished from later, alkaline, sills which range in age from Jurassic to mid-Cretaceous and were emplaced into the Hawasina and Haybi complexes during a later stage of the development of the Oman Tethys. Early Mesozoic volcanic rocks, of comparable age and type to the Haybi Volcanics and related to the early stages of continental rifting and the development of ocean basins, occur throughout the eastern Tethys region from the eastern Mediterranean to Oman. They occur in Yugoslavia and central Greece; for example, in the Othris area (the Agrilia Formation (Hynes 1974)), and also in the Pindos area (Terry 1971). The Mamonia complex of SW Cyprus contains alkaline volcanics interbedded with Triassic limestones (Bear 1958: Rocci and Lapierre 1969; Pearce 1975; Lapierre and Rocci 1976), which Robertson and Woodcock (1979) interpret as a rifted conti-

nental margin sequence. In southern Turkey above the Antalya ophiolite nappe there is a thrust sheet of alkali volcanics associated with Upper Triassic sediments, known as the Alakir Cay Series (Marcoux 1970; Robertson & Woodcock 1981). In the Baer-Bassit area of NW Syria there are two sequences of Lower Mesozoic volcanic rocks; a lower group of Triassic tholeiites and, overlying them, alkaline rocks interbedded with mid-Jurassic marine sediments (Parrot 1977). Other, less well known, occurrences are in eastern Turkey (Perincek 1980) and in the Zagros ranges of southern-central Iran (Ricou et al. 1977). However, Arvin (1982) has shown that the lavas in Ricou's melange nappe beneath the Neyriz ophiolite have depositional contacts with shallow water Upper Cretaceous (Senonian) limestones. These rocks range from basalts (alkaline to tholeiitic in composition) to rhyolites (ignimbrites) and some of them are related by Arvin (op. cit.) to subduction in the Zagros area in the mid to late Cretaceous. A feature of many of the Triassic volcanic sequences is an association with both shallow water "exotic-type" limestones and deep water pelagic sediments and their structural location in complex thrust zones beneath late Cretaceous ophiolite nappes (Ricou 1976). The Haybi Volcanics occur in a number of small separate outcrops in northern Oman (Fig. 2.6) as well as further north in the Dibba Zone (Lippard et al. 1982). They are here divided into two groups: (i) a varied sequence of lavas and pyroclastic rocks ranging from ultramafic ankaramites to felsic trachytes that are markedly alkaline in composition, and (ii) basaltic lavas, with compositions ranging from mildly alkaline to tholeiitic.

2.2.3.1 Haybi alkaline volcanics The alkaline volcanics are best developed in the Sumeini area, in the north-western mountains, but also occur in small outcrops in Asjudi and in Wadis Jizi. Ahin and Hawasina (Fig. 2.6). The Sumeini succession consists of up to 400 m of

Evolution of the Oman Tethys

23

7 HAYBI (J. Aswad)

3

i_

W. JlZl

,,. ,~..,

~ ".o.

(KL) (UT)

-

.-,

8 4

A-T

J. AJRAN

--~T

2

......... ;f.f.f..
J. GHAWlL

1 ~ ~ T Halfa

Halfa Fm ASJUDI ~- ~ (KL)

SUMEINI

""':; f : : : , -

T

(218)

ii

6

W. AHIN (233) (223) (220) Base not (216) seen

A

~:

,'~-:,~~

9

W. BANI W. HAWASINA UMAR (Shakhbut) K " (TI(Th)9

Base not seen

A

Nephelinite

s

!;'."d

~

y///J.//~..

I I :H Exotic Limestones (mainly U. Triassic Cherts (Halfa Fm) Lava/limestone conglomerates Pillow breccias/tuffs Pillow basalts Microsyen~te sill Massive basalt lavas I

[

I

I

Massive "greenstones" with occasional pillow structure

I ~ ~ Graniteblocks

,

.... .... ......

,

i

"(

in tuffs

r / / / / / / / f

( ) (Ma) KL UT

Fossil ages Thrust contact Radiometric ages Th Tholeiitic Lower Cretaceous T Alkali-transitional Upper Triassic A strongly alkalic

N T

-r

t

Trachytes Ankaramites

Ill

Dykes / Agglomerates

Haybi volcanics

~

SemailNappe ~

:A-T-:-.-:.:

: ::: :._-:

Major outcrops

~ ! . i::_ [:.-~' )( i~Zi H l i : )B : i L~ o_ c: ~K~- . :~i - ; ::-" V

,:'.-i'.:.-.~ -~

Major faults

1 Sumeini

6

Wadi Ahin

2 Jebel Ghawil

7

Haybi

3 Wadi Jizi

8 Wadi Bani Umar

4 Jebel Ajran

9

Wadi Hawasina

".\ 5 Asjudi

'~

,~t~! ~ SALAHI BLOCK : -

4

A-T

A

5

.,-,.

-,::,,,:

6

"

V

V Haybi volcanic blocks in melange

SARAMI BLOCK N .- ,;.j. -

Basalt types A Alkalic TTransitional Th Tholeiitic

'-"',--" . . ..v) . . .C.. ' ' ,b"-'., ,~, WUQBAH BLOCK ~ ~--.;~-__,~.,-',',;..-'_'.-!~.~ '-,. -~TL':,'_' ~-,,--":, ;,

~!.':,;,~-

r - T h ",',.,_

HAY LAYN

B,LooK :,'i, ,- ..-.~.:_

,

,::.,..

\\,-A-T:Th -,-,-

v

Fig.2.6.Stratigraphicsectionsand distributionof the HaybiVolcanicsin the northernOmanMountains.

\1

1\-_ _

/

24

Chapter 2

porphyritic ankaramites with some interbedded trachyte lavas. In the middle of the sequence there is a small body of nephelinite, which may be a lava flow. In the Asjudi section there are some 60 m of trachytic lavas overlain by up to 100 m of ankaramitic basalts. In the Jebel Ghawil and Wadi Hawasina sections a few metres of ankaramites occur at the base of continuous sequences with overlying alkaline to transitional basalts (Fig. 2.6). The ankaramites occur as lava flows, agglomerates, ruffs, dykes and sills. The flows are massive with blocky jointing and occasional scoriaceous glassy tops. They are interbedded with agglomerates containing angular lava blocks up to a metre across in a tuffaceous crystal-rich matrix. Some of the blocks are highly vesicular and glassy, although the majority are compact and holocrystallirie. The coarse pyroclastic beds grade locally into finer grained tufts. A tuffaceous agglomerate horizon in the Sumeini area contains 10-20 cm sized blocks of pink to red coloured, medium to coarse grained, alkali feldspar rich rock (albitites, syenite and albite granite). The volcanic rocks are cut by numerous thin (0.1-1 m thick) dykes and

Plate 2.1. Euhedral titanaugite phenocrysts in a Haybi Volcanics ankaramite; Wadi Ahin [PPL]. From Searle (1980).

2.2. Phlogopite phenocrysts in a Haybi Volcanics ankaramite" Wadi Ahin [PPL]. From Searle (1980).

Plate

irregular sheets of ankaramite similar in composition to the extrusives. These rarely form more than 5% of the total rock volume and the dykes have varying trends from predominantly NNE-SSW in Sumeini, to WNW-ESE in Asjudi and NNWSSE in Wadi Ahin. The ankaramites contain abundant (10-30% modal) fresh, black titanaugite phenocrysts, up to 15 mm across, that are pale green to brown coloured in thin section and show normal and sector zoning (Plate 2.1). Individual crystals show the complete range in composition from WosoEn3,~Fsll to Wo52En34Fs14 with Mg*(78-70) decreasing and TiO2(2.034.19cA) and Ai203 (5.48-9.42%) increasing from core to margin. Some of the flows, and more particularly the ankaramite dykes, contain large (<20ram) euhedral phenocrysts of bronze coloured Ba and Ti-rich phiogopite (Mg*74, TiO2 7.85, BaO 0.93) and, more rarely, smaller rounded phenocrysts of olivine, completely pseudomorphed by carbonate and serpentine, and, in one sample, brown kaersutite (Mg*42, TiO2 5.7%) (Plate 2.2). The rocks range from glassy to aphanitic types, in which small microphenocrysts of titanaugite, titano-

Evolution of the Oman Tethys magnetite, sphene and apatite are set in a brown almost isotropic iron-stained matrix, to fine grained, holocrystalline types with intergranular and intersertal textures composed of titanaugite, titanomagnetite, apatite, sphene and interstitial plagioclase (AnTo_~5). The plagioclase is turbid due to partial replacement by clays, iron oxides and carbonate. The nephelinite is a fine grained rock with 2-3 mm sized phenocrysts of altered nepheline that, although it is pseudomorphed by fine grained sericite and zeolite minerals, can be recognized by the characteristic hexagonal and square shapes (Plate 2.3). The rock contains microphenocrysts and groundmass grains of green aegirine-augite and sphene set in a fine grained matrix of secondary carbonate and zeolites. The trachytes are massive to autobrecciated felsic lavas containing small, 1-3 mm sized, white or pink coloured feldspar phenocrysts some of which have rounded shapes suggesting partial resorption during eruption. The phenocrysts are altered to fine grained sericite and carbonate. The groundmass is a fine grained cryptocrystalline "felsite" composed largely of altered feldspar, quartz, carbonate, magnetite, haematite and chlorite. Devitrification textures, such as spherulites and perlitic cracks, suggest an original glassy state for most of the rock. Some of the trachytes are brecciated, probably as a result of autobrecciation during eruption. Other more compact, coarsely crystalline rocks composed of altered euhedral feldspar laths and interstitial quartz may be parts of flows or, more likely, in cases where local cross-cutting relations occur, minor sills or plugs. A 30 m thick quartz microsyenite sill near Lihaban in Wadi Bani Umar cuts the upper part of the Haybi volcanic sequence. Blocks of red or pink coloured, medium to coarse grained, feldspar-rich rocks occur in some of the ankaramitic and trachytic tufts. Most of them are composed of an allotriomorphic-granular aggregate of turbid, iron-stained albite traversed by calcareous and ferrugineous veins, but less altered examples contain granular quartz and have hypidiomorphic or granophyric textures typical of acid plutonic rocks. They are most likely Na-metasomatized alkali granites or syenites co-magmatic with the late Triassic volcanics but the.y could conceivably be older granitoid basement rocks brought up as accidental xenoliths by the explosive alkaline volcanism. The localized occurrences of coarse pyroclastics, the rapid thickness variations and the local dyke swarms of varying

2.3. Zoned six-sided nepheline phenocrysts in a matrix containing small aegerine-augite crystals; Haybi Volcanics nephelinite from Wadi Sumeini [PPL] From Searle (1980).

Plate

25

trends coupled with a lack of pillow lavas or sub-aqueous tufts, point to the Haybi alkaline volcanics having formed largely from subaerial eruptions and in local, central-vent eruptive centres. 2.2.3.2 Basalts Fine grained basaltic pillow lavas are the commonest rock type in the Haybi Volcanics. They form a 700 m thick sequence on Jebel Aswad in the Haybi area and are particularly well developed around the Hawasina window especially in Wadis Bani Umar and Hawasina (Fig. 2.6). They also crop out in Wadi Jizi and on the west side of the mountains on both Jebel Ajran and on Jebei Ghawil where they overlie the alkaline rocks described in the preceding section. In the field the Haybi pillow basalts were mapped as a single unit, but subsequent geochemical studies led to the recognition of two distinct types; one of alkaline to transitional basalts and the other tholeiitic in composition (Searle et al. 1980). The basaltic eftusives comprise massive to pillowed flows, pillow breccias and occasional tufts. Sedimentary intercalations include conglomerate horizons containing a mixture of lava and shallow water limestone clasts and thin interbeds and interpillow infillings of fine grained lime-mudstone, cherty mudstone and red jasper. Large blocks of Upper Triassic Exotic limestone occur in the volcanics, often as massive blocks or lenses several tens of metres thick at or near the top of the lava sequence. On Jebel Abiad in Wadi Bani Umar a 20-30 m thick bed of massive limestone, that can be traced for over 3 km along strike, is conformably both under- and overlain by the lavas. Furthermore, on both Jebels Ghawil and Ajran the limestones are stratigraphically conformable with the volcanic rocks. Numerous smaller, metre-sized limestone masses are enclosed in the lavas and were apparently displaced from the site of deposition and slid into the volcanic pile during or between eruptions. The Haybi basalts show all gradations from massive flows to close-packed pillows and pillow breccias. The flows are generally thin, 5-10 m thick, and the pillows small (<1 m across) and rounded or bun-shaped, although there are some that are composed of larger tubes or bolsters up to several metres long. The lavas are mostly non-vesicular and predominantly aphyric

Chapter 2

26

Aswad) area, some 25 km to the NW of Wadi Hawasina, the Haybi lavas are more vesicular and are interbedded with considerable volumes of hyaloclastite pillow breccias and tufts. The pyroclastic beds are composed of whole and broken pillows in a tuffaceous matrix containing angular lithic fragments and devitrified glass shards. The lava fragments are a mixture of porphyritic and aphyric types and have a range of vesicularities from zero to over 50% by volume. In addition to these volcaniclastic deposits, there are interbeds of redeposited breccia and conglomerate containing a mixture of angular to rounded lava blocks and limestones in a sandy tuffaceous matrix. These beds are often matrix-supported and are interpreted as debris flows. The abundance of pillow breccias and conglomerates suggests that some of the eruptions occurred on steep, unstable slopes with eruption depths perhaps ranging from near sea-level to several hundred metres (Fig. 2.7). In summary, the majority of the Haybi basalts are submarine but appear to have formed at depths ranging from sea level down to over a kilometre. Locally pillow lava piles built up to sea level and shallow water limestones were deposited on them. Volcanism and contemporaneous limestone deposition occurred in an unstable environment on the edges of the carbonate platforms where slide blocks of limestone in volcanic rocks, mixed limestone/lava conglomerates and basaltic pillow breccias formed. Elsewhere quiet effusions of pillow lavas with thin interbeds of pelagic sediment took place in a deep water environment.

or sparsely porphyritic with olivine, clinopyroxene and plagioclase phenocrysts. The olivines are completely altered to chlorite and serpentine minerals but contain small relict spinel inclusions. The olivine-phyric types are rare, and also usually contain small plagioclase microphenocrysts. Where plagioclase is the dominant phenocryst, it may occur alone or be accompanied by clinopyroxene. The feldspar phenocrysts are altered to a mixture of sericite, clays, quartz and carbonate but the pyr0xenes may be fresh or only partly chloritized and are colourless low-Ti (TiO2 0.77-1.31%) augites (Wo44_4~En4w41 Fs15_~). The groundmass of the basalts is generally fine grained with variolitic and intersertal textures containing altered feldspar microlites with delicately preserved hollow and skeletal forms. Coarser grained varieties, with sub-ophitic textures, occur in the centres of large pillows and in massive flows and contain altered plagioclase, some fresh augite and opaque grains. In general, the original igneous textures are well preserved although the matrix is completely replaced by a fine grained mixture of clays, chlorite, quartz, haematite and, most abundant, carbonate. Veins and amygdales are infilled with carbonate. The metamorphic facies is sub-greenschist, probably zeolite facies, but no diagnostic minerals have been identified. The alteration effects are attributed to ocean floor hydrothermal metamorphism that occurred during and immediately after eruption in the late Triassic. In Wadi Hawasina, near Suri, there is a 200 m thick sequence of compact Haybi pillow lavas whose low vesicularity suggests quiet, relatively deep water lava effusions, probably at depths below 1 kin. On the other hand, in the Haybi (Jebel L a g o o n a l m e g a l o d o n t facies

sea level i 1

Massive shallow water limestones

1

1 ~

l~'-m~ I ~I "~., I I I [ I I 11 i I i I i 1 11 ! 1 I t ~ ,I I 1 I l-i I

I I I I

L I 1

I

Pillow brecc,as Limestone

I

I

I I 1,. I I 1 I ] I

'

-

,,

-

I

, ,

Coa res

I ~X.'~

,,,t_z , : ~

Interbedded

~--~"'..

~~

".-',.~-,-~" ,

,

~, ' " - : ' " - - ' ; :

'".""

"'"

~ C , . ' - '

m e s t o n e breccias .Pillowbreccias/tuffs

/

400

.,-,

massive and pillowbasalts

- Om

i ~

~Turbidite

grainstones

.'-" ' . . : ~ ~ ~ - ~ ~ " < 4 ~ _Mixed l i m e s t o n e / ~,~... :,:..._ ~ ~ . ' ~ ; ~ , _ " ~ lava b reccias ,,!=\-,.~, "=L_'~.,, . ~ ' ~ _ _ . ~ : ' -" -, . . . ~. . ~ - , " ' . ~ ~ --Halobla9 limestones : ,,.,.~z~ ~-w.,-,,~,, ~~.--------:-1~ and cherts / 600

Fig. 2.7. Schematic reconstruction of Upper Triassic Exotic Limestone - - Haybi Volcanics relationship.

2.2.3.3 Age of the Haybi Volcanics Potassium-argon (K-Ar) age determinations of separated biotires from the alkaline lavas and dykes in the lower part of the Haybi Volcanic sequence give mid to late Triassic ages ranging from 233 + 8 to 216 + 8 Ma (Table 2.2) (Searle et al. 1980; Lippard & Rex 1982). (A block from the Batinah Melange gives a younger age of 200 + 8 Ma but is suspect as this sample has a low K20 content which appears to be the result of slight chloritization of the biotite. Such alteration could cause loss of radiogenic argon and may account for the slightly discrepant age.) A mean age of 220 +5 Ma for the other five samples places them in the late Triassic, in good agreement with the

occurrence of Upper Triassic limestones in the overlying basalt lavas, and suggests that the whole Haybi Volcanic sequence is of late Triassic age. 2.2.3.4

Geochemistry

(i) Strongly alkaline volcanics Of all the Haybi volcanics, the ankaramites appear to be the least affected by secondary alteration. SiO2 varies from 38.7 to 44.6 wt % and they have higher TiOe (2.49-3.90%) and P205 (0.68-1.10%) contents reflecting the presence of titanaugite, sphene and apatite phenocrysts. The alkalis are relatively high with K20 > Na20 and they are both leucite and nepheline

E v o l u t i o n o f the O m a n

Tethys

27

Table 2.2. K-Ar age determinations of biotites in alkaline igneous rocks from the northern Oman mountains. Data from Lippard & Rex (1982).

Sample number

Location

Lat.

Long.

%K

1. 2. 3. 4. 5. 6. 7. 8. 9. 10. 11. 12. 13. 14.

Wadi Ahin Wadi Ahin Wadi Ahin Wadi Ahin Asjudi Wadi Jizi Wadi Hawasina Wadi Hawasina Jebel Ajran Jebel Ghawil Jebel Ghawil Jebel Ghawil Jebel Ghawil Dibba

24~

56~

23~ 24~ 23~

56~ 56~ 56~

24~ 24~

56~ 56~

25~

56~

6.80 7.25 7.64 7.02 7.09 6.61 5.29 5.42 6.92 6.51 6.88 6.13 5.96 7.47

OM4909a OM4909b OM4909c OM4909d OM1818 AB1 AB2 OM8625 OM8621 OM1790 OM1791 OM8603 OM8606 AE28

vol 4~ rad. (• 4cm~/gm-l) 0.6582 0.6688 0.6959 0.6273 0.6377 0.5328 0.3501 0.3502 0.3642 0.2383 0.2553 0.2262 0.2197 0.2853

%40rad.

Age (Ma) (+Error (o))

93.4 93.1 89.0 93.6 95.1 94.2 90.9 90.4 91.5 81.3 84.7 85.2 87.7 84.5

233 + 9 233 + 8 220 _+ 8 216 + 8 218 +_9 200 + 8 162 + 6 159 _+6 129 _+5 92 _+4 93 + 4 93 + 4 92 + 4 96 + 4

1-6 Biotite phenocrysts from ankaramite dykes and lavas, Haybi volcanics. (1-5 in situ, 6 block in Batinah Melange) 7-13 Biotites from alkali peridotite sills. 14 Biotite from crystal tufts, Aruma Group. Potassium and argon analyses by the techniques described by Rex & Dodson (1970). Analyst: D. C. Rex, Earth Sciences Department, Leeds University. Constants: kc = 0.584 x 10-1C~yr X~ = 4.72 x 10-1~ -l 4~ = 1.19 • 10 2

normative, although modal nepheline has been identified in only one rock. The presence of phlogopitic biotite phenocrysts with 7-8 wt % K 2 0 is an indication that the high potassium contents are primary. The rocks are markedly enriched in those incompatible trace elements (e.g. Ba, Sr, Zr, Nb, Th, U, Ta and Hf) which are typically abundant in alkali basalts (Harris 1957; Gast 1968). In addition, they have steep, light rare earth element-enriched chondrite-normalized R E E patterns (La/Yb(N) - 3 0 ) (Fig. 2.9), a typical feature of most within-plate alkaline rocks (Kay & Gast 1973). The Haybi trachytes and the Na-metasomatized syenite/ granite blocks have variable compositions (Table 2.3), but much of the variation, particularly in SiO 2 contents and Na/K ratios, can probably be attributed to secondary alteration. Despite this, they are clearly undersaturated and nepheline normative (up to 10% ne) although modal nepheline has not been identified. The high total alkali contents (11-12 wt %) and extreme enrichments of some incompatible trace element abundances; e.g. Zr and Nb which range up to over 200 ppm and 350 ppm respectively, are typical for felsic alkaline rocks. (ii) Alkali-transitional and tholeiitic basalts Analyses of the Haybi pillow basalts show the presence of two distinct chemical types, termed "transitional basalts" and "tholeiites" by Searle et al. (1980) (Table 2.4). On the basis of the field sampling, it appears that the former type is the more abundant; for example, they form the whole of the 700 m thick Jebel Aswad section in Wadi Haybi. The two types are interbedded in the Wadi Bani U m a r section in the Hawasina window where the tholeiites occur predominantly at the top of the sequence above the Jebel Abiad limestones. The basalts have in general suffered extensive secondary changes during sea-floor weathering and low-grade hydrothermal metamorphism. This has resulted in loss or gain of most of the major elements, and, as a result, normative calculations and oxide variation diagrams, such as the alkalis/silica and

FMA plots that are conventionally used to distinguish chemical types, are of little value. Most of the rocks plot in the alkali basalt field on the alkalis/silica diagram (Fig. 2.8a) mainly due to silica loss during carbonation and Na20 gain during albitization of the plagioclase, in contradiction to their transitional and tholeiitic character as revealed by the trace element plots and ratios (Fig. 2.8b & c). Selected elements, Ti, Zr, Y and Nb, which are believed to be relatively stable under conditions of low grade metamorphism, have been widely used to characterize altered basalts according to their original chemical type and tectonic environment of formation (Pearce & Cann 1973). The most common basalt type in Haybi pillow lavas has Zr contents of 90 to 170 ppm; TiO2 of 1.3-2.0 wt % and Y/Nb of 0.8-1.9. These values are typical for mildly alkaline to transitional "within-plate" basalts (Pearce & Cann 1973). The other basalt type has lower Zr (24-55 ppm) and TiO2 (0.6-1.2%) contents and is highly depleted in Nb giving Y/Nb ratios greater than 4. They are clearly tholeiitic in character. The differences extend to other incompatible elements, such as Ta, Th and U, as well as the R E E (Table 2.4, Fig. 2.9). The two types have different chondrite-normalized R E E patterns; the aklaki-transitional basalts are, in general, light REE-enriched (CeN/Yb N > 1.1), whereas the tholeiites have lower overall R E E abundances and are LREE-depleted (CeN/Yb TM <0.7). La shows somewhat anomalous enrichment in the two samples analysed (Fig. 2.9), probably due to severe alteration (Frey et al. 1974). 2.2.3.5 Tectonic setting of the Haybi volcanics

On the discrimination criteria of Pearce & Cann (1973) based on the stable elements Ti, Zr and Y, the alkaline rocks in the Haybi Volcanics are typical within-plate magmas. Such rocks are found in several tectonic settings, including continental rifts and on intra-plate oceanic islands. Nephelinites and Krich rocks, such as the biotite-bearing ankaramites in the

28

Chapter 2

H a y b i s e q u e n c e , are most c o m m o n in the continental setting (King 1970; C a r m i c h a e l et al. 1974), but do occur on m a n y ocean islands, e.g. the C a n a r i e s , Tahiti and Mauritius. Whereas the early alkaline lavas were p r o b a b l y e r u p t e d subaerially, the H a y b i pillow basalts are largely the products of submarine eruptions. T h e p r e s e n c e of i n t e r b e d d e d shallow water marine limestones and the local a b u n d a n c e s of pyroclastic and epiclastic rudites show that, after initial subsidence, s o m e - o f the volcanic piles built up to sea-level. The c o n f o r m a b l e relations

of the volcanic rocks and the Exotic limestones led both Glennie et al. (1974) and Searle et al. (1980) to propose that the H a y b i basalts f o r m e d the f o u n d a t i o n s to oceanic islands. On the Ti-Zr-Y and T i - Z r plots (Pearce & C a n n t973), the alkalitransitional basalts overlap the mid-ocean ridge basalt ( M O R B ) and within-plate basalt ( W P B ) fields (Fig. 2.10). T h e y are e n r i c h e d in most of the incompatible trace e l e m e n t s relative to N-type M O R B (Fig. 2.11) and resemble the " e n r i c h e d " or " E - t y p e " M O R B that is found in " h o t - s p o t "

Table 2.3. Whole-rock composition of Haybi alkaline volcanics. TRACHYTES

A N K A RA MI TES 0M4667

0M4668

0M4683

Si02 TiO2 A!203 Fe203 FeO MnO MgO CaO Na20 K20 P205 LOI

38.00 3.83 12.70 6.56 5.65 0.18 5.96 12.00 2.23 3.29 1.06 7.38

40.30 3.73 13.20 4.11 7.30 0.19 5.38 12.07 2.55 3.14 1.08 5.27

44.60 2.95 12.60 6.66 4.36 0.12 4.77 12.56 4.46 1.24 0,14 5.07

52.60 0.40 20.40 1.85 1.77 0.15 2.00 1.83 6.28 5.02 0.05 7.75

Total

98.84

98.32

99.53

100.10

Zr Y Nb Rb Sr Ba Hf Th Ta Sc Co La Ce Nd Sm Eu Gd Tb Tm Yb Lu

260 28 98 35 1059 1143

282 28 115 44 1024 1796 7.0 8.7 6.7 15.3 23.2 74.5 147.3 64.1 11.7 3.66 9.70 1.16 0.25 1.69 0.21

387 36 161 5 1023 742

757 37 254 96 776 181 13.15 21.7 13.0 0.15 4.4 135.8 (414.0) 221.0 (255.5) 55.1 (87.5) 8.1 (39.9) 2.18 (28.3)

(227.1) (170.3) (101.7) (57.6) (47.5) (35.1) (22.5) (7.35) (7.68) (6.18)

0M4666

0M4665

SYENITE

0M4699

0M4660

0M4682

55.50 0.45 20.10 1.81 2.01 0.19 1.11 0.97 4.71 7.48 0.06 5.36

58.80 0.91 17.10 5.05 0.30 0.03 0.36 1.52 4.86 7.89 0.57 1.79

61.00 0.04 19.60 2.61 0.13 0.11 0.14 0.74 7.10 4.80 0.10 2.15

57.80 (I.54 16.30 3.91 0.43 0.09 0.62 5.78 8.88 0.09 0.05 5.39

100.11

99.19

98.52

99.88

554 41 323 74 572 594

667 48 155 74 155 1211

2034 87 343 131 277 597

1485 9 72 2007 753

0.88 (16.9) 2.99 (13.6)

C I P W Norms

or ab an lc ne di ol wo mt hm il ap c

10.2

17.0

15.1 7.4 10.3 29.7 0.9

15.6 1.5 11.9 30.5 2.4

7.8 1.3 7.4 2.5

6.1 7.2 2.6

1.4 26.2 14.0

29.6 34.1 8.7

46.3 28.8 4.4

47.5 40.3 1.6

28.6 59.5 3.0

0.5 67.4 4.5

6.8 26.3

10.3

6.0

0.6

5.0 3.4

4.5

3.3

0.9 0.3 0.5

2.7

2.6

0.8 0.1 1.4

0.9 0.1 2.2

6.3 6.0 2.7 5.7 0.3

Major and traces Zr-Ba by XRF at Memorial University, Newfoundland. Traces Hf-Co and REEs by INAA at Open University. (Chondrite-normalized values of REEs given in brackets) CIPW norms based on analyses recalculated to 100% volatile-free.

5.1 0.7 1.4 0.9

1.3 O.7 2.2 0.1 0.2 1.6

8.4 0.1 3.9 1.0 0.1

Evolution of the Oman Tethvs ~3C~.C~3C~.~.3C3Cc,",

29

3C

~C ,C

0 '

~

~

r,-,

~I

('1

~",

u/",

hi

Z-O0

~',

~

~

~

3C

~

~.

"2',d ~ ~'i ~

~

cq

,-,

nc ~ ~-i ~

4 ~ .

--4"

~ ~

~

~

,.~ ~', ~ ~

~

~C

~

"1"

.-~ ~ ~ ~

~

t'*",

~

~ ~-i--/ ~

.-.f"

~|

~1

3C

%

~

.-2 ~

~

~

~ ~

~",

c,i ~ -'-=

~

~

,

~

~

~

~

~

>

~ C

~

~

~

~

~

~

~i ~

~

~

~

,

,~.~--

~

~

~

"T-

>--,

~_

~

~i~ ~

~ .

~ ~

~ .

.

.

.

.

.

.

.

.

~ ~

~

~,~

3C

.

.

.

.

~

~ ~

.3~

~.

.

~

~ =

.~.

~: ~

6.~

~

-.-

~

c_:=

.

tr~, ".C

~'~6 3C

~

~,c,=

._

.--

-~

~.~

0 0 0 0 ~ > ~ .

,r"4

~t-,9 3 c

~c

ur;

~.1

.,.c. ~

~

g

~

'

. ~ ~

<

~

M ,

m ~

~

,

% g ~

~

g

~

.

r~ ,

~

3c o~

~

. ,

~

~

~

~

q

~,~

~

~

1

~

,

E E

...~ ~

~ ~!

~

~

.~

,

~ : ~ ~

~

,

~,~

,,C

u,-,~

~'-,~,

r--,c

~

"5

~

u,-. ,c. r.--. e~ ,c. ,r-~ -~ ~i

~

F-

,~

~ u,",

r---

,,~

,,.C

,~,-, "4-

~',

,"- r---

-9..4

,,-,

~

~

~

,

.~

t-,"-,

'

~

~

9=

~

.~o?,

~-=--=

~ ~ < C< q _ ~J~ <

~ ~ ~ ,~ ,~ ,~ ,~ . . . .,~ ~

t,. e"

=

~

~

~

r"',

s

OC

3,C

C'4

~

r'",

ut",

r,",

~

',,,~

~

~

u",

~

~

~

~

,,C

3,C

NNNN

~ ' 0 0 0 0 0 0 0 --I"(",,I

s OC

e-.,

'~f "

~

. ~. --.

,,', ~,

"4 E

I'~

'~:

OC

"~"

~

~,,~ ~

~

,:3",

I' ~

--. ~! - -

~, ~., - -

c.!

,g.

~-, ,~ r-4 .c :m

~

~

~

::3",

r"',

o<

~

~C,

t"'~

~

~ ~

~,=-~

~:~r<06

~r
~

r

,,~

,.,.;

....

~

-

~-'-

::3".

r*",

9 e..t

L)

.~

.z ~

:.-3 - - :-5 ~ .

.

.

.

.

.

.

.

.

.

.

~

~

Cq

.

.

.

.

.

.

.

.

.

.

,...

~

~

"~

F_...,-.

,r-

tt

= 4,,,,"

,,~: , ~-'T ~ , . ,,,C_ ,',,C~ ("-| ,,~.~,c,..

,.,0

0'~

m ,4 i

[-

C'4

9 "

3[: ~

r

" ~d

~C: L~, Ip.-- ~

~'~ r~,

",E; OC

~;~;-.:~r-.i~

~ 3C

e~, O'~

~r"l ur

.=

e~", Lr]~ ~ ~ eq, ~f"

~

~..

~,-..~ ~

~

"'-

"-

.= r<

r,i~

t,.~

~

0 0 0 0 0 0 0 0

,C,,..

Chapter 2

3~

(c) (b)

4.11

100 --

(a) -

+

9

+ 9

9

O

9 O

80

9

9 o o,i

9 9

0+

0

~

5 m

+

-vo .

~

#

0000 0

z

oH"

Transitional basalts

19

I-

o e"T'O

+

9169

Alkali basalts

ALKALINE~

40

+ Tholeiites

c-~

-4o o,i

z

++ +

10-

0.05

50 Si0 2

~ 4

I 100

0

0.15

9 1000

Zr (ppm)

Fig. 2.8. Major and trace element plots for the Havbi basalts. Solid circles = alkali basalts, open circles = transitional basaits and crosses = tholeiites. (a) A l k a l i - Silica diagram with dividing'line after Mcdonald & Katsura (1964). (b) Nb/Y vs Zr/P20~ plot with fields after Floyd & Winchester (1975). (c) Y/Nb vs Zr plot.

(a) After Pearce and Cann (1973)

"Within-plate" basalts Field D Ocean floor basalts Field B Low K tholeiites Fields A & B ~I~ Calc-alkali basalts

Ti/100

100 .///

-k Ankaramites OM 639 OM 637 [] Trachyte OM 4666 | Alkali basalt OM 4674 9 Transitional basalts OM 4659 OM 567 OM 4650 9 Depleted tholeiites OM 580 OM 565

9

Y.3 Ankaramites

9 Alkali - transitional basalts

9 Depleted tholeiites

. Trachytes (lower figure only)

, " ~ Field of Jurasslc-Cretaceous alkali stlls

MORB Mid-ocean ridge basalts WPB Within-plate basalts lAB Island-arc basalts

(b) After Pearce (1980)

Ti 0 2 (wt %)

//

10

/

A A

..

.

...

/.

0

I I La Ce

I Nd

I

I

i

i

Sm Eu Gd Tb

i

1

I

Tm Yb Lu

Fig. 2.9. Rare-earth element plots for the Haybi Volcanics.

/

10

I

j

l

100

1000 Zr (ppm)

Fig. 2.10. Ti-Zr-Y and TiO2-Zr plots for the Haybi Volcanics.

E v o l u t i o n o f the O m a n areas such as Iceland and 45~ on the Mid-Atlantic Ridge (Tarney et al. 1980) and in geologically young oceanic areas such as the Red Sea-Gulf of Aden-Afar region (Treuil & Joron 1974). The tholeiites, by contrast, are at least as depleted in all the same elements as average MORB and, although some plot in the MORB field on Fig. 2.10, they most resemble island-arc tholeiites (IAT). In summary, the Haybi volcanics exhibit a wide range of magma types, from an undersaturated alkaline ankaramitetrachyte suite, to transitional basalts and tholeiites. All these lava types were erupted in a geologically short interval (<20 Ma?) in the late Triassic. The presence of mid-Triassic deep water sediments in the Hawasina basin and the association of Haybi Volcanics with both pelagic sediments and shallow water limestones, suggests that they formed as oceanic islands or platforms in a rifted deep water basin perhaps underlain by thinned continental to transitional oceanic crust. A present day analogy would be the late Cenozoic volcanism of the southern Red Sea and adjacent areas where alkaline, transitional and tholeiitic basalts have been erupted over the last 30 Ma (Gass 1970; Gass et al. 1973) in an area of rifted and attenuated continental crust during the early stages of ocean basin development. Mineralogical and geochemical studies of late Triassic radiolarities and mudstones associated with the Haybi Volcanics show a high terrigeneous input favouring an origin in a small ocean 100

+ Ankaramite (OM 4668) | Alkali basalt (OM 4674) O Transitional basalt (OM 4650) 9 Depleted tholeiite (OM 580/796)

10

MORB

U ncertai n-----~wl values (close to detection limits)

I

I

1

I

1

I

Sr Ba Th Ta Nb Ce (120)

(20) (02)(0.18)

(4)

I

1

I

[

I

P Zr Hf Sm Ti

I

[

Y Yb

(10) (0.1L;~k,)(90) (2.4) (3.3)(1.5%)(30)

(3.4)

I

I

Sc Cr (40) (250)

MORB values (ppm except where stated)

Fig. 2.11. "Spider" diagram comparing trace element contents of Haybi basic lavas with N-type MORB (Pearce 1980). Ankaramite is strongly enriched in all incompatible trace elements. The alkalitransitional basalts are enriched in the elements Sr-P. The tholeiites are depleted in most incompatible elements.

Tethys

3I

basin like the Red Sea rather than in a distal oceanic environment (A. H. F. Robertson pers. comm.). 2.3 The Oman Passive Margin

By the late Triassic the Oman Mountain area had developed into a passive NE-facing continental margin with well-defined shelf, slope and rise-basin depositional environments. The most distal sediments, represented in the Hawasina assemblage, show a change from dominantly terrigeneous-derived quartz turbidites in the late Triassic and early Jurassic to dominantly calciturbidite-chert deposition by the mid-Jurassic (Graham 1980a). Igneous activity was much reduced after the late Triassic, with only minor amounts of alkaline rocks emplaced into the outer part of the margin in the mid-late Jurassic to mid-Cretaceous. In the following section, the nature of the continental margin sediments is discussed, mainly in terms of sedimentary facies and environments, followed by a brief description of the late Mesozoic alkaline igneous activity.

2.3.1 Shelf and slope facies sediments

Uplift and erosion of the Arabian platform at the end of the Triassic was followed by subsidence and a resumption of shallow water limestone deposition in the shelf sequence autochthonous carbonates of the Sahtan Group (Glennie et al. 1974) (Fig. 2.12). The Lower Jurassic succession contains quartz arenites which indicate the uplift and rejuvenation of clastic source regions to the southwest and west (Glennie et al. 1974; Murris 1980). In the Middle and Upper Jurassic shallow water shelf limestones, lacking detrital terrigeneous material, were deposited. Following an erosional unconformity at the top of the Jurassic, Upper Tithonian to Valanginian porcellanitic lime-mudstones represent a regional deepening of the shelf area. By the Albian, shallow water conditions had returned and the Albian to Cenomanian is represented by upwardshoaling cycles of shallow water calcareous shales and rudistid limestones. The allochthonous Upper Triassic to Cenomanian Sumeini Group sediments (Fig. 2.12) were interpreted by Glennie et al. (1974) as outer shelf and continental slope deposits that formed to the northeast of the autochthonous shelf carbonate succession. The lower part of the sequence consists of Upper Triassic dolomitic limestones, calcareous shales and cherts with some conglomerate horizons. Some of the lowest units are of shallow water facies, including local reef limestones, whereas the higher units are deeper water deposits; calcarenites, wackestones and lime-mudstones with interbedded slump conglomerate beds composed largely of blocks of shallow water limestones derived from the shelf area. The overlying Jurassic to Cretaceous succession consists of decimetrebedded doiomitic lime-mudstones, turbiditic grainstones, intraformational slump conglomerates and some cherts. The dominant lime-mudstones are laminated with only a sparse fauna of radiolaria and pelagic crinoids. Intraformational unconformities, representing gaps in the succession where downslope slumping has occurred, are common throughout the sequence. The litho- and biofacies of the Sumeini Group sediments are compatible with sedimentation in moderately deep water conditions on the continental slope. The Sumeini Group has a total thickness of about 2000 m, compared to over 3000 m for the contemporaneous shelf sequence (Glennie et al. 1974).

Chapter 2

32 2.3.2 Rise and basin facies sediments; the Middle and Upper Hawasina

The Jurassic to mid-Cretaceous sediments of the Hawasina

Assemblage occur in the upper Guwayza, Sidr and Nayid formations, which together form the middle and upper parts of the Hamrat Duru Group, and in the Wahrah Formation which forms a tectonically higher thrust slice and was interpreted by Glennie et al. (1974) as a more distal equivalent of the Hamrat Duru Group (Fig. 2.12). All these units are composed dominantly of calciturbidites with interbedded shales and cherts. The upper part of the Guwayza Formation comprises up to 350 m of redeposited Middle Jurassic turbiditic limestones, including massive calcirudites, oolitic and pelletoidal packstones and grainstones and some lime-mudstones. Almost all the clasts are of shallow water origin and palaeocurrent directions, which show mainly eastward transport directions (Fig. 2.13), suggest that most of the material was derived from the contemporaneous carbonate shelf area to the west. The source of occasional reef limestone clasts, including coral-algal and calcisponge boundstones, is problematic but they may have been derived either from areas near the shelf edge. not presently exposed in the autochthonous or parautochthonous sequences, or from Permian or Upper Triassic reefs, such as the Permian Exotics, which could have formed off-margin highs that persisted into the mid-Jurassic. Oolitic limestones occur in the shelf sequence and provide a ready source for the oolites in the mid-Jurassic turbidites. A minor non-calcareous component includes quartz grains and iithic fragments, including schist and phyllite clasts, which could have been derived from areas of exposed pre-Permian basement or by erosion of the Triassic and Lower Jurassic sands. Graham (1980b) interEvents in shelf e v o l u t i o n

preted the Guwayza limestones in terms of five major lithofacies that are typical of the inner and middle parts of a complex submarine fan system (Fig. 2.14). In this model, based on Walker & Mutti (1973), the calcirudites and coarse grained calcarenites represent channelized deposits in the inner fan area. These grade outwards into fine grained calcarenites which represent downslope middle fan deposits. Interchannel pelagic deposition is represented by the lime-mudstones. The lower part of the Wahrah Formation, which is equivalent in age to the Guwayza, consists largely of fine grained calcarenites, calcilutites and calcareous shales and is interpreted as a distal facies probably formed in the outer fan area (Graham 1980b). The Guwayza limestones pass upwards into the generally finer grained sediments of the Sidr Formation which consists of up to 150 m of variably silicified limestones and cherts (Fig. 2.13). These Upper Jurassic to Lower Cretaceous deep water facies deposits are part of a regional deepening and sediment starvation event which occurred across the whole Oman continental margin at this time (Glennie et al. 1974). Only small amounts of sediment were transported into the Hawasina basin and, as a result, the limestones in the Sidr Formation are finer grained and thinner bedded than those in the Guwayza, although grading and turbidite features are well developed. Silicification is extensive and ranges in intensity from nodular beds to totally replaced porcellanitic cherts. They are interbedded with rare primary cherts that are, in general, more vitreous and thinner bedded than the replacement cherts. The primary cherts consist of whole and broken radiolarian tests with occasional sponge spicules in a near-isotropic haematitic clay matrix. Some of the cherts are graded with an upward

SHELF

Influx of clastics in d e e p - w a t e r basin -Shallow water

u-: _ : ;- -

<:_ ~ "-~_~ .~

carbonates

LOCI-IShoaling upward sequences

O ~ p_ W

L L7

TT.-j~i .---~ - ~ : 7 ~ ~ - ~ I

-----~

--

deepening

Clastic influx

,~

[K

I

RISE H a m r a t Duru Gp

:~176176176

BASIN Wahrah Fm

H a l i w , H a l f a Fms.

7:

t

-~ ~

~

" --~

--~=~- Z2

Regional

SLOPE

-l~!i:!i!-

i~

,~I:_\

.:~1

, ,-=~[! ~ k ~ _ ,

=

100m 7

:3 < F ~ C ~ - L - L ~

0J

/

sea level

km 10-

20

30-

C o n t i n e n t a l Crust

~ _.~

T r a n s i t i o n a l Crust

-

~

O c e a n i c Crust -

Fig. 2.12. Jurassic-Cretaceous sedimentary facies in the Oman Mountains. (Hypothetical cross-section is based on that of a modern passive margin).

Evolution of the Oman Tethys

33

HAMRAT DURU [,~ Inset map: OUtcroC~sn~fra?aMwg;i?asediments 0

X,~OHAR

% ~

~

HawasinaWindow

'ASINA WINDOW

Hamrat Duru~ Range ~ Secondary ~ chert .,

~

....p

Chert Shale Calcarenites

~

I ]

Fine

I

Coarse

Calcirudites

1

0-J

increase in clay content and decrease in radiolarian size suggesting that they too were deposited as distal turbidites. The Sidr cherts are overlain by the calciturbidites of the Nayid Formation (Glennie et al. 1974) (Fig. 2.13). The contact is gradational over several tens of metres and seems to be markedly diachronous from as old as Berriasian in the Hamrat Duru range south of the mountains, to as young as AlbianCenomanian in the Hawasina Window (Glennie et al. op. cit.). As a result, the Nayid limestones are more than twice as thick in the Hamrat Duru area compared to the Hawasina. The Nayid Formation consists largely of thin bedded pelletoidal and lithoclastic packstones, wackestones and lime-mudstones. The oolitic grainstones that are typical of the Jurassic Guwayza calciturbidites are absent reflecting a change in the nature of the corresponding platform sequence. Complete Bouma sequences are common with typical turbiditic internal and bottom structures, including abundant trace fossils that indicate deposition in moderately deep water (Glennie et al. 1974). There are a n u m b e r of coarse conglomerate beds interbedded with the fine grained calciturbidites that are composed largely of blocks of shallow water limestones. These may correlate with similar conglomerates of the same age in the Sumeini Group and are interpreted as having formed by the

Fig. 2.13. Stratigraphy of the upper Hawasina sediments (Hamrat Duru Group) in the Hawasina Window and Hamrat Duru Range. Rose diagram shows palaeo current directions taken in both areas (data from Glennie et al. (1974) and Graham (1980a,b)).

slumping of shelf-derived material down the continental slope and into deep water (Glennie et al. 1974; G r a h a m 1980a). Thin sequences of outer margin, basin-plain facies, largely pelagic sediments belong to the Haliw and Halfa Formations of Giennie et al. (1974) and occur in the upper part of the Hawasina Assemblage and in the Haybi complex. These sediments are rarely more than 50 m thick and are usually highly disturbed by folding and faulting. In Wadi Jizi, near Khan, an unusually complete, c. 150 m thick, sequence of pelagic limestones and cherts overlies the late Triassic Haybi basalts. Although they are cut by numerous thrusts, the sediments show a consistent upward younging sequence, based on radiolaria and pelagic bivalve faunas, from the Upper Triassic to the Lower Cretaceous (Valanginian-Hauterivian) (E.A. Pessagno pers comm.)*. These are the most distal facies sediments found in the Hawasina sequences and are largely millimetrebedded radiolarian cherts with occasional interbeds of thin * Lower Cretaceous forms that have been identified by Pessagno include Pantellum riedeli, Sethocapsa trachystraca, Xitus sp., Dictyomitra lacrimula, Obecapsula sp., Holocryptocanium sp., Archaeodictyomitra apiara, Parvicingula boesli, Tharnarla conica, Pseudodictymitra sp., Parvicingula citae, Parvicingula sp., Mirifusus sp. and Cecops septemporata.

Chapter 2

34 FACIES A

FACIES B

FACIES C

FACIES D

calcilutite. T h e y are the result of distal turbidite and pelagic s e d i m e n t a t i o n , p r o b a b l y at abyssal depths, and were interp r e t e d by G l e n n i e et al. (1974) as the deposits forming on the floor of the H a w a s i n a O c e a n .

/ o ooo! 0

~ 9 o

o

!~

o

o

o

2.3.3 Late Mesozoic alkaline igneous activity T h e Triassic to m i d - C r e t a c e o u s s e d i m e n t s and Triassic H a y b i Volcanics of the u p p e r part of the H a w a s i n a A s s e m b l a g e and the H a y b i complex contain a variety of intrusive igneous rocks that were a p p a r e n t l y e m p l a c e d in the Jurassic and C r e t a c e o u s , during passive c o n t i n e n t a l margin s e d i m e n t a t i o n . Most comm o n are sills which, in some of the d o m i n a n t l y chert sequences, m a y form up to 30% of the rock volume. Two types of intrusion can be recognized: fine grained basalts and dolerites; and coarse grained mafic-ultramafic sills c o m p o s e d of alkali gabbros and peridotites (Searle 1984). T h e fine to m e d i u m grained mafic sills are by far the more c o m m o n type. T h e y range from 0.5 to 30 m thick, have thin chilled aphanitic margins and cause n a r r o w zones of baking and discoloration of the s u r r o u n d i n g sediments. T h e i r intersertal to sub-ophitic textures are c o m p o s e d of colourless to pale green titaniferous augite and plagioclase. Fe-Ti o p a q u e oxides and trace biotite, apatite and sphene are usually present, although secondary alteration, particularly of the feldspar, is extensive. C a r b o n a t e veins and cavity infillings are c o m m o n . The thickest bodies are c o m p o s e d of coarse grained dolerite with large, up to 5 cm long. dark brown kaersutites ( M g ' 5 8 - 6 1 , T i O 2 6 . 1 % , Ba 0 . 2 9 % ) in a matrix of titanaugite, t i t a n o m a g n e t i t e and partly altered plagioclase (An55_45) which may be r i m m e d by clear s e c o n d a r y albite.

c::)9 c7

'

0

Calcirudites Matrix and

I

Granule and pebble calcirudites and calcarenites

pebblesupported

Dominantly calcarenites (Thinning upward cycles) coarse ~-fine

FACIES E Calcilutite-calcareous mudstones and shales deposited between turbidite units SHELFSLOPE

INNER FAN

MIDDLE FAN

OUTER FAN

FACIES

Debris flow

- ~ . Grain flow

and slump deposits

- ~ - Classic turbidity _ ~. Distal current deposition turbidites + pelagics

Fig. 2.14. Facies types and distribution in the Jurassic and Cretaceous calciturbidites of the Hawasina sediments (taken from Graham (1980b)).

Table 2.5. Alkali basalt, alkali gabbro and peridotite sills - - representive analyses. 1 0M7991

2 7992

3 7923

4 7933

5 7934

6 7935

SiO2 TiO2 A1203 Fe203* MnO MgO CaO Na20 K20 P205 LOI

49.41 2.55 16.86 11.23 0.13 4.46 5.31 5.07 3.27 0.84 4.9

43.88 3.20 13.14 13.64 0.45 7.83 12.17 3.40 0.45 0.86 9.5

48.60 51.42 1.74 2.42 16.57 16.30 1 0 . 4 1 10.00 0.17 0.08 7.85 7.73 9.20 4.68 4.10 5.42 0.09 1.65 0.32 0.67 4.9 7.3

49.69 2.31 16.34 9.92 0.13 6.32 7.55 5.56 0.50 0.72 8.5

49.19 53.70 48.30 2.73 1.51 1.56 16.50 13.70 15.50 10.91 1(I.22 10.65 0.14 0.09 0.11 10.01 4.13 2.62 3.67 7.14 7.09 4.42 6.94 7.06 0.59 0.11 0.14 0.79 0.72 0.39 6.4 1.9 6.2

Total

99.14

99.01

99.09

99.60

99.02

98.95

99.88

366 37 94 53 438

280 29 76 9 683

126 21 23 3 362

436 31 94 17 163

408 30 92 11 257

410 29 93 12 305

No

Zr Y Nb Rb Sr Ba Cr Ni

7 4662

8 4663

9 4792

10 1829

11 8611

12 4627

13 8626

14 8608

15 8625

16 4675

46.87 2.58 16.90 9.53 0.16 4.38 6.89 4.44 3.02 0.94 4.4

45.26 2.22 18.14 9.95 0.19 3.60 7.96 2.67 4.33 0.79 3.9

43.68 2.32 18.32 9.49 0.17 3.54 10.72 1.65 3.94 (/.83 3.3

42.69 2.18 13.77 111.94 0.17 10.99 10.27 1.96 1.74 0.64 4.1

40.37 3.30 9.25 11.30 0.16 11.15 13.92 1.65 0.66 5.9

39.26 38.95 1.80 1.96 6.98 5.55 1 1 . 5 5 13.02 0.16 0.17 21.33 23.33 10.09 6.83 1.10 0.77 0.93 1.01 0.53 0.46 4.85 6.5

38.66 1.80 7.10 13.27 0.19 22.74 8.39 0.36 0.66 0.47 6.6

99.46

100.14

98.51

98.17

99.01

97.80

99.16

98.90

99.41

133 22 44

121 24 30

454

321

340 37 108 56 721

303 36 138 18 886

228 28 75 55 919 690

186 115

361 118

304 35 145 102 1079 1918 14 16

246 31 65 4 304 357 242 98

187 25 55 28 475 752 1006 825

158 17 45 30 380 645 775 918

137 18 82 24 591 694 631 696

126

1-8 Alkali basalt/dolerites; 7991-2 Wadi Ahin; 7923-35 Wadi Jizi: 4662-3 Asjudi. 7992 & 7935 are chills. 9-11 Kaersutite gabbros; 4792 Wadi Ahin; 1829/8611 Jebel Ghawil. 12-13 Fine grained olivine gabbro/dolerites from margins of differentiated sills. 14-16 Wehrlites; 8608 Kaersutite-bearing; 8025 Poikilitic biotite: 4675 Equigranular. Analyses by XRF at Edinburgh 1-6; MUN 7-10, 12, 16, OU 11, 13-15. * Total Fe as F%O3. Major elements in wt. % oxides, trace elements in ppm.

Evolution of the Oman Tethys The mafic-ultramafic sills are less common than the mafic ones, but are highly distinctive rocks. They are medium to coarse grained and in many cases composed entirely of remarkably fresh poikilitic wehrlites and alkali pyroxenites with up to 40% unzoned, partly serpentinized olivines (0.2-2 ram) (Foss) and 30-50% euhedral, colourless to pale brown, sector-zoned diopside-titansalites (0.3-4 ram) (Wo49.sEn41.7Fss.s- Wo52.2En33.oFs14.8; A120 3 3.6-9.8%; TiO2 1.8-5.1%) (Lippard 1984). The olivines and clinopyroxenes are enclosed in large interstitial plates of phlogopitic biotite (Mg'75-78, TiO2 5.4--6.8%, Ba 0.2%0.96%) and kaersutitic hornblende (Mg'64-71, TiO2 1.6-3.7%, Ba 0.06-0.09%), each of which may reach a grain size of 3 cm. The texture of the rocks is typically "orthocumulate" with cumulus olivines and pyroxenes enclosed in intercumulus biotite and hornblende. The sills have narrow, 5-10 cm wide, chilled margins which, although somewhat altered, can be identified as minute grains of titanaugite, clinopyroxene and sphene in a carbonated altered aphanitic matrix. The largest body on Jebel Ghawil is about 35 m thick and is compositionally zoned with an ultramafic base passing upwards into olivine gabbro and then kaersutite gabbro which forms the top of the sill (Searle 1984). At the lower contact some 50 cm of the country rock basalt tuff is converted into a massive hornfels. The base of the sill consists of a chill a few centimetres wide overlain by a metre-thick equigranular, hypidiomorphic-texture, olivinerich zone containing small rounded olivines, pseudomorphed by serpentine, titanaugite, biotite, titanomagnetite, apatite and sphene.

N

JEBEL GHAWIL

l

GULF OF OMAN. scat OMAN

::::::::::::::::::::::: 1 km

35

Representative analyses of the alkali basalt and alkali gabbro-wehrlite sills are given on Table 2.5. The alkaline nature of the suite is indicated by the high alkali, TiO2, P205 and incompatible trace element contents. All the basic rocks are nepheline normative and plot in the "within-plate" basalt field on the Ti-Zr-Y diagram (Fig. 2.10). The alkali peridotites (pyroxenites and wehrlites) have high MgO, Ni and Cr contents, resulting from the abundance of olivine and calcic pyroxene, combined with relatively high contents of incompatible elements, such as Ti, P, K, Sr, Ba, Zr, Nb and REEs. The whole suite has steeply inclined, light REE-enriched patterns that are typical of alkaline rocks (Fig. 2.16) with total REEs increasing in the sequence peridotite-olivine gabbro-kaersutite gabbro. The compositions of the chilled margins of the peridotite sills and the evidence for crystal settling suggest that the parent magmas were of picritic composition with up to 30% entrained olivine crystals at the time of emplacement. A high volatile content in the magmas is indicated by the coarse grain size of the sills and the abundance of hydrous minerals. K-Ar age determinations on biotites from the ultramafic sills in three localities (Fig. 2.15) give ages of 160_+ 6Ma (mid-late Jurassic) (Wadi Hawasina), 129+5Ma (early Cretaceous) (Jebel Ajran) and 93-92_+4Ma (mid-Cretaceous) (Jebel Ghawil) (Table 2.2). From these data, none of which conflicts with the stratigraphic ages of the rocks intruded by the sills, it appears that the sills were intruded over a period of some 70 Ma (Lippard & Rex 1982). The youngest sills, from the Jebel Ghawil area, give Cenomanian-Turonian ages that are within error of the formation age of the Semail ophiolite. In the

i

Wadi Jizi JEBEL AJRAN Wadi Ahin

Semail ophiolite n a p p e ~ Haybi complex

~ - --:. . . . . .

~

~

Hawasina allochthon Autochthonous & paratochthonous carbonates

Key to inset maps

~

Serpentinized harzburgite (base of semail nappe)

~

Metamorphics

,,~,, 1

l

WADI

HAWASINA 9

._

_

Exotic limestones ii

~

Haybi volcanics

~

Hawasina sediments

~o

Alkali sills

"~ "

Bedding (normal, inverted, vertical) Foliation Thrust Fault

\

25Km I

Fig. 2.15. Geological setting of the dated alkali wehrlite sills (taken from Lippard 1984)

36

Chapter 2

Kaersutite gabbro [] Olivine gabbro -'1- Peridotite (alkali wehrlite)

100

10

succession, and the conglomerates and Globotruncana-bearing marls of the Coniacian-Campanian Muti Formation (Glennie et al. 1974). This Turonian depositional hiatus, known as the "Wasia-Aruma break", is of regional extent throughout the Arabian Gulf region (Murris 1980). Glennie et al. (op. cit.) suggested that the Muti conglomerates formed by the slumping of uplifted and eroded shelf carbonates into an intracratonic deep-water basin that formed on the former shelf area. The Muti Formation (c. 300 m thick) grades westwards into finer grained marls and shales of the Campanian Fiqa Formation (>1300 m (Glennie et al. 1974) which is nowhere exposed in the mountains but found in the subsurface to the west and southwest. The coarse grained Juweiza Formation (c. 3000 m) of late Campanian to early Maastrichtian age overlies the Fiqa beds and contains abundant Hawasina and Semail clasts (Glennie et al. op. cit.). The relationships and lithologies of these Aruma Group sediments were interpreted by Glennie et al. (1974) as indicating the deposits of a deep water ensialic basin or "foredeep" that migrated westwards diachronously in front of the advancing nappes of the Oman Mountains. 2.4.2 The Hawasina Melange

0

1

I

La Ce

I

Nd

I

I

I

I

Sm Eu Gd Tb

I

1

1

Tm Yb Lu

Fig. 2.16. Chondrite-normalized REE patterns for alkali peridotite and gabbro sills. From Lippard (1984). northern U.A.E. part of the mountains, in the Dibba zone, there are alkaline ankaramitic tufts, dated as 96 + 4 Ma (Lippard & Rex 1982), which contain blocks of alkali peridotite and syenite. These volcanics are interbedded with red radiolarian cherts and occur in a melange of probable mid-late Cretaceous age equivalent to the Aruma Group (Searle et al. 1983). It is possible that this Cenomanian alkaline igneous activity is related to tectonic instability of the Oman margin due to the onset of subduction and ocean basin closure prior to ophiolite obduction and nappe emplacement (see next section). The late Mesozoic alkaline igneous activity in the Oman Mountains is petrologically similar to the mid-late Triassic Haybi alkaline volcanics, but represents a later and quite distinct magmatic event. Incontrast to the early phase, which was related to rifting and continental break-up, the second and longer phase seems to have occurred largely during passive margin development. Alkaline rocks are presently found off some passive margins, such as the Tertiary to Recent volcanics of the Canary and Cape Verde island groups off northwest Africa which are formed of Cenozoic alkaline rocks built partly upon an older Mesozoic oceanic foundation (Robertson & Stillman 1979; Stillman et al. 1982).

2.4 Mid-late Cretaceous Syn-tectonic Events on the Continental Margin - a Prelude to Late Cretaceous Nappe Emplacement 2.4.1 The Aruma Group An important break in sedimentation occurred in the Arabian shelf sequence between the Cenomanian limestones of the Wasia Group, at the top of the autochthonous carbonate

The Hawasina Melange forms a discontinuous unit above the Hawasina Assemblage at the base of or imbricated within the Haybi complex. The boundaries are generally tectonic and thickness estimates are difficult because of tectonic disruption and thickening due to late Cretaceous thrusting and folding; however, Graham (1980a) suggests that locally it reaches 1500 m although thicknesses of 50 to 250 m are typical. The deposits range from chaotic "block-on-block" megabreccias with blocks ranging from metre- up to kilometre-size to olistostromes composed of smaller (<10 m) blocks contained in a sedimentary pelitic matrix. The matrix consists of individual beds 0.5 to 20 m (mean 3.5 m) thick and forming amalgamated units up to 70 m thick. They may be interbedded with normally bedded cherts or turbiditic sandstones. The blocks are largely of Hawasina sediments with a predominance of Upper Triassic to Lower Cretaceous cherts (Halfa, Haliw Formations) and Upper Triassic limestone breccias and fine grained Halobia limestones (AI Aridh Formation), together with late Triassic Haybi Volcanics and Permian and Upper Triassic Exotic limestones. The matrix of the olistostromes is a red or occasionally green shale or shaly siliceous mudstone which contains poorly preserved radiolaria. These have been identified successfully only in samples from the Jebel Qamar area of the Dibba Zone where a mid-Cretaceous (Albian-Cenomanian) fauna has been found containing Pseudodictyomitra vesalansis, Tharnarla sp., Strichomitra sp., Eucyrtis sp., Zifondium sp., Staurosphaera hindei and Stichomitra sp. (E. A. Pessagno pers. comm.). Some of the largest Exotic limestones (such as Jebel Harmali in Wadi Hawasina and Jebel Misht, Fig. 2.4) appear to occur as giant blocks ("olistoliths') in the melange. Beneath Jebel Misht there is an olistostrome sequence in which there is an upward increase in the size and content of limestone blocks and on top of which the Exotic itself appears to have slid, although the contact at the base of the limestone is sheared due to late Cretaceous thrusting. In general, as a result of tectonic deformation, the pelitic matrix of the melange is pervasively sheared and the blocks rotated but largely undeformed apart from brecciation at the margins. The Hawasina Melange is interpreted as a sedimentary deposit formed by debris flow and gravity sliding of blocks of outer margin and ocean basin sequences into a deep water and

Evolution of the Oman Tethys

"Hawasina O c e a n " , by stretching out and relocating each thrust sheet of the Hawasina so that the highest and most distal units lay progressively further to the NE of the shelf edge. From this they calculated the "half-width" of the ocean to have been between 450 and 1200 kin, the latter figure resulting from the fact that there are no direct correlations between the various Hawasina units and therefore considerable gaps must exist. On the other hand, it is equally likely that not all adjacent units are proximal-distal equivalents (Graham 1980b) and there may have been lateral facies variations along the margin. For example, from the general stacking order of the nappes, it is equally likely that the AI Ayn or Wahrah formations may overlie the Hamrat Duru, although where both are present the AI Ayn overlies the Wahrah despite the fact that it is made up of more proximal-type sediments. Graham (1980b) calculated from the extent and stacking order of the Sumeini and Hawasina nappes in the Hawasina window that these rocks were originally deposited over a cross-strike distance of about 370 km; that is the minimum distance from the shelf edge to where predominantly pelagic sediments were deposited on oceanic crust. Reconstructions are most reliable for the late Triassic (Fig. 2.17) where the maximum amount of information is available and take no account of subsequent JurassicCretaceous sea-floor spreading within the ocean basin to the NE. Evidence of any Jurassic-Cretaceous ocean floor is scanty and it may well have been largely lost by subduction during the mid-Cretaceous (see below). The southern "Afro-Arabian'" continental margin of the

tectonically unstable basin. It apparently formed prior to ophiolite obduction (there is a complete lack of any detritus that can be related to the Semail ophiolite) and may be the result of the initiation of northward-directed subduction in the mid-Cretaceous off the Oman continental margin (Graham 1980a; Searle & Graham 1982). Searle et al. (1983) consider the allochthonous sedimentary melange in the Jebel Qamar area of the Dibba Zone, which contains interbedded alkaline tufts dated at 96 _+4Ma (Cenomanian) that are not seen in the Oman sections, as a distal equivalent of the syn-tectonic Aruma Group.

2.5 T h e O m a n T e t h y s - a S u m m a r y The preceding section shows that the Oman area developed as a passive or Atlantic-type continental margin throughout most of the Mesozoic from the Permo-Triassic (260-230 Ma) to the mid-Cretaceous (c. 90 Ma). The margin faced towards the NE with the shelf edge lying beyond the present limits of the autochthonous basement outcrops (Fig. 2.17). The Triassic to Cenomanian Sumeini Group and Hawasina sediments formed on the continental slope and rise, whilst further to the NE lav marginal oceanic crust formed during Triassic rifting and composed of Haybi Volcanics, Upper Triassic Exotic limestones and a cover of Triassic to early Cretaceous pelagic sediments. Glennie et al. (1974) attempted a simple palinspastic reconstruction of the margin and the adjacent ocean basin, their

Ocean sands Carbonatebanks Ba sement horst,

~

/

Zone

I

) I

/

.~

\

/.J

k

Permian and ~ :\ Triassic Exotic \~ -\ Limestones ~ ,

Guwayza

,.:,..

@ Y/: ! \ . 1_~: . 7 \ . !i!::" .;~-/%z .'N,: - _- : ~ : : : ; . .

-~" J ~ 7

k

z2

U pp er Triassic Exotic Limestones

. , "'-- T . . . . ' . ' . ' . ' . . '

A, A n' IX, I:.X,,

.~ : : : ~ : . : . ~,,,~ : ~ 7-T"~,~{...-... .~/!7.,.,~.:~7.~i:.ii:5 HaybiVolcanics

$7 ~/'

~"

AI Ayn

V ~

\

~ \ %~,..~

O

37

O

Sumeini ~

,,

"Oman Embayment"

~'//

Basin ~

~

/~/

~"~ . : ~:.:..:!)-.7 ..

~

-,

~ i . . . : - : : - : . . v : . . ~ - ? "" :::!::: ::-: :-."i~.:..-,, -...'.....:... ..:.:.~ \ \

~

9.- ."..-..:.:=-.:v.~,\ .,,,

~ ', \

Autochthon outcrop

_ ~

(

x-

Predominant sediment transport direction Fig. 2.17. Schematic palaeogeographic reconstruction of the Oman continental margin.

0 \

\

"

1

100 km I

38

Chapter 2

Tethys ocean in the Middle East appears to have contained major irregularities, due to the geometry of the Triassic rifting, which has had important consequences to later collisional events in the late Cretaceous and Cenozoic. In the central and southeastern Oman Mountains facies changes and palaeocurrent directions in the continental margin sediments show a broadly NW-SE trending margin facing to the NE. In the northern mountains and the Dibba Zone in particular, there is evidence of a NE-SW trending and SE-facing margin (Searle et al. 1983) which may have formed along an offset rifted zone. This arrangement of NE-SW "strike-slip" and NW-SE "rifted" sectors produced a deep embayment in the southern margin of the Tethys, here called the "Oman Embayment" (Fig. 2.17). This feature has been largely responsible for the absence of continent-continent collision between Oman and the Makran, in contrast to the Zagros mountains where continental collision took place along the Zagros thrust zone possibly as early as late Cretaceous but certainly by the mid-Cenozoic. In the Oman Mountains passive margin sedimentation ceased in the Cenomanian to be followed by the deposition of Coniacian-Campanian deep water clastics and olistostromes (Aruma Group). These deposits indicate tectonic instability and rapid subsidence of the margin. This appears to have resulted initially in the mid-Cretaceous from partial subduction of the margin in a trench followed by the development of a foredeep formed by nappe loading at the outer part of the continental margin. The general lack of ophiolite detritus in the Aruma Group deposits, except in the upper part, and the predominance of Hawasina and slope and shelf carbonate debris, suggests that these sediments were largely derived from the continental margin. The high pressure (blueschist facies) metamorphism of the continental margin rocks near Muscat (Section 4.3.1) suggests downbending of the outer continental margin to c. 20 km depth which can only be easily explained by subduction processes (Lippard 1983). In addition, the ocean floor that originally lay immediately to the NE of the Oman margin was most likely late Triassic in age, formed soon after the original rifting, yet little or none of this remains, except possibly in the Haybi complex and protoliths of the high grade amphibolites in the Semail metamorphic sheet, suggesting that it was either largely removed by subduction or never formed. The mid-Cretaceous age of the Semail ophiolite, and its geochemistry (Chapter 3), suggest that it formed in a marginal basin setting above and to the north of a north to northeastward-dipping subduction zone which dipped away from the Oman continental margin (Pearce et al. 1981). The ophio-

lite nappe was detached by intra-oceanic thrusting in the Turonian (c. 90 Ma) possibly above the subduction plane (Searle & Malpas 1980), although it is more probable that a detachment surface developed along the junction between the new midCretaceous ocean floor and the older Triassic crust (Rothery 1982). Later Coniacian-Campanian nappe emplacement resulted from progressive underthrusting of the continental margin and the incorporation of the margin sediments in imbricated nappe complexes beneath the Semail Nappe, followed by uplift and the emplacement of nappes across the continental platform by gravity-driven processes (Chapter 4). There are regional, as well as local, tectonic considerations which suggest that northward-directed subduction away from the Oman continental margin is the most likely mechanism for ophiolite obduction and nappe emplacement in the Oman Mountains and the destruction of the Oman Tethys. In the first place~ the evolution of Tethys has been dominated throughout by the development of northward-dipping subduction zones, a process that is continuing today in the Aegean and Makran areas. Secondly, Smith (1971) and Dewey et al. (1973) show that the relative motions of Africa and Eurasia, based on Atlantic opening, require that rapid convergence occurred in the eastern Tethys between Oman and Iran-Afghanistan between 95 and 70 Ma. Thirdly, the emplacement of mid to late Cretaceous calc-alkaline magmatic arc complexes in central Iran points to the northward subduction of "Arabian" Tethys ocean floor beneath that area from this time onwards (Berberian et al. 1982). Subduction beneath southeastern Iran in the Makran region, has been continuous at least from the late Cretaceous to the present (McCall & Kidd 1982; Berberian et al. 1982), apparently without continental collision occurring. If the present convergence rate beneath Makran of 4 cm/a (Jacob & Quittmeyer 1979) is extrapolated back to 75 Ma then about 3000 km of ocean crust has been subducted beneath the Makran since that time. It is likely that the Makran subduction was initiated after nappe emplacement in the Oman Mountains. It is most probable that the Oman subduction system failed because of the inability of the Arabian continental crust to be subducted in the Semail subduction zone. A northward jumping of the subduction therefore occurred at about 70 Ma, although it is possible that, for a short period, both northward-dipping subduction zones operated simultaneously, as shown by McCall & Kidd (1982). As a result of these complex events, the Gulf of Oman appears to have survived as an unconsumed remnant of the late Mesozoic Tethys (Hutchinson et al. 1981).

Chapter 3 T h e S e m a i l Ophiolite

3.1 I n t r o d u c t i o n

southeast (Fig. 3.1). The sheet-like form of the nappe has been confirmed by a regional gravity survey (Shelton 1984) which shows that the highest gravity values (+ 100-150 mgal) occur at the eastern or northern limits of the present outcrops and that the nappe extends beneath most of the Batinah coastal plain probably with its tapering trailing edge close to the present Gulf of Oman coastline. Marked structural thinning and imbrication of the ophiolite along the western or southern edges of

The Semail Ophiolite, also referred to in the literature as the O m a n or N o r t h e r n O m a n Ophiolite, forms a huge thrust sheet or n a p p e , the Semail Nappe, that is over 600 km long, up to 150 km wide and between 5 and 10 km thick. It crops out over an area of about 20,000 km 2 along the length of the O m a n M o u n t a i n s from near D i b b a in the north to near Sur in the

I

I 58~ Sheeted dykes and extrusives

t Crustal Gabbros and cumulate peridotites ! sequence

i iiiiiiiiiiiiill

Tectonite peridotite = Mantle sequence Semail thrust (Base of ophiolite nappe)

.7.1 ,:,-i' "..99i',..'.,".

--....

.'."..'.].

Thrust fault (Barb on overthrust block) 25 ~ N m

',;..

High-angle faults (Tick on downthrow side, where known)

-.,..

Major faults along edges of ophiolite "blocks"

4

Numbered ophiolite block (see figure caption) Stratigraphic or uncertain contact of ophiloite

~176176176 ,,.. ,..

Eroded edge of ophiolite outcrop

'.

"!'i.".".

.....,~ +.* ,., :.. i.','-'~

:t iii- ,, t...: t

i9

'::i 0

I

50km

I

24~

: ~:<~

8

iiiiiiii)

' ': .........~

% ~'L.....

, [ ~

c<~:r

i

t;: ~ : ~ :i:::;~84~ : : : , :

,,..__..-,.. ,,,<;. :,,-

:: ::~::;!~)/

Fig. 3.1. The Semail Nappe and its division into 12 major blocks (areas in km 2 in parentheses): 1. Khor Pakkan (870), 2. Aswad (1700), 3. Pizh (2610), 4. West Jizi (590), 5. Salahi (940), 6. Sarami (560), 7. Wuqbah (1470), 8. Haylayn (1380), 9. Muqniyat (680), 10. Rustaq (600), 11. Bahlah (810), 12. Ibra (6740).

Chapter 3

40

Km O--

6

Volcanics Sheeted Dykes

2 1 9 . . -..,...

4

5--

(""": "

:::3, --X - - \

5

ilUi i i i il

Layered Series 9

Fig. 3.2. Column sections through the Semail ophiolite showing thickness variations for different blocks. 1. Fizh, 2. Salahi, 3. Sarami, 4. Haylayn, 5. Rustaq, 6. Ibra (Hopson et al. 1981). Maximum and minimum thicknesses for Layered Series shown, for the other units maxima only. Maximum thicknesses of total section is given in km at base of each column. Random dashes - - Mantle Sequence, horizontal l i n e s - Layered Series, stipple - - Sheeted Dyke Complex, v ' s - - Lower Volcanics, inverted v ' s - Upper Volcanics. (High level intrusives not included in these sections.)

'

222'N

Palaeomoho 10 . . . . ~--" "'/,Z'~.

(

. ,',', -" Y i

. '_", " - 5 . . . . .

? :Li

-:i'bY;: Mantle Sequence

~.~ ~., _ . " . ",-':3<

15 - - (~LL'-,'

'.i- "5 i'.,.'~

,-, .,. 'j,.,.,?~

(7.5)

". :_;'".-'i:" ,;~,"

(9)

(14) . . . . . .

-

20--

, ,,? J

- . - , , ,y - , y :

,

:..,,I"-. .,.,

'.\L2

-,,-,<,.

(11.5)

,

(18)

(16)

the outcrop suggest that the leading edge of the nappe during emplacement was not far beyond the present outcrop limits (Section 4.3). As a result of syn- and post-emplacement deformation, the Semail Nappe has been broken up into about a dozen major tectonic "blocks" of between 600 and 7000 km z in area (Fig. 3.1). Within each block the original internal stratigraphy of the ophiolite is more or less intact. The major blocks in the northern part of the mountains are separated by either (i) cross-strike fault zones, e.g. Wadis Ham, Hatta, Jizi and Ahin, some of which may have originated as sea-floor (transform?) faults, or (ii) by imbricate thrusts or reverse faults as in the Wadi Jizi sector. The occurrence of late Cretaceous tectonic melange along these fault zones shows that thinning and break-up of the nappe, largely by the upward protrusion of melange from beneath (Lippard et al, 1983; Robertson & Woodcock 1983b, Section 4.3.3.2), was a syn-emplacement event. South of Wadi Ahin, Glennie et al. (1974) recognized that the ophiolite blocks occur on either flank of the HawasinaJebel A k h d a r anticline and suggested that they formed as synclinal structures of mid-Cenozoic age. However, Graham (1980b) argued that the separation of the ophiolite blocks around the Hawasina Window was largely the result of late Cretaceous syn-emplacement break-up of the nappe. He showed that northward-directed thrusting on the northwestern edge of the window was the result of the ophiolite nappe sliding northwards off an uplift whilst at the same time on the opposite flank the general direction of movement was towards the southwest. The deeply incised relief and almost complete exposure of the ophiolitic rocks in the rugged mountainous terrain comprises elevations of 500 to 1500 m with occasional summit heights up to 1800 m. This relief, combined with the effects of tilting and folding of the nappe during and after emplacement, has provided complete stratigraphic sections through the ophiolite that are at least 16 km, and perhaps as much as 20 km, thick in the Fizh (Smewing 1980a) and Ibra (Hopson et al. 1981) blocks (Fig. 3.2). By contrast, gravity data (Shelton 1984) suggest that the maximum structural thickness of the nappe in the centres of the largest blocks is between 5 and 10 km and that the basal thrust is more or less flat-lying so that it

cuts across the often moderately to steeply dipping internal stratigraphy (see Section 4.3.3.3 for further discussion). 3.1.1 Internal subdivisions

The internal stratigraphy of the Semail ophiolite conforms in all respects with the classic Penrose Conference definition (Anon 1972, Table 3.1) of an ophiolite complex and comprises all ten units of Moores' (1982) expanded "ophiolite association". Table 3.2 shows the presently used subdivisions of the Semail ophiolite and compares these with those previously employed by Allemann & Peters (1972) and Glennie et al. (1974). The ophiolite is broadly divided by the Petrological M o h o into two major units: Table 3.1 For almost fifty years "ophiolite'" was used and abused until, under the unifying paradigm of plate tectonics, the participants of the GSA Penrose Conference provided the following definition (Anonymous 1972): "Ophiolite refers to a distinctive assemblage of mafic to ultramafic rocks, It should not be used as a rock name or as a lithologic unit in mapping. In a completely developed ophiolite the rock types occur in the following sequence, starting from the bottom and working up: 1 Ultramafic complex, consisting of variable proportions of harzburgite, Iherzolite, and dunite, usually with a metamorphic tectonite fabric (more or less serpentinized); 2 Gabbroic complex, ordinarily with cumulus textures commonly containing cumulus peridotites and pyroxenites and usually less deformed than the ultramafic complex; 3 Maficsheeted dyke complex; 4 Mafic volcanic complex, commonly pillowed. Associated rock types may include (i) an overlying sedimentary section typically including ribbon cherts, thin shale interbeds, and minor limestones; (ii) podiform bodies of chromite generally associated with dunite; and (iii) sodic felsic intrusive and extrusive rocks. Faulted contacts between mappable units are common. Whole sections may be missing. An ophiolite may be incomplete, dismembered, or metamorphosed ophiolite. Although ophiolite generally is interpreted to be oceanic crust and upper mantle, the use of the term should be independent of its supposed origin."

The Semail Ophiolite

4I

Table 3.2 The stratigraphic subdivisions of the Semail Ophiolite.

Allemann & Peters (1972)

Gabbro complex "Layered Zone" Peridotite and serpentinite complex

Max.

Glennie et al. (1974)

Extrusives and pelagic sediment (E) Diabase dyke swarm (D) Hypabyssal gabbroid unit (HG) Gabbros (G) Transition zone gabbros and peridotites (PG) Peridotites and serpentines (P) Semail Thrust Metamorphic Sheet

Metamorphics Oman Melange

This volume

thickness

Pelagic and ophiolitic sediments ~(Os) Extrusive sequence2 (El, Eu)

100 m 2000 m

Sheeted Dyke Complex (D) High-level Intrusives (G, Tr)

1500 m 600 m

[Late Intrusive Complexes (G', Tr')[ Layered Series (Cg, Cpg, Cp) (Petrological Moho) Mantle Sequence (MS)

q ,~

/

4000 m* 8-10000 m

(Banded Unit) [Biotite granites iGr)] Semail Thrust ] Metamorphic Sheet Haybi Complex Basal Serpentinite

l

Notes 1 Subdivided by Fleet & Robertson (1980) and Robertson & Woodcock (1983b} into Suhavlah and Zabvat Formations. 2 Subdivided by Alabaster et al. (1980, 1982) and Alabaster (1982) into Geotimes, Lasail, Alley, Cpx-o and Salahi Units * Pallister & Hopson (1981) record 5900 m of cumulate rocks in the Ibra area of the SE Mountains

(1) A lower "Mantle Sequence", so called because it is believed to represent the upper part of the sub-oceanic mantle (Allen 1975; Smewing 1980a), composed largely of variably serpentinized peridotites, mainly tectonized harzburgites with associated lherzolites and dunites. The maximum exposed thickness of the Semail Mantle Sequence is estimated to be 1012 km. (2) An upper "Crustal Sequence" consisting of a magmatic assemblage of cumulate peridotites and gabbros (the Layered Series) overlain by non-layered plutonic rocks (the High-level Intrusives), a Sheeted Dyke Complex and an Extrusive Sequence of lavas interbedded with and overlain by fine grained pelagic sediments (Suhaylah and Zabyat Formations). These units, up to and including parts of the lava sequence, are cut by several suites of cross-cutting plutonic and hypabyssal rocks collectively called Late Intrusive Complexes. Radiometric dating of plagiogranites from both the High-level Intrusives and Late Intrusive Complexes (Tilton et al. 1981) and the dating of radiolarian faunas from the pelagic sediments (Glennie et al. 1974; Tippit et al. 1981) show that the upper part of the Semail ophiolite represents a section through mid-Cretaceous, c. 95 Ma old, oceanic crust. Although it shows considerable variability, ranging from less than 4 km to locally as much as 9 km thick, the crustal section is on average about 6 km thick so that, unlike many other ophiolite complexes (Moores & Jackson 1974), it is generally of comparable thickness to normal oceanic crust (Raitt 1963; Christensen & Salisbury 1975) (Fig. 3.4). In addition, Christensen & Smewing (1981) showed that, after taking into account the fracturing and high porosities of the rocks, the seismic structure of the Semail ophiolite is similar to that of modern oceanic crust and upper mantle. That the Semail ophiolite represents an exceptionally well-preserved section through "normal" oceanic lithosphere probably formed at a spreading axis in an "open" ocean basin was the consensus opinion of the "Oman Ophiolite" volume published in the Journal of Geophysical Research (vol. 86, B4) by R. G. Coleman, C. A. Hopson and their co-workers. This work was largely based on a study of the ophiolite in the Ibra

area of the SE mountains. This view has been challenged by members of the Open University group working in the northern part of the mountains (Pearce et al. 1981; Alabaster et al. 1982) who, while recognizing that the main "axis" sequence of the ophiolite must have formed at a constructional plate margin, have suggested that the complex magmatic evolution and the geochemistry of the dykes and lavas are more compatible with its formation in a short-lived marginal basin above a subduction zone. 3.1.2 External relations of the Semail Nappe The base of the Semail Nappe is everywhere marked by a major thrust zone that separates the ophiolite and its basal metamorphic sheet from the underlying rocks. The lower part of the Mantle Sequence consists of up to 500 m of mylonitized harzburgites/lherzolites and dunites. Parts of these lithologies are markedly banded on a metre-scale and comprise the "Banded Unit" of Searle (1980). These rocks are more highly deformed than the rest of the Mantle Sequence and possess fine grained mylonitic layers that are superimposed on the earlier high temperature (>1000~ "asthenospheric" tectonic fabrics that formed beneath the spreading axis. The later "lithospheric" deformation occurred at moderate temperatures (800-1000~ and high shearing stresses during the early "intra-oceanic" stage of ophiolite detachment (Searle 1980; Searle & Malpas 1980; Boudier & Coleman 1981; Boudier et al. 1982). Searle & Malpas (op. tit.) suggested that the formation of the Banded Unit was contemporaneous with the upper amphibolite facies metamorphism at the top of the underlying Metamorphic Sheet. This Metamorphic Sheet forms a separate thrust slice attached to the base of the Semail Nappe and is composed of up to 500 m of amphibolite to greenschist facies rocks which show a marked decrease in metamorphic grade from the top (800+50~ to the base (c. 450~ of the sequence (Searle 1980; Searle & Malpas 1980; Ghent & Stout 1981). These metamorphic rocks are interpreted as having formed by the dynamothermal metamorphism of ocean floor

42

Chapter 3 PELAGIC SEDIMENTS

Km

---

0-

Zabyat gm Suhaylah Fm

EXTRUSIVE SEQUENCE

5 Salahi Unit 4 Cpx-~ Unit 3 Alley Unit 2 Lasail Unit 1 Geotimes Unit

P Picrite dykes R Rhyolite Lavas F Felsite sills

SHEETED DYKE COMPLEX ____-LATE INTRUSIVE COMPLEXES Plagiogranites ~

Isotropic gabbros

HIGH LEVEL INTRUSIVES

- - - - - Laminated gabbros

~ ~

~

.I. /

LAYERED SERIES Cumulus Layering 9 "~" ~ G a b b r o cumulates ~ Ultramafic cumulates

' '

.~.....

/

~

~

i i

PETROLOGICAL MOHO

MANTLE SEQUENCE

H Harzburgite Lher.zotite

OY<ES

/

/

lJ-/ / / /

" / /

..........

/ / / / / ~ ~ - ' ~ ~ / / ~ ~ ~ ~ - / ~ / ~ ~ _ ~ - " -

"C-D"2// / J ~ J ~ , , " " f / ~ : ~ / ~ X , ' ~ P x

' .................

Px Pyroxenite G N Gabbronorite

_~.~-Tectonic Foliation

: ............

: ::;:'~::'N;E~

' METAMORPHIC

i---L---z---z---z_--L'--z__-z_--=_-z_--i__ -z_-7_ ~ - 7 _ _ ~-=_----_._ ~I_GREENSOHISTS ~

SHEET

20 Fig. 3.3. Generalized section through the Semail ophiolite. and continental margin rocks that were overridden by the Semail Nappe during successive stages of its emplacement. Radiometric ages of the amphibolites are 96-85 Ma (Allemann & Peters 1972; Searle 1980; Lanphere 1981) which suggest that obduction commenced shortly (<10 Ma) after the formation of the ophiolite. The amphibolites and the lower part of the ophiolite are cut by sheets of deformed potassic granite, with a continental crustal geochemical and isotopic signature, that are dated at 85 Ma (Searle 1980). The underlying greenschist facies rocks, which are mainly of metasedimentary origin, are always separated from the amphibolites by a thrust and give younger radiometric ages (80-70 Ma, Lanphere 1981) than the amphibolites. It is suggested that they were formed from continental margin sediments that were attached to the base of the nappe at a later stage in the emplacement (obduction) process.

The underlying allochthonous and autochthonous rocks of the Hawasina series and Hajar Supergroup are largely unmetamorphosed except in the Muscat area where they suffered a late Cretaceous blueschist-greenschist facies metamorphism during nappe emplacement (Lippard 1983; Michard et al. 1984). In the northern Oman Mountains the thin sub-ophiolite metamorphic sheet is only discontinuously developed at the base of the Semaii Nappe and is sometimes represented by a tectonic melange composed of blocks of metamorphic rocks in a sheared serpentinized peridotite matrix, the latter clearly derived from the base of the overlying ophiolite nappe. All stages are seen, from small protrusions of serpentinite into the metamorphic sheet to its complete break-up and engulfment as isolated rafts in a serpentinite melange. The serpentinite

The Semail Ophiolite I Rhyolites

Pelagic sediments

J basalts 1 t (vesicula'

l

Basaltic

. ,~a'~~

2

3

4

.~ii!ii::::::!::i::i::i!::::ii::t" :.

Layer 1 2km/s

Extrusive Sequence

|

/

Sheeted dolerite

~ dykes lGabbr~176

5-

Layer 2

\

\

5 km/s

Sheeted Dyke Complex

b~,~--,o~JCo m plexe s -P2,-.->~ I

High-level Intrusives

Gabbro

..'"

~

mainly melange

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

WADI JlZl

Layer 3b 7.1 km/s

i!!!

k I I \ I I I 7 I I Vp km/s

.

! !/!!i!:: ~

-,,

/ / Perid ,_/~, otite //

.

56~176 E . . . . . . . . . . . ....

Ampiib~

"Seismic Moho"

'*Petrological Moho" Tectonite ectonite harzburgite ~rzburgite Mantle Sequence + dunite 85 I I

..:

[.

Layered S e r i e s ~ 6-

50'N

Ze i

Layer 3a 6.7km/s 4

WADI HATTA

5:6: '0'E

-

lite \

43

1

Layer 4

8.1 km/s

9 I

Fig. 3.4. Stratigraphic column (1), P-wave velocity profile (Christensen & Sinewing 1981) (2) and metamorphic facies (3) of the Semail

ophiolite compared to the seismic structure of oceanic crust (Raitt 1963). Details of the metamorphic facies overlap can be found in Section 3.10. melange has been called the "Basal Serpentinite" by Searle (1980) and contains thrust slices of weakly metamorphosed (sub-greenschist facies), but highly deformed, Haybi volcanics, Hawasina sediments and Exotic limestones in a complex tectonic unit known as the Haybi Complex (Searle 1980, Section 4.3.3.2). The Haybi Complex not only structurally underlies the Semail Nappe but also occupies cross-strike fault zones or "corridors", such as Wadi Jizi, between some of the ophiolite blocks (Fig. 3.5). In some of the corridors, notably in Wadi Sakhin just to the south of Wadi Ahin, it is continuous with the Batinah Melange that locally overlies the Semail Nappe along its eastern margin (Lippard et al. 1983; Robertson & Woodcock 1983b). In some places; for example, at Suhaylah in Wadi Jizi, the melange is in depositional contact with underlying pelagic sediments that in turn lie on top of the ophiolite (Woodcock & Robertson 1982a), elsewhere it has a tectonic contact. The Semail Nappe and this discontinuous melange cover are in turn tectonically overlain by several thrust slices of Hawasina-type sediments, the Batinah Sediment Sheets (Woodcock & Robertson 1982b), which form the upper part of Woodcock & Robertson's "Batinah Complex". On the flank of the mountains the Semail Nappe and the other allochthonous units are unconformably overlain by Maastrichtian sediments. In the sub-surface to the west of the mountains the late Campanian Juweiza Formation (Glennie et al. 1974) contains abundant clasts of igneous rocks probably derived from the Semail Nappe suggesting that it had been emplaced into its present position on the Arabian continental edge by that time. Coleman (1981) describes evidence for lateritic weathering of the ophiolite before Maastrichtian sedimentation in the Ibra area of the SE Mountains, and the same is found along the western flanks of the northern mountains in the United Arab Emirates (U.A.E.) (P. W. Skelton pers. comm.).

l

56 ~ 3 0!' ,E: , v ,. : . ~9.;.~. .i ~. .b - . o .",~.

r

9 : 9 :::

:i: ::

/

iiiii:, ~

;

:i! /

~

:: :

'

;

::): : :

;:::: :.'..

~ ,:~' ~ T ' a , ,o~

::) ':,1 '

Y

0

o"

,

.!!iiiiiiiii::::::: .........

..

10 km I

0

Corridor

l

I

SEMAIL NAPPE ~ Extrusive Sequence ~Sheeted Dyke Complex Layered Series.and High level Intrusive Complexes

~ ~

WADI AHIN

Normal stratigraphic contact I

High-angle fault

9

Thrust fault

Mantle Sequence

Fig. 3.5. Three melange-filled fault zones in the northern part of the Oman M o u n t a i n s - Wadis Hatta, Jizi and Ahin.

3.2 The Mantle Sequence At the base of all complete and coherent ophiolite complexes there is a unit of tectonized ultramafic rocks which, because it has petrological, geochemical and geophysical characteristics similar to those envisaged for oceanic upper mantle, is often referred to as the "Mantle Sequence". This term was first used by Allen (1975) to describe the tectonized peridotites at the centre of the Troodos Massif on Cyprus. The Mantle Sequence of the Semail Ophiolite (Smewing 1980a) comprises the lower half to two thirds of the Semail Nappe and consists of variably serpentinized tectonized peridotites, dominantly harzburgites that locally grade into lherzolites and have subordinate bodies of dunite, the whole being cut by numerous, but volumetrically insignificant, dykes, veins and irregular pods of ultramafic and mafic rock types (Fig. 3.3). Precise estimates of thickness are difficult to make

44

Chapter 3

because, although the harzburgites and dunites, the two major rock types in the sequence, can be readily identified in the field, they are irregularly disposed and no major tectonic breaks or repetitions of the sequence have so far been identified by their dislocation. The best estimates that can be made are by simply extrapolating the dip of the overlying crustal units, and in particular the plane of the Petrological Moho (Section 3.3), down section across the width of the Mantle Sequence outcrop. This method assumes no major repetitions of the sequence produced by thrusting, an assumption supported by detailed studies along several well-exposed wadi sections which have failed to locate major thrusts, and also that the dip remains constant across the section (a less certain criterion in areas affected by large-scale folding). Using these assumptions, variations in the maximum thickness of the Mantle Sequence ranging from 5-7 km in some of the smaller blocks (e.g. Rustaq (Browning 1982)) to 8-10 km in the large Fizh block (Rayy-Ragmi section (Smewing 1980a)) and 9-12 km in the Ibra-Wadi Tayin area (Boudier & Coleman 1981) are indicated. Near to the base of the sequence, particularly in the Rayy-Ragmi section, the peridotites are more lherzolitic in composition; this, although not a major occurrence, is petrologically significant and is described below as the "Basal Lherzolite". Irrespective of the thickness of the Mantle Sequence, the lowermost peridotites, for as much as 500 m above the basal thrust, have been more intensely deformed than the rest of the sequence. In places these highly deformed rocks have a markedly banded appearance due to metre-thick alternations of harzburgite and dunite and have been called the "Banded Unit" by Searle (1980). The origin of the Banded Unit and its role in the obduction process are discussed in Section 4.2.2. The Mantle Sequence forms 60-70% of the outcrop area of the Semail Nappe and gives rise to a distinctive dark rugged mountainous terrain (Plate 3.1) with summit heights of 10001500 m that is deeply dissected with numerous steep gullies feeding into a dendritic to rectangular drainage pattern of steep-sided, flat-bottomed wadis. The peaks and slopes are covered in a near in situ scree of weathered blocks and rubble so that massive outcrops are largely confined to the lower slopes and floors of the larger watercourses where good waterwashed exposures occur. Where fresh, the peridotites are pale

Plate 3.1. Typical rugged Mantle Sequence (harzburgite) landscape. Distant peak is - 1200 m above sea level and about 10 km distant.

grey-green rocks but they weather to various shades of brown with an iron-rich desert varnish on exposed outer surfaces. The sequence consists of 85-95% harzburgite (locally gradational into lherzolite) and 5-15% dunite. The harzburgites are medium to coarse grained rocks composed of 75-85% olivine and 15-25% orthopyroxene. Chrome spinel (chromite) is a ubiquitous accessory forming 0.5-2% of the mode. Clinopyroxene, although generally low in abundance (<1%), can, with concomitant decrease in orthopyroxene content, form up to 5% of the mode. Usually between 50% and 80% of the primary silicate minerals have been altered to mixtures of lizardite and chrysotile with accessory magnetite. Brucite is very occasionally found and chromite grains, although unaffected by the serpentinization, can display marginal oxidation. On weathered surfaces the harzburgites have a rough texture due to the more resistant nature of the orthopyroxene and are dark brown owing to the weathering colour of the serpentinized olivine, the orthopyroxenes are conspicuous as grey-green laths or "brassy" bastite plates up to 1 cm long and the chromite occurs as lustrous black euhedral or spindle-shaped crystals. Weathering is rarely more than a few centimetres deep and on fresh surfaces the rock is grey-green and the three mineral phases are less easily distinguished. The pervasive tectonite fabric (foliation) of the harzburgites is seen in outcrop as a planar alignment of the orthopyroxene and spinel grains. In areas of more intense deformation, a lineation, defined by elongate spinel grains, is evident. Although extremely homogeneous on a regional scale, at outcrop the harzburgites display a segregation banding sub-parallel to the foliation into olivine and orthopyroxene-rich bands that are usually between 5 mm and 5 cm and always less than 10 cm wide and have diffuse margins (Plate 3.2). These bands or ratio layers can rarely be traced for more than a few metres across an outcrop. The banding is bilaterally symmetrical, preferential mineral grading in any one direction does not occur; rather, the olivine and orthopyroxene-rich bands grade into one another. As this compositional banding is usually parallel to the foliation, it is most probably the result of metamorphic segregation and can be thought of as a gneissic banding (Dick & Sinton 1979; Browning 1982). The dunites are composed of >98% olivine and <2% red-

The Semail Ophiolite

45

Plate3.2. Orthopyroxene-rich and olivine-rich segregation layers in tectonized harzburgite. From Bartholomew (1983).

Plate 3.3. Irregular pod of (lighter weathering) dunite in tectonite harzburgite, base of Wadi Bani Kharus section. From Browning (1982).

Plate 3.4. 'Banded Unit" at the base of the Semail Nappe. Interlayered lighter coloured dunites and harzburgites. Height of cliff is about t50 m, the large dunite mass near the top is about 100 m across and 25 m thick.

brown chrome spinel grains which are conspicuous on weathered surfaces. The rocks generally have a light brown colour at outcrop and can be readily distinguished from the darker harzburgites (Plate 3.3). The dunite bodies occur sporadically throughout the Mantle Sequence and, although varying widely in shape and size, they always have sharp contacts with the

surrounding harzburgite. Locally they contain chromitite layers or pods containing up to 100% chrome spinel, but all gradations through chromitiferous dunites and olivine chromitites occur. The tectonite fabric found in the harzburgites generally continues into the dunite and is not deflected as it crosses the boundary. Indeed, although the size of the dunite

46

Chapter 3

bodies varies widely, from a few metres up to several hundreds of metres in width, their shape is directly related to the amount of deformation they suffered; when highly deformed, particularly towards the base of the nappe in the Banded Unit, they are elongated in the foliation direction (Plate 3.4).

3.2.1 Harzburgites In thin section the harzburgites can be seen to have a predominantly mosaic porphyroclastic texture (terminology after Harte 1976) grading into porphyroclastic to coarse porphyroclastic textures with large (<10 mm) strained and kinked orthopyroxene porphyroclasts showing very fine clinopyroxene exsolution lamellae parallel to (100). The porphyroclasts lie in a finer grained (<5 ram) mosaic of equant, polygonal olivine grains (Plate 3.5). Chrome spinel is present as small (<0.05 mm), scattered, eu-subhedral crystals and as larger (<2 mm) grains with typically cuspate boundaries (Plate 3.6). In thin section the chrome spinels are the only clear indicators of the orientation of the foliation plane. In some sections elongate grains occur in two preferred orientations; a few lying at a different orientation to the major alignment. Bartholomew (1983), who identified this phenomenon, interpreted it as indicating that some chrome spinel grains retained an alignment parallel to a relict earlier foliation plane whereas most had been rotated into a later plane. By studying three orthogonal sections from oriented field samples, Bartholomew (op. cit.) was able to define two foliation planes, which he termed D~ and D2, in a number of areas (Section 3.2.4). He noted that areas of weak D2 deformation are characterized by coarse porphyroclastic textures with euhedral or "holly-leaf" spinels, whereas areas of strong D 2 fabrics have more deformed mosaic porphyroclastic textures and the spinel grains are highly elongated. In the more strongly deformed specimens disrupted mosaic-porphyroclastic textures are present, the shape of the orthopyroxene porphyroclasts becoming markedly more tabular and often occurring as an aggregate of small grains rather than a single large one. The olivine mosaic becomes liner grained (<2 mm) and the spinels occur as discontinuous stringers of small anhedral grains with increased deformation. Olivine grain sizes and shapes vary considerably both within a single thin section and between specimens. The olivine grains vary from highly strained porphyroclasts (<5 mm) with markedly undulose extinction to recrystallized neoblasts (0.11 ram) with straight extinction. The grain shapes are largely controlled by the deformation history for, as deformation progressed, the olivines underwent a cycle of (i) crystallization with grain growth, (ii) crystal strain with the breakdown of large grains to a groundmass of small grains and (iii) recrvstallization of favourably oriented grains. So, in any one thin section, olivine grains at all stages of this cycle may be present. Harzburgite olivines are unzoned with a limited composition range FO~9.2-~)1.7 (mean Fo,~.~,). They contain 0-0.29'~ MnO, mostly in the range 0.11-0.16%, 0. ll-0.78% NiO. mostly around 0.4%, and <0.09% CaO, usually 0.01-0.03%. Although the possibility of systematic spatial trends in olivine compositions has been investigated (Brown 1983), none have been confirmed in the two traverses through the Mantle Sequence detailed below, except in the Basal Lherzolite at the base of the Ragmi-Rayy traverse where the mean Fo content decreases to Fogo.o. Orthopyroxenes in the harzburgites are colourless, unzoned enstatites. Brown (1982) maintains that, as with the associated olivines, the orthopyroxenes occur as both porphyroclasts and

neoblasts. He describes the porphyroclasts as being up to 12 mm long and usually tabular although more irregular "retort-shaped" grains occur. The cleavage is often distinctly curved, the extinction undulose and clinopyroxene exsolution lamellae along the (100) cleavage are ubiquitous (Plate 3.7). Although most of the lamellae are deformed, some exsolution blebs transgress deformation features suggesting that exsolution was, at least in part, post-deformational. Brown (1983) describes the orthopyroxene neoblasts as being usually 12 mm across and occurring as clusters of straight-sided polygonal grains. Their cleavage is undeformed, extinctions are straight and unstrained and they lack exsolution lamellae. In contrast, Bartholomew (1983) suggests that there are no orthopyroxene neoblasts and that, as the mineral deforms less readily than olivine and reacts in a more brittle manner to strain, the large porphyroclast grains are fractured to produce the smaller, equant grains. There is additional evidence suggesting that Bartholomew (1983) is correct in that whereas olivine neoblasts contain fluid inclusions indicating recrystallization, such inclusions are extremely rare in orthopyroxene grains. Harzburgitic orthopyroxene compositions are plotted on Fig. 3.6: only a few contain more than 5 mol% Wo and the complete composition range is En,~l.2_s3.a Fsm. 2 7.s W07.6-0.5. The compositions show a peak at Mg'gl.2 with the bulk of the analyses falling in the range Mg'9o.7_,~4 (Mg' = 100 Mg/(Mg + Fe) mol %). Minor elements show the following ranges: TiO2 0.02-0.18%; AI20~ 0.78-5.01%" Cr203 0.29-0.87%" MnO 0.08-4).17%; CaO 0.45-3.95% and Na20 0-4).13%. The only identified trend in composition is in the Basal Lherzolite where there is a concomitant decrease in Mg' (89.7-90.2) and an increase in the A1203 contents of the orthopyroxenes. Clinopyroxene has an average modal abundance of about ().5r in most harzburgites. It occurs as rare, ragged,

Wo 50

40

30

20

10

En / Fo

....1 i 95

~ l. 9'0

I i 85

I i 75

I i 65

~ Fs ~-- Fa

Fig. 3.6. Compositional fields of olivine, clmo- and orthopyroxenes from harzburgites. Coexisting ol-opx-cpx assemblages ./(fined bv tielines.

The Semail Ophiolite

Plate 3.5. Harzburgite showing clinopyroxene exsolution lamellae in kinkbanded orthopyroxene porphyroclast. Length of field of view - 3 mm [XN].

Plate 3.6. Chrome spinel with cuspate

boundaries in porphyroclastic textured harzburgite. Length of field of view 3 mm [XN].

Plate 3.7. Orthopyroxene porphyroclast

in harzburgite showing curved cleavage and clinopyroxene exsolution lamellae along the (100) plane, extending to the irregular margin of the porphyroclast. Length of field of view - I mm [XN].

47

48

Chapter 3

subanhedral grains up to 2 mm across, often with orthopyroxene exsolution lamellae. Weak compositional zoning is occasionally present. Strain features are similar to those in the orthopyroxenes and some recrystallization of porphyroclasts is seen. The mineral is a green chrome diopside with a normal composition range of En46.9_5o.5 Fs~.~_a.3 W045.7 50.4 (Fig. 3.6) (two exceptional analyses have more magnesian, less calcic compositions of En54.2_sa.s Fs3.2_4.7 Wo41.1_42.1). Clinopyroxene analyses from a single specimen can cover almost the whole of this range. Minor element oxides have the following ranges: TiO 2 0.04-0.3%" A1203 2.19-3.66%" Cr20~ 0.33-1.60%; MnO 0.06-0.13% and Na.O 0.02-0.20%. Compared to the associated orthopyroxenes, the coexisting clinopyroxenes in the harzburgites are richer in Cr20~, Na20, TiO2 and AI203 and poorer in MnO. The chrome spinels usually occur as small (<1 ram) grains that vary in colour from deep red to black in clinopyroxenepoor harzburgites to orange brown to yellow in the clinopyroxene-bearing harzburgites and lherzolites. Some large grains up to 2 mm in diameter have been identified near the top of the sequence. The mineral occurs in several forms: (i) as exsolution lamellae in enstatite; (ii) sub to anhedral, sometimes vermicular, grains at enstatite/olivine grain boundaries, (iii) sub to anhedral, "tadpole"-shaped grains in linear strings subparallel to the foliation and (iv) large euhedral or "holly-leaf" shaped grains. Some have pull-apart cracks indicating low temperature brittle fracturing. Within a single thin section a complete range of spinel textures from minute "wormy" to euhedral grains is commonly present. The wormy spinel is present throughout the sequence but the euhedral grains give way to more elongate types in areas of high strain. The hollyleaf variety could have formed by the deformation by pressure solution of large euhedral crystals (F. G. Christiansen pers. comm.) or have crystallized from droplets of chromite melt. Numerous chrome spinel analyses from the harzburgites are reported by Brown (1983) and Browning (1982). Browning (op. cic) suggests that the only internal zoning of the crystals is in respect of minor constituents such as MnO, TiO2 and Na20 but Roberts (1985) recognizes major element zoning in large spinels from the top of the sequence with A1 increasing at the expense of Cr towards the rims. Brown (op. cit.) shows that all the Semail harzburgite chrome spinels have low Fe ~+, falling close to the A1-Cr side of the AI-Cr-Fe 3+ triangle (Fig. 3.7), and on the Mg/(Mg+Fe) vs Cr/(Cr+A1) plot fall mostly within the Alpine peridotite spinel field of Irvine & Findlay (1972) (Fig. 3.7). Both Brown (1983) and Browning (1982) concentrate on the variations in spinel chemistry and stress the interdependence of Mg/(Mg+Fe) (range 0.44-0.76) and Cr/ (Cr+AI) (range 0.12-0.73) contents. TiO2 (0.03-0.17%) and MnO (0.09--0.25%) show no correlation with Cr/(Cr+Al) within their restricted ranges although MnO shows a negative correlation with Mg/(Mg+Fe) which may represent substitution of Mn for Fe z+ in the spinel lattice (Brown 1983). In the Basal Lherzolite unit at the base of the nappe, the paler coloured spinels are markedly more magnesian, aluminous and NiO rich with higher Mg/(Mg+Fe) (>0.7) and lower TiO2 and Cr/(Cr+AI) (<0.4) contents than those from the rest of the sequence. Variations in the mineral chemistry of the harzburgites have been investigated along two traverses - Ragmi-Rayy in the Fizh block and along Wadi Bani Kharus in the Rustaq block (Fig. 3.8). Along the 19 km Ragmi-Rayy traverse through an 8-10 km thick Mantle Sequence section, 27 samples were collected whilst along the 13.5 km long and 6-7 km thick Bani

Kharus Mantle Sequence traverse 10 samples were analysed. The results are plotted on Fig. 3.8 which also shows the analytical error bars and the composition range within a single specimen for comparison (data from Browning (1982) and Brown (1983)). There are two striking features of these traverses: (i) the general homogeneity and lack of variation throughout most of the Rayy-Ragmi traverse and the whole of the Wadi Bani Kharus traverse, and (ii) the marked change in composition in the lower 1-2 km of the Rayy-Ragmi section in the Basal Lherzolite (Section 3.2.1.1). Whole rock compositions of the harzburgites are given on Table 3.3. The low abundances of "magmaphilic" elements such as AI, Ca, Ti, Na and K, as compared with postulated pristine upper mantle compositions, indicate that the harzburgites are residua from which a melt fraction has been extracted by partial melting. One sample (OM1499) is of a fresh unserpentinized harzburgite, the remaining samples are serpentinized to a greater extent. The alteration may cause preferential loss of some elements, particularly Ca, and cannot be simply corrected for by recalculating to a 100% anhydrous composition (Coleman & Keith 1971). Nonetheless, the harzburgites have relatively uniform values of Mg' (90.0-91.0%), Cr203 (0.35-0.41%) and NiO (0.30-0.34%). Small variations in SiO2, MgO and FeO* are due largely to the varying modal proportions of olivine and orthopyroxene; for example, OM1499 is a relatively orthopyroxene-rich sample and has a correspondingly higher SiO2/MgO ratio than most. Despite the effects of serpentinization, the rocks in the Basal Lherzolite unit in the Rayy area clearly have higher A1203, CaO and TiO2 contents which are in accordance with their higher modal contents of clinopyroxene and the more aluminous and titaniferous nature of the chrome spinels. The highly depleted nature of the harzburgites in terms of incompatible trace elements is emphasized by REE analyses of two whole rocks (carried out by isotope dilution mass spectrometry) which have concentrations between 0.14 and 0.007 x chondritic. The chrondrite-normalized plots (Fig. 3.9) show "V-shaped" patterns similar to those that have been found for other ophiolite depleted peridotites (Frey 1970; Menzies 1976). An explanation of the origin of these patterns has recently been put forward by Prinzhofer & Allegre (1985). 3.2.1.1 The Basal lherzolite Some ophiolite mantle sequences are composed wholly or mostly of lherzolites (e.g. those of the western Mediterranean region (Beccaluva et al. 1984)) but the majority, and particularly those of the eastern Mediterranean and Middle East, are harzburgitic. In many complexes there are local lherzolite zones or patches in a dominantly harzburgite sequence, as described by Menzies & Allen (1974) for the Troodos and Othris (Greece) ophiolites. Indeed, Browning (1982) suggests that a 2 km thick band in the middle of the Wadi Bani Kharus section has a slightly more lherzolitic character. He recognized samples from this zone with 3 mm thick lenticles of clinopyroxene that also have more aluminous orthopyroxenes than the other peridotites but was unable to confirm whether this is a major occurrence or local lherzolitic patches. The Table Mountain section of the Bay of Islands ophiolite in Newfoundland has a 1 km thick basal lherzolite zone underlying approximately 10 km of harzburgites (Malpas 1978). Malpas (op. cit.) concludes that the Bay of Islands lherzolites represent only partially depleted upper mantle. Browning (1982) came to the same conclusion about sample OM9015 from the base of the

The Semail Ophiolite

49 AI

0.8

-

Fe 3 .

.

0.6Cr Cr + AI

0.4

0.2 Fe 3.

Fig. 3.7. Chrome spinels from harzburgites. Alpine peridotite field of Irvine & Findlay (1972) outlined.

0

I

I

1

0.8

0.6

0.4

Mg/Mg

Cr

+ Fe

(a) I

m

Km 0-

t

I

1

H

t--t

89

1

I

I

1

I

I

I

I

90

91

92

90

91

92

93

94

Fo

Mg' opx cpx

(b) Km O-

,.. qJ,o

j

9O0 9 99

95

I

I

I

1

I

I

I

1

1

2

4

0.4

0.6

0.8

0.1

0.3

0.5

0.7

AI203 opx 9 wt% cpx &

9 ,~

M9 chr Mg + Fe

Cr

chr

Cr + AI

f A

OOi~ .

J

l

6-

I

I

I

I

I

I

I

I

1

1

1

I

I

90

91

92

91

92

0

5

0.4

0.6

0.8

0.3

0.5

0.7

Fo

Mg' opx 9 cpx A

AI 2 0 3 opx 9

wt%

cpx ~

McJ__ chr Mg+ Fe

Cr chr Cr+ AI

Fig. 3.8. Mineral composition variations with depth in harzburgites along the (a) Rayy-Ragmi and (b) Wadi Bani Kharus traverses. Error bars based oil microprobe analysis of each parameter given at top. Points joined by horizontal tie-lines show compositional variation within one

Chapter 3

50 Table 3.3 Mantle Peridotites.

Harz burgites 0M1499

0M2264

Dunites

0M2265

0M9337

0M9304

0M9310

0M9556

0M3707

SiO 2 TiO2 AI203 FeO* MnO MgO CaO Cr203 NiO LOI

45.04 0.03 0.93 8.25 0.15 43.51 1.12 0.32 0.35 -0.12

40.01 0.03 0.75 7.29 0.12 40.87 0.62 0.36 0.30 10.05

40.27 nd 0.75 7.32 0.11 40.76 0.53 0.37 0.29 9.06

40.23 nd 1,07 9.46 0.14 38.60 0.70 0.37 0.26 8.48

41.58 nd 0.69 7.84 0.12 42.56 0.78 0.37 0.31 6.17

41.63 nd 0.57 7.30 0.12 43.06 0.49 0.38 0.31 6.69

36.11 nd 0.11 6.51 0.10 44.59 0.15 0.32 0.35 13.25

35.81 nd 0.17 8.54 0.14 43.40 0.16 0.68 0.23 11.60

Total

99.58

100.40

99.46

99.31

100.42

100.55

101.49

100.73

CIPW norm-t an di hy ol mt il

2.52 2.32 25.96 66.02 1.58 0.06

Co Cr Cu Ni S V Zn

111 2193 3 2766 863 22 29 (REE)

1.55 1.04 20.52 72.65 1.55 107 2453 nd 2335 769 35 38

0.03 0.39 0.97 95.65 1.40

103 2550 5 2286 621 38 35 (REE)

112 2528 44 2083 757 39 47

111 2535 13 2430 300 40 40

109 2621 8 2420 553 34 39

111 2150 3 2750 863 22 29

120 4650 7 1815 2310 28 32

*Total Fe as FeO. +FeO/Fe20~=9, calculated anhydrous. OM1499 OM2264 OM2265 OM9337 OM9304 OM9310 OM9556 OM3707

Fresh, unserpentinized harzburgite, Wadi Jizi (loose block) Harzburgite, Wadi Bani Kharus, 40 m below Petrological Moho. Harzburgite, Wadi Bani Kharus, 10 m below OM2264. Harzburgites, Wadi Ragmi. Dunite, AI Juwayf (discordant body near base of Mantle Sequence). Chromitiferous dunite, Wadi Ragmi (top of Mantle Sequence).

Rare earth elements (istope dilution technique at the Earth Sciences Dept. O.U.) OM1499 OM2265

La 0.026 (0.078) 0.0058 (0.018)

Ce 0.055 (0.064) 0.0115 (0.0134)

Nd 0.026 (0.041) 0.0059 (0.0094)

Sm 0.0051 (0.025) 0.0014 (0.0067)

sequence in the Rayy area for which he calculated a model whole rock analysis, based on the point-counted mode and the composition of the constituent phases olivine (73% modal), orthopyroxene (21%), clinopyroxene (3%) and chrome spinel (3%), of SiO 2 42.76, TiO2 0.04, AI20 3 2.84, Cr20 3 0.53, FeO* 8.89, M n O 0.15, NiO 0.27, MgO 43.22, C a O 1.28 and NaeO 0.03. Browning (op. cir.) concluded that, although it is richer in the basaltic components Ca, Ai and Ti than the typical highly depleted harzburgites, the rock is not sufficiently rich in these components to represent fully fertile mantle. Fig. 3.8(a) shows that in the lowermost 1-2 km of the Ragmi-Rayy traverse there is a change in the compositions of the mantle mineral phases towards less magnesian olivines (F089_9o) and orthopyroxenes (Engo), the latter having higher

Eu 0.0011 (0.014) 0.00063 (0.0081)

Gd 0.0068 (0.025) 0.0025 (0.009)

Dy 0.016 (0.0475) 0.0102 (0.03)

Er 0.019 (0.086) 0.012 (0.052)

Yb 0.031 (0.14) 0.027 (0.123)

AI20 3 contents ( > 4 % ) than in the rest of the sequence, and, most significantly, the chrome spinels become more magnesian and aluminous marked by a colour change from dark red-black to yellow-brown (Roberts 1985). There is also a concomitant increase in the modal clinopyroxene content from < 1 to > 3 % . Major element analyses (Table 3.4) confirm a progressive change from harzburgitic to lherzolitic compositions towards the base of the sequence. In particular, the lowermost sample (OM9013), which was collected within 500 m of the base, has markedly higher Ti, AI and Ca contents than all the other analysed samples of mantle peridotites from the Semail ophiolite. Spray (1984), in discussing the decoupling of ophiolites from their in situ oceanic position, suggests that they represent

The Semail Ophiolite

estimate the temperatures and pressures of partial melting or, at least, the conditions at which the residual mantle material was last in equilibrium with the melt. But disequilibrium and lower temperature and pressure sub-solidus re-equilibration make these 'geometers' unreliable. We include a summary of the data despite having doubts as to their usefulness and accuracy. Calculated geothermometric and geobarometric data for the Semail Mantle Sequence, taken from Browning (1982), is given in Table 3.5. Individual orthopyroxene compositions, even from a single sample, produce a wide range of pressures (5-53 kb) and temperatures (900-5600~ These variations are far in excess of the likely errors (Mercier 1980), but not surprising in view of the known heterogeneity of the orthopyroxenes with respect to CaO, A1203 and Cr203 contents. Single clinopyroxene calculations give a similar wide range of values (844--1207~ 5.0-38.3 kb), although the core of a large clinopyroxene grain from the lherzolite (OM9015) gives 1207~ and 38.3kb, whereas the recrystallized groundmass grains in the same specimen give 844-906~ and 5-11.8kb. The results of two-pyroxene thermobarometry on the two most cpx-rich rocks are also given on Table 3.5. It appears that there is no correlation between the temperatures and pressures obtained by these methods and the present depth of the samples within the Mantle Sequence. Browning (1982) notes that on a P-T plot most of the data points fall on the oceanic geotherm of Mercier (1980). The olivine-spinel geothermometer gives 700-850~ using the method of Fabries (1979) and 650-800~ using that of Roeder et al. (1979). The lherzolite (OM9015) gives up to 1000~ using the latter calibration mostly because the freeenergy values adopted by Roeder will give higher values for Al-rich rocks. These relatively low temperatures are clearly the results of sub-solidus re-equilibration (Browning 1982; Brown 1982).

0.1

0.01

I

I

La Ce

I Nd

I

I

I

Sm Eu Gd

1

I

I

Dy

Er

Yb

Fig. 3.9. REE/chrondrite plots for two Semail harzburgites. relatively newly-formed oceanic lithosphere and that the plane of the peridotite solidus (c. 1200~ between depleted harzburgite above and less depleted lherzolite below is a mechanically weak horizon along which decoupling is likely to occur (see Section 4.3 on emplacement mechanisms). 3.2.1.2

5I

Geothermometry and geobarometry

The Mantle Sequence rocks contain a number of potential geothermometers and geobarometers which have been used to Table 3.4 Traverse through the Basal Lherzolite, Wadi Rayy. 0M9026

0M9022

0M9020

0M9016

0M9013

m* SiO2 TiO2 A1203 FeO* MnO MgO CaO Cr203 NiO LOI

c.2000 40.80 nd 1.03 7.67 0.12 40.20 1.07 0.39 0.35 8.42

1750-2000 40.42 nd (0.004) 1.08 7.56 0.13 40.57 0.89 0.33 0.31 9.20

c. 1 5 0 0 42.09 nd 1.31 7.59 0.13 39.16 1.12 0.37 0.29 8.06

750-10(10 41.40 nd (0.013) 1.72 7.69 0.13 39.38 1.75 0.49 0.29 7.17

< 500 40.97 0.04 (0.047) 2.16 7.49 0.13 37.70 2.36 0.37 0.28 7.48

Total

100.05

100.49

100.12

100.02

98.98

5770 8357 2691 2400

24 6300 7000 2231 2415

7518 8714 2558 2294

81 9794 13428 3327 2295

280 12494 18428 2554 2236

ppm$

Ti A1 Ca Cr Ni

m* estimated height of sample above basal thrust in metres. TiO 2 contents: nd--not detected by major element analysis (values in brackets based on recalculated trace element analysis, values given in lower table). -~ total Fe as FeO. $ parts per million, elements recalculated to 100% anhydrous.

Chapter 3

52

Table 3.5. Geothermometry and geobarometry of the Mantle Sequence.

1. Single pyroxene

Mercier (1980) Orthopyroxene

Sample number

Depth*

P(kb)

OM2307 2262 2538 6676 2306 2540 2297 2309 2315 2317 2320 2328 2333 2339 2342 9015

0.0 0.0 0.1 0.2 0.3 0.5 0.6 0.9 2.0 2.6 3.3 4.0 4.4 5.1 5.5 9.0

26.3-28.6 19.8-28.0 16.2-24.7 15.6-26.3 24.8-40.5 16.9-24.0 12.7-23.5 22.3-29.8 17.8-24.7 10.6-47.2 15.0-30.4 20.4-24.7 20.5-25.2 12.6-27.0 4.4-21.9 5.9-52.7

2. Two pyroxenes

Wells (1977)

OM6676

888~

9015

865~

T(~ 1227-1250 1060-1210 1007-1149 1098-1209 1234-1486 1016-1190 1106-1279 1199-1273 1082-1165 993-1466 990-1259 1088-1139 1111-1142 915-1227 989-1158 908-1563

Clinopyroxene P(kb)

T(~

17.3 13.3-18.4 9.1-19.2

1011 192-943 897-1054

26.2

1204

5.0-38.3

844-1207

Powell (1978) 10 kb 20 kb 30 kb 10 kb 20 kb 30 kb

1039~ I052~ 1064~ 1029~ 1045~ 1060~

*Depth within mantle sequence below petrological Moho datum (km) Data taken from Browning (1982)

3.2.2 Dunites and chromitites

Within the dominantly harzburgitic host of the Mantle Sequence are dunitic bodies of varying shapes and sizes that may, or may not, contain chromitites. Descriptions of their field relations given by various workers differ widely, although there is general agreement on the petrography, composition and origin of the dunites (Table 3.6). The authorities whose observations are recorded on Table 3.6 studied different parts of the Semail Nappe so, if they are correct, the shape, size and abundance of the dunite masses, and their distribution and occurrence within the sequence, vary considerably from place to place. It is immediately obvious that, in the Wadi Tayin-Ibra traverse studied by Boudier & Coleman (1981), the dunitic bodies are larger and far more abundant than elsewhere. For, in the northern area the dunite masses do not exceed 600 m in length, yet one body 14 km long has been mapped in the Ibra area. Just as the size and abundance of the dunite masses is variable, so also is their shape. All authors describe the masses as being irregular with anastomosing offshoots. The contacts are identified as sharp or rapidly gradational showing complex interfingering with the enclosing harzburgite; in several cases blocks of harzburgite are enclosed in the dunite. Our own observations agree most clearly with those of Bartholomew (1983) who noted that the contacts are always sharp, irregular and intricate. Most workers agree that contacts indicating intrusion of dunite into harzburgite are present but note that the shape and deformation of the dunite masses are commonly related to the harzbur-

gite foliation. Boudier & Coleman (op. cit.) identify two types of dunite masses; those that are concordant with, and those that are discordant to, the foliation. Both Browning (1982) and Brown (1982) note that some dunites interrupt the harzburgite foliation, whereas Bartholomew (1983) maintains that the foliation cuts through all the dunite bodies but is only obvious where the deformation is strongest. He identifies the shape of the dunites as being directly related to the intensity of the deformation they have suffered (Fig. 3.10). Intrusive relations have been so widely observed that there can be little doubt that the dunites were part of, or resulted from, magmatic bodies that invaded the harzburgite. The relations of these intrusive events to deformation are far less obvious. The simplest case of post-deformational dunite intrusion, as identified by Boudier & Coleman (1981), does not seem to occur in the northern area. Here, although observations are in general agreement, interpretations vary. Brown (1982), noting that the igneous features within the dunite, such as the cumulus textures of chromite layers, were undisturbed, took this as evidence that at least some of the dunites were post-deformational. Bartholomew (1983) confirms these observations but maintains that the centres of the ]east deformed dunite bodies were protected from deformation because the stress was absorbed by the easily recrystallized olivines in the marginal zones. Bartholomew (op. cit.) identified the deformation (foliation) in the harzburgite as stronger in some areas than others. The dunites were intruded at various stages of the deformation history so that those intruded earliest would usually be the most deformed. This model of

The Semail Ophiolite

(a)

53

(c) :'~-'-'." 5"-".-.".'.'.'

.'.'.

Interdigitated contact

[ Harzburgite

Anastomozing

sheets/dykes/veins of dunite

Dunite

(b)

Foliation trace Fig. 3.10. Progressive deformation of a dunite body (taken from Bartholomew (1983)).

Table

3.6. Summary of information on dunite bodies in the Mantle Sequence

Authority-area studied

Abundance, size and shape

Occurrence within sequence

Relation to harzburgite

Rock types and textures

1. Boudier & Coleman (1981) Wadi Tayin Ibra

Abundant Sharp, interforming 50% of fingering outcrop at base of sequence. Largest 14 x 3 km. Shape varies with deformation, markedly flattened at base of sequence

Most abundant near top and base of sequence

Two types identified concordant with and discordant to the foliation. Discordant type cuts foliation

2. Brown (1980, 1982) Northern mountains, mainly Wadi Ragmi and Farfar areas

5-15% of sequence, 1-600 m in length. Irregular, lenticular, anastomosing

Mainly near top of sequence in uppermost 500 m but some large bodies near base

3. Browning < 5 % of sequence, (1982) Wadi Bani 1-10 m. Irregular, Kharus, Rustaq anastomosing block

4. Bartholomew Very variable in (1983) Northern abundance part of (c. 10%) 1-600 m mountains in size. Shape controlled by deformation irregular when undeformed, elongated in foliation when deformed

Contacts

Mineralogy

Origin

Adcumulate dunites with podiform chromite

FOgl t)2

Precipitates from primitive picritic tholeiite magma with some bodies partial melt residue

Some interrupt foliation. Rarely enclose blocks of harzburgite

Complete range from dunite to chromitite. Olivinechrome spinel adcumulates with rare cpx and hbl oikocrysts

Fo~5 , ~ . 3 chrome spine cpx hbl p[ag Fe-Ni sulphides

Magmatic precipitates within rising basaltic magmas

Interfingering, Mainly within rapidly lowermost 2 km gradational but of sequence well-defined

When undeformed seem to cut harzburgite structure. Elongated in foliation when deformed

Adcumulate dunites

Fo,~2

Magmatic

Interdigitating, Very variable. sharp Five areas studied showed wide variation in size, shape and abundance

Invariably have foliation parallel to that of enclosing harzburgite

Interfingering, sharp

Chapter 3

54

magmatic emplacement into a varying and variable stress field, we believe, accounts for all the field observations. Although dunites with occasional and isolated chrome spinel grains forming only 1-2% of the mode are the most common rock type in these bodies, all petrographic varieties from dunite through chromitiferous dunite and olivine chromitite to chromitite occur. In all cases the chromitite is enclosed in a dunitic host although this may only be a few centimetres thick and intensely sheared. Usually, it is the largest dunite bodies that contain the largest chromitite masses and, in general, these lie in the uppermost 1.5 km of the Mantle Sequence just below the Petrological Moho (Brown 1980, 1982). In a detailed study of these deposits Brown (1982) identified structures and textures suggesting that both the olivines and chrome spinels were precipitated from a basic, probably picritic, melt. In particular, the chromitite commonly occurred in layers that showed igneous sedimentation features such as cross-bedding, slumping and differentiation compaction (Plate 3.8). The identification of these textures as dominantly adcumulate stands up to subsequent investigations although Christiansen & Roberts

(in press) reinterpret the "harrisitic" textures of Brown (1982) as due to the formation of olivine neoblasts, The compositional range from dunite to chromitite is due to a concomitant change in the modal abundances of olivine from 99 to 10% and chrome spinel from 1 to 90% (Brown 1982). Brown (op. cit.) also records the occasional presence (<3% modal) of colourless Mg-Cr rich diopsides (Ena~-s0.5 Fs2_5 Wo4a..~-ag) and pargasitic amphiboles (Mg'~9.2-93.7) which occur as interstitial grains poikilitically enclosing chrome spinels. Brown (op. cit.) also notes that in one chromitite specimen there is c. 20% of plagioclase (An76._79) which, although he does not comment on its origin, probably represents basaltic melt trapped between the chromite grains. The dunite olivines fall in the range FO91.5_96.3 with NiO contents of 0.35-0.84%. The olivines lowest in MgO and NiO occur in the chromite-poor dunites, and these oxides become progressively more abundant through to chromitite suggesting some sub-solidus exchange between the two minerals. Compared to those in the harzburgites, the dunite olivines are higher in Fo but lower in NiO. Similarly, the chrome spinels

(a)

(b)

f

-

41~,

Plate 3.8. (a) Cross-laminations of fine chromitite and dunite layers between a chromitite layer above and a dunite layer below9 (b) Possible load structures, protrusions of dunite extend up into an

overlying chromite layer. From Brown (1982).

:'~~.,~, *

~t

9

"

,

k"

II

t

.

dun

er

~

": ~ ~ ~i"~ " ,84

The Semail Ophiolite are higher in Mg/(Mg+Fe > ) (0.38-0.86) and marginally higher in Cr/(Cr+AI) (0.44-0.83) than the harzburgite spinels (Fig. 3.11). This relation is emphasized when chrome spinels from spatially adjacent harzburgite and dunite are compared (Brown 1980, 1983). With increasing depth below the Petrological Moho along the Ragmi-Rayy traverse, Brown (op. cit.) noted that the dunite chrome spinels became progressively higher in Cr/(Cr+AI) and lower in TiO2. 3.2.2.1 Chrom#e mineralization Chromitite occurs as pods, lenses and layers in dunite bodies within the Mantle Sequence and as thin layers or lenses within the basal ultramafic cumulates of the Layered Series (Peters & Kramers 1974). Brown (1982) showed that the biggest chromite deposits are found in the mantle dunite bodies immediately beneath the Petrological Moho, where individual pods reach up to 15 m in thickness and are 40-50 m long. He initially estimated that the largest deposit in the northern mountains (Maharah 1 in Wadi Ragmi) contained c. 600,000 tonnes of

Plate 3.9. Composite gabbroic dyke in harzburgite with clinopyroxene-rich rim and plagioclase-rich centre. Dyke - 4 cm in width. From Bartholomew (1983).

55

chromite ore at an average grade of 42% Cr203. More recently, accurate mapping reduced the estimate to <150,000 tonnes (S. Roberts pers. comm. 1985). Brown (op. cit.) notes that, although chrome grades (Cr203 content and Cr/Fe ratios in chrome spinels) generally increase with depth within the Mantle Sequence, there is a markedly concomitant decrease in ore tonnages at deeper levels. The chrome spinels in chromitites contain rare inclusions of pentlandite and low concentrations of platinum-group elements (average values Pd 8, Pt 14, Rh 6, Ir 48 and Ru 135 ppb) (Page et al. 1982).

3.2.3 Mantle dykes There are numerous but volumetrically small mafic and ultramafic dykes, veins and small pods that cut the Mantle Sequence. They usually occur as steeply-inclined, 1 cm to 1.5 m wide, bodies with sharp but unchilled margins (Plate 3.9). All cut through the dunite bodies and many transect the tectonite foliation of the host peridotites. No dykes were seen to intersect the basal thrust in the northern mountains but

56

Chapter 3 A

0.8

0 0

i

0.7-

0.6-

0.5-

+ 0.4

m

/

/ I

O

/

/ /

/ /

l

/ /

/

/

/

// /

B

/

t

/

0.2-

/

//

I

0.3t-

/

I

/

/

!

/

/

/

I

1

I

\ \

//

i

0.1

I

I

1

I

I

I

0.9

0.8

0.7

0.6

0.5

0.4

M g / (Mg + Fe 2+)

Fig. 3.11. Chrome spinels from dunites (o) and chromitites (shaded field). Tie-lines connect mineral pairs from coexisting dunites and harzburgites. Field of harzburgite spinels outlined by dashed line.

Pallister (1981) describes basaltic dykes in the Wadi Tayin area that cut the metamorphic sole. Some dyke types are apparently confined to the Mantle Sequence but others extend upward into the Layered Series although no single dyke has yet been traced continuously up into the Sheeted Dyke Complex. Attempts have been made to classify the Mantle Sequence dykes on the basis of their compositions, trends and relations to the deformation of the harzburgites. Brown (1982) subdivided them into ultramafic and mafic varieties and noted that the former are 3-30 cm wide, coarse-grained ortho- and clinopyroxenites, websterites and olivine websterites that are only present in the lower part of the sequence. The mafic dykes are 5-50 cm wide olivine-orthopyroxene, clinopyroxene and hornblende gabbros, varying in texture from microgabbroic to pegmatitic, that mostly occur near the top of the sequence. He identified four mafic dyke generations with varying mineral assemblages: (i) ol > cpx > plag, (ii) plag > cpx > ol, (iii) plag > cpx > opx and (iv) plag > cpx and noted that, in the case of dykes formed by multiple intrusion, the last melt to be emplaced is usually the most feldspathic. Bartholomew (1983) identified one dyke, devoid of plagioclase in its lower section, that became progressively richer in that mineral upwards thereby changing from ultramafic to mafic in composition. Browning (1982) was able to subdivide the dyke types in the Rustaq block still further. He agreed with Brown's (op. cit.) general subdivision into upper mafic and lower ultramafic dykes but further divided the latter into clinopyroxenite and orthopyroxenite types and also identified a suite of "late" gabbronorite dykes that occur throughout the sequence cutting all other rock types and extending up into the layered rocks above.

9 Mafic dykes 9 Ultramafic dykes [] Gabbronorite dykes (Wadi Bani Kharus only) Fig. 3.12. Mantle dyke trends: A, Wadi Ragmi; B, Farfar; C, Wadi Bani Kharus. Symbols are poles to the dykes and line shows mean strike of sheeted dykes in each area.

The trends of the Mantle Sequence dykes (Fig. 3.12) have been measured but the coherence of these data is diminished as Brown (1982) and Browning (1982) relate them to composition whilst Bartholomew (1983) correlates with deformation. Brown (op. cit.) notes that the trends of the dykes are scattered but that the mean coincides with the general strike of the overlying sheeted dyke complex in both the Ragmi (120 ~ and Farfar (020 ~ areas. Browning (op. cit.) reports that for the Rustaq block the majority of the dykes trend between 320360 ~ with steep dips whilst the gabbronorite dykes strike at 040 ~ and dip NW. Bartholomew (op. cir.), from studies in several areas, identifies three main dyke orientations, in order of decreasing age: (i) 340 ~ with steep dips, (ii) 040 ~ with steep

The Semail Ophiolite

Plate 3.10. Pyroxenite dyke in harzburgite showing ptygmatic folding with extreme thickening of fold noses and attenuation of fold limbs. From Brown (1982).

Plate 3.11. Shearing of plagioclase-rich gabbroic dyke. Bar indicates $2 foliation in the harzburgite, coin is 2.5 cm in diameter. From Bartholomew (1983).

Plate 3.12. Photomicrograph of

orthopyroxenite, showing sutured margins between grains and associated fine marginal recrystailization. Length of field of view 8 ram. From Browning (1982).

57

Chapter 3

58

dips and (iii) 120~ with shallow dips. He records that types (ii) and (iii) cut the layered gabbros. He also states that the majority of the dykes are only affected by late-stage brittle deformation but where deformation is most intense at the base of the sequence the dykes are most deformed so that they lie parallel to the foliation, are boudinaged and in places tightly, almost isoclinally, folded. The deformed dykes show two deformation styles. The more competent pyroxene-rich dykes have been deformed by flexural shear processes into tight folds or boudins (Plate 3.10); pyroxene-poor dykes, whose competence is closer to that of the enclosing harzburgite, have been rotated by shearing towards the foliation plane (Plate 3.11). The pyroxenite dykes are coarse grained (1-5 cm) with large interlocking pyroxenes showing sutured grain boundaries and associated fine marginal recrystallizaton (Plate 3.12). Most are composed wholly of pyroxenes but some contain small rounded grains of olivine and occasional small black chrome spinels. There is a complete gradation from orthopyroxenites (>95% modal opx) through websterites to clinopyroxenites. Porphyroclastic texture is developed in dykes which have suffered deformation and folding and the pyroxenes show curved cleavages, kink bands, undulose extinction and kink-controlled exsolution blebs. However, strained porphyoclasts of olivine and pyroxene in dykes composed largely of strain-free grains may be xenocrysts derived from the dyke walls (Brown 1982). Orthopyroxene compositions show a range En,s7.,-s3.5 Fsl~.4_ms Wos.s--l.l and clinopyroxenes have a mean composition of En47.1 Fs4.s Woas.1. Brown (1982) found one clinopyroxenite dyke that contains about 15% modal brown hornblende (Mg' 86.o).

(a)

(b)

Mg' (cpx)

9O !

85 I

The gabbro dykes are likewise mostly coarse grained and exhibit a wide range of petrographic types. They sometimes show a remarkable banding with phase ratio layering parallel to the walls. Poikilitic textures, identical to those of cumulates, are characteristic such as: (i) rounded and embayed olivine and orthopyroxene grains in clinopyroxene oikocrysts, (ii) olivine and clinopyroxene enclosed in plagioclase with reaction rims of hornblende, (iii) plagioclase enclosed in clinopyroxene or hornblende and, rarely, (iv) hornblende enclosed in plagioclase. Chrome spinels (mean composition Mg/(Mg+Fe) = 0.42, Cr/(Cr§ = 0.57) and Ni-Fe sulphides are present as accessory phases. The olivines in the gabbro dykes range Fo~5 9,, clinopyroxenes En52.2 4~9 Fs5.4-4.3 Wo46.x-42.2 and the plagioclases Ans~s7. These compositions are typical for cumulate olivine gabbros from the base of the overlying Layered Series (Fig. 3.13) and support the idea that the two are cogenetic and that the gabbro dykes represent fractionated liquids that solidified within the Mantle Sequence before entering the main magma chambers. The gabbronorite dykes show a similar variety of phase assemblages and textural types to the gabbro dykes (Browning 1982). Some are fine grained (<1.5 mm) throughout, some have coarse grained margins and show a wide range of grain sizes (0.5-7 ram). Olivine is absent from these rocks and they consist of orthopyroxene (En76(~715 Fs26.7 21.s WO1.6-1.8), clinopyroxene (En44.2_43 ~ FSl,. l s9 Wo4~.1~47..) and plagioclase (An~7 ~_,~ 4). The unusual association of highly calcic plagioclase with relatively fractionated (Mg'(opx) 73-78, Mg'(cpx) 81-83.5) pyroxenes is a distinctive feature of these rocks and distinguishes them clearly from the gabbro dykes and the cumulate gabbros, and websterites of the Layered Series

8O I

Mg' (opx) 75 80 I I

70 I

75 I

X•

95

95

85 I

•

90

/

\ I

\\

An

9 \

B

-

90

\\

\

85

,

\

An \

\

\

\

J

k

\

N

X \ X

80

o

85

N

\

\ N

\

\

N N \

\

\

\

N

k

CY

80

\

,y

\

X \

75 -

\

75 •

70

X Fig. 3.13. Pyroxene-plagioclase plots for mantle dykes. (a) Mg' (cpx) vs An. Open circles - - gabbronorite dykes, closed circles -- gabbro dykes. A -- field of Wadi Ragmi orthopyroxene-bearing cumulates. B - - field of other cumulates. (b) Mg' (opx) vs An. Dots - - gabbronorite dykes, crosses -- Wadi Ragmi cumulates.

The Semail Ophiolite

59

Table 3.7. Composition of selected mantle dykes

Orthopyroxenites

Websterite

Gabbronorite Gabbro pegmatites

0M1471 0M2334 0M2330

0M2345

0M2325 0M2343 0M3698

SiO2 TiO2 Ai20~ FeO* MnO MgO CaO Na,O K,O P205 LOI

55.43 nd 1.62 6.95 0.16 32.50 2.23 nd nd nd 0.89

54.49 nd 1.81 6.88 0.16 32.67 2.20 nd nd nd 1.42

53.18 nd 1.48 7.66 0.18 32.42 2.08 nd nd nd 2.07

52.7(I 0.10 3.79 7.43 0.17 20.32 14.70 nd nd nd (1.56

44.02 0.05 19.77 6.00 0.11 10.66 13.01 0.95 0.06 nd 5.33

48.82 0.05 18.19 6.64 (I.14 12.05 13.41 nd nd nd 0.57

47.06 0.2(I 19.06 7.28 0.15 9.58 13.73 1.36 nd nd 0.52

Total

99.78

99.63

99.07

99.77

99.96

99.87

98.94

0.24 3.54 47.38 14.91 29.32 2.55 1.26 0.09 0.07

0.24 11.59 46.11 17.95 8.86 12.66 1.40 0.38 0.10

44 318 185 207 34 146 47 nd

21 125 48 48 131 211 39 nd

-

-

CIPW norm,

or ab an di hy ol mt il ap

4.44 5.26 85.19 3.10 1.33

4.99 4.76 82.11 6.14 1.32

4.13 5.07 79.22 9.31 1.49

10.33 49.40 36.13 1.67 1.42 0.19

0.37 8.43 51.95 12.66 7.36 17.31 1.20 0.10

Co Cr Cu Ni Sr V Zn Zr

68 4464 8 812 nd 124 45 nd

21 4502 9 774 nd 109 40 nd

74 4153 6 798 nd 145 47 nd

412 846 559 412 17 254 46 nd

13 103 12 185 67 220 42 nd

*total Fe as FeO OM1471 OM2334 OM2330 OM2345 OM2325 OM2343 OM3698

n d - not detected

-FeO/Fe203 = 6.7

Orthopyroxenite dyke, Wadi Bani Kharus. Orthopyroxenite vein in harzburgite, Wadi Bani Kharus. Orthopyroxenite vein in dunite, Wadi Bani Kharus. Websterite dyke in harzburgite, AI Maydah, west Rustaq block. Gabbronorite dyke, Wadi Bani Kharus. Gabbro pegmatite dyke (fine grained facies) cutting harzburgite, AI Maydah. Gabbro pegmatite dyke, Wadi Ragmi.

(Fig. 3.13). It supports the contention that the gabbronorite dykes represent a later magmatic event perhaps related to some of the late intrusive complexes and the upper lava units (Sections 3.6 & 3.8). The crystallization of highly calcic plagioclase from the mafic melt is usually taken to indicate crystallization having taken place under conditions of high PH2o (Johannes 1978) but it is not clear why this should be so in the case of gabbronorites which contain no primary amphiboles and it may simply be a function of the unusual composition of these magmas with perhaps low contents of Na20 and correspondingly high Ca/Na ratios. Whole rock compositions of mantle dykes are presented in Table 3.7.

3.2.4 Structure The structures impressed upon the Mantle Sequence whilst it was an in situ part of the oceanic lithosphere have been partly

overprinted by those produced during detachment of the ophiolite from its oceanic setting before emplacement onto the Arabian continental margin. Only the constructive margin structures produced by continuous and progressive deformation at high temperatures (1000-1200~ and relatively low to moderate pressures are described and discussed in this section. The field and petrographic evidence discussed previously indicates that the structures and textures of the Mantle Sequence rocks were largely formed by plastic deformation at high temperatures. The plastic deformation of rocks is best considered as a flow process in which the transport of matter is driven by stress. At constant volume, the flow of plastic solids is analogous to that of viscous fluids and is essentially a shear process with flow depending on two conflicting processes, work hardening and recovery. In work hardening, the more deformed a material becomes, the more difficult it is to deform. Recovery allows deformation to proceed usually by crystal slip processes. However, in material undergoing

6o

Chapter 3

steady-state plastic deformation at the high temperatures and strain rates envisaged for the lower part of the oceanic crust and subjacent upper mantle, the strain energy is sufficiently high to allow recrystallization instead of recovery. Newlycrystallized grains are rapidly strain-hardened and the process repeats itself with periods of strain-hardening alternating with periods of recrystallization. Nicolas & Poirier (1976) compare this process to metallurgical "hot working". At higher stresses the viscosity and shear stress are interdependent with viscosity decreasing as shear stress increases. Durham et al. (1977) described this non-Newtonian steady state flow as power law creep. Bartholomew (1983) states that dislocation is the most common mechanism for high temperature creep. In this mechanism, the main process causing strain during power law flow is dislocation slip, the dislocation being between adjacent, parallel planes. The model invoking creep by dislocation slip was first proposed by Weertman (1968) and discussed in detail by Nicolas & Poirier (1976). An alternative mechanism whereby recrystallization (piezo crystallization) is stress-induced was proposed by Av6 Lallement & Carter (1970), but Nicolas & Poirier (op. cit.) argue that dislocation creep and associated annealing recrystallization best fit the available petrographic evidence from ophiolite mantle sequence textures. If it is assumed that the dominant mechanism for steady state flow in the Mantle Sequence rocks of the Semail ophiolite is power law creep, then the main process of flow will be dislocation creep caused by slip along a specific slip dislocation plane in a crystal in a specific slip direction. The amount and direction of slip is usually expressed as the Burgers vector. Within a crystal the energy of dislocation is proportional to the square of the Burgers vector's length. As the Burgers vector is directly proportional to the unit cell dimensions, slip most readily occurs in the direction of the smallest dimension of the unit cell. Thus, it is possible to predict operative slip planes and directions for the olivines and orthopyroxenes in the Mantle Sequence tectonites. For forsteritic olivine the expected slip directions would be either [100] or [001] and most likely slip plane would be (010) as this is the only plane on which Si-O bonds do not have to be broken (Poirier 1975). From experimental deformation of individual olivine crystals, albeit at much faster strain rates than those calculated theoretically for the asthenosphere (10 -~ in contrast with 10-~3 to 10-~5 sec-l), Carter & Av6 Lallement (1970) identified three major olivine slip systems that are largely dependent on temperature. From this it was noted that slip in olivine at upper mantle astheno spheric temperatures and pressures should be in the (010) [100] system (i.e. slip in the [100] direction on (010) planes). Experimental studies by A. Nicolas and his co-workers (e.g. Nicolas et al. 1973; Poirier & Nicolas 1975; Nicolas & Poirier 1976) show that the development of preferred orientation textures in olivine-rich rocks is a two-stage process. Below 30% strain the initially anisometric grains are rotated mainly by grain boundary sliding. Thus, the grains rotate bodily with the [010] crystallographic axis rotating towards the direction of applied stress (a~). At greater than 30% strain, intercrystalline (010) [100] slip is dominant and dislocation creep causes the [100] axis to become orientated parallel to the shear direction, and [010] perpendicular to the shear plane. So, the olivine [100] crystallographic axis becomes parallel to the minimum principal compressional strain axis. Most workers identify that the Mantle Sequence lineations are parallel to the olivine [100] axis. For enstatite, another orthorhombic silicate, the shortest

Burgers vector is parallel to the [001] crystallographic axis in the (100) plane and therefore the (100) [001] slip system should be most active during dislocation creep. Nicolas & Poirier (1976) have shown that at high temperatures orthopyroxene has a much higher resistance to creep than olivine. So, only at higher stresses will orthopyroxene deform by dislocation creep, and only then will its [001] crystallographic axis become parallel or sub-parallel to the olivine [100] axis. As orthopyroxene recrystallization textures are uncommon in mantle sequence rocks it is evident that the high creep resistance of orthopyroxene inhibits polygonization and recrystallization. The shortest Burgers vector for clinopyroxene is in the (100) [001] system and, like enstatite, it also has a high creep resistance. Bartholomew (1983) used a 4-axis universal stage to determine the active slip systems in olivine and orthopyroxene by identifying the crystallographic orientation of mineral grains. Then, by determining the angular relations between crystallographic axes and crystal shape (foliation and lineation) orientations, he was able to identify whether the specimens had undergone pure or simple shear deformation. After reviewing experimental and descriptive evidence from natural peridotites, he concluded that, at high temperature, deformation occurred by simple shear causing the crystal lattice to rotate towards the shear direction, as proposed by Nicolas & Poirier (1976). In simple shear the amount of shearing involved can be identified by measuring the angle (a) between the slip and mineral elongation (lineation) directions. This relation is illustrated for progressive dextral shear on Fig. 3.14. After obtaining crystallographic orientation data from the Mantle Sequence rocks, Bartholomew (1983) noted that "for most areas specimens have been deformed in a plastic shear flow regime in which the shear (slip) direction has changed". He proposed that the intensity of the last active slip direction (D2) determines the variation in the orientation of the olivine slip [100] direction. In areas of intense D2 deformation, foliation planes have a constant trend and no relicts of the earlier (D1) shearing have been preserved. However, in areas of weak D2 shearing, the D2 foliation is weak and relict D~ foliations are preserved. Bartholomew (op. cit.) studied the mantle structures along a series of wadi sections spaced at intervals of 2030 km along strike. He noted that the dominant shear sense and intensity of shearing varied along individual traverses and that no structural correlations could be made between traverses (see also Smewing et al. 1984). The D2 deformational features for each area are different and the dominant shear sense, variation in intensity of shearing with depth below the Moho and the direction of shearing are all unique for each traverse. Bartholomew (op. cit.) suggests that this implies that both the D~ and D2 events in different parts of the Semail ophiolite took place at different times. This is of critical importance when considering the relation of these structural features to the oceanic spreading environment. For the detailed studies indicating that the penetrative deformation of the Semail Mantle Sequence was an extremely complex process, the reader is referred to Bartholomew (1983). Here, these data are examined to see if they will provide insights into the style and nature of asthenopheric flow beneath and/or adjacent to the spreading axis. First, it is relevant to note that Bartholomew's investigation led him to agree with Nicolas & Violette (1982) who assumed that ophiolitic mantle structures were asthenospheric flow structures and that once the mantle cooled below 1000~ it was no longer ductile. So, if the limit to ductile deformation is taken as the

The Semail Ophiolite asthenosphere-lithosphere boundary, the structures preserved in the Mantle Sequence will be those imposed during the nearhorizontal flow away from the spreading axis, the structures produced during the vertical rise of the mantle material having been obliterated. This structural data can be used to decide whether the asthenosperic flow makes a "right-angle turn" of the type envisaged by Langseth et al. (1966) in which the uprising mantle passes through a 90~ turn to flow horizontally away from, and in a direction normal to, the ridge axis, or whether the mantle rises diapirically to then diverge radially from centres during the horizontal flow stage as has been proposed by Juteau et al. (1977) and Nicolas & Violette (1982). To explain the processes involved in producing the two

6I

shearing events, D~ and D 2, Bartholomew (1983) proposed the following model. He envisaged that, as the asthenosphere flows more or less horizontally away in all directions from a number of diapiric centres on the same ridge, the divergent flow systems would meet and interact with one another in the regions between the centres (Fig. 3.15). If the diapiric centres were active at different times, then the younger ones would produce D2 flow planes that would fold and shear those of an earlier centre. The intensity of the D2 shearing would depend on how close that part of the oceanic lithosphere, subsequently preserved as an ophiolite, was to the diapiric centre when the mantle material became 'fossilized' on passing through the asthenosphere-lithosphere boundary, and the spreading rate. If the ophiolite were originally close to a diapiric centre, then

0

r

Shear direction Less resistant phase eg Olivine ~More

resistant phase eg Orthopyroxene

Fig. 3.14. Progressive shear in harzburgite (taken from Bartholomew 1983).

Fossilized

flow

asthenospheric

SPREADINGDIRECTION Active magma chamber Extinct magma chamber

Crust I t Mantle j ~

sequence1t

Asthenosphere Lithosphere I boundary

a., /

J <

diapir

z..._._

Earlier diapir

11

-- ~

~

~--L_ V

Fig. 3.15. Envisaged diapiric flow system in the sub-spreading axis asthenosphere showing how flow lines of an earlier diapir (thin arrows) could have been shear folded by those of a later diapir (bold arrows). After Bartholomew (1983).

62

Chapter 3

the shearing away from that centre would be intense compared to that in the area distant from it. At a slow-spreading ridge the process of magma generation and movement of asthenospheric shearing would be slower and weaker but the orientation of the asthenosphere-lithosphere boundary would be steep and the diverging flow planes from the diapiric centres would be 'frozen in' much nearer the centres. Although there is no theoretical reason why the divergent flow from the diapiric centres should be horizontal, they commonly are in laboratory experiments (Elder 1977) and the mean orientation of the shear planes in the Semail Mantle Sequence seem to imply that they were. However, these planes show no consistencies in their shear sense. This can be explained by invoking surges of asthenospheric mantle moving at different rates and depths away from a diapiric centre and so producing variations in both shear sense and intensity of deformation with depth and along strike as found by Bartholomew (1983). Another factor that could complicate the flow system would be the presence of a barrier that could divert the asthenospheric flow. The most likely barrier in a constructive plate margin setting would be a transform fault plane. Here asthenospheric mantle encounters lower temperature lithospheric mantle and the asthenospheric flow would be directed away from the spreading axis along the transform zone. Although the diapiric divergent flow model explains many of the structural features of the Semail Mantle Sequence, as yet we do not have sufficient data to place the various areas studied in detail precisely within this model framework. 3.2.5 Alteration

The compositional and textural complexity of the Mantle Sequence has been increased by the variable alteration of the various rock types which, in ultramafic rocks, comes under the broad generic heading of "serpentinization". Although no special study has been made on the alteration of the Semail Mantle Sequence, data on the extent of the alteration and the processes involved were needed for economic (Brown 1982), geophysical (Shelton 1984) and hydrogeological (Stanger 1986) investigations. The account presented here is taken from these studies although it is emphasized that the data

Plate 3.13. Serpentinized harzburgite showing development of "chicken-mesh" replacement. Length of field of view - 4 mm [XN].

could be atypical for the Mantle Sequence as a whole because the specimens were collected for specific studies, e.g. many are unusually chromitiferous or show other compositional or structural abnormalities, and collections were made almost entirely from water-worn outcrops in wadis whose courses may be fault-controlled which may, in turn, affect the degree of alteration. 3.2.5.1 Alteration products

The six primary minerals of the Mantle Sequence; olivine, orthopyroxene, clinopyroxene, chrome spinel, plagioclase and amphibole (in decreasing order of abundance) have, to varying degrees, been altered to secondary assemblages of lizardite, chrysotile, iron oxides, secondary amphiboles, chlorite and carbonates. Only the chrome spinel, although commonly displaying marginal alteration to ferritchromite, is inert during serpentinization. Olivine, and to a lesser extent orthopyroxene, is first altered along grain boundries and fractures. With extensive alteration, a "chicken-wire" or mesh texture (Wicks & Whittaker 1977) is developed in which a network of serpentine encloses a tesselated array of polygonal cells (Plate 3.13). Each cell has a rim of fibrous serpentine and a core of either primary olivine or orthopyroxene or serpentine. Usually, there is only a rim to each cell composed of the 'qizardite" polymorph (Wicks & Whittaker op. cit.). In those few examples with serpentine centres, the serpentine is near isotropic "serpophite", identified by Wicks & Whittaker as microcrystalline lizardite. Veins of chrysotile varying in width from 1 mm to several 10's cm crosscut the mesh textures and, in places, replace the lizardite of the mesh rims. The Mantle Sequence chrome diopsides may be altered to tremolite which occurs as fibrous aggregates pseudomorphing the grains and as veins. It has higher Mg' than the parent clinopyroxene, relatively high Cr203 content (<1%), low A120~ (<0.9%) and negligible Na and K. Chlorite, which is found in the matrix of some chromitites, is bright green in hand specimen but colourless to pale green in thin section. It is very low in iron and high in chromium (<3.8% Cr203) and is either sheridanite or clinochlore in Hey's (1954) chemical classification. All of the alteration products have higher Mg/(Mg+Fe) ratios than the minerals from which they were derived. The

63

The Semail Ophiolite alteration processes must therefore involve the release of iron from silicate minerals which results in the formation of magnetite and/or haematite depending on the prevailing redox conditions (e.g. Neal & Stanger 1984). Irregular stringers of these iron oxides are common in serpentinite mesh and vein structures. The carbonates, magnesite, dolomite and calcite, occur in veins and shear zones throughout the Mantle Sequence. Magnesite is by far the most common (c. 80%) carbonate in the peridotites, the remainder being mostly dolomite, whereas calcite alone is the cementing matrix in recent wadi gravels (Stanger 1986). Typically magnesite occurs as pure white, very fine grained, massive to concretionary textured, millimetre to metre thick veins many of which show slickensiding. Instances of present day magnesite precipitation are known, but only from freshly exposed basal Mantle Sequence areas where they are closely associated with metastable huntite, a magnesite precursor (Stanger 1986). Silicified serpentinite forms hard, resistant outcrops with sparkling lustrous fracture surfaces that occur sporadically in the highly altered rocks along the basal thrust of the Semail Nappe and along a few fault zones through the Mantle Sequence. There is no doubt that the protolith of the silica-rich rock was intensely brecciated serpentinite, for occasional chromite grains occur with a matrix of red-brown ferruginous cherty silica within a micro-vein network of goethite (Stanger 1985). Various terms (e.g. birbirite and amqat) have been coined for this rock type and Glennie et al. (1974) proposed that it was produced by "selective leaching of magnesium under tropical weathering conditions". Alternatively Stanger (1985) concludes that "silicification was a low temperature chemical replacement feature requiring special conditions of formation and not a weathering phenomenon". He further suggests that total serpentinization of the protolith and reactivation along fracture zones, permitting large scale circulation of slightly acid, CO2-rich groundwaters under reducing conditions at temperatures of 40-50~ were necessary pre-requisites for silicification. The silicification was contemporaneous with serpentinite dissolution and formed around residual iron oxides and hydroxides. During the dissolution process nickel is released from the serpentinite and arsenic from the residual iron minerals, producing late-stage Ni-As and Ni silicate mineralization. 3.2.5.2 Physical and chemical effects of serpentinization The majority of the Mantle Sequence rocks are between 50% and 80% serpentinized. As there should be a direct correlation between the density of the ultramafic rocks and their degree of alteration, Shelton (1984), in his evaluation of the gravity field produced by the Semail Nappe, enquired into the spatial variations in density and serpentinization. Fig. 3.16 plots percentage alteration (as deduced from the loss of ignition at 1000~ against specimen density. The regression line was constructed to pass through the unaltered harzburgite (OM1499), which has a density of 3.329 g cm -3, and has a 96% fit at a fully hydrated serpentinite density of 2.68 g cm -3. OM1499 comes from the basal section of the harzburgite of Wadi Jizi, it preserves the strongest deformation fabric of any sample collected and it appears that the cataclastic deformation made the rock impermeable to serpentinizing fluids. The densities of highly altered rocks range from 2.55 to 2.76 g c m - 3 reflecting the complexity and variability of the alteration processes. Based on an average p-wave velocity through the Mantle

100 -

80

c

60

< ~

4O

O

20

0

1 2.6

I 2.8

I 3.0 Density g/cc

I OM1499 I 3.2 3.4

Fig. 3.16. Alteration vs density plot for harzburgites (after Shelton 1984). Sequence of 7.0 km s e c - I (Glennie et al. 1974), the derived density suggests that the average alteration of the rocks is 42%. The Basal Serpentinite unit of Searle (1980) at the base of the Semail Nappe is intensely sheared and generally 100% serpentinized. It appears darker at outcrop and on aerial photographs than the Mantle Sequence proper and is readily identifiable on Landsat MSS (Multispectral Scanner) images. Rothery (1982, 1984), using the spectral response of this unit, was able to identify other areas of intense serpentinization. Locally high degrees of serpentinization occur along imbricate or high-angle fracture zones or lineaments that cut through the Mantle Sequence and throughout most of the leading edge of the Semail Nappe. However, variations in degree of serpentinization in areas smaller than about 200 m 2 cannot be identified using such imagery. Although there is a crude correlation between serpentinization and magnetic properties (the more intensely serpentinized rocks having higher magnetic intensity and susceptibility), this relationship is .insufficient to allow the use of aeromagnetic measurements as a guide to degree of serpentinization (Shelton 1984) (Fig. 3.17). Neal and Stanger (1984) and Stanger (1986) suggest that there are two types of serpentinization in the Semail ultramafic rocks. The first, termed "higher temperature alteration serpentinization", is relatively uniform in its effects. The second, which they identified as "low temperature precipitation serpentinization", is clearly related to the movement of present day meteoric waters through the rocks. The higher temperature alteration affects the entire Mantle Sequence and is unrelated in any way to past or present day water tables or to the weathered zone although it is most intense (100%) along major thrust and fracture zones. It is possible that this type of alteration took place either whilst the ophiolite was an in situ part of the oceanic lithosphere or during its detachment. Rothery (1982, 1984) argues that the correlation between serpentinization (as mapped by remote sensing) and emplacement-related structures shows a large part, if not all, of the serpentinization to be due to either syn-

64

Chapter

f

\\ \

3

k\\\\

\\\x\\\

\k

/

\ ~ ~ _ \ L

/

/ ii:!if.:.',.:..~:,~!:c:~:::: :? :.. .:.:: . : .:. .

\'<"-~

/

I\\

"-- - - - - ------ \

\

\

I

) I

//

\

(L)

/

. . . . \

)

i11

_

/

/

/

Degree of Serpentinization LOW (<50%) L----] Moderate High (c 100%) CS Crustal sequence MS Mantle sequence H Magnetic "high" L Magnetic "low" Fig. 3.17. Degree of serpentinization and magnetic anomaly contours for the Rustaq block (after Shelton (1984)). For scale and orientation see enclosure one. or post-emplacement processes. Although commonly pseudomorphing mylonitized and brecciated primary phases in the Banded Unit, the serpentinite minerals are not deformed. It is likely that some serpentinization took place during the later stages of nappe emplacement particularly as the high chloride concentrations in the serpentinite (Stanger 1986) suggest that the all-pervasive serpentinization was produced partly by the passage of sea water. In contrast, the oxygen isotope studies on the Semail Mantle Sequence (Dunlop & Fouillac 1984) indicate that the serpentinization was mainly produced by meteoric water at low temperatures (<50~ as has been found in other ophiolite peridotites (Magaritz & Taylor 1974; Wenner & Taylor 1978). The low temperature precipitation serpentinization of Stanger (1986) occurred at temperatures of c. 30-45~ It has been produced by transient serpentinizing fluids moving along fracture zone pathways through the Mantle Sequence. The fluids originated as meteoric waters and the serpentinization must be related to the post-obduction part of the ophiolite history producing late stage, cross-cutting veins. Changes in the prevailing hydrogeological conditions ensure the irregularity of low temperature serpentinization, which formed at high rock to water ratio, both in intensity and periodicity. Just how much of the serpentinization belongs to the high and low temperature types is, however, difficult to determine.

3.2.6 Summary The following consensus model is invoked to explain the origin of the Mantle Sequence: Relatively "fertile" aluminous lherzolite asthenospheric upper mantle, convectively rising and diverging beneath the constructive margin, would begin to melt at c. 50-70 km depth (O'Hara 1968) to produce a Mg-rich picritic partial melt which would rise adiabatically along with the enclosing high temperature but sub-solidus residual mantle (Section 3.9). The percentage of partial melt would increase (to about 20-30%) so that by the time the system had reached depths of 20-30 km the melt fraction would have coalesced into discrete magma bodies or diapirs. As the magma-harzburgite system continued to rise only forsteritic olivine and chrome spinel would be precipitated from the melt that would remain during its ascent within the expanded olivine stability volume of the hydrous basalt tetrahedron (O'Hara op. cit.). In the three-component system of (i) tectonized harzburgite, (ii) melt and (iii) crystal precipitates, sooner or later the melt will escape into an overlying magma chamber leaving the precipitated olivine and chrome spinel to form the dunite-chromitite masses. As the upper mantle was convecting at temperatures of c. 1000~ (the temperature necessary to extract melts from the peridotite), the system was subjected to continued and progressive deformation. As a result, not only the harzburgites

The Semail Ophiolite but also the dunites show tectonite fabrics that were preserved as the rocks passed through the asthenosphere-lithosphere boundary during more-or-less horizontal flow. In some areas two asthenospheric deformation phases can be recognized and are probably related to a complex interference pattern of diapiric flow beneath the spreading axis. Cross-cutting ultramafic and mafic dykes, some of which are deformed, are related to several stages of later ocean-floor magmatism but pre-date the detachment of the ophiolite from the oceanic lithosphere. The latter event is recorded in the mylonitization and shearing of the basal peridotites in the Banded Unit. Serpentinization of the peridotites was a late-stage, largely static event that occurred either during the last stages of late Cretaceous emplacement of the Semail Nappe or by postemplacement exposure to meteoric ground waters and has continued to the present day.

3.3 The Petrological Moho The contact between the Mantle Sequence and the overlying crustal layers of the ophiolite is a marked structural and petro-

Plate 3.14. General view of the Moho in Wadi Bani Kharus.

Plate 3.15. Gabbro sheets and veins near the top of the Mantle Sequence at Wadi Bani Kharus.

65

logical boundary between tectonized harzburgites below and overlying cumulate peridotites and gabbros that is generally taken to represent the sub-oceanic "Petrological Moho" (Greenbaum 1972; Malpas 1976, 1977; Smewing 1980a). The name was used to distinguish it from the "seismic" or "geophysical" Moho (Cann 1970; Moores & Vine 1971) which occurs at the boundary between ultramafic and mafic rocks and which, where there are thick ultramafic cumulates developed, does not necessarily coincide with the base of the Crustal Sequence (Fig. 3.3). An intermediate "transition" or "critical zone" between these two boundaries, that has been recognized in some ophiolite complexes (Coleman 1977; Malpas 1977) as being composed of ultramafic rocks of mixed residual and cumulate origin, does not occur in the Semail ophiolite; although Pallister & Hopson (1981) refer to the 100-200 m of dunites at the base of the cumulate sequence in the Ibra area as "transitional", they propose a cumulate origin for these rocks. On the other hand, Nicolas & Prinzhofer (1983) recognize both in the Semail and several other ophiolites a zone of residual dunite, that is dunites formed by the melting out of orthopyroxene from a harzburgite and containing layers of wehrlite and gabbro formed by magmatic impregnation, at the

66

Chapter 3

top of the Mantle Sequence. We cannot, in general, agree with their interpretation for these rocks and suggest that, at least for the Semail ophiolite, all the large dunite masses, both within the Mantle Sequence proper and within the Layered Series, are of magmatic intrusive-cumulate origin (see Section 3.4.3). Within the area under investigation the nature of the Petrological Moho boundary shows major structural and petrological changes along strike, often over distances of a few kilometres, The top of the Mantle Sequence is sometimes sharply truncated by the overlying layered rocks whose planar layering can be markedly discordant to the tectonite fabric of the former. The lower cumulates may or may not possess the same tectonite fabric as the underlying harzburgites (Bartholomew 1983; Smewing et al. 1984). There is sometimes a more diffuse type of boundary, either where there are sills or pods of cumulate dunite, wehrlite or gabbro at the top of the Mantle Sequence, as described by Browning (1982) in Wadi Bani Kharus (Plate 3.14), or where brecciation of the harzburgite has occurred and it forms a type of magmatic breccia in a dunite, wehrlite or gabbro host, as described by Brown (1982) from the Wadi Ragmi area. Here, for several tens of metres below the contact, there are 2-3 m thick gabbro sills that lie sub-parallel to the Moho and both they and the host harzburgite are cut by numerous small (<5 cm) gabbro dykes and veins. These increase in abundance upwards until in the top few metres of the contact zone they form a breccia of harzburgite net-veined by gabbro (Plate 3.15). Browning (1982) noted that in the Rustaq block there is a correspondence between the compositions of the minor intrusive bodies, specifically the pods, in the top of the Mantle Sequence and the lower cumulates. Where ultramafic pods and dykes are common then ultramafic cumulates predominate immediately above the contact and, conversely, where gabbroic minor intrusives are abundant, basal gabbroic cumulates are usually present (Fig. 3.18). Where exceptional thicknesses (maximum c. 900 m in the western part of Rustaq block) of ultramafic cumulates occur, they often form lensoid masses that occupy depressions, some apparently fault-bounded, in the Mantle Sequence surface. There also seems to be a correlation between the proportion of ultramafic cumulates present in any section and the total thickness of the layered rocks. In places where the cumulates are thin (<2 km) they consist largely of gabbroic rocks (e.g. Wadi Bani Kharus), but where the sequences are thicker (>2 kin) the amount of ultramafic rocks is greater. The great variability in the thickness of the Layered Series (<0.5 km to >5 km) can be partly attributed to pre-existing structural relief on the top of the Mantle Sequence of perhaps 1-2 km which led to ultramafic cumulates being locally ponded in depressions.

RUSTAQ BLOCK

/ / // ~

~ CS ~. ", "-~ '" -'- ~ ,... f "----\ \ /,.4 t ~ ,-,~. [" -'--~ ~,~." ,..~. L,~ ! ~",

/ .7,~

\

~F

\\ \

\

,'--, ~ ",

E

I

C \

(

MS \ -,,

t ~ ~..

"., "

)L -"- l ~ /

,t

/

/ CS Crustal sequence

/

MS Mantle Sequence

<

\

Edge of ~ ~ _ ~ ~ - ~ ~ .... ophiolite outcrop -

) ~

Cumulate peridotite

....

A

B

~ = . ~

C

Fault

D -

300m I ,

F

E

!

:

- -

,,

,

200 I

Iiilt

100

!ii:

Layered Series Cumulates Gabbros

1 Dykes

Mantle Sequence

I Doterite

[

i Harzburgitetectonites

Troctolites

~ Gabbro

@

Dunite pods

~

Wehrlites

~ Pyroxenite

Q

Gabbro pods

~

Dunites

Fig. 3.18. Sections through the Petrological Moho (top of harzburgite tectonite) in the Rustaq block. From Browning (1982).

3.4 The Layered Series There is abundant evidence from well-preserved ophiolite complexes and the large suite of plutonic rocks that have been recovered from the ocean floor that the lower part of the oceanic crust consists of layered peridotites and gabbros that crystallized in crustal magma chambers. Geophysical evidence for magma chambers beneath present day ridges is, however, only conclusive for fast-spreading ridges such as the East Pacific Rise where seismic refraction and reflection profiling indicate the presence of crustal magma chambers with half-widths

perpendicular to the ridge axis of at least 6 km (Rosendahl et 1976: Herron et al. 1980; Orcutt et al. 1984). All models of ocean crust formation that are based largely on ophiolite studies (Cann 1970, 1974; Greenbaum 1972; Dewey & Kidd 1977; Gass & Smewing 1981) postulate the existence of large magma chambers at the constructive margin. In these models the sub-axial magma chambers are seen as dynamic and longlived systems which are continually replenished by new mantle-derived melts from below and depleted by dyke injection

al.

The Sernail Ophiolite

67

Fig. 3.19. The Layered Series - - outcrop and vertical sections. Columns show the distribution of mafic and ultramafic cumulates. 1. Wadi Fayd, 2. Wadi Ragmi, 3. Wadi Jizi, 4. Wadi al Hilti, 5. Wadi Shafan, 6 & 7 Haylayn area, 8. Rustaq block west, 9. Wadi Bani Kharus, 10. Rustaq block east, 11. Ibra area. (Data taken from Pallister & Hopson 1981, Smewing 1981, Browning 1982, Dahl et al. 1983 and Bartholomew 1983).

68

Chapter 3

and lava eruption at the top. Detailed studies of the layered plutonic sequences in many well-known ophiolite complexes such as Troodos, Cyprus (Greenbaum 1972; Allen 1975), Bay of Islands, Newfoundland (Church & Riccio 1977; Malpas 1978; Elthon et al. 1982), Vourinos, Greece (Moores 1969: Jackson et al. 1975), Antalya, Turkey (Juteau 8,: Whitechurch 1980) and the Semail ophiolite (Hopson & Pallister 1980; Pallister & Hopson 1981; Smewing 1980a, 1981; Browning 1982, 1984; Dahl et al. 1983), do indeed show that they consist of cyclic ultramafic-mafic units with complex crystallization sequences that can be interpreted as the results of repeated influxes of relatively primitive magma into the chamber or chambers.

WADI

,

,

I W ~ AL HILT;

9 ~

3.4.1 The Layered Series of the Semail Ophiolite The Layered Series of the Semail ophiolite consists of a complex association of interbedded layered peridotites and gabbros, ranging in thickness from less than 0.5 km to nearly 6 km (average c. 2.3 km), that can be traced throughout the length of the Semail Nappe outcrop (Fig. 3.19). In over 25 km thickness of measured sections, 74% of the series consists of gabbros and the remainder is mainly dunite and wehrlite. The layered rocks rest disconformably on the underlying Mantle Sequence along the Petrological Moho (Section 3.3) and pass upwards with a gradational contact into the overlying nonlayered High-level Intrusives (Section 3.5). The Layered Series forms rugged mountains with comparable summit heights (1000-1800 m) to the adjacent Mantle Sequence terrain, but with a more rounded profile. The gentler slopes are covered in large boulders of a near in situ scree deposit that are coated in a red-brown desert varnish. On the steeper slopes on the sides of deeply incised wadis, massive outcrops of fresh rock are generally found.

,.

9

"~ "h [ ~ ~ ,~

WADI BANI

Fig. 3.20. Variations m the attitude of cumulate layering in the Layered Sequence. The map shows the limits of the gabbro outcrop with the generalized trends of the layering compared to the strike of the sheeted dyke complex (double lines). Stereograms show equal area plots of poles to layering in selected areas where the attitude of the Petrological Moho (cross) is known (open circles show inverted layering in the Wadi Ragmi area) Data taken from Browning (1982), Bartholomew (1983) and Oman Ophiolite Project 1:100,000 maps.

3.4.1.1 Structure The most persistent and characteristic feature of this series is the rhythmic igneous layering on scales of 0.5 cm to about 2 m. The layers are mostly mineral graded although isomodal layers do occur. Layer boundaries are defined by sharp ratio or, more rarely, phase contacts (Plate 3.16). Individual layers can be traced for up to several hundreds of metres along strike but invariably pinch out or are truncated. The layering dies out upwards and becomes irregular in orientation at the top of the sequence. The transition zone with the high-level intrusives contains compositionally homogeneous gabbros which show only a weak igneous lamination.

Plate 3.16. Rhythmic layering on a cm scale in olivine-gabbro cumulates. The dark bases to each unit are rich in olivine and clinopyroxene.

As noted by Smewing (1980a) and Browning (1982), the attitude of the layering is variable in any one area or section and is only rarely concordant with the plane of the underlying Moho and the dip of the overlying ophiolite units. Browning (1982) and Rothery (1983) both showed that the layering is more conformable at the base near the Petrological Moho and become steeper in dip and more discordant towards the top of the sequence. This upward steepening, they suggest, reflects the shape of the walls of the magma chamber. Fig. 3.20 shows the generalized strike of the layering in the area between Wadi Hatta and Rustaq along with stereographic plots that show the variability of the layering attitude in local areas. (The effects of syn- and post-emplacement folding have not been taken into account and may be important in the Wuqbah, Haylayn and Rustaq blocks where there has been folding of the ophiolite nappe on NW-SE axes). The angle between the generalized strike of the layering and that of the sheeted dykes, as shown on Fig. 3.20, varies from 0 to 90~ A spreading ridge model with a continuous magma chamber beneath the ridge would predict that this angle should be close to zero throughout. The fact that it is not, is taken to support a multiple magma chamber model of discrete, finite length magma chambers along the spreading axis, the limits of the chambers occurring where the strike of the layering is at high angles to the sheeted dyke trend (Browning 1982; Smewing et al. 1984). Instances of this closure are inferred in the Wadi Shafan area at the SE end of the Sarami block and in the Wadi Bani Kharus section of the Rustaq block (Browning 1982).

The Semail Ophiolite [100]

[010]

69

~)01]

Plane of layering [010] Maximum(-)

OM 2429 (OG)

OM 193 (OG)

OM 33 (OG)

OM 84 (OG)

OM 41 (W)

OM 43 (W)

OM 36 (OG)

Fig. 3.21. Olivine petrofabric diagrams for cumulate rocks (OG - olivine gabbro, W - wehrlite). All more or less undeformed except OM193 and OM33 which have less well-defined [010] clusters. Analyses by N.I. Christenson, University of Washington, Seattle.

Chapter 3

7~

The layered rocks show all the structural and textural features of magmatic cumulates. We use the term 'cumulate' here in a non-genetic sense whilst recognizing that the relative importance of mechanical (sorting by differential settling) and physico-chemical (sorting by differential nucleation, growth rates and diffusion) processes in the formation of layered igneous rocks is in dispute (Campbell 1978; McBirney & Noyes 1979; Irvine 1980). Nonetheless, we consider the cumulate terminology of Wager et al. (1960) to be the most appropriate. Petrofabric analysis of the olivines in undeformed cumulates shows preferred orientations ([010] axes lying perpendicular to the layering and (001) and (100) planes lying in the plane of the layering (Fig. 3.21)) that are consistent with simple settling processes. Preferred orientations of [001] and [100] axes suggest that deposition was additionally controlled by some current activity (Smewing et al. 1984). Small-scale sedimentary structures, such as cross-bedding, load casts, flame structures and slump folds, are found although Smewing (1980a) considers "that they are less common in the Semail cumulate rocks than in most continental stratiform intrusions". Browning (1982) made a detailed study of a 37 cm thick rhythmically graded olivine gabbro layer from Wadi Shafan. It is composed throughout of medium to fine grained laminated meso- and adcumulates. A sharp ratio contact 6 cm from the base marks the boundary between olivine-rich (>54% modal ol) and olivine-poor (<30%) parts of the layer. The upper third contains the least amount of olivine ( < 5 % ) and here there is textural evidence that it has reacted with the magma to form clinopyroxene overgrowths. In the lower part of the layer all three cumulus phases have similar grain sizes whereas in the upper part the grain size of plag > cpx > ol (Fig. 3.22). Plagioclase and clinopyroxene both show reverse size grading in the lower part of the layer whereas olivine decreases in size throughout. There is also an overall decrease in the "'lamination index" (a measure of the degree to which the minerals show preferred orientations) upward through the layer. In terms of variations in mineral compositions, the unzoned olivines show a small decrease in Fo content with height (Foss to Fos3.s, returning to Fos4.6 at the base of the overlying layer),

clinopyroxenes show no systematic variation within a small composition range (Mg',,~.,~ ~,.3) and similarly, the plagioclases show a small but non-systematic compositional variation (Anss.4~s4.6). The lack of large or systematic cryptic variations, plus the evidence for some size sorting and flow differentiation, appears to favour an accumulative origin for the layer, possibly involving a density flow. However, Browning (1982) considers that in situ crystallization, modified by infiltration metasomatism (Irvine 1980), cannot be ruled out as a possible mechanism of layer formation. Infiltration metasomatism is a process in which the cumulus minerals are changed in composition by reaction with the intercumulus liquid that is migrating upward due to the compaction of the layered series.

Plate 3.17. Deformed olivine-plagioclase accumulate. The deformation is most clearly seen in the central plagioclase crystal as thin irregular cracks. The remainder of the rock is olivine with some interstitial clinopyroxene. Field of view ~3mm [XN].

At the base of some sections the layered rocks show evidence of sub-solidus ductile deformation (Plate 3.17) giving them tectonite fabrics that are generally concordant with those of the underlying Mantle Sequence (Smewing 1980a, Browning 1982: Bartholomew 1984: Smewing et al. 1984). Usually

O_ -, 71 t'L\\Y/~,\\'~,,~ '//,\\

\ : i /~ s " - _ \ . ~

txl

,\_\,l_\,x,l_

-.

'Top of Layer

X-/k-

\\-:-,,

I

/

9,7,',_',~;,_,,':, '--.",","_'\'-,.,,,-,,-4

./ If,,

PX

:

PLAG

,/ /"

" ~ L " L't'-" \ / \ ' ,

.,-c'.,;,/"-.,.','-.. --

50

lii~il'~

l--ll

'

Ratio contact Base of layer

I!~il) i-,c-, &',', ii~~i,-';

1

0

50 100 Modal proportions (%)

0

1 2 Maximum diameter (mm)

0

1

2

Minimum diameter (mm)

"Lamination index" (plag)

Fig. 3.22. Modal, grain size and "lamination index" wtriations through a rhythmically graded olivine gabbro laver in Wadi Shafan (taken from Browning 1982) In the grain size variations plots olivine is represented by the solid line, clinopyroxene by the dot-dash line and plagioclase by the dashed line.

The Semail Ophiolite this deformation affects only the lowermost few metres of the cumulates but can sometimes be traced for up to 1000 m above the Petrological Moho. The igneous layering in the deformed rocks is usually preserved although the minerals are visibly strained and usually show some evidence of grain size reduction. Olivine petrofabric analysis (Fig. 3.21) of tectonized gabbros and peridotites show orientations of the crystallographic axes consistent with moderate to high temperature (>900~ deformation (Bartholomew 1983: Sinewing et al. 1984). The tectonite fabrics always die out up section in the layered sequence which suggests that they were produced at or close to the spreading axis by the same processes that formed the tectonite fabrics in the Mantle Sequence (Section 3.2.4). In some places the layered rocks are cut by sub-vertical ductile shear zones up to 5 m long and 5 to 30 cm wide (Plate 3.18). These generally show displacements of the layering of only a few centimeters. Amphibolite facies recrystallization in the shear zones gives rise to mylonitized hornblende gabbros with porphyroclastic textures (Plate 3.19). Smewing (1980b) relates these structures in the Wadi Ragmi area, where they trend NW-SE and have a sinistral shear sense, to Riedel-type

7I

shears associated with movements along a probable transform fault zone. Cross-cutting intrusives in the Layered Series include irregular bodies of wehrlite, gabbro and pegmatitic gabbro, that are most likely mobilized cumulates rich in interstitial fluid that were squeezed out of the crystal pile and intruded at higher levels. Where the enclosing rocks are deformed, these bodies carry the same tectonite fabric as their host. Finer-grained basic dykes that cut the layered sequence invariably cut the tectonites although they are both cut by, and cut, the later shear zones. In some areas, these dykes appear to be mostly related to the later magmatic episodes (Section 3.8) although Rothery (1982, 1983) suggests that vertical dolerite dykes, that cut steeply dipping layered gabbros in the Wuqbah block, can be traced from the overlying sheeted dyke complex and are the result of lateral intrusion of dykes parallel to the ridge axis beyond the end of the active magma chamber. Some discordant Late Intrusive Complexes (Section 3.6), ranging from peridotite to trondhjemite in composition, cut the upper part of the Layered Series and in the Haylayn area the lower cumulates are cut by sheets of biotite granite (Section 4.3.2).

Plate 3.18. Small-scale ductile shear zone in the cumulate gabbros of the Wadi AI Abyad area. Coin diameter 2.5 cm. From Bartholomew (1983).

Plate 3.19. Porphyroclastic texture of amphibolite facies mylonite from a shear zone in the layered gabbros near Muslaf. From Smewing (1980b). Field of view 1.5 cm.

Chapter 3

72 3.4.1.2 Petrology and mineralogy

The primary minerals in the Layered Series are plagioclase, clinopyroxene, olivine, chrome spinel, orthopyroxene, hornblende and titanomagnetite. Olivine and chrome spinel occur as early formed cumulus grains, the pyroxenes and plagioclase form both cumulus and intercumulus growths and hornblende and titanomagnetite occur as relatively minor late-stage intercumulus phases. Various combinations and proportions of these minerals lead to a great variety of rock types. Dunites (ol + chr adcumulates), wehrlites (usually ol + chr orthocumulates with intercumulus cpx) and olivine clinopyroxenites (ol + cpx ortho- and adcumulates) are the dominant ultramafic rock types but the early crystallization of orthopyroxene in the Wadi Ragmi area leads to a local development of olivine websterites (ol + cpx + opx adcumulates). Some of the ultramafic rocks contain up to 10% interstitial plagioclase. Brown hornblende is also a minor constituent of some ultramafic rocks as an overgrowth on clinopyroxene or as large poikilitic grains enclosing olivine, chrome spinel and, sometimes, clinopyroxene. Meso- and adcumulate textures predominate in the olivine gabbros (ol-cpx-plag cumulates) and the less common two-pyroxene gabbros or olivine gabbronorites (ol-cpx-opxplag) and troctolites (ol-plag). Completely olivine-free rocks are rare and mainly occur in the transition zone with the highlevel intrusives where there are relatively fractionated gabbros containing intercumulus hornblende and titanomagnetite. 9 Wadi Ragmi o Wadi Jizi 9 Wadi Shafan

0.4--

~, 9

MnO 0.3--

j~' A 9

,0.2r- 9 ~ ,

0.1~oC~~

The olivines in the ultramafic rocks are euhedral to subhedrai 1-5 mm sized grains that are often poikilitically enclosed in clinopyroxene, or more rarely, plagioclase and hornblende. In the gabbros the olivines are smaller in size (<3 mm) and tend to be anhedral. They often show partial resorption textures exhibiting a reaction relation with the magma to form clinopyroxene overgrowths. Many of the olivines show partial or complete replacement by secondary phases such as serpentine, magnetite and rarely brucite and talc. The fresh grains are mostly unzoned and their compositions vary from Fo92_85 in dunites, FO9o_83in wehrlites and olivine pyroxenites to Fogtr_67 (mean Fo85) in gabbros. As far as minor elements in the olivines are concerned, NiO (0.02-0.54%) decreases and MnO (0.12-0.48%) increases with decreasing Fo (Fig. 3.23) and CaO is low (<0.05%) throughout. Textural relations show that olivine and chrome spinel crystallized as the frst-formed liquidus phases. The chrome spinels are present in dunites, wehrlites and some gabbros mostly as a disseminated accessory phase but are abundant enough in the dunites to form thin chromitite layers 0.5-2.0 cm thick. In rocks devoid of an intercumulus phase, the euhedral chrome spinels are concentrated at the edges of the olivine grains but, where intercumulus minerals are present, they are usually included in them. The chrome spinels in the layered sequence show a wide compositional range (Mg/(Mg+Fe 2+) = 0.12-0.63, Cr/ (Cr+AI) = 0.43-0.63) and tend towards higher Fe/Mg values than the spinels from the Mantle Sequence (Fig. 3.24). Individual grains are often complexely zoned, particularly with respect to Ti, Mn and Cr. There is no consistent relationship between chrome spinel compositions and the Fo contents of the coexisting olivines (Browning 1982). Clinopyroxenes in the cumulate rocks vary from small (<5 mm) lath or prismatic shaped crystals in gabbros to larger (5-20 mm) equidimensional grains in ultramafic rocks which poikilitically enclose earlier-formed olivines and chrome spinels. In the wehrlites, websterites and pyroxenites they are mainly diopsides (Mg'(100 Mg/(Mg+Fe 2+) mol %) 94-84) and in the gabbros they are diopsidic augites (Mg' 9z-74). The complete and rather restricted range in cumulus compositions

~

100

I

I

1

I

I <+ 80

0

NiO

0.3

/',

9

O

I

8

oo

9

~ 60

--K) 9

o t~~149 # 0

\

o

0 CO

0 9163

i

\ ,

O

0.4 O 0

Wadi Ragmi ~ 9Dunites 4- WehrHte Wadi Shafan 9Mainly gabbros i u Dunites Rustaq block 0 Wehrlites c Harzburgites r

-,aNn

amnnnm.~._ 9 9

O

AI~I~jhOO

~

9

ffMantesequencechrome Q

0.2 40-

k ~

0.1

kk

9 AkAA A

I

I

1

I

90

85

80

75

9

I 70

Fo Fig. 3.23. Minor element (MnO & NiO) vs Fo plots for cumulus olivines from the Layered Series.

/4

..... --.--

2O

I 60

Chrornitites Dunites Har-zburgites

I 40 Mg' ~00 Ug/(ag + Fe~

I 20

Fig. 3.24. Chrome spinel compositions from the Layered Series.

73

The Semail Ophiolite

plagioclase is often altered to a-fine-grained mixture of Ca-A1 silicates clearly associated with the serpentinization of the olivine; in the gabbros it is invariably fresh. The full range of feldspar compositions in the layered sequence is An96 to An60. They are usually unzoned or show slight reverse zoning (rarely more than 2 mol% An). K20 contents are low (<0.06%) and FeO varies from 0.2 to 0.55% but shows no systematic relationship with An. In 400 m of basal peridotite and gabbro cumulates in the AI Maydah section (western Rustaq block) the cumulus and intercumulus plagioclases have overall ranges of An,~2 ss and Ans~77 respectively (Browning 1982). In 1400 m of gabbros in Wadi Bani Kharus the plagioclases range from a mean value of Ango at the base to An76 at the top of the section (Browning 1982). Brown titanium (2-4.5% TiO2) hornblende (Mg'v~-ss) occurs as an occasional primary intercumulus phase in the wehrlites and in gabbros as an overgrowth on clinopyroxene. The brown hornblende is partly replaced by colourless (low Ti) more magnesian hornblende (Mg'~s-9o) which retains the high AI and Na contents of the original. Green hornblendes and actinolite hornblendes with < 2 % TiO2 are secondary after clinopyroxenes which for the most part show only minor replacement along cleavages and grain boundaries. Only in the upper part of the layered gabbros are the pyroxenes completely actinolitized and late-stage veins filled with actinolite occur. Fig. 3.27 shows the complete range in amphibole compositions from the Layered Series.

is from Wo47En49Fs 4 to Wo41En45Fs14. With decreasing Mg' the clinopyroxenes show small minor element variations with decreasing Cr203 (max. 1.14%) and generally increasing TiO2, MnO and Na2 O contents (Fig. 3.25). Zoning of individual grains can be highly complex with a tendency for reverse zoning in Mg' and TiO2 and both normal and reverse zoning in Na20 (Browning 1982). Cumulus orthopyroxenes are largely restricted to the Wadi Ragmi section, where they crystallized after clinopyroxene but before plagioclase, although late-stage cumulus orthopyroxene (Mg's0) occurs after plagioclase in the Wadi Bani Kharus gabbros (Browning 1982). The orthopyroxenes in the lower part of the Wadi Ragmi section range from Mg'ss_7] and have a maximum W 9 content of 3.4 mol%. Minor element contents are generally low; A1203 < 1.88%, MnO < 0.32%, TiO2 < 0.21% and Na20 < 0.05%, with TiO2 and Na20 lower than in the coexisting clinopyroxenes. Cr203 in orthopyroxenes ranges from 0.48% to 0.06% and is markedly less abundant in intercumulus than cumulus grains. Cr20 3 shows a poor positive correlation with Mg', whereas TiO2 and MnO have negative trends (Fig. 3.25). Co-existing ol-cpx-opx assemblages from the Wadi Ragmi section show near-parallel tie-lines (Fig. 3.26) suggesting equilibrium crystallization during melt fractionation. Plagioclase occurs as an intercumulus phase in some ultramafic cumulates and is an abundant (up to 70% modal) cumulus phase in all gabbros. In the feldspathic peridotites the (=) 0.6O

9

(b) o.41 _

0.3-

3

TiO2 0

0.2~176176 ~~

9

0.2-

A 9176176 9 I

9

0.~

I

1

o

I

I

I

i

I

~O

O

o.1;

0

0.4

I

0.30.2-

0.6

9

9

I

0.g MnO 9. ".L ". - . _ . . . ~,, . ~ ' - F a" .. , , ~ 0.1 1

O9

IO 00 9

o 9 j,.

N

.'.; ".:.

9

I

~a20

9

9 :

1

0.2 GO

I e 1.4

--O

9

I

I

I

0.4-

Oo

03-

000 9

1.0

o

9

,Z

o

at 9

0r203 C'

9

9 A 9

of,"

0.2-

kO

06-

9

o.1-

&O 9

9

9

O9 O9

9

0.2I 90

I 85

A

>

Mg'J-lO0 Mg/(Mg .Fe)]

I

8O

I

75

I 85

9

1 80

I

75

9

I

70

Mg'~O0 Mg/(Mg+ Fe)~

Fig. 3.25. Minor element (TiO2, MnO, Na,O & Cr:O3) vs Mg' plots for cumulus pyroxenes (a) clinopyroxenes, (b) orthopyroxenes from the Layered Series (symbols as in Fig. 3.23).

Chapter 3

74

Wadi Shafan

Wadi Ragmi

/ ..,. y

7

.\,m/ \\

OO

..../

/ Field of cpx's / from Wadi Ragmi Wadi Jizi

~,~,

Hd

Di

FS

En

En

f

FO

""

90

"" 'i 8O

7O

Fig. 3.26. Layered Series pyroxenes plotted on the pyroxene quadrilateral. Coexisting ol-cpx-opx compositions from Wadi Ragmi connected by tielines.

%

[]

1.0

[]

!I

Q. O O. _..--

o

s ~,

E

v

,

--Is

Z

0.5

sJ ~

s

/

/

0.93 9

I

~-,__

0"?,~s0.$71 ~

ss J

sJ 0.64 e

,,s'~

~

s 9 0.17

.....--"

~

,,

0.78

P

4.84 9

3.36 e tt 117 1.67e 1.95 / " .~ -~n

0.06

[]

~

1.74 e~l~, - U n n7 .... :"\ -0 09 3.27_--~" 3.70 " ,~ "2.88

~ ",' u.u,,-, 0.56_l-10-11 013 "1-10.11 I ~ 0.26 9 9 9 ,s 0.24

-'.4u

0"04ee0"69""'s

0 73 " 9 - . , . . " ; ~ ' - ,

~s S s

/' 1 25

s ~

A-T

\

s S

s ss

e l .00

0

1533 "

9

0.31

]

0.5

I 1.0

I 1.5

TS rh 2.u

AI (mol prop) Fig. 3.27. Compositions of amphiboles from the Layered Series (dots) and Mantle Sequence (open squares). TiO2 contents (wt 9~-) given by each analysis. End-member compositions: A-T Actinolite-tremolite, TS - - Tschermakite, E - - Edenite and P - - Pargasite. Field of amphiboles from high-level and late intrusive complexes outlined.

The Semail Ophiolite

75

Table 3.8. Analyses of layered series rocks Olivine clinopyroxenites

Dunite

Wehrlites

0M944

OM1210 0M1216 0M1214 0M1212 0M1206 0M942 5f Sf Sf Sf Sf Sf

sf

Olivine gabbros 0M930 R

0M1205 Sf

Pvroxene gabbros

0M1207 0M929 Sf R

0M1215 Sf

SiO2 TiO~ A1203 Fe203 FeO MnO MgO CaO Na20 K20 P2Os LOI

40. l0 0.20 7.31 2.64 5.24 0.12 30.90 5.79 0.29 0.03 0.01 7.52

46.20 (/.25 4.21 2.14 5.96 0.15 24.83 13.35 0.26 nd nd 2.77

48.10 0.19 2.60 2.66 5.69 0.15 24.73 14.13 0.16 nd 0.01 2.5(/

49.40 0.22 2.76 1.10 2.75 0.09 23.23 18.00 0.2(/ nd 0.09 2.45

50.70 0.16 1.47 3.21 0.90 0.10 22.83 18.38 0.12 nd nd 1.35

48.30 0.17 12.50 1.18 3.85 0.10 16.22 16.06 0.48 nd 0.01 1.83

46.00 0.19 16.60 0.79 3.11 0.08 14.04 14.90 0.68 nd nd 3.68

49.30 0.28 15.30 1.17 5.56 0.14 11.11 14.90 1.45 nd nd 0.79

48.40 0.48 15.70 1.50 4.52 0.13 10.11 13.97 1.77 0.06 0.02 2.61

Total

100.42

100.36

100.92

99.20

99.72

100.70

100.01

100.00

99.27

0.18 2.45 18.53 7.93 7.88 51.30 3.82 (/.38 0.02

2.26 10.6(I 45.36 5.49 32.61 3.19 0.49

1.34 6.32 49.76 11.89 24.01 3.82 0.36 0.02

1.69 6.63 64.45 1.29 22.01 1.10 0.42 0.21

1.01 3.47 68.92 6.52 16.33 1.31 0.30

4.03 31.73 37.39 9.62 13.37 1.7(/ (/.32 (I.02

5.75 42.24 25.(/(I 6.73 15.15 1.15 0.36

12.27 35.24 30.98 9.84 8.65 1.70 0.53

0.36 15.09 34.98 27.67 9.54 6.58 2.19 0.92 0.05

1210 57 2 440 9 7 169 42 6

4220 10 4 532 8 6 106 14 6

835 17 10 210 68 6 95 26 6

555 27 I0 310 169 5 60 18 7

71 274 69 13 125 126 12 129 31 9

35 570 67 13 163 156 13 187 28 26

or

ab an

di hy ol mt i[

ap Co Cr Cu Ga Ni Sr Y V Zn Zr

122 2670 4 2 1650 2 2 3 32 4

3530 45 6 1040 34 4 47 40 6

2250 26 4 1200 50 4 32 42 9

969 40 6 1390 58 6 60 44 16

46 3225 52 2 426 6 4 91 15 4

Rare earth elements La 0.2 (0.61) 0M930 0.4 (1.22) 0M929 0.6 (1.83) 0M1215 1.2 (3.66) 0M944

OM944 OM1210 OM1216 OM1214 OM1212 OM1206 OM942 OM930 OM1205 OM1207 OM929 OM1215

Ce [2.2]

Nd nd

nd

nd

1.3 (1.50) 2.9 (3.35)

[0.6] 3.1 (4.92)

Sm [0.3]

Eu [0.02]

Gd nd

Tb [0.02]

Tm [0.02]

0.2 (0.98) 0.2 (0.98) 1.1 (5.42)

0.04 (0.52) 0.27 (3.51) 0.49 (6.36)

nd

[0.04]

[0.07]

0.08 (1.54) [1.8] 0.3 (5.77)

[0.05]

[0.3]

[0.21]

Yb 0.21 (0.95) 0.16 (0.73) (/.28 (1.27) 1.14 (5.18)

Lu nd

Hf nd

Ta 0.04

Th 0.22

Sc 4.4

[0.03]

nd

nd

[0.09]

46.1

0.04 (1.18) 0.17 (5.00)

0.36

nd

[0.03]

17.7

0.67

0.03

[0.12]

41.4

Ol-chr adcumulate (dunite) ol (Fog0) 99, chr 1 Ol-chr orthocumulate, intercumulus cpx (wehrlite) ol (FOsT) 57, opx (Mg',n) 42, chr 1 Ol-chr orthocumulate, intercumulus cpx-hbl (wehrlite) ol (Fos7) 57, cpx (Mg's~) 35, hbl 7, chr 1 O1 orthocumulate, intercumulus cpx (wehrlite) ol (Fos6) 54, cpx 46 Ol-cpx orthocumulate, intercumulus cpx-plag (plagioclase wehrlite) ol (Fos2) 30, cpx (Mg's6) 67, plag 3 Ol-cpx mesocumulate (olivine clinopyroxenite) ol (Fos~ 5) 21, cpx (Mg'85) 79 Ol-cpx adcumulate (olivine clinopyroxenite) ol (Fos9) 10, cpx (Mg'~2) 90 Ol-cpx adcumulate (olivine clinopyroxenite) ol (Fo8s) 15, cpx (Mg'~) 85 Ol-cpx-plag adcumulate (olivine gabbro) ol (Fos3) 11, cpx (Mg's6) 48, plag (Ans7) 41 Ol-cpx-plag adcumulate (olivine gabbro) ol (Foss) 10, cpx (Mg'ss) 45, plag (AnsT) 45 Ol-cpx-opx-plag adcumulate (two-pyroxene gabbro) ol (Fos,) 5, cpx (Mg's6) 32, cpx (Mg'~2) 8, plag (An79) 55 Cpx-plag adcumulate (gabbro) cpx (Mg%) 50, plag (An75) 50

Key to locations of samples: Sf = Shafan, R = Ragmi. ( ) = Chondrite normalized, [ ] analysis unreliable. Small ( < 2 m m across) p r i m a r y grains of t i t a n o m a g n e t i t e with ilmenite exsolution lamellae occur in the most differentiated gabbros chiefly f o u n d in the transition zone at the top of the l a y e r e d sequence. Fine grained (0.02 m m ) s e c o n d a r y mag-

netite is a c o m m o n b r e a k d o w n product of altered olivines and pyroxenes. T h e highest t e m p e r a t u r e alteration in the layered sequence is r e c o r d e d by the partial r e p l a c e m e n t of olivine by talc +

76

Chapter 3

magnetite and the pyroxenes by green hornblende. In their upper part, the layered gabbros are increasingly affected by a pervasive actinolite facies alteration (c. 450~ with clinopyroxenes partly replaced by actinolite and plagioclase largely unaltered, that is continuous with the alteration of the overlying high-level gabbros (Section 3.5). The layered rocks are also cut by late-stage greenschist facies (c. 350~ veins with fibrous actinolite, epidote, quartz and Fe-Cu sulphides. Grego~ & Taylor (1981) demonstrated that the alteration of the layered gabbros is the result of hydrothermal circulation penetrating deep into the crustal section and that even fresh looking rocks show plagioclase-clinopyroxene oxygen isotopic disequilibrium due to sub-solidus exchange with heated isotopically modified seawater.

OM 1215

*0

3.4.1.3

Geochemistry

Whole rock analyses of a variety of rock types from the Layered Series given in Table 3.8 show that there is a wide range in compositions from dunites with more than 40% MgO, though wehrlites and pyroxenites, to gabbros with as little as 10% MgO. Compatible trace elements, such as Ni and Cr, that are concentrated in early-formed olivines and pyroxenes, decrease from highest values in the ultramafic rocks to lower but still signifcant contents in the mafic cumulates. With a few exceptions, all the rocks have extremely low contents of incompatible elements (K, P, Zr, Y, REE) which suggest that they are dominantly adcumulates from which all or most of the intercumulus melt has been expelled. This is supported by the REE patterns for two cumulate gabbros which show low concentrations of rare earths (1-3 x chondrite) and positive Eu anomalies indicative of plagioclase accumulation (Fig. 3.28). The ultramafic rocks have even lower REE concentrations, less than or equal to chondrite.

3.4.2 Large scale cyclic units, crystallization sequences and cryptic variations

Alternations of ultramafic and mafic lithologies giving rise to large-scale (10s to 100s m thick) cyclic units, are common in the Semail Layered Series and, although predominantly occurring in the lower part of the succession, can occur at high levels in the sequence (Fig. 3.29). Sinewing (1980a) maintains that some of the thickest cyclic units can be traced for up to 10 km along strike but the variation in successions from area to area prevents correlations over greater distances. Fig. 3.29 shows details of parts of the sections in Wadi Ragmi (containing 17 cycles in 700 m of section), Wadi Jizi (6 cycles in 110 m), Wadi Shafan (8 cycles in 1400 m) and Wadi Bani Kharus (14 cycles in 900 m), along with the variations in mineral compositions (taken from Sinewing 1980a, 1981; Browning 1982). The most common order of crystallization of the cumulus mineral phases in the Layered Series is olivine (+ chrome spinel) followed by clinopyroxene and then plagioclase (ol--~cpx--~plag) which gives rise to the rock sequence dunitewehrlite-olivine gabbro-gabbro from base to top of the cyclic unit. Some of the cycles are complete but many are partial with some members of the lithological sequence missing. However, two other orders of crystallization; ol---~cpx--~opx--~plag, giving a dunite-olivine websterite-two pyroxene olivine gabbro-gabbronorite sequence, and ol--~pla'g--~cpx, giving dunite-troctolite-olivine gabbro-gabbro, occur in some areas, e.g. in Wadi

.....~ OM 930 9 OM 944

1

I

La Ce

I

I

Nd

I

I

I,I

Sm Eu (Gd)Tb

I

I

I

(Tm)Yb Lu

Fig. 3.28. REE/chondrite plot for cumulates compared to the high-

level gabbros (shaded field). OM1215, O M 9 2 9 - gabbros, O M 9 3 0 olivine clinopyroxenite. OM944 - - dunite.

Ragmi where websterites are common, particularly in the lower and middle parts of the sequence and in the eastern part of the Rustaq block, where troctolites occur. As has already been noted, there are only limited overall cryptic variations of the cumulus phases in the layered sequence as a whole (olivines FO92_67, clinopyroxenes Mg'94_74 , ortbopyroxenes Mg's~_T1, plagioclases An,~w6o). Within each of the ultramafic-mafic cyclic units there is usually an upward trend towards more fractionated mineral compositions, i.e. lower Fo, Mg' and An contents. The base of each cycle is marked by an abrupt shift to less fractionated compositions (most marked in the Cr'=(100 Cr/(Cr+AI)) contents of the clinopyroxenes) suggesting a sudden influx of relatively primitive melt into the chamber. Browning (1982, 1984) showed that these "zig-zag" patterns of mineral composition variation are also found in the 900 m thick sequence of three-phase gabbro cumulates in Wadi Bani Kharus (Fig. 3.29). Here, there are three major cycles each showing positive gradients for incompatible major and minor elements (Mn in olivine, Na, Ti and Mn in clinopyroxene) and negative gradients for compatible elements (Fo and Ni in olivine, Mg' and Cr' in clinopyroxene). Each of the major cycles comprises several minor cycles most of which show the same "normal" trends as the major cycles but some (e.g. cycles 9 and 10 in Fig. 3.29) show reverse trends. Browning (op. cit.) suggests that these reverse trends are due either to eruption of evolved magma or a small primitive magma input leaking into the chamber during fractionation. In the case of compatible parameters, there is a progressive decrease in their value in each successive reset at the beginnings of both the major and minor cycles. Browning (op. cit.) noted that these trends are characteristic of a magma chamber undergoing open system fractionation and tending towards a steady state composition.

The Semail Ophiolite

77

I

I I

I

/

o

t I I I I I I

I I I I I I I I

I I I I

I I I I

,

i / , ~

~~i

I I I !J._

-~ "~ '

I I I I

I I I I

I I I I

I I I I

I I I I

~

L.,

I

,

i

I

/r,.i

I I I

I I I

I I

i I

"~"

~

I

I I I I

I I I

I I I I

1

II

,I

I

1~

,I

I

I I I I I

I

I

I I I I

I I I I

II I ...J o I~I ~ _Ill

-,,_ "~

i

_/=..IM

o,~,

' ,']"~/"

I I

i

I I I

I.

I

~--[o I I I/

, I

,~1

I--i~. iiIu ,~ ~:~ ..XX

-

I I I I

I l I !

I

I I

-I

I

II

L.

I

-

I

"i'I/

i'~l~

,X::

i

I

I

*-,

'

~X

/I

"

I I I

~

I

I I i I I

I I

I I

I I

I ~ I I

I ~ I

,

I ill

r-~c ) 17 (D I-I I

m

e"

,~ I

I I

IJ

I I I I I I

I t

9

I I I ! i

~

ip

~- ~

,

,

I

I

I~

I~I

tt "t....-4.~L~oJ../ I-

,

I

I"

I ~ I

I

I

I

I

I

I

I

I

I

I

I

I

I

I

I

I

,

G9

<-'r

~

it

~

~

I

~

,

,

I~176

,

~

,

"~I~I

' .

.

.

.

t,:q

I ,-,I I I

,

U_

~~-

,i ~

~

l'-

~o

-~

~.~

._

o

f.,,

~

T-

(w) ,,oqol~l. a ^ o q e lq6!aH

I I

I I

I

I

, ~i..~

I I I .I

I

I

|

I

III~

I

I I I II

I

~ I

.',

'.I

~(.~.

1 o

_o

b~ 2 T--

1 I

I I I I i

I

I

1 ~oa

I --

.--

cO

cO

(D

I --

2 2 8 I I

I I I

I

I

I

I

III

1

I

lilJ_ii iTfll O rr

il

"F cr~

~ (W)

II

llll I

I

~

w

,OLIOS, '

I

'~' I

II II

II

--:::: . . . . . . . . . . .

~

~

II

I

I I I

if:, ii]iIl

: : : : : : iiiNiiiliiii~

aAoq'e l q G ! a H

I r

I I I

I

~

~o

~

I I I

.{D .0

0

I~--=:~ :]":i~fll !IHN

-~ CO

I

I

~

~

w

~ o

w ,,oqol,~,, a ^ o q e ~.qf!aH

0

o

~

~

~

CD

78

Chapter 3

3.4.3 Magma chamber processes 3.4.3.1 Origin of the cyclic units The cyclic nature of the Layered Series points to open system fractionation with breaks in fractionation trends and resets in mineral chemistry to less evolved compositions resulting from repeated influxes of unfractionated magmas. The strongly bimodal character of most of the ultramafic-mafic units, with a sharp boundary between the peridotite and gabbro members is marked by a sudden appearance of abundant cumulus plagioclase. This has been explained by Browning & Smewing (1981) and Smewing (1981), based on the density contrast model developed by Huppert & Sparks (1980b) which maintains that when a relatively dense picritic magma is injected from below into a basaltic magma chamber, the picritic melt is temporarily ponded at the bottom of the magma chamber. It there undergoes olivine crystallization and accumulation which produces a dunite layer that forms the basal part of the cyclic unit. Fractionation of olivine causes the density of the lower liquid layer to be reduced. When its density becomes equal to that of the overlying evolved magma, mixing takes place. Three-phase gabbro cumulates then form to produce the mafic part of the cyclic unit. In the case of the Wadi Bani Kharus cycles, where gabbro cumulates were precipitated throughout, it seems that the successive pulses of more primitive magma that entered the chamber were already sufficiently differentiated to mix immediately with the existing magma (Browning 1982, 1984).

3.4.3.2 Magma chamber size and shape The lack of major discordances in layering or intrusive contacts in the Layered Series means that it is not possible to identify precisely the edges of individual magma chambers. However, the evidence from the discordance between the strike of the layering and the trend of the Sheeted Dyke Longitudinal intrusion of axis dyke into steeply dipping cumulates Centre of older magma chamber

Complex plus the non-perpendicular relationship with mantle lineations suggest that discrete magma chambers existed along the length of the spreading axis. This model is supported by the occurrence of dykes from the sheeted dyke complex that cut steeply dipping layered cumulates (Rothery 1982, 1983) suggesting that an earlier magma chamber had cooled and solidified before a later chamber developed beneath another part of the ridge axis. As to the cross-strike width of the chambers, Browning (1982) found an average dip of the lowest part of the layered sequence in Wadi Bani Kharus of 36~ relative to the moho datum which, for a 2 km thickness of cumulates, suggests a maximum half-width for the magma chamber of 2.75 km and for a 4 km thickness a maximum halfwidth of 5.5 km (assuming constant dip). This is an area where the strike of the layering is at a high angle to the direction of the spreading axis and therefore probably lay near to the edge of a magma chamber, which may account for the relatively small calculated widths. Smewing (1981) considered that, in the Wadi Ragmi area, individual cyclic units can be traced for up to 10 km parallel to the spreading direction and from this inferred a minimum halfwidth for the magma chamber of 5 km in this area. Pallister & Hopson (1981) suggest a half-width for the magma chamber of 18 km, based on an average 15~ vergence between the dip of the layers and the contact with the underlying mantle tectonites. They were unaware of any upward steepening of the layering in their section, although they conceded that there were considerable variations in dip, and proposed a broad, funnel-shaped cross-section for the chamber, similar to the model of Greenbaum (1972) for the Troodos ophiolite. This model assumes a more or less constant angle of dip throughout the cumulate sequence. In most places in northern Oman, however, the layering usually steepens up section relative to the general dip of the ophiolite (Smewing 1980, 1981; Browning 1982, 1984; Rothery 1982, 1983; Sinewing et al. 1984) suggesting a bowl-shaped flat-bottomed chamber with steep

Compositional layering relating to an older, solidified magma chamber \

\

\\1

\~1~

IIII II/I ///

Extrusives Sheeted Dyke Complex High-level Intrusives

/...

Edge of magma chamber

Compositional layering 9 in Cumulates

Active magma cham

Active ridge plane

Fig. 3.30. A magma chamber model (modified after Browning 1982. 1984).

The Semail Ophiolite upwardly concave walls (Fig. 3.30), similar to that proposed by Casey & Karson (1981) for the Bay of Islands ophiolite. This leads to a smaller chamber width than that calculated from the constant dip model; for example, upward steepening of the layering to near-vertical in the cumulate gabbros of the Wuqbah block led Rothery (1982) to suggest a chamber half-width of only c. 2.5 km which, in view of the uncertainties involved, we believe to be a reasonable, if minimum, estimate. As to the lengths of the chambers along the ridge axis, we have no reliable data to constrain these but suggest a maximum of 10 km.

3.4.4 Fractionation trends and primary magma compositions Although it has previously been recognized that the order in which the major cumulus minerals crystallize differs from one ophiolite complex to another (Church & Riccio 1977), it is clear from a regional study of the Semail ophiolite that these orders can change even within a single complex. In addition, the commonest sequence of crystallization in most ophiolites (e.g. Vourinos (Moores 1969; Beccaluva et al. 1984), Betts Cove and Bay of Islands (Church & Riccio 1977; Malpas 1978) and Marum, New Guinea (Jaques 1981; Jaques et al. 1983), as well as the Semail) is of clinopyroxene and sometimes orthopyroxene crystallization preceding plagioclase which is not compatible with the low pressure crystallization behaviour of MORB magmas (Stolper 1980). Browning (1982) discusses the various mechanisms that may be responsible for the different crystallization orders shown by the Semail Layered Series. He rejects variation in P,ot~ as it is likely that similar thicknesses for the Layered Series in the west and east of the Rustaq block reflect similar crustal depths to the Petrological Moho. The observed abundance of intercumulus amphibole is contrary to that expected if the crystallization order was controlled by Pn_~o and he concludes that the order must reflect contrasting parent magma compositions. The most common ol-->cpx---~plag sequence probably results from magmas with higher CaO/ AI20 3 ratios than those which produce the less common ol---~plag---~cpx (MORB) sequence. The early crystallization of A

Q4'/ ~

^_

20 Kbar

-~" "~ " " - ~ - - ~

CMS2 50 DIOPSIDE

40

30

~ CaO/AI203 = 0.65~

20

10

ENSTATITE

79

orthopyroxene apparently represents an even more CaO-rich and Al20~-depleted parent magma. These conclusions are based on the phase relations in the CMAS system (O'Hara 1968) (Fig. 3.31). Olivine fractionation alone cannot alter the CaO/AI20~ ratio and, as this is the main process occurring as the melts rise through the upper mantle, the contrasting primary magma compositions are most probably produced at the site of mantle partial melting. They may be caused by different degrees of melting of a homogeneous mantle or heterogeneities in the composition of the source mantle. Source heterogeneity may be an original feature or the result of an earlier melting episode which depleted the source region particularly in A1203. In general, it seems that the magmas that underwent ol----~plag-->cpx crystallization were derived from a less depleted or more '~fertile" source than those which display the ol-->cpx-->plag and particularly the ol-->cpx-->opx-*plag crystallization sequences (see Section 3.9 where a quantitative model of the partial melting process is developed).

3.5 T h e H i g h - L e v e l Intrusive In most complete ophiolite complexes, between the underlying layered cumulate sequence and the overlying sheeted dyke complex, there is a generally thin and irregular unit of "highlevel" gabbros that are distinguishable from the underlying plutonic rocks by their variable texture and the absence of well-defined igneous layering. These rocks were first described by Wilson (1959) from the Troodos Massif on Cyprus and were interpreted by Bear (1960) as having been produced by a basic melt chilling against the roof of a magma chamber. More detailed studies of these rocks on Troodos by Allen (1975) supported the roof-chill origin, and this interpretation has been extended to similar rocks in other ophiolite complexes (Coleman 1977; Dewey & Kidd 1977). Reinhardt (1969 and in Glennie et al. 1974) described a unit of medium grained, non-layered gabbros in the Semail ophiolite and termed it the "hypabyssal gabbroid unit". Aldiss (1978) showed that the Semail high-level gabbros were very similar to the corresponding rocks of the Troodos complex and that they contained small bodies of differentiated intermediate to acid plagiogranites (we follow Coleman & Peterman (1975) and use the general term plagiogranite to collectively describe the quartz diorites, tonalites and trondhjemites found in the upper levels of ophiolite plutonic complexes). Aldiss (op. cit.) proposed that these rocks formed by the crystallization of hydrous melts in a magma chamber roof zone where magmatic volatiles were concentrated and seawater convecting through the oceanic crust entered the magma via the stoping of hydrothermally altered blocks of sheeted dykes. The High-level Intrusives are distinguished from the petrologically similar but later cross-cutting plutonic bodies, the "Late Intrusive Complexes" (described in Section 3.6), by their stratigraphic position between the Layered Series and the Sheeted Dyke Complex and by their non-intrusive but generally gradational contacts with these units.

3.5.1 Field relations Fig. 3.31. Projection from olivine onto the CS-MS-A plane of the CMAS tetrahedron (O'Hara 1968) showing the compositions of the Semail harzburgites (dots), Lherzolite (open circle) and Browning's primary magma composition (triangle). 1 bar and 20 kbar phase boundaries shown and projected fractional path of the Semail magmas (dashed arrowed lines) (from Browning 1982).

The Semail High-level Intrusives form a discontinuous unit, up to 700 m thick but generally less than 200 m and sometimes only a few metres thick or absent, at the top of the plutonic sequence (Fig. 3.32). The sequence of rock types is often complex and variable from section to section. The transition

80

Chapter 3

-F

111111 7.';',"-'

T i

"1 \ I - L

-i

i,

+++4-+

-;:b~

",, t ' - x

~

'L" ( 5 - ' , "

t

2,'_.'2'z~l

T

,-2 ,',, C

I i " ; l ~ / Ii

;' -",)~d

:;.,',.1 r,-',,-t.

x/i

_-

i

_~.~,

,?;7:7? + +§ ~ § .,~ 9,4"I- i , l ^ \ \\\\

\', \ \ "1

+

Xx\\\/

"-'\ \ \ \

I "/~',',",\

t

it

.'-',jiLl

mt

"- 7".. '.',q

'.: ,','.d

I

ol

ol ~ - ~ Dykes

ol

•

Isotropic gabbro

~

Laminatedgabbro

ol

Layeredgabbro

I magnetite-bearing gabbros mt

m ~ Plagiogranite ~ Xenoliths ~ ~ Coarsepegmatitic gabbro-diorite

I olivine-bearinggabbros

from layered to non-layered rocks occurs gradational!y over a transition zone a few tens of metres thick where the compositional layering, which is often steeply dipping in the uppermost cumulates, dies out up section. There usually are no marked mineralogical changes across the contact but olivine is rare in the high-level gabbros and disappears up section where titanomagnetite becomes more abundant giving rise to ferrogabbros with up to 20% opaque oxides. Hopson & Pallister (1981) recognize an intermediate unit of laminated but non-layered olivine gabbros and gabbros between the cumulate gabbros

Plate 3.20. (a) Wadi Salahi, plagiogranite bodies at the base of the Sheeted Dyke Complex. High-level gabbros poorly exposed in the foreground.

Fig. 3.32. Representative sections through the high-level intrusive sequence (modified after Browning 1982).

proper and the high-level gabbros but we find that patches of laminated gabbro occur throughout the high-level sequence and frequently grade into non-laminated rocks. The plagiogranites occur as light-coloured veins and dykes throughout the gabbros increasing in abundance upwards where they coalesce to form larger 5-50 m thick sheet-like bodies in the upper part of the sequence (Plate 3.20(a)). Here they locally intrude into the base of the sheeted dyke complex although they may in turn be cut by later basic dykes (Plate 3.20(c)). The largest plagiogranite bodies are characterized by sharp,

The Semail Ophiolite

intrusive contacts and often contain xenoliths of gabbros or, more commonly, metadolerite with angular to rounded shapes and representing partially assimilated blocks of dyke stoped from the overlying sheeted complex. In addition, some plagiogranite bodies consist of several distinct intrusions that can be identified by differences in grain size, textures and xenolith populations.

3.5.2 Petrology 3.5.2.1 Gabbros and diorites The high-level gabbros are characterized for the most part by medium grained (1-5 mm) hypidiomorphic to ophitic textures (Plate 3.21). They are non-layered but occasionally laminated with planar orientations of the constituent minerals, particularly plagioclase. Pegmatitic segregation patches and veins, with up to 5 cm grain size, are common and have diffuse, gradational or sharp and cross-cutting contacts with the finer grained host. The pegmatites contain patches of leucogabbro and diorite which Aldiss (1978) interprets as residual liquids that became trapped in the solidifying gabbro. The high-level gabbros are composed largely of unaltered zoned plagioclase (An84_5~) and partly altered diopsidic augite (Wo46.5._49.5En4~48Fs5.5_lO.5). In contrast to the cumulate gabbros, where zoning of the mineral phases is very limited and

Plate 3.20 (Cont.) (b) Fine-grained gabbro-diorite netveined by plagiogranite. (c) Plagiogranite cut by sheeted dykes near Fujayg, Wadi Sarami (Figure for scale). Plate 3.21. High-level gabbro. Zoned plagioclase showing patchy alteration and clinopyroxene showing partial marginal replacement by actinolite. Hypidiomorphic granular texture.

8I

82

Chapter 3

often reversed, the feldspars, and to a lesser extent the pyroxenes, in the high-level gabbros are strongly normally zoned. Brown and green hornblendes with a wide range of compositions (Fig. 3.33) occur as an interstitial phase or as pyroxene overgrowths. The Ti and Al-rich compositions reflect primary magmatic crystallization. The pyroxenes are partly altered to fibrous green actinolite (uralitization) containing inclusions of fine grained sphene and magnetite as additional secondary alteration products. Small (<0.5 mm) altered olivines, often enclosed in clinopyroxene, occur in the lower part of the sequence where the rocks are predominantly Mg-rich and Fepoor gabbros with only 1% or less of primary Fe-Ti oxides. Higher in the sequence olivine-free gabbros predominate and grade into ferrogabbros with 5% to 20% titan 9 (with ilmenite exsolution lamellae) that have grain sizes up to 1.5 mm. The opaques form subhedral grains that crystallized after plagioclase but together with clinopyroxene. They are commonly altered to a mixture of titanomaghemite and sphene. The gabbros and ferrogabbros grade in turn into diorites and ferrodiorites that consist of zoned intermediate plagioclase (Ans(~_30), uralitized pyroxene, prismatic green actinolitic hornblende, Fe-Ti oxides and up to 5% interstitial quartz or microgranophyre. In these intermediate rocks alteration is generally more intense than in the associated gabbros for, whereas in the latter the calcic plagioclase is invariably fresh, in the former it is commonly replaced by a fine grained mosaic of secondary minerals, with prehnite and epidote preferentially occurring in the calcic cores, whereas albite and quartz are found in the outer margins. Likewise, the ferromagnesian and opaque minerals are totally replaced by a fine grained mixture of fibrous actinolite, sphene and magnetite.

E

1.0-

+

[:3P

.' '*) , Green hornblende

Brown hornblende

"Z 0.5

2

u " 5t u

iio

9

Actinolite

0 ~A

I Q5

l 1.0

AI IV

I 1.5

Ti-I 20

Fig. 3.33. A m p h i b o l e compositions from the high-level intrusives. End-member compositions: A - - actinolite, T - - tschermakite, E - - edenite, P - - pargasite (Leake 1978).

TONALITES

70 Si02

6O

GABBROS

& A&

50

9 I

9o

~' eo

99

DIORITES

TRONDHJEMITES

t

I

I

I

I

1

I

Z FeO 10

9149149 I

TiO2

I

2 i

MgO

9

I

9

el 9

~

d~

ere

ot

10 I

I

I

"

9e

e ~ b Jb

aen

-J

I

1

I

9

'~"

"~,-t'..

9

ol

AAA~

9

t e ~ o olll

9

9 9149149

9

QI

CaO 10 AI203 20

3.5.2.2

Plagiogranites

The plagiogranites are equigranular medium to fine grained mesotype to leucocratic rocks with 50-90% zoned subhedral plagioclase (An36_9) and 10-40% modal quartz. The textures range from hypidiomorphic granular, where quartz forms a granular mosaic partly enclosing earlier formed feldspars, to granophyric where a graphic intergrowth of albite (An9~3) and quartz predominates. The plagioclase is rarely fresh and usually has turbid altered andesine-oligoclase cores and partly altered rims and overgrowths of clearer albite. Up to 10% prismatic to fibrous actinolite occurs with fine grained magnetite inclusions and probably replaces original clinopyroxene. Also present are larger (1-2 mm), probably primary, opaques, mainly Ti-magnetite and ilmenite intergrowths partly replaced by sphene, small apatites, often as inclusions in quartz, and occasional zircons. Secondary hydrothermal minerals (epidote, sphene, prehnite and quartz are the most common) occur in interstitial cavities and as veins in a granular mosaic in the more completely altered parts of the rocks.

3.5.3

Geochemistry

The high-level intrusives, whose rock types range from gabbros to trondhjemites, have a wide compositional range with SiO 2 contents from 45% to 77% (Table 3.9, Fig. 3.34). The composition of the gabbros and ferrogabbros overlap with the field of "ocean floor gabbros" in the AFM plot (Fig. 3.35). The least fractionated olivine gabbros have the highest Mg* (100 Mg/(Mg + Fe)), Ni and Cr contents and overlap with the compositions of the underlying cumulate gabbros (Fig. 3.35).

Na20 5 I

I

1

1

K20 0.5

I

I

9 9

I

I

I

1 9149 9 9 000~

9 9149149 1

I

i

I

I

oo

9

P205 0.2 0.1

9

I

9

9 I

J

el

00

I

I

.I

Zr 300 (ppm) 200 9

100 1

10

9 ~MI 9

20

9

1

30

1

I

40 50 DI(Q +or +ab)

9

9

I

1

I

I

60

70

80

90

Fig. 3.34. Major oxides and Zr versus Differentiation Index (DI) for the high-level intrusive sequence. 9 -- high level gabbros, 9 = plagiogranites.

83

The Semail Ophiolite Table 3.9. Geochemical analyses of high-level intrusives.

Gabbros

Ferrogabbros

Ferrodiorite

Quartz diorites

0M1220 0M3074 0M1219 0M3082 0M1290 0M3038 0M984 Sf Fo Sf F F A R

R

47.34 0.40 15.01 8.71"

49.10 2.51 14.20 5.62 8.37 (/.16 5.27 9.16 3.65 0.12 (I. 1(I tl.39

53.39 1.59 13.44 12.50"

54.95 1.5(1 12.45 9.45*

(/.13 8.00 12.43 2.39 (I.(/8 (/.(/5 3.22

50.70 (/.85 15.4(I 6.47 7.31 0.12 5.63 1(I.23 1.64 0.(/4 (t.(/5 1.46

54.19 1.57 13.65 11.47*

P205 LOI

46.80 0.42 17.70 1.71 4.65 0.12 9.26 14.62 1.64 0.14 0.02 2.65

(I.14 5.89 7.47 4.30 (t.21 0.12 1.28

0.12 6.72 7.34 3.94 ti.16 (1.11 1.01

Total

1(t(I.61 99.30

99.9(t

99.56

10(I.29

100.32

11.12 (1.24 13.88 34.55 12.85 14.7(t

1.99 0.71 30.88 22.01 18.47 11.05

2.01 1.24 36.38 17.33 15.63 2(1.16

1.49 0.95 33.34 18.52 14.15 24.34

0.07 1.(t(I 39.(19 12.74 24.58 16.2(t

9.38 1.61 (1.12

8.15 4.77 (I.24

1.97 2.98 0.28

2.16 3.(/2 0.26

1.62 2.85 0.28

63

53

48

nd

19

19

nd

5

nd

15 10 14

210

148

nd 44 213

nd nd 21

34 22 92

29 7 67

33 21 95

76 71 44 nd 298

180 275 42 22 114

SiO~ TiO2 A1203 Fe203 FeO MnO MgO CaO Na20

48.47 0.44 15.34 8.51"

0.14 10.74 13.06 1.81 0.06 0.02 2.66

48.30 0.43 17.80 1.46 3.70 0.09 8.87 14.30 1.94 0.09 0.(15 2.91

99.74

99.95

99.93

99.83

Q or ab an di hy ol mt il ap

0.83 13.91 40.64 25.53 1.68 11.42 2.49 (1.80 0.05

0.36 15.31 32.66 25.80 5.40 14.69 1.51 0.76 0.05

0.53 16.41 39.60 24.67 6.14 6.62 2.12 0.82 0.12

1.18 17.01 32.25 23.98 7.17 13.60 1.46 0.84 0.09

Ba Co Cr Cu Ni Sc Sr V Y Zn Zr

37 503 57 121 42 193 152 9 27 19

K20

21

* total Fe as Fe203

330 128 103 10 50 14

0.14 1(t.63 12.55 2.01 0.2(I 0.(/4 1.50

49.30 (1.71 13.4(/ 3.44 6.85 0.15 10.07 12.16 2.22 0.12 0.06 2.12

(/.71 18.78 26.25 26.91 16.00 2.46 4.99 3.25 0.14

25 34 92 35 99 38 197 134 13 19 28

369 127 143 12 24 32

48.53 (/.76 14.55 9.16"

(/.47 20.22 28.74 26.5(t 10.17 6.03 1.58 1.44 (1.12

0M169

28 53 222 31 87 47 102 192 23 28 41

0M3029 0M3117 0M3018 0M3073 0M1225 J J F Fo Sf

0.16 7.(13 8.81 4.62 0.17 0.12 (/.63

56.50 0.59 19.4(/ 1.58 1.79 0.07 0.92 9.59 7.20 0.01 0.10 2.20

57.2(/ 1.27 15.30 3.96 5.66 0.16 3.31 5.59 4.76 0.38 (1.15 1.77

99.88

99.95

99.52

0.06 56.80 20.60 7.08 7.34 2.23 2.29 1.12 0.24

10.76 2.25 40.27 19.26 6.11 10.58 5.74 2.41 0.35

39 47 39

26 87 26 38

186 21 21 39

103 8(12 14 39 19

Key to location of samples: F = Fizh, Fo = Forest, A = Ahin, J = Jizi, Sf = Shafan, R = Ragmi.

84

Chapter 3

Table. 3.9. C o n t i n u e d Plagiogranites Tonalites

Aplite

Trondhjemites

OM3110 J

OM3107 J

OM168 R

OM3077 F

OM4944 A

OM3011 F

OM3024 J

OM3000 A

OM4722 J

OM3119 J

SiO~ TiO2 AI20 3 Fe203 FeO MnO MgO CaO Na~O K20 P205 LOI

62.14 0.89 14.80 7.15"

65.10 0.97 14.00 4.16 2.90 0.07 1.40 3.78 5.65 0.10 0.16 0.82

65.57 0.55 13.92 5.89*

71.81 0.40 11.63 3.50*

72.14 0.31 12.98 2.44*

0.10 1.21 3.1/I) 6.93 0.18 0.22 2.03

66.80 0.59 13.60 4.15 2.29 0.1/5 0.85 3.62 5.82 0.1/5 0.24 1.61

70.13 0.57 13.71 1.97"

0.06 2.28 5.82 4.73 0.25 0.15 1.58

64.70 0.68 14.20 1.65 4.13 0.09 1.56 3.51 5.84 0.41 0.23 2.02

//.1/4 0.81 3.69 7,76 0.01 11.12 1.111

11./14 1.68 4.33 4.41 0.117 t).119 1.31

/).113 1.33 2.35 6.65 0.112 I).tt4 1.1/0

72.50 I).45 13.211 1.98 1.26 0.04 0.61 2.11 5.51 0.22 0.07 1.011

Total

99.85

99.02

99.01

99.60

99.53

99.82

99.27

99.29

Q or ab an di hy mt il ap others

17.64 1.48 40.02 18.42 7.48 7.57 2.74 1,69 0.35

18.56 2.42 49.41 11.33 3.83 7.23 2.39 1.29 0.54

22.89 2.59 47.80 12.55 4.63 1.73 6.03 1.84 0.12

16.15 1.116 58.63 6.35 5.98 5.15 2.26 1.04 1/.52

26.29 0.30 48.1/6 11.47 3.83 0.34 5.83 1.12 I).57 0.13

18.92 I).116 65.66 2.55 6.4/I 2.96 t).75 1.118 0.28

35.02 (1.41 37.31 11.74 7.49 3.43 1.35 11.76 /).21

27.56 0.12 56.27 5.52 4.81 3.25 //.75 0.59 11.tl9

93

13 42

45

18

4

7

35

nd 11 86

2

4

99

56

51 nd 2411

63 nd 180

5 7 140

-

Ba Co Cr Cu Ni Sc Sr V Y Zn Zr

46 9 10 190 41 3 147

Rare earth elements La 0M1220 0.8 (2.44) 0M1219 1.2 (3.66) 0M1290 2.1 (6.40) 0M169 3.0 (9.15) 0M1225 5.2 (15.8) 0M4944 6.11 (18.6) 0M4722 7.89 (24.0)

nd nd 37

23 9 15

129 35 65 4 182

151 51 58 7 209

Ce 2.1 (2.43) 3.9 (4.51) 5.5 (6.36) 10.7 (12.41 13.5 (15.6) 17.7 (20.5) 22.2 (25.7)

Nd 2.1 (3.33) 4.3 (6.83) 5.8 (9.21) 11.2 (17.8) 12.7 (20.2) 15.5 (24.6) 19.6 131.1)

30 12 16 109

5 14 84 11 65 6 139

68 11 122

Sm 0.8 (3.94) 1.3 (6.40) 2.4 (11.8) 3.9 (19.2) 4.1 (20.2) 5.0 (24.6) 5.31 (26.2)

Eu 11.45 (5.84) 0.58 (7.53) 11.86 (11,2) 1.34 (17.4) 1.47 (19.1) 1.65 (21.4) 1.45 (18.8)

Gd 1.6 (5.80) 2.3 (8.33) 3.3 112.111 5.3 (19.2) 5.8 (21.111

Tb 11.25 (4.81) 11.34 (6.54) 0.55 (10.61 1.01 (19.4) 1.112 (19.61 1.26 (24.2) 1.27 (24.4)

Tm 0.16 (4.71) [0.28] 11.43 112.61 11.77 (22.6) 0.8 (23.5) [0.81] [0.94]

Yb 11.88 (4.1111) 1.25 (5.68) 2.11 (9.59) 3.86 (17.61 4.06 (18.5) 5.19 (23.6) 6.22 (28.3)

OM3071 Fo

OM3085 F

73.111 I),45 11.90 1.33 1.84 /)./13 0.87 1.94 5.35 1/.32 0.117 1.14

75.5/) 0.31 12.90 11.51 1.35 0./13 0.66 1.711 6.68 0.06 nd 0.92

77.20 0.08 13.27 1t.54"

98.95

98.34

1t1/).61

100.79

34.37 1.3/I 46.62 111.1/1

35,13 1.89 45.27 7.52 1.35 3.10 1.93 11.86 0.16

31.79 0.36 56.52 5./15 2.8tl 1.86 0.74 0.59

28.69 0.18 68.10

31

nd+ 42 nd nd nd 4 55 43 16 nd 328

20 7

1.53 2.87 0.86 11.16 /1.23 66 68 nd nd 9 9 85 14 77 10 246

Lu 11.13 (3.82) 0.19 (5.59) 0.32 (9.41) 11.61 (17.91 0.64 (18.8) 11.87 (25.6) 1.08 (31.8)

nd nd 56 111 28 100 nd 257

/).1/1 0.13 /).72 8.30 0.03 0.02 0.40

1.62

(1.15 0.115 1.1/9

nd 13 31 18 nd 97

Ta [0.14]

0.79

0.03

[0,17]

37.7

1.14

0.06

[0.15]

47.3

2.59

0.20

1t.18

38.5

0.42

25.0

3.117 0.51

U

Sc

Hf Th 11.51 0.02

4.2

11.66 0.72 0.37

6.53

1.29

i n d = not detectable * total Fe as Fe203 Key to location of samples: F = Fizh, Fo= Forest, A =- Ahin, J = Jizi, Sf = Shafan, R = Ragmi. ( ) indicates chondrite normalized values, [ ] analysis unreliable.

11.8 11.67

41.7

14.1 8.9

The Semail Ophiolite

85

Table 3.10. Results of petrological mixing calculations Fractionating phases Parent

Product

Plag

Cpx*

Gabbro (OM1219)

Ferragabbro (OM169)

Gabbro (OM1219)

Ferrodiorite

35.4 (Ansi) 35.5 (An,~,,) 36.0 (An~_,) 40.9 (An30

Ferrodiorite (OM1225) Ferrodiorite (OM1225)

(ON 1225) Tonalite (OM4944) Trondhjemite (OM4722)

Mt*

Liquids Il*

l

2

32.8

31.9

29--35+

35.4

26.7

25-31+

12.4

5.5

2.0

42.7

40-52-

13.8

6.9

2.1

34.4

34--42w

Liquid 1 = % melt remaining from major element least squares modelling. Liquid 2 = % melt remaining from REE modelling using Rayleigh fractionation equation * Mineral compositions Clinopyroxene: SiO2 52.8, TiO2 0.39. AI203 2.9. FeO 4.64. MnO 0.1, MgO 16.67, CaO 22.26, Na20 0.21. Magnetite and ilmenite taken as stoichiometric Fe304 and FeTiO3 Olivine and apatite were included but not used by the programme All values in percentages. + Using D(cc.Nd.Yb) = 0.1 and D iEu) = (}.2 $ U s i n g D(cc.yb ) = 0.7, D {Nd~ = 0.75 and D ~EL,I = 0.85 w U s i n g D {Ce.Nal = 0.5, D (Vh) = 0.6 and D I E , ) = 1.5

They have low incompatible element contents (e.g. Zr <30 ppm) and "humped-shaped" convex-upward REE patterns relatively enriched in the middle REE but lack positive Eu anomalies (Fig. 3.36). These features indicate a non-accumulative origin and, as the rocks do not show well-developed cumulate textures except igneous lamination, it seems likely that early-formed crystals were separated from the interstitial residual liquid by a process such as filter-pressing. The more differentiated (Mg* = 72.5-50) Fe and Ti-rich gabbros contain higher concentrations of incompatible elements (Zr 40-100 ppm) and have negligible Cr and Ni contents. They have similar shaped REE profiles to the mafic gabbros but with higher overall abundances and small negative Eu anomalies suggesting that they are residual liquids formed by fractional crystallization. As an example of the likely relationship between the mafic gabbros and ferrogabbros, the composition of the most Fe and Ti-rich ferrogabbros can be modelled as the 30-35% residual liquid formed by the removal of approximately equal proportions of clinopyroxene and plagioclase (Ans0) from a gabbroic parent melt (Table 3.10). Such a tholeiite differentiation trend towards iron enrichment is generally considered to represent fractionation under anhydrous and reducing (low fO2) conditions. The plagiogranites, despite severe secondary alteration, clearly have a wide range of primary compositions from quartz diorites (<60% SiO2) through tonalites to trondhjemites with 67-77% SiO 2. In general, the trondhjemites have lower MgO (<1.5%) and higher differentiation indices (DI = total Q + a b + o r ) (>70) than the tonalites. Most of the plagiogranites show good correlations between major and trace elements and differentiation index (Fig. 3.34). P2Os attains its highest values (up to 0.25%) in the tonalites and then decreases in the trondhjemites, suggesting late-stage fractionation of apatite. There is little petrographic evidence for secondary enrichment in silica in rocks with SiO2 contents as high as 75% which appear to have more-or-less primary magmatic compositions. Fractional crystallization of plagioclase (An3o), clinopyroxene and Fe-Ti oxides from a ferrodiorite can

produce tonalite and trondhjemite compositions as c. 35% residual liquids (Table 3.10). These results are supported by the higher incompatible element contents (Zr 122-328 ppm) of the plagiogranites and their REE patterns which show increasing negative Eu anomalies in progressively more differentiated compositions (Fig. 3.36). Secondary alteration of plagiogranites, mainly patchy replacement by epidote-prehnite-quartz and albitization of the feldspars causes a loss of K and an increase in Na and some redistribution of other elements. Na20/K20 ratios range from 14 in the least altered to over 700 in the most severely altered rocks. Fig. 3.35. An A (Na20+K20)F(FeO+

FeaO3x0.9)M(MgO) plot for the highlevel intrusives. F

r~2~-') Cumulate gabbros and wehrlites 9 High-level gabbros FG Ferrogabbro FD Ferrodiorite Plagiogranites

A

9 FG

FD &A

86

Chapter 3

30--

~

OM 4722 ~

~

20 '-'OM 4944 ~ OM 1225 ~

#

~

-

-

~

-

~

-

~ / t ~ ~ ' ~ /

4

~

~

OM 4944 T,:mal,t~ OM 1225. Forrod,o*de OM169 F. . . . . . . . . . ht.....

~

.~

~

0 r

OM 4722 Trondhjem,b,

1

OM 1290 F~rroqabbr,)

OM 1290 5 --

Gabbros ~'qk~'-o

M 1219

OM 1220 1

OM 1220-~

I

1

La Ce

I

Nd

I

1

I

I

Sm Eu Gd Tb

I

I

I

Tm Yb Lu

Fig. 3.36. REE/chondrite plots for the High-level lntrusives.

3.5.4 Petrogenesis The transition from mineralogically unaltered to altered gabbros approximates to the cumulate/high-level intrusive boundary and there is abundant evidence that both late-stage magmatic and hydrothermal fluids played an important role in the alteration of the upper part of the plutonic complex. The more differentiated rocks show more intense alteration suggesting that the bulk of the alteration occurred by magmatic fluids concentrated in the residual melts. However, 61sO isotope studies of whole rocks and minerals from the high-level intrusives (Gregory & Taylor 1981; Stakes et al. 1984) clearly show that seawater penetrated the roof zone of the magma chamber, either directly along fractures by hydrothermal circulation or, more probably, by the incorporation and dehydration of xenoliths of hydrothermally altered dykes. Stakes et al. (o19 cit.) obtained 6~SO values of +7.4 on quartzes in "axis sequence" plagiogranites from Wadi Ragmi and Wadi Jizi (Musafiyah) that indicate post-magmatic re-equilibration with 1SO-enriched hydrothermal fluids at 350-400~ They also found that the host plagiogranite and a metadolerite xenolith from the same body have whole rock 61~Os of 8.42 and 5.49 respectively and are thus seen to be out of isotopic equilibrium. It is not clear from the evidence available whether the highlevel intrusives represent a single fractionation series, as suggested by the continuous trends on variation diagrams and by the crystallization-residual liquid modelling described in the preceding section, or two separate trends with one branch leading to Fe-rich gabbros and the other to a plagiogranite end member. Spulber & Rutherford (1983) have shown experimentally that, under hydrothermal conditions (1-3 kb pressure, fH20 - 0.6 P~uid and fH2 < 0.33 Pnuid), Fe-Ti oxides crystallize closer to the basalt solidus than under anhydrous conditions. The timing of oxide precipitation is critically dependent on the fO2; under buffered relatively low fO2 conditions, they crystallize after about 50% fractionation of clinopyroxene and plagioclase which produces an overall trend that begins with Fe-enrichment and then moves rapidly to Sienrichment at the onset of oxide precipitation. At higher fO~

conditions, the silica and alkali-rich granitic melts are produced earlier with no intermediate stage of iron-enrichment. Thus, the various differentiated rocks in the high-level intrusives could have formed in isolated magma pockets some of which initially fractionated under relatively anhydrous conditions to form ferrogabbros, whereas others fractionated in more hydrous and oxidizing conditions to form plagiogranites, the course of fractionation being controlled by the time at which the fluids interacted with the hot crystal mush/residual liquid system. The most popular model for the formation of ophiolitic plagiogranites is as c. 10% residual liquids which formed by low pressure crystal fractionation under hydrous conditions from a subalkaline low-K tholeiitic magma (Coleman & Peterman 1975; Aldiss 1978, 1981). We have demonstrated that this mechanism appears to fit the field relations and geochemistry of the Semail high-level intrusives. Other models for their origin; e.g. by the local melting of hydrous gabbro or amphibolite (Pedersen & Malpas 1984) or by liquid immiscibility (Dixon & Rutherford 1979) are not supported by the data. Some authors (Bischoff & Dickson 1979; Mottl & Holland 1978) have suggested that the characteristic potassium (and Rb, Ba) poor nature of plagiogranites results from post-magmatic interaction with heated seawater. This does not appear to be true for the Semail plagiogranites because, apart from the absence of pseudomorphs after K-rich phases, the oligoclases have higher K20 contents (Oro.7_3.2) than the andesine cores (Oro ~0 5) showing that potassium is concentrated during the later stages of fractionation and retained in the last-formed feldspars. The plagiogranites contain up to 0.4% K20 in the least altered samples, although K is clearly lost in the more altered rocks. Thus, the low-K character of the plagiogranites appears to be an original feature but one which is enhanced by secondary alteration.

3.6 Late Intrusive Complexes Widely distributed along the length of the ophiolite outcrop are cross-cutting plutonic complexes (Fig. 3.37) composed of peridotites, gabbros, diorites and plagiogranites, that are similar petrologically to the axis sequence high-level intrusives but which intrude up into the upper crustal units of the ophiolite (Sinewing 1980a; Browning & Sinewing 1981; Browning 1982). Browning & Smewing (op. cit. ) divided the late intrusives into two groups: an older series of large (1-6 km 2) differentiated gabbro to plagiogranite plutons and a younger group of smaller (<1 km diameter) peridotite-gabbro intrusions. The first type are mostly equidimensional or only slightly elongated in plan whereas the second type are commonly elongated, and can consist of a string of small bodies lying along faults that usually trend between 110~ and 135~ A feature of many of the late intrusive complexes is that they are composed of multiple intrusions often of strongly contrasting compositions. Xenolithic contact zones, involving earlier and later members of the same complex or between the plutonic rocks and the country rocks in the roof and wall zones of the intrusion, are common. Many of the rocks show evidence of hydrous crystallization with primary amphiboles present in the peridotites and gabbros as well as in the associated diorites. In general, however, the gabbros, diorites and plagiogranites of the late intrusives are petrographically indistinguishable from those of the axis sequence high-level intrusives.

87

The Semail Ophiolite 1 WADI FAYD

CS

D

"AXIS" OPHIOLITE SEQUENCE

E

~ ~

"~

D

E Extrusives D Sheeted dykes G High level intrusives CS Cumulate sequence

E ~ . . ~ ~, ::::::::::::::::::::: {i.I :.).: :.:.' :: : :. :.1, : ?:i:..::!:. : : :~

LATE INTRUSIVE COMPLEXES

Cg

:

~

/

Gabbro-Diorite

~1 /.

2 WADI RAGMI

onalite-Trond hjemite

/

~

/

Peridotite

~\ CS

'~'~k

CS

5 LASAIL

G

!

E

\

D

E

O l

Dykes Maior faults (tick on downthrown side) 5km J

u

3

ZABIN

8 BUDIT

4

9 WADIYAH

f rlLrl ~L.UL~ ~ L ~,~ 4~t

S~AMpIX'~, (

~

Mantle sequence

D

6 LEMARER

9 ::\"~.,'~N~.~'~-~

.

~

Sheeteddky es

Cumulate sequence ~

,,~ . . . . . . . . to Line m~rus~ve u o r n p ~ e x e s

7 JEBEL FAYYAD

~ 7 D~,"t'-............~ Gh~\~,uzayn 4,.h~,.

:

t tonalite g gabbro

Mashin

-

D

~

'/"~---~-~-~,. ~ E

D ~

~

X ~ 9 ~ ~ , ~

~i~?i~ ~~

Laves

~ ~ : 5

D X

trondhjemite~ p peridotite '

~:~ ~

i ~

3.6.1 Large gabbro-diorite-plagiogranite bodies These are composed mostly of layered and massive gabbros and diorites with subordinate volumes of plagiogranite and have outcrop areas from 1 km 2 to over 10 km 2 (Table 3.11). The Lasail complex, located just to the south of Wadi Jizi, is a typical example and consists of a core of gabbros transitional into diorites overlain by an irregular sheet of tonalite which crops out around the margins of the complex (Fig. 3.38, Plate 3.22). It is emplaced into the sheeted dykes and the base of the laves and along its southern margin is faulted against layered gabbros of the main axis sequence. The intrusion has a domal structure with the central gabbros dipping radially away from the centre of the intrusion. The gabbros are at least 150 m thick but the base is nowhere exposed. The overlying tonalite

Fig. 3.37.

Distribution of the Late Intru-

sive C o m p l e x e s

complexes

and maps of ten individual

sheet is up to 80 m thick and in places highly xenolithic. Numerous acid sills, both aplitic and pegmatitic, up to 2 m thick cut the gabbros which are in places net-veined by fine grained felsite. The gabbros are fine to medium grained (0.2-2 mm) rocks with fresh to slightly altered zoned plagioclase (An86-55) and interstitial clinopyroxene (Wo47.sEn41Fs11.5) that is partly replaced by actinolite. The subhedral feldspars show an igneous lamination and there is sometimes an indistinct modal layering of the proportions of feldspar and pyroxene. The rocks contain between 1% and 10% modal Fe-Ti oxides. The laminated gabbros grade upwards into vari-textured diorites with a slightly coarser (average 2 mm) grain size and composed of more sodic zoned plagioclase (An75_40), secondary actinolite and opaques.

Chapter 3

88

Table 3.11. Gabbro-Plagiogranite Late Intrusive Complexes. 1 Wadi Fayd Wadi Ragmi Zabin (1)

61/2 80%G 20%P 61/: 100%G 12 100%G

(2)

].1//2

Jebel Shaykh LasailWadi Barghah Lemarer (Wadi Sarami) Jebel Fayyad

25? 10

100C~G 40%G 60%P

4V2 55%G 45%P 13

Maydan Budit Wadiyah 1. 2. 3. 4. 5.

2

50%G 50%P

4 65%G 35%P 4 40%G 60%P 21/2 20%G 80%P

3

4

5

HLG-D Cg D-E Cg-HLG HLG-D-E D-E

WNW-ESE E-W to NW-SE N-S E-W to NW-SE E-W E-W to WNW-ESE

D-E

(NW-SE)

D

(NW-SE)

(NW-SE) E-W

WNW-ESE to NNW-SSE NW-SE E-W to WNW-ESE NW-SE NW-SE to N-S WNW-ESE WNW-ESE to NW-SE

(Cg-HLG)-D Cg-HLG-D Cg-HLG-D

Area km 2. % amounts of gabbro-diorite (G) and plagiogranite (P). Host to intrusion: Cg cumulate gabbros, HLG high-level gabbro, D sheeted dykes, E extrusives. General trend of complex (in brackets if weakly defined). Trends of associated late dykes.

Lasail ."""..-"" \ ~6~.. .............

:X~X\

v

\

v v

V V

v

V

V

v \

si:84i

5"

.-'M:'.

V

9176

:(i:. a01:

J

"-'/

,,

" " : /J

40 t

. . . . . . . ::::::::::::::

0

1 km

I

I

~:32--"

.......

Inclined sheets in lavas A

.;. . . .

root of intrusion? / - -

.

.

.

.

;.~

Wadi Barghah Fault

~

B

[ V ] Volcanics (mainlyGeotimes Unit) ~

Sheeted dyke complex ._a_Dip/strike of dykes

~

Cumulate gabbros ~

Dip/strike of layering

Late Intrusive Complex ~ Tonalite ~\,,~ Xenoliths Dip/strike of lamination ~Gabbro/diorite ....1_ or tonatite sheets _L

Dip/strike of andesite inclined sheets

~-

Felsite dykes

~

Late peridotite

Fault .......... Wadi

Fig. 3.38. Map and idealized cross-section of the Lasail late intrusive complex9

The Semail Ophiolite

89

Plate 3.22. Edge of the Lasail late intrusive complex. Gabbros and plagiogranites at left are intruded into the lavas which make up the more pointed ranges beyond.

The tonalite sheet is a massive, columnar-jointed intrusion that has a fine to medium grain size with, for the most part, a hypidiomorphic granular texture. Most of the rock shows severe effects of hydrothermal alteration with complete replacement of the primary feldspars and abundant late-stage epidote and prehnite in veins and cavities. It contains abundant dark xenoliths up to 5 m long which can form up to 50% of the rock volume (Plate 3.23). The largest xenoliths are tabular but smaller ones tend to be equidimensional with rounded shapes indicating assimilation. Where it is highly xenolithic, the tonalite is darker than normal and clearly some contamination of the acid magma has taken place. The xeno-

liths consist largely of fine grained actinolite, epidote, albite and quartz. Fine grained acid dykes and sheets cut the country rocks surrounding the Lasail complex and a set of andesitic cone sheets is centred on the intrusion (Fig. 3.53, Section 3.8). The large, dominantly gabbro-plagiogranite late intrusives are typified by the Lasail complex (other examples are Jebel Fayyad in Wadi Sarami, and several bodies between Wadi Fizh and Wadi Hatta; Jebel Shaykh, Zabin and Fayd). They probably represent the roof zones of moderately large magma chambers which fed the Lasail Unit extrusives (Section 3.8). The lower parts of the chambers are not exposed but they are up to 6 km across and possibly 2-3 km deep.

3.23. Lasail late intrusive complex. Plagiogranite with super-abundant metadolerite xenoliths.

Plate

90

Chapter 3

3.6.2 Peridotite-gabbro complexes The second group of late intrusive complexes includes the small (1.5 km long, up to 200 m thick) NW-SE trending Mashin intrusion (Fig. 3.37) in the northwest corner of the Haylayn block (Browning 1982). This body consists of a complete differentiation series from wehrlite (50 m thick) at the base, overlain by layered and non-layered gabbros and diorites (140 m) and plagiogranite (35 m) at the top. The poikilitic wehrlite has a faulted contact with the underlying cumulate rocks so that the original floor of the intrusion is not seen. The maximum 200 m thickness of plutonic rocks is only seen at the northwest end of the intrusion. The sequence is much thinner at the southeast end where the felsic rocks are absent and it consists of about 30 m of wehrlites overlain by 10 m of layered gabbronorites. The Mashin poikilitic wehrlite consists of up to 80% rounded, 1-5 mm sized olivines (F085-88) poikilitically enclosed in intercumulus clinopyroxenes (Wo44EnsoFs6) and orthopyroxenes (Wo3En86Fs11) up to 10 mm across. Much of the rock is composed of secondary amphiboles with brown pargasitic hornblende rimming and replacing clinopyroxene and both the pyroxenes and hornblende being replaced by colourless fibrous tremolitic hornblende. In the upper part of the wehrlite, plagioclase (An83-85) is an intercumulus phase and forms schlieren of feldspathic wehrlite that are elongated parallel to the gross lithological banding, of the intrusion. A two-metre thick layered olivine gabbro overlies the wehrlite and has an adcumulate texture with small (1 mm) olivines (Fo84) and laminated cumulus plagioclase (An75) with intercumulus clinopyroxene (Woa6En46Fss) and brown hornblende. This passes up into 90 m of coarse grained gabbros gradational into diorites. These have variable grain sizes and are composed of euhedral to subhedral, normally-zoned plagioclase (Anss_20) with interstitial clinopyroxene (mean Wo44En40Fs16) and hornblende both of which are extensively altered to actinolite. Fe-Ti oxides are present in only small amounts (up to 5%) and no ferrogabbros are present. The laminated gabbro-diorite unit is succeeded up section by 30 m of isotropic diorites and at the top of the sequence is a 35 m thick sheet of plagiogranite. This is medium to fine grained with a hypidiomorphic granular to granophyric texture and is heavily altered. The plagiogranite is intruded into the Sheeted Dyke Complex and contains small xenoliths of metadolerite. In addition to the main intrusion, fine to coarse grained acid

Plate 3.24. Dark peridotite sill cutting Sheeted Dyke/Geotimes Unit contact zone near Ghuzayn.

sheets cut all levels of the intrusion and extend into the surrounding rocks. Most other members of this group consist of small (<1 km 2 in area) intrusions composed largely of peridotite, mainly coarse-grained wehrlite, usually with subordinate amounts of gabbros and more differentiated rocks in marginal zones. The majority of these late peridotites are elongated dyke-like intrusions up to 1.5 km long and a few hundreds of metres wide. Most have near-vertical contacts but some were emplaced as flat-lying sills, particularly where they occur in the extrusive sequence (Plate 3~24). They are commonly aligned along NWto WNW-trending faults. They cut the Sheeted Dyke Complex and nearly all the lava sequence, up to and including the Alley Unit. Some cut the gabbro-plagiogranite bodies described in the preceding section. Wehrlites, grading into lherzolites, are the most common rock type and are easily recognized by their massive, dark brown, speckled appearance. I n the cores of spheroidally weathered masses they can be extremely hard, fresh rocks composed of up to 80% subhedral to anhedral olivine (Fo85), 1.5-5 mm across, with interstitial diopsidic clinopyroxene (Wo41En52Fs7) and variable amounts of orthopyroxene (WoaEn85Fs13) containing minute chrome spinel euhedra. Dark brown to colourless pargasitic to tremolitic hornblende commonly occurs interstitially and partly replaces the clinopyroxene. Poikilitic textures are by far the most common, but some examples have an intergranular texture with subhedral pyroxenes more or less the same size as the olivines. Plagioclase wehrlites contain either cumulus or intercumulus plagioclase (An92-80) and with decreasing olivine content there is also a gradation into olivine pyroxenites (both clino- and orthotypes). The associated gabbro~ which usually form a narrow zone a few metres wide at the margins of the body, are medium grained (1-2.5 mm) hypidiomorphic to intergranular textured with zoned plagioclases (An90-65) and clinopyroxenes (Wo42EnsoFs8 to Wo43.5En45FSil.5), partly altered to green actinolite, and with accessory Fe-Ti oxides. Small differentiated intrusions, like the Mashin complex, appear to represent high-level magma chambers in which latestage magmas collected and tractionated and which fed the Alley Unit lavas (Section 3.8.1). The small, dominantly ultramarie bodies, some of which are intruded high up into the lava sequence are feeders for the Cpx-phyric unit and may be crystal mush intrusions produced by the mobilization of cumulates from the bases and walls of these chambers.

The Semail Ophiolite 3.6.3 Geochemistry and petrogenesis W h o l e rock analyses of r e p r e s e n t a t i v e peridotites, gabbros, diorites and plagiogranites, including both tonalites and t r o n d h j e m i t e s , f r o m the late intrusive c o m p l e x e s are p r e s e n t e d on Table 3.12. As already n o t e d , they are similar p e t r o g r a p h i cally and g e o c h e m i c a l l y (in m a j o r e l e m e n t s at least) to m a n y of the rocks f o u n d in the axis s e q u e n c e cumulates and high-level intrusives (Fig. 3.39). A n A F M plot (Fig. 3.40) bears this out and shows a b r o a d o v e r l a p of the field occupied by the two suites, the m a i n e x c e p t i o n appears to be the f e r r o g a b b r o s and f e r r o d i o r i t e s which are only f o u n d in the axis s e q u e n c e . As a result, the late intrusive c o m p l e x e s show a flatter, m o r e "calca l k a l i n e " t r e n d on this diagram with a greater a b u n d a n c e of i r o n - p o o r i n t e r m e d i a t e rocks of quartz diorite c o m p o s i t i o n leading directly to tonalites and t r o n d h j e m i t e s . This t r e n d implies crystallization u n d e r m o r e h y d r o u s conditions of higher f H 2 0 and fO2 which appears to be b o r n e out by the g r e a t e r a b u n d a n c e of h y d r o u s minerals in the late intrusive suite. T h e v o l u m e of acid differentiates is g r e a t e r in the late intrusives c o m p a r e d to the high-level intrusives and, in par-

91

ticular, the large ( > 1 k m 3) xenolithic tonalite bodies s e e m to be typical of this group. In general, the late plagiogranites have a similar range of m a j o r e l e m e n t c o m p o s i t i o n s to the axis plagiogranites, but differ significantly in their trace e l e m e n t c o n t e n t s ; in particular, they are m a r k e d l y m o r e d e p l e t e d in i n c o m p a t i b l e HFS e l e m e n t s (e.g. Zr c o n t e n t s of the "late" plagiogranites are 30-125 p p m c o m p a r e d to 120-430 p p m for the "axis" o n e s (Fig. 3.41)) and they are generally m o r e d e p l e t e d in light r a r e - e a r t h e l e m e n t s giving sloping rather than the flat c h o n d r i t e - n o r m a l i z e d patterns of the axis s e q u e n c e plagiogranites (Fig. 3.42), although they show a wide range of R E E c o n t e n t s f r o m < 5 x to - 2 0 x c h o n d r i t e . T h e fractionation trends, as exemplified by plots of Ti vs Zr, for the w h o l e rocks from the Lasail and Mashin complexes, are s h o w n on Fig. 3.43, w h e r e they are c o m p a r e d with the t r e n d of the axis l a y e r e d s e q u e n c e and high-level intrusives. The late intrusive c o m p l e x trends, s h o w i n g l o w e r Zr values (mainly less than 100 p p m Zr) in the f r a c t i o n a t e d rocks, follow the same trends as the Lasail, Alley and Cpx-phyric Unit extrusives (Section 3.8) thus s u p p o r t i n g the field e v i d e n c e that the late intrusives p r o b a b l y r e p r e s e n t the crystallization products of high-level m a g m a c h a m b e r s from which these later volcanic

Table 3.12. Geochemical analyses of Late Intrusive complexes. Peridotites 0M1331 Z

Gabbros 0M1437 0M7119 0M7069 0M7037 0M1285 0M1313 0M7070 0M1283 0M1332 0M7129 0M7079 0M7112 0M7127 Sf Hw Hw J Z K Hw Z Z Rq L B Hw

SiO, TiO~ A1203 Fe202 FeO MnO MgO CaO Na~O K~O P205 LOI

4 0 . 6 7 41.57 0.09 (I.20 3.68 5.34 10.27" 9.28* 0.17 32.58 4.12 0.76 0.01 0.01 8.51

0.15 31.33 4.41 1.00 0.06 0.02 5.97

42.40 0.20 5.90 2.86 7.41 0.18 31.18 4.85 0.33 0.03 0.00 4.73

Total

100.87

99.33

100.07 98.19

0.06 6.96 7.15 11.81 4.67 66.66 1.47 0.18 0.03

0.38 9.06 10.61 10.09 3.84 63.37 1.31 0.41 0.02

0.18 2.79 14.53 7,53 17.93 47.86 4.15 0.38

3023

2641

1187 9

1274 37

2617 56 941 26 111 8 63 15

Q or ab an di hy ol mt ii ap others Cr " Cu Ni Sr V Y Zn Zr

3 1

14

45.7(I 0.29 5.20 1.25 3.91 0.13 21.0(I 16.63 0.10 0.00 0.01 3.97

0.69 13.74 54.10 23.22 1.81 0.55 0.02 ne 0.08 2345 253 929 8 162 11 24 15

47.30 0.17 6.3(I 2.41 6.47 0.17 25.00 7.62 0.54 0.03 0.00 2.91

47.90 (I. 11 3.68 1.53 5.27 0.15 24.73 13.04 0. I0 0.01 0.08 3.77

46.70 0.28 10.50 1.45 5.79 0.13 18.26 11.46 0.78 0.05 0.04 4.33

48.8(I 0.33 10.70 0.79 4.47 0.14 14.58 14.63 1.41 0.11 0.00 3.59

51.00 0.31 13.80 1.49 6.34 0.13 12.77 12.64 0.7(I 0.02 0.02 1.38

51.21 0.28 13.80 8.32*

98.92

100.37 9 9 . 7 8

99.55

0.18 4.57 14.68 18.29 32.30 22.28 3.49 0.32

0.06 0.85 9.56 43.10 17.51 22.91 2.22 0.21 0.19

0.30 6.60 25.00 25.13 17.29 19.09 1.36 0.53 0.09

0.65 11.93 22.55 39.70 5.94 13.42 1.14 0.63

2061 63 741 21 168 7 51 11

2050 28 592 10 131 5 42 6

1293 112 445 72 137 8 38 14

1150 6 187 140 180 11 11 9

Key to locations: B = Bani Umar, Hw = Hawasina, J

=

0.14 10.80 11.39 1.14 0.02 0.01 2.20

50.00 0.82 15.90 2.47 5.66 0.22 7.38 10.20 3.08 0.03 0.06 3.76

47.30 (I.73 18.20 5.80 5.5(I 0.13 7.51 10.07 1.90 0.29 0.01 2.66

50.90 0.29 16.40 2.50 4.43 0.13 8.76 12.12 1.29 0.10 0.03 1.99

52.20 0.57 15.30 4.93 5.27 0.16 5.21 7.00 4.07 0.38 0.07 3.36

100.60

99.31

99.58

100.10 98.93

98.52

1.71 (I.12 5.92 34.46 22.57 32.27

4.00 0.12 9.93 33.45 19.90 30.12

4.17 0.59 10.91 38.67 17.17 23.38

0.69 2.25 34.44 23.66 9.88 22.08

1.48 0.59 0.05

1.13 0.55 0.02

1.27 0.55 0.07

1.86 1.08 0.17

578 48 202 49 208 11 69 13

675

108 36 43 77 175 9 29 15

18 74 20 169 312 22 70 38

230 49 7 11

0.18 26.06 29.48 16.88 16.60 3.29 1.51 1.56 0.14 188 12 80 176 114 13 11 27

1.71 16.07 40.28 7.90 20.82 6.74 2.04 1.39 0.02

45 147 661 8 10 10

Jizi, K = Khabiyat, L = Lasail, Rq = Rustaq, Sf = Shafan, Z = Zabin.

Chapter 3

92 Table 3.12 continued Quartz diorites

Diorites 0M7128 Rq

Tonalites

0M7152 J

0M4948 Sr

0M8593 L

0M7086 H

61.66 0.88 15.90 6.73*

Trondhjemites 0M4954 Sr

0M7153 J

0.09 2.91 7.06 3.26 0.57 0.06 1.38

62.40 0.80 13.80 3.51 5.38 0.14 2.29 3.85 5.63 0.45 0.03 1.97

64.70 0.63 13.70 3.40 3.27 0.08 1.52 4.56 4.06 0.31 (I.06 2.70

66.30 0.63 12.40 4.12 2.41 0.(14 1.20 4.88 2.31 1.91 0.12 2.88

0M7114 B

0M4729 S

0M7015 Sr

0M7087 H

0M7035 J

68.10 0.58 13.50 2.54 1.91 0.07 1.86 5.29 3.61 0.20 0.04 1.83

69.50 0.63 12.30 4.50 1.70 (I,02 1.11 3.14 3.61 0.27 0.12 2.46

70.10 0.53 12.00 2.76 1.90 0.13 0.80 3.06 4.41 0.28 0.11 2.55

73.80 0.31 11.90 2.91 1.40 0.04 0.52 2.10 5.57 0.23 (/.04 0.97

76.70 0.20 11.50 0.81 1.28 0.04 0.19 2.00 4.81 0.20 0.05 1.10

SiOz TiO2 A1203 FezO, FeO MnO MgO CaO Na20 K20 P205 LOI

58.00 1.12 17.50 1.45 2.01 0.04 3.14 10.88 4.00 0.20 0.09 2.08

57.80 1.09 13.90 6.65 3.61 0.13 3.19 5.81 3.13 0.88 0.10 2.56

59.20 0.40 14.50 2.25 3.79 0.26 3.91 7.24 3.86 0.18 0.05 3.15

Total

100.51

98.85

98.69

100.28

100.25

98.99

99.19

100.28

99.36

98.63

_99.80

98.88

Q or ab an di hy ol mt il ap others Cr Cu Ni Sr V Y Zn Zr

10.37 1,18 33.84 29.21 19.34 0.95

14.92 5.20 26.48 21.28 5.79 17.92

15.61 1.06 32.66 21.71 11.24 9.21

20.46 3.37 27.58 24.86 8.05 9.57

15.35 2.66 47.63 11 .(16 6.50 8.40

27.81 1.83 34.35 18.25 3.24 4.56

34.86 11.23 19.55 17.86 4.34 1.02

29.72 1.18 36.89 16.68 7.33 1.76

39.42 1.60 30.54 14.79

36.18 1.66 37.31 12.13 1.87 1.68

36.09 1.36 47.13 6.80 2.62 0.08

42.97 1.18 40.70 9.20 0.34 1.73

1.31 2.13 0.21

1.84 2.07 0.24

3.26 0.76 0.12

2.07 1.67 0.14

5.10 1.52 0.07

4.93 1.20 0.14

5.97 1.20 0.28

3.68 1.10 0.09

4.00 1.00 0.26

117

43

28

3.74 0.59 0.09 hm 0.33 57

1.17 0.38 0.12

54

3.72 1.20 0.28 hm 1.94 62

195 154 25

238 123 40

112 121 26

263 60 34

393

523 46 62

200 44

134 13 30

124

37

185 35 28

40

49

31

43

83

57

116

99

60

93

22 8 10 420 221 33 68

Rare earth elements La 0M7079 0M7112 0M4948 0M7153 0M7114 0M7015 0M7087

0.1 (0.3) 0.3 (0.91) 1.3 (3.96) 3.28 (10.0) 1.0 (3.05) 4.0 (12.3) 1.56 (4.76)

Ba 44 11 308 275 30 36 68

2.76

56

35

Ce

Nd

Sm

Eu

Gd

Tb

Tm

Yb

Lu

Hf

Th

Ta

Sc

1.0 (1.16) 1.3 (1.50) 4.25 (4.91) 9.56 (11.0) 2.9 (3.39) 10.7 (12.4) 4.91 (5.68)

[1.9]

0.7 (3.45) 0.6 (2.96) 1.49 (7.34) 3.05 (15.0) 1.7 (8.37) 3.7 (18.2) 2.12 (10.41

0.27 (2.51) 0.22 (2.86) 0.70 (9.09) 1.08 (14.0) 0.56 (7.27) 1.11 (14.4) 0.72 (9.35)

0.81 (2.93) 0.65 (2.311

0.18 (3.46) 0.17 (3.27) 0.46 (8.52) 0.85 (16.3) 0.60 (11.5) 0.97 (18.6) 0.66 (12.71

0.11 (3.24) 0.12 (3.53) 0.35 (10.3) 0.64 (18.9) 0.43 (12.61 0.68 (20.0) 0.29 (8.53)

0.87 (3.95) 0.74 (3.36) 2.24 (10.2) 4.02 (18.3) 2.96 (13.41 4.54 (20.6) 3.16 (14.4)

0.13 (3.82) 0.12 (3.53) 0.40 (11.8) 0.72 (21.2) 0.51 (15.0) 0.72 (21.2) 0.57 (16.8)

0.29

[0.10]

[0.01[

44.3

0.42

[0.06]

[0.02]

38.7

0.54

26.5

[1.0] 4.0 (7.30) 8.58 (13.6) 3.5 (5.56) 10.3 (16.3) 5.28 (8.78)

2.6 (9.42) 5.0 (18.1) 2.24 (8.12)

1.75 2.75

0.34

0.13

19.0

1.89

0.23

0.06

21.8

3.03

0.32

0.13

12.4

2,31

0.28

0.09

15.1

Key to locations of samples: B = Bani Umar, H = Hatta, Hw = Hawasina, J = Jizi, K = Khabiyat, L = Lasaii, Rq = Rustaq, S = Salahi, Sf = Shafan, Sr = Sarami, Z =- Zabin. ( ) indicates chondrite normalized values, [ ] analysis unreliable. s e q u e n c e s w e r e e r u p t e d . It is n o t a b l e t h a t t h e M a s h i n i n t r u s i o n s h o w s a r e l a t i v e l y o r d e r l y t r e n d with i n c r e a s i n g i n c o m p a t i b l e e l e m e n t c o n t e n t s in t h e s e q u e n c e p e r i d o t i t e - w e h r l i t e - g a b b r o d i o r i t e - p l a g i o g r a n i t e , w h e r e a s t h e Lasail c o m p l e x s h o w s a m u c h less r e g u l a r t r e n d with t h e v o l u m e t r i c a l l y i m p o r t a n t i n t e r m e d i a t e to acid rocks ( d i o r i t e s a n d t o n a l i t e s ) s h o w i n g a w i d e s p r e a d o f c o m p o s i t i o n s . It is m o s t likely t h a t this c o m plexity is t h e result of t h e c o n s i d e r a b l e c o n t a m i n a t i o n s u f f e r e d by t h e acid r o c k s as a result of t h e i n c o r p o r a t i o n o f a n d r e a c t i o n with basic x e n o l i t h s . M a n y o f t h e late t o n a l i t e s a n d

d i o r i t e s w e r e p r o b a b l y p r o d u c e d by c o n t a m i n a t i o n o f t h e acid m a g m a s r a t h e r t h a n as n o r m a l f r a c t i o n a t e s of a g a b b r o p a r e n t . T h e s e c o n t a m i n a t e d rocks are c h a r a c t e r i z e d by low i n c o m p a t ible e l e m e n t c o n t e n t s n o t m u c h d i f f e r e n t f r o m t h o s e o f t h e associated gabbros and the metabasic xenoliths. Fractionation t h r o u g h i r o n - p o o r q u a r t z d i o r i t e s to t o n a l i t e s giving rise to t h e c a l c - a l k a l i n e t r e n d s m a y i n d e e d be t h e d i r e c t result o f t h e i n t r o d u c t i o n of w a t e r i n t o t h e m a g m a s by t h e d e h y d r a t i o n a n d partial a s s i m i l a t i o n of h y d r o t h e r m a l l y a l t e r e d x e n o l i t h s of d y k e s a n d lavas.

The Semail Ophiolite

93

ppm

TRONDHJEMITES 9 a, _ 70SiO2

60 -

GABBROS ~-ck~, c"@Z~-':c _ _ ' ~ - .

50-

DIORITES ~ ~o ..d o . . . .

~

TONALITES

o

ZFeO 1 0 -

o / /

TiO2

l

~

<._r~-_ 9

) c

F

1 --

MgO

10 -

CaO

10

AI2 03 20 -

Na20

5 -

K20

0.5 --

~w~O--

4~

- -- -~0- -- @ ~ C

O

0.2-P2 05

\

0.1. O~----r~J~.Q ~~ ~ ~

9

O0 9 9

300Zr

X Peridotite A Wehrlite $ Gabbr0 o Diorite 9 Quartz diorite [3 Tonalite .

~

~ - -

200 -

// /

100I~-~ ;; .7~S,- o%o- 1 - ~ 10 30

-c-

9

50 D.I. (or + ab + an + Q)

70

I 90

Fig. 3.39. Major oxides and Zr vs DI plots for the Late Intrusive Complexes (fields of high-level intrusive rocks outlined). O = gabbros, 9 = plagiogranites.

I

Fig. 3.40. AFM plot for the Late Intrusive Complexes (generalized trend of the high-level intrusives (dashed line) shown for comparison).

Axis P l a g i o g r a n i t ~

7153 "E

-

OM 7035~

7035 7087 1 ~ 7114 4954

OM 7015

9 /Quartz

TONALITES

diorit~s) :

l\ o ~o

9I 50 " - An

40

30

~ ,i.

.

~

10

e-

~ '20

)M 7153 Axis high-level gabbros

2

q~_

~

4948 i ~

~1 ~ 10

Ab

cc

5

oM 7079t OM7~12~

(b)

OM 7114 er-

)M 7079 OM 7112L K20

ga

Th Ta Ce Hf

Zr Sm Y Yb

Fig. 3.41. Axis (open symbols) and late intrusive complex plagiogranites (closed symbols) from the Semail ophiolite. (a) Normative Or-An-Ab plot (O'Conner 1965) showing the tendency for the axis plagiogranites to be more fractionated. (b) Multi-element comparison ("spider") diagram. Normalized against "average ocean ridge granite" (ORG) taken from Pearce et al. (1984).

I

I

La Ce

I

Nd

I 1

I

Sm Eu Gd

I

Tb

'1

I

I

Tm Yb Lu

Fig. 3.42. REE/chondrite plot for the Late Intrusive Complexes with compositional field of similar axis sequence rocks shown for comparison.

Chapter 3

94 10000

MASHIN

Table 3.13. Sheeted dyke trends (means and standard deviations). See Fig. 3.44.

% //~O"x.,,\ / X /9 "\

Ti

'~/ O0 0

7" \

(ppm)

/t.-"~

,

1000

/ 1 x,

0

//\

0

.U___~_./

[]

n

s

o

Aswad block (all excluding W. Hatta) W. Ajib (W. Hatta)

23 13

173 117

18 14

10 20 (45) 14 13

135 106 (115) 144 168

33 17 (13) 18 16

(31) 19

(175) 165

(14) 22

24

172

12W

14W (24) 15W

177W (192) 172W

23W (12) 29W

llW 16W (20) 16 12

178 162 (163) 169 136

20 20 (20) 7 14

65

152

17

10 (44)

136 (130)

32 (8)

76 56

158 177

18 28

Fizh block

I

W. Fayd/Ashar W. Ragmi (433327228) W. Fizh W. Khabiyat- W. Bani Umar (including Zabiyat section) W. Jizi (Thuqbah section) "Alley block"

I

LASAIL 10000

tO

9

Ti (ppm)

/

0r

~ \

I

1000

Location~area \

i ,7

,/~)

j

_A_

J

Salahi block

\\x x

OoO

x\ \

0

\

0

Northern part: W. Lasail to W. Salahi W. Forest - W. al Hilti (451026~m) W. Ahin - Khabutayah

k.J"

Sarami block

[] A 9 9

Wehrlite Gabbro Diorite Plagiogranite

I

I

10

100

Zr (ppm)

r " x T r e n d o f axis high-level intrusives // " - ' ~ Field of cumulate (._.-. J sequencegabbros

W. Sakhin - W. Shidah W. Sarami (477026453) W. Kanut- W. Shafan Ghuzayn Haylayn block (all) Rustaq block

(W. Bani Kharus 571725977)

Fig. 3.43. Ti vs Zr plots showing fractionation trends for the Mashin and Lasail late intrusive complexes.

3.7 The Sheeted Dyke Complex A sheeted dyke complex between underlying plutonic rocks and an overlying volcanic sequence is a common and significant component of most intact ophiolite complexes. The term "Sheeted Intrusive Complex" was introduced by Wilson (1959) to describe part of the Troodos complex on Cyprus that was composed of essentially 100% near vertical dykes. Gass (1968) and Moores & Vine (1971) later proposed that the Troodos sheeted dyke complex represents 100% extension and is the strongest evidence that the ophiolite was formed at an ocean ridge spreading axis. In the 1970s sheeted dyke complexes were described from other ophiolites with some of the best examples in Newfoundland (Church & Stevens 1971; Williams & Malpas 1972; Rosencrantz 1983), southern Chile (Stern et al. 1976) and western Norway (Sturt & Thon 1978). Statistical studies of chilled margin distributions, dyke widths and sequences of intrusion of the Troodos and other sheeted dyke complexes (Betts Cove, Newfoundland and Smartville, California) by Kidd & Cann (1974) and Kidd (1977) suggested a narrow zone of dyke injection only a few tens of metres wide at the ridge axis. On the Troodos massif facing directions of chilled margins were used to deduce the position of the palaeoridge axis with respect to the present ophiolite outcrop. Whilst not doubting the validity of the theoretical model on which this work was based, several workers (e.g. Gass & Smewing 1981; Gass 1982) have queried Kidd & Cann's find-

Wuqbah block

Maydan syncline W. Mu'aydin

Readings in brackets refer to single dyke sections, UTM grid references given. ings and urge caution in the use of chilling statistics to deduce palaeospreading directions. More recent field studies on Troodos indicate that the predominance of east-facing chills identified by Kidd & Cann (op. cit.), from which they deduced that the ridge lay to the east of the island, is an oversimplification. Early work by the Cyprus Geological Survey (e.g. Bear 1960) indicated that the Troodos sheeted dyke complex was composed of several petrologically different rock types and Desmet et al. (1978), who studied the geochemistry as well as the structure of the Troodos dykes, confirmed the composite nature of the complex and identified up to four separate phases of the dyke injection. In a later paper (Desmet et al. 1980) these authors argued against the model of spreading at a midocean ridge. The Sheeted Dyke Complex of the Semail ophiolite, first briefly described by Reinhardt in Glennie et al. (1974; p. 228), has been studied in detail in the Ibra block in the southeast part of the mountains by Pallister (1981) and in the Maydan area of the Wuqbah block, west of the central part of the mountains, by Rothery (1982, 1983) who described the relations between the dykes and the underlying high-level and cumulate gabbros. Smewing (1980a) recorded that the general trend of the sheeted dykes in the northern Oman mountains is between N-S and NW-SE although in the northern part of the

The Semail Ophiolite Fizh block, around Wadi Ragmi, the dykes trend WNW-ESE, a feature described by Smewing (1980b) as a "leaky" transform fault zone along which local spreading took place. Pearce et al. (1981) showed that both N-S and WNW-ESE trending dykes occur throughout the length of the sheeted dyke outcrop between Wadi Hatta and Rustaq. They recognized three types of dyke on geochemical grounds, described as "axis", "arc" and "clinopyroxene-phyric arc" types, which they related to the Geotimes, Lasail and Cpx-phyric lava units (Section 3.8.1), but were unable to identify any correlation between dyke composition and trend.

95 The individual dykes range from <0.1 m to 4.4 m in thickness but most are 0.5-1 m thick. The jointing is blocky with prominent joint sets parallel and perpendicular to the dyke margins (Plate 3.27). The rocks are mostly hard, compact and massive although there are occasional breccia zones, composed of angular to sub-rounded blocks of dyke rock in an epidote-rich matrix, that are 5 to 50 cm wide and can even affect whole dykes (Rothery 1983). These are associated with zones of quartz and epidote veining and usually trend parallel to the strike of the dykes. They are interpreted as the channelways of hydrothermal fluids.

3.7.1 Field relations and structure

As the name implies, the Sheeted Dyke Complex consists of 100% near vertical dykes (Plate 3.25) that have lava and gabbro screens in the upper and lower contact zones. These contacts are gradational over tens of metres and indicate cogenetic relations between the sheeted dykes and the high-level gabbros below and the lowermost pillow basalts above. The Sheeted Dyke Complex gives rise to a terrain of sharp, elongated ridges (Plate 3.25) whose grain is parallel to the general trend of the dykes. It can be traced as a consistent unit throughout the length of the Semail ophiolite outcrop from just north of Wadi Hatta to the Ibra area in the southeast, a strike distance of nearly 350 km (Fig. 3.44). In the northern part of this area the thickness of the sheeted dyke unit varies from 600-2000 m with an average of about 1500 m. Pallister (1981) estimated that the thickness of the sheeted dykes in the Ibra syncline is 1200-1600 m. The greater part (>95%) of the Sheeted Dyke Complex consists of parallel, steeply dipping dykes. Most contacts between dykes are planar and parallel although there are transgressive, usually thin, cross-cutting dykes (Plate 3.26). Some dykes have small apophyses where small fragments of wall-rock are plucked from the adjacent dykes but xenoliths are rare. Correspondences of the shapes of opposing margins, even of irregular dykes, indicate emplacement by infilling of a tension fracture rather than by forceful injection. Very few dykes, except near the top and base of the unit, are seen to branch or die out vertically or laterally.

t

.....

D / ]r ~'~ASWA ~,~BLOCK ~"

/

t L

ZABIYAT

"~X-~-~'t

\'s',~',~,~C \ ~ . ~ . .

,

BLOC. . . .

/ / RUSTAQ BLOCK

L'

......

C__

/

L\. \

/'~ U

/

'

L_] .................. Meat dyke trends tsee Table 3 7 I ,

,'

<-: : *,,,r,au ~ ~ A; Ar~madl

"9

D,rect,on of spread,rig a•

~---~_/,?

a c c o r d i n g to c h l l h n g data

Rose d,aqrams ShOW dyketrends ~or each

I

~,,s..

f*,4~r,asured dyke Soc[iOrlS IS~ Tabiu 3 T2 :

,~

\

LIN

,see Palhster 1981} / s,~,'9.', matot ,)Oh OII1# D OC,

I 1 ~

10 ,eadmqs

~

.

f~

/

Fig. 3.44. The Sheeted Dyke Complex - - distribution and trends.

Plate 3.25. (a) Sheeted Dyke Complex, (b) Typical jagged ridge (60-70 m high) formed by the Sheeted Dyke Complex.

96

Plate 3.26. (a) Close up of water-washed outcrop of sheeted dykes in Wadi Shafan. "Late" dyke to left of hammer has an irregular shape with matching angular contacts and a small finger-like apophysis into the adjacent dolerite screen. (b) Thin cross-cutting dykes in the Sheeted Dyke Complex.

Plate 3.27. Close up of sheeted dykes with three dykes and two chilled contacts both facing to the right. Pencil 5 cm long.

Chapter 3

The Semail Ophiolite At its upper and lower contacts the 100% sheeted nature of this complex gives way, over a vertical distance of tens of metres, to lavas and plutonic rocks respectively. On a local scale, the mutually intrusive relations between the lower part of the dykes and the various members of the high-level intrusive suite vary considerably. Dykes become more abundant, from 0-10% Of the rock volume, upwards in the high-level gabbros, at the top of which there is a sharp increase in the proportion ~ dykes from 10% to 100% over a vertical distance of up to 70 m (Plate 3.28). Rothery (1982, 1983) showed in a detailed study of a small part of the base of the dyke complex that the contact has an irregular relief of up to 50 m over a horizontal distance of 800 m. The dykes in the transition unit often have complex anastomosing shapes with branchings and blind offshoots (Fig. 3.45) and, where the gabbro forms 3070% of the host rock, the dykes are up to 6 m thick, somewhat thicker than in the sheeted complex proper. Rothery (op. cit.) recognized rooting of the dykes in the underlying massive gabbros where oppositely facing chills of dolerite in a coarser grained gabbro host are up to 10 m apart and can be traced upwards into 1-5 m thick dykes in an analogous way to the relations described by Allen (1975) at the base of the Troodos sheeted complex (Fig. 3.46). The lower part of the sheeted dykes is also intruded by later members of the high-level intrusive suite, often by plagiogranite bodies that may contain abundant xenoliths of metadolerite. These may, in turn, be cut by further dykes parallel to the main swarm, testifying to a complex sequence of intrusive events at this horizon. Below the high-level intrusives, isolated dykes or dyke sets cut the underlying cumulates. In the upper part of the cumulate gabbros these sometimes trend parallel to the main sheeted complex, a good example being the WNW-trending

97

dykes in the Wadi Ragmi area. Rothery (1982, 1983) describes widespread ridge-parallel dolerite dykes from the sheeted complex cutting steeply-dipping layered gabbros in the Wuqbah block as evidence for discontinuous magma chambers along the spreading axis. Lower down in the cumulates and in the mantle sequence, fine grained dykes are rare and usually do not trend parallel to the sheeted dykes (Browning 1982). Many of these are more likely related to the later stages of magmatism, postdating the main "axis" event that produced the Sheeted Dyke Complex. The upper contact between the sheeted dykes and the Geotimes lavas at the base of the extrusive sequence is a sharp transition zone, the nature of which appears to depend largely on whether the lower lavas are predominantly pillowed or massive sheet flows. In the latter case, the sheeted dykes are usually abruptly overlain by an up to 50 m thick sequence of 515 m thick columnar-jointed flows cut by 10-20% dykes. Where pillow lavas predominate, a more gradational transition zone of 30-70% pillow lava screens with irregular thin dykes can be as much as 100 m thick above which the abundance of dykes drops to just a few percent. The sheeted dyke/lava contact of the Semail ophiolite is generally more similar to that

~ D y k e marginwithchill [~ Plate 3.28. Base of sheeted dyke ridge. Thick dykes with gabbro screen at base pass upwards into 100% thinner dykes.

Gabbro host

~

30~ Slope

Fig. 3.45. Field sketch of part of the sheeted dyke - - gabbro transition zone (taken from Rothery 1982). The numbers identify discrete dykes from left to right and do not imply an order of intrusion.

Chapter 3

98

iiiii!!!!!!!i! : iiiiiiiiii: lii!!!iiii iii!iii! ;iiiiiii

iiiiiii!!ii,i,iiiiiiiiiiliiiiiiii, i ii

ii!iii ii

iiilii

...... i i!i'! : : : : : : : : : : : : : : : : : : : : !i!!iiiii!ii!::~'!!iiiiiiiiii!iiii

.'..'..!: .9

99

::% : : : ; ; ;7

[9 :.-'::':";.";i!::..x,..,,/,~..)

9

iii!i?.::.U--c !i. / i.......:i::....

xkk'

11~ slopeI ,711 magma chamber

:::::::::::::::::::::::::::::::::::::

:2' . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

i

Fig. 3.46. Rooting of dykes at the base of the Sheeted Dyke Complex. (a) Model proposed by Allen (1975) for the T r o o d o s Complex. (b) Field relations seen on gentle slopes in the Maydan area (Rotherv 1982). Bhmk - - dolerite dyke. close stipple - - fine grained gabbro-diorite coarse stipple - - coarse gabbro-diorite.

described by Rosencrantz (1983) in the Bay of Islands ophiolite, where there is a narrow transition zone only a few tens of metres thick, rather than that of the Troodos complex where the Basal Group of lavas with abundant dykes is several hundreds of metres thick (Wilson 1959). Above the transition zone basic dykes are rare in the bulk of the lavas and can usually be related to higher lava units such as the Alley and Salahi units (Section 3.8.1). Local dykes and inclined sheet swarms are associated with the Lasail lavas and are focused on central intrusive complexes 9 In the vicinity of the intrusive centres Alabaster (1982) suggests that they locally form sheeted dyke complexes in their own right. A detailed traverse of the sheeted dykes at Zabyat on the west side of the Alley is shown on Fig. 3.47. It contains eighty

sheets which are either complete dykes or parts of split dykes and is divided into two subsections by a fault zone of poor and confused exposure 9 The mean dyke thickness in this section is 0.63 m with individuals ranging up to a maximum of 2.05 m. In a nearby section (Zabayat II) the mean dyke thickness is 0.72 m. These values are clearly less than the original dyke widths due to the splitting of earlier dykes by later ones. Elsewhere, average dyke widths are measured as 1.08 m and 0.95 m (Table 3.14) and the dyke thickness seems to depend on position within the sheeted complex with the dykes generally increasing in thickness towards the base. The Zabyat data are in good agreement with those of Pallister (1981) who showed for the AI A h m a d i outcrop in the Ibra area a mean dyke width of 0.67 m and suggested a pre-splitting average of

W Section A 15 17 14 16 1

2

3

r

4

l5

9",,

'

,'

10

11

12

13

T177 ~

22 18

19

23 25 27 24 26

20 21

T169 o

8

30

31

T163 o

Section B 45 35-38

41

44

46

51 49 52

248

53 50

E

Chilled margin

55 54 4

0 9

58 57

56

63 62

59

60

,

61

66

64

65

71 70

68

67

69

79 78

73

72

74

75

76

77

10 metres I

T Dyke trends

l~/ Hydrothermal breccia veins /I Epidote veins

Fig. 3.47. The Zabyat sheeted dyke section. The tick on the chilled margins shows the direction of chill. T h e eastern end of section A and the western end of B are separated by a 20 m fault zone of poor and confused exposure. A n o t h e r fault zone occurs between dykes 9 and 10.

The Semail Ophiolite

99

Table 3.14. Chilled margin counts on sheeted dyke sections in northern Oman. A verage

Locality

Gridref. Strike~dip

CHW CHE TDTH OWC

dyke-width (m)

Zabiyat 1, Zabiyat 2 Wadi Lasail Wadi Forest Wadi Ahin 1 Wadi Ahin 2 Wadi Sarami 1 Wadi Sarami 2

358932 365936 390805 500665 523573 512561 723522 772458

51 78 40 22 21 15 56 21

0.63 0.72

370/81W 352/85W 341/77W 009/73E 003/65E 017/73E 350/70W 341/72W

46 69 35 19 14 13 53 16

52 105 80 38 62 34 140 60

5.15 6.1 6.7 7.3 (20) 7.1 2.7 13.5

1.08"

0.95*

CHW, west-sided chills: CHE, east-sided chills: TDTH, thickness of section (metres): OWC = (CHW-CHe)/(CHW+CHE) • 100 (after Kidd 19771. * Excluding thin dykes (<30 cm wide)

c. 1.0 m. These values are similar to Rosencrantz's (1983) mean dyke width of 0.9 m for the Bay of Islands ophiolite, but contrast with Kidd's (1977) suggested original dyke widths of 1.8 m for Troodos and 0.6 m for the Betts Cove ophiolite. The frequencies of chilled versus unchilled margins of the dykes in the Zabyat section are as follows: Of the seventy-four dykes for which there are adequate data, fifteen have no chilled margin (these are screens or "septa" of Pallister (1981)), twenty-one have a single chilled edge and thirty-eight are double chilled. Of the seventy-five contacts, twenty-three are "back to back" chills where a dyke has been emplaced along the margin of an earlier one, in which case it is impossible to determine the relative ages of the dykes. The high proportion of back to back chills suggests that there is a m a r k e d preference for later dykes to emplace along pre-existing dyke margins. Study of eight dyke sections between Wadi Jizi and Wadi Sarami, spread along a strike distance of 75 km, is shown on Table 3.14. The most northerly sections at Zabyat were analysed in the greatest detail, for elsewhere most of the thin (<30 cm) double-chilled dykes were not counted and hence the "one-way chilling" ( O W C on Table 3.14) statistic has less significance in these cases. In the Zabyat sections, where all the dykes were counted, the O W C values are 5.15 and 6.1 which compare with the average of 4.6 obtained by Kidd & Cann (1974) for the Troodos sheeted dykes and which Kidd (1977) showed was consistent with the intrusion of 95% of the dykes in an axial zone less than 50 m wide. All the measured sections in northern O m a n show a preponderance of westsided chills (Table 3.14) which suggests that the spreading axis lay to the east of the present outcrop. This contradicts the result obtained by Pallister (1981) who found that the dominant facing direction of the chills in the Ibra area was to the east. It is possible that the spreading axis could have lain between the northern outcrops and the Ibra area, which are presently separated by a cross-strike distance of over 100 km, but Boudier & Coleman (1981) suggest, on the basis of the sense of shear in the mantle tectonite, that the Ibra area also lay on the western flank of the ridge. However, although a position on the western side of the ridge appears most likely for the whole of the Semail ophiolite, multiple spreading axis and spreading centre jumps, as observed on present day ridges (Ballard & van Andel 1977), will result in complex changes in the facing directions of the dykes both along and across the strike of the ridge and may well invalidate any simple relationship between chilling directions and the position of the ridge axis.

ASWAD BLOCK Wadi Ajib (Wadi Hatta) N

FIZH BLOCK NORTH ~ , F i e t d and mean for Wadi Ragmi (n=45) N

N

N

SARAMI BLOCK Ghuzayn area

WUQBAH BLOCK NORTH (Maydan syncline)

FIZH BLOCK SOUTH 9West of Alley ~ East of Alley ( ~ Field and mean for Thuqbah (Wadi Jizi) (n=31) N

SALAHI BLOCK N

N

N

WUQBAH BLOCK SOUTH and MUQNIYAT BLOCK

YLAY N BLOCK Field and mean for Wadi Bani Kharus (n=44) (Rustaq block)

Fig. 3.48. Equal area plots of poles to dykes. Line shows mean trend for each area.

Chapter 3

IOO

are neither sufficiently precise nor clear in their origin to prove whether these changes represent original variations in dyke trends along the ridge axis or whether they are the result of tectonic deformation during nappe emplacement.

3.7.2 Dyke trends The trends of the sheeted dykes throughout the Semail Nappe are plotted on Fig. 3.48. There is clearly a dominance of NW to N trends, but with a subsidiary peak in a WNW direction with a broad spread of intermediate values (note that there are very few strike readings between 020 ~ and 090~ After correcting for tectonic deformation, Pallister (1981) obtained a mean trend of 170~ for the Ibra area. The WNW trend (mean 115~ is largely confined to the northern part of the Fizh block between Wadis Hatta and Zabin along a strike distance of 35 kin. Smewing (1980b) explained the WNW trending dykes in this area as having been emplaced in a spreading direction oblique to the main axial direction along a "leaky" transform fault zone. The presence of both NW and N and WNW-trending dykes in the Fizh block (Fig. 3.48) argues against subsequent tectonic rotations to explain these variations. This is supported by palaeomagnetic data (Section 3.12.2) which suggest that the WNW dykes were emplaced in their present orientation (Shelton 1984). There is a general swing in the trends of the dykes from N-S (170-190 ~ in the southern part of the Fizh block and in the Salahi and West Jizi blocks, to 170-150 ~ in the Sarami, Wuqbah and Haylayn blocks and 130-140 ~ in the Rustaq block (Fig. 3.49). This swing follows the change of strike of the ophiolite from N-S to E-W around the arcuate shape of the mountains. There are some variations within blocks, such as the change in strike of the dykes at the eastern end of the Sarami block from NNW to NW near Ghuzayn. Unfortunately, palaeomagnetic results (Section 3.12.2) from these areas

Ajib (W. Hatta)

The dykes in the sheeted complex have equigranular medium to fine grained ophitic, sub-ophitic and intergranular textures. They are mostly aphyric to sparsely plagioclase- and/or pyroxene-phyric basalts, the main exception being olivine and clinopyroxene-phyric picrite basalts that were emplaced late in the sequence and often cross-cut the main dyke swarm. In general, there is a correlation between grain size and dyke thickness, thin dykes (<50 cm) and small apophyses of larger dykes are black to dark grey at outcrop and microcrystalline with fine grained (< 1 mm) intergranular basaltic textures. Dykes larger than about half a metre wide have a paler grey or green colour and exhibit sub-ophitic to ophitic dolerite textures with grain sizes of 1-2 mm. The chilled margins of dykes where they intrude other dykes are from 0.1 to 5 cm (average about 2 cm) wide and show a progressive decrease in grain size towards a usually knife-sharp contact. The chilled edges are usually darker in colour than the remainder of the dyke and have vitrophyric to aphanitic textures. The primary igneous mineralogy of the dykes consists of calcic plagioclase, clinopyroxene and Fe-Ti oxides (titan 9 netite and ilmenite) but most of these show an almost complete replacement by a greenschist facies assemblage of actinolite, pumpellyite, chlorite, epidote, albite, prehnite, quartz,

W. Khabiyat Ragmi Zabiyat Lasail Forest

Fayd

Aswad

3.7.3 Petrography and mineralogy

+

Bani Kharus

Sarami Ghuzayn

;(W. Jizj~/ Sal;hi

th~fal

Haylayn ~L 1

I

200

..I

180

I

Z"

I9

00

9 9

oo

~

O9

160 9 0O 140 9

n

@@

t.--

+5 120 E <

'

100

"I' : I

ol

80

I

"TransformZone"

60 I 25 O0

I

I 50

I

I 40

Fig. 3.49. Variation of mean dyke trend with latitude.

1

I 30

I

I

I

I

20 10 Latitude

I

I 2400

1

I 50

I

I

40

I

I

30

I

The Semail Ophiolite sphene and secondary Fe-Ti and Fe-oxides and hydroxides. Rare core relicts of clinopyroxene have average compositions of Wo4]En42Fs17and locally at the base of the sheeted complex some dykes contain an assemblage of calcic plagioclase and actinolitic hornblende which represents higher grade actinolite facies conditions (Section 3.10). Most dykes retain their primary igneous textures but there are some that are more severely altered to equigranular textured quartz and epidoterich epidosites. These are often green coloured compared to the more normal grey dykes that are usually characterized by dominantly albite-actinolite-chlorite assemblages (Rothery 1983). The epidosite dykes are sometimes veined or brecciated and the intense alteration is attributed to extreme hydrothermal activity. Gradations between the grey and green dykes can be recognized which have intermediate alteration assemblages such as quartz-chlorite-sphene. In the upper part of the dyke complex the distinction between the grey and green dykes is less clear and they have a predominantly red colour, matching that of the overlying lower pillow lavas, that is probably due to an increase of secondary haematite at this level. The dykes contain up to 5% modal opaque minerals mostly as small (up to 0.5 mm, average 0.2 mm) skeletal grains of secondary Ti-poor titanomaghemite (Fe3_• with mean x = 0.64) sometimes with exsolution lamellae of ilmenite (Shelton 1984). Some alteration to haematite can be recognized and magnetic studies (Section 3.12.2) suggest that haematite is the most stable magnetic carrier in the dykes. Secondary sphene forms in association with the oxidation alteration of the Fe-Ti oxides. Smaller grains (up to 0.05 mm, average 0.02 mm) of magnetite are associated with altered ferromagnesian minerals and probably formed by the replacement of clinopyroxenes by actinolite and chlorite. The picritic dykes are distinguished from the majority of the dykes within the sheeted complex by their strongly porphyritic texture, a generally less altered appearance and occasional cross-cutting relations with other dykes. Pearce et al. (1981) studied the geochemistry of these dykes and, on their low Ti and Zr contents, related them to the Cpx-phyric lava unit (Section 3.8.2). These authors found that of fifty-two dykes studied by them, five were of this low Ti-Zr type and that these occur with both 180 ~ "axial" and 120~ "transform" trends. In the Wadi Ragmi area the picritic dykes are more common than elsewhere and in some cases cut the layered sequence, but they are widely distributed with examples occurring as far south as Ghuzayn in Wadi Hawasina. The picritic basalts contain 10-40% phenocrysts of olivine, usually completely altered to epidote and chlorite, colourless to green diopside, zoned from (Wo45EnsoFs~) to (Wo4o.5En44.sFs15), and less common, partly altered calcic plagioclase (An88_95). The pyroxene phenocrysts are about twice as abundant as the olivines and have a greater size range (0.5-10 mm). The phenocrysts form glomerophyric clusters and have rounded or irregular shapes indicating resorption. The groundmass is medium to fine grained (4 mm to less than 1 mm) and largely composed of altered plagioclase with fine grained secondary actinolite, sphene and epidote, interstitial chlorite and late quartz and prehnite. 3.7.4 Geochemistry: For comparative purposes, analyses of Sheeted Dyke Complex rocks are given together with those of the Extrusive Sequence (see Table 3.18, p. 111) The sea-floor hydrothermal metamorphism that has affected all of the sheeted dyke complex has changed the major and some trace element compositions of the rocks from their pri-

IOI

mary magmatic values. These effects are discussed in greater detail in Section 3.10, but as far as most of the dykes are concerned the changes have resulted in increases in losses on ignition and Fe203/(Fe203+FeO) due to hydration and oxidation and, in most of the normal grey dykes, an increase in Na at the expense of Ca due to albitization of the original calcic plagioclase. The green epidosite dykes show a complimentary increase in Ca and Si and depletion of the alkalis. K20 contents are low (<0.1%) in most of the dykes as a result of the low-K nature of the parent tholeiite magma enhanced by K loss during metamorphism. Most of the dykes appear to be relatively fractionated rocks with low Mg* (MgO/ (FeOt+MgO)), Cr and Ni contents and show a wide range of compositions from basalts to andesites (Table 3.18, Fig. 3.50). In terms of stable element compositions (e.g. TiO2 0.58-2.2% and Zr 45-300 ppm) they show an overlap with the range of compositions of the immediately overlying Geotimes basalts (Fig. 3.51), which supports the field evidence that there is a close genetic relationship between the two units with many of the dykes acting as feeders for the Geotimes Unit lavas. Very few dykes in the sheeted complex appear to correspond in composition to the Lasail basalts (Fig. 3.51). This is contrary to the finding of Pearce et al. (1981) who divided the majority of the dykes into "axis" and "arc" types which they suggested corresponded to the Geotimes and Lasail lava units respectively. Fourteen dykes from Zabyat that were collected over just 20 m of continuous section, in which the relative ages are known from cutting relations, show a wide range of compositions, spanning most of the range of the whole suite, and there is no correlation between the sequence of intrusion and the 9

65--

o~

SiO 2 55

--

9

9 9

9

...,.i'.:..... 9 9 00

9 OO 9

45 I

I

14 12 FeO"

oO

9 99

10

~

8

I

9

OoO

9 9 ~0 ~ 9

9

I

o9

8

.,S.o ~ MgO

70 6O Mg"

50

%0~176 ~

6 4

9

--

9

o~

9

I o

~lJo 1

I

L 200

I 300

go 00-0

0

9

.L-: ee

40

ep~

e

30 I 100

Zr

Fig. 3.50. Compositional variation within the Sheeted Dyke Complex. SiO2, FeO* (total Fe as FeO), MgO and Mg* (100MgO/MgO+FeO*) versus Zr (ppm).

Chapter 3

102

15 Field of Geotimes Basalts.

/

10

/ /

/

\ \

OO

9

9

9

9

\

\

/

r 0 T-

//

x

/

E r

/

9

/

~0

9

9

9

/ /

O_

/

///

9

9

9

9

9

O0

\\

/

/ \

\~

,ID/

//O

/

//

0\

/

/ \

/

I

J

./i" ~ ~ ""

/

/.4

9

Oo~

9 _/~

9

Fig. 3.51. Ti versus Zr plot for the sheeted dykes compared to the Geotimes and Lasail basalts.

~~Las2, ;asalts 0

I

I

50

100

I

Zr(ppm)

I

150

compositions of the dykes (Table 3.15). This suggests that small m a g m a batches of different compositions were being generated more or less randomly in the underlying axial m a g m a chamber. The picritic dykes ("clinopyroxene-phyric arc" type of Pearce et al. (1981)) have higher Mg*, Ni and Cr and lower Zr ( < 4 0 ppm) and T i O , (0.2-0.62%) contents and comprise a distinct group of more primitive rocks compared to

200 the rest of the dykes. Interpretation and modelling of the R E E and other trace element data of the dykes are considered together with the associated lava groups in Section 3.8.5. Six whole-rock samples of the sheeted dykes have been analysed for oxygen isotopes by D. S. Stakes (pers. c o m m . and in Stakes et al. 1984). They have yielded 6 1 8 0 W R values ranging from 6.27 to 9.6% (Table 3.16). These values are higher than

Table 3.15. Intrusive sequence versus compositions of sheeted dykes from the Zabyat section (see Fig. 3.47).

Sp. no.

OM8579

Relations to other dykes

OM8582

Cut by 8576 and 8580 Cut by 8582 and 8585 Cut by 8573 and 8575 Cuts 8574, cut by 8 5 7 6 Cuts 8 5 7 5 and 8579 Cuts 8579, back-to-back chill with 8581 Back-to-back chill with 8580, cut by 8582 Cuts 8 5 8 1

OM8585

Cuts 8 5 8 3

OM8583 OM8574 OM8575 OM8576 OM8580

OM8581

OM8586

OM8573

Cuts 8574

Dyke o'pe

Zr (ppm)

Ti (ppm)

Dolerite screen

45

4869

1.09

Dolerite screen

61

6360

1.20

Single-chilled dolerite Single-chilled dolerite Double-chilled dolerite Double-chilled dolerite

49

5820

1.32

65

7860

1.71

63

7140

1.89

48

5100

1.21

Double-chilled dolerite

88

10200

2.56

Double-chilled dolerite Double-chilled dolerite Double-chilled porphyritic (cpx, plag-phyric) dolerite Basalt, vein-like apophysis

36

3600

0.86

56

6900

1.78

49

462(/

0.92

70

7860

1.72

All medium to coarse grained aphyric dolerites except 8586 and 8573. Order of intrusion is from top to bottom. FeO ~ = total Fe.

FeO'/MgO

The Semail Ophiolite

10 3

Table 3.16. Oxygen isotope analyses of Semail dykes (data from D. Stakes).

No.

Locality

Secondary mineral assemblage

6'aOwn

"Axis" or type 1* "Axis" or type 1* "Axis" or type 1* "Axis" or type 1* "Arc" or type 2*

Act, chl

6.39

Act, chl, qtz, preh

6.27

Act, ep, sph

7.81

Act, chl, pump, preh, sph Act, chl, sph, pump

9.60

,)

Act, chl, sph, preh, pump

9.42

Chi, ep, qtz, sph (fresh calcic plagioclase) Chl, ep, qtz, sph Chl, ep, qtz, sph

5.98

Rock type

Dyke class

OM5060 Wadi Hatta

Aphyric dolerite

OM5066 Wadi Ragmi

Aphyric dolerite

OM5080 Wadi Sarami

Aphyric dolerite

OM5085 Wadi Kanut

Aphyric dolerite

OM5089 Haylayn

Aphyric dolerite

OM5099 Wadi Fizh

Aphyric dolerite (some fresh cpx)

7.00

"Late" dykes OM5104 Ghuzayn OM7041 Wadi Fizh OM7124 Ghuzayn

ol-cpx-phyric basalt (ol replaced by epidote + chlorite, cpx fresh) ol-cpx-phyric basalt ol-cpx-phyric basalt

ol - olivine; plag - plagioclase; cpx - clinopyroxene; act - actinolite; * dyke classification of Pearce et al. (1981)

chl- chlorite; qtz- quartz;

the accepted primary magmatic ones (c. 5.8"/0,, for M O R B ) and are ascribed to interaction, at temperatures of less than 500~ between the dykes and l~O-enriched hydrothermal fluids that had previously reacted with the underlying gabbros (Gregory & Taylor 1981; Stakes et al. 1984). There is no obvious correlation between the metamorphic mineral assemblage and the 6 ~ O values although the most severely altered epidosites were not analysed in the study. The later picritic dykes have lower 51sO values (5.59-5.98% o), closer to the primary magmatic values than those of the aphyric dykes (Stakes et al. 1984). This supports the contention that these dykes were emplaced "off-axis", after the main axial phase of magmatism and were only affected by a late stage of less intense, hydrothermal alteration.

3.8 The Extrusive Sequence The upper part of the Semail ophiolite, that overlying the Sheeted Dyke Complex, consists of up to 2000 m of pillowed basaltic lavas with some andesite to rhyolite extrusives that are cut by basic to acid hypabyssal intrusives. The extrusives are both intercalated with and depositionally overlain by deep water sediments (Section 3.11), the whole representing the products of submarine volcanism and sedimentation at or close to a mid-Cretaceous spreading axis. 3.8.1 Volcanic stratigraphy The division of the Semail Extrusive Sequence into five units (Table 3.17) on the basis of field and petrological characteristics has been developed by Alabaster et al. (1980, 1982),

preh - prehnite; ep - epidote;

5.59 5.84

sph - sphene; pump - pumpellyite.

Pearce et al. (1981) and Alabaster (1982). Alabaster et al. (1980) originally proposed a three-fold division into basal Geotimes, Lasail and upper Alley units. Later, Pearce et al. (1981) added the upper "Cpx-0" (clinopyroxene-phyric) and uppermost Salahi units. All these units are described in detail by Alabaster (1982). 3.8.1.1 Geotimes Unit The G e o t i m e s Unit is the lowermost lava unit which was named from a locality in Wadi Jizi (GR437226 sS~ first featured on the cover of issue no. 8, volume 20 of Geotimes Magazine. This unit directly overlies the Sheeted Dyke Complex with a gradational contact described in Section 3.7. It is composed of 7501500 m of red to brown coloured basaltic to andesitic pillow lavas with some massive flow units and occasional pillow breccias and hyaloclastite tufts. The thickest flows, which reach 6 m, often possess a massive columnar-jointed central part and have pillowed tops and bases. The pillows are mostly large tubes or bolsters often over 3 m long and up to 1.5 m in cross section (Plate 3.29). The pillow surfaces often show a rectilinear pattern of cooling joints. The majority of the Geotimes lavas are non-vesicular and aphyric. Pillow interiors and the massive flows have intersertal to sub-ophitic textures whereas pillow margins exhibit quench skeletal and variolitic textures with small microphenocrysts of clinopyroxene and albitized plagioclase. Where fresh, the pyroxenes are augites with an average composition of Wo42En43Fs15. The groundmass is altered to chlorite (iron-rich smectite), quartz, haematite and Fe-oxides and hydroxides; many are so heavily iron-stained that they appear almost opaque in thin section. The occasional, small vesicles are infilled with chlorite, quartz, epidote, prehnite and stilbite.

Chapter 3

IO4 Table 3.17. Characteristics of the Semail Extrusive Sequence.

R E E-chrondrite Incompatible trace normalization elements (Fig. 3.59) (Fig. 3.61)

Magma type/ Petrogenesis/ Tectonic setting

100--190

Enriched relative to N-type MORB especially Th, Ta and Nb.

LREE-enrichcd: LaN/YbN>5. LREE>MREE >HREE

Mildly alkaline to transitional within-plate basalts. Erupted well offaxis possibly from E-W fissures. Small degrees of melting of an undeplcted mantle.

10-40

Highly depleted relative to N-type MORB especially Ta, Nb, Zr and others.

LREE-depleted Las/YbN<0.3.

Island-arc tholeiite suite. Picritic basalts and basalts erupted from N-S to WNW-ESE fissures in interseamount graben.

Unit Rock types (thickness) (phenocrvsts)

Clinopyroxenes (Fig. 3.60)

Geochemistry FeO*/MgO

770,

Zr

SALAH1 (up to 100m)

Basalt (cpx, plag, op)

Salite Wo45En3, (high Na, Ti)

1-3

1.2-1.9

CPX-O (up to 50 m)

Picrite basalt Diopside Wo42En51 (ol, cpx, plag, chr) (very low Na, Ti)

1

0.2~1.5

ALLEY (up to 500 m)

LASAIL (up to 500 m)

Basalt (cpx, plag, op)

Augite Wo43Enat (very low Na, Ti)

1-2.5

0.45-1.1

25-55

LREE<MREE
Intermediatc to acid extrusives form local centres above high-level magma chambers. Hydrous melting (second-stage?) of depleted mantle.

Transitional MORB to island-arc tholeiite suite. Moderate (20-30%) degrees of melting of a slightly depleted mantle. Erupted to form seamount centres. Intermediate to acid rocks formed by closed-system fractionation in high-level magma chambers.

0.2--0.9 Andesite, dacite, rhyolite (pitchstone) (plag, op, qtz)

2.5-11.5

Olivine basalt (ol, Diopsidic augite cpx, plag) Wo41Enso (low Na, Ti) Andesite, (rhyolite) (plag, op)

1

55-120

O,34). 9

20-60

Depleted relative to N-type MORB especially Ta and Zr.

LREE-depleted LaN/Yb-~0.6 LREE<MREE =HREE

60-200

Slightly depicted relative to N-type MORB.

Eruption of relatively Slightly LREEdepleted LaN/YbN evolved magmas at a N-S spreading axis. 0.7-1.0 Fractionated in large, opensystem magma chambers.

Transitional Group- intermediate characteristics bctwccn Lasail and Gcotimes basalts. GEOTIMES (7501500 m)

Basalt, andesite (cpx, plag, op)

AugitcWoa2En42

1.5-5.0

(low Na, Ti)

Plate 3.29. (a) ' G e o t i m e s ' U n i t , W a d i Jizi. (b) C l o s e - u p o f s a m e .

0.9-2.2

The Semail Ophiolite 3.8.1.2 Lasail Unit

The Lasail Unit comprises green coloured basaltic lavas that are usually highly altered and which, in outcrop, disintegrate easily compared to the more compact and harder Geotimes pillows. The striking colour difference between the Geotimes and Lasail lavas can be used as a sharp line of demarcation in the field. The Lasail pillows are usually smaller (< 1 m in cross section) and more bun-shaped than those of the Geotimes (Plate 3.30). They are composed of sparsely to moderately porphyritic basalts with small (1-2 mm) olivine (altered to epidote and chlorite), clinopyroxene (partly altered to actinolite) and plagioclase (altered to prehnite) phenocrysts. The pyroxenes are distinctly more magnesian than those in the Geotimes lavas and are diopsidic with an average composition of Wo43En49Fs 8. The groundmass of the Lasail basalts is typically non-vesicular with fine grained skeletal to intersertal textures. Chlorite, epidote and prehnite are common in interstices, cavity infillings and as alteration products of phenocryst and groundmass minerals. Other secondary phases present are quartz, calcite, sphene, Fe-oxides and hydroxides and pyrite. The Lasail basalts range up to 750 m in thickness, but, unlike the Geotimes lavas which are found at the base of the extrusive sequence in nearly all areas, they are restricted in their distribution (Fig. 3.52). In add!tion, although the contact between the Geotimes and Lasail Units is usually sharp, as in the Wadi Jizi (Lasail) area, some sections (e.g. Wadi Hatta and Mahab areas) have an intervening series of "transitional" lavas up to 300 m thick with intermediate geochemical characteristics (Fig. 3.52). As an additional complication ia the Wadi Hatta area, Lasail-type basalts locally occur at the~oase of the lava sequence (Alabaster 1982). The Lasail basalts and the immediately underlying Geotimes basalts are cut by numerous andesitic cone sheets. These range up to 2 m thick and locally form up to 40% of the rock volume (Plate 3.31). There are up to four sets of cone sheets focused on distinct centres in the Lasail area (Fig. 3.53). The andesites are typically orangy brown in colour with a blocky jointing and are hard and compact rocks that are generally non-vesicular, although zones of small vesicles, infilled with epidote, quartz prehnite and pyrite occur along the margins. The rocks are fine grained with groundmasses composed of altered feldspar microlites, chlorite, haematite and quartz and with sparse microphenocrysts of altered titanomagnetite and plagioclase.

Io5

Alabaster (1982) showed that the cone sheets of the Lasail area were emplaced after the eruption of the Lasail basalts but before the Alley Unit. In the same area there are localized felsite sills also associated with the Lasail Unit activity. These are up to 12 m thick and in some places as many as six individual sheets are emplaced one upon the other or separated by thin screens of basalt. The felsite sills often have flow banded margins with elongate vesicles showing the direction of flow and massive columnar jointed interiors. They are aphyric to sparsely porphyritic with altered plagioclase and rare quartz phenocrysts with a fine grained, often spherulitic groundmass. Hydrothermal alteration, marked by breccia zones rich in chlorite, epidote and pyrite, is common along the sill margins and occasionally in rubbly central zones. The felsite sills were emplaced after the inclined sheets and represent the last phase of Lasail Unit magmatism in the Lasail area. In contrast to this area, where intermediate to acid composition hypabyssal rocks are abundant, in the Wadi Hatta and Mahab (Wadi Sarami) areas the top of the Lasail Unit comprises a series of massive columnar-jointed brown andesite flows, up to 40 m thick, with a characteristic speckled appearance in outcrop. In other areas, such as Wadi Fizh, the Lasail basalts and the andesite flows are interbedded (Plate 3.32). In all these areas where there are abundant andesite extrusives, the cone sheets are rare or absent. The localization of the Lasail Unit intrusive-extrusive complexes to restricted areas, 15-20 km across and separated by roughly 20-30 km along strike, led Alabaster et al. (1980) to propose that they represent "seamount" central vent-type volcanic complexes built up at intervals on top of the underlying oceanic crust composed of the Geotimes Unit and the Sheeted Dyke Complex. This model was further developed by Pearce et al. (1981), although the compositional variations in the sheeted dykes recognized by them did not appear to relate to the seamount areas. Alabaster (1982) identified the main Lasail Unit centres as, from north to south, (1) the Wadi Hatta-Fizh complex, which overlies the supposed "transform" section of the ophiolite where the predominant dyke trend is WNW-ESE, (2) the Wadi Jizi-Lasail area, (3) the MahabJebel Fayyad complex in Wadi Sarami and (4) the GhuzaynMabrah complex, southeast of Wadi Hawasina (Fig. 3.52). He noted a close association between these areas and the occurrence of massive sulphide mineralization in the lavas, often at the base of the Lasail Unit basalts (Section 3.10.4). Another

Plate 3.30. Typical Lasail Unit pillow lavas sandwiched between two subhorizontal andesite sheets. The lavas are cross-cut by a feeder dyke. From Alabaster (1982).

Chapter 3

Io6

blocky, autobrecciated flow tops and bases. Within the Alley Unit there are a wide range of petrological types ranging from basalts to rhyolites. The basic rocks are mostly microporphyritic with phenocrysts of altered plagioclase, replaced by albite and stilbite, and generally fresh colourless augite (average composition Wo40En44Fs16) often in small glomerophyric clusters. They have fine grained intersertal and variolitic textured groundmasses composed of a network of altered plagioclase laths and interstitial celadonite, chlorite (smectite), carbonate, stilbite and mesolite. The green coloration produced by celadonite is a distinctive feature of many of the Alley basalt-s. In two areas, in the Wadi Jizi-Lasail area and near Mabrah in the Haylayn block, there are small volume dacite to rhyolite lava flows, consisting of grey or pink coloured felsite and black flow-banded pitchstone, in the upper part of the Alley Unit. It seems likely that these rocks were originally mostly composed of glass and that the pitchstone layers represent relict, nondevitrified material (Plate 3.35). The flows are up to 25 m thick and are massive to autobrecciated and highly heterogeneous, some are composed of large pillow-shaped masses up to 2.5 m in diameter. In the Wadi Jizi-Lasail area there are four separate outcrops of the acid volcanics, each forming a small ridge or isolated peak that represents a single extrusive body from 30 m to 300 m in length. The largest outcrop forms a NW-SE

feature of the "seamount" areas is the presence of large (up to 6 km 2) differentiated plutonic complexes that intrude into the Sheeted Dyke Complex and the Geotimes lavas. These late intrusive complexes, such as that located 5 km southwest of the Lasail ore body and Jebel Fayyad in Wadi Sarami, are the intrusive equivalents of the Lasail Unit extrusives. In contrast, in the intervening areas, called "interseamount" by Alabaster et al. (1980) and where Alley Unit lavas directly overlie those of the Geotimes Unit, there are no large late intrusive complexes nor seemingly any major sulphide mineralization. 3.8.1.3 Alley Unit

The Alley Unit, which overlies either the Geotimes or Lasail Units, comprises up to 500 m of grey or brown coloured pillow basalts that can be highly vesicular with distinctive amygdale infillings of zeolites (Plate 3.33). Where the Alley basalts directly overlie the Geotimes basalts there is usually a 2-3 m thick umber horizon at the contact. The Alley basalts exhibit a variety of pillow forms (Plate 3.33, Plate 3.34), from bunshapes to tubes, whereas some flows are sheet-like with ~Oo oOo*~j

Wadi Hatta

0 Fig. 3.52. The Extrusive Sequence Outcrop, stratigraphy and the distribution of the "Seamount" areas, -

Wadi Fayd

1 km

/1S (WadiRagmi

-

key to columns

~ SalahiUnit

--. Mn-Fe ''~\\\ Wadi F i z h ~

Cpx-~ Unit Alley Unit (Rhyolite in upper part) Ferromanganoan Umbers Mn- Fe Massive Flows ', Lasail Unit ::)0) Pillow basalts oo : o ~ --Massive sulphides Fe-Cu Transitional Unit R

The "Alle'

ii :~iilhll/I

[.... =-J,Fe-Cu

Wadi~/./~ Lasail ~

I

]

~

Mn-Fe

,

, Mn-Fe

R

R

~ t

.

-

.

-

.

.

.

.

~

Geotimes Unit i

Wadi Salahi ,

West Wadi, Wadi

Sheeted Dyke Complex

i ,~'1~

Mn-Fe LF_-~-= ~

(~.~Sb "Seamount" areas (see text) Massive sulphides and gossans

\

sX \2

t' ~lilEil

\

25km I

Wadi Mu'aydin ~

Mn-Fe

~

" ~kah

~ a b r a h

The Semail Ophiolite trending ridge where there are associated inclined sheets and dykes of felsite cutting the underlying Alley basalts close to the base of the acid extrusive. A local set of NE-SW trending enechelon felsite dykes occurs just to the south of the small acid extrusive centre in Wadi Lasail (Fig. 3.53). The acid lavas are aphanitic to vitreous rocks with occasional microphenocrysts of sodic plagioclase and quartz. Small vesicles are common and are often infilled with secondary chalcedonic quartz and epidote. Small (<2 km), differentiated plutonic bodies, like the Mashin intrusion, that are the probable plutonic equivalents of the Alley Unit extrusives, are described in Section 3.6.

\

3.8.1.4

107

Cpx-r Unit

The Cpx-r (Clinopyroxene-phyric) Unit consists of picritic or ankaramitic basalts that form a series of small-volume flows associated with localized dyke sets often along NW-SE trending faults. The associated dykes, which are petrologically similar to the lavas, have been described previously in Section 3.7. The association of small, "late" peridotite intrusives and the picrite dykes and lavas is discussed in Section 3.6.2.

3.8.1.5 Salahi Unit The uppermost Salahi Unit is only locally found at the top of the lava pile between Wadis Salahi and Forest (al Hilti) in the Salahi block where it forms up to 100 m of pillowed and mai columnar-jointed lava flows (Plate 3.36). Some of the Salahi basalt flows are exceptionally thick, up to 30 m, and the flow centres are coarse grained (4 mm) with ophitic textures. Large colourless salitic augites (average composition Wo44En39Fs14 ) enclose altered plagioclase that has been replaced by albite, prehnite and epidote. There is abundant interstitial chlorite which, with prehnite and carbonate, infills cavities and veins. Some feeder dykes, that are up to 3 m wide and trend 120~ to 150~, cut the underlying Alley Unit in the Salahi area. Alabaster (1982) also identified a few dykes with the same composition as the Salahi flows in the Sheeted Dyke Complex.

\

Wadi Jizi

3.8.2 Geochemistry

/

ILj lie

t

j

Andesite cone sheets Horizontal

F

Rhyolite extrusives at top of Alley Unit

> 40 ~ dip

Felsite sheets in Lasail Unit

~- 20-40 ~ dip

k

Lasail late intrusive complex

< 20 ~ dip

i~)

Felsite dykes

[~

Focal centres of cone sheet-radial dyke swarms Lasail gossan

Fault (tick on downthrown side)

Fig. 3.53. Detailed map of acid and intermediate minor intrusive and extrusive rocks in the Lasail area.

All the Semail hypabyssal and extrusive lavas have suffered ocean-floor hydrothermal alteration which has been responsible for the almost complete replacement of primary mineral assemblages by secondary phases and which, in the case of the lavas, was probably superimposed on the effects of low temperature submarine weathering. The effects of these changes on bulk chemical compositions have to be taken into account in reconstructing primary magmatic compositions and petrogenetic relationships. Most of the major elements are highly susceptible to change and may be lost or gained depending on the nature and intensity of the alteration process (e.g. Cann 1971; Wood et al. 1976; Humphris & Thompson 1978). High Fe203/FeO ratios and volatile contents (expressed as losses on ignition (LOI) in the analyses) well above typical magmatic values are found in all the rocks. The effects of alteration extend to some geochemically important trace elements, especially the light ion lithophile (LIL) elements, such as Ba, Sr and Rb (Hart 1973). In order to compare the altered lavas with their modern or unaltered equivalents, recourse has to be made to those elements that are believed to be stable at low to moderate grades of metamorphism. The most important of these are the high-field strength (HFS) elements, such as Ti, P, Zr, Nb, Ta and Y and the middle and heavy rare earths, and some transition metals, notably Ni and Cr (Cann 1970; Hart 1970; Pearce & Cann 1971, 1973; Thompson 1973). Even so, there are conflicting data for some of these and Mattey et al. (1980) demonstrated that altered pillow interiors of sea-floor basalts showed small increases of Ti, P and Zr compared to the pillow rims. The light rare earth elements (LREE) (e.g. La, Ce, Nd) are even more controversial, some authors (e.g. Hellman et al. 1979; Ludden & Thompson 1979) have shown enrichment of La and variable behaviour of Ce in severely altered basalts, whereas Menzies et al. (1979) and Hajash (1984) report no changes of LREE contents during experimental seawater and hydrothermal greenschist facies alteration of

I08

Chapter 3

basalts. It is certain that the temperature of alteration and retrograde reactions are critical factors. Alabaster (1982) concluded, on the basis of major oxide versus Zr plots that all the major elements, except possibly Ti and P, had been mobile during alteration of the Semail lavas. Fig. 3.54 bears this out, where even the TiO2 versus Zr and P205 versus Zr plots show considerable scatter when compared, for example, to the near-linear plot of Y versus Zr. Nonetheless, despite alteration effects, some of the general trends seen on this figure reflect magmatic processes. For

example, the FeOt/MgO ratio, which is an index of fractionation, shows a broad positive correlation with Zr. Another test of whether the FeOt/MgO of the lavas reflects the primary magmatic compositions is the plot of the whole rock ratio against that of the least differentiated (lowest Fe/Mg) clinopyroxene composition in the same rock (Fig. 3.55). The positive correlation (which is close to unity using an equilibrium distribution coefficient (FeOt/MgO cpx/liquid) of 0.29 taken from Thompson (1974)) suggests that the whole rock ratios do indeed reflect the primary magmatic values.

Plate 3.31. Lasail Unit lavas c u t by numerous inclined andesitic cone sheets.

Plate 3.32. Blocky andesite flow overlying basaltic lava (poorly exposed). A columnar jointed andesite sheet has been intruded along the contact between the lavas.

Plate 3.33 Alley unit pillow lavas, with characteristic large zeolitefilled amygdales and interpillow cherty sediments.

Plate 3.34. Radially-jointed Alley Unit pillow basalt.

Plate 3.35. Part of glassy rhyolite flow in the Alley Unit. Dark patches near hammer head are rhyolitic glass (pitchstone).

Plate 3.36. Massive, columnar-jointed Salahi Unit basaltic andesite

lava flow.

The Semail Ophiolite (a)

(b)

140

-

9 o 9o

60 - 9 SiO2

9

55 -

o 9

1412 -

9

o

o9

~q_>o

Ca 9

9

o

6

o

9

I

I

9 o~e~Ld~. 00 9 . 9 ~c?o~ 9 |%,~I 9 o 99 9 9 o o Qo~ oo~ o9 o u o~O ~ o ..o._~..+,.

~,~.,..~

99

K20 0.5 0.1

9 o

I

9

0

l

.+...

o

--

O~

I

9

~" ~'0

9

P+Os

0.4 02

--

0 9.

, 9

l

Y

60

20

o

-

9149 .~sql

~ill~'m'~'' I

,~-,

o o + . o --

9

.-o

I

9

04 02

~~

9

6

I~

_

9 I

I

4 -

9 9

-'-

o

oO 0

4 FeO t MgO 3 -

"~"o

9

9

9

O0

9

o 9

_ ".p 5,D _ o o~o9 ~ 9 o311,OOor 9 o--Q ~ 9 ~, 9 --'

I

100

9

o t'_s

I

9

09

1

f

9

o

o9

o

-

~q'~m~~149 _

o~

oOO0

0

MnO

9

-

C

8

,"

C

9 C800

CO

-~O 9 :~

91

100

]

--

10

9

~ 0 0

::

I

o.- . . . . , ,

+'I$~-

2

o

9

12

-

%

oo

.,

9

9 9

~ 1 4 9 1 94 9 9 1 7 6 oo 9 ee

no~

o

o~

10

o

oa

2

MgO

~ 9

L~

I

FeO t

9

9 ooo

--

~ "9

o ~ l,

16 A'20+

~o9

Zr(ppm)

I

I

200

300

..

9 l.~'~-'o o ~"'- 9

9 ~ 9 ~

0

2 !

-

I 100

Zr(ppm)

I 200

300

Tables 3.19-3.24 show a selection of analyses of the Semail lavas and minor intrusive rocks mostly taken from Alabaster (1982). Only those analyses showing the least effects of secondary alteration and therefore closest to the original magmatic compositions are given.

O O

-

o

/o

9 / 0

/

o9

FeOt/MgO (cpx)

r

I

L

1

2 FeO t / MgO

, 1 3

( w h o l e rock)

Fig. 3.55. FeOt/MgO for whole rocks and coexisting clinopyroxene phenocrysts (symbols as in Fig. 3.57 p. 116)+

Fig. 3.54. Geochemical variation diagrams for the sheeted dykes and lavas. 0 = sheeted dykes, 9 = lavas.

3.8.2.1 Sheeted dykes and Geotimes Unit The Geotimes lavas and the underlying Sheeted Dyke Complex are mostly affected by "spilitic" greenschist facies alteration that is reflected geochemically by losses of SIO2, which shows a general decrease with increased losses on ignition, and increases in Na20 (up to 7%) with low contents of C a 9 ( < 1 0 % ) resulting from the exchange of Na for Ca during albitization of the plagioclase. Ca is clearly lost from these rocks during alteration and is not retained in secondary Carich silicates such as epidote or prehnite. Oxidation ratios (Fe2Os/(Fe203+FeO)) of 0.55-0.95 and losses on ignition (1.5-8%) of the Geotimes lavas are generally higher than the corresponding values (0.29-0.67; 1.5--4%) in the sheeted dykes probably largely reflecting the effects of early sea-floor weathering and lower grade alteration of the lavas. K20 contents of the dykes and lavas are generally low ( < 0 . 1 % ) but with occasional higher values up to 1.5% in the lavas. This may be the result of the superimposition of retrograde low-temperature alteration on the greenschist assemblage (Hart 1974) or be associated with the formation of K-rich clays during seafloor weathering. The sheeted dykes and Geotimes lavas broadly overlap in composition on the TiO2-Zr plot and are mostly relatively fractionated basalts, basaltic andesites and andesites; the low MgO ( < 7 % ) , high FeOVMgO (>1.5) and negligible Cr and Ni contents all indicating considerable degrees of pre-eruption

Chapter 3

IIO Table 3.18. Analyses

of Sheeted

Dyke

Complex

rocks.

Basalts 0M5089 Hy

A ndesites 0M657 J

0M5112 Fo

0M163 R

0 M 5 1 0 8 0M5065 0M7040" OM5111 0M5115 0M5994" 0M7122 0 M 5 0 9 7 A H F Fo Y Sr Hw Hw

0 M 5 0 7 4 0M5094 0 M 6 8 4 R Hw J

SiO2 TiO_, A1203 F%O3 FeO MnO MgO CaO Na,O K20 PeOs LOI

50.43 0.58 15.82 2.31 5.74 0.13 8.49 11.40 2.36 0.05 0.04 2.23

50.20 0.84 14.9(/ 2.86 5.55 0.17 8.10 11.86 2.66 0.10 0.05 2.85

49.25 0.76 15.90 2.67 5.83 (/.16 8.71 10.44 3.83 (I.(12 0,07 3,50

50.30 1.24 14.9(I 3.51 6.97 0.17 6.79 11.36 3.11 0.08 0.09 1.83

51.20 1.72 14.20 5.07 5.71 0.19 5.73 6,13 5.14 (I.49 0.16 3.61

49.51 1.58 14.81 5.33 6.82 0.20 6.07 9.12 4.03 0.10 0.20 1.71

49.50 1.53 14.40 5.85 6.48 I).30 5.95 9.68 3.16 (I.08 (I.17 1,71

50.20 2.(17 13.5(I 5.59 7.94 0.20 5.68 3.89 5.85 0.18 (I.16 3.22

53.30 1.64 13.80 7.13 3.59 (I. 17 4.11 4.56 6.15 0.02 0.13 4.(19

51.6(I 2.24 14.30 6.67 6.61 0.20 4.43 4.07 5.05 0.19 0.21 3.15

50.00 1.91 14.3(/ 7.22 5.96 0.36 4.93 6.94 4.32 nd 0.25 3.06

54.20 1.72 13.70 6.40 4.78 (I.25 3.75 5.10 5.33 0.01 (I.21 2.78

55.60 1.6(I 14.00 6.61 4.98 0.16 3.53 3.53 5.81 0.10 (I.39 2.74

58.00 1.21 13.90 6.23 3.39 0.09 2.07 6.35 3.66 0.14 0.38 3.16

60.70 1.00 14.00 4.80 4.11 0.13 2.93 2.30 6.09 0.07 0.32 2.73

Total

99.62

100.15

100.14

100.26

99.35

99.84

98.91

98.48

98.70

98.23

99.24

98.23

99.05

98.59

99.18

Q or ab an di hy ol mt il ap

0.30 20.06 33.83 19.65 23.15 0.62 1.15 1.14 1.10

0.59 22.52 28.45 24.49 13.18 5.02 1.19 1.6(1 0.12

0.47 26.29 26.44 24.09 9.95 6.90 1.48 2.35 0.21

3.04 45.71 14.96 13.14 4.13 13.64 1.55 3.43 0.40

0.61 35.11 ~,.6"~'~9 18.70 3.73 13.88 1.78 3.09 0.49

0.49 27.74 . . . . 8 "~ "~5 18.98 18.97 2.85 1.75 3.(11 0.42

1.12 52.34 1(I.63 7.22 1.46 20.71 1.96 4.16 0.40

0.13 55.49 1(I.66 10.50 13.43 4.65 1.53 3.32 0.33

3.30 (I.06 47.63 14.19 9.30 19.97

4.14 0.61 51.07 12.29 2.74 22.34

17.67 0.87 32.70 22.27 7.13 14.63

11.31 0.42 52.16 9.44 0.50 19.49

1.59 3.45 0.52

2.71 3.16 0.96

1.36 2.43 0.95

1.25 1.92 0.76

82 5 52 162 277 34 21 80

nd nd 4(I 196 377 43 15 120

38 51 38 148 303 42 106 100

nd nd 36 112 430 48 16 139

nd nd 24 41 386 45 150 145

Cr Cu Ni Sr V Y Zn Zr

0.12 29.40 26.87 21 .(13 2.12 17.61 1.21 1.48 0.17

135 39 70 178 208 24

14

21

47

37 48

54

Key to locations (Tablcs 3.18-3.231:

Table 3,19, Geotimes

Unit.

38 97

0.01 1.18 45.(/8 16.67 2.67 27.46

38.34 20.60 11.55 18.76 4.47 1.87 3.80 0.62

1.91 4.49 0.52 9 10 40 153 208 55 32 153

nd 20 21 130 405 46 195 124

nd nd 38 91 256 48 94 142

= A h i n , F = F i z h , F o = Forcst. 11 = tlatla, lily = llawasina, tlv = Itaylayn block. Rustaq. S = Salahi. Sr = Sarami. Y = Yanbu.

A

Analyses

of Extrusive

Sequence

nd 35 43 88 216 70 7 214

nd 18 60 207 81 94 nd 297

J = Jizi, M = Mahab,

R =

4 76 5 91 82 30 80 124 Ragmi. Rq

rocks.

Basalts

A ndesites

Transitional basalts

0M5827 H

0M5831" H

0 M 5 9 3 8 0M7043 0M4279" Fo F J

0M5756" Fo

0 M 7 0 6 7 0M5745" Hw Fo

0M7003 0M7106 0 M 5 9 4 0 Sr Hw F

P205 LOI

50.7 1.22 15.1 6.84 3.82 0.15 3.21 6.65 5.42 0.02 0.10 5.43

47.2 1.35 17.2 7.51 3.84 0.20 3.56 7.81 5.31 0.03 0.14 4.82

48.2 1.38 15.0 7.(12 3.39 0.20 5.90 6.90 4.67 0.80 0.12 4.75

5(I.6 2.23 14.4 11.19 2.51 0.23 3.17 4.82 6.72 0.11 0.14 3.32

52.1 2.09 15.3 8.55 3.(13 0.14 3.76 4.30 6.54 0.04 0.26 2.58

49.4 1.88 16.5 8.09 3.84 0.24 3.74 4.13 5.86 0.81 0.16 3.97

52.6 1.59 12.8 7.96 3.18 0.20 3.56 5.94 5.33 11.03 0.15 6.25

57.9 1.47 14.4 6.22 2.58 0.11 2.86 2.48 6.56 I).112 0.20 3.38

56.5 1.87 13.8 7.78 3.33 0.19 2.31 6.00 5.45 (I.(12 0.24 2.35

54.1 1.86 14.3 9.41 2.(17 0.13 2.00 3.95 7.58 0.03 0.22 3.37

50.0 1.05 15.3 6.04 2.96 0.48 7.04 5.95 5.30 0.22 0.12 4.14

5(I.2 0.81 17.1 4.11 3.85 0.12 5.12 7.95 4.31 0.01 0.04 6.411

5(I.5 (I.91 15.4 3.26 5.67 0.18 7.0(I 8.54 3.29 0.21 0.07 3.42

48.7 0.95 15.3 3.95 5.17 (1.14 7.45 9.55 4.04 0.14 0.11 3.83

Total

98.66

98.78

98.33

99.43

98.70

98.62

99.58

98.18

99.84

99.111

98.30

100.02

98.46

99.33

O or ab an di hy ol nat hm il ap

3.75 I).12 46.48 17.05 12.24 2.42

0.18 43.42 24.11 11 .(14 I).28 2.70 9.23 1.24 2.59 0.33

4.81 40.18 17.91 12.54 2.86 4.40 7.71 1.82 2.67 0.29

11.65 57.18 8.86 11.11 1.10 1.18 2.39 9.6(I 4.26 0.33

1.33 0.24 56.11 12.45 6.17 6.64

7.7(I t).18 45.28 10.96 13.74 2.53

I I. 12 0.12 56.53 9.98 0.95 6.81

12.37 0.12 46.19 13.16 11.74 0.32

1.95 11.18 64.77 4.96 10.41 0.20

1.94 11.(16 36.46 27.28 9.43 10.84

2.37 1.26 28.28 27.06 12.47 18.37

6.32 3.63 3.03 0.36

4.49 3.24 2.84 0.48

5.94 3.69 3.56 0.57

1.72 8.32 3.56 (!.52

5.96

4.8(t

I).83 34.41 23.36 18.83 1.3i 9.56 5.77

1.54 (I.(19

1.76 (I.17

1.82 11.26

Cr Gu Ni Sr V Y Zn Zr

12 71 16 96 356 33 77 78

nd nd 23 175 416 37 69 81

37 47 36 182 13 36 88 91

nd 2 34 118 536 53 102 127

nd nd 29 55 292 52 51 145

SiO2 TiOe AIeO3 Fe20~ FeO MnO MgO CaO Na~O K,O

9.39 0.46 2.35 0.24

4.22 5.70 4.02 (1.46

4.85 5(/.28 16.56 2.46 4.63 2.58 7.82 2.81 3.62 0.38 1(1 6 16 97 271 48 126 141

nd 9 25 42 481 39 78 106

11 21 11 62 .~.6 "~'~ 50 69 154

nd nd 34 117 188 56 67 157

nd nd 33 58 284 59 59 185

1.32 45.62 17.61 9.114 6.47 5.112 7.21 1.17 2.(t3 11.29 72 27 48 145 278 27 50 57

0M5823" H

nd 52 4(1 116 253 20 56 43

0M5855 0 M 5 8 1 0 H H

69 77 58 156 25(I 22 104 54

116 74 58 71 261 24 65 57

=

The S e m a i l O p h i o l i t e

III

Table 3.20. Lasail Unit Basalts

Andesites

0M5807 H

0M5858 H

0M5782 F

0M4206" J

0M7016 Sr

SiO 2 TiO, AI20 3 Fe203 FeO MnO MgO CaO Na20 K20 P205 LOI

50.70 0.77 14.30 3.30 3.80 0.13 6.6O 10.42 4.40 0.02 0.05 4.60

47.90 0.83 14.60 4.47 3.49 0.16 7.41 13.13 2.29 0.01 0.11 4.00

48.60 0.73 13.90 3.76 3.78 0.16 9.37 10.06 3.68 0.02 0.11 4.65

47.00 0.48 15.90 5.08 3.26 0.15 5.92 10.96 1.75 1.11 nd 7.88

41.80 0.55 14.70 7.14 1.84 0.17 4.24 21.08 nd nd 0.13 6.89

Total

99.09

98.41

98.82

99.88

98.54

Q or ab an di hy ol mt il ap

0.12 37.57 19.39 25.74 2.32 3.80 4.83 1.48 0.12

Cr Cu Ni Sr V Y Zn Zr

137 32 55 58 238 19 47 39

2.50 0.06 19.69 30.01 27.86 7.37 6.59 1.60 0.26 212 nd 64 130 216 20 37 30

0.12 31.51 21.61 22.21 4.26 8.41 5.52 1.40 0.26 581 43 335 106 174 19 55 43

4.48 6.59 14.88 32.42 17.46 7.92

213 14 58 192 226 16 32 33

Felsites

0M4280 J

0M4203" J

0M4375 J

50.60 0,70 15.10 3.31 3.34 0.28 7.19 11.11 3.91 0.17 0.05 4.74

52.00 0.80 14.50 7.07 5.47 0.17 5.95 2.97 3.11 1.02 0.04 5.29

56.80 1.00 13.00 5.90 3.93 0.17 3.22 6.30 2.28 0.81 0.07 6.01

74.70 0.22 11.60 2.64 1.67 0.06 0.57 2.35 2.83 0.84 0.03 3.03

100.49

98.40

99.50

99.54

100.62

13.64 3.03 6.13 26.74 14.71 18.31

23.71 4.81 19.39 22.97 6.30 6.11

48.67 1.83 4.94 23.82 11.40 2.04

43.32 0.24 37.25 11.04 0.94 1.28

10.42 1.54 O. 10

8.60 1.91 0.17

3.81 (I.42 0.07

2.07 0.45 0.12

1.00 32.92 23.04 24.88 3.32 3.91 4.77 1.32 0.12

7.40 0.92 268 685 67 498 212 10 13 20

0M7072" Hw

241 nd 56 56 197 17 53 35

nd 6 5 210 435 24 74 45

nd 234 26 470 387 38 25 59

nd 83 10 163 nd 43 18 95

0M4348" J 75.50 0.24 11.40 2.87* 0.03 0.37 2.54 4.43 0.04 0.05 3.15

8 20 53 2 115

* Total Fe as Fe203

Table 3.21. Alley Unit. Basalts

A ndesites

Rh volites

0M5967" J

0M5746" S

0M5966 J

OM7010 M

0M7095 J

0M4385 J

0M5669" J

0M5667" J

0M4238 J

SiO2 Ti Or ' A1203 Fe203 FeO MnO MgO CaO Na,O K20 P205 LOI

45.00 0.72 15.70 5.43 3.33 0.15 7.45 8.63 2.73 0.09 nd 9.92

48.40 0.51 14.80 4.88 2.58 0.17 5.88 9.22 2.18 0.52 0. I0 9.14

52.00 0.80 13.60 4.45 2.78 0.16 4.23 7.76 2.77 0.66 0.08 9.28

55.30 0.88 14.90 7.03 2.41 0.13 4.08 4.20 6.6(I (I. 06 0.09 3.99

54.40 1.09 13.2(I 6.54 2.61 0.10 2.7(1 7.(13 2.27 (I. 38 (I. 14 7.89

56.10 0.93 14.2(I 5.59 2.69 0.11 3.35 6.4(/ 3.06 0.82 0.11 6.48

61.70 0.90 13.20 6.66 2.(11 0.15 2.22 2.27 7.18 0.07 0.10 2.55

67.6(1 0.2(1 10.50 1.46 2.27 0.09 0.24 3.47 2.76 0.7 l 0.04 9.24

79.2(I 0.2(I 8.3(I 2.58 0.27 0.05 (I. 76 //.96 4.28 0.41 0.03 1.28

Total

99.14

98.39

98.56

99.67

98.34

99.54

99.(11

98.57

98.32

Q or ab an di hy mt hm il ap

1.56 0.54 23.30 30.58 9.85 14.85 7.94

9.01 3.12 18.75 29.54 12.70 9.17 7.19

13.96 3.96 23.78 23.06 12.09 5.47 6.54

0.98 0.24

1.54 0.19

23.43 2.28 19.53 25.13 7.32 3.44 5.67 2.74 2.10 (I.34

18.62 4.85 25.93 22.63 6.58 5.31 6.34 1.22 1.77 (I.26

13.43 11.42 61.36 3.63 5.52 3.03 4.4(I 3.69 1.73 (I.24

41.(16 4.26 23.69 14.37 2.45 2.18 2.15

1.38

3.90 0.36 56.03 10.90 7.33 6.79 5.66 3.15 1.68 0.21

50.99 2.46 36.83 2.27 1.85 1.(17 (I.46 2.31 /I.39 0.07

Cr Cu Ni Sr V Y Zn Zr

83 48 54 41 301 24 59 46

8(I 55 44 190 280 22 59 27

nd 99 15 178 337 30 84 62

nd nd 37 81 347 34 51 65

nd nd 26 39 314 41 53 68

nd 11 26 196 11 45 57 93

nd 99 15 178 337 30 84 62

45 7 45 34 298 24 44 54

//.38 0.1(I

nd 12 12 21 nd 41 62 82

112

Chapter 3

Table 3.22. Cpx-O Unit Picritic basaits 0M173 R

0M5104 Hw

Table 3.23. Salahi Unit Basalts.

0M7041 F

0M7121" 0M7124 0M7135 Hw Hw Rq

0M8514" S

0M8515 S

0M8518 S

0M5759 S

SiO2 TiO2 A1203 Fe203 FeO MnO MgO CaO Na20 K20 PzOs LOI

51.04 0.18 10.00 1.58 6.35 0.16 15.57 10.59 0.55 0.06 0.00 3.81

51.13 0.44 15.78 2.29 6.26 0.20 8.58 10.27 2.21 0.37 0.05 2.84

51.90 0.20 13.90 2.52 5.24 0.16 10.10 10.27 1.20 0.07 0.04 3.34

52.80 0.44 15.00 2.62 5.34 0.17 7.72 7.31 3.35 0.52 0.04 3.74

50.00 0.37 15.60 2.07 5.20 0.16 9.35 10.58 1.58 0.31 0.05 3.79

48.50 0.62 18.80 2.18 4.97 0.08 7.29 13.77 1.84 0.01 0.06 2.14

Si02 TiO 2 Ai203 Fe203 FeO MnO MgO CaO Na20 K20 P205 LOI

48.20 1.80 14.50 6.97 5.96 0.23 4.33 6.05 5.53 1.03 0.46 3.54

49.00 1.93 14.40 7.64 5.74 0.23 4.90 5.04 5.38 0.74 0.40 3.71

49.10 1.76 14.40 6.27 4.29 0.19 5.82 6.66 5.08 1.02 0.28 4.01

48.60 1.57 14.40 5.79 5.87 0.19 5.28 8.00 4.12 0.98 0.27 3.35

Total

99.89

100.59

98.44

99.05

99.07

100.26

Total

98.59

99.11

98.89

98.42

Q or ab an di hy ol mt il ap

1.75 0.25 6.10 25.45 21.29 43.58

0.50 2.25 19.23 32.95 15.54 27.33

6.55 0.43 10.64 33.89 15.55 31.29

1.48 3.23 29.79 25.61 9.97 27.75

2.10 1.93 14.05 36.33 14.95 28.70

6.17 43.58 11.87 2.10 12.36

4.41 45.93 13.08

5.90 35.93 17.42

1.22 0.86 0.12

1.16 0.40 0.10

1.19 0.88 0.10

1.08 0.74 0.12

5.53 10.25 3.47 1.10

7.50 7.74 1.80 11.17 3.70 0.95

6.10 42.83 13.64 0.35 13.97

1.18 0.40

or ab an ne di hy ol mt il ap

Cr Cu Ni Sr V Y Zn Zr

1257 9 375 117 226 6 51 10

Ba Cr Cu Nb Ni Sr V Y Zn Zr

409 nd nd 37 9 128 319 45 46 189

245 nd nd 38 7 168 408 42 34 178

24 27

299 41 130 64 229 9 41 13

153 7 80 149 248 13 124 23

231 65 113 104 179 13 144 22

0.06 15.90 43.94 20.66 12.68 4.38 1.04 1.20 0.14 144 14 49 246 182 16 6 38

fractionation. On the TiO2-Zr plot (Fig. 3.56) they have a positive trend at relatively low TiO2 (0.6-1.5%) and Zr (50100 ppm) contents with a change in slope to a flatter and more dispersed trend between 100 and 200 ppm Zr with scattered TiO 2 values ranging up to 2.2%.

5.84 9.19 3.38 0.67 165 27 nd 27 nd 150 336 32 33 144

16.96 7.06 1.49 8.55 3.04 0.65 145 57 19 20 17 372 280 32 75 120

and the felsite sills are high-silica rhyolites (8i02 72.0-75.5%). On the FeOt/MgO versus Zr and TiO2 versus Zr plots (Figs. 3.57, 3.56) the Lasail Unit shows a fractionation series that differs markedly from that of the Geotimes Unit with trends towards high FeOt/MgO, low TiO2 and relatively low values of Zr (95-150 ppm) in the most differentiated rocks.

3.8.2.2 Lasail Unit

The Lasail basalts have suffered a different and more intense type of alteration than the Geotimes basalts; in particular, they have high C a O contents (up to 22.4%) which clearly reflects the high abundances of secondary epidote and prehnite in many of the rocks. The rocks with highest Ca contents have clearly suffered the greatest losses in SiO2, N a 2 0 and MgO and show m a r k e d increases in F e 2 O J F e O and LOI. Despite the effects of such alteration, it can still be identified that the Lasail basalts are less fractionated than the Geotimes basalts with distinctly lower F e O t M g O (<1.25), TiO2 (0.3-0.9%) and Zr (20-50 ppm) contents and they contain significant amounts of Ni ( - 5 0 - 7 0 ppm) and Cr ( - 2 0 0 - 3 0 0 ppm). These relatively unfractionated compositions correlate with the presence of pseudomorphs after olivine (absent in 'he Geotimes Unit) and the more magnesian nature of the clinopyroxene phenocrysts in the Lasail basalts. The associated inclined sheets are mainly basaltic andesites and andesites (SIO2 52-57%, M g O 6 - 3 % , FeOt/MgO 2.0-3.0)

3.8.2.3 Alley Unit

The Alley basalts often have high LOIs (up to 10%) which reflect the high abundances of secondary hydrous minerals, particularly zeolites, in many of the rocks. K20 contents are variable ( 0 - 7 . 2 % ) , the exceptionally high values being due to the presence of K-bearing celadonite in the groundmasses. The basalts exhibit a wide range of compositions (FeOt/MgO 1-3.5, Zr 20-65 ppm) and are part of a complete differentiation trend from basalt to high-silica rhyolite which follows similar paths to the Lasail Unit on the FeOt/MgO versus Zr and TiO2 versus Zr plots (Figs. 3.57, 3.56). 3.8.2.4 Cpx-phyric Unit

The Cpx-phyric Unit contains the most "primitive" or unffactionated fine grained rocks in the Semail ophiolite suite which have low FeOt/MgO (0.75-1.75), Zr ( < 4 0 ppm) and TiO2 ( < 0 . 5 % ) and generally high MgO, Cr and Ni contents (one

The Semail Ophiolite

II3

Table 3.24. REE analyses of dykes and lavas.

Sheeted dykes OM5994 OM7040 Geotimes Unit OM4279 OM5745 OM5756 OM5831 Transitional OM5823 I.asail Unit OM4206 OM7072 OM4203 OM4348 Alley Unit OM5746 OM5967 OM5669 OM5667 Cpx-O Unit OM7121 OM7135 Salahi Unit OM8514

La

Ce

Nd

Sm

Eu

Gd

Tb

Tm

Yb

Lu

Hf

Ta

Th

Sc

6.7 (20.4) 6.0 (18.3)

18.3 (21.2) 17.2 (19,9)

17.5 (27.8) 16.2 (25.7)

5.7 (28.1) 5.2 (25.6)

2.27 (29.5) 1.63 (21.2)

7.5 (27.2) 7.1 (25.7)

1.38 (26.5) 1.27 (24.4)

0.94 (27.6) 0.99 (29.1)

5.11 (23.2) 4.93 (22.4)

0.78 (22.9) 0.78 (22.9)

3.93

0.26

0.36

29.4

3.54

0.16

0.32

17.1

5.2 (15.9) 6.2 (18.9) 6.2 (18.9) 4.0 (12.2)

16.9 (19.5) 17.6 (20.3) 15.6 (18.0) 12.0 (14.0)

14.9 (23.7) 16.5 (26.2) 14.7 (23.3) 10.5 (16.7)

5.2 (25.6) 5.0 (24.6) 4.8 (23.6) 3.7 (18.2)

1.86 (24.2) 1.67 (21.7) 1.75 (22.7) 1.68 (21.8)

6.7 (24.3) 7.1 (25.7) 6.9 (25.0) na

1.16 (22.3) 1.23 (23.7) 1.21 (23.3) 0.92 (17.7)

0.80 (23.5) 0.92 (27.1) 0.83 (24.4) 0.56 (16.5)

5.13 (23.3) 4.80 (21.8) 4.64 (21.1) 3.56 (16.6)

0.81 (23.8) 0.75 (22.1) 0.72 (21.2) 0.62 (18.2)

3.87

0.32

0.48

28.7

3.79

0.22

0.38

23.9

3.54

0.20

0.34

29.5

2.56

0.12

nd

28.5

2.18 (6.65)

6.3 (7.28)

6.1 (9.68)

1.87 (9.21)

0.80 (10.4)

[1.72] (6.23)

0.49 (9.42)

[(/.31] (9.12)

2.05 (9.32)

0.34 (10.0)

1.37

0.10

nd

28.4

1.1 (3.35) 1.45 (4.42) 3.7 (11.3) 3.1 (9.45)

4.0 (4.62) 4.69 (5.42) 8.7 (10.1) 10.1 (11.7)

3.6 (5.71) 4.55 (7.22) 7.6 (12.1) 9.5 (15.1)

1.3 (6.40) 1.54 (7.59) 2.8 (13.8) 3.6 (17.7)

0.86 (11.2) 0.62 (8.05) 0.95 (12.3) 0.94 (12.2)

1.9 (6.88) na

0.31 (5.96) 0.40 (7.69) 0.72 (13.8) 0.91 (17.5)

0.20 (5.88) [0.26] (7.65) 0.56 (16.5) 0.81 (23.8)

1.23 (5.59) 1.59 (7.23) 3.41 (15.5) 5.33 (24.2)

0.23 (6.76) 0.28 (8.23) 0.58 (17.1) 0.92 (27.1)

0.90

0.03

0.17

26.4

1. l 1

0.05

nd

35.3

2.19

0.14

0.32

27.1

3.25

0.16

0.34

11.2

1.7 (5.18) 2.5 (7.62) 2.3 (7.01) 3.4 (10.4)

3.0 (3.47) 5.8 (6.71) 7.0 (8.09) 9.4 (10.9)

2.8 (4.44) 6.9 (10.9) 6.5 (10.3) 8.9 (14.1)

1.1 (5.42) 1.9 (9.36) 2.6 (12.6) 3.4 (16.7)

0.48 (6.23) 0.68 (8.83) 0.92 (11.9) 0.98 (12.7)

0.37 (7.11) 0.53 (10.2) 0.69 (13.3) 0.91 (17.5)

0.36 (10.6) 0.42 (12.4) 0.51 (15.0) 0.74 (21.8)

1.77 (8.05) 2.50 (11.4) 3.57 (16.2) 4.97 (22.6)

0.29 (8.53) 0.43 (12.6) 0.58 (17.1) 0.82 (24.1)

0.73

0.02

0.12

44.8

1.45

0.08

na

32.4

2.23

0.11

0.28

28.7

3.04

0.15

0.33

14.6

0.6 (1.83) 1.2 (3.66)

1.5 (1.73) 4.4 (5.09)

2.0 (3.17) 4.5 (7.14)

0.8 (3.94) 1.3 (6.40)

0.33 (4.29) 0.65 (8.44)

0.63

[0.02] [0.08] 43.1

0.92

0.03

[0.12] 37.8

3.86

1.38

3 . 7 1 24.1

26.4 (80.5)

50.9 (58.8)

26.8 (42.5)

5.9 (29.1)

1.98 (25.7)

3.3 (12.0) 4.8 (17.4) 1.8 (6.52) na 3.7 (13.4) 4.8 (17.4) 1.2 (4.35) 1.7 (6.16) 6.6 (23.9)

0.28 (5.38) 0.37 (7.12)

0.22 (6.47) 0.21 (6.18)

1.25 (5.68) 1.27 (5.77)

0.19 (5.59) 0.21 (6.18)

1.07 (20.6)

0.56 (16.5)

3.57 (16.2)

0.55 (16.2)

( ) Chondrite normalized values. [ ] Analysis unreliable. olivine-phyric dyke contains 15.78% M g O , 886 p p m Cr and 333 ppm Ni (Smewing 1980a), but most contain 7 - 1 0 % M g O , 130-3 p p m Cr and 50-130 p p m Ni). Most of the analysed samples (Table 3.22) are dykes but the lavas that have been identified are petrographically identical and there is reason to believe that they have the same range of chemical compositions. 3.8.2.5 Salahi Unit

The Salahi lavas are relatively fractionated basalts ( F e O t / M g O 1.5-2.75), with low to negligible Ni and Cr contents that are geochemically distinct from the rest of the Semail lavas. The high L I L element (K, Ba, Rb) contents, that are clearly not entirely the result of secondary enrichment, plus relatively high Ta, Nb and L R E E ' s are characteristic of this unit which, with Nb/Y ratios of 0.6--0.9, indicate a mildly alkaline m a g m a type. M a j o r elements, even though showing some alteration effects, conform to the pattern and some of the rocks are

slightly ne normative. On the TiO2 versus Zr and TiO2 versus F e O t / ( F e O + M g O ) plots (Figs. 3.56 & 3.57) they show a positive trend towards high TiO2 ( > 2 % ) and Zr ( > 2 0 0 ppm) values with very little change in fractionation index. 3.8.3 Tectonic setting

The tectonic setting of the Semail ophiolite, as deduced from the lava geochemistry, has been discussed by Pearce (1980), Pearce et al. (1981), A l a b a s t e r (1982) and A l a b a s t e r et al. (1982). On the TiO2 versus F e O t / M g O plot (Fig. 3.58) the basic lavas and dykes from the Semail extrusive sequence cover a b r o a d swathe of compositions b o u n d e d by the " A b y s sal Tholeiite" and the "Island A r c Tholeiite" trends as defined by Miyashiro (1975). Most of the scatter is p r o d u c e d by the sheeted dykes and G e o t i m e s Unit basalts, whereas the Lasail and Alley basalts show a low Ti trend which is closer to those of the island arc tholeiite series. The various basalt types ( Ge o t i m e s , Lasail, Alley, Cpx-

Chapter 3

I 14 (a)

9

2.0 -

9

..

s

9

9

9

(a)~; r ~

~!

1.5z,,/~a)9

9

1.0-,/"Y //

9

(&

"~

9

:~ : . 9

t2

9

~.)

./.-.)

~: ', ....

vJ> ( ~ (';.) 2

9

9

) (~) 9

9

(&)A'(zX/X~ " " ~ J ( ~ )" \ A zz~'~&) (A'),,\

9 Geotimes lavas " Sheeted dykes ;/.) Sheeted dykes (data from Pearce

/~

05

%)

1.0

et al

Cpx-phyric basalts 0 8asalts Andesites Felsites rhyolites

1981)

O5

~:'0

', \ \

~

~

Lasail and Alley trends

0 1

0

1

I

I

1

50

100

150

200

0 t

,

I

I

50

100

,.,J

150

Zr (ppm)

Zr (ppm) (d) 20-

~D 15

(b)

I.O[9

9 Lasail basalts ~ Andesite inclined sheets ~_Felsite sills

9 ~ 9

0.5--

1.0

.'T 0.

0

I

1

I

50

100

150

Zr

(ppm)

0

l

I

I

I

50

100

150

200

Zr (ppm)

Fig. 3.56. TiO2 vs Zr plots for the different lava units. (a) Geotimes Unit and sheeted dykes, (b) Lasail Unit, (c) Alley and Cpx-O Units and (d) Salahi Unit. phyric and Salahi) can be compared to one another and with standard magma types from different tectonic settings by using geochemical pattern or "spider" diagrams (Fig. 3.59) where a whole range of elements with varying geochemical characters are plotted by normalizing against a standard composition. Following Pearce (1980) we have used his "N-type" MORB (mid-ocean ridge basalt) as the standard. A feature of these plots is that modern subduction-related calc-alkaline and tholeiite arc basalts (CAB and IAT types) can be distinguished from most other magma types by two features; first, their high abundances of LIL elements (Sr-Th on Fig. 3.59) relative to the HFS elements (Ta-Ti) and second, their exceptionally low contents of Nb and Ta (Pearce 1975, 1980; Saunders & Tarney 1979). Unfortunately, the LIL elements are those most likely to be affected by alteration so that the first discriminant cannot

always be confidently applied to altered rocks. Within-plate basalts (WPB) can usually be distinguished by an enrichment of all the incompatible elements (both LIL and HFS) relative to MORB with high contents of Nb and Ta being particularly characteristic. These features are shown by the Salahi basalts (Fig. 3.59) in which Nb is enriched 8-10x MORB and over 30x its value in some of the other Semail basalt types. For this reason, Alabaster et al. (1982) suggested that the Salahi basalts formed during continent-arc collision by melting of enriched mantle that underlay the continental margin. The Lasail, Alley and Cpx-phyric basalts all have similar and typical island-arc tholeiite (IAT) patterns with especially low Nb and Ta contents and with all the elements Ta-Yb well below MORB values (Fig. 3.59). The Geotimes basalts have similar shaped patterns but, because of their more fractionated

The Semail Ophiolite X Sheeted dykes

Basalts, basak~c andesltes

o Geolimes Unit

Basalts basalbc andesdes, andesltes

§ Transitional Unit

r

[

Basalts I Andes~tes i Dac~tes, rhyohtes

[ Basalts i Andes,ies i Daotes. rhyohlos ] PIcrlte basalts

Alley Unit Alley Unit Umt Cpx- ~ Umt

§ AHey

8

I

i

f

u Sa~ahl Umt

FeO t MgO

lous parts of the mid-ocean ridges (e.g. Iceland), and may occur along some fracture zones. Alkali basalt sills have also been drilled on the floors of some back-arc basins; e.g. the Skikoku and Minami-Daito basins in the W. Pacific (Marsh et al. 1980; Dick 1982) and in the Caribbean (Donnelly et al. 1973). They are usually intruded "late", often into pelagic sediments above the igneous basement, although interdigitations of alkaline and tholeiitic basalts occur (Dick op. cir.). Geochemically these back-arc basin basalts have the same "within-plate" compositions as the Salahi basalts.

Basalts

r o Lasail Umt Lasatl UnFt Lasad Unit

II5

i Basalts

3.8.4 Clinopyroxene compositions §

As pointed out by Nisbet & Pearce (1977) clinopyroxene is often the only primary igneous phase to be preserved in altered and metamorphosed basalts and they provide several chemical criteria by which clinopyroxenes from different tectonic settings could be recognized. Clinopyroxenes are preserved in all the Semail lava units and some of the dykes and show significant differences between different units that have important petrogenetic implications (Alabaster 1982; Alabaster et al. 1982). The least fractionated clinopyroxenes, diopsides with up to 1.5% Cr203, occur in the Lasail basalts and in the Cpx-phyric lavas and dykes, whereas the Geotimes and Alley basalts contain augites (Fig. 3.60). The pyroxenes in the Salahi basalts are distinct and show an alkaline trend towards salite compositions and have the highest TiO2 (0.9-2.6%) and Na20 (0.30.5%) contents. Despite the almost complete overlap on the En-Fs-Di-Hd plot (Fig. 3.60), the clinopyroxenes in the Alley and Cpx-phyric units can be distinguished from those in the Geotimes and Lasail units by lower Na and Ti contents (Fig. 3.60). This difference, perhaps highlighted by the greater precision of fresh mineral over whole rock analyses in altered rocks, suggests that there is an important geochemical difference between the lower (Geotimes and Lasail) and the upper (Alley and Cpx-phyric) lavas. This distinction is reinforced by the different REE patterns shown by these two groups discussed in the next section.

r

D

. 2

v 0

V

Av v I

vX

X~

~

"

C}

X X (')

X

LJ

O

1

I

50

100

Zr (ppm)

I

I

150

200

Fig. 3.57. FeOt/MgO vs Zr plot for the Semail extrusives.

nature, are closer to MORB values for most HFS elements except Ta and Nb. They are closest to some marginal basin basalts that show transitional behaviour between IAT and MORB magmas (Saunders & Tarney 1979; Saunders 1983). On the basis of these data, Alabaster (1982) and Alabaster et al. (1982) proposed that all the Semail lavas, with the exception of the last erupted Salahi Unit, show a "subduction component" and suggested that the ophiolite was formed above a subduction zone, probably in a marginal basin setting (Pearce et al. 1981). Alkaline magmatism in the ocean basins is normally associated with within-plate ocean islands (e.g. Hawaii) or anoma[ ] Salahi Unit

~

Lasail Unit a n d e s i t e s

Cpx-~Unit

9

Lasail Unit basalts

A l l e y Unit a n d e s i t e s

|

T r a n s i t i o n a l Unit

A l l e y Unit basalts

O G e o t i m e s Unit 9

Sheeted dykes

AT

9 _

9

/

TiO2 (wt %)

/ /

~

"_..,

,-

o,--.

A_

o

9

,,-

o ,, ~

-% To

1 o

0

1

1

1

2

I

FeOt/MgO

3

i

4

Fig. 3.58. TiO, vs FeOt/MgO plot for the Semail extrusives and dykes. Trend lines (after Mivashiro 1975) are for abyssal tholeiite (AT) and three intra-oceanic island arcs (K - - Kyushu, To -- Tofua, I-B -- Izu-Bonin).

[ [

6

Chapter 3 (c) 1

0.5

0.5

( 0.25

0.25

Th Ta

Ce

P

Zr

,,

,,,,

Hf Sm Ti Y Yb

,

Sc Co Cr

I ~2,

,~,

, ,,

Th ~' Ce P Zr Hf Sm Ti la

,,,

Y Yb Sc Co

, Cr

(d)

Fig. 3.59. "Spider diagrams" (normalized against N-type MORB of Pearce 1980) for the Semail extrusives. (a) Geotimes (open circles) and transitional basalts, (b) Lasail basalts, (c) Alley (open diamonds) and Cpx-o basalts and (d) Salahi basalts.

0.25

1

Wo5o En;,. Wo~oEn5o

/..O-

,~

A U-G/-TE X X ~ Wo, En ,, Fs4,

0.04 F

004 F ~ " ~

~

/

.4

/ .' / / ' / / ,

/

~

It.

70

Salahi

:-~

80 100 Mg/Mg+Fe

""4

-~o.02t:~ F..

"

L__+--.-J~

I ................

,

|

90

,

70

~

on, ,

.~'.

".~."

..............,...--.....~

80 100 Mg / Mg+Fe

90

o o3

('.~'.) Alley Cpx-~Unit Unit ::ii'..".ii:'

o' 5~ aa I

f-- 2

Lasail Unit

~ 0.01 .........

~

Geotimes

--

",

~ X " ..... r--~. \

Unit

,o.,,...a,

1

1

1

I

I

1

I

I

I Cr

i

J

Th Ta

l Nb

I Ce

I P

~ Zr

I

I

J

J

J

Hf Sm Ti y Yb

I I Sc Co

3.8.5 Rare-earth elements

Fs~oo

/ 3o o.O2F

1

Wo:, En ~,, Fs,,

Wo~,,En:,

o

1

WosoFsso

Emoo

-~

I

Th Ta Ce P Zr Hf Sm Ti y Yb Sc Co

"................ ."i :~-- ~i'.~

Unit

70

80 100 Mg / Mg+Fe

90

Fig. 3.60. Clinopyroxene compositions from the Semai[ extrusives.

The REE abundances, normalized against chondritic or some other standard values, are a widely used and useful way to examine petrogenetic relationships between and within suites of igneous rocks, particularly with regard to partial melting and fractional crystallization modelling (Hanson 1980). The Geotimes basalts and basaltic andesites and most of the sheeted dykes have fiat to concave-upward REE patterns with relatively uniform values from Nd to Lu but are slightly depleted in Ce and La (Fig. 3.61a). These patterns are similar to those of the high-level gabbros and ferrogabbros described in Section 3.5 and support a cogenetic relationship between the two groups. The overall REE abundances in the Geotimes basalts range from 10x to 25x chondrite and they have (La/ Yb)N (chondrite-normalized ratio of La over Yb) of 0.63 to 0.97. The Lasail basalts show similar patterns (Fig. 3.61b) but at lower REE abundances (5-10x chondrite) and with slightly greater LREE depletions (La/Yb(N) 0.56-0.61). One of the samples has a positive Eu anomaly probably due to a high abundance of secondary epidote. The Geotimes and Lasail basalts are sufficiently similar geochemically that they could be related by fractional crystallization and magma mixing from a

The Semail Ophiolite (9

3~r 92

1200 14.9 0.46

68 5400 3.60 0.43 94 1200 5.44 0.64 62 6540 3.15 0.58 46 4320 1.10 0.67 65 5580 230 0.75 53 348O 1 92 0.63

(a)

Zr 137 165 145 153 154 141 81 83

3O

2O

Ti

67 6340 10

I

1

5La Ce

Nd

[

I

I

I

I

Sm EuGd Tb

I

27

~

9840 2.99 0.81 8160 2.44 0.97 12540 2.85 0.68 13440 2.85 0.88 8820 2.86 0.87 11280 2.97 090 8100 2.98 0.74 7380 1.54 0.72

3060 1.19 0.64

10 1620 1.04 0.05 38 3720 0.95 0.62 23 2640 1.00 0.32

2.60 0.71 2.06 0.78

65

6120

43

29 4620

4860

1.47 0.71 1.40 0.66

54

5640

123

1.0

I

Tm Yb Lu

La Ce

Nd

Sm Eu Gd Tb

Tm Yb Lu

(b) 30-

15 1440 6.98 0.39

20 --

~ ~

t

-

~

~

2

:

:

:

~

~

(d) 100-

~ 59 6000 2.87 0.73 72 5800 261 066

~

10' ~

57 5700 1.17 0.63

/"

8.

/Jl._\

/- ~ ~ _ _ , ~ _ . ~

/

~0~

J l 35 4200 0.68 0.61

~ 37~0 065 OOl

~

20 2880 1 32 0.61

,0o283,9, i

20

9540 2.50 5.79

I

La Ce

1

Nd

I

I

I

I

Sm EuGd Tb

I

I

I

Tm Yb Lu

I 10 La Ce

I

Nd

1

I

I

I

Sm Eu Gd Tb

1

I

I

Tm Yb Lu

Fig. 3.61. REE/chondrite plots for the Semail extrusives. (a) Geotimes Unit basalts (open circles) and andesites (closed circles) and transitional basalts. (b) Lasail Unit basalt (closed circles), andesites (half-closed circles) and felsite. (c) Alley Unit basalts (open diamonds), andesites and rhyolites (half-open diamonds) and Cpx-0 Unit basalts (closed diamonds) and (d) Salahi Unit basalts. common parent. This is supported by the occurrence of "transitional" lavas that are both in the field and geochemically, intermediate between the two groups (Alabaster 1982). The andesite cone sheets and lavas in the Lasail Unit have REE patterns parallel with the basalts but at higher concentrations (10 to 15 • chondrite) and can be modelled by about 50% fractionation of cpx and plagioclase from the basalts. The basalts from the Alley and Cpx-phyric units show distinctly different REE patterns (Fig. 3.61c) to the Lasail basalts with more marked LREE depletions (La/Yb(N) = 0.32-0.75) at similar total REE abundances. The chondrite-normalized patterns have more uniform slopes with constant decreases in values for Lu to La with the result that they have HREE/ LREE ratios (e.g. Sm/Yb(N)) <1 compared to >1 for the Lasail basalts which gives rise to the more pronounced "dogleg" pattern of the latter. The Cpx-phyric Unit picritic dykes and lavas are the most depleted rocks with La and Ce contents close to or even below chondritic values. (These REE/chondrite patterns are rather irregular largely due to the problems of analysing such low concentrations of these elements by the

INAA method). Modelling of REEs and other trace elements in the Alley differentiation suite shows that the andesites can be produced by 40% to 60% basaltic (olo.2cpxo.4plago.4) fractionation from the associated basalts and that the dacites and rhyolites, which have marked negative Eu anomalies in the most differentiated compositions, can be modelled by a further 30% fractionation of cpxo.splago.45Fe-Ti oxideso.1 from the andesites. As already noted, the Salahi basalts are completely different geochemically from the rest of the Semail lava suite and this is confirmed by their markedly LREE-enriched (La/Yb(N) = 56) patterns (Fig. 3.61d) with convex-upward shapes typical of a|kalic basalts (Kay & Gast 1973).

3.9 Petrogenesis The overwhelming abundance of basic and ultrabasic rocks, the high temperature (c. IO00~ deformation of the harzburgites, the sheeted nature of the dyke complex, the preponder-

II8

Chapter 3

FIZH BLOCK

I~ll; 9

po

"'..

Jllal ill". I".I I"'.I I J"..i l l"-.l "x'-" Ill, \lJillI, U,~ / , ~ / ! l\/,:~'-~-~'/ l ~ l j l l l l , I .." ; l.. l l l.-l l l." l i,l l l ,l l l 9l l l.. 9 .' I9 I l, l l l,"l '-..~.~--,-, iiiii!i!!i! .-.'i.-'.2.'.-'..".--'...'" ~ ;- "- "2-"-".."."-".'-. -"-,"." ,'."-"-:" 9

--'.

"

~

..

A

20km I

V///A

Crust

Supra solidus mantle

Depleted mantle

Lherzolite solidus ............. Trace of Iherzolite solidus

Boundary between fully and partly depleted mantle a~ Petrological moho

.

~

~

Undepleted '",,, mantle

(D

_c: s {/3 oc.~ J

t

' '\\\1

Asthenosphere

Fig. 3.62. Postulated spatial distribution of supra-solidus mantle in a section normal to the spreading axis. In this model the highest temperature would be reached in the central part (a"-b") of the rising limb and here extensive partial melting would leave a fully residual harzburgite. Temperature would decrease progressively from a" to a' and b" to b'. It is considered that the basal part of the Fizh Block Mantle Sequence represents a'-a" type mantle. Inset shows model of Whitehead et al. (1984) in which viscous asthenosphere rises between two spreading lithospheric plates that thicken away from the boundary. Above a certain level the asthenosphere passes through a zone in which the partial melt forms and collects. Gravitational instability leads to regularly spaced concentrations of melt which rise diapirically to form crustal magma chambers. ance of pillowed lava flows and the absence of intra-ophiolite terrestrial sediments have all been used in support of the thesis that ophiolites are fragments of oceanic lithosphere formed at constructive plate margins (e.g. Gass 1977). The Semail Nappe has all these characteristics, in most cases better preserved than in other ophiolites. On the broadest scale, it is believed that mantle convection currents rise and diverge beneath constructive plate margins (Oxburgh & Turcotte 1968; Nicolas & Rabinowicz 1984) and that magmas produced by the partial melting of the convecting mantle move upwards to form the oceanic crust. Here we use data from our study of the Semail nappe in an attempt to define more closely the physical parameters and style of mantle convection and the magmatic processes that occur within and because of it. Whilst a constructive margin is active,

subjacent convection is a continual process. Therefore the model we are investigating has to fit with a convective continuum in which the magmas are in a passive responsive role moving as space is created by the convective driving forces. There is no evidence, for instance, that the individual dykes in the Sheeted Dyke Complex were forcefully injected. Our Semail data can only add to the uppermost 10 km or so near the axial part of any convective model. Furthermore, although the Mantle Sequence is the lowest unit in a nearhorizontal ophiolite stratigraphy, the deformation and partial melt processes recorded therein were "frozen in" as the mantle passed through its solidus on moving away from the constructive plate margin (Figs. 3.62 & 3.66) and therefore can be used to infer what happened during sub-spreading axis mantle ascent.

The Semail Ophiolite 3.9.1 Convective processes

In the Fizh Block, where the thickest Mantle Sequence is preserved, it has been noted (Section 3.2.1.1) that beyond a depth of c. 10 km below the Moho, the previously uniformly harzburgitic composition becomes progressively more lherzolitic. This change in mantle composition with depth has been identified in other ophiolites (e.g. Newfoundland, Malpas, 1978). Although there are other possibilities such as subsolidus re-equilibration and metasomatic modification, the most realistic explanation is that the mantle is becoming progressively less refractory with "depth" (horizontal distance from the axis of the rising plume) and that a smaller percentage of melt was extracted therefrom. Even the least refractory specimens from the Fizh Block are depleted compared to pristine mantle lherzolite, indicating that some partial melting and melt extraction had occurred. Projection of the Fizh composition depth curve suggests that pristine mantle should be encountered at c. 15 km depth below the Moho. This field based calculation does not identify the percentage melt fraction with depth relation. Indeed, this relation could well be complex for near a hot axial constructive margin the movement of depressurizing melt towards the surface would tend to leach the lherzolite wall rock of fusible components such as clinopyroxene and spinel and thereby invalidate the calculation of percentage melt from geochemical data. Two main models have been proposed to explain the kinematic behaviour of uppermost mantle beneath constructive margins. The "dyke-intrusion" model of McKenzie (1967) and Cann (1970, 1974) envisages that melt within a narrow vertical filament of partially molten mantle feeds into a high level, crustal magma chamber whilst the refracting mantle is plated onto the lithospheric hanging walls. The "right-angle turn" model of Langseth et al. (1966) invoked the vertical upwelling of asthenospheric mantle beneath a constructive margin. The right-angle turn of the flow to a horizontal orientation occurs shortly before the mantle passes through the asthenospherelithosphere boundary. Both models are entirely plausible and either could produce the observed melt fraction-depth relations. Studies of Mantle Sequence foliations and lineations led Bartholomew (1983) to conclude that the mantle does not convect by steady uniform flow normal to the ridge axis but as discrete diapiric masses (see Fig. 3.15). This conclusion is in keeping with structural data from other ophiolite mantle sequences (e.g. Juteau et al. 1977; Nicolas & Rabinowicz 1984). Bartholomew (op. cit.) does not indicate how large these mantle diapirs were. He does however state that foliation and lineation orientations vary markedly from one wadi section to another. And, as the wadi sections are mainly eastwest and some 10 km apart, it seems likely that the mantle diapirs were in the order of 10, rather than 1-2 km diameter. Expanding Bartholomew's model, we suggest that beneath a spreading axis the mantle rises, at least for the last few 10's of kilometers, in a diapiric fashion. Diapirism along the axis is intermittent in both time and space so that one part of the axis is active whilst another is dormant. The model envisaged is very similar to that of Whitehead et al. (1984) but instead of magma bodies rising diapirically (inset Fig. 3.62), we envisage diapiric masses of asthenospheric mantle. Within these mantle diapirs the initial small melt percentage will increase during adiabatic decompression as the diapir rises and discrete melt bodies, themselves diapiric, will form. For the west side of a N-

II9

S constructive margin (see Section 4.3.3) the "dyke-injection" model should produce a tectonic fabric with S-folds. Conversely, in the same situation, the right-angle turn model would produce Z-folds. Although searched for in the Oman Mantle Sequence (Bartholomew op. cit.), no unequivocal fold closures were found. 3.9.2 Partial melting processes

It can be confidently assumed that the mantle, from which the picritic parent* magmas which gave rise to the Crustal Sequence of the Semail Ophiolite were ultimately derived, was of Iherzolitic composition. We have noted earlier (Section 3.2) that the Semail Mantle Sequence consists predominantly of highly refractory cpx-poor harzburgites which have uniformly depleted compositions (TiO2 - 0 % , A1203 < 1 . 0 % , CaO < 1 . 0 % ) down to depths of 8-10 km below the Petrological Moho. Below this, in the lowermost 1-2 km of the exposed sequence, there are less refractory cpx-harzburgites and lherzolites (TiO2 < 0 . 0 5 % , A1203 < 2 . 1 6 % , CaO <2.36%) in the Basal Lherzolite (Section 3.2.1.1). However, these rocks are still depleted in basaltic components compared to "fertile" mantle compositions and are only capable of producing c. 10% of basaltic and c. 15% of picritic melts (Table 3.25). Table 3.25. Least squares solution for partial melting of Semail Iherzolite OM9013 to produce the Primary Magma picrite leaving the average harzburgite as residuum. P.M.

SiO~ TiO~ A1203 FeO MgO CaO Cr20~ NiO Quantity -

Picrite

Harzburgite

Lherzolite

Productg

50.8 0.6 13.15 7.5 14.2 10.8 0.15 0.04 11.13

40.5 0.0 1.(19 8.38 44.0 0.98 0.4 0.3 89.01

44.8 0.4 2.36 8.18 41.2 2.58 0.37 0.28

45.26 0.55 2.43 8.29 40.75 2.07 0.37 0.27 100.1

All compositions are normalized to 100%. Sum of squares of residuals = 0.69, the residuals are the differences between the calculated product+ and the actual target (lherzolite).

In order to model the partial melting process, educated guesses must be made of the compositions of both the mantle and the primary* magmas produced within it. High pressure experimental studies (Green 1973; Green et al. 1979; Stolper 1980) suggest that picritic liquids, with 15 to 18% MgO, are in equilibrium with spinel lherzolite mantle at 15-20 kb and are capable, upon extensive low pressure fractionation, of producing MORB-like magmas. High SiO2 picrites, that are probably parental to IAT magmas, can be produced at the same depths but by the hydrous melting of a more refractory lherzolite (Green et al. op. cit.).

* We identify a primary magma as being an equilibrium partial melt of mantle peridotite. A parent magma is considered to be parental through magmatic processes to a cogenetic suite of igneous rocks. A primitive magma is used in a relative sense for magmas having high liquidus temperatures and compositions close to that of the parent magma.

[zo

Chapter 3 A

"Fertile" mantle Iherzolites

10

e~,~~'-

sp

20 Kbar

..~30

\ ~plaag/

1bar

'~20

-\

P

Primary mantle (• 2.5 chondrite)

1

Parental mantle to Lasail basalts (4% melt extracted from 1 )

2

Parental mantle to Alley basalts (model A) (4% melt extracted from 2 )

3

Parental mantle to Alley basalts (model B) (95% harzburgite + 5% basalt (see text))

Lasail Unit 10--...

Unit

/ ~10

,'15 /

Semail harzburgite (OM2265, OM1499)

Ce N cs~V

L

50 CMS2 DIOPSIDE

l

4O

L

I

3O

I

20

~\

10

,k 2

MS

ENSTATITE

Fig. 3.63. Projection from olivine onto the CS-MS-A plane of the CMAS tetrahedron (O'Hara 1968) showing the compositions of the Semail harzburgites (dots), lherzolite (open circle) and Browning's primary magma composition (triangle). 1 bar and 20 kbar phase boundaries shown and projected fractional path of the Semail magmas (dashed arrowed line) (after Browning 1982).

5000

0.1 Partial melts 10, 15, 20, 25% melting

~_~)

9 Residues 20, 25, 30% melting

15(11)

0.01 DEPLETED MANTLE

Lherzolite 1

T499

Lherzolite 2 Partial melts 10, 15, 20, 25% melting

(1.4)3~_~

| Residues 15, 20, 25% melting

25(4.0) +(2.8)10

PRIMARY MANTLE

Lherzolite 3 -~-Partial melt 10% melting -~ Residue

(For residues % cpx given in brackets)

1000

\ \ - X \ \ Mantle and Crustal \ \ . ~,~, Fraetionation "~

E {3. Q.

0.1

L,px-~ Unit

~[

\ ~.

I,

".t'\

Lasail Unit

x_

z

100 .......

x

\i: \ "

Alley Unit

\ ~ I ~ Ge~ "~. o~ Unit \ ~ ~'.

\\ 10

\

-

1

10

I

I

1

10

Fig. 3.65. Partial melting modelling in terms of CeN and YbN (see text for details). Fields of Lasail and Alley Unit basalts shown for reference.

\\.

v

o

YbN

1

Y(ppm)

50

Fig. 3.64. Cr vs Y log plot for the Semail lavas. Lower lavas (Geotimes, Transitional and Lasail Units) shown by Solid symbols, upper lavas (Alley and Cpx-O Units) by open symbols. See text and Table 3.27 for details of the partial melting modelling of the depleted mantle compositions. The plot suggests that Geotimes and Lasail magmas were derived by 25-30% partial melting of a slightly depleted lherzolitic mantle whereas the Alley and Cpx-r Unit melts must derive from a more depleted mantle unless unacceptably high degrees of partial melting are invoked.

Browning (1982) and Brown (1982) calculated the compositions of Semail "parent magmas" based on mass-balance calculations of the average crustal section through the ophiolite. The resulting compositions have values of Mg* that are too low to be in equilibrium with olivine Fogo (the composition in the basal ultramafic cumulates). This deficiency can be corrected by adding an appropriate amount of olivine (+ chrome spinel) equivalent to the dunite bodies in the Mantle Sequence (these are interpreted as the precipitates from the rising picrite magmas whose remaining melt fraction escaped into the overlying magma chamber (Section 3.2.2)). Browrfing (op. cit.) only allowed 100 m of dunite (equivalent to 1.5% of the crustal section), whereas Brown (op. cit.) added 350 m (6.7% of the crustal section). In view of the fact that dunites comprise between 5 and 10% by volume of the Mantle Sequence down to a depth of at least 10 km (Section 3.2.2), we consider that at least 500 m equivalent thickness of dunite should be added. Using this minimum value of 500 m results in a primary magma composition with MgO - 1 5 . 5 % , Mg* = 0.78, Cr = 1200 ppm and Ni = 350 ppm. These primitive parent magmas contrast with the basaltic composition " S A V E " (Semail AVErage) of Pallister (1984) which derived from a mass-balance calculation of the Jebel Dimh section in the Ibra area (Table 3.26). Browning (1982) presented a major element mixing calculation which showed that 17% melting of a spinel lherzolite (a particular sample from the Bay of Islands ophiolite was used)

The Semail Ophiolite

I2I

Table 3.26. "Primitive" and "Parent" Magma Compositions. Lasail basalts 0M5984

SiO, TiO, A1203 FeO* MnO MgO CaO Na,O K20 P20 s rag* Zr Y Cr Ni

50.24 0.59 16.64 8.45 0.17 9.02 11.07 3.55 0.23 0.04 65.6 35 17 110 64

()y.t-O basalts 0M5782

51.80 0.78 14.80 7.64 0.17 9.99 10.72 3.92 0.02 0.12 70.0 43 19 581 335

0M5985

0M7131

0M7124

49.95 0.52 11.74 8.03 I).26 15.45 11.94 1.79 0.00 0.04 77.4 22 13 450 314

51.57 0.65 15.91 7.68 0.10 9.08 12.30 2.58 0.37 0.05 67.8 39 20 90 58

50.00 0.37 15.60 7.06 0.16 9.35 10.58 1.58 0.31 0.05 70.3 22 13 231 113

Semail primary magmas Browning 11982

SiO, TiO~ A1203 FeO* MnO MgO CaO Na20 K20 P20 5 mg* Zr Y Cr Ni

51.99 0.62 13.46 7.14 (I.13 13.46 10.87 2.16 0.09 77.0

Brown (1982)

50.8 0.60 13.15 7.54 0.13 14.23 10.80 2.48 0.09 0.05 77.1 34 16 1069 290

0M173

50.90 0.20 10.10 7.69 0.15 15.78 10.04 0.69 0.03 0.00 78.5 10 6 1257 375

Primary MORB's

"5"AVE'" Pallister r

51.1 0.60 16.60 7.2 0.12 9.2 12.8 2.3 0.12 0.05 69.5

Stolper 11980)

Green et al. + 11979)

46.86 0.72 14.59 10.30 0.12 15.98 9.29 1.94 0.05

48.3 0.6 13.7 7.9 0.12 16.7 10.9 1.65 0.01

73.4

79.0

* All analyses recalculated to 100% anhydrous with Fe as FeO. rag* = 100 Mg/(Mg+Fe 2+) tool prop (after adjusting Fe3+/(Fe3++Fe 2+) to 0.1). t 83% DSDP glass + 17% olivine.

could produce his Semail primary magma leaving a harzburgite residuum. However, just as in the case of the Semail lherzolite (OM9013), the Bay of Islands specimen has probably been partly depleted by melting. We therefore believe it more likely that the primary magmas were produced by 20-30% melting of more fertile lherzolite mantle in accordance with the views of O ' H a r a (1968, 1970), Frey et al. (1978) and Green et al. (1979). All the parental magma compositions that have been calculated for the Semail Ophiolite have high 8i02, low TiO2 and high CaOAI/203 ratios compared to parental magmas proposed for M O R B (Table 3.26, Fig. 3.63). Browning (1982), using the experimental evidence cited by Ford (1976), Green (1973) and Green et al. (1979), suggested that these compositions are the results of high degrees of melting, probably at elevated water pressures, of a previously depleted lherzolite. The high CaO/AI203 ratio (>0.65, probably c. 0.85) of the magma, which will not be significantly affected by olivine fractionation during its ascent through the mantle, will allow pyroxenes to crystallize before plagioclase (Fig. 3.63). We have already seen (Section 3.8) that the lower (Geotimes and Lasail Units) and upper (Alley and Cpx-phyric

Units) lava units have distinct trace element contents and ratios, for example different R E E patterns, that cannot be easily related by fractionation or other high-level magma chamber or crustal processes to a common parent magma. These differences most likely reflect different primary magma compositions produced by complex partial melting processes in the mantle. The most primitive (Mg* >65, Cr >150 ppm, Ni >50 ppm) lava and dyke compositions (Table 3.26) are found in the Lasail Unit (representing the "early" magmas) and the Cpxphyric Unit (representing the "late" magmas). These rocks are olivine-clinopyroxene-phyric and therefore probably accumulative but Mg/Fe ratios of the phenocrysts and whole rocks suggest that they are in, or close to, equilibrium (Section 3.8.2). Although, as noted by Alabaster (1982), the fields occupied by these two units overlap on the Cr/Y plot (Fig. 3.64), the Cpxphyric Unit is generally displaced towards lower Y contents at equivalent Cr values. Figure 3.64 also shows, in terms of Cr/Y space, the probable trends for basaltic ( o l + c p x + p l a g ) fractionation, the possible primary magma compositions and equilibrium mantle compositions. Using this type of modelling, it seems that the parent picrite magmas (c, 1200 ppm Cr)

Chapter 3

I22

can be derived by 20-30% melting of typical lherzolite sources (c. 2500 ppm Cr). The REEs, and particularly the chondrite-normalized values for Ce, Yb and Ce/Yb, have been used to model possible partial melting processes for the two magma suites. The most primitive Lasail basalts, containing Ti 2600-4200 ppm, Zr 2535 ppm and Cr 250 ppm, can be modelled as the products of 20-30% fractionation from a parent picrite with YbN = 5.4 and (Ce/Yb)N = 0.56 (Fig. 3.65), whereas the primitive Cpx-phyric basalts (Ti 2600-3100 ppm, Zr 20-27 ppm, Cr 150 ppm) can be modelled as the products of 25--40% fractionation of a picrite with YbN = 4.3 and (Ce/Yb)N = 0.4. In modelling the partial melting process, it is currently assumed that melting commences in the spinel lherzolite field at 50-70 km depth. The major phases present at this depth are believed to be olivine, clinopyroxene and orthopyroxene in the approximate proportions Olo.6 opxo. 2 cpxo. 2. Melting occurs in the ratio 01o.2 opxo. 2 cpxo. 6 and, once the clinopyroxene is exhausted, the harzburgite residue (olo~ opxo.2) is assumed to be incapable of further melting. Using these constraints and published partition coefficients (Table 3.27), the Lasail parent magma can be successfully modelled as the 25% melt of a slightly depleted (relative to primitive mantle) lherzolite of composition O1o.61 opxo.2 cpxo.~9, Y b N - 1 . 7 and (Ce/Yb)N=0.6. The earliest depletion

could be caused by the extraction of 4% of a LREE-enriched melt ((Ce/Yb)N=l.89). The residue of the melting is a cpxharzburgite (olo.75 opx~, 2cpxo,os, YbN = 0.16, (Ce/Yb)N = 0.09) (Table 3.27, Fig. 65). The more LREE-depleted Cpx-phyric parent magma can be produced in two ways: (a) As a 20% melt of a more depleted lherzolite (o10.62 opxo. 2 cpxo.ls, YbN = 1.2, (Ce/Yb)N = 0.32), which could be produced by two small-volume (c. 3%), LREE-enriched melt extractions from the primitive mantle. (b) As a 10% melt of the residual cpx-harzburgite, formed by the melting of the Lasail parent magma which retained 5% of the melt (Table 3.27, Fig. 3.65). This process is an incremental, second-stage melting model similar to that proposed by Duncan & Green (1980) for the formation of the Upper Pillow Lavas of the Troodos ophiolite. Melting such already depleted scarting materials probably depends on the availability of water and other volatiles which can be produced by dehydration of the downgoing oceanic crust during subduction causing hydration of the overlying mantle wedge (Hawkesworth et al. 1979; Saunders & Tarney 1979). The second-stage melting model for the Semail upper lavas and late intrusives is preferred because it agrees with the sequence in which the magmas were produced and their relative volumes with the late eruptives amounting to about 30% of the lava sequence.

Table 3.27. Partial melting modelling. Cr (ppm)

Ni (ppm)

Ybv

(Ce/Yb)N

6.5

0.61

1. Axis sequence (Lasail basalts)

Average basalt

250

80

1' c. 30% ol+cpx+plag fractionation (ol-cpx>plag) Parent magma (picrite basalt)

1200

350

5.4

(I.58

2000

1.7

0.55

3100

2550

0.16

0.15

15(1

6(1

5.6

0.43

1' 25% partial melt Lherzolite 1 (ol..~jopxo.2cpx, l,,)

2600 { residue

Harzburgite 1 (oi..75opxozcpx..,,5) 2. Late sequence (Cpx-o basalts)

Average basalt

T c. 40% oi+cpx+plag fractionation (cpx>ol>plag) Primary magma (picrite basalt MODEL ]

Lherzolite 2 (ol..~,2opxo2cpxois)

12(X)

350

4.3

0.40

2200

1.2

(I.32

3100

0.17

(1.119

244(I

0.42

0.17

2670

0.06

0.08

1' 20% partial melt 2700

J, residue Harzburgite 2 (o1r 73opx,i2cpxo.v) MODEL2 Lherzolite 3 (01.7,opx, zcpx,, ~,,)

3000

10c/c partial melt 3(X)0

{ residue Harzburgite 3 (olo.T7OpXo2cpxo.03)

32(X1

Lherzolite 1 Primary mantle (Cr 2500 ppm. Ni 1800 ppm. Yb N 2.5, (Ce/YbN 1) minus a 4% melt fraction. Lherzolite 2 Primary mantle minus two 3% melt fractions. Lherzolite 3 95% residual harzburgite 1 + 5% Lasail parent magma. Fractional crystallization modelled by C~/Co=F (D-~). Partial melting by the equilibration melting equation (C~/Co = 1/(D(1-F)+F)). Partition coefficients from Irvine (1978) and Frev et al. (1978).

The Sernail Ophiolite Spreading~axis

I . -::.7>-Z:~:,S<<-.~:~,~...~< ~C<" ~-.*~.-.~:: <>-Z:*L<:-(>)'.~:-:- .- -- .1

Crustal

}~,jllJJli.!}J.lJ!):.')".'!)!,].[.I.!'!',i,l.[!. }jljlill},l{lll' .... :::::[':]}:! \ Petrological Moho ".."..".. '.--:~!!!!!!!i!iiiii!:::::::: I

Dunite ........... chromitite--~ ~

pods

'L' Z A J _'_ _ _ _ " ' " - ' ~ 7 ~ ~___.~- ~ ~-C r:~ t~:

- - - ~ . - - ~ 0 - - -

,~ / r T~c,ton,te harzourgim foliation

o

~-~,~-"- ~ ~i

- -

reate~ o, Llmospnere_.sthenospnere v <:) ~ boundar, a s

... :. .. :~ . ..2.. : , < t . < ~,

9~ - :

,LD

~ ~ -

~ \ . % ,~ ~ 9 /' O/

O

\~

~A

0 ~

..

0 (

o 00~ Mantle flow lines

l.

(

" 0

o

0

coalesce

and begin to rise Outer limit of zone of / abundant (>50/0) partial melting ~

/ /

/

t

Mantle upwelling

/

/

/ f

Initiation of partial melting

/

1

0 10km I I Approximate horizontal and vertical scale

Fig. 3.66. Model of mantle melting and diapiric uprise beneath the spreading axis.

The (Ce/Yb)N values of the Semail harzburgites (Fig. 3.65) are similar to those prediced by the multi-stage partial melting process. However, the marked "V"-shaped patterns with depletion of the middle REEs cannot be explained by partial melting and may be the result of later introduction of LREEs now residing in fluid inclusions or in crystal lattice defects (S. Roberts pers. comm.).

3.9.3 Magma fractionation Contrasting primary magma compositions are the most likely cause of the different crystallization sequences found in the Semail Layered Series as well as between the Layered Series and some of the Late Intrusive Complexes (Sections 3.4 & 3.6). Such variability is most likely to be the result of earlier melting episodes in the mantle, as described below. In general, magmas produced by the partial melting of already partially depleted mantle will have relatively high CaO/AI203 ratios and will give rise to ol--~cpx---,plag crystallization sequences at

I23

low pressures. Ol--,plag---~cpx, the most common MORB sequence, is rare in the Semail cumulates and results from magmas with low CaO/AI20~ caused by melting of a less depleted source (Browning 1982). Ol--,opx--*cpx--~plag sequences, such as seen in the middle part of the Wadi Ragmi section, may result from the melting of an even more depleted source. Likewise, many of the Late Intrusive Complexes show the early crystallization of orthopyroxene. At elevated water pressures, for advanced degrees of melting, the CaO/AI203 ratio of the primary magma will not vary but it is likely to be richer in S i O 2 and MgO and more depleted in incompatible elements (Ford 1976). The picrite melts, formed by the initial melting of upwardflowing mantle at c. 50 km depth, rise together with the residual mantle until at c. 30 km the melts coalesce to form small diapirs (Gass 8,: Smewing 1981). Above this level the diapirs rise adiabatically through the host peridotite which itself continues to rise (but probably more slowly than the melt) owing to the mantle convection beneath the spreading axis (Fig. 3.66). Because the olivine stability volume expands with decreasing pressure and the prevailing hydrous conditions, picritic melts will remain within this volume as they ascend through the upper mantle, during which ascent only olivine and chrome spinel will be precipitated. This process necessitates a high melt to rock ratio to prevent much magmawall rock reaction taking place and accounts for the dunite bodies which form 5-10% of the Mantle Sequence (Section 3.2). Variable amounts of fractionation during ascent through the upper mantle is the most likely cause of the variations in the phase assemblages of the basal cumulates which range from dunites to gabbros (Section 3.4). Where the magmas arrive at the crustal magma chambers close to the three-phase gabbro cotectic, then they will tend to mix freely with the existing melt in the chamber and give rise to gabbroic cumulates at or close to the base of the layered sequence. Conversely, where relatively more primitive (ultrabasic) melts enter the chamber, they are less likely to mix with the magma already there and form a basal picritic layer which precipitates a layer of basal ultramafic cumulates (Section 3.4.3). Repeated influxes of primitive melt give rise to the cyclic units so common in the basal cumulate sections (Section 3.4.2). Browning (1982, 1984) has demonstrated that the cryptic mineralogical variations through the Semail Layered Series, with numerous compositional resets and little overall trend towards more fractionated compositions, is compatible with fractionation of the magmas in a large, open-system magma chamber. This leads to the production of large volumes of moderately evolved basaltic magmas which were discharged from the chamber to produce the Sheeted Dyke Complex and Geotimes iavas. An important feature of this type of system is the decoupling of the compatible and incompatible elements (O'Hara 1977: O'Hara & Matthews 1981) so that the erupted products have high concentrations of the latter relative to their only moderately fractionated major element compositions (Section 3.8). On the other hand, the later Lasail and particularly Alley lava units seem to have been erupted from smaller, higher level magma chambers which became easily detached from their feeder zones and evolved under more closed system conditions giving rise to a higher proportion of more evolved (acidic) magmas. Using such open versus closed system fractionation schemes, it is possible to account for the relative volumes of the axis and late acidic differentiation products and their compositional differences.

I24

Chapter

3.10 Ocean-floor Metamorphism The effects of ocean-floor metamorphism have been widely reported from present day oceanic basalts and plutonic rocks (Cann 1969, 1979; Bonatti et al. 1975; Humphris & Thompson 1978; Moody 1979; Fox & Stroup 1981) and from ophiolite complexes (Gass & Smewing 1973; Spooner & Fyfe 1973; Coleman 1977; Elthon & Stern 1978, and many others). The well-known characteristics of this type of metamorphism are: 1 That it is "static", i.e. imposed on the rocks without accompanying deformation and often leaves primary igneous textures preserved. 2 An increase in metamorphic grade with depth from brownstone (submarine weathering) to zeolite facies in the uppermost lavas, greenschist facies in the lower lavas and sheeted dyke complex to lower amphibolite facies in the underlying gabbros indicating metamorphic gradients of over 200~ (Gass & Smewing 1973; Spooner & Fyfe 1973). 3 The presence of disequilibrium mineral assemblages indicating incomplete reactions; for example, the preservation of relict igneous phases and lower grade assemblages partly overprinting higher grade ones and vice versa. Evidence from active hydrothermal vents near present day ridges (Corliss et al. 1979; Rona 1984), heat flow studies (Lister 1972) and theoretical modelling (Spooner & Fyfe 1973; Wolery & Sleep 1976) suggest that convectionally-driven circulation of heated seawater is the agent responsible for oceanfloor metamorphism. This is supported by laboratory studies of basalt-seawater interactions (Mottl & Seyfried 1980) and stable isotope studies of altered ocean-floor and ophiolite rocks (Heaton & Sheppard 1977; Gregory & Taylor 1981; Stakes & O'Neil 1982).

3

are descriptively inappropriate. In particular, as first noted by Miyashiro et al. (1971), at the intermediate grades of oceanic metamorphism calcic plagioclase persists throughout the amphibolite facies although the coexisting clinopyroxenes are altered to actinolite and other low temperature metamorphic minerals. We therefore follow Elthon & Stern (1978) and use the term "actinolite facies" to describe those high-level gabbros and dolerites characterized by the metamorphic assemblage calcic plagioclase-actinolite rather than lower amphibolite facies which more appropriately describes the calcic hornblende-intermediate plagioclase assemblage of moderate pressure regional metamorphic schists. By the same token, the term "greenschist facies" is also inappropriate for non-schistose rocks and should, we believe, be replaced by the term "greenstone facies". We realize however, that this has no historical precedent and therefore have retained the original term. Other alternatives, such as epidote-albite facies, are more cumbersome and, in our view, less appropriate. Liou et al. (1974) found experimentally that the upper greenschist facies assemblage of chlorite + albite + Al-poor amphibole becomes unstable at about 475~ and 2 kb which is taken to define the boundary between the greenschist and actinolite facies. Nitsch (1971) showed that the assemblage prehnite + pumpellyite + chlorite + quartz is stable up to 345 _+ 20~ where it is replaced by the assemblage actinolite + chlorite + epidote + quartz. Thus, 350 ~ is taken as the temperature boundary between the upper and lower greenschist facies. Reactions that mark the lower limit of the greenschist facies, such as chlorite replacing smectite and the appearance of epidote, occur in the temperature range 200-250~ (Sigvaldason 1962: Tomasson & Kristmansdottir 1972). The boundary between the zeolite and brownstone facies is less well defined and in our view it seems best to regard the brownstone facies as the 0~ conditions of seafloor weathering, although Cann (1979) suggests that the temperature of the brownstonezeolite facies boundary should be between 50-100~

3.10.1 Metamorphic facies; nomenclature and conditions Most descriptions of oceanic crust metamorphism in both present day in situ situations and in ophiolite complexes use terms, such as greenschist and amphibolite, derived from studies of regionally metamorphosed continental rocks which

3.10.2 Ocean-floor metamorphism in the Semail ophiolite All of the upper crustal units of the Semail ophiolite have suffered pervasive medium to low grade hydrothermal meta-

Table 3.28. Ocean-floor metamorphic facies. The Brownstone facies is usually limited to the uppermost 100 m of exposed lavas. Facies terminology and limiting temperatures

A lternative facies terminology

Diagnostic metamorphic minerals

K-rich Fe-illite + Mg-rich smectite

Brownstone ...... 20__50oc Zeolite

Zeolites, smectite + albite + sphene + calcite

...... 200-250oc Lower greenschist

Albite + epidote + chlorite + sphene + quartz + calcite prehnite-pumpellyite

...... 350oc Upper greenschist

lower actinolite

Albite + actinolite (2.5-5% A!203) + epidote + chlorite +sphene + quartz

...... 475~ Upper actinolite

lower amphibolite

Calcic plagioclase + actinolite (>5% A1203)

The Semail Ophiolite

I2 5

PRIMARY IGNEOUS

[' "130

r>

0O

1

:u

--t

q m

m

.< q

-~ q

m

m

_ -4

q m

m

,

Depth km SALAHI UNIT

I I

Cpx-~ UNIT ALLEY UNIT

I LASAIL UNIT

q rn

I

I

i

i

J I

I

I

I

200-250~

I

S

Lower Greenschist facies

i I

GEOTIMES UNIT

I

0

i i

o

c. 350 ~ C

3-

SHEETED DYKE COMPLEX

1 HIGH LEVEL INTRUSIVES

T

A

1::u

Upper Greenschist facies

~ ,~ S

c. 475~

o ~. O

i i

Actinolite facies

'

__

~~

I

Fig. 3.67. Metamorphic facies and distribution of mineral phases in the upper part of the Semail ophiolite.

morphism of ocean-floor type (Coleman 1977; Smewing 1980a; Pallister 1981; Pallister & Hopson 1981). Gregory & Taylor (1981) have shown, on the basis of ~sO isotope studies, that seawater alteration or re-equilibration affected much of the layered sequence but left little or no mineralogical record. Alabaster et al. (1980), Alabaster (1982) and Alabaster & Pearce (1985) have demonstrated a complex metamorphic history for the Semail lava sequence involving several phases of alteration related to cycles of magmatic/hydrothermal activity. The metamorphic assemblages are here described from lowest to highest grade which, in general, represent increasing depth in the ophiolite succession (Fig. 3.67). 3.10.2.1 Brownstone facies

The brownstone facies is characterized by low temperature clays, particularly a K-rich dioctahedral Fe-illite, formed by the alteration of olivine and interstitial glass under oxidizing conditions, and Mg-rich trioctahedral smectite formed under more reducing conditions (Cann 1979). Alabaster (1982) has identified iron smectites in the groundmass of the Geotimes lavas which he claims are relicts of an early brownstone facies now mostly overprinted by the lower greenschist facies assemblage that is typical of these rocks.

3.10.2.2 Zeolite facies The Alley Unit basalts contain abundant zeolite minerals, such as stilbite and mesolite, along with celadonite they are found infilling and lining vesicles and sometimes replacing igneous minerals. Stilbite is particularly common and replaces plagioclase and occurs as large sheaf-like radiating aggregates in the groundmass of these lavas. It is also sometimes found as vesicle and cavity infillings in the Geotimes Unit. Mesolite and celadonite are confined to the Alley Unit, the latter occurring as a dark green groundmass mineral and as vesicle lining. In these rocks chlorite is partly replaced by smectite and it appears that the zeolite facies assemblage overprints an earlier lower greenschist facies metamorphism (Alabaster & Pearce 1985). The calcium zeolite laumontite is present in cavities in the upper part of the Lasail Unit which indicates retrogressive metamorphism of these rocks at temperatures of 150-250~ (Liou 1971). 3.10.2.3 Greenschist facies

Most of the sheeted dykes and the extrusive rocks of the Semail ophiolite have been affected by greenschist facies metamorphism. The metamorphic mineral assemblages are complex and there is a general heterogeneity and irregular

I26

Chapter 3

distribution of many of the secondary minerals. This, combined with the presence of relict igneous minerals (mainly clinopyroxenes) in relatively low grade metamorphic assemblages suggests that equilibrium in the temperature range 250450~ was only rarely and locally attained. Olivines are altered to mixtures of chlorite and epidote, clinopyroxenes to actinolite and chlorite and plagioclase to albite, prehnite, epidote and quartz. Sphene and disseminated Fe-oxides replace titanomagnetites, and epidote, chlorite, prehnite, Fe-oxides and pyrite occur in the groundmass and infill cavities and veins. The actinolites in the upper greenschist facies rocks have AI203 c o n t e n t s of 2.06-4.36% which overlap with the range of values in the upper actinolite facies rocks. Actinolite replacing clinopyroxene is common in the sheeted dykes but less so in the extrusives where chlorite is the dominant alteration product of ferromagnesian minerals. The chlorites have a wide range of compositions (Fe/(Fe+Mg) 0.2-0.6); fibrous green Fe-chlorites (Fe/(Fe+Mg) >0.4) mainly replace pyroxene and/ or actinolite. The chlorite that replaces olivine is less iron-rich (Fe/Fe+Mg 0.2-0.4) (Fig. 3.68). The general absence of actinolite and the greater preservation of clinopyroxenes in the volcanics suggests that temperatures of metamorphism were lower than in the sheeted dykes. Ellis & Green (1979) suggest temperatures of about 350~ for the transition although secondary amphibole formation from clinopyroxene is also suppressed by high f 9 2 (Liou et al. 1974). Epidotes are common throughout the sheeted dykes and lower lava units (Geotimes and Lasail Units). In the dykes they are most abundant in the green epidosite dykes and breccias that were channelways for hydrothermal fluids. In the Geotimes lavas epidote occurs mainly as cavity and vein infil-

0.7 cr

0.6

O5

1O 1~ O1

O5

o

0.5 Fe Fe + Mg

a

2 ~F-

4 0 ~ 94

z <

E-

0.4

6

n

03

9

6o

0.3

So 2

J

f

0.2

sj 0.1

I 5

6

7

Si O Olivine pseudomorphs 9 Groundmass

Fig. 3.68. Classification of chlorites (Hey 1954) from the Semail ophiolite. Olivine pseudomorphs: 1. Cumulate gabbro, 2. Cpx-o basalt, 3. Lasail basalt. Groundmass: 1. Sheeted dykes, 2. Geotimes basalt, 3. Lasail basalt, 4. Cpx-r basalt, 5, Salahi basalt, 6. Alley basalt.

lings where it forms coarse fan-shaped aggregates with a pale yellow-green pleochroism. In the Lasail lavas it occurs as an abundant groundmass mineral and also in a brown granular form that replaces olivine. There is no compositional difference between the various epidotes, all are Fe-rich types with pistacite contents of 25-29%. Pale green to colourless epidote occurs with quartz and chalcedony as vesicle infillings in felsite sills and rhyolite lavas in the Lasail and Alley Units. Sphene is widespread in the greenschist facies assemblages as an alteration product of Fe-Ti oxides and occasionally wellformed crystals in cavities and hydrothermal veins. Haematite is particularly common in the Geotimes lavas but, associated with oxy-hydroxides, is present throughout the lava succession. Prismatic to fibrous prehnite occurs in interstitial cavities and in veins and fractures in the Sheeted Dyke Complex. It occurs as vesicle and vein infillings in the Geotimes lavas but is most abundant in the Lasail basalts where it occurs as large crystalline aggregates up to 5 mm across replacing the groundmass, in vesicles and as occasional pseudomorphs after plagioclase. The prehnites have variable Fe203 contents (0.37.1%) even within a single crystal (Alabaster 1982). Boles & Coombs (1977) suggest a minimum temperature for the appearance of prehnite at 90~ and that the mineral is stable up to at least 400~ (Liou 1971). Fe-rich pumpellyite has been identified in a few samples of the sheeted dykes as sheaf-like aggregates of acicular crystals apparently replacing actinolite. Mevel (1981) has pointed out that the non-coexistence of epidote and pumpellyite in oceanic basalts and she suggests that the pumpellyite-prehnite facies in ocean-floor metamorphism is most usually represented by the assemblage of prehn i t e + epidote, which is the typical association of the Lasail basalts (Alabaster 1982). Greenschist facies mineral assemblages, dominated by epidote, prehnite and sphene, occur as cavity and vein infillings in the high-level intrusives, particularly the plagiogranites where they may affect the whole rock. Epidote-quartz-sulphide veins occur at deeper levels in the gabbros showing that mineralizing fluids passed through these rocks along joints. Local zones of high intensity alteration in the overlying sheeted dykes and lavas represent the channelways of the same fluids as they passed upwards through the oceanic crust. These can be recognized as epidote-quartz veins and epidosite dykes and as quartz-chlorite-sulphide veins and stockworks beneath the massive sulphide deposits (Section 3.10.4). The stockworks are comprised of mineralized dykes and lava fragments, dominated by quartz-chlorite alteration, in a silicified matrix of quartz, pyrite, sphene and chalcopyrite. Relict magnetite and ilmenite occur as inclusions in the pyrite in both the stockwork and the overlying massive ore body. The stockwork zone may be cut by later lower temperature veins of quartz, haematite and calcite. Seyfried et al. (1979) and, more recently, Mottl (1983) have shown that the intense quartz-chlorite alteration of metabasalts in the stockwork zones is the product of high water/rock ratios (>50) and suggest that they probably form as the upflow limbs of hydrothermal circulation systems where the fluids are highly channelized. 3.10.2.4 Actinolite facies

Most of the high-level gabbros of both the axis sequence and the late intrusive complexes and some of the dykes at the base of the sheeted complex are characterized by an association of primary calcic plagioclase, the unaltered nature of which is

The Semail

evident from optical and microprobe studies which show normal compositional zoning in the range Ango_so, and a secondary Al-poor brown, or more commonly green, actinolite or actinolitic hornblende replacing clinopyroxene or, in rare cases, hornblende. Primary Fe-Ti oxides, mainly titanomagnetites or magnetites with exsolution lamellae of ilmenite, are largely unaltered in these rocks. The AI203 contents of the secondary amphiboles range from 2.4 to 5.9% suggesting that the boundary between the upper and lower actinolite (greenschist) facies amphiboles placed at 5% AI20 3 by Elthon & Stern (1978) does not apply to the Semail rocks. Analyses of clinopyroxene-actinolite pairs from a high-level gabbro show increases in AI, Fe, Ti, Na and K and a decrease in Ca content with alteration (Alabaster 1982). Mg/(Mg+Fe) decreases from 0.79 in the pyroxenes to 0.73-0.63 in the actinolites. The secondary amphiboles are often fibrous and clearly replacive although they occasionally have subhedral prismatic forms and are difficult to distinguish from primary brown-green hornblendes, although the primary amphiboles usually have markedly higher A120 3 and TiO2 contents than the secondary ones. The high-level gabbros often show patchy replacement by greenschist facies assemblages such as saussuritization of the feldspars and by the presence of cavity and vein infillings of epidote, prehnite and quartz. In the associated diorites and plagiogranites greenschist facies assemblages predominate and the feldspars are usually completely altered except for oligoclase-albite rims and overgrowths.

3.10.3 Summary In the upper part of the Semail ophiolite there is a general increase in metamorphic grade from Brownstone and lower greenschist facies (350-250~ in the lavas and from upper greenschist (350-475~ to actinolite (475-550~ facies in the sheeted dykes and high-level gabbros (Fig. 3.67). The probable temperatures of the facies boundaries suggest an overall temperature gradient through the 4 km of section of about 150~ However, the pattern is complicated by several features: (i) Most of the rocks show evidence of several phases or pulses of metamorphism, usually of a retrogressive nature with lower grade assemblages overprinting higher grade ones. Commonly the highest grade event is a bulk rock alteration which is partly overprinted by lower grade alteration confined to dispersed vein and cavity infilling. For example, most of the zeolite facies alteration in the Alley Unit lavas is a late-stage vein and cavity alteration whereas the groundmass of the rock remains in lower greenschist facies. Likewise, the actinolite facies gabbros often contain greenschist facies vein assemblages. These features indicate several episodes of different temperature hydrothermal fluids and renewed convection triggered by new cycles of intrusive and extrusive activity. Alabaster (1982) and Alabaster & Pearce (1985) suggest that the metamorphic history of the lavas is complex because of the intense but localized hydrothermal activity related to the Lasail and Alley Unit magmatism. (ii) The presence of lower greenschist assemblages (chloritealbite-epidote) in all the lavas units, up to and including the uppermost Salahi Unit, shows that the whole volcanic section was subjected to temperatures of at least 200-250~ This implies a high geothermal gradient near to the surface. Alabaster & Pearce (1985) suggest that the late pervasive greenschist facies overprint of the entire eruptive sequence occurred at a

Ophiolite

I27

late stage and was caused by the "sealing in" of the hydrothermal system by the sedimentary cap overlying the lavas. They cite Davis & Lister (1977) who showed theoretically that temperatures at the base of an impervious sediment cap only a few metres thick overlying young oceanic crust could reach 200~ (iii) There are areas of unusually high grade metamorphism where stratigraphically higher units show mineral assemblages typical of lower ones. For example, in the Lasail area, the Lasail basalts which form the hanging-wall to the massive sulphide ore-body contain actinolite and show upper greenschist facies alteration more typical of the Sheeted Dyke Complex. In most other areas they are characterized by an epidoteprehnite-chlorite lower greenschist assemblage.

3.10.4 Massive sulphide deposits The Extrusive Sequence contains numerous gossans and other showings (including slag heaps and other evidence of ancient Sumarian working) of massive base metal sulphide mineralization (Bailey & Coleman 1975: Coleman et al. 1979). Following on from a prospecting programme using geophysical techniques and borehole sampling, three of the largest deposits in the Wadi Jizi area, at Lasail, Bayda and Aarja, are currently being mined by the Oman Mining Company. The geological setting of the ores has been described b)' Alabaster et al. (1980) and Alabaster (1982). Ixer et al. (1984) have made a detailed study of mineralogy and geochemistry of the ores and Alabaster and Pearce (1985) discuss the relationship between the hydrothermal metamorphism of the lavas and ore deposition. The deposits are typical exhalative '~Cyprus-type", such as those described from the Troodos massif (e.g. Constantiniou & Govett 1973) and subsequently recognized as forming at hydrothermal vents on modern spreading axes (Francheteau et al. 1979; Oudin et al. 1981). Haymon et al. (1984) have described fossil worm tubes, similar to those found in modern vent communities, in the Bayda massive sulphides. Chen & Pallister (1981) show that lead isotope data from the Fe-Cu sulphides indicate a close genetic relation with the igneous rocks of the ophiolite suggesting that the source of the sulphide is the oceanic crust (Section 3.12.1). The gossans overlying the exposed ore bodies form areas of brightly coloured ground up to 100 m across (Lasail) and contain a variety of secondary minerals, including limonite, jarosite, haematite, azurite, malachite and native copper. The underlying mineralized stockwork zones of altered volcanics beneath the ore bodies are characterized by quartz-pyritesphene-chalcopyrite mineralization in amygdales and crosscutting veinlets. The primary ore bodies themselves are essentially massive pyritic ores containing varied amounts of chalcopyrite and sphalerite. They consist largely of massive early-formed euhedral or colloform pyrite with some associated haematite-magnetite and chalcopyrite, secondary chalcopyrite-sphalerite-bornite (mostly in fractures in the pyrite), and late-stage replacements or veins of haematite-quartz (Ixer et al. 1984). Apart from quartz, other gangue minerals present are calcite, gypsum, chlorite and epidote. Within the ore, values of iron and copper are erratic and copper contents >0.5 wt % only occur where chalcopyrite is a conspicuous phase. Except locally, where sphalerite is present, zinc values are generally low (<0.15 wt %), as are the concentrations of Pb, C d , Mo, Ni and As, but Co values are relatively high ranging up to 500 ppm. All the major massive sulphide bodies are located at the top of the Geotimes Unit lavas. In all cases the footwall lavas

I28

Chapter 3

beneath the deposits are silicified and brecciated to form a chlorite-quartz-pyrite stockwork and form an irregular surface upon which the massive sulphide was deposited. At Lasail the ore body is overlain by barren Lasail Unit basalts and cut by up to 40% in volume of unmineralized andesitic inclined sheets, dipping at c. 25 ~ W. The ore body is roughly saucer-shaped, slightly elongated N-S and dips 30--40~ E with a down-dip extension of at least 200 m and has a maximum thickness of about 50 m. Towards the west the ore body thins rapidly and passes laterally into a 1 m thick umber. The Bayda deposit is located' 8 km N of Lasail and there the ore body consists of a number of irregular pod-shaped masses each of which is underlain by a stockwork zone. They are strung out along a NW-SE fault and cut by E-W faults. The hanging wall at Bayda is composed of Lasail Unit basaltic andesites. At the nearby Aarja deposit, 1 km to the SE, the ore body is cigar-shaped, approximately 75 m long in a NW-SE direction and plunges at 30 ~ to the southeast. It is located close to a NW-SE fault but, unlike Bayda, the ore body is not affected by later faulting. In contrast to the Lasail and Bayda deposits, the hanging wall lavas at Aarja belong to the Alley Unit. This and other differences, such as a higher Zn content of the ore and mineral paragenesis indicate lower temperatures of formation (c. 275~ as opposed to >320~ (Ixer et al. 1984)) and that the Aarja deposit formed later and from cooler hydrothermal fluids than those at Lasail and Bayda. As pointed out by Alabaster (1982), these deposits and the other major gossans are located within the so-called "seamount" areas where they are associated with Lasail and Alley Unit intrusive-extrusive complexes and late gabbro-plagiogranite plutonic bodies (Section 3.6.1). Alabaster & Pearce (1985) have developed a model in which they envisage that the mineralization occurred on the flanks of the magma chambers where the high-level intrusions provided the heat necessary for driving the hydrothermal circulation system and produced a fracture system that allowed concentrated discharge of the orebearing fluids onto the sea-floor. Another important factor in the formation of an ore deposit was the presence of a preexisting depression into which the metal-rich discharge could be ponded. The lateral passage from massive sulphide into FeMn oxy-hydroxide sediment (umiSer) seen at Lasail can be attributed to fractionation around the hydrothermal vent or vents as the ore-forming solutions mixed with seawater and became progressively more dispersed and oxidized (Rona 1984). Minor sulphide-bearing veins occur in faults and joints throughout the crustal sequence, particularly in the sheeted dykes and high-level gabbros. Some of the largest of these mineralized zones, in the Maydan area of the Wuqbah block, are associated with NW-SE faults that control the late intrusive complexes (Rothery 1982).

3.11 Metalliferous and Pelagic Sediments Metalliferous and associated pelagic sediments, which are typically found in the upper parts of the ophiolite sequence, provide strong evidence for a deep-water submarine origin for the lavas. Most typical are umbers which are dark coloured ferromanganoan mudstones of amorphous Fe-Mn oxy-hydroxides that are often rich in trace metals such as Cu and Ni (Robertson & Hudson 1973). The chocolate brown to redbrown eoloured umbers of the Semail ophiolite form lenses (up to 5 m thick) that were deposited on the lava surface and

were then buried by either later eruptions or post-lava sediments. Above the lava sequence the umbers give way up section to paler coloured radiolarian mudstones of the Suhaylah Formation which have higher silica and lower metal contents. In places, grey laminated calcilutites overlie the mudstones. All these fine grained pelagic sediments are locally interbedded with, or overlain by, fine to coarse grained clastic mudstones, siitstones, sandstones and conglomerates (Zabyat) Formation), composed largely of ophiolite debris and probably formed by the erosion of submarine fault scarps on the ocean floor. The Semail sediments have been studied in detail by Fleet & Robertson (1980) and Robertson & Woodcock (1983b). 3.11.1 Fauna and age

The Semail umbers and radiolarian mudstones contain up to 80% radiolarian tests that are usually moderately well-preserved where they are infilled and replaced by chalcedony. They occur in a matrix of ferromanganese oxides and hydroxides or micrite and some of the tests are incompletely replaced by carbonate making identification more difficult. The calcilutites contain scattered and broken radiolarian tests as well as some planktonic foraminifera but these are usually completely replaced by carbonate and difficult to identify. Glennie et al. (1974) identified the foraminifera Rotalipora sp. of Cenomanian age in a pelagic lime-mudstone overlying pillow lavas in Wadi Jizi, whilst from further north in Wadi Ragmi they found Globotruncana sp., including G. signali and G. scheegansi, in radiolarian cherts and silicified iime-mudstones interbedded with pillow lavas. Glennie et al. (1974) suggested that this fauna indicated a Coniacian age for these rocks but these forms are now known to have a greater age range extending back to the Cenomanian. More extensive faunas were subsequently described, mainly from the Wadi Jizi area, by Tippit et al. (1981) (Table 3.29). They found that the interlava sediments in the lowermost Geotimes Unit yield an early Cenomanian (97.5-95 Ma) radiolarian fauna, whereas those in the overlying Alley Unit give Cenomanian to Lower Turonian (95-90 Ma) ages. In the supralava Suhaylah Formation the lower umbers, which lie 1.5 m above the contact with the Geotimes Unit, have an early Cenomanian (97.5-95 Ma) fauna, whereas the overlying calcilutites yield a fauna with much younger Coniacian-Santonian (88.5-85 Ma) ages. This suggests that the sediments at Suhaylah in Wadi Jizi, although only 15-20 m thick, were deposited over a time interval of about 10 Ma. This agrees with the fine grained nature of these sediments and the possible tectonic position of the Suhaylah area on an upfaulted part of the sea-floor. 3.11.2 Stratigraphy and field relations

The Semail metalliferous and pelagic sediments are described in terms of (i) interlava occurrences within and between the Geotimes, Lasail, Alley and Salahi lava units, (ii) the supralava Suhaylah Formation and (iii) the submarine clastics of the Zabyat Formation. 3.11.2.1 lnterlava occurrences

Thin (<1 m thick) umbers, sometimes associated with orangebrown ochres (Fe-rich but Mn-poor deposits formed by submarine weathering and oxidation of sulphides (Constantinou & Govett 1972)), occur throughout the Geotimes basalts but,

The Semail Ophiolite

I2 9

Table 3.29. Radiolarian faunas from Semail pelagic sediments.

Rock type

Locality

Stratigraphic position

Laminated calcilutite

Suhaylah village, Wadi Jizi

Suhaylah Fro., 16 m above umber sample

43442~833

Fauna

Age

Hemicrvptocapsa polyhedra Crucella cachensis A rchaeospongoprun um

Coniacian Santonian -

Zone 12-13

rItDISCVeHse

Umber

Suhaylah village, Wadi Jizi 4344~6833

Umber

Mulayyinah, Wadi Jizi 441126858

Umber

Mulayyinah, Wadi Jizi 441126858

Umber

44072~857

Red calcareous mudstone

Wadi Lasail 440526844

Red chert

Wadi Lasail 440526844

Umber

Umber

Wadi Jizi 436226854

Rakah 457026181

Suhaylah Fm., 1.5 m above Geotimes Unit lava contact

Top umber, just below serpentine conglomerate Lower umber, below upper Alley Unit basalt flow

Dictvomitra formosa Praeconocarvomma universa Dictvomitra duodecimocostata Q uinq u ecapsula ria spinosa Holoco'ptocanium tuberculatum H. geyserense Pseudodictvornitra pseudomacrocephala Tharnarla veneta T. pulchra Hemico'ptocapsa prepolyhedra Novixitus Dictyomitra formosa Hernico'ptocapsa polyhedra

Amphipyndax stocki A rtostrobium sp. Pseudodictvomitra tiara A canthocircus heuyi Novixitus sp. Umber resting A lievium antiguum on Alley acid Acanthocircus squaboli centre volcanics Hemicrvptocapsa polyhedra Tharnarla veneta Xitus spicularis Interbedded in A rtostrobium tina Alley Unit Praeconocarynima basalts lipmanae Interbedded in (i) Hemicryptocapsa polyhedra Alley Unit Lithomelissa petila basalts (ii) Tharnarla pulchra Pseudodictyomitra Pseudomacrocephala Interbedded Pseudodictvomitra Pseudomacrocephala in Geotimes Hernicryptocapsa Unit basalts prepolyhedra Novixitus Schichomitra sp. (aft. S. magna) Halesium Interbedded sexangulum in Geotimes Rotaforma hessi Unit basalts Holocryptocanium tubercalatum Tharnarla elegantissima Petasiforma foremanae Acanthocircus multidentatus A. heuyi

Early Cenomanian Zone 10, subzone 10a

Turonian to Santonian Zone 11-13

Cenomanian to Turonian

Cenomanian to early Turonian Zone 10, subzone 10a to Zone 11, subzone 1la Cenomanian Zone 10 (undiff)

Early Turonian

Early Cenomanian to early Turonian Early Cenomanian Zone 10, subzone 10

Early Cenomanian Zone 10, subzone 10

I30

Chapter 3

in general, sediment deposition and hydrothermal mineralization is of relatively minor importance in the Geotimes Unit. A typical example, at Zuha in Wadi Salahi, occurs where the upper part of the Geotimes lavas are silicified and veined with haematite over a lenticular zone measuring 30 m by 50 m. This is overlain successively by a brightly coloured 3 m thick ochre and unmineralized pillow lavas, that pass upwards into a second brecciated and mineralized zone capped by another c. 5 m thick ochre horizon which is laminated at the base with partings of volcaniclastic silt. The ochres, which are interpreted by Fleet & Robertson (1980) as oxidized sulphides, are pseudoconglomeratic and rubbly and pass laterally into redbrown silty manganiferous umbers. The metalliferous sediments are overlain by a massive, columnar-jointed lava flow which marks the top of the Geotimes Unit in Wadi Salahi. The Lasail Unit lavas contain little or no peiagic sediments and were apparently erupted relatively rapidly; however, umbers, up to 1.5 m thick and traceable for up to 200 m along strike are common in the higher Alley Unit. A 1.5-2 m thick umber occurs at the base of this unit where it rests directly on the Geotimes Unit indicating a hiatus in eruptions that is elsewhere represented by the Lasail Unit. The umbers within the Alley Unit are typically finely laminated and silicified with occasional lamellae of volcaniclastic silt; some are strongly bioturbated. Ferruginous mounds occur in the Wadi Salahi area (Fleet & Robertson 1980). These are up to 2 m high, 3 m in diameter and 10-20 m apart and either rest directly on the underlying lava, which is silicified and contains ferruginous veins, or on umber. The mounds themselves comprise orange coloured cemented ferrugineous and siliceous oxides capped by laminated, highly manganiferous (up to 36% MnO), soft black to dark brown umber. Fleet & Robertson (1980) tentatively interpret these structures as the result of organic (bacterial?) activity. The sediments within the lavas of the uppermost Salahi Unit consist of layers of grey to pink, finely laminated calcilutites and siliceous mudstones which occur as up to 3 m thick horizons between lava flows. The sediments are sometimes hornfelsed to a hard flinty chert, up to 20 cm thick, at the base of the thickest (30 m) Salahi Unit flows. In general, it seems that the intra-lava sediments become thicker and more abundant higher in the lava succession indicating that volcanism was waning, becoming less voluminous and more intermittent with time. The Geotimes and Lasail units were erupted relatively rapidly and continuously with very little accumulation of sediments between flows, the only major hiatus occurring between the two units when extensive hydrothermal mineralization and associated metalliferous sedimentation took place (Section 3.10.4). The Alley Unit contains numerous umber horizons indicating some significant time-gaps between eruptions, whereas the Salahi Unit eruptions were even more infrequent allowing up to 2-3 m of pelagic sediments, including fine grained limestones, to accumulate between flows. Important sediment horizons, up to 2-3 m thick and traceable over several tens of kilometres along strike, occur at the boundaries of the Geotimes-Alley and Alley-Salahi Units indicating major time breaks. The Alley Unit sediments yield Cenomanian to early Turonian (95-90 Ma) faunas compared to the early Cenomanian (97.595 Ma) ages obtained from the Geotimes Unit (Table 3.29). No dates are available on the inter-lava sediments in the Salahi Unit.

3.11.2.2

Supralava sediments

The thickest developments of fine-grained pelagic sediments, up to 30 m thick, overlie the Semail lavas and were assigned to the Suhaylah Formation by Fleet & Robertson (1980). At the type locality in Wadi Jizi (Fig. 3.69) the sediments directly overlie the Geotimes Unit iavas suggesting either (i) that erosion of the higher lava units occurred before sediment deposition, or (ii) that this area was part of a structural high, probably a horst block, that was never covered by the upper lavas. The latter explanation is the more likely as the basal umbers at Suhaylah contain an early Cenomanian fauna the same age as that of the underlying Geotimes lavas. Above this, there is no evidence of a break in sedimentation with the remainder of the Suhaylah Formation, the upper part of which contains a Coniacian-Santonian fauna (Table 3.29). In other areas, e.g. at Mulayyinah in Wadi Jizi and the Wadi Salahi and Wadi Ahin sections, the lowest sediments of the Suhaylah Formation rest on the Alley Unit. The supralava pelagic sediments consist of dark umbers, radiolarian mudstones and calcilutites, similar to many of the interlava occurrences. At Suhaylah (Fig. 3.69) Mn-poor redbrown umbers occur at the base of the succession where they rest on rubbly weathered and brecciated lavas and are overlain by dark brown silicified, laminated umbers that pass upwards into paler coloured, more argillaceous umberiferous mudstones that contain bands of porcellanous chert. The umber succession ranges from 3 to 9 m thick and is overlain by 1 m of calcareous radiolarian mudstones that pass up into 6.5 m of pink to grey coloured calcilutites which have small scale current ripples and are extensively burrowed. At the top of the sequence there are mudstones and volcaniclastic silts that Robertson & Woodcock (1983a) assign to the Zabyat Formation. Between five and ten kilometres to the north of Suhaylah, on the west side of the Alley (Fig. 3.69), similar pelagic sediments, dominantly thin-bedded pink and grey non-calcareous radiolarites and radiolarian mudstones, are interbedded with and overlie lava breccias of the Zabyat Formation. At Mulayyinah, 6.5 km east of Suhaylah, up to 18 m of pelagic sediments overlie rhyolitic agglomerates and lavas at the top of the Alley Unit. Between the volcanics and sediments there are laterally discontinuous lava breccias and conglomerates which contain acid and basic lava fragments in a purple to brown mudstone matrix. The overlying umbers and radiolarian cherts are finely laminated and silicified with only minor intercalations of volcaniclastic silt. A few hundred metres along strike to the east, away from the acid centre, the basalts at the top of the Alley Unit are overlain by 3-4 m of red umbers and brown to grey calcareous mudstones that contain a basalt lava flow. The section is capped by a 15-20 m thick conglomerate composed of basalt, dolerite and gabbro cobbles. Thus, in the Wadi Jizi area, within a few kilometres, the Suhaylah Formation shows markedly contrasting lava-sediment relationships that result from (i) early faulting of the sea-floor, probably along both N-S and NE-SW trends, that produced horst blocks such as that at Suhaylah, and (ii) localized central-type volcanic activity forming acid centres such as that at Mulayyinah producing an irregular sea-floor topography. In Wadi Salahi (Fig. 3.69) the relations are simpler than in the Wadi Jizi area, probably because the sea-floor here was apparently smoother and unfaulted. Here a lens of pelagic

The Semail Ophiolite

~

V

q

V

A,,eyv

~::.:. ~

A :J::::t

---

~

1

:.:?.~!.?.}~f..~: ,:.,?.,~:~Z).,~.5'" v .:: ,' -,~:',,~,v',, i.:: "i:'" v .::::

, ..

- ~

J,,;,,;~;;-~-~",' '.,~:'ti;,:'{:i-'~: '~-

-

-

-

- -- /

~:'~::i

I

':

9 ::::.:":

~~/~,4qp\

0 I

o,-

A

5 km I

I

I

ZabyatFm [~Q-] Coarse rudites k=~..--I v-volcanic clast

4 t: ::::: ] J:L ]

I_~~ I chert clast

I

[~.'~.;'~tFine rudite

/ ~1

[~:~t Sandstone, siltstone ~-----q Mudstone, shale Suhaylah Fm Semail laves ~ Mudstone [ . ~ Volcan!clastic silts

IG-~4 ]----O] I{~:~,--~ /~2) I

~

Ca,c,,ut,te

Radiolarite I11111III I11111lllUmber 2

....:

I3I

-

-

[!~G] cong,om~rate t@.~.t

I~g,c~lVolcanic breccias+[:i:: ::ii:.:t / c----' l autobreccias / ..L~I l-----J Pillow lava /L~s p_ /~:::~-Q~

di S a l a h i i:i)

A

I l : 2 ~ Suhay'ah/Zabyat Fm

~ V ~ A ~

I Major outcrops(see opposite):

:::::~/.. v , J A / :

I 1 Suhaylah

west'side o,

~

13 Mu~ayyio.ah I~i daW ~ Ahin' . . .. . . . .

V/

~v .

1

~4

I~ " J ~ ~ d{ sot~," s~ I "

Massivesulphides

I~"~a i[L~sa,

~L

I sheete ykeoom0.ex

/ 'u,onicroc s / [

i

Majorfaults (Zickondownthrowside)

3

l---Q:)

/

~2

I~:::

||

i:?: :

~b--I /-C~ :b_

--:

-o

/

\A /::

t?

/

V \ /~:. :: 4 ....

~

a~ gossan z~_~_ Zuha

/ /

. Af.::

~:~

I[ - - - ] ~ ,

'

'~.~-_~.:1

v ~

:

.

.

t

..'5.):,i:~. /

~

20m7

//

I9

24000..,

56o30, e

/ ] /

"

.

AlleyUnit ~ <

.

.

.

.

.

.

.

.

.

.

c,. ~ , AlleyUnit \x--,J Geotimes /~'T,--"). Unit 0

Fig. 3.69. The Zabyat and Suhaylah Formations - major outcrops and stratigraphic sections (modified after Fleet & Robertson 1980). sediments that can be traced for 2 km along strike overlies the Alley Unit basalts. The sediments have a maximum thickness of 12.5 m and show a systematic variation in relation to their position within the lens. At the edge of the outcrop they consist of non-calcareous radiolarian mudstones with bands and nodules of replacement chalcedony. These pass laterally, towards the centre of the lens, into interbedded umbers and argillaceous, graded volcaniclastic silts. The lavas underlying the maximum thickness of umbers are intensely brecciated and impregnated with Fe-oxyhydroxides. The sediments of the Suhaylah Formation record a general sequence of umbers overlying lavas or lava breccias passing upwards into more siliceous and less metal-rich mudstones and then, in some places, into fine grained pelagic limestones. In addition to this vertical sequence, there is often a lateral passage from umbers to siliceous sediments and calcareous sediments. This suggests that, at the time of deposition of the Suhaylah Formation, the Semail sea-floor was for the most part above the carbonate compensation depth (CCD) and that acidic solutions caused local solution of carbonate around hydrothermal vents where there is an association of chert deposition with metalliferous sediments. A. H. F. Robertson (pers. comm.) also suggests that the upward passage from siliceous to calcareous deposition (the opposite of that found on modern mid-oceanic ridges which sink with time away from the spreading centre) may result from uplift of the ocean floor

during the earliest stage of emplacement of the Semail Nappe. As already noted, in some areas the supra-lava pelagic sediments are interbedded with and overlain by ophiolitic clastic sediments, conglomerates, sandstones, siltstones and mudstones, called the Zabyat Formation by Woodcock & Robertson (1982a). These rocks, which have been described in detail by Robertson & Woodcock (1983b), are mainly found in the Alley-Wadi Jizi and Wadi Ahin areas where there is a close association with sea-floor faulting. By contrast, the clastic rocks are absent or poorly represented in the intervening, less strongly faulted areas such as Wadi Salahi. The Zabyat Formation consists of cyclic rudite-arenite-lutite sequences from 30 cm to 5 m thick forming amalgamated units up to 100 m thick. The coarse grained deposits contain angular to sub-rounded clasts of altered basalt with less abundant dolerite, gabbro and, rarely, chert and umber. The clasts range from 5 cm to 2 m in size and the deposits are massive and unstratified with no fabric or clast orientation (Plate 3.37). In general, the greater the clast size, the lower is the clast/matrix ratio and they grade from clast-supported to matrix-supported. The sediment blocks include some unusually large (up to 10 m across and 2 m thick) bedding plane rafts. The rudites are interbedded with sandstones that are up to 0.6 m thick and show normal grading and small-scale cross-bedding and convoluted laminations. The mudstones sometimes grade upwards into finely laminated siltstones and red brown mudstones with

I32

Chapter 3

Plate 3.37. (a) Umber horizons in the Alley Unit. Figure is standing on an umber "mound'. (b) Zamyat Formation conglomerate beds (debris flows) interbedded with calcareous shales in Wadi Jizi. scattered lava clasts up to several centimetres in diameter. The matrix-supported texture, the lack of internal stratification and the absence of sorting all suggest that the coarsest grained sediments were deposited as debris flows. The finer grained rudites, arenites and lutites, which are better sorted and graded, were probably deposited as grain flows or turbidites. A. H. F. Robertson (pers. comm.) suggests that the fine grained detrital component of the mudstones is consistent with volcaniclastic silt derived by weathering and erosion of the ophiolite lavas. There is a correlation between clast content and local basement lithologies showing that the Zabyat Formation clastics were mostly locally derived from within a few kilometres of the site of deposition. The occurrence of medium to coarsegrained dolerite and gabbro clasts, which often become more abundant higher in the succession, suggests that faulting penetrated the upper plutonic levels of the oceanic crust. This may have allowed local access of seawater to produce some serpentinization of the deeper level mafic and ultramafic rocks. In the Alley area, described by Smewing et al. (1977), the coarse grained clastics directly overlie the Semail lavas (Geotimes and Alley Units) and, locally, even the Sheeted Dyke Complex, and there is a progressively fining upwards sequence through fine grained clastics to pelagic sediments. This suggests that the N-S faulting in this area, parallel to the spreading axis, developed during or soon after lava eruptions and was less active thereafter as the fault scarp was gradually buried. This contrasts with the situation in the Wadi Jizi and Wadi Ahin areas where the clastics overlie up to 30 m of pelagic sediments and the clast size and proportion of dolerite and gabbro clasts increase up section. This suggests that the NESW faults in these areas developed later than the N-S ones and that erosion bit deeper into the ophiolite section with time, probably as the result of continued faulting keeping pace with erosion. Robertson & Woodcock (1983b) suggest that the deposition of the Zabyat Formation in those areas occurred during "early emplacement" faulting and uplift of the sea-floor

concomitant with the beginning of nappe emplacement. It is worth noting the general similarity of the Zabyat Formation clastics to deposits found in ridge-parallel and transform faultzones in other ophiolite complexes, such as the Ligurian (Italy) and Corsican ophiolites (Barrett & Spooner 1977) and the Arakapas fault zone on the Troodos massif of Cyprus (Simonian & Gass 1978).

3.11.3 Geochemistry of the pelagic sediments Most of the metalliferous sediments are comparable geochemically to the dispersed Fe, Mn and trace metal-rich sediments forming on present day spreading ridges (Rona 1984). The dark coloured umbers are most enriched in iron and manganese with high Cu, Ni, Cr, V and other trace metals (Table 3.30). There is a decrease in metal content with increasing SiO2 in the paler coloured siliceous mudstones that tend towards more normal pelagic clay compositions. The AI203 content is a direct measure of the background detrital component and increases with height above lava base in the mudstones of the Suhaylah Formation (A. H. F. Robertson, pers. comm.). The composition of the fine-grained detrital sediments suggest that they are mostly derived from altered volcaniclastic silt eroded from the Semail lavas. They have gained K, Ce, Ba and Nb during sea-floor weathering, erosion and transport and SiO 2 from biogenic sources; there is no evidence for a terrigenous component in the Semail sediments. This agrees with the Pb isotope data of Gale et al, (1981) [Pb 2~176 37.87-38.07, Pb 2~176 15.42-15.50, Pb 2~176 18.01-18.29], which suggests that all the lead in the umbers is of leached basaltic origin. Contents of Ni and Cr are generally low in the pelagic sediments suggesting that, in keeping with other fault zones in ophiolites, such as the Arakapas fault zone on Cyprus, ultramafic rocks were not extensively exposed on the sea-floor in the Wadi Jizi and Wadi Ahin fault zones during sediment deposition. The umbers are enriched in light rare earth elements with

The Semail Ophiolite

I33

Table 3.30. Umber compositions from the Semail ophiolite. 0M231 S

0M241 Y

0M274 J

200-

0M254 j

49.81 0.23 5.20 26.06 1.80 3.29 12.94 0.02 0.01 0.65

51.15 0.16 5.42 27.04 4.53 0.87 8.24 0.11 0.34 0.92

57.63 0.16 5.10 13.62 5.51 0.82 16.24 0.08 0.16 0.37

73.51 0.41 9.09 8.10 2.84 2.15 0.79 0.35 2.15 0.07

Total

100.01

98.78

99.69

99.46

Ba Cr Cu Ni Pb Sr V Y Zn Zr

16 6 289 310 34 nd 138 628 349 63

80 4 336 233 100 8 239 217 233

174 64 117 98 44 345 182 30 191

42

9(1

La Ce Nd Sm Eu Tb Tm Yb Lu Th Ta Hf Sc Co

80.3 (244.8) 32.9 (38.0) 67.3 (106.8) 13.1 (64.5) 3.58 (46.5) 1.96 (37.7) 0.86 (25.3) 5.04 (22.9) 0.89 (26.2) 4.03 0.33 1.46 6.8 26.0

63 28 306 190 126 85 460 81 196 53 58.5 (178.4) 20.7 (23.9) 59.7 (94.8) 12.6 (62.1) 3.76 (48.8) 2.00 (38.5) 0.91 (26.8) 5.78 (26.3) 0.86 (25.3) 2.8 0.25 1.05 4.4 15.7

OM231 Purple brown mudstone, Geotimes Unit. OM241 Chocolate brown umber, base of Alley Unit. OM274 Brown umber, top of Alley Unit. OM254 Dark brown mudstone, Suhaylah Formation. ( ) Chrondrite normalized values

marked negative Ce anomalies (Fig. 3.70) which can be attributed to absorption of the R E E s from seawater onto the fine grained colloidal metal oxides and hydroxides during their precipitation. The siliceous mudstones have similar patterns to the umbers but at lower R E E contents owing to dilution by clays and biogenic silica. Sr isotopic compositions of the umbers, with initial 87Sr/86Sr ratios of 0.7074-0.7081 (Gale et al. 1981), are indistinguishable from late Cretaceous seawater suggesting that the strontium was also abstracted from seawater at the time of sediment deposition. By and large the O m a n umbers and other metalliferous and pelagic sediments are similar to those from the Troodos (Cyprus) ophiolite (Robertson & Hudson 1973), the rare earth contents (Robertson & Fleet 1976) and isotopic compositions (Gale et al, 1981) of which are closely comparable to the Oman examples. The only difference noted by Fleet & Robertson (1980) is that the O m a n umbers tend to be more siliceous than those on Cyprus which would tend to make them less suitable as a source for pigments.

Dark brown umber

~-

Purple brown umber

~-

Umber

#

1

SiO2 TiO2 A1203 Fe203 MnO MgO CaO Na20 K20 P205

"

Ferruginous vein in lava Pink brown mudstone

100

-

tO r (..) o o rr 50

20

I

1

La

Ce

Nd

Sm Eu

Tb

Tm

Yb Lu

Fig. 3.70. Rare earth element/chondrite plots for metallic and pelagic sediments (data provided by A. J. Fleet).

3.12 Isotopic and Magnetic Studies 3.12.1

Isotopic studies

Isotopic studies on rocks from the Semail ophiolite have been carried out by a n u m b e r of workers. Lead isotopes have been measured by Chen & Pallister (1981), Gale et al. (1981) and Hamelin et al. (1984), strontium and neodymium isotopes by Lanphere et al. (1981), McCulloch etal. (1980, 1981) and H. P. Dunlop (pets. c o m m . ) and oxygen and hydrogen isotopes by Gregory & Taylor (1981), Stakes et al. (1984) and H. P. Dunlop (pers. c o m m . ) (Table 3.31 & 3.16). These data can be discussed under three headings: (i) dating the ophiolite, (ii) identifying its primary magmatic signature and possible tectonic site of formation and (iii) modelling and geothermometry of the sub-sea floor hydrothermal alteration processes in the oceanic crust. (i) McCulloch et al. (1980, 1981) obtained iternal isochrons by the 147Sm-~4-~Nd method using coexisting plagioclase, pyroxene and whole rock analyses of fresh gabbros which gave dates of 130+ 12 Ma (Ibra) and 100___20 Ma (Wadi Fizh, Rustaq). There appears to be no reasonable explanation for the discrepancy between these two ages particularly as plagiogranites from both areas give similar U-Pb zircon ages of 95 _+2 Ma (Tilton et al. 1981). The younger gabbro age is in reasonable agreement with the Cenomanian age of the pelagic sediments (Section 3.11). (ii) McCulloch et al. (1980, 1981) reported a range of ZNd values (+7.6 to +8.6) for the ophiolite, including a harzburgite tectonite and most members of the crustal sequence, which demonstrated the cogenetic nature of the suite and its "oceanic affinity", although they recognized that the values overlapped with M O R B , marginal basin, island arc and, to a lesser degree, some ocean island rocks. Dunlop ( p e r s . c o m m . , briefly

Chapter 3

I34

Table 3.31. Summary of published isotopic data on the Semail Ophiolite. Pb 2062r

Pb 2r

Pb 2 0 8 , 2 0 4

17.0%18.01 18.14 18.06 18.22-19.06 18.14 18.09 18.02-18.14 18.58 18.57

15.43-15.46 15.45 15.45 15.48-15.54 15.46 15.44 15.44-15.46 15.63 15.61

37.66-37.85 37.85 37.80 38.13-38.78 37.97 37.93 37.83-37.97 38.42 38.35

Hamelin et al. 11984) Dolerite dyke Gabbro Gabbro

18.50 17.91 18.00

15.49 15.42 15.45

37.98 37.70 37.83

Gale etal. (1981) Umbers

18.01-18.29

15.42-15.50

37.87-38.07

7 2tl4

Lead Chen & Pallister (1981)

Sulphides Basalt Calcite vein Dole rite dykes Plagiogranite High-level gabbro Cumulate gabbros Cumulate dunite Harzburgite

Lamphere et al. 119811 (100 Ma)

Strontium Sheeted dykes Plagiogranite High-level gabbros Cumulate gabbros Harzburgite

Gale et al. 119811

S r '~7'~0

Sr s'7''~ (95 ~PVIa)

0. 7034-0.7060 0.7040-0.7045 0.7034-0.7039 0.7027-11.7070 0.71156

Umbers

Gregory & Taylor (1981)

Stakes et al. (1984) and D. S. Stakes (pers. comm.)

Oxygen (values given as + 61SOw• relative to SMOW) Basalts 10.7-12.7 Alley Unit rhyolite Cpx-O Unit basalts

Alley Unit basalt Lasail Unit basalt Geotimes Unit basalts

Dolerites dykes

4.%11.4

Plagiogranite

5.2-13.6

High-level gabbros Cumulate gabbros

3.7-5.9 4.5-6.4

0.7073-0.71181

(OM5957) (OM5945) (OM7096) (OM7029) (OM5942) (OM5838) (OM5835)

Late dykes Axis dykes (scc Table 3.16) Axis plagiogranite Metadolerite xenolith in plagiogranitc G abbro (OM5132) Cumulate gabbros

reported in Dunlop & Fouillac 1984) has obtained " M O R B type" ENd values of +9.9 and +10.4 on chrome spinel and diopside respectively from a Mantle Sequence chromitite, +6.5 to +9.6 for a range of plutonic rocks and +6.1 to +9.1 for basaltic lavas from the Crustal Sequence. When combined with ZSr data on the same rocks, the lavas are displaced to high ZSr values as a result of secondary alteration but, perhaps more significantly, the late intrusive complex rocks have higher Zsr than the axis sequence plutonics (Fig. 3.71). It is not clear whether this difference is the result of the greater alteration of the late complexes or whether it reflects different

21.6 16.3 13.3 12.7 7.9 12.4 11.7 8.5 5.6-6.5 6._3--9.6 8.4 5.7 6.2 5.1, 5.7, 6.9

primary magmatic values (analyses are underway on pyroxene separates from the early and late magmatic suites to resolve this problem (G. R. Davies p e t s . c o m m . ) ) . The potassic granite, that is intruded into the lower part of the ophiolite (Section 4.3.2), has an isotopic signature that is clearly different from that of the ophiolite (Fig. 3.71) indicating derivation from an ancient continental crustal source (Dunlop & Fouillac 1984). Hamelin et al. (1984), who incorporated the earlier Pb isotope data of C h e n & Pallister (1981) in their study, suggested that the lead isotopic signature of the Semail ophiolite, with

The Semail Ophiolite +10 p O + - - - ~-A~B--... I ..-~G\ , G*

!:" o

,oP

HG Ae

t35

AG

.~. AU ~

oj,

e AU

*

LUQ

Fig. 3.71. Sr-Nd isotopic diagram for the Semail ophiolite Key to symbols: Plutonic rocks, open circles- axis sequence, closed symbols - late intrusive complexes. W-wehrlite, G-gabbro, HG-hornblende gabbro, D-diorite, FD-ferrodiorite, P-plagiogranite. Basalts, GU - - Geotimes Unit, LU - - Lasail Unit, AU - - Alley Unit. Stars mark data points from McCulloch et al. (1980, 1981), other data from H. M. Dunlop (unpub.). Fields of mid-ocean ridge basalts (MORB) and island-arc basalts (IAB) shown. Inset figure: KG - - biotite granite (OMl18), shaded a r e a - Semail ophiolite. For discussion see text.

Aw

-YNd(95) ~10 0

KG+ --5

20

pb2~176 comparable to MORB and pb2~176 slightly enriched relative to MORB, could be interpreted by an origin "in an interarc basin or premature arc". 6~sO values for mineral separates from the Mantle Sequence rocks give high temperature (>900~ magmatic compositions (opx=5.7% o, cpx=5.6, o1=5.4 and chrome spinel=4.7) (Dunlop & Fouillac 1984). 8D values of fluid inclusions in chrome spinels ( - 5 6 to -79) are typical of magmatic water, whereas whole rock alSO values of harzburgites (mean = 6 . 1 % o ) are the result of low temperature modification by serpentinization from an estimated primary value of 5.4%o. There is no evidence in the Mantle Sequence for the high to moderate temperature hydrothermal seawater alteration that is so prevalent in the crustal sequence. (iii) Gregory & Taylor (1981) and Stakes et al. (1984) use oxygen isotope studies of the Semail ophiolite to model the hydrothermal alteration by circulating seawater in the Crustal Sequence. These authors have shown that alSO is enriched over primary magmatic values in the high-level gabbros but depleted in samples beneath that level. Gregory & Taylor (1981) showed that gabbros which appeared fresh in thin section have suffered sub-solidus isotopic exchanges with heated seawater and that in places this penetrated down to the Petrological Moho. Oxygen isotope data on the Semail dykes is discussed in Section 3.7.4.

3.12.2 Magnetic studies The palaeomagnetic investigation of the Semail Nappe (Shelton 1984) had the principal aim of determining the displacement of the ophiolite with respect to the Arabian continent. A substantial displacement would have required that the ophiolite originated in a major ocean basin. Subsidiary aims were to study any rotation of the ophiolite nappe as a whole and to look for differential rotation between blocks. In addition, measurements of magnetic property data allowed comparison with drilled and dredged oceanic samples and provided an estimate of the relative contribution of each part of the stratigraphy to the sea-floor magnetic signature. The Ibra block to the south of the Semail Gap is the subject of the only other

0 ! -10

I

~20

+40 I 0

Z'Sr(95)

I

,60 I +10

I +20

I +30

palaeomagnetic study (Luyendyk et al. 1982; Luyendyk & Day 1982) the interpretation of which is often at variance with the broader-based study of Shelton (1984). 3.12.2.1 Sheeted dyke results Both Shelton (1984) and Luyendyk et al. (1982) measured a scatter of NRM directions with the only notable concentration centred on the present day field direction. The measurements of Luyendyk et al. (1982) were not structurally corrected at this stage as they derive from a limited area within the Ibra block and have only small differences in attitude. The NRM directions derived in the more broadly based survey came from dykes with large differences in attitude and require structural correction (to give the palaeo-position or direction) in order to consider the distribution of results. Shelton (1984) found that such a correction led to increased dispersion reinforcing the evidence of the NRM field-position distribution that a viscous remanent magnetization (VRM) with a present-day field direction, has overprinted the original magnetization. In order to isolate a more original and useful 'stable' direction, demagnetization studies were carried out. After stepwise alternating field demagnetization Shelton (1984) identified the following types of magnetic behaviour: (i) most samples contain a random or northward oriented low coercivity component destroyed by cleaning fields of less than 100 Oe, (ii) commonly two components of mixed coercivity are found, one a viscous (VRM) or chemical (CRM) remanent magnetization, the other often directed to the south and east of the presentday direction and stable in demagnetizing fields of up to 600 Oe. (iii) Less often, a hard VRM or CRM stable above 1000 Oe is the sole component and (iv) occasional samples show total instability. Similar experiments by Luyendyk et al. (1982) attributed the unstable behaviour (median destructive field <100 Oe) to coarse-grained samples as they found that the MDF was inversely related to magnetic grain size and concentration, both favoured by slower cooling of magma. The remaining 70% of their sample set had MDFs ranging up to and above 1000 Oe and were classified as belonging to either a reversed or normal group. Sheiton (1984) found that the

I36

Chapter 3

higher stability magnetization corresponded to the northern grouping (assumed due to VRM), preventing the isolation of a more original vector possibly situated in the southern quadrant. Thermal demagnetization in both studies points to the presence of haematite as the most stable magnetic carrier but this appears to correspond to reversed polarity magnetization in the southern part of the ophiolite (Luyendyk et al. op. cit.) and normal polarity .magnetization in the more northerly biased sample set of Shelton (1984). The lower coercivity magnetization is identified as due to titanomagnetite and its low temperature (<200~ oxidation product titanomaghemite (Fe, Ti, D)304. Petrographically, oxidation is displayed as high temperature (>600~ trellis texture exsolution of ilmenite-rich lamellae (Plate 3.38). As haematite may form under a wide range of conditions ranging from high-temperature oxidation close to the spreading axis through to present day subaerial alteration, the significance of the sheeted dyke magnetic data is in doubt. The crustal section of the ophiolite has been subjected to an unusually varied series of sub-sea floor metamorphic events. These have produced a variety of magnetic minerals which would all be influenced by the same magnetic field (or its reversed equivalent if there were a polarity change within a few million years of the spreading event). However, these diverse minerals have varied stabilities which, allied w.ith the

range of grain size in the dykes, control subsequent low temperature alteration and produce a spread of directions which may correspond to situations ranging from the igneous cooling age to the present day. Luyendyk et al. (1982) discuss the possibility that the intermixing of dykes with normal, reversed or both polarities is due to remagnetization by later dyke intrusion. They find that this is unlikely as demonstrably late cross-cutting dykes also have dual polarity and consider that the polarity distribution must be affected by hydrothermal metamorphic events. 3.12.2.2 Layered gabbro results Shelton (1984) found that mean vectors for each gabbro site (location means), show no preference for the present day field direction and as such represent a more stable magnetization than the dykes. If a structural correction is made by restoring local layering to horizontal, the large amount of scatter is somewhat reduced suggesting that a significant proportion of the gabbros have been sampled close to the base of the magma chamber where layering is more likely to be sub-horizontal (Section 3.4.1.1). Structural correction using a mean dip for each block (derived from the plane of the Moho, sheeted dyke orientations and sea-floor sediment attitudes) produces approximately the same slight improvement. Nonetheless, the location means remain almost randomly scattered. Shelton (1984) showed, by stepwise alternating field demagnetization to isolate the stable magnetization vector, that the gabbros have high stability. Random low coercivity components were easily removed by fields up to 100 Oe above which two components of mixed coercivity were detected, one of which could be removed by fields of 800-1000 Oe. The field positions of demagnetized specimens tended to originate in the south-east quadrant if stable, or if less stable, to move towards that quadrant. The majority of the gabbros have very sparse opaques. A large proportion of the better defined grains analyse as a variety of Fe, Cu and Ni sulphides. Sulphides are also found in the cracks in serpentinized olivines but the most common mineral in this situation is magnetite. Shelton (1984) identified the gabbro high stability carrier as haematite and tentatively attributed the 800-1000 Oe coercivity component to magnetite resulting from the alteration of olivines and pyroxenes. Luyendyk & Day (1982) positively identify magnetite from thermal demagnetization, and state that it carries a normal polarity magnetization. The high between-site dispersion in the stable gabbro palaeo-positions observed by Shelton (1984) is best accounted for by postulating two polarities in the magnetization. During magnetization the vectors approach their stable positions from two opposing directions and classifying the results by block (Fig. 3.73) shows a demarcation between the southern and northern parts of the ophiolite. Results from the southern block (8, 10 and 12 in Fig. 3,72) agree with the 50 Ibra sites measured by Luyendyk & Day (1982) which cluster around declination 340 ~ with low positive and negative inclinations (I = 7.6~ 3.12.2.3 Polarity reversals and timing of magnetization

Plate 3.38. X-ray map of Fe-Ti oxide grain in high Ti sheeted dyke.

Upper image showing titanium, lower silica.

It appears that dykes in the north of Oman are of normal polarity whilst those in the south have mixed polarity, by contrast, gabbros in the north of Oman have reversed polarity whilst those from the south are normal.

The Semail Ophiolite

I37

I

0

58~ 26ON --

BLOCK NAMES 1 Khawr Fakkan 2 Aswad 3 Fizh 4 Salahi 5 West Jizi 6 Sarami 7 Wuqbah 8 Haylayn 9 Muqniyat 10 Rustaq 11 Bahia 12 Ibra 13 Muscat 14 Sumeini

Shah

24 ~

AI Arid ~,

0

,

50Km

O

,

-.~ 180

I

l

Fig. 3.72. Outcrop extent and structural units of the Oman ophiolite showing the blocks into which the Semail Nappe can be divided. After Rothery (1982).

Fig. 3.73. Stable gabbro palaeo positions (corrected using block tilts) labelled by block number. O = positive inclination, 9 = negative. 9 = present day axial dipole approximation. 3 = Fizh, 4 = Salahi, 6 = Sarami and Ghuzayn, 8 = Haylayn, 10 = Rustag and Ibra. Key to blocks in Fig. 3.72. From Shelton (1984).

There are four possible scenarios, presented in Fig. 3.74, that would account for the switch in polarities between the northern (Fizh) and southern (Ibra) blocks and the difference in polarity between gabbros and dykes of the same block. They depend on (1) the relative ages of the Fizh and Ibra blocks, (2) the order in which gabbros and dykes acquire their dominant magnetization and (3) whether the system was subjected to a short normal event in a reversed period or a reversal in a normal period. These points are discussed in detail in Shelton (1984) and can be summarized as follows:

1 The relative ages of the blocks are difficult to ascertain, the Sm-Nd isochrons (McCulloch et al. 1980) are too inaccurate to distinguish between blocks and the U-Pb analyses (Tilton et al. 1981) show that the Semail Nappe has a virtually uniform age along its length. 2 The relative timing of the dominant magnetization is based on the demagnetization conclusions of Shelton (1984), Luyendyk et al. (1982) and Luyendyk & Day (1982). The sheeted dykes appear to show a mirror image of assigned magnetic carriers between the Fizh and Ibra

Fizh dykes = normal

Fizh older than Ibra dykes reset during hydrothermal alteration

Ibra older than Fizh dykes reset during hydrothermal alteration

Fizh older than Ibra gabbros slower to cool

Fizh spreading. Gabbros and dykes -normal

iiiiNiiiiil Ibra spreading. ii!iiii!iiil GabbrOSnormal and dykes=

Fizh gabbros cool through Curie temperatures

Ibra gabbros coot = normal

Ibra spreading. Gabbros and dykes :

Fizh spreading. Gabbros and dykes : reversed iiiiiiiiiiii hydrothermal Fizh dykes reset byalteration = normal

Ibra block spreading.

Fizh dykes reset by ~i:-i::::: Ibra gabbros hydrothermal alteration iiiiilNiii cool : normal : normal

Ibra dykes reset by R I hydrothermal alteration = mixed

Ibra block spreading. Dykes = reversed (+ normal)

Ibra older than Fizh gabbros slower to cool

~1

Ibra dykes reset by hydrothermal alteration : mixed

Fizh dykes : normal

Fizh gabbros cool through Curie temperatures

::::::::::::::

(a)

(b)

(c)

(d)

Fig. 3.74. Four alternatives to account for the dominant magnetization differences between the northern and southern parts of the ophiolite. The. diagrams speculate on the order and nature of magnetization processes with time (top to bottom with increasing time). The alternatives depend on which block is the oldest, whether the dykes acquire permanent magnetization before the gabbros and the polarity scheme between ophiolite formation and emplacement.

I38

Chapter 3 areas. In the interpretation of the Ibra dykes, haematite carries reversed and magnetite + maghemite carry normal polarity but in the case of the Fizh dykes, haematite + maghemite carry normal and magnetite carries reversed polarity. It is clearly crucial whether the haematite is the product of low (secondary) or high temperature (primary) oxidation. Luyendyk et al. (op. cit.) hold that the normal polarity is carried by high temperature magnetite and that the haematite is secondary. Shelton (1984) also suggests that haematization follows maghemitization and is therefore secondary. Of the six gabbro samples thermally demagnetized by Luydendyk & Day (1982) only magnetite blocking temperatures of 550 to 580~ were found. Luyendyk & Day (op. cir.) do not recognize haematite as a carrier and there is insufficient evidence to determine which of the gabbro carriers recognized by Shelton (1984) is the earliest. There is, at present, no obvious means of determining whether the time taken for the layered gabbros to cool through their Curie points is longer or shorter than the time taken for the hydrothermal alteration in the dykes to cease resetting their magnetization. The interpretation also depends upon the magnetic history of the period in which the ophiolite was created and emplaced. The radiometric and palaeontological age data constrains this to the period 100 to 75 Ma most of which falls within the "Cretaceous quiet zone", so called because of the absence of any globally traceable magnetic anomalies. There appears to be uniformly normal polarity between reversals M0 (118.2-119 Ma) and 33r (78.5-82.9 Ma) according to Harland et al. (1982) who exclude reversals or excursions of less than 0.03 Ma. The evidence for reversals in the period of interest is very slender (Lowrie et al. 1980).

To summarize the foregoing with reference to the options shown in Fig. 3.74, the dyke magnetic data favour options (B) and (C). If the Fizh block is older than Ibra, options (A.) and (C) apply and for the time of formation (Cenomanian) it seems more likely that options (A) and (B) represent the polarity history. This only eliminates option (D); the others are equally possible if the data are of equal validity. The polarity decision is subjective, but the alternative (C+D) requires two reversals in what appears to be a normal period rather than one change of polarity. The evidence concerning the relative ages of the blocks must be considered the weakest, and it is here proposed that the admittedly uncertain data of the Sm-Nd dating which makes Ibra older than Fizh is correct. This then agrees with both the dyke evidence and the favoured polarity sequence to make option (B) the preferred model. This does not mean that the position of the ophiolite to the east of the spreading centre is incorrect; rather it suggests that the present outcrop distribution does not correlate with distance from the palaeo-ridge, the present spatial relationship of the blocks being due to emplacement and subsequent tectonics. The isolated reversal responsible for resetting the Ibra dykes and giving a southerly polarity to the Fizh gabbros could not have been very long-lasting or it would have shown on more late Cretaceous magnetostratigraphies. As it appears that the present relative positions of the ophiolite blocks are only related to the axis direction and not the spreading direction, the maximum width of spreading can be as little as that contained in the widest block. If account is taken of the ophiolite buried beneath the Batinah and a correction is made for block

tilt, a maximum width of 50 km is obtained. At an intermediate spreading rate of 3 cm/a, only 1.7 Ma is required to produce this width of ocean floor. This is long enough to reset the Ibra dykes by hydrothermal alteration but short enough to lie within the spread of plagiogranite ages. 3.12.2.4 Rotation and relative displacement of the ophiolite Smewing (1980b) noted the presence of a 30 km belt south of Wadi Hatta in which dykes striking at 115~ predominate. He suggested two possible transform zone situations which could account for the sigmoidal strike pattern from north to south. One involves an acute ridge/transform angle in which dykes are rotated by left-lateral shear into the plane of the transform, the other requires an obtuse ridge/transform angle and has the dykes injected parallel to the qeaky' transform direction. Shelton (1984) sampled dykes across this zone and despite the obscuring VRM imprint found no evidence for relative rotation of the dykes in the Wadi Ragmi transform zone area. The dykes appear to have been emplaced in their present relative orientations as envisaged in the 'leaky' transform model. The palaeomagnetic data for the sheeted dykes is not considered sufficiently 'clean* to resolve either relative latitudinal movement or relative rotation. However the Ibra gabbros of Luyendyk & Day (1982) have a mean declination of 339~ representing a clockwise rotation of only 3~ from the Upper Cretaceous African reference pole of Gough et al. (1964). Restoration of the anticlockwise rotation of Arabia with respect to Africa due to the opening of the Red Sea (approximately 7~) would increase the rotation recorded by the southern gabbros to 10~ W. Luyendyk et al. (1982) and Luyendyk & Day (1982) were fortunate in that their dyke data, after separation into northerly and southerly vectors, passed the reversal test. The circle of 95% confidence of the normal and reversed means overlapped. The more precise Ibra gabbro mean also falls within the combined Ibra dyke circle of confidence. With these encouraging data they went on to derive a palaeomagnetic pole which is far-sided (285~ 60.5~ but near-sided in relation to a reference pole for Upper Cretaceous Africa (259~ 62.0~ from Gough et al. (1964). Their interpretation is that, after relative southward movement of the ophiolite onto the Arabian foreland, the emplaced ophiolite has moved northward with Afro-Arabia. This is, in general, entirely consistent with the known path of Afro-Arabia and the geology of Oman. In detail there are inconsistencies in that Luyendyk et al. (op. cir.) quote 13.1 + 11~ as the southward displacement between the ophiolite and Arabia. This is equivalent to 1450 km N-S or over 2000 km in a NE to SW emplacement direction. This amount of initial separation would require unrealistically high spreading rates to accomplish closure. It is considered that an initial separation equivalent to 3 to 5~ of latitude, as allowed in the quoted error, is more appropriate. The gabbro data of both Luyendyk & Day (1982) and Shelton (1984), despite the problems of defining palaeo-horizontal within the unit, give uniformly low inclinations in accordance with an equatorial site of origin. The southerly blocks (including Luyendyk & Day 1982) have a mean inclination of +6 ~ corresponding to a palaeolatitude of 3~ From the reconstructions of Smith & Briden (1977) and Firstbrook et al. (1979) a reference locality taken as Sohar (present latitude 24.3~ longitude 56.3~ would have been at latitude - 3 . 0 to -3.5 ~ (S) at 100 Ma. By 80 Ma the locality had crossed the equator to

The Semail Ophiolite +3.0 or 4.0~ The southern block gabbros thus suggest 3 to 5 ~ of southerly displacement of the ophiolite relative to Arabia. The northern block gabbros resist such interpretation for, although their inclinations are low, they are southerly and positive (88% of those in the SE quadrant). This would require either large rotations (e.g. 130 ~) or, more possibly, that the inclinations are in error by up to 20 ~ In summary, the gabbro data indicate that the southward movement of the ophiolite relative to Arabia was small (<500 km N to S) and that there has been little overall rotation of the ophiolite during that emplacement. This limited displacement lends further support to a marginal basin origin for the ophiolite.

I39 nT

Km

0-

Sources of marine magnetic anomalies

Comparison of magnetic data and relative contribution to a marine anomaly measured 5kin above. (A) Luyendyk & Day (1982), Ibra area, (B) Shelton (1984). Total = magnetisation per metre of crossstrike in amperes. Ibra upper gabbro unit includes isotropic gabbros and upper half of layered gabbro unit. + = geometric mean, * = based on lava At = 2km, dykes At = 2kin. The anomaly relief at sea-level is calculated according to the method of Luyendyk & Day (op. cit.). Az = NRM At/(z + At/2) where z is the depth to the top of the layer, Az is the vertical magnetic component and NRM is the volume magnetization (which has been doubled to simulate the contrast of adjacent crust with opposite magnetization).

Table 3.32.

NRM A m -I

Qn

At rn

Total A

1.51 0.73 1.03 0.27

>1.6 0.8 2.8 3.3

1000 1000 3700 2000

1510 730 3811 540

Anomaly relief nT

A

Ibra lavas Ibra dykes Upper gabbros Lower cumulates

110 45 174 18

B

Semail lavas

3.09+

1.2 1500-2500 4635-7725

412"

Semail dykes

0.59+

0.5 1500--2000 885-1180

59*

Semail gabbros

0.33+

3.1

43*

3500 1155

0-

2-

4-

4-

111111/11145,,30 o,

6-

i i l 412 (8O%)

~ ')'(-',q

8 - i~'.:[}::~4i~

The contribution of a layer of oceanic crust to an anomaly measured at sea-level will largely depend on three factors: (1) depth to the layer mid-point (z + At/2) and thickness of layer At, (2) the volume intensity of N R M and coherence of direction of magnetization and (3) timing of the magnetization. There are several possible pitfalls when constructing a multilayer source for marine magnetic anomalies from ophiolite data. The most serious is that ophiolites are atypical ocean crust in that they were not subducted and are perhaps more representative of marginal basins than deep oceans. The relative thickness of the ophiolite units varies within and between ophiolites (Moores & Jackson 1974; Moores 1982), but there are available representative seismic averages from the oceans. If an acceptable structural model can be found there remains the possibility that the magnetic parameters from the ophiolites will not apply to 'normal' oceanic crust. In the Semail nappe, there are certainly post-emplacement influences on the magnetization, e.g. the V R M in the dykes and low temperature serpentinization in the gabbros. There will also be small

nT

Km

2-

'-

3.12.2.5

Shelton 1984

Luyendyk & Day 1982

172(50~

8",', 5,2, L{~S,

10-

10.

.

.

.

9 "- .,,

.

18(5%)

12-

43 (8%)

12-

,,

345 total

514 total

Fig. 3.75. Comparison of contributions to a hypothetical marine anomaly based on the interpretations of (a) Luyendyk & Day (1982) and (b) Shelton (1984).

differences in chemistry due to the supra-subduction zone setting and large differences in the history of hydrothermal alteration (principally restricted to the dyke and lava units). There is now abundant data from ocean drilling of lavas and a little from the dykes (Anderson et al. 1982). Thus the ophiolite data of most interest is that from the sheeted dykes and cumulates. Table 3.32 includes the interpretation of the Ibra section as a possible anomaly source from Luyendyk & Day (1982). They comment that their dispersed dyke directions would substantially diminish the contribution of the dyke layer. As this dispersion results from a polarity change it offers no challenge to the principles of anomaly modelling. The Ibra section has an unusually thick gabbro stratigraphy and this has been split into two sub-units based on their magnetic properties. The upper unit includes the isotropic gabbros, plagiogranites and all but the lowest 2 km of the layered gabbros. This sub-unit appears to provide a more substantial contribution than the extrusives to the surface anomaly. In the present work only the high-level (isotropic) gabbros were found to have primary oxides. These are generally restricted to a few hundred metres thickness and as such have a negligible effect. The remainder of the gabbros appeared uniformly sparse in magnetic minerals and have here been treated as a single unit. Figure 3.75 illustrates the alternatives of (a) Luyendyk & Day (1982) from the Ibra area and (b) a synthesis from the remainder of the ophiolite (Shelton 1984). The latter results are less controversial than those of Luyendyk & Day (op. cir.) in that they suggest that the extrusives are responsible for the major portion of marine magnetic anomalies.

Chapter 4 Ophiolite Detachment, Emplacement and Subsequent Deformation 4.1 I n t r o d u c t i o n In Chapter 3, the Semail ophiolite is described and the conclusion reached that it is oceanic lithosphere formed on the western flank of a N-S constructive margin that was above a NNE-inclined subduction zone that lay some 300-750 km from the Arabian (Oman) continental margin. Here, we are concerned with (i) the detachment of the ophiolite from its position as in situ oceanic lithosphere, (ii) how it became emplaced on the eastern passive margin of the Arabian continent and (iii) what happened to it subsequently. The evidence that identifies the time constraints and the structural, magmatic and metamorphic processes involved is presented and the deductions therefrom collated into a dynamic model. A temporal sequence will be followed so that, after reviewing the evidence for the supra-subduction site and origin of the ophiolite, data relative to detachment processes are presented before those concerning the subsequent emplacement and deformation. As products and processes falling within this detachmentemplacement phase are not presented elsewhere, they are fully described here. Uranium-lead isotopic ages on plagiogranites (Tilton et al. 1981) indicate that the Semail ophiolite was formed about 95 Ma ago. Most plagiogranite ages are between 93.5-95.9 Ma, although older ages of 97.0 and 97.9 Ma have been determined and seem to be real. On the basis of the spatial distribution of these isotopic ages Tilton et al. (op. cir.) suggest that the ophiolite formed on the western flank of the N-S spreading axis: this is supported by the statistical predominance of west-side oneway chills on the sheeted dykes of the northern mountains (Lippard 1983) and the general eastward inclination of layered gabbros (e.g. Pallister & Hopson 1981). Lanphere (1981) identified that the amphibolites of the subophiolite metamorphic sheet have a weighted mean 4~ total fusion age of 90.0 _+ 3.0 Ma. These amphibolites are thought (Lanphere op. cit.; Searle & Malpas 1981) to represent metamorphic products welded to the base of the ophiolite as it was detached from its in situ position as oceanic lithosphere. The oldest plagiogranite (97.9 Ma) and the youngest possible metamorphic (87.0 Ma) ages together indicate a maximum of 10.9 Ma between formation and detachment. This in turn implies detachment reasonably close to the spreading axis. Just how close is difficult to determine but the minimum across strike width of the ophiolite, including that part beneath the coastal plain, is about 50 km. Previous estimates of the across strike width of the ophiolite vary from 130 km.(Tilton et al. 1981) to 275 km (Pallister 1981) based on the present spatial distribution of ophiolite blocks whose present arrangement is the result of the final emplacement processes and may bear no relationship to the original spreading situation. Intra-oceanic thrusting led to the deformation of the base of that part of the oceanic lithosphere that was to become the Semail Nappe. Cataclastic and mylonitic textures were produced in the basal peridotites (Banded Unit) and a subjacent Metamorphic Sheet was produced by the thermodynamic alteration of the overridden oceanic crust. There are therefore two aspects of the detachment, the metamorphic processes and products and the structural evidence identifying detachment directions.

The major problem here is that the effects of this early detachment event are overprinted by those produced during the subsequent emplacement of the ophiolite onto the Arabian continental margin. The Metamorphic Sheet beneath the ophiolite, in which there is an inverse thermal gradient from high-temperature amphibolites immediately subjacent to the peridotite to lower temperature greenschist facies rocks further therefrom, has been mapped as a single sheet. However, detailed study reveals that it can be divided into an upper, amphibolite facies unit whose structure and geochemistry suggest was produced during intra-oceanic detachment, and a lower, mainly greenschist facies, unit that was added during subsequent emplacement. So, after a general description of the Metamorphic Sheet as a whole, only that part relevant to intra-oceanic detachment processes is described and discussed here; the lower, greenschist facies unit is described in Section 4.3.1. Similarly, the structure of the peridotite Banded Unit contains three elements: those produced during constructive margin processes (Section 3.2.4), those produced during detachmelat discussed here, and those superimposed later (Section 4.3).

4.2 Detachment As identified above there are two rock units, the Metamorphic Sheet and the Peridotite Banded Unit, that provide evidence on detachment.

4.2.1 The metamorphic sheet Polyphase-deformed, schistose metamorphic rocks are present throughout the Oman mountains at or near to the base of the Semail Nappe (Fig. 4.1, Plate 4.1) often forming a separate thrust sheet or as blocks in tectonic melanges (Searle 1980; Searle & Malpas 1980, 1982; Ghent & Stout 1981). Where absent, the metamorphic rocks have been removed by later thrusting or faulting. Glennie et al. (1974) showed that the metamorphic rocks occur in tectonic contact with the Semail ophiolite and recognized two distinct units: an upper group of "polymetamorphic" rocks composed of amphibolites, quartzites and calc-silicate marbles formed by an early high temperature "hornblende granulite to pyroxene hornfels" facies metamorphism. This, they indicated, was overprinted by a greenschist facies event. They also identified a lower group of "monometamorphic" quartzites, phyllites, mica and chlorite-epidote schists, some of which show clear evidence of sedimentary or volcanic protoliths, that had only suffered a greenschist facies metamorphism. Glennie et al. (op. cit.) suggested that the greenschist facies event in the two units was the same and occurred during late Cretaceous nappe emplacement. Allemann 8: Peters (1972), using the same metamorphic terminology as Glennie et al. (1974), suggested that only the monometamorphic greenschist facies was of late Cretaceous age. The late Cretaceous ages of 87 - 85 + 5 Ma they obtained on the amphibolites were ascribed to the greenschist facies retrogression of the rocks. It is relevant to note here that the greenschist facies rocks of the Saih Hatat area south of Muscat which Glennie et al. (1974) thought were part of the pre-

Ophiolite Detachment, Emplacement and Subsequent Deformation N

NW

SE

1 DIBBA ZONE J, Qamar ,

Musafi

W. Ham ~.,~...- ~ , _ ~.

i

2 SUMEINI "

Uwaynah Masafi

I4t

.~_ ~~I

DIBBA

Semail Nappe Amphibolites [ T ~ Greenschists

idhn 1 "i~N'~{

3 ASJUDI

Haybi complex Melange (olistostrome)t~_~ ] Haybi volcanics ~ Hawasina

4 HAYBI

o fParautochthonous ~

o,oc ,.ooous

2Wadi Hatta

Pre-Permian basement F ~ / j

5 WADI HAWASlNA SUMEINI 2

\ -\ ~:~i~ ! - 84

,

i

6 WADI TAYIN-MUSCAT

blueschists \

\

ASJUDI 3

0

50 Km

~ _ . -....

IFAH \

Semail Nappe

~ Fig. 4.1.

Sub-ophiolite metamorphic "sole"

........

i!iii!iiii!-i,i ~/

. , . .....

.~i!~-il/~~i)ii I

Saih Hatat "blueschists"

Outcrops and schematic cross-sections showing the distribution of sub-ophiolitic rocks.

Plate 4.1. Sumeini area metamorphic rocks. Masses of exotic limestone are thrust over Haybi Volcanics which make up the fore and middle ground.

Chapter 4

I42

Table 4.1. Radiometric ages of Oman metamorphic rocks. Rock type

Locality

Method

Amphibolite

Wadi Tayin

Ar4"-3~hornblende

"Partial melt" in amphibolite Amphibolite ("retrograded") Quartz-mica vein (greenschist facies) Greenschist

Sumeini

K-Ar hornblende

96.5-+5.7 89.2-+2.0 89_+3

U.A.E.

K-Ar biotite

85+5

U.A.E.

K-Ar muscovite

87_+5

Wadi Tayin

Ar4~-3'~whole rock

Mica schist (blueschist facies)

As Sifar

K-Ar phengite

Permian basement are now known to be of mid to late Cretaceous ages (100 to 70 Ma) (see Table 4.1) and belong to a regional high pressure blueschist to greenschist metamorphism that affects the lower allochthonous and upper autochthonous units of the area; this will be described and discussed in Section 4.3.1. The Metamorphic Sheet is present throughout the length of the mountains from the Dibba-Masafi area in the north to Wadi Tayin in the southeast (Fig. 4.1). It occupies a consistent tectonic position at the base of the Semail Nappe, often occurring where there is a thick overlying Mantle Sequence. The thrust contact with the peridotites is commonly disturbed by downward protrusions of serpentinite into the metamorphic rocks. This can occur on all scales from local protrusion only a few metres across to areas in which the whole Metamorphic Sheet is broken up into a tectonic melange with serpentinite matrix. In the northern area, the metamorphic rocks are best developed and most intact in the Sumeini area where there is a steep, inverted metamorphic gradient from upper amphibolite

-~-

.

.

.

.

Banded amphibolites Cpx and garnet -bearing in upper part Some calc-silicate, marble and quartzite bands

.

JI|NI~\

Local protrusions of / s e r p e n t i n i t e along thrust

Sepentinized harzburgites and dunites (Banded Unit) -.~-Semail Thrust - -

- ,

v

Cpx-bearing quartzites Banded muscovite and piemontite-bearing quartzites

----_ .~_------~ .-~---_--: ~---~-Z-: .-_------~ ......

Sericite quartz mylonites '"-and phyllonites along thrust ----j~_-----~ Amphibolites

~

'-'~vtXl\

x.r~l

X..

Quartz-mica schists with layers of graphitic schist Strong mylonite fabrics and quartz veining (pre-FO (in parts)

Garnet amphibolites with pods and veins of plag-hbl gneiss (partial melt) Medium grained amphibolites. becoming coarser-grained upwards

Quartzite mylonites

Thin bedded quartzite quartz-mica schist and some marbles, and biotite schists

,,v,

tO0

a~-

,~ , ~ , o ~ ~ ,

Weakly metamorphosed (lower greenschist facies) olistostrome, Grades into unmetamorphosed rocks away from thrust

m

1o

Fig. 4.2. Representative sections through the metamorphic sheet in the Sumeini area (after Searle 1980).

Age ( M a )

76.0_+2.2 70.8+8.6 101_+4 l(X)+4 80+2

Reference

Lanphere1981 Searle (1980) Analyst D. C. Rex Allemann & Peters (1972) Allemann & Peters (1972) Lanphere (1981) Lippard (1983) Analyst D. C. Rex

facies rocks (up to 250 m thick) at the top closest to the peridotite contact, to epidote amphibolite and upper greenschist facies rocks (up to 300 m thick) at the base (Fig. 4.2). There is a tectonic contact between the amphibolite and greenschist facies rocks and occasionally the two facies are tectonically interleaved. Elsewhere, where tectonic disruption is even more severe, the metamorphic rocks occur either as imbricated thrust slices within the Haybi Complex, with unmetamorphosed rocks above and below (as on Jebel Ghawil) or as detached blocks in serpentinite melange (the Basal Serpentinite), as in the Wadi Ahin-Haybi and Asjudi areas. In addition, these metamorphic rocks occasionally occur as tectonic inclusions in the multi-component Haybi and Batinah melanges. Plate 4.2 shows a selection of the metamorphic facies and textures taken from Smewing et al. (1982). 4.2.1.1 Amphibolite facies rocks Dark coloured, banded amphibolites predominate in the upper part of the Metamorphic Sheet where they have a maximum thickness of about 300 m. They are generally finely laminated, schistose rocks consisting largely of brown-green hornblende (Mg'48-66, AI20~ 11-13%, TiO2 1-1.5%) with lesser amounts of plagioclase (An6~_4o), quartz and sphene. Grain sizes range from 0.2 to 6 mm and the rocks often have a pronounced banding owing to alternating 5 to 15 mm thick mafic and felsic layers. The coarsest grained rocks are found at the top of the sequence and have porphyroblasts of colourless to green salitic pyroxene (Wo48.sEn32_29Fs2o_23,4--5.5 wt% A1203) and red garnet (Alm~0Gr33Pr14Sp3). The hornblendes show a trend towards higher Ti and Fe/Mg contents towards the top of the sequence in the direction of higher metamorphic grade (Searle 1980). Most of the amphibolites show some evidence of secondary recrystallization and granulation, leading to a reduction in grain size, accompanied by retrogression to greenschist facies assemblages with brown-green hornblende replaced by actinolite and biotite, garnet by chlorite and calcic plagioclase by albite and sericite. Some of the amphibolites, that are cut by fine grained rodingite 'veins' composed of hydrogrossular, prehnite, clinozoisite and carbonate, have high CaO and low SiO2, A1203, and Na20 contents (Table 4.2). The highest grade amphibolites are often gneissic in appearance and contain deformed layers, augen and pods of coarse-

Ophiolite Detachment, Emplacement and Subsequent Deformation

I43

Plate 4.2. (a) Amphibolite containing brown hornblende, quartz, plagioclase and opaques. Two generations of amphibole: elongate prisms forming the main fabric and (later?) equidimensional poikiloblastic grains. (b) Epidote amphibolite containing epidote, hornblende, quartz, plagioclase, chlorite and muscovite. Orientated grains of hornblende and epidote with interstitial granular quartz and retrogressive chlorite. The quartz contains numerous inclusions of sericite. (c) Metaquartzite [XN] containing quartz, muscovite and biotite (trace). Amphibolite facies quartzite composed largely of strained quartz showing a strong parallel orientation. (d) Greenschist composed of actinolite, chlorite, epidote, opaque oxides, sphene, quartz and carbonate. Composed largely of granular epidote and actinolite with patches of fibrous chlorite and some granular quartz. (e) Garnet mica schist [XN] composed of quartz, muscovite, biotite and garnet porphyroblasts with quartz inclusion trails. Matrix of highly strained interlocking quartz anhedra with folded trains of minute mica flakes. (f) Similar to (e). Enlargement of an elongate folded garnet porphyroblast. Taken from Sinewing et al. (1982).

Chapter 4

x44 Table 4.2. Amphibolite compositions.

0M586

0M765

0M1533

0M1590

51.7 1.25 14.6 4.91 3.98 0.16 5.44 10.85 4.17 0.20 0.07 1.82

0M1592

0M1576

Mean

Range

47.0 1.45 14.1 3.95 7.28 0.18 8.27 12.73 2.10 1.05 0.14 1.99

44.1 1.20 9.8 4.88 7.61 0.22 9.37 18.25 0.96 0.18 0.07 2.70

48.1 1.30 14.3

41.6 -54.5 0.73- 1.72 12.1 -15.6

Major oxides

SiO2 TiO2 A1203 Fe203 FeO MnO MgO CaO Na20 K20 P2Os LOI Total

54.52 1.10 15.60 10.04"

47.59 1.48 14.19 9.70*

0.16 5.70 7.99 2.93 0.55 0.13 1.22

0.14 6.74 13.67 3.24 0.56 0.29 2.34

51.6 0.63 15.4 3.46 5.95 0.17 6.95 7.71 4.44 0.80 0.05 2.03

100.09

100.19

99.20

99.15

100.25

99.30

253

159

211

137 9 168 205 13 43

67 3 314 226 26 76

108 17 208 318 23 69

10.14" 0.18 6.98 12.65 2.72 0.64 0.10 1.85

8.40-12.32 0.13- 0.23 5.44- 8.29 9.99-13.67 1.48- 2.17 0.20- 1.51 0.05- 0.29 1.34- 2.34

120

168

80-253

92 4 38 287 26 72

81 9 252 304 26 68

64-137 3- 17 112-446 182-457 13- 40 26-123

Trace elements

Ba Cr Nb Ni Rb Sr V Y Zr

13 9 205

6 460

26 59

30 106

Rare Earths

La Ce Nd Sm Eu Gd Tb Tm Yb Lu Hf Th Ta Sc Co

5.9 (18.0) 8.75 (10.1) 7.62 (12.1) 2.8 (13.8) 1.03 (13.4) 3.96 (14.3) 0.71 (13.7) 0.4 (11.8) 2.76 (12.5) 0.46 (13.5) 2.02 0.3 0.26 35.5 36.3

13.5 (41.2) 25.6 (26.7) 17.3 (27.5) 4.2 (20.7) 1.56 (20.3) 5.68 (20.6) 0.84 (16.1) 0.47 (13.8) 2.8 (12.7) 0.49 (14.4)

OM586, OM765 Fine grained amphibolites, Wadi Jizi OM1533 } OM1590 "Fresh" amphibolites, Sumeini. Garnet-bearing with veins and patches of "partial melt". OM1592 OM1576 Rodingitized amphibolite, Sumeini. Mean and ranges of 10 amphibolites from Sumeini, Wadi Jizi and Hawasina sections (excludes rodingitized and other altered samples). * Total iron as FeO ( ) Chondrite normalized values. Data taken from Searle (1980)

grained, granoblastic, leucocratic rock c o m p o s e d largely of albite (Anlo-3) and up to 10% h o r n b l e n d e and clinopyroxene. T h e s e acid segregations, a l t h o u g h themselves d e f o r m e d , are often c o n c e n t r a t e d in fold hinges. The rocks were i n t e r p r e t e d by Searle (1980) and Searle & Malpas (1980) as partial melts f o r m e d at the p e a k of m e t a m o r p h i s m . This origin is supported by their g r a d a t i o n a l contacts with the host amphibolites and their tonalitic compositions. T h e a m p h i b o l i t e s are locally i n t e r b e d d e d with calc-silicate and quartzite layers. T h e calc-silicates contain calcite, grossular garnet, diopside and pargasitic h o r n b l e n d e (Fig. 4.3, Plate 4.3b), with rare woilastonite and orthoclase. T h e y are often

b a n d e d on a centimetre-scale with alternating layers richer in silicates and carbonates. T h e quartzites are c o m p o s e d largely of fine grained strained quartz with accessory h y p e r s t h e n e , diopside and muscovite. As the m a j o r i t y of the a m p h i b o l i t e facies rocks are metabasaltic in c o m p o s i t i o n (Table 4.2) and are t a k e n to r e p r e s e n t m e t a m o r p h o s e d lavas or tufts; the calcsilicates and quartzites are believed to be i n t e r b e d d e d calcareous and siliceous sediments. T h e mafic a m p h i b o l i t e s have a "stable" trace e l e m e n t composition (Zr/Y 1.5-3.5, Y/Nb 2.3-5.3, L a / Y b ( N ) 1.44-3.24) characteristic of ocean-floor basalts i n t e r m e d i a t e in character b e t w e e n M O R B and within-plate tholeiite. Searle & M a l p a s

Ophiolite Detachment, Emplacement and Subsequent Deformation 1,4

I45

--

Glaucophane {!i

["-] ,~kARGASITE 1.0

Crossite II 99 I I

--

\ /~

LP ~ H P

/ \

\

1 t /n HI

0.6-

I

el mJ

il

/ ~o2N~ / X

~r

A-~-" ~ O o o)

Hornblende

17,1 <

i5 /@ ! I iI [] !

l~

X

\ Actinolite I

0.2--

EDEN ITE

I

1

&-----~ 9

ACTINOLITE

ITI

6.0

~[~)

I

7.0

Si

FR 8.0

[] Glaucophane epidote schists 9 Amphibolites 9 Partial melts in amphibolites 9 Amphibolite facies marbles --"5 Kb line" Separating low and high pressure amphibolites

Fig. 4.3. Composition of metamorphic amphiboles (for discussion of glaucophane epidote schists see Section 4.3.1). (1982) considered that they represent oceanic basalts of Jurassic or early Cretaceous age that were largely consumed by midCretaceous subduction. However, the trace element contents, including R E E patterns (Fig. 4.4), of the amphibolites are so similar to some of the transitional-tholeiitic basalts from the 50

-

::::::::::::::::::::::

O M :::::::::::::::::::::::::~, ================================== 766 .::::::::::::: ................x :::::::::::::::::::::::::::::::::::::

iiiiiliiiii~i~ii;ii!iiiiiiiii!i!iii!iiiiiii~. :::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::: ..::::::.~

O M :::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::

586 Plate 4,3. (a) Folded quartz-mica schist. Field of view - 4 mm. (b) Pargasite marble (pargasitic hornblende and calcite). Field of view 4 ram. (c) Partial melt in amphibolite. Hornblende-albitized plagioclase-magnetite, granoblastic texture. The feldspar is cut by prehnite veins. Field of view - 4 mm.

10

~

I

I

La Ce

Range in REE compositions for Haybi votcanics

I Nd

I

I

I

I

Sm Eu Gd Tb

I

I

I

Tm Yb Lu

Fig. 4.4. REE/chondrite normalized plot for amphibolites compared to the range of composition for the Haybi Volcanics (data from Searle & Malpas 1982).

Upper Triassic Haybi volcanics (Section 2.2.3) that it is difficult to discount these as possible amphibolite protoliths as originally suggested by Searle (1980). This has important palaeo-tectonic implications because it implies that the midCretaceous ocean-floor spreading which produced the Semail ophiolite may well have been initiated within an older, late Triassic, oceanic lithosphere. The highest grade of metamorphism in the Metamorphic Sheet is represented by the assemblage hornblende-diopsidic augite-almandine garnet-labradorite in metabasic rocks which

Chapter 4

I46

is typical for the upper amphibolite-lower granulite facies transition. Using the coexisting garnet and clinopyroxene geothermometer methods of Ellis & Green (1979) and Saxena (1979), that take into account the Ca content of the garnet (Table 4.3), the Oman amphibolites gave temperatures of 800 + 50~ Ghent & Stout (1981), using cpx-gt geothermometry and the Na content of the pyroxene as a possible

b a r o m e t e r , s u g g e s t p e a k m e t a m o r p h i c c o n d i t i o n s o f 750~ a n d 4 - 5 k b f o r t h e f o r m a t i o n o f t h e h i g h e s t g r a d e a m p h i b o l i t e s in t h e W a d i T a y i n a r e a . F l e e t & B a r r e t t (1978) d e f i n e d a l o w ( < 5 k b ) - h i g h p r e s s u r e d i v i d e b a s e d o n t h e AI~V/AI v~ r a t i o s o f a m p h i b o l e s . T h e h o r n b l e n d e s in t h e O m a n a m p h i b o l i t e s h a v e AI~V/AI vl = 3.0 a n d t h u s fall in t h e l o w p r e s s u r e field. T h e y also f o l l o w a l o w p r e s s u r e t r e n d o n t h e AI vl vs Si p l o t o f R a a s e

Table 4.3. Compositions of coexisting clinopyroxenes and garnets from Oman amphibolites. OM17,53 cpx(4)

SiO2

gt(4)

0M1593 cpx(2)

gt(4)

0M1592 cpx(2)

gt(4)

0M1649 cpx( 1)*

gt(3)

Cr20.~

48.92 0.46 4.91 10.25 0.17 10.51 22.04 1.00 0.05 0.04

38.14 0.26 21.07 22.65 1.27 3.96 12.73 0.10 0.04 0.03

49.15 (I.39 4.48 13.03 0.31 9.29 21.79 0.93 0.00 0.02

38.27 0.14 20.58 24.07 1.59 3.49 12.29 0.00 0.00 0.02

49.52 0.38 4.65 11.16 0.22 10.61 22.13 0.75 0.00 0.01

38.29 0.15 20.93 23.37 1.27 4.05 12.31 0.00 0.00 0.02

49.92 0.54 5.24 11.06 0.25 10.15 21.64 0.92 0.00 0.00

39.24 0.32 20.91 23.21 1.22 4.51 11.58 0.00 0.00 0.00

Total

98.35

100.25

99.39

I(X).45

99.43

100.39

99.72

100.99

TiO~ AleO~ FeO MnO MgO CaO N a~O K,O -

Cation formula units (based on 6 oxygens for both minerals) Si A1 Ti Fe Mn Mg Ca Na K

1.878 0.222 0.013 0.329 0.006 0.601 0.906 0.075 0.002

1.489 0.970 0.008 0.740 0.042 0.230 0.533 0.008 0.002

1.889 0.203 0.011 0.419 0.010 0.532 0.897 0.069 0.(100

1.498 (I.956 0.008 0.788 0.053 0.204 0.516 0.000 O.(XX)

1.885 0.209 0.011 0.355 0.(X)7 0.602 0.902 0.055 0.000

1.495 0.963 0.004 0.763 (}.042 0.236 0.515 0.000 0.0(X)

1.888 0.234 0.015 0.350 0.008 0.572 0.877 0.067 0.000

1.514 0.951 0.009 0.749 0.040 0.259 0.479 0.000 0.000

Geothermometry

KI)

(XFe/XMg)gt

5.87

4.92

5.48

4.72

785 776

849 804

815 797

885 813

(XFe/XMg)cpx Ellis & Green (1979) Saxena ( 1979),-

* Clinopyroxene inclusion in garnet porphyroblast, others are separate but adjacent pyroxene and garnet grains. Numbers in brackets refer to number of analyses averaged. + Temperatures in ~ and calculated at P=5 kb. Error +50~

Table 4.4. Deformational features and temperature, / pressure conditions of D i and D, stage metamorphic rocks.

Main rock type

Deformation stage

Foliation

Schistosity

Metamorphic facies T & P conditions Upper amphibolite - lower granulite transition T = 800+50~ P = <5 kb

Amphibolites, tonalitic segregations,

DI

F~ isoclines (rarely seen)

$1 axial planar schistosity (only preserved in F, fold hinges)

calc-silicates quartzites

D,

F: isoclines

$2 axial planar schistosity (dominant fabric)

M~

static greenschist facies retrogression

Ophiolite Detachment, Emplacement and Subsequent Deformation

I47

(1974) (Fig. 4.3). However, as the amphibolites seem to have formed beneath a 15-20 km thick ophiolite slab (see Section 4.3.3.3), minimum pressures of formation of about 5 kb seem likely. 4.2.1.2

Amphibolite structuresand metamorphicfabrics

The types of foliation and schistosity seen in the amphibolites and therefore related to the detachment processes are shown in Table 4.4 together with the main rock types present and the calculated metamorphic T and P conditions. The amphibolites show a dominant $2 foliation which overprints an earlier S1 fabric only occasionally preserved in the cores of F2 isoclines (Plate 4.4). The foliations (Fig. 4.5) are usually parallel or sub-parallel to any lithological banding which in turn is parallel to the thrust contact and the foliation at the base of the overlying peridotites (Searle 1980). The schistosity and fold axis plots for four areas, presented in Figure 4.5, show that the metamorphic structures produced during detachment processes vary in orientation. Accepting that these structures were produced by the underthrusting of a slab of oceanic lithosphere, the plots record the relative movement, with respect to this slab, of the portion so detached. It should be noted that although in all cases, the detachment direction has a westerly component, it varies from 287 ~ in the north (Dibba zone) to 223 ~ in the southernmost outcrop measured (Haybi-Hawasina), some 350 km to the south. Indeed, the movement azimuth swings from WNW, through W, to SW going south along'the metamorphic outcrop. Even allowing for the rotation of the various ophiolite blocks on emplacement (see Section 4.3.3), these directions vary significantly and suggest that detachment may have been a polyphase process involving movement in different directions, along different planes, at different times.

Plate 4.4. Banded amphibolites showing dominant $2 schistocity and F~ isoclines. Wadi Hawasina.

DIBBA ZONE 9 Lineations on S2 foliation planes

B

D

c

oo

SUMEINI [] Mean schistosity pole Sumeini South ~ with great 9 Mean schistosity pole Sumeini North f circles

Fig. 4.5.

WADI HAM with [] Mean schistosity pole Wadi Ham East t great 9 Mean schistosity pole Wadi Ham West ~ circles

Equal area stereographic plots of amphibolite facies $2 schistosity and F 2 fold axes from four selected areas.

HAYBI-HAWASINA ~'~

\

Trend of major fold axes

Direction of displacement 9 Pole to S2 schistosity o F2 fold axes B C D modified after Searle (1980) unpublished data of D. A. Rothery A and S. J. Lippard.

I48

Chapter 4

4.2.2 Peridotite banded unit

Another apparent manifestation of the detachment process was the production of a c. 150 m thick, intensely mylonitized zone at the base of the Mantle Sequence. In this zone the high temperature (c. 1200~ low stress constructive margin deformation (Section 3.2.4) has been obliterated by a later, lower temperature (c. 700~ large strain-high stress (1-2 kb deviatoric) deformation. This deformation produced closely spaced mylonite zones within the harzburgite-dunite-lherzolite host that are parallel to the plane of the underlying Semail thrust which separates the zone from the subjacent amphibolite facies metamorphic rocks (Section 4.2.1). Searle (1980), using the term "banded ultramafic unit" identified that the mylonite zones in the unit are totally serpentinized, only a few centimetres across and separated from each other by less serpentinized, less deformed zones of similar width in which a porphyroclastic texture is dominant (Plate 4.5). The banding within the unit has, in places, suffered recumbent and/or small-scale chevron folding. The more intense serpentinization of this zone, compared to the rest of the Mantle Sequence, gives the unit a darker appearance that makes it readily identifiable in the field, on aerial photographs and on satellite imagery (Rothery 1982). The unit consists of the same rock types in much the same abundance as in the rest of Mantle Sequence (Section 3.2). Dunite masses are markedly flattened and elongated parallel to the banding and this is probably due to extreme deformation of the type described by Bartholomew (1983) and illustrated in Figure 3.9. The downward increase in lherzolitic character of the Mantle Sequence (Section 3.2.1.1) coincides with the banded zone, and the significance of this is discussed in Section 3.9. In the mylonite zone the harzburgite consists of minute (2050 ~tm) recrystallized grains of olivine and orthopyroxene enclosing rare elongated olivine and orthopyroxene porphyroclasts, the latter being strongly elongated. Throughout, the dominant mineral is olivine with orthopyroxene forming a few percent and chrome spinel usually less than 1%. The downward increase in lherzolite character, indicated by whole rock, orthopyroxene and spinel geochemistry has not been identified petrologically. Searle (1980) and Searle & Malpas (1980)

maintain that this mylonitized (Banded) basal part of the Mantle Sequence is at most 150 m thick. In contrast, Boudier & Coleman (1981, p. 2584) maintain that "the strain related to this deformation is particularly high in the lower 500 m of the cross-section and is concentrated in mylonite bands; as much as 2 km higher in the cross-section, the imprint of this high stress deformation is restricted to close spaced microstructures in the olivine". Although unable to confirm Boudier & Coleman's findings in this respect, we do agree that petrofabric analysis indicates that the higher rocks in the zone were displaced westwards with respect to those below. Because the structures and textures in the amphibolite to granulite facies rocks of the Metamorphic Sheet and in the Banded Unit were produced at similar temperatures and pressures, their collective evidence can be used in deducing detachment processes. Regarding temperatures, the amphibolites were formed at temperatures of 820 ~ 50~ (Table 4.3) and 750~ according to Ghent & Stout (1981). The temperature of deformation in the Banded Unit is variously described as 950~ and c. 1000~ by Boudier & Coleman (1981) and Boudier et al. (1983) respectively. As there are considerable inherent errors in the various methods of identifying palaeotemperatures, a temperature regime of 750~176 seems reasonable. Ignoring the contribution of serpentinization, the possible sources of this thermal energy are twofold: (i) that remaining in the uppermost, recently generated oceanic lithosphere that was, at most, 10 Ma old and (ii), that produced by friction between the two slabs. Opinions vary widely as to the proportion of heat derived from each of these sources and these, in turn, depend on varying thermodynamic constants used. Several authorities however (e.g. Malpas 1979; Ghent & Stout 1981) conclude that frictional heat was unlikely to exceed c. 200~ and it is tentatively concluded here that, within the temperature regime of 750-1000~ the lower figure is more likely and that some 200~ of this was frictional, the remainder being provided by residual heat in the overlying hot ophiolite. At the end of Section 4.2.1 it was noted that the structure within the amphibolites of the Metamorphic Sheet gave displacement directions of between 287 ~ and 223 ~. This southward swing of displacement direction is continued south-

Plate 4.5. Banded Unit. Mylonitized and serpentinized bands follow the tectonite fabric. There are also cross-veins marked by trains of magnetite grains. Field of view - 3 ram.

Ophiolite Detachment, Emplacement and Subsequent Deformation wards along the ophiolite outcrop into the Rustaq* Block where Boudier et al. (1983) identify north to south (180 ~ movement in the banded peridotite. In contrast, the movement as deduced from the banded peridotite in the Tayin Block is E-W, indicating, as do other criteria, that there is a major structural discontinuity along the line of the Semail Gap (Fig. 4.1). Although there can be little doubt concerning the direction of movements deduced from structural studies, (e.g. Searle 1980; Searle & Malpas 1980; Boudier & Coleman 1981; Boudier et al. 1983), there is no evidence that these were synchronous. If one coherent slab was being thrust over another, it is difficult to see how such a wide diversity of movement was produced especially when the structural trend of the detached slab, as indicated by the strike of the sheeted dykes is more or less constant. The 90 ~ change between Rustaq and Tayin can be accounted for by a major structure along the Semail Gap, it is the progressive change from 287 ~ to 180~ going south along the ophiolite outcrop of the northern mountains that is more difficult to explain. The only suggestion we can offer is that between the coherent top of the ophiolite and the more intensely deformed base of the Mantle Sequence there is a thrust plane(s) along which rotational movement of the lower part could take place.

4 . 3 E m p l a c e m e n t a n d related processes At 90 +_3 Ma the detached slab of oceanic lithosphere, with a thin amphibolite sheet now welded to its base, started to move. Palinspastic reconstructions (Graham 1980) indicate that it will have travelled a minimum of 300 km with respect to the underlying formations before being emplaced on the north-eastern Arabian continental margin some 10 Ma later. The constraints for this relative movement fall into three categories: (i) the metamorphic rocks from which an age of emplacement can be calculated (ii) the igneous rocks whose geochemistry identifies that underlying continental crust is present and (iii) the complex structural deformation which affected other rocks of the Tethyan ocean floor as well as the ophiolite. Included in this last section will be the processes that produced the melanges associated with the Semail Nappe. These data will be presented in this order as Sections 4.3.1, 4.3.2, 4.3.3, where the metamorphic, igneous and structural processes respectively form the main themes. There are however many complications to this oversimplified division for, not only are the three processes complex in themselves, but they overlap in time and space and further complicate each other.

4.3.1 Metamorphism We are here primarily concerned with (i) the epidote amphibolite-greenschist facies rocks that occur in the lower part of the sub-ophiolite Metamorphic Sheet. However, the metamorphic history of the area is complicated by (ii) the imprint of a highpressure blueschist facies metamorphism that is itself overprinted by a greenschist event; both occurred during the emplacement process. Here we describe and discuss the two metamorphic events separately although some of the pro-

*We have retained Open University block terminology rather than use the new names introduced by the French workers. We deplore the completely unnecessary complications introduced by our French colleagues who seem to change place and formation names wherever they work.

I49

cesses; e.g. those of retrogressive greenschist facies metamorphism, occur in both.

4.3.1.1 Sub.ophiolite epidote amphibolites and greenschist facies

schists There are various lines of evidence which indicate that the metamorphism that produced the epidote amphibolite-greenschist facies rocks of the metamorphic sole took place well after that which produced the amphibolites. The most obvious, although possibly the least trustworthy evidence is that the radiometric dating of the amphibolites gives an age of 90 + 3 Ma whilst the greenschists fall in the age range of 85-70 Ma. The clearest geological evidence is the rare occurrence of amphibolite blocks, with the main amphibolite F 2 foliation, enclosed in greenschists. The deformation of the blocks clearly predates that of the greenschist host. The epidote amphibolite and greenschist facies rocks form the lower part of the metamorphic sheet and comprise a varied group of metasedimentary and metavolcanic schists showing polyphase folding and deformation. Their maximum structural thickness is about 300 m but they are everywhere cut by numerous small thrusts and unit boundaries are commonly thrusts. There is usually a sharp tectonic and metamorphic break between the greenschists and the underlying rocks, commonly one of the Haybi Complex units (e.g. Haybi volcanics, sedimentary olistostrome melange) or Hawasina sediments. The most important lithologies in this unit are garnetbiotite-muscovite schists, quartzites, marbles and phyllites, developed from sedimentary and volcanic protoliths (Plate 4.3). Typical mineral assemblages in the metabasic rocks are: (i) blue-green actinolitic hornblende + epidote + plagioclase (An3~-lo) (epidote-amphibolite facies) and (ii) crossite-actinolite + epidote + albite (upper greenschist facies). Banded iron and manganese-rich quartzites contain small pink piemontites, Mn-rich garnets and rare stilpnomelane and are probably metamorphosed pelagic cherts. Marble bands, commonly interbedded with the quartzites, contain calcite, epidote, zoisite and grossularite garnet. Metapelites are graphitic schists and phyllites composed of fine-grained biotite and sericite, which, with increasing quartz content, grade into quartz-rich semi-pelitic schists that occasionally contain small colourless to pink garnet porphyroblasts. The metabasaltic rocks, generally subordinate to the metasediments, sometimes preserve relict pillow or pyroclastic textures but some are coarser grained and probably represent intrusives. In summary, the greenschists were formed by the metamorphism of a varied series of sandstones, shales, limestones and cherts with interlayered basic volcanic horizons or sills suggesting derivation from the Hawasina series or possibly locally the Haybi volcanics. The greenschists are macroscopically folded in most outcrops and show up to five phases of deformation. The earliest and usually dominant ($1) foliation is axial planar to early isoclines and generally parallel to lithological boundaries. Some of the greenschists show well-developed mylonite fabrics, probably related to thrust movements, which have strong quartz rodding formed during D2 folding on more or less horizontal axes. D3 is represented by a static phase of mineral growth on $2 foliation surfaces. D4 and D5 are late minor fold phases with steeply dipping axial planes and not associated with any new growth of metamorphic minerals. D 4 folds are open to moderately tight structures with non-penetrative crenulation cleavage, D5 folds are open box-folds or kink bands. All the fold structures are truncated by thrust planes

Chapter 4

I50

Table 4.5. Types of foliation and schistocity developed in epidote amphibolite - greenschist facies metamorphic rocks. Main rock type

Deformation stage

Foliation

Schistosity

Greenschists

D~

FI isoclines

$1 axial planar schistosity/ cleavage, mainly parallel to bedding and dominant fabric in the rocks

D~

F~ isoclinal to semi-recumbent folds with mainly upright axes

$2 cleavage poorly developed

M3

Static growth of greenschist facies minerals

D4

F3 crenulations

D5

F4

kink-bands

formed during the late stages of nappe emplacement. The D 4 and D5 folds are similar to the common fold styles in the underlying Hawasina sediments and may be the same age. The types of foliation and schistosity displayed by these rocks are tabulated in Table 4.5. The greenschist metamorphic and deformation events are later than those in the amphibolites which resisted greenschist facies deformation because they were effectively welded to the base of the Semail Nappe at the time. However, the agents of metamorphism for the greenschists were the same as those for the amphibolites - namely, heat from an overlying hot ophiolite slab and frictional heat and deformation produced by the movement along the underlying thrust plane. During the emplacement stage the residual heat from the overlying ophiolite would be much reduced. Fifteen to twenty million years after its generation, the temperature of a 10-20 km slice of upper oceanic lithosphere would have reduced to c. 200-400~ (Sclater & Franchetau 1970), although the thermodynamic (shear heat) input due to the movement between the two rock masses would be much the same - about 200~ 4.3.1.2.

High-pressure metamorphism - the Saih Hatat blueschists

Schistose rocks containing abundant blue amphiboles occur in the northern part of the Saih Hatat tectonic window south of Muscat (Fig. 4.1). The high pressure metamorphism, believed to have occurred at up to 8 kb and 350-400~ (Lippard 1983) is overprinted by a greenschist facies event, particularly seen in metabasites where early-formed blue amphiboles are rimmed by actinolite and included in late- to post-kinematic porphyroblasts of epidote and albite. Most of the rocks affected by the high pressure metamorphism are metasediments; pelites and psammites, that may be metamorphosed equivalents of the Hawasina series (Boudier & Michard 1981) or part of the pre-Permian basement (Glenhie et al., 1974). Quartz-mica schists are the most common rock type and are composed of highly-deformed, elongate

Metamorphic facies

Epidote amphibolite to upper greenschist facies metamorphism

$3 non-penetrative axial planar cleavage No fabric developed

quartz and phengitic white mica with occasional glaucophane and albite. These semi-pelites are interbedded with bands of AI-Fe-rich pelite containing up to 30% red almandine garnets and 20-30% dark-blue glaucophane - crossites in a matrix of fine grained quartz, phengite and late chlorite. The garnets (unzoned almandines AlmyyPyllGrllSpl) are poikiloblastic with inclusions of dark green chloritoid, magnetite and quartz. They are partly altered to chlorite around the margins. The blue amphiboles plot on the boundary between glaucophane and crossite (Fig. 4.3) and show some zoning with Fe and Ca increasing and Al decreasing from core to margin. Layers of mafic schist, representing volcanics or basic sills, consist of epidote and unzoned, untwinned albite (Ano.5_~) porphyroblasts that contain numerous blue-green amphiboles. These are zoned glaucophane-crossites which may be rimmed by a narrow zone of actinolite. Up to 30% of the matrix consists of late interstitial chlorite containing minute inclusions of sphene. The early-formed blue-green amphiboles form inclusion trails that define an F~ fabric predating the crystallization of albite and epidote porphyroblasts that are weakly aligned in the F2 foliation. All the rocks are affected by the syn-metamorphic S~ foliation which is axial planar to minor FI folds and is generally parallel to bedding in the metasediments. In incompetent lithologies (phyllites, mica schists) this is transposed by a penetrative $2 cleavage/schistosity which is axial-planar to most of the major flat-lying recumbent folds. In addition some phyllites show an $3 crenulation cleavage. The dominant S~_2 or $2 foliation bears a well-defined stretching lineation developed during F2. This lineation has a constant orientation over most of the Saih Hatat area and trends 010 ~ which is interpreted by Michard et al. (1984) as the emplacement direction of the nappes in this area. K-Ar ages of 100+4 and 80+4 Ma (Table 4.1) of phengites from the Oman blueschists near As Sifar suggest that the high pressure metamorphism occurred both during and before emplacement of the Semail Nappe. In the northern Saih Hatat area the schists form part of a complexely folded, imbri-

Ophiolite Detachment, Emplacement and Subsequent Deformation cate nappe sequence and are interleaved with apparently unmetamorphosed Permo-Triassic carbonates and pre-Permian basement rocks (Michard et al. 1984). The high pressure metamorphism thus occurred outside the area and most likely took place in that part of the continental margin that was depressed to depths of c. 20 km beneath the nappe pile. Later (Section 4.3.3), it will be argued thdt this was due to underthrusting beneath a northward-dipping subduction zone. The greenschist facies overprint on the blueschists could be the result of rapid isostatic uplift, following the cessation of subduction, during which temperatures of 300-400~ were maintained.

4.3.2 Igneous activity associated with emplacement (biotite granites) Igneous rocks that fall into this catagory have to be clearly intrusive into the ophiolite and also into the subjacent amphibolite sheet. The only rocks that meet these criteria are small intrusive bodies of white-weathering, fine-grained, foliated, biotite granite and aplite that occur in the lower part of the Semail Nappe (Browning & Sinewing 1981; Browning 1982), mainly in the Khor Fakkan ( U A E ) Block and in the SE part of the Haylayn Block near Rustaq. In both areas they have been emplaced into the upper part of the Mantle Sequence and near the base of the Layered Series. Searle (1980) records the presence of a biotite aplite (sp OM 118) cutting the amphibolite sheet and the overlying harzburgites at Sharm in the UAE. In the Haylayn area the granites form elongate bodies with m a x i m u m dimensions of c. 500 m x 100 m, that generally have steep contacts with the host peridotites and gabbros and are elongated W N W - E S E along high-angle faults (Browning 1982). They are mostly compact and massive rocks but near the contact with the country rocks show minor brecciation with quartz and prehnite veining. The granites are fine to medium grained ( < 2 ram) leucocratic rocks with granoblastic textures. Zoned plagioclase (An4{~25) and orthoclase porphyroblasts, between 1-2 mm across, occur in a fine grained (0.05-1 ram) mortar textured, quartzo-feldspathic matrix. The mode is c. 40-50% quartz, 45-55% feldspars and < 5 % mafic minerals. Small biotite flakes are sometimes included in the K-feldspars but they occur mainly as interstitial flakes and shreds forming stringers parallel to the foliation. The rocks show evidence of sub-solidus, post-deformation alteration, under greenschist facies conditions. The calcic cores of the plagioclase are saussuritized and the K-feldspar is partly replaced by myrmekitic albite-quartz intergrowths. The biotites are part to completely chloritized and there are patches of secondary fine-grained green amphibole (tremolitic hornblendes Mg' 73, A1203 3.8-6.1%, M n O 1.2-1.9%), epidote, prehnite and magnetite in the groundmass. The freshest sample, an aplite from the Sharm area of the U A E (OM118), has the highest K, Rb and Ba contents. These elements have been lost from the more severely altered samples where the K-feldspars and biotites are replaced by secondary minerals and where chloritization of the biotite has increased the Mg and Fe contents at the expense of alkalies. Biotite from O M l 1 8 has yielded a K-Ar age of 85 +_3 Ma (K20 5.44%, vol4~ 1.8301, 4~ 78%, decay constants as on Table 4.1 ; Analyst D. C. Rex, Leeds University). This age overlaps with the ages of the Metamorphic Sheet (Section 4.2.1, Table 4.1), being slightly younger than the mean age (90 + 3) of the amphibolites and somewhat older than that of the greenschists (80-70 Ma); this accords with the

I5I

Table 4.6. Analyses of biotite aplite (Searle 1980) and biotite granite (Browning 1982). OMll8

0M2697

SiO, TiO, AL203 Fe:O3 MnO MgO CaO Na,O K~O P20 5 LOI

73.27 0.26 13.69 2.62 {).09 1.17 2.85 2.52 3.25 0.08 {}.90

72.93 0.34 12.58 3.22 0.17 3.17 3.19 2.27 1.02 0.21 1.89

Tot al

100.60

100.97

38.23 1.O4 19.27 21.39 13.76 3.92 1.80 0.50 0.19

44.76 2.48 6.10 19.44 14.63 9.42 2.01 0.65 0.50

-

Q C or ab an hy mt il ap Rb Ba Sr Zr Y Nb Cr Ni ka Ce Nd Sm Eu Yb Tm Yb Lu Th Ta Hf U Sc Cs

76 2{11 117 108 20 8 19 11 35.3 (1{17.6) 73.2 (84.6) 28.1 (44.6) 5.2 (25.6) 1.17 (15.2) 0.62 (11.9) 0.31 (9.121 1.64 (7.45) 0.29 (8.55) 11.05 0.92 2.65 3.6 13.6 2.6

24 117 124 111 28 14 44 89 24.5 58.1 24.5 4.8 1.10 0.77 0.40 2.42 0.40 9.11 1.13 3.16 3.{1 18.6 1.{}

(74.7) (67.2) (38.9) (23.6) (14.3) (14.8) (11.8) (11.0) (11.8)

field evidence in that the granites cut the amphibolites but not the underlying greenschists. Geochemically the biotite granites are peraluminous calcalkaline types that are enriched in trace elements, particularly the LIL elements (K, Ba, Rb, Cs, Th, U), compared to the plagiogranites from the ophiolite (Table 4.6, Fig. 4.6). They are markedly light REE-enriched with steep slopes between La and Eu and flatter trends from Eu to Lu on a chondritenormalized plot (Fig. 4.7). These R E E patterns are typical for granitic rocks that have been produced by the melting of the upper continental crust (e.g. average greywacke). The fresh specimen (OMl18) plots close to the minimum melt composition on the normative Q-Plag ( A n + A b ) - O r plot at p H 2 0 = 4 kb (Fig. 4.8). An origin by partial melting of continentally derived material is likewise supported by the strontium and

152

Chapter 4 Q

10

I

I

K2O Rb

I

I

Ba

Th

I

I

Ta Nb

I

I

I

I

I

9

Ce

Hf

Zr

Sm

Y

Yb

PLAG (Ab + An)

Fig. 4.6. Rock/ORG (average ocean ridge granite of Pearce et al. 1984) for biotite granites (open circle and circle with dots) compared to typical axis plagiogranite (solid circles).

100 -

JOM 118 OM

~

5

.

:i:i:i:i:i:::i:!ii:i:i:i:!:!:i:i:i:i!:i::ii i!:!:i:i:i:i:i:i:i:i:!!i:i:?!:i:i:i

10

I

I

ka Ce

I NO

I

l

I

I

Sm Eu Gd Tb

I

I

I

Tm Yb Lu

Fig. 4.7. Chondrite-normalized plot of rare earth elements for two biotite granites compared to plagiogranites from the Semail ophiolite (shaded region). For analyses see Table 4.7.

neodymium isotopic compositions (eNdS5 ~ t -- --4.1, es~85 = +66.9) which suggest derivation from a Lower Palaeozoic, Na isotopically-enriched continental source (TcnuR = 480 Ma; Dunlop & Fouillac 1984). Pearce et al. (1984), discussing the

Or

Fig. 4.8. Biotitegranites plotted on a normative Ab + An - Or - Q plot showing phase boundaries and minimum melt compositions at 4kb.

origin of these granites in terms of trace element discrimination diagrams for determining the tectonic settings of granitic rocks, assigned them as a "post-collisional" type despite the fact that the biotite granites fall into the "volcanic arc granite" fields on most of the plots (e.g. Yb vs Ta, (Y+Nb) vs Rb and (Yb+Ta) vs Rb). Geological and geochronological evidence indicates that they are early "syn-collisional" in relation to the obduction of the Semail Nappe onto the Arabian continental margin and should therefore be termed "syn-obduction" These granitic rocks differ markedly from the tonalitic partial melts found in the amphibolites of the Metamorphic Sheet which were produced earlier by melting of ocean-floor basic rocks (Searle & Malpas 1980, 1982). The isotopic and chemical compositions suggest that they were produced by the partial melting of continental or continentally derived rocks. This combined with their syn-obduction age, makes it most likely that the melts that formed these rocks were produced by the partial melting of the continental lithosphere due to heat emanating from the still-hot ophiolite during its emplacement. There is little doubt that sufficient heat can be provided as metamorphic mineral pair geothermometry (summarized by Spray 1984, p. 260) indicates that temperatures up to 875~ have been produced at the base of ophiolites. How was the necessary heat budget obtained? Jaegar (1961) calculated that the maximum contact temperature between a hot upper slab and a lower cooler one in a static situation would be about 0.5T (where T is the temperature difference between the two). A c. 10 Ma old oceanic lithosphere at a depth of 15-20 km and 300 km from the ridge axis (3 cm/yr half spreading rate) would have a residual temperature of 800-850~ (Sclater & Francheteau 1970). The contact temperature on the underlying rocks produced by such a static, hot ophiolite slab would be 400450~ clearly insufficient for the production of a granitic melt. But if the heat source is continuously moving, the temperatures at the contact rise and become closer to that of the overlying hot slab (Jaegar 1942). So with this additional thermal input and another (<200~ provided by shear heating along the thrust plane (Graham & England 1976) partial melt temperatures could be achieved.

Ophiolite Detachment, Emplacement and Subsequent Deformation SW

NE

(a) 110 Ma AFROARABIA

initiation of s u b d u c t i o n

A

EURASIA

(c) c.95 Ma <

.--

300-750km

---

.

a : n e w l y c r e a t e d s u p r a s u b d u c t i o n zone o c e a n i c lithosphere

>a B

%

~'~a = 105-175 km p r o d u c e d b e t w e e n 9 7 . 9 - 9 3 5 Ma

and then detail evidence which supports them. This approach is primarily to present an understandable process from a superabundance of structural detail. Structurally, the Oman Mountains are an Alpine forelandtype fold and thrust belt in which nappes of Mesozoic Tethyan rocks were thrust onto the continental rocks of the northeastern Arabian plate margin. The Tethyan rocks so emplaced include (i) Triassic to Cretaceous continental slope, rise and abyssal plain sediments (the Sumeini Group and the Hawasina Assemblage), (ii) Triassic alkali to transitional basic volcanics (the Haybi Volcanics) that are associated with Triassic to

Fig. 4.9. Suprasubduction zone setting of the Oman ophiolite. Initiation of subduction at (A) is calculated to have occurred at c. 110 Ma. Subsequent seafloor spreading related to this subduction started at (B) at 98 Ma. The minimum length of subducted lithosphere (c) is 85 km on magmagenetic grounds and 140 km based on present-day trencharc separations. This implies that subduction was initiated a minimum of 5 Ma prior to new oceanic crust generation i.e. - 1(13 Ma. Fig. 4.12 shows our consensus model in detail.

c

S

.:~T~.k ~ ~- :C7~t~/ :,~\-:~

I

A

I

INEIIranl

4.3.3 Structural evolution .....

It has already been stated in the introduction to this Chapter (Section 4.1) that the Semail ophiolite represents oceanic lithosphere, produced between 97.9-93.5 Ma ago on the western flank of a N-S constructive margin above a NNE inclined subduction zone that lay some 300-700 km NNE of the Arabian (Oman) continental margin; detachment of this slab that was to become the O m a n ophiolite from its in situ position as oceanic lithosphere took place 90 + 3 Ma ago. Palaeomagnetic evidence (e.g. McElhinny et al. 1968, Briden 1967) suggests that some 110 Ma ago, the AfroArabian plate, which had been stationary in the South Polar area for the previous 100 Ma, started to move northeastward. Plate tectonic constraints require that this movement be accommodated by subduction of the plate's leading edge. This took place in the oceanic realm of Tethys, the Neotethys part of which was created by the divergent movement of Eurasia and Afro-Arabia that started c. 200 Ma ago (Section 2.2). On the evidence here presented at least some of this subduction was to the south and west of the oceanic lithosphere that was to become the Semail ophiolite. Briden & Gass (1974, Figs. 3 & 4) indicate that the movement of Afro-Arabia took place between 110 and 45 Ma ago and on reconstructing their polar wander curve it appears that the plate travelled 1750 km at an average rate of 2.7 cm yr -1 during this time. This, in turn, implies that just over 400 km of oceanic plate had been subducted before Semail ophiolite magmatism commenced at c. 95 Ma. Magmagenetic considerations indicate that the initial mantle melting that produced basic magma takes place at depths of 50-70 km and this, assuming a 45 ~ subducting angle, requires at least 85 km of oceanic lithosphere be subducted before magmatism is initiated. At a rate of 2.7 cm yr- ~it would take 3.1 Ma to subduct 85 km of oceanic plate, this is the minimum period required, more regional considerations suggest that 400 km were subducted during a 15 Ma period. Here we attempt to establish the order of deformation events and systematically restore them to produce a palinspastic reconstruction of the original palaeotectonic situation. As the main objective of this work is to describe and discuss the Semail N a p p e , we present our structural conclusions first

t53

::::i i

.

I

ii i i

Uphioiltes '\/ ~/,;/,r,,N~ I fin S Semail \ ~ I/~.'N ~ure T Troodos \ ( / / / H Halay \. // ~ H::r~z "\\, . / z ~ / . . - ~ . Present day f~ Spreading \/...t-" outline of Arabia ~; axis

84 A

:i'iii:!i:):;ii:i:~/'~\ ~ ~L- i l!

//t

,

./'J//W~Q.eental ~ /

:'-!i::.:.:!::,::?;:~~."?:~?.: co

/ " 4/2- ,. /<,.

/

2\ /0/O' %

0

I

100km

l

/

I~]

Mid-Cretaceous oceanic lithosphere

~

Older Tethyan oceanic lithosphere

Fig. 4.10. Plate tectonic setting envisaged for NE Arabia in the mid Cretaceous. Lower figure, enlarged from box in upper figure, shows new oceanic crust (destined to become the Semail Nappe) forming NE of the marginal subduction zone.

I54

Chapter 4

L o w e r C r e t a c e o u s pelagic sediments and large exotic blocks of P e r m i a n and U p p e r Triassic shallow w a t e r c a r b o n a t e s t o g e t h e r with a s e d i m e n t a r y olistostromal m e l a n g e (the H a y b i Complex) and (iii) the basic and ultrabasic oceanic b a s e m e n t rocks of the Middle C r e t a c e o u s Semail ophiolite (the Semail N a p p e ) . T h e main d e f o r m a t i o n s t o o k place during the T u r o n i a n to C a m p a n i a n (91-73 Ma); the nappe structures were then u n c o n f o r m a b l y covered by Maastrichtian marine sediments. F r o m the late P r e c a m b r i a n to the C r e t a c e o u s the area was part of a stable c o n t i n e n t a l region r e p e a t e d l y subjected to m a r i n e incursions during which shallow water c a r b o n a t e s were deposited. Permo-Triassic rifting caused continental b r e a k - u p so that the area to the n o r t h e a s t of the present m o u n t a i n s b e c a m e the s o u t h e r n passive margin of Tethys. S o m e t i m e b e t w e e n 110-100 M a ago, a N to N E - d i p p i n g subduction zone d e v e l o p e d . It is p r o b a b l e that this zone was initiated some 800 km north of the continental margin. Figure 4.10 shows the envisaged situation. W e here present a sequence of figures (Figs. 4.11A-M) which we believe represent critical stages in the ophiolite n a p p e e m p l a c e m e n t . T h e y incorporate all the k n o w n radiometric age data and biostratigraphic constraints, the latter r e t i m e d using the recently published timescale of H a r l a n d et al. (1982). In cases where several age dates have b e e n m e a s u r e d for the same p h e n o m e n o n (e.g. the

a m p h i b o l i t e facies m e t a m o r p h i s m ) , we infer that the process c o n t i n u e d over a time span incorporating the m e a s u r e d ages. C a r e has b e e n t a k e n to scale the c o m p o n e n t s of the model, the c o n t i n e n t a l margin configuration is t a k e n from that determ i n e d for the US East Coast by G r o w et al. (1979). W e believe that this s e q u e n c e of cross sections represents a realistic reconstruction as it incorporates the postulate that over the entire area (initial width 1100 km), the rate of cons u m p t i o n of crust will be a s m o o t h function. Previous interpretations often d e m a n d e d unrealistically high displacement rates and rapid changes b e t w e e n overall c o n s u m p t i o n and spreading. H e r e the figures are based on the proposition that the subduction rate will increase due to the increasing length of cold d e s c e n d i n g slab ( M c K e n z i e 1969). This increase in rate will then diminish and level off as the d e s c e n d i n g slab reaches regions of higher t e m p e r a t u r e and/or s e g m e n t s are shed. It is here c o n s i d e r e d that s p r e a d i n g is most likely to be initiated at the time of the steepest gradient in the subduction rate, w h e n it is most difficult for the crustal system to a c c o m m o d a t e this change. T h e effect of spreading, itself a s m o o t h function, is to reduce the a m p l i t u d e of the crustal c o n s u m p t i o n rate curve (Fig. 4.12). The d e s c e n d i n g side of this curve is constrained in time by the f o r m a t i o n of the a m p h i b o l i t e s but there is little to constrain the a m p l i t u d e on geologic evidence p r e s e n t e d here. This s c h e m e b e c o m e s invalid b e t w e e n - 8 8 to 78 M a by which

Fig. 4.11. (a) 105 Ma Subduction commences approximately 800 km from the continental margin. Initial subduction rate = 3 cm/a. (b) 100 Ma The subduction rate has increased due to the increasing length of cold descending oceanic lithosphere (Fig. 4.11). By 100 Ma approximately 180 km of lithosphere has been subducted and as the rate increases, spreading is about to start 100 km oceanwards of the SZ. Overall closure rate across the area now - 5 cm/a. (c) 96.5 Ma At the steepest gradient in closure rate, spreading is initiated (98 Ma) and rapidly increases to an estimated 3 cm/a/flank. By 96.5 Ma, c. 64 km of new oceanic crust has been created above subducted older oceanic crust. Assuming a 45 ~ subduction angle, this extends a distance of 300 km from the SZ i.e. more than sufficient to give the required geochemistry. Over 400 km of the original Tethyan oceanic crust has by now been subducted and the outer limits of the Hawasina marginal basin sediments are approaching the subduction zone. The resulting compressional forces initiate thin-skinned thrusting in the Halfa and Haliw sediments. Overall closure rate for the area is c. 8.0 cm/a. (d) 93 Ma Subduction has continued to the extent that continental crust is now involved in the subduction zone. This effectively chokes the subduction process (McKenzie 1969) and halts spreading. The resulting stresses cause the oceanic lithosphere to break at the western margin between old and new oceanic crust. Amphibolite facies metamorphism is produced along the dislocation zone in its initial movement (Section 4.2). The Hawasina sediments originally extending over a distance of 340 km have been imbricated and as a result thickened by c. 3.4:1 (Graham 1980b). The thrusting has yet to involve the shelf and slope sediments. Closure rate is now c. 6.5 cm/a. (e) 91 Ma The marginal continental crust is still subsiding but much more slowly. The closure is principally taken up by the western flank of the new oceanic lithosphere overriding its older, denser counterpart. Amphibolite facies conditions continue in the later parts of the dislocation shear. Closure rate is now = c. 3 cm/a. (f) 88 Ma The continental margin is further depressed by the combined effects of the slowing subduction and overriding of the new oceanic crust. The shelf-edge sediments are incorporated in the imbricate stack of margin sediments. Blueschist metamorphism is being impressed on parts of the subducted continental margin. (g) 85 Ma The older oceanic segment is now completely overriden and the hot slab of newly created oceanic lithosphere lies directly over continental crust at depths of 15-25 km (Lippart 1983). In these conditions the temperature and pressure are locally sufficiently high to partially melt the continental crust and provide the biotite granites

magmas (Section 4.3.2). Away from these localized melt zones blueschist conditions exist in the continental margin crust (Section 4.3.1.2). Closure over the area has been slowed and as a result the system is in overall compression. (h) 80 Ma Greenschist facies metamorphic rocks are incorporated in the metamorphic sole of the new oceanic lithosphere as it begins to override the marginal basin sediment stack (Section 4.3.1.1). The outer parts of the continental margin are re-equilibrating. Isostatic rebound of the outer margin, following cessation of closure across the area, is recorded by the overprinting of a greenschist metamorphism on the blueschist margin rocks (Lippart 1983). This uplift has two principal effects, it (1) raises the new oceanic crust above sea level giving it high potential energy and stress and (2), causes a sympathetic downwards flexure further onto the margin into which sediments (Muti) are rapidly deposited (Burruss et al. 1983). (i) 78 Ma With continuing uplift of the outer margin the new oceanic crust shears and the western flank moves over and deforms the Hawasina sediments. The failure in the new oceanic lithosphere was most likely to take place at the weakest point - - the palaeo-ridge axis. The greenschist facies metamorphism is now fully established in the metamorphic sole and continues in the outer continental margin crust. Muti (Glennie et al. 1974) sedimentation continues. (j) 76 Ma The flexure in the continental margin is shown here at its most extreme. Continued rebound at the outer continental margin causing the western flank of the new oceanic lithosphere to be emplaced as the Semail Nappe and the eastern flank to move eastward off the high. There must, by this time, have been subduction along the north-eastern Tethys margins to accommodate the changes at this southwestern margin. Considerable melange deposition, shown here as volumetrically equivalent to the erosional loss of oceanic crust, accompanies the separation. The Muti sediments are overthrust by the Semail Nappe. (k) 74 Ma The processes described in (j) continue, the re-equilibration of the continental margin tilts the Semail Nappe to the E and NE leading to further erosion, particularly of the leading edge (Juweiza Fm of Giennie et al. 1974). Greenschist overprinting in the outer margin ceases and the eastern flank continues to retreat eastwards. (1) 72 Ma Further tilting, erosion and thinning of the ophiolite occur as the continental margin re-equilibration continues. (m) 70 Ma Further re-equilibration of the continental margin and retreat of the originally overriding eastern flank of the new oceanic crust allow deeper coastal waters and deposition of Maastrichtian sediments on the margin and ophiolite trailing edge.

Ophiolite Detachment, Emplacement and Subsequent Deformation

I55

A lo5 Ma

o

2o/:4-:',':'.'.'.','.'.'.'.',,,'5,'.', 4 :'.'.'.-.'.-.'." -" -1

I

.

.

.

.

.

.

.

.

B loo Ma 0

~ , ' ~ ' . ' , r ,

, ' - r ~ [ T T . . ' . . r q ~ q , ' , ' . ~

.

.

.

.

.

.

.

.

.

.

.

. -

-

%-

.

.

.

.

.

.

.

.

.

.

.

.

.

~

.

.

:"

=.;

. . . . . .

.-.-

,

~

,

~

. . . . .

:_ . . . . .

".:'"

~,

C 96.5 Ma .

.

--

j

~ ' "

.

"

"

9

.

,

-

"::.:.":-:-:

.-

......

:

. . . . . . .

,~

. ",:

-'

-.>,.-!..._--

.~.~::-+..:

:..-....,

::::::::::::::::::::::::::::::::

..

, .......

..

~ / "

~ , ~ ~:

D 93Ma

o ::::::::::::::::::::::::::::::::::::::::

.-...:.

.....

1:,:.:.:.:.:,:.:.:.:.:.:.:.~,1191 Ma

oP-,a

......... ~-

.

,o

G 85Ma

~'~"

H 8OMa 0

.

20

.

~,.,

, . . ,

.

-,,

....

-:

.............

I 7s~

Tethyan oceanic crust with km markers at 105 Ma retained throughout to demonstrate relative movement

J76Ma 0

- -

2O

Upper Cretaceous oceanic crust with half million year time markers

K 74Ma

Continental crust and carbonate cover

~o~::::::::::::::::::::::::::::::::::::::::::: :

~

Amphibolite metamorphic facies rocks welded to Upper Cretaceous oceanic lithosphere during shear

IL 72Ma 0

'

.

Imbricate stack of marginal basin sediments

. . . . . . .

20

B 1 1 I

M 70Ma o

~

~

[

~/.~;i,r,,,)~

Blueschist metamorphism Biotite Granite intrusion Muti sediments Melange

Fig. 4.11. Stages in ophiolite development, detachment, emplacement and subsequent deformation. Descriptive captions for individual cross sections (A-M) given overleaf.

Chapter 4

I56 Steepest rate of increase in closure across area initiates spreading Overall closure rate for region . . . . cm/a 14--

Closure rate without smoothing effect of spreading

.....

Spreading event /

/

\

\ -~

[

12 -

/ /

10 -

/

\

\

/

/

amphibolite dates

\

I subduction + spreading

/

I

-

~

~

~.,.

subduction overthr+usting

-

J closure

t 1 -

x\

/ / I

\

\

spreading

I ~

spreading event

i

4

i

\ 6

bined stratigraphic thickness of 1500-2000 m and with generally coherent internal stratigraphies (Section 2.3, Fig. 4.13, Plate 4.7). The higher thrust slices of Hawasina sediments are thinner and less continuous than the lower ones and consist of a highly deformed mixture of Upper Triassic to Lower Cretaceous distal cherts and thin bedded pelagic limestones (Halfa and Haliw Formations), large masses of Permo-Triassic limestones (Oman Exotics) and redeposited Upper Triassic limestone breccias and pelagic sediments (AI Aridh Formation). Basaltic sills and volcanic layers, dated as Middle Jurassic to mid Cretaceous (Lippard & Rex 1982), are present in some of these sediments. Tectonic thickening (x 2.5-3.0 in the lower and more in the upper units) of each thrust sheet is the result of secondary thrust/reverse faulting and associated folding. The underlying Sumeini Group limestones form the lowermost thrust unit, the base of which is not seen, in the core of the window (Jebel Reis). The folds are usually asymmetrical tight to isoclinal, semi-recumbent or reclined structures and the thrusts are usually parallel to bedding with small displacements on the short overturned fold limbs. The folds have NW-SE axes parallel to the overall trend of the Hawasina Window. Folding of competent sandstone and limestone layers is by flexural slip followed by flattening (Plate 4.8). Incompetent

I 104

I 102

I 100

I 98

\.1_ 96 Ma

/ _~" 94

I 92

I 90

, p;,,,

I 88

Fig. 4.12. Postulated changes in the rate of consumption of crust during the Late Cretaceous at the NE Arabian margin.

',

-,,

,, ,,, ) .

"',,,

9 " "'[,~ H~',l

I I

?/

X ~

,,./ [ ^~v,:'///'/ ~".~..

I

Semad

"

I

/// .... / HM

time subduction is occurring outside the system in the northern Tethys and the margin is capable of extension. In Figures 4.11I-M the section of emplaced oceanic lithosphere is generalized. The space form of the Semail Nappe varies (Section 4.3.3.3) and has often been drastically affected by post-emplacement folding, thrusting, uplift and erosion. Similarly, as the intent is to illustrate the overall mechanism, it is not possible to accurately portray the thrusting in the Haybi, Hawasina and Sumeini units at this scale. They are shown schematically to demonstrate the structural setting and timing of this complex thrust system. Details of the stratigraphies and their relative palaeoenvironments have been given in Chapter II, the structural details are given in the following section.

Sor.~a,I

Semail nappe HL'I Hawasma melango Serpentmrte melange

/ ///j Hamrat Duru

...... ,,j /*/" .

0 I

4.3.3.1

Thrusting in the sedimentary nappes

The telescoping of the Oman continental margin and the Hawasina basin along a northeasterly axis produced the largely sedimentary allochthonous units that both underlie (Hawasina Assemblage) and overlie (Batinah Sediment Sheets) the Semail Nappe. These units consist of numerous thrust slices of Mesozoic hemi-pelagic and pelagic sediments that show complex geometries, several phases of syn-emplacement folding with associated cleavage formation and some low grade metamorphism. In the central Oman Mountains the sub-ophiolite sedimentary nappes are best exposed in the Hawasina Window (Plate 4.6). Here, the Hawasina Assemblage consists of three lower thrust sheets of relatively proximal facies turbiditic sediments (Hamrat Duru, Wahrah and AI Ayn Formations) with a corn-

, .

/ / /,

5km i Semad

SW A

NE B

Fig. 4.13. Tectonic map of the Hawasina Window showing major thrust faults and simplified cross-section of structural styles. Arrows indicate facing directions of main folds. After Graham (1980a, b).

Ophiolite Detachment, Emplacement and Subsequent Deformation

I57

Plate 4.6. Landsat MSS image at a scale of 1: 1000,000. The northern Oman ophiolite outcrops. From the top of the image the Aswad Block with

Wadi Hatta at its lower limit containing Hawasina and Tertiary outcrops. The large block beneath comprises the Fizh (and Sumeini) blocks, the major break in the centre of the chain is Wadi Jizi. South of Jizi is the Salahi block in which the Moho, dividing the lighter Crustal Sequence than the darker Mantle Sequence, is particularly well defined. In the lower right of the image the Hawasina Window is flanked by the Sarami block to the north, the Haylayn block to the east and Wuqbah block to the west. Inland the Tertiary outcrops of Jebels Hafit and Awaynah are prominent and on the Batinah coastal plain the linear outcrop marks the base of the Tertiary to the east of the Mountains. shales are deformed by homogeneous shear and in most cases have developed a secondary cleavage. In the centimetre-thick bedded cherts, buckling and flattening have resulted in the formation of kink folds. The thrusting of the sedimentary nappes, the Sumeini and Hawasini allochthons, was "thin-skinned" and no basement rocks are involved, both the Hawasina Assemblage and Haybi Complex can be described as "duplexes" as defined by Elliot (1977); Searle (1984). Typically, they are highly variable in thickness; for example, in the Hawasina Window the Hawasina thrust sheets form a tectonically thickened and folded nappe pile probably 2-3 km thick that thins rapidly to the southeast so that near Rustaq the Semail Nappe rests on only a

thin Hawasina unit or directly on top of the autochthonous Ha jar Supergroup. The Hawasina nappes were partly overridden by the Semail Nappe, as in the Hawasina Window, but elsewhere, e.g. in the Hamrat Duru Range to the SE, were emplaced in front of it. An estimated tectonic thickening factor (tectonic versus original stratigraphic thickness) of 2.53.0 for the Hawasina thrust sheets in the Hawasina Window compares to only 1.1-1.5 in the Hamrat Duru area (Graham 1980a) indicating the different tectonic regimes of these two areas. G r a h a m (1980 a, b) recognized four phases of deformation, based on studies of meso- and megascopic fold structures, in the Hawasina sediments:

~58

Chapter 4

Plate 4.7. Guwayza Formation in the Hawasina Window bedded limestones with lenticular conglomerate bands.

Plate 4.8. Flat-lyingF 2 folds in sandstone/shale sequence.

1 F~ tight, isoclinal to semi-recumbent SW-facing folds. 2 F 2 tight, semi-recumbent to upright NE-facing folds (most of the major folds in the centre of the window: eg. the Jebel Reis anticline a r e F 2 structures). 3 F3 recumbent folds facing W or SW that are mostly found along the southwest side of the window. 4 Gentle open folding on upright axes. On the northern edge of the window F~ folds are flattened and refolded by F2, locally they are preserved as isolated fold hinges with a n F 2 foliation surrounding them. Both F l and F 2 folds are associated with a penetrative cleavage in less competent lithologies. The cleavage increases in intensity downwards in the nappe pile and is particularly strongly developed

in the pelites and semi-pelites of the Zulla Formation which have been converted into chlorite-muscovite-biotite schists at the base of the Hamrat Duru thrust sheet. Graham (1980b) concluded that deformation phases F~_3 occurred during late Cretaceous nappe emplacement. The earliest FI, SW-directed structures probably formed during the initial assembly of the Hawasina nappes before emplacement onto the continental margin. The later NE-directed F 2 structures are difficult to explain. They may have formed as a result of excessive piling up of the nappes in this area as a result of emplacement over a ramp or basement ridge (possibly the shelf-slope break of the continental margin?), which led to thrusting in a direction opposite to that of the main emplacement. The northern edge of the window is complex with

Ophiolite Detachment,

Emplacement

imbricated slices of serpentinite, Haybi Complex and upper Hawasina units. The thrusts are vertical or steeply dipping to the south and the rocks are overturned and downward-facing to the north. These northward-facing structures are probably related to the F 2 structures in the Hawasina nappes. On the SW margin of the window the Hawasina nappes are locally thrust over Haybi Complex and serpentinite on NE-dipping thrust planes. This has probably resulted from rethrusting leading to a reversal of the normal sequence of nappes; "out of sequence" thrusting e.g. Searle (1984). Graham (1980b) concluded that the NE and SW-directed structures on opposing flanks of the Hawasina Window resulted from syn-emplacement uplift of the window causing sliding of the Semail Nappe on either flank. In summary, the following sequence of events in the formation of the Hawasina Window is envisaged. Early thrusting and assembly of the Hawasina nappes with associated SW-facing (F1) folds preceded emplacement onto the continental edge. Rethrusting of the Hawasina nappes in a duplex type structure with a structural thickness of 3-5 km in the Hawasina Window resulted in fracturing of the overlying Semail Nappe and its sliding off on either flank leading to NE-directed thrusting and imbrication on the northern side and SW-directed thrusting on the southern side. The late thrusting and folding ( F 2 and F~ fold phases) resulted in reorganization of the nappes and "out of sequence" thrusting. Graham (1980b) considers that the broadly anticlinal structure of the window is largely a late Cretaceous feature, although probably accentuated by midTertiary folding (F4 phase). Another series of thrust sheets of Hawasina-type sediments, the Batinah Sediment Sheets, structurally overlie the Semail Nappe and Batinah Melange along the eastern edge of the mountains between wadis Jizi and Ahin. The Batinah Sediment Sheets consist of three nappes, called by Woodcock & Robertson (1982b) the northern, central and southern sheets. These authors divide the sediments within the sheets into lower Sakhin (Upper Triassic-Lower Jurassic) and upper Salahi (early-mid Jurassic to mid Cretaceous) Formations which are equivalent to parts of the Hamrat Duru Group and Wahrah Formations of the Hawasina (Giennie et al. 1974). In the northern and southern Batinah sheets the rocks are largely inverted, in the central sheet they are predominantly rightway-up. The southern sheet apparently tectonically underlies the central sheet, but the relationship between the central and northern sheets is unclear. The bedding of the sediments in the northern sheet dips gently at about 15~ towards the SE. Mesoscopic open folds plunge gently northwards and are westwardfacing, suggesting that the northern sheet forms the inverted limb of a major west-facing recumbent fold and that it was therefore emplaced from east to west. The central sheet dips to the E or SE and the beds are folded into eastward-plunging gently dipping folds which predominantly face towards the south. Large mesoscopic folds have NE-trending axes and, in one case, deform the basal Tertiary sediments which unconformably overlie the Batinah Sheets. The southern sheet is divided into two klippen. The northernmost has a steeply northward-dipping contact with the underlying central sheet with beds that are inverted and dip to the NE with NE-trending folds. In the southern klippe the structures are more complex with areas of both right-way-up and inverted strata and often moderate to steep dips. Fold axes trend NE-SW and face towards the NW. Thus, whereas the northern and central sheets were emplaced towards the west or southwest, perhaps as the inverted and right-way-up limbs of a

and Subsequent Deformation

I59

major recumbent fold, the southern sheet was apparently emplaced at a later date towards the NW. As discussed by Woodcock & Robertson (1982a), the origin and provenance of the Batinah Sheets poses a palinspastic problem. Glennie et al. (1974) and Graham (1980b) recognized their similarity to the Hawasina sediments and suggested that they are parts of the Hawasina allochthon previously overridden by the Semail Nappe and then emplaced on its trailing edge. Woodcock & Robertson (1982a) consider that the deformation state of the sediments is incompatible with their ever having been beneath the ophiolite nappe and suggest that either (1) they represent parts of the Oman continental margin never overridden by the ophiolite, or (2) they are parts of an opposite (Iranian?) margin to the Tethys. The latter idea seems unlikely as there is no evidence for complete closure in the late Cretaceous of the Oman Tethys of which the present Gulf of Oman appears to be an unconsumed remnant. Their first alternative also has problems in that the continuity and lateral extent of the Semail Nappe over a strike distance of over 600 km make it appear unlikely that any part of the Oman margin escaped being buried by the Semail Nappe during its emplacement. In fact the structures cited by Woodcock & Robertson (1982b) as evidence for a superficial history for the Batinah Sheets during their emplacement; boudinage in competent lithologies, en-echelon tension gashes filled with calcite and a lack of tectonic cleavage, are typical of much of the Hawasina sediments which structurally underlie the Semail Nappe. Thus, in our view, it is most likely that the Batinah Sheets were part of the Hawasina allochthon initially overridden by the ophiolite nappe but then emplaced onto its trailing edge during the last stages of nappe emplacement. Emplacement of the Batinah sediment sheets clearly postdated formation and emplacement of the Batinah Melange and probably occurred as late as the Campanian. In both the Musandam Peninsula and the northern part of the Saih Hatat window parts of the "autochthonous" carbonate succession are clearly involved in the thrusting, although the latest thrust movements of the Ruus al Jebel limestones over the Hagab thrust were mid-Tertiary (Hudson et al. 1954; Searle et al. 1983). Thus, only the basement rocks of Jebel Akhdar and the southern part of Saih Hatat are considered to be truly autochthonous. Even this has been queried by S. Hanna (pers. comm. 1983) who believes that a major thrust underlies those areas, possibly as deep as 10-15 km, and that the Jebel Akhdar and Saih Hatat anticline are major culminations formed above footwall ramps (Dahlstrom 1970) (Plate 4.9). 4.3.3.2 Tectonic Melanges - Haybi and Batinah complexes

Tectonic melanges, associated with late Cretaceous nappe emplacement, are found both below and above the Semail Nappe and in fault zones that cut through it (Lippard et al. 1983) (Fig. 4.14). The sub-ophiolite melange, which is a part of the Haybi Complex, is a sheared serpentinite, called the Basal Serpentinite by Searle (1980), containing metre to kilometre sized blocks of the Metamorphic Sheet, Haybi volcanics, Exotic limestones and Hawasina sediments. Although local concordance of structures between blocks indicates passive break-up by serpentinite protrusion (Searle 1980), dispersal and rotation of the blocks has led to preferred orientations which may be used as an indicator of tectonic movement directions (e.g. NE-SW in the Haybi corridor). There are all gradations between more-or-less 100% serpentinite and a

160

Chapter 4

Plate 4.9. Landsat MSS image at a scale of 1:1000,000. Central to the image are the autochthonous culminations of Jebels Akhdar and Nahkl. To the north are the Haylayn and Rustaq ophiolite blocks. East across the Semail Gap is the Ibra block and south the Bahala block. Between that and the Muqniyat block to the west are the Exotic limestones of Jebel Khawr. The southern part of the Hawasina Window is visible between the Muqniyat and Haylayn blocks and part of the Saih Hatat Window is shown in the upper right of the image. The arcuate culminations of Jebels Salakh and Madamar at the lower edge of the image mark the limit of thrusting. The Jebels are of autochthonous Wasia Group, between them and the ophiolites are the fold and thrust belt of allochthonous Hawasina sediments.

block-on-block melange containing little or no matrix (Plate 4.10). The block-on-block melange is most common in imbricate thrust slices and narrow fault zones between the ophiolite blocks (Fig. 4.14). As the melange is traced to higher tectonic levels through the ophiolite it contains an increasing amount of ophiolite material and grades upwards into the supra-ophiolite Batinah Melange (Lippard et al. 1983; Robertson and Woodcock 1983a). This melange overlies the Semail Nappe along the eastern edge of the mountains and its main features are:

1 It locally rests with a depositional contact on the top of the ophiolite where there is a gradation from the highest Semail pelagic sediments into the melange with exotic blocks in a sedimentary matrix. 2 It is for the most part, a block-on-block melange with little or no recognizable matrix. 3 Blocks of the Semail ophiolite within the melange correlate with the local basement although they are mixed with a variety of exotic blocks mostly clearly derived from

Ophiolite Detachment, Emplacement and Subsequent Deformation .

.

.

.

.

1

'"

=======================~::: .: ..... 56 ~ 10' E

~~

G o m pie xes - mainly melange

WADI HATTA

-,: . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . " ..... 11 56 ~ 20' E

'

,, W A D I OlZl

i: i!iiii iiiiiiiiiiiii ~ . : : ~ . ~ - ~ , o %

::::::::::::::::::::::::::

O0 (d:

iiiiiiii:ii}iiiiiii!i}:

~"

l,

,

}

o .~2>

56 ~ 30' E ,~:,'~:.~,.,,.;~-.o.. I ," t ~ 0.

i i!

S E M A I L N~,'PP E

~<::~

~ " o~

.....

Plate 4.10. (a) Batinah Melange. Block-upon-block structure of varicoloured sediments, volcanics and serpentinite. (b) Batinah Melange in Wadi Ahin. Top of ophiolite at left of picture marked by dark umbers and massive Salahi Unit flows dipping to right (east). Melange forms most of middle ground with typical hummocky topography containing upstanding masses of exotic limestones.

I

Extrusive Sequence Dyke Complex

~Layered Series and H i g h level Intrusive Complexes ~

WADI AHIN

Wiii:!r ~ . ........................

::::.,~-d:: :~. . . .

~Sheeted

o.,

.o.O:oO ~, o;

:::::::::::::::::::::::::::

~

I6t

Normal stratigraphic contact

~ I.

High-angle fault

9

T h r u s t fault

The Batinah Melange forms the lower part of the Batinah Complex (Woodcock & Robertson 1982a) and is tectonically overlain by the Batinah Sediment Sheets.

Mantle Sequence

Fig. 4.14. Structural setting of Haybi and Batinah melanges in three between-block corridors.

beneath the ophiolite nappe (metamorphic rocks, serpentinite, Exotic limestones). 4 It shows marked thickness variations with the maximum developments in the W a d i Jizi and Wadi Ahin fault zones and least in the intervening areas where it may be locally absent (Fig. 4.14). Robertson and Woodcock (1983a) concluded that the Batinah Melange formed beneath the Semail Nappe during its emplacement and was emplaced as a near-solid slurry through fault zones onto the ophiolite surface. This process can be seen in an arrested stage in the Wadi Jizi and Wadi Ahin (Sakhin) corridors. The lower part of the melange was extruded onto the ophiolite whilst pelagic sedimentation was still continuing.

4.3.3.3

The Semail Nappe

It appears that the Semail Nappe was emplaced as a largely intact sheet or slab until folding and faulting broke it up during the last stage of emplacement onto the continental margin. Glennie et al. (1974 p.346) noted that . . . "structurally the Semail (Nappe) forms a number of plates that are more or less detached from one another . . . some are separated by straight, vertical transcurrent or scissor faults others by major fold axes . . . almost every plate has a slightly different attitude from its neighbours" (Plate 4.6). These authors suggested that the tectonic dislocation of the nappe was largely the result of Tertiary folding and faulting and subsequent erosion. Although this may be true in some cases, it is clear that the majority of the ophiolite "blocks" (as they are here termed) were separated from one another during, rather than after, late Cretaceous emplacement. The reasons for the early breakup of the Semail Nappe appear to have been: (i) irregularities of the surface over which it travelled; e.g. in the Hawasina Window which was the site of a major tectonic culmination

I62

Chapter 4

owing to the piling up of the underlying Hawasina nappes, which in turn was probably located over a basement high; and (ii) collapse of the ophiolite blocks into pre-existing tectonic depressions. Once break-up had been initiated it was accentuated and possibly accelerated by (iii) the upward movement of low density serpentinite-matrix melange and sedimentary rocks along fault zones between the blocks. The edges of the blocks are marked by three types of structure: (a) Imbricate faults; these mainly occur at the leading edge of the nappe and arrested its forward movement. As compressional features, they are rarely associated with wide melange zones, but usually with a thin layer of sheared serpentinite. The largest and most obvious example of an imbricate structure is the West Jizi block which structurally underlies the Fizh and Salahi blocks at the west end of Wadi Jizi. (b) Cross-strike fault zones; these trend from NE-SW to NW-SE and are high-angle fault zones which cut obliquely across the strike of the ophiolite approximately parallel to the probable direction of emplacement of the nappe. They are usually straight and consist of several parallel faults and have large apparent variations in throw along their lengths. In most cases (Wadi Jizi, Wadi Ahin (Sakhin corridor), Ghuzayn and Rustaq corridors) the fault zones are occupied by 5 to 10 km wide zones of melange grading downwards into a broken formation and then largely intact sequences of Haybi Complex and Hawasina sedimentary units (Fig. 4.14). The rocks within the corridors clearly become more disaggregated and chaotic as they are traced upwards towards the surface of the nappe. (c) Major anticlinal uplifts; such as the Hawasina-Jebel A k h d a r and Saih Hatat anticlines (Plate 4.9). In the central and southern mountains ophiolite blocks lie on either flank of these structures and are folded into subsidiary flanking synclines and anticlines parallel to the main fold axes (Fig. 4.16). In many cases it is difficult to distinguish between late Cretaceous and mid-Tertiary ages for these structures. The thickness of the Semail Nappe varies markedly but since the acquisition of the gravity data (Shelton 1984) estimates of the depth of ophiolite can be made. Enclosure Two shows the Simple Bouguer A n o m a l y field for a large part of the region. The subsequent interpretation of these data suffer from two sources of inaccuracy, (i) the difficulty in determining the regional gravity field and (ii), the variability in density contrast between the ophiolite and its surroundings. The first problem is due to the fact that the area of interest is over a passive margin which gives rise to a field having variations of the same order and frequency as that due to the ophiolite nappes. The second is due to the variable degree of serpentinization that affects the Mantle Sequence. This can give rise to ophiolite densities that are less than those of the underlying carbonates (Section 3.2.5.2). Shelton (1984), allowing for these problems, gives a number of 2D gravity interpretations of the present form of the ophiolite nappe (Fig. 4.15). They show the nappe as having an essentially flat basal thrust which cuts up-section at the trailing edge. As a result the nappe does not extend further than 15 km from its trailing edge outcrop limits. This is interpreted as further evidence that the final stages of emplacement were dominated by gravity-gliding mechanisms. The thickness of the ophiolite nappes appears less in the central mountains and an indication of this is given in Table 4.7. The preferred density contrast of 0.52 g/cc shows the approximate depth of ophiolite beneath the maximum gravity anomaly on that

O l

1 FIZH 60 70

50

40

30

20

10

Okra 0km

2 SHAFAN . . . . . . . .

.

.

.

.

3 GHUZAYN

4 RUSTAQ

Fig. 4.15.

True-scale cross sections of the Semail Nappe deduced from the gravity data. Zero km distance taken from coastline. From Shelton (1984).

Table 4.7. Maximum thickness of principal ophiolite blocks calculated with the horizontal thin sheet approximation for an average block width of 20 km. Values are given in km for maximum, preferred and minimum density contrasts in g cm -~. From Shelton (1984). Block Name

N. FIZH S. FIZH SALAHI SARAMI SHAFAN+ GHUZAYN HAYLAYN RUSTAQ

Max ~ Anomaly

131-145 121-12b 110-? 73-? 66-68 88-92 87-'? 114

U I T M posn Calculated thickness (km) as A e of max 0.80 0.52 0.25

2725/430 2710/445 2710/445 2680/450 2645/475 2635/500 2610/530 2597/575

4.2-4.7 3.9-4.0 3.3-? 2.3-'? 2.0-2.1 2.8-2.9 2.7-'? 3.6

6.7-7.5 6.1-6.4 5.5-'? 3.6-'? 3.2-3.3 4.3-4.5 4.2-'? 5.7

=

16.0-18.0 14.5-15.2 12.8-2 7.9-? 7.1- 7.3 9.9-10.3 9.7-? 13.3

* SBA calculated at 2.67, less modelled regional with zero datum set equal to -62 mgal observed. * Refers to unexposed (ophiolite) block under Batinah cover off Sarami Block. particular block. The extreme density contrasts illustrate the possible range of thicknesses.

4.3.3.4 Nappe emplacement; overview The Oman Mountains thrust belt, which broadly coincides with the present extent of the mountains, is about 700 km long and more than 200 km wide. The thrust belt is arcuate through more than 90 ~ of arc and is concave to the northeast towards the Gulf of O m a n ; a shape which appears to be an original

Ophiolite D e t a c h m e n t , E m p l a c e m e n t and S u b s e q u e n t D e f o r m a t i o n feature of its emplacement. The thickness of the nappe pile is at least 20 km but varies considerably along and across strike and many of the units are highly lenticular. There is no obvious basal decollement, although shaly horizons in the lowermost Triassic units of the Hawasina Assemblage may have acted as detachment planes at the base of the Hawasina nappes. Regional considerations and tectonic emplacement direction indicators, such as orientations and facing directions of thrust planes and related folds, show that the thrust nappes of the Oman Mountains were emplaced broadly from NE to SW. In the Dibba Zone, at the northern end of the thrust belt, emplacement was directed towards the WNW (Lippard et al. 1982). Searle et al. (1983) and Searle (1980) recognized an E to

W emplacement direction in the Sumeini area, and Graham (1980) used major and minor fold orientations and the elongation directions of blocks in tectonic melanges to infer NE to SW emplacement in the Hawasina Window. Further SE, in the Saih Hatat area south of Muscat, Boudier et al. (1982) showed that here emplacement occurred towards the SSW. From these results, it appears that the directions of nappe emplacement were radially directed around the present arcuate shape of the mountains (Fig. 4.16). The distances of nappe transport are difficult to estimate and depend on palinspastic reconstructions of the Oman continental margin and the width of the adjacent Tethys ocean prior to nappe emplacement. Glennie et al. (1974) estimated a distance

I

56 ~

I63

57 ~

58~

Sheeted dykes and extrusives

.). "... ......

( Crustal Gabbros and cumulate peridotites ! sequence

! ".,." .'..'.',

Tectonite peridotite = Mantle sequence

:..:.i

Semail thrust (Base of ophiolite nappe)

:.:. "~". ...'.... ..

Thrust fault (Barb on overthrust block)

:., " 'i

25 ~ N

High-angle faults (Tick on downthrow side, where known) Major faults along edges of ophiolite "blocks"

4

Numbered ophiolite block (see figure caption)

'-.,.,......... ",.,.~......,. Stratigraphic or uncertain contact of ophiloite

Eroded edge of ophiolite outcrop

....,.'..'.~ ..' ..'.,~

'.. :," ."

24~

::"~"~

0

I

" .. .~. ....

9 I "J.

...

50 k m

I \'.....'....'..

9

k

.

.f

,.__,1Fig. 4.16. S t r u c t u r a l s u m m a r y o f N o r t h e r n O m a n .

~ _

~

/---<, 9149

!9 ?~

r64

Chapter 4

of 400-1200 km for the "half-width" of the Hawasina Ocean, that is the distance between the shelf-edge and the spreading ridge where the Semail ophiolite was generated. Graham (1980b) obtained a distance of about 350 km for the original widths of the Sumeini and Hawasina thrust sheets in the Hawasina Window taking into account an estimated average tectonic thickening of 2.4 during emplacement. By adding 50 km, which is the present maximum cross-strike width of the Semail Nappe, it seems that the Semail ophiolite may have been translated at least 400 km from its site of formation. It must be remembered that this is an approximate figure as the tectonic deformation of the upper Hawasina units and Haybi Complex rocks is too great to give a reasonable estimate of their original palaeogeographic extent. Palaeomagnetic data on the Semail dykes and gabbros (Section 3.12) introduce a further constraint suggesting that the Semail Nappe moved through less than 400 km of latitude during its emplacement. The stacking sequence of the nappes developed from NE to SW from the oceanic area towards the continental foreland. As illustrated in Fig. 4.12, the Semail Nappe was detached from the oceanic lithosphere in the Turonian-Santonian (91-83 Ma), probably along a low-angle thrust. The stratigraphic thickness of the Semail ophiolite suggests that this thrust extended at least 15-20 km into the sub-oceanic mantle (Searle & Malpas 1980). The early lineation directions in the basal peridotites and metamorphic sheet are different from the later emplacement vectors implied from fold axis and thrust plane orientations suggesting at least two phases of movement; an early intra-oceanic stage ("displacement" stage of Searle & Malpas (1980)) and a later continental ("emplacement") stage of thrusting. The youngest continental margin sediments involved in the thrusting beneath the Semail Nappe are of Cenomanian age (Glennie et al. 1974). Emplacement of the Mesozoic continental margin rocks (Hawasina, Sumeini units) commenced soon after this, probably associated with the formation of sedimentary melanges (olistostromes) during early tectonic instability (Graham 1980a). The absence of ophiolite melange between the Hawasina thrust sheets confirms that the Hawasina nappes were assembled before they were overridden by the Semail Nappe. Thrusting progressed towards the continental margin and the Sumeini thrust sheets were displaced beneath the Hawasina nappes. All these units were emplaced onto the continental margin and overrode the Coniacian-Santonian (89-83 Ma), Muti formation which had formed earlier in response to subsidence of the continental shelf following Turonian uplift and erosion (Glennie et al. 1974). Sedimentary debris was shed off the Hawasina nappes to form the Fiqa and Juweiza formations. In the Santonian-Campanian (c. 80 Ma) the sedimentary nappes were overridden by the Semail Nappe. The upper parts of the Hawasina thrust sheets were considerably broken up and tectonic melanges of Mesozoic pelagic sediments and Triassic ocean floor rocks, including Exotic limestone masses, were formed. The metamorphic sheet was broken up and dispersed in a serpentinite melange derived from the Semail Nappe and Haybi Complex. The plane of thrusting between the ophiolite and the underlying units at the base of the Semail Nappe (Basal Serpentinite) approximately followed the initial intra-oceanic displacement surface, but locally cut up through the ophiolite section towards the leading and trailing edges of the nappe and along cross-strike fault zones. Melange formed beneath the nappe was locally forced up these fault zones and extruded onto the nappe surface to form the Batinah Melange (Robertson & Woodcock 1982a).

Many of the features of the Semail Nappe; its wedge-shape, imbrication at the leading edge, extension at the trailing edge and a decollement layer of deformed serpentinite at the base, are consistent with gravity-driven movement during the last stage of emplacement across the continental margin. 4.3.3.5 Post emplacement tectonics

The Upper Cretaceous nappe emplacement event was caused by the interaction of the Oman passive margin with the subduction zone. Subduction was transferred at that time to the northern margins of the Tethyan ocean. The next major compressional event occurred in the lower Miocene and must correspond to the collision of the Arabian margin with the accretionary wedge of that northern subduction zone. During the nappe emplacement event the depression of the inner margin due to a combination of the isostatic rebound of the outer margin and the emplacement of the Hawasina and Semail Nappes resulted in a transgressive period. Sedimentary formations in the north-eastern UAE (Skelton & Nolan 1985) indicate an eastward transgression onto the western flank of the ophiolite during early to mid-Maastrichtian times. The Qahlah Formation, a beach environment deposit, lies unconformable on lateritized ophiolite. The succeeding Simsima Formation represents lower foreshore to open marine shelf facies. The sequence of post emplacement events is more clearly distinguished in the north of Oman and Musandam peninsula. Due to erosion following mid to late Tertiary uplift there are no Lower Tertiary deposits in the central part of the mountains which means that the late Cretaceous or mid-Tertiary age of many of the structures is in dispute (Glennie et al. 1974; Searle et al. 1983; S. Hanna p e r s . c o m m , 1983). Mid-Tertiary thrusting has only been proved on the Hagab thrust on the western side of the Ruus al Jebel (Musandam Peninsula) (Hudson et al. 1954; Searle et al. 1983) and involved translation distances of no more than five to ten kilometres. This approximately E-W compressional phase is present throughout the Middle East and is generally dated as Lower Miocene. It gave rise to large scale open folding on upright axes, mainly parallel to the arcuate trend of the mountains. Ricateau and Rich6 (1980) attribute the creation of the Musandam ridge and the Oman mountains to this phase. They also invoke vertical "movements of compensation" arising from the release of compression. Seismic studies of the Gulf of Oman do show continuing subsidence with flat-lying sedimentation from the Upper Miocene onwards. Beneath the Upper Miocene levels there is abundant faulting indicating compression in the Lower Miocene and distension prior to the Pliocene. The planated surface of the Musandam peninsula appears to have suffered submergence then uplift of the southern part in the Pleistocene that produced a tilt which has given rise to fluvial terraces, Quaternary raised beaches, inclination of the Musandam peneplain and the submergence of the valleys of the NE coast (Ricateau & Riche 1980). Vita-Finzi (1982) estimates a subsidence of 30 m in the past 10,000 years for this area. The Dibba Thrust Zone, despite showing little strike-slip movement, still affects the region. The tilting affecting the Musandam, possibly indicative of its engagement with the subduction zone, is disconnected from northern Oman. Here the Late Cretaceous sedimentation has been followed by progressive uplift as illustrated by the elevated, but undated, gravel terraces at 300 m on the flanks of J. Nahkl and NW Saih

Ophiolite Detachment, Emplacement and Subsequent Deformation Hatat. Recently, the Oman area appears stable; the 'subrecent' shells of Lees (1928) at 375 m.a.s.1, thought to indicate recent uplift, are apparently not in situ (Vita-Finzi 1982), and carbon dating of shells close to the present high-water mark shows less than 1 m of vertical movement in the past 5000 years. Following the mid-Tertiary tectonism the uplift has resulted in considerable erosion. Glennie et al. (1974) estimate that over 4 km thickness of Cenozoic molasse off the Batinah coast represents 100,000 km 3 of erosion from the Lower Tertiary, Semail and Hawasina nappes. If the nappe area is taken to be

I65

approximately 30,000 km 2 then over 3 km of erosion has taken place. It is probable that the erosion was more pronounced on the steeper seaward gradient of the uplifted margin than on the gentler inland slopes. This process is seen in the present Red Sea margin drainage pattern and satisfactorily explains the relative lack of erosional material west of the mountains. In the centre of the Jebel Akhdar area erosion has stripped off up to 10 km of rocks including the whole of the Permian-Upper Cretaceous carbonate sequence (3 km thick) beneath which it has exposed the pre-Permian basement rocks.

References

ANON 1963. Geologic Map of the Arabian Peninsula. USGS Missc.

ABBAS, S.G. & AHMED, Z. 1979. The Muslimbagh Ophiolites, Pakistan. In: FARAH, A. & DE JONG, K.A. (eds) Geodynamics of Pakistan. Geol. Surv. of Pakistan, Quetta, 243-250. ABBATE, E., BORTOLO1TI, V. & PASERINI, P. 1976. Major structural events related to ophiolites of the Tethyan Belt. Ofioliti 1, 5-32. ABBOrrS, I.L. 1978. High-potassium granites in the Masirah ophiolite of Oman. Geol. Mag. 115,415-425. ABBOrrS, I.L. 1981. Masirah (Oman) ophiolite sheeted dykes and pillow lavas: geochemical evidence of the former ocean ridge environment. Lithos 14,282-294. ADAMIA, S.A., CHKOTUA, T., KEKELIA, M., LORDKIPANIDZE, M., SHAVISTIVILIC, I. & ZAKARIADZE,G. 1981. Tectonics of the Caucasus and adjoining regions: implications for the evolution of the Tethys ocean. J. Struct. Geol. 3,437-447. ADm, D. 1978. Rodingites from the Neyriz ophiolitic complex. Neues.

Geol. Investigations Map 1-270A. ANON 1972. Penrose field conference on ophiolites. Geotimes 17, 24-25. ARTHURTON, R.S., FARAH, A. & AHMED, W. 1982. The late Cretaceous Cenozoic history of western Baluchistan, Pakistan - The northern margin of the Makran subduction complex. In: LEGGETr, J.K. (ed) Trench- Forearc Geology, Spec. Publ. Geol. Soc. London 10, 373-385. ARVIN, M. 1982. Petrology and geochemistry of ophiolites and associated rocks from the Zagros suture, Neyriz, lran. Unpubi. PhD. Thesis, University of London. ASRARULLA, AHMED, Z. & ABBAS, S.G. 1979. Ophiolites in Pakistan: an introduction. In: FARAH, A. & DEJON6 (eds), Geodynamics of Pakistan, Geol. Surv. Pakistan. Quetta, 181-193. AUBOUIN, J. 1965. Geosvnclines. Developments in Geotectonics 1. Elsevier Publishing Co., Amsterdam-London-New York, 335pp. AUGE, T. & ROBERTS, S. 1982. Petrology and geochemistry of some chromitiferous bodies within the Oman ophiolite. Ofioliti 23, 133-154. AVE LALLEMANT, M.G. & CARTER, N.L. 1970. Syntectonic recrystallisation of olivine and modes of flow in the upper mantle. Bull. geol. Soc. Am. 81, 2203-2220. BAILEY. D.K. 1980. Volcanism, Earth degassing and replenished lithosphere mantle. Phil. Trans. Roy. Soc. Lond. A297, 309-322. BAILEY, D.K. 1982. Mantle metasomatism - continuing chemical change within the Earth. Nature 296,525-530. BAILEY, E.H. (ed). 1981. Geologic Map of the Muscat - Ibra Area, Sultanate of Oman. J. Geophys. Res. 86 (B4), (enclosure). BAILEY, E.H. & COLEMAN, R.J. 1975. Mineral deposits in the Samail Ophiolite of Northern Oman. Geol. Soc. Am. Abstr. Programme 7 (3), 293. BALLARD, R.D. & VAN ANDEL, T.H. 1977. Morphology and tectonics of the inner rift valley at Lat. 36~ on the Mid-Atlantic Ridge.

Jb. geol. Palaont. Abh. ADIB, D. & PAMIC, J. 1980. Ultramafic and mafic cumulates from the Neyriz ophiolite complex in SE parts of the Zagros range (Iran). In: PANAYIOTOUA. (ed.) Ophiolites, Proceedings of the International Ophiolite Symposium, Cyprus 1979, 392-397. AGRAWAL, O.P. & KACKER, R.N. 1980. Nagaland ophiolites, India - a subduction zone ophiolitecomplex in Tethyan orogenic belt. In: PANAVIOTOUA. (ed.) Ophiolites, Proceedings of the International Ophiolite Symposium, Cyprus 1979, 454-461. ALABASTER, T. 1982. The interrelationship between volcanic and hydrothermal processes in the Oman ophiolite. Unpubl. PhD Thesis, Open University, 408pp. ALABASTER, T. & PEARCE, J.A. 1985. The interrelationship between magmatic and ore forming hydrothermal processes in the Oman ophiolite. Econ. Geol. 80 (1), 1-16. ALABASTER, T., PEARCE, J.A., MALLICK, D.I.J. & EL BOUSHI I.M. 1980. The volcanic stratigraphy and location of massive sulphide deposits in the Oman ophiolite. In: PANAYIOTOUA. (ed.) Ophiolites, Proceedings of the International Ophiolite Symposium, Cyprus 1979, 751-757. ALABASTER, T., PEARCE, J.A. & MALPAS, J. 1982. The volcanic stratigraphy and petrogenesis of the Oman Ophiolite Complex. Contrib. Min. Pet. 81, 168-183. ALAVI-TEHRANI, N. 1976. Geology and petrography in the ophiolitic

Bull. geol. Soc. Am. 88,507-530. BARBERI, F. & VARET. J. 1977. Volcanism in Afar: small scale plate tectonic implications. Bull. geol. Soc. Am. 88, 1251-1266. BARBER1, F., FERRARA, G., SANTACROCE, R. & VARET, J. 1975. Structural evolution of an Afar triple junction. In: PILGER, A. & ROSLER, A. (eds) Afar Depression of Ethiopia. Nagele Obermillar, Stuttgart, 38-54. BARRETr, T.J. & SPOONER, E.T.C. 1977. Ophiolitic breccias associated with allochthonous oceanic crustal rocks in the East Ligurian Apennines, Italy - A comparison with observations from rifted oceanic ridges. Earth Plan. Sci. Lett. 35, 79-91. BARTHOLOMEW, I.D. 1983. The primary structures and fabrics of the

range NW of Sabzevar (Khorassan/lran), with special regard to metamorphism and genetic relations of the ophiolite suite. Dissertation Univ. Saarbrucken, 147pp. ALAVI-TEHRANI, N. 1977. Geology and petrography in the ophiolite range NW of Sabzevar (Khorassan/Iran) Geol. Surv. lran Rept. nr. 43. ALAVI-TEHRANI, N. 1980. The distribution of ophiolites in Iran and their significance. In: Rocct, G. (ed) Ofioliti Special Issue Tethyan Ophiolites Vol. 2 Eastern Area, 315-334. ALAVl, M. 1980. Tectonostratigrahic evolution of the Zagrosides of Iran. Geology 8, 144-149. ALDISS, D.T. 1978. Granitic rocks ofophiolites. Unpubl. PhD Thesis, Open University, 198pp. ALDISS, D.T. 1981. Plagiogranites from the ocean crust and ophiolites. Nature 289,577-578. ALLEGRE, C.J. et al. 1984. Structure and evolution of the HimalayanTibetan orogenic belt. Nature 307, 17-22. ALLEMANN, F. & PETERS, T. 1972. The ophiolite-radiolarite belt of the Northern Oman Mountains. Eclog. geol. Heh,. 65,657-697. ALLEN, C.R. 1975. The petrology of a portion of the Troodos plutonic complex, Cyprus. Unpubl. PhD Thesis, University of Cambridge. AMERICAN COMMISSION ON STRATIGRAPHIC NOMENCLATURE (ACNS) 1961. Codes of stratigraphic nomenclature. Bull. Am. Assoc. Petrol. Geol. 45,645-665. ANDERSON, R.N., HONNOREZ, J., BECKER, K., ADAMSON, A.C., ALT, J.C., EMMFRMANN, R., KEMPTON, P.D., KINOSHITA, H., LAVERNE, C., MOTrL, M.J. & NEWMARK, R.L. 1982. DSDP Hole 504B. The first reference section over 1 km through Layer 2 of the oceanic crust. Nature 300, 589-594,

Upper Mantle and Lower Crust from ophiolite complexes. Unpubl. PhD Thesis, Ope~ University, 523pp. BEAR, L.M. 1960. The geology and mineral resources of the AkakiLvthrodondha area. Cyprus Geol. Sum,. Dept Mem. 3, 122pp. Basal Group. Cyprus Geol. Surv. Dept. Ann. Report of 1958, 12-14. Bt-AR, L.M. 1960. The geology and mineral resources of the AkakiLythrodondha area. Cyprus Geol. Surv. Dept Mem. 3, 122. BEBIEN, J., BI,ANCHET, R., CADET, J.P., CHARVET, J., CHOROWITZ, J., LAP1ERRE, H. & RAMPNOUX, J.P. 1978. Le volcanisme Triasique des Pinarides en Yougoslavie, sa place dans l'evolution geotectoniques peri-Mediterraneenne. Tectonophysics 47, 159-176. BECCALUVA, L., OHNENSTETTER, D., OHNENSTE'FrER, m. a VENTUREI,LI, G. 1977. The trace element geochemistry of Corsican ophiolites. Contrib. Min. Petr. 64, 11-31. BECCAI,UVA, L., MACCIOTTA, G., PICCARDO, G.B. & ZEDOU, O. 1984. Petrology of [herzolitic rocks from the Northern Apennine ophioIRes. Lithos 17,299-316. BERBERIAN, M. & KING, G.C.P. 1981. Towards a palaeogeography and tectonic evolution of Iran. Can. J. Earth Sci. 18,210-265. BERBERIAN, F., MUIR, I.D., PANKHURST, R.J. & BERBERIAN, M. 1982. Late Cretaceous and early Mesozoic Andean-type plutonic acti167

I68

References

vity in northern Makran and Central Iran. J. geol. Soc. Lond. 139, 605-614. BERGOU6NAN, H. 1975. Relation entre les edifices Pontique et Taurique dans le Nord-Est de l'Anatolie. Bull. Soc. Geol. France (7), XVII, No 6. BERNOUILLI, D. & JENKYNS, H.C. 1974. Alpine, Mediterranean and Central Atlantic Mesozoic facies in relation to the early evolution of Tethys. In: DoTr, R.K. & SHAVER, R.H. (eds) Modern and Ancient Geosynclinal Sedimentation Soc. Econ. Palaeontologist Mineralogists Spec. Publ. 19, 12%160. BEYDOUN, Z.R. 1966. The geology of the Arabian Peninsula, Eastern Arabian Protectorate and part of Dhofar. US Geol. Surv. Prof. Paper 560H. US Govt Printing Office, Washington, 49pp. BEZZI, g . & PICCARDO, G.B. 1971. Structural features of the Ligurian ophiolites: petrologic evidence for the "oceanic" floor of the Northern Apennines geosyncline: a contribution to the problems of the Alpine-type gabbro-peridotite associations. Mere. Soc. Geol. Ital. 10, 55-63. BICKLE, M.J. & NISBET, E.G. 1972. The oceanic affinities of some alpine mafic rocks based on their Ti-Zr-Y contents. J. Geol. Soc. 128, 267-271. BICKLE, M.J. & PEARCEJ.A. 1975. Oceanic mafic rocks in the Eastern Alps. Contrib. Min. Pet. 49, 177-189. BIEHLER, J., CHEVALIER, C. & RICHATEAU, R. 1975. Carte geologique de ia peninsule de Musandam. BRGM Orleans, France. BISCHOFF, J.L. & DICKSON, F.W. 1979. Seawater-basalt interaction at 200~ and 500 bars: Implications for origin of sea floor heavy mineral deposits and regulation of seawater chemistry. Earth Plan. Sci. Lett. 25, 385-397. BLACK, P.M. & BROTHERS, R.N. 1977. Blueschist ophiolites in the melange zone, northern New Caledonia. Contrib. Min. Pet. 65, 6%78. BLANEORD, W.T. 1872. Note on Muskat and Mussandam on the coast of Arabia. Records Geol. Surv. India 5, 75-77. BLUMENTAL, M.M. 1963. Le system structurale du Taurus subanatolien. Livre de Memoire du Professeur Paul Fallot Mem. Soc. Geol. Fr., tome 2, 611-662. BODINIER, J.L., DUPREY, C., DIOSTAL, J. & CARME, F. 1981. Geochemistry of ophiolites from the Chamrousse complex, Belledonne Massif, Alp. Contrib. Min. Pet. 78,379-388. BOLES, J.R. & COOMBS, D.S. 1977. Zeolite facies alteration of sandstones in the Southland Syncline, New Zealand. Am. J. Sci. 277, 982-1012. BONAI~fI, E., HONNOREZ, J., KIRST, P. & RADICATI, F. 1975. Metagabbros from the Mid-Atlantic Ridge at 06~ contact-hydrotherreal-dynamic metamorphism beneath the axial valley. J. Geol. 85, 61-78. BOSSELLINI, A. & ROSSI, D. 1974. Triassic carbonate buildups of the Dolomites, Northern Italy. In: LAPORTE, L. (ed) Reef~ in Time and Space Soc. Econ. Palaeontologists Mineralogists Spec. Publ. 18, 209-233. BOSSELLINI, A. & WINTERER 1975. Pelagic limestone and radiolarite of the Tethyan Mesozoic: a genetic model. Geology 3, 279-282. BOUDIER, F. & COLEMAN, R.G. 1981. Cross section through the peridotite of the Samail ophiolite, southeastern Oman mountains. J. Geophys. Res. 86, 2573-2592. BOUDIER, F. & MICHARD, A. 1981. Oman ophiolites. The quiet obduction of ocean crust. Terra cognita 1, 109-118. BOUDIER, F., NICOLAS, A. & BOUCHEZ, J.L. 1982. Kinematics of oceanic thrusting and subduction from basal sections of ophiolites. Nature 296, 825-828. BOUDIER, F., NICOLAS, A., BOUCHEZ,J-L., CRAMBERT,S., DAHL, R. 8, JUTEAU, T. 1983. Les ophiolites des nappes de Semail (Oman): structures internes des massifs de Nakhl et de Rustaq. Sci. Geol. Bull. 36, 17-33. BOULIN, J. 1981. Afghanistan structure, greater India concept and eastern Tethys evolution. Tectonophysics 72, 261-287. BOYER, S.E. & ELLIOTI', D. 1982. Thrust Systems. Bull. Amer. Assoc. Petrol. Geol. 66 (9), 1196-1230. BRAACKMAN, J.H., LEVELL, B.K., MARTIN, J.H., POTTER, T.L. & VAN VLIET, A. 1982. Late Palaeozoic Gondwana glaciation in Oman. Nature 299, 48-50.

BRIDEN, J.C. 1967. Recurrent continental drift of Gondwanaland. Nature 215, 1334-1339. BRIDEN, J.C. & GASS, I.G. 1974. Plate movement and continental magmatism. Nature 248,650-653. BROOKFIELD, M.E. 1977. The emplacement of giant ophiolite nappes: I. Mesozoic-Cenozoic examples. Tectonophysics 37,247-303. BROWN, G.F. 1971. Tectonic map of the Arabian Peninsula. US Geol. Surv., Saudi Arabian Project Report 134. BROWN, G.F. & COLEMAN, R.G. 1972. The tectonic framework of the Arabian Peninsula. 24th Init. Geol. Cong. Montreal Proc. 3. BROWN, M.A. 1980. Textural and geochemical evidence for the origin of some chromite deposits in the Oman ophiolite. In: PANAYIOTOU A. (ed.) Ophiolites, Proceedings of the International Ophiolite Symposium, Cyprus 1979, 714-721. BROWN, M.A. 1982. Chromite deposits and their ultramafic host rocks in the Oman ophiolite. Unpubl. PhD Thesis, Open University, 263pp. BROWNING, P. 1982. The petrology, geochemistry, and structure of the plutonic rocks of the Oman ophiolite. Unpubl. PhD Thesis, Open University, 404pp. BROWNING, P. 1984. Cryptic variation within the Cumulate Sequence of the Oman ophiolite: magma chamber depth and petrological implications. GASS, I.G., LIPPARD, S.J. & SHELTON, A.W. ( eds. )Ophiolites and Oceanic Lithosphere. Spec. Publ. Geol.Soc. London 14, Blackwell Scientific Publications, Oxford, 71-82. BROWNING, P. & LIPPARD, S.J. (eds) 1982. Open University Oman Project Map Sheet 45 - The Wadi Hawasina - Rustaq Area. Directorate of Overseas Surveys. BROWNING, P. & SMEWlNG, J.D. 1981. Processes in magma chambers beneath spreading axes: evidence for magmatic associations in the Oman ophiolite. J. geol. Soc. Lond. 138, 27%280. BRUNN, J.H. 1960. Mise en place et differenciation de l'association pluto-volcanique due cortege ophiolitique. Rev. Geogr. Phys. et geol. dyn. 3, 115-132. BUNGENSTOCK, H. 1966. Seismiche Unterschungen im nordlichen Teil des arabischen Meers (Golf von Oman). Erd. Kohl. Erdgas. Petrochem. 19,237-243. BURRUSS, R.C., CERCONE, K.R. & HARRIS, P.M. 1983. Fluid inclusion petrography and tectonic-burial history of the A1 Ali No 2 well: Evidence for the timing of diagenesis and oil migration, northern Oman foredeep. Geology 11,567-570. CAMP, V.E. & GRIFFIS, V.E. 1982. Character, genesis and tectonic setting of igneous rock in the Sisten suture zone, E. Iran. Lithos, 15,221-239. CAMPBELL, I.H. 1978. Some problems with the cumulus theory. Lithos 11,311-323. CANN, J.R. 1969. Spillites from the Carlsberg Ridge, Indian Ocean. J. Petrol. 10, 1-19. CANN, J.R. 1970. New model for the structure of the oceanic crust. Nature 226,928-930. CANN, J.R. 1974. A model for oceanic crustal structure developed. Geophys. J. R. astr. Soe. 39, 169-187. CANN, J.R. 1979. Metamorphism in the ocean crust. In: TALWANI,M., HARRISON, C.G. & HAYES, D.E. (eds). Deep drilling results in the Atlantic Ocean Crust, Maurice Ewing Series 2, A.G.U., Washington, Geodynamics Project, Scientific Report 48,230-238. CAPREDI, S., VENTURELLI, G. & TOSCANI, L. 1982. Petrology of an ophiolitic cumulate sequence from Pindos, Greece. Geol. J. 17, 223-242. CAPREDI, S., VENTURELLI, G., BOCCHI, G., DOSTAL, J. & RossI, A. 1980. The geochemistry and retrogression of an ophiolitic suite from Pindos, Greece. Contrib. Min. Pet. 74, 18%200 CARMICHAEL, I.S.E., TURNER, F.J. & VERHOOGAN, J. 1974. Igneous Petrology McGraw-Hill, New York, 739pp. CARNEY, J.N. & WELLAND, M.J.P. 1974. Geology and Mineral Resources of the Oman Mountains. Great Britain Institute of Geological Sciences, Overseas Division, Report 27, 49pp. CARTER, A.J. 1850. Geological observations on the igneous rocks of Muscat and its neighbourhood, and a limestone formation at their circumference. J. Bombay Branch Roy. Asiat. Soc. 3, 118-129. CARTER, N.L. & AVE LALLEMANT, H.G. 1970. High temperature flow of dunite and peridotite Bull. geol. Soc. Am. 81, 2181-2202. CASEY, J.F. & DEWEY, J.F. 1984. Initiation of subduction zones along

References transform and accreting plate boundaries, triple junction evolution, and forearc spreading centres - implications for ophiolitic geology and obduction. In: GASS, I.G., LIPPARD, S.J. 8` SHELTON, A.W. (eds) Ophiolites and Oceanic Lithosphere Spec. Publ. Geol. Soc. London 14, Blackwell Scientific Publications, Oxford 269-290. CASEY, J.F. 8` KARSON, J.A. 1981. Magma chamber profiles from the Bay of Islands ophiolite complex. Nature 292, 295-301. CHEN, J.H. 8` PALLISTER,J.S. 1981. Lead isotope studies of the Samail ophiolite, Oman. J. Geophys. Res. 86, 2699-2708. CHRISTENSEN, N.I. 8` SALISBURY,M.H. 1975. Structure and constitution of the lower oceanic crust. Rev. Geophys. Space Physics 13, 57-86. CHRISTENSEN, N.I. 8` SMEWlNG, J.D. 1981. Geology and seismic structure of the northern section of the Oman Ophiolite. J. Geophys. Res. 86, 2545-2555. CHRISTIANSEN, F.G. 1985. Deformation fabric and microstructures in ophiolitic chromitites and host ultramafics, Sultanate of Oman. Geolog. Rundsch. 74 (1), 61-76. CHRISTIANSEN, F.G. 8` ROBERTS, S. (1986). Formation of olivine pseudo-crescumulates by syntectonic axial planar growth during mantle deformation. Geol. Mag. 123, 73-79, CHURCH, W.R. 8` RICCIO, L. 1977. Fractionation trends in the Bay of Islands ophiolite of Newfoundland; polycyclic cumulate sequences in ophiolites and their classification. Can. J. Earth Sci, 14, 1156--1165. CHURCH, W.R. 8` STEVENS, R.K. 1971. Early Palaeozoic ophiolite complexes of the Newfoundland Appalachians as mantle-ocean crust sequences. J. Geophys. Res. 76, 1466. ~/' COLEMAN, R.G. 1971. Plate tectonic emplacement Of upper mantle peridotites along continental edges. J. Geophys. Res. 76, 1212-1222. COLEMAN, R.G. 1977. Ophiolites." Ancient oceanic lithosphere? Springer-Verlag Berlin, 229pp. COLEMAN, R.G. 1981. Tectonic setting of ophiolite obduction in Oman. J, Geophys. Res. 86, 2497-2508. COLEMAN, R.G. 8` KHTH, T.E. 1971. A chemical study of serpentinization- Burro Mountain, California. Petrol. 12, 311-328. COLEMAN, R.G. 8, PETERMAN, Z.E. 1975. Oceanic plagiogranite. J. Geophys. Res. 80, 1099-1108. COLEMAN, R.G., HUSTON, C.C., EL-BOUSHI, I.M., AL-HINAI, K.M. 8` BAILEY, E.H. 1979. The Semail ophiolite and associated massive sulphide deposits, Sultanate of Oman. Evolution and Mineralization of the Arabian Shield Institute of Applied Geology, King Abdulaziz University, Jeddah, Kingdom of Saudi Arabia, Bull. 3, Pergamon Press, 179-192. CONSTANTINOU, G. 8` GovErr, G.J.S. 1972. Genesis of sulphide deposits, ochre and umber of Cyprus. Inst. Mining Metallurgy Trans. 81, B34-B46. CONSTANTINOU, G. & GOVETT, G.J.S. 1973. Geology, geochemistry and genesis of Cyprus sulfide deposits. Econ. Geol. 68,843-858. CORLISS, J.B., DYMOND, J., GORDON, L.I., EDMOND, J.M., WON HERZEN, R.P., BALLARD, R.D,, GREEN, K., WILLIAMS, D., BAINBRIDGE, A., CRANE, K. 8` VAN ANDEL, T.H. 1979. Submarine thermal springs on the Galapagos rift. Science 203, 1073-1083. CRAWFORD, A.R. 1972. Iran, continental drift and plate tectonics. 24th Int. Geol. Congress, Montreal Section 3,106-112. DAHL, R., JUTEAU, T., BOUDIER, F., NICOLAS, A., BOUCHEZ, J-L. 8` CRAMBERT, S. 1983. Ophiolites des nappes de Semail (Oman): Nouvelles donnees de terrain sur les parties plutoniques superieures des massifs de Rustaq et de Nakhl. Sci. Geol. Bull. 36 (1), 35-59. DAHLSTROM, C.D.A. 1970. Structural geology in the eastern margin of the Canadian Rocky Mountains. Bull. Can. Petrol. Geol. 18, 332-406. DAVIES, H.L. 1971. Peridotite-gabbro-basalt complex in eastern Papua: an overthrust plate of oceanic mantle and crust. Bur. Miner. Resour. Aust. Bull. 128, 48pp. DAVIES, H.L, 1980. Folded thrust fault and associated metamorphics in the Suckling-Dayman massif, Papua New Guinea. Am. J. Sci. 280A, 171-191. DAVIES, H.L. 8, JACQUES, A.L. 1984. Emplacement of ophiolite in Papua New Guinea. In: GASS, I.G., LIPPARD, S.J. 8` SHELTON,

I69

A.W. (eds.) Ophiolites and Oceanic Lithosphere. Spec. Publ. Geol. Soc. London 14, Blackweil Scientific Publications, Oxford, 341-349. DAVIS, E.E. & LISTER, C.R.B. 1977. Heat flow measured over the Juan de Fuca Ridge: evidence for widespread hydrothermal circulation in a highly heat transportive crust. J. Geophys. Res. 82, 4845-4860. DAVOUDZEDEH, M. 1969. Geologie und Petrographie des Gebietes nordlich van Nain, Zentral Iran. Dissertation ETH Zurich, 90pp. DAVOUDZEDEH. M. 1972. Geology and petrography of the area north of Nain, Central Iran. Geol. Surv. lran. Rept. 14, 89pp. DEANS, A.G. 1950. The Radiolaria of the Hawasina Series of Oman. Proc. Geol. Soc. 61,206-217. DELALOYE, M. & DESMONS, J. 1980. Ophiolites and melange terrains in Iran: a geochronological study and its palaeotectonic implications. Tectonophysics. 68, 83-111. DESMET, A., LAPIERRE, H., ROCCl, G., GAGNY, C.L., PARROT, J-F. & DELALOYE, M. 1978. Constitution and significance of the Troodos sheeted complex. Nature 273, 527-530. DESMET, A., GA~NY, C.L., LAPIERRE, H. & Roccl, G. 1980. Organisation spatio-temporelle du complexe filonien du Troodos: Son enracinement dans la chambre magmatique. In: PANAYIOTOUA. (ed.) Ophiolites, Proceedings of the International Ophiolite Symposium, Cyprus 1979, 66-72. DESMONS, J. 1977a. Ophiolites and melange in the Makustan-Fanuj area southeast of Iranshar (eastern Baluchistan, Iran). Unpubi. Rept. Geol. Surv. Iran. DESMONS, J. 1977b. Ophiolites and melange in the Tchehel Kireh sheet (west of Zahedean SE Iran). Unpubl. Rept. Geol. Surv. Iran. DESMONS, J. & BECCALUVA,L. 1983. Mid-ocean ridge and island arc affinities in ophiolites from Iran: palaeogeographic implications. Chem. Geol. 39, 39-63. DESMONS, J., DELALOYE,M., DESONET, A., GAGNY, eL., ROCCI, G. 8` VOLDET, P. 1980. Trace and rare earth element abundances in Troodos lavas and sheeted dykes, Cyprus. Ofioliti 5, 35-56. DEWEY, J.F. 1974. Continental margins and ophiolite obduction: Appalachian - Caledonian system. In: BURKE, C.A. & DRAKE, C.L. (eds) The geology of continental margins Springer-Verlag, New York, 933-950. DEWEY, J.F. 1975. Finite path evolution: implications for the evolution of the rock masses at plate margins. Am. J. Sci. 275-A, 260-284. DEWEY, J,F. 1976. Ophiolite obduction. Tectonophysics 31, 93-120. DEWEY, J.F. 8` BIRD, J.M. 1971. Origin and emplacement of the ophiolite suite: Appalachian ophiolites in Newfoundland. J. Geophys. Res. 76, 3179-3206. DEWEY, J.F. 8` KIDD, W.S.F. 1977. Geometry of plate accretion. Bull. geol. Soc. Am. 88, 960-968. DEWEY, J.F. 8` SHACKLETON,R.M. 1984. A model for the evolution of the Grampian tract in the early Caledonides and Appalachians. Nature 312. 115-121. DEWEY, J.F., PITMAN, W.C., RYAN, W.B.F. 8` BONIN, J. 1973. Plate tectonics and the evolution of the Alpine system. Bull. geol. Soc. Am. 84. 3137-3180. DICK, H.J.B. 1982. The petrology of two back-arc basins of the northern Philippine Sea. Am. J. Sci. 282, 644-700. DICK, H.J.B. 8` SINTON, J.M. 1979. Compositional layering in alpine peridotites: evidence for pressure solution creep in the mantle. J. Geol. 87, 403-416. DixoN, S. 8` RUTHERFORD, M.J. 1979. Plagiogranites as late stage immiscible liquids in ophiolite and mid-ocean ridge suites - an experimental study. Earth Plan. Sci. Lett. 45, 45-60. DUBERTRET, L. 1953. Geologie des roches du NW de la Syrie et du Hatay (Turquie). Notes et Mem. sur la Moyen-Orient 6, (Mus. National d'Histoire Naturelle). DUNCAN, R.A. 8` GREEN, D.H. 1980. Role of multistage melting in the formation of oceanic crust. Geology 8,222-226. DUNLOP, H.M. 8` FOUILLAC,C. 1984. O, H, Sr, Nd isotope systematics of the Oman Ophiolite. Ophiolites "Through Time". Conf. Nantes abst. p. 25. DONNELLY, T.W., MELSON, W.K.R. 8` ROGERS, J.J.W. 1973. Basalts and dolerites of Late Cretaceous age from the central Caribbean.

I70

References

In: EDGAR, N.T., SAUNDERS,J.B. et al. (eds) Initial Reports of the Deep Sea Drilling Project, 15, US Govt. Printing Office, Washington, 989-1011. DURHAM, W.B., GOETZE, C. & BLAKE, B. 1977. Plastic flow of oriented single crystals of olivine. II. Observations and interpretations of the dislocation structures. J. Geophys. Res. 82, 5755-5770. ELDER, J.W. 1977. Model of hydrothermal ore genesis. Spec. Publ. Geol. Soc. Lond. 7, 4-13. ELLIOTT, D. 1976. The motion of thrust sheets. J. Geophys. Res. 81 (5), 949-963. ELLIOTr, D. & JOHNSON, M.R.W. 1980. Structural evolution in the northern part of the Moine thrust belt, NW Scotland. Trans. Roy. Soc. Edinburgh : Earth Sci. 71, 69-96. ELLIS, D.J. 8, GREEN, D.H. 1979. An experimental study of the effect of Ca upon garnet-clinopyroxene Fe-Mg equilibrium. Contrib. Min. Pet. 71, 13-22. ELTHON, D. & STERN, C. 1978. Metamorphic petrology of the Samiento ophiolite complex, Chile. Geology 6,464--468. ELTHON, D., CASEY, J.F. & KOMOR, S. 1982. Mineral chemistry of ultramafic cumulates from the North Arm Mountain massif of the Bay of Islands Ophiolite: evidence for high-pressure crystal fractionation of oceanic basalts. J. Geophys. Res. 87, 8717-8734. FABRIES, J. 1979. Spinel-olivine geothermometry in peridotites from ultramafic complexes. Contrib. Min. Pet. 69,329-336. FALCON, N.L. 1967. The geology of the northeast margin of the Arabian Basement Shield. Adv. Sci. 24, 31-42. FALCON, N.L. 1974. Southern Iran: Zagros Mts. In: SPENCER, A.M. (ed.) Mesozoic-Cenozoic Orogenic Belts, data for orogenic studies. Spec. Publ. Geol. Soc. London 4, 199-211. FALVEY, D.A. 1974. The development of continental margins in plate tectonic theory. J. Aust. Pet. Explor. Assoc. 14, 95-106. FARHOUDI, G. & KARIG, D.E. 1977. Makran of lran and Pakistan as an active arc system. Geology 5,664-668. FERRARA, G., INNOCENTI, F., RICCI, C.A. & SERRI, G. 1976. Oceanfloor affinity of basalts from North Apennine ophiolites: geochemical evidence. Chem. Geol. 17, 101-111. FIRSTBROOK, P.L., FUNNELL, B.M., HURLEY, A.M. & SMITH, A.G. 1979. Paleoceanic reconstructions, 160-0 Ma. DSDP Publication. FLEET, M.E. & BARNE1T, R.L. 1978. AllVAlV~ partitioning in calciferous amphiboles from the Frood Mine, Sudbury, Ontario. Contrib. Min. Pet. 16, 52%532. FLEET, A.J. & ROBERTSON, A.H.F. 1980. Ocean ridge metalliferous and pelagic sediments of the Semail nappe, Oman. J. geol. Soc. Lond. 137, 403-422. FLOYD, P.A. & WINCHESTER, J.A. 1975. Magma type and tectonic setting discrimination using immobile elements. Earth Plan. Sci. Lett. 27, 211-218. FORD, C.E. 1976. Phase relations in the system CaO - MgO - AI203 SiO2 - H20. Pros. in Exp. Pet. 3, 266--272. Fox, P.J. & STROUP, J.B. 1981. The plutonic foundation of the oceanic crust. In: EMILIANI, C. (ed) The Oceanic Lithosphere, The Sea VIII, 119-218. FRANCHETAU, J., NEEDHAM, H.D., CHOUKROUNE, P., JUTEAU, T., SEGURET, M., BALLARD, R.D., FOX, P.J., NORMARK, W., CARRANZA, A., CORDOBA,D., GUERRERO,J., RANGIN, C., BOUGAULT, H., CAMBON, P. & HEKINIAN, R. 1979. Massive deep-sea sulphide ore deposits discovered on the East Pacific Rise. Nature 277, 523-528. FREY, F.A. 1970. Abundances in alpine metamorphic rocks. Phys. Earth Planet. Interiors 3, 323-330. FREY, F.A., BRYAN, W.B.A. ,s, THOMPSON, G. 1974. Atlantic ocean floor, geochemistry and petrology of basalts from Legs 2 and 3 of the Deep Sea Drilling Project. J. Geophys. Res. 79, 5507-5027. FREY, F.A., GREEN, D.H. & RoY, S.D. 1978. Integrated models of basalt petrogenesis. J. Petrol. 19, 463-513. GASS, I.G. 1968. Is the Troodos Massif of Cyprus a fragment of Mesozoic ocean floor? Nature 220, 39-42. GASS, I.G. 1970. The evolution of volcanism in the junction area of the Red Sea, Gulf of Aden and Ethiopian Rifts. Phil. Trans. Roy. Soc. Lond. A267, 369-381. GASS, I.G. 1977. The evolution of the Pan African crystalline basement in NE Africa and Arabia. J. geol. Soc. Lond. 134, 129-138. GASS, I.G. 1982. Ophiolites. Sci. Amer. 247, 2, 122-131.

GASS. I.G. & SMEWING, J.D. 1973. Intrusion, extrusion and metamorphism at constructive margins: Evidence from the Troodos Massif, Cyprus. Nature 242, 26-29. GASS, I.G. & SMEWING, J.D. 1981. Ophiolites: obducted oceanic lithosphere. In: EMILIANI, C. (ed.). The Sea Vol VII; Oceanic lithosphere. Wiley lnterscience, New York, 339-362. GASS, I.G., MALUCK, D.I.J. & Cox, K.G. 1973. Volcanic islands of the Red Sea. J. geol. Soc. Lond. 129,275-310. GASS, I.G., SMITH, A.G. & VINE, F.J. 1975. Origin and emplacement of ophiolites. In: Geodynamics today, a review of the Earth's dynamic processes. Roy. Soc. London, 55-64. GAST, P.W. 1968. Trace element fractionation and the origin of tholeiitic and alkaline magma types. Geochim. Cosmochim. Acta 32, 1057-1086. GALE, N.H., SPOONER, E.T.C. & Po~s, P.J. 1981. The lead and strontium isotope geochemistry of metalliferous sediments associated with Upper Cretaceous ophiolitic rocks in Cyprus, Syria and Oman. Can. J. Earth Sci. 18, 1290-1302. GANSSER, A. 1955. New aspects of the geology of Central Iran. Proc. 4th. World Petrol. Cong. Rome, section l/A/5,279-300. GANSSER, A. 1976. The Ophiolite Melange: a World-wide problem on Tethyan examples. Eclog. geol. Helv. 67,479-507. GEALEY, W.K. 1977. Ophiolite obduction and the geologic evolution of the Oman Mountains and adjacent areas. Bull. geol. Soc. Am. 88, 1183-1191. GEALEY, W.K. 1980. Ophiolite obduction mechanism. In: PANAYIOTOO A. (ed.) Ophiolites, Proceedings of the International Phiolite Symposium, Cyprus 1979, 228-243. GHENT, E.D. & STOUT, M.Z. 1981. Metamorphism at the base of the Samail ophiolite, southeastern Oman Mountains. J. Geophys. Res. 86, 2557-2571. GHOSE, N.C. & SINGH, R.N. 1981. Ophiolite belt of Nasa Hills and its relationship with Indus Suture and Andaman-Nicobar Islands. In: Three decades of development in Petrology, Mineralogy and Petrochemistrv in India, Geol. Surv. India, Jaipur. GIRARDEAU, J., MARCOUX, J., ALLEGRE, C.J., BASSOULLET, J.D., TANG YOUKING, XIAO XUCHANG, ZAO YOUGONG & WANG XIBIN 1984. Tectonic environment and geodynamic significance of the NeD-Cimmerian Donqiao ophiolite, Bangong Nujiang suture zone, Tibet. Nature 307, 27-31. GIRARDEAU, J., NICOLAS, A., MARCOUX, J., DUPRE, B., CHANG CHENG-FA, WANG XIBIN & ZHAO YO-GANG. 1981. The Xigaze ophiolite (Tibet): a peculiar oceanic crust. Terra Cognita Spec. Issue, Ophiolites and Greenstone Belts, 17-18. GLENNIE, K.W. 1977. Outline of the geology of Oman. Mem. h.-ser. Soc. Geol. Fr. 8, 25-31. GLENNIE. K.W. & HUGHES CLARK, M.W. 1974. Late Cretaceous nappes in the Oman Mountains: a reply to J.D. Moody. B A A P G 58, 895-898. GLENNIE, K.W., BOEUF, M.G.A., HUGHES CLARK, M.W., MOODYSTUART, M., PILAAR, W.F.H. & REINHARDT, B.M. 1973. Late Cretaceous nappes in the Oman Mountains and their geologic evolution. Bull. Amer. Assoc. Petrol. Geol. 57, 5-27. GLENNIE, K.W., BEOUF, M.G.A., HUGHES CLARK, M.W., MOODYSTUART, M., PILAAR, W.F.H. & REINHARDT, B.M. 1974. The geology of the Oman Mountains, Konin. Neder. Geol. Mijnbouw. Genoot. Verdh. 31, Part 1 (text), 423pp., Part 2 (illustrations), Part 3 (enclosures). GORIN, G.E., RACZ, L.G. & WALTER, M.G. 1982. Late PreCambrianCambrian sediments of Huqf Group, Sultanate of Oman. Bull. Am. Assoc. Petrol. Geol. 66 (12), 2609-2627. GOUGH, D.I., BROCK, A., JONES, D.L. & OPDYKE, N.D. 1964. The paleomagnetism of the ring complexes of Marangudgi and the Mateke Hills. J. Geophys. Res. 69, 2499-2507. GRAHAM, G.M. 1980a. Evolution of a passive margin and nappe emplacement in the Oman mountains. In: PANAYIOTOUA. (ed.) Ophiolites, Proceedings of the International Ophiolite Symposium, Cyprus 1979, 414-423. GRAHAM, G.M. 1980b. Structure and sedimentology of the Hawasina Window, Oman mountains. Unpubl. PhD Thesis, Open University, 458pp. GRAHAM, C.M. & ENGLAND, P.C. 1976. Thermal regimes and regional metamorphism in the vicinity of overthrust faults : an example of shear heating and inverted metamorphic zonation for southern

Referen ces California, Earth Plan, Sci. Lett. 31, 142-152. GREEN, D.H. 1970. A review of experimental evidence of the origin of basaltic and nephelinitic magmas. Phys. Earth. Planet. Interiors 3, 221-235. GREEN, T.H. 1973. Experimental melting studies on a model uppermantle composition at high pressure under water-saturated conditions. Earth Plan. Sci. Lett. 19, 37-53. GREEN, D.H., HIBBERSON, W.O. & JAQUES, A.L. 1979. Petrogenesis of mid-ocean ridge basalts. In: MCELHINNEY, M.W. (ed). The Earth, Its origin, structure and evolution. Academic Press, London, 265-297. GREENBAUM, D. 1972. Magmatic process at ocean ridges : evidence from the Troodos Massif, Cyprus. Nature, Phys. Sci. 238, 18-21. GREENWOOD, J.E.G.W. & LONEY, P.E. 1968. Geology and mineral resources of the Trucial Oman Range. Inst. Geol. Sciences, Overseas Division Unpubl. Report 108pp. GREENWOOD, W.R., HADLEY, D.G., ANDERSON, R.E., FLECK, R.J. & SCHMIDT, D.L. 1976. Late Proterozoic cratonization of southwestern Saudi Arabia. Phil. Trans. Roy. Soc. Lond. A280, 517-527. GREGORY, R.T. & TAYLOR, H.P. 1981. An oxygen isotope profile in a section of Cretaceous oceanic crust, Samail ophiolite: evidence for 6180 buffering of the oceans by deep (>5 kin) seawaterhydrothermal circulation at mid-ocean ridges. J. Geophys. Res. 86, 2737-2755. GROW, J.A., BowiY, C.O. & HUTCHINSON, D.R. 1979. The gravity field of the US Atlantic continental margin. In: KEEN, C.E. (ed.) Crustal properties across passive margins, Tectonophysics, 59, 27-52. HAJASH, A. JR. 1984. Rare-earth element abundances and distribution pattern in hydrothermally altered basalts: experimental results. Contrib. Min. Pet. 83,409-412. HALL, R. 1980. Contact metamorphism by an ophiolitic peridotite from Neyriz Iran. Science, 208, 1259-1262. HALL, R. 1981. Ophiolite-related contact metamorphism: skarns from Neyriz, Iran. Proc. Geol. Ass. 92, 231-240. HALL, R. 1984. Ophiolites: Figments of oceanic lithosphere.'? ln: GASS, I.G., LIPPARD, S.J. & SHELTON, A.W. (eds.) Ophiolites and Oceanic Lithosphere. Spec. Publ. Geol. Soc. London 14, Blackwell Scientific Publications, Oxford, 393-403. HALLAM, A. 1976. Geology and plate tectonic interpretation of sediments in the Mesozoic radiolarite-ophiolite complex in the Neyriz region, southern Iran. Bull. Geol. Soc. Am., 87, 47-52. HAMELIN, B., DUPRE, B. & ALLEGRE, C.J. 1984. The lead isotope systematics of ophiolite complexes. Earth Plan. Sci. Lett. 67, 351-366. HANSON, G.N. 1978. The application of trace elements to the petrogenesis of igneous rocks of granitic compositions. Earth Plan. Sci. Lett. 38, 26-43. HANSON, G.N. 1980. Rare earth elements in petrogenetic studies of igneous systems. Ann. Rev. Earth Planet. Sci. 8,371-406. HAREMBOURE, J. & HORSTINK, J. 1967. Mesozoic nappes in the Oman mountains, a hypothesis. Unpubl. Report P.D. (Oman). HARLAND, W.B., Cox, A.V., LLEWELLYN, P.G., PICKTON, C.A.G., SMITH, A.G. & WALTERS, R. 1982. A geologic time scale. Cambridge University Press, Cambridge, 131pp. HARRIS, P.G. 1957. Zone refining and the origin of potassic basalts. Geochim. Cosmochirn. Acta 32, 12, 195-208. HART, E.W. 1970. A phenomenological theory for plastic deformation of polycrystalline metals. Acta. Metall. 18, 599-610. HART, R.A. 1973. A model for chemical exchange in the basalt seawater system of oceanic layer II. Can, J. Earth Sci. 10, 799-816. HARTE, B. 1976. Rock nomenclature with particular relation to deformation and recrystallization textures in olivine-bearing xenoliths. J. Geol. 85, 279-288. HAWKINS, T.R.W., HINDLE, D. & STRUGNELL, R. 1981. Outline of the stratigraphy and structural framework of southern Dhofar (Sultanate of Oman). Geol. Mijn. 60, 247-256. HAWKESWORTH, C.J., NORRY, M.J., RODDICK, J.C., BAKER, P.E., FRANCIS, P.W. & THORPE, R.S. 1979. 143Nd/144Nd, 87Sr/86Sr, and incompatible element variations in calc-alkaline andesites and plateau lavas from South America. Earth Plan. Sci. Lett. 42, 45-57. HAYMON, R.M., KOSKI, R.A. & SINCLAIR. C. 1984. Fossils of hydro-

171

thermal vent worms from Cretaceous sulfide ores of the Samail Ophiolite, Oman. Science 223, 1407-1409. HEATON, T.H.E. & SHEPPARD, S.M.F. 1979. Hydrogen and oxygen isotope evidence for sea-water hydrothermal alteration and ore deposition, Troodos Complex, Cyprus. Spec. Publ. Geol. Soc. Lond. 7, 42-57. HELLMAN, P.L., SMITH, R.E. & HENDERSON, P. 1979. The mobility of the rare earth elements: evidence and implications from selected terrains affected by burial metamorphism Contrib. Min. Pet. 71, 23-44. HERRON, J.J., STOFFA, P.L. & BUHL, P. 1980. Magma chambers and mantle reflections- East Pacific Rise. Geophys. Res. Lett. 7, 989-992. HEY, M.H, 1954. A new series of chlorites. Min. Mag. 30, 277-292. HOPSON, C.A. & PALLISTER, J.S. 1980. Samail ophiolite magma chamber: I, evidence from gabbro phase variation, internal structure and layering. In: PANAYIOTOUA. (ed.) Ophiolites, Proceedings of the International Ophiolite Symposium, Cyprus 1979, 402-4O4. HoPSON, C.A., COLEMAN, R.G., GREGORY, R.T., PALLISTER, J.S. & BAILEY, E.H. 1981. Geologic section through the Samail ophiolite and associated rocks along a Muscat-Ibra transect, southeastern Oman mountains. J. Geophys. Res. 86, 2527-2544. Hsu, K.J. & BERNOUILLI,D. 1978. Genesis of the Tethys and Mediterranean. In: Hsu, K.J. & MONTADART, L. et al. Initial Reports of the Deep Sea Drilling Project 42. Washington U.S. Govt. Printing Office, 943-949. HUDSON, R.G.S. 1960. The Permian and Trias of the Oman Peninsula, Arabia. Geol. Mag. 97,299-308. HUDSON, R.G.S., BROWNE, R.V. & CHARON, C. 1954a. The structure and stratigraphy of the Jebel Qamar area, Oman. Proc. geol. Soc. Lond. 1513, 49-54. HUDSON, R.G.S., MCGRUGEN, A. & MORTON, D.M. 1954b. The structure of the Jebei Hagab area, Trucial Oman. Quart. J. geol. Soc. Lond. 110, 121-152. HUDSON, R.G.S. & CHATrON, C. 1959. The Musandam limestone (Jurassic to Lower Cretaceous) of Oman, Arabia. Paris Mus. Nat. d'Histoire Naturelle, notes et memoire Moyen-Orient 7, 69-93. HUMPHRIS, S.E. & THOMPSON, G. 1978. Hydrothermal alteration of oceanic basalts by seawater. Geochim. Cosmochim. Acta 42, 107-125. HUPPERT, H.E. a SPARKS, R.S.J. 1980a. Restrictions on the compositions of mid-ocean ridge basalts: a fluid dynamical investigation. Nature 286, 46-48. HUPPERT, M.E. & SPARKS, R.S.J. 1980b. The fluid dynamics of a basaltic magma chamber replenished by influx of hot, dense, ultrabasic magma. Contrib. Min. Pet. 75,279-289. HUTCHISON, I., LOUDEN, K.E., WHITE, R.S. & VON HERZEN, R.P. 1981. Heat flow and the age of the Gulf of Oman. Earth Plan. Sci. Lett. 56,252-262. HYNES, A. 1974. Igneous activity at the birth of an ocean basin in eastern Greece. Can. J. Earth. Sci. 11,842-853. IRVINE, T.N. 1978. Density current structure and magmatic sedimentation. Carnegie Institution of Washington Year Book 77, 717-725. IRV1NE, T.N. 1980. Magmatic infiltration metasomatism, double-diffusive fractional crystallisation, and adcumulus growth in the Muskox intrusion and other layered intrusions. In: HARGREAVES, R.B. (ed), Physics of magmatic processes, Princetown University Press, 325-384. IRVINE, T.N. & FINDLAY, T.C. 1972. Alpine-type peridotite with particular reference to the Bay of Islands igneous complex. Pub. Earth Phys. Branch Dep. Energy, Mines & Res., Canada 42, 97-128. IRVING, A.J. 1978. A review of experimental studies of crystal/liquid trace element partitioning. Geochim. Cosmochim. Acta 42, 743-770. IXER, R.A., ALABASTER,T. & PEARCE,J.A. 1984. Ore petrography and geochemistry of massive sulphide deposits within the Semail ophiolite Oman. Trans. lnstn. Min. Metall. 93, 114-124. JACKSON, E.D., GREEN, H.W. & MOORES, E.M. 1975. The Vourinos ophiolite, Greece: cyclic units of lineated cumulates overlying harzburgite tectonite. Bull. geol. Soc. Am. 86,390-398.

I72

References

JACOB, K.H. & QuIrrMEYER, R.C. 1979. The Makran region of Pakistan and Iran, trench-arc system with active plate subduction. In: FARAN, A. & DEJONG, K.A. (eds) Geodynamics of Pakistan, Geol. surF. Pakistan Quetta, 305-318. JAEGER, J.C. 1942. Moving sources of heat and the temperatures at sliding contacts. J. proc. R. Soc. N.S. W. 76, 203-224. JAEGER, J.C. 1961. The cooling of irregularly shaped igneous bodies. Am. J. Sci. 259, 721-734. JAMIESON, R.A. 1980. Formation of metamorphic aureoles beneath ophiolites- evidence from the St. Anthony Complex, Newfoundland. Geology 8, 150-154. JAQUES, A.L. 1981. Petrology and petrogenesis of cumulate peridotites and gabbros from the Marum ophiolite complex, northern Papua New Guinea. J. Petrol. 22, 1-40. JAQUES, A.L., ROBINSON,G.P. & TAYLOR, S.R. 1983. Geochemistry of cumulus peridotites and gabbros from the Marum ophiolite complex, northern Papua New Guinea. Contrib. Min. Pet. 82, 154-164. JENKYNS, H.C. 1979. Pelagic Environments. In: READING, H.G. (ed) Sedimentary environments and facies. Elsevier, New York, 314-371. JOHANNES, W. 1968. Experimental investigation of the reaction forsterite + H~O --~ serpentine + brucite. Contrib. Min. Pet. 19, 30%315. JOHANNES, W. 1978. Melting of plagioclases in the system Ab-AnH20 and Qz-Ab-An-H20 at Pn2o = 5 kbars, an equilibrium problem. Contrib. Min. Pet. 66, 295-303. JUTEAU, T. 1970. Petrogeneses des ophiolites des Nappes d'Antalya (Taurus Lycien oriental, Turquie): Leur liaison avec une phase d'expansion oceanique active au Trias superieur. Sci. Terre 15, part 3,265-288. JUTEAU, T. 1975. Les ophiolites des nappes d'Antalya (Taurides occidentales, Turquie). Petrologie d'un fragment de l'ancienne coute oceanique tethysienne. Sci. Terre Nancy Mere. 32,692pp. JUTEAU, T. 1980. Ophiolites of Turkey. In: Roccl, G. (ed) Ofioliti Special Issue Tethyan Ophiolites Vol. 2 Eastern Area, 19%237. JUTEAU, T. & WHITECHURCH, H. 1980. The magmatic cumulates of Antalya (Turkey): Evidence of multiple intrusions in an ophiolitic magma chamber. In: PANAYIOTOU A. (ed.) Ophiolites, Proceedings of the International Ophiolite Symposium, Cyprus 1979, 377-391. JUTEAU, T., LAPIERRE, H., NICOLAS, A., PARROT, J-F., Ricou, L.E., RoccI, G. & ROLLET, M. 1973. Idees actuelles sur la constitution, l'origine et l'evolution des assemblages ophiolitiques mesogeens. Bull. Soc. Geol. Fr. 15,478--493. JUTEAU, T., NICOLAS, A., DUBESSY, J. FRUCHARD, J.C. & BOUCHEZ, J.L. 1977. The Antalya ophiolite complex (western Taurides, Turkey): a structural model for an oceanic ridge. Bull. geol. Soc. Am. 88, 1740-1748. KARAMATA, S. 1980. Metamorphism beneath obducted ophiolite slabs. In: PANAYIOTOU A. (ed.) Ophiolites, Proceedings of the International Ophiolite Symposium, Cyprus 1979, 219-227. KARAMATA,S., MAYER, V. & PAMIC, J. 1980. Ophiolites of Yugoslavia, In: Roccl, G. (ed) Ofioliti Special Issue Tethyan Ophiolites Vol. 1 Western area, 105-125. KAY, R.W. & GAST, P.W. 1973. The rare-earth content and origin of alkali-rich basalts. J. Geol. 81,653-682. KIDD, R.G.W. 1977. A model for the process of formation of the upper oceanic crust. Geophys. J. Roy. Astron. Soc. 50, 149-183. KIDD, R.G.W. & CANN, J.R. 1974. Chilling statistics indicate an ocean floor spreading origin for the Troodos Complex, Cyprus. Earth Plan. Sci. Lett. 24, 151-155. KING, B.C. 1970. Vulcanicity and rift tectonics in East Africa. In: CLIFFORD, T.N. & GASS, I.G. (eds) African magmatism and Tectonics, Hajner, New York, 263-283. KRUMSlEK, K. 1976. Zur Bewegung der Iransch-Afghanisten Platte. Geol. Rundsch. 65, 90%929. KUHN, O. 1929. Beitrage zur Palaeontologie und Stratigraphie von Oman (Ost Arabian). Ann. Nat. Mus. Wien. 43, 13-18. KNIPPER, A.L. & KHAIN, E.g. 1980. Structural position of ophiolites of the Caucasus. In: Rocci, G. (ed) Ofioliti, Special Issue: Tethyan Ophiolites Vol. 2, Eastern Area, 297-314. LANGSETH, M.G., LE PICHON, X. & EWING, M. 1966. Crustal structures of midocean ridges. J. Geophys. Res. 71,6351-6356.

LANPHERE, M.A. 1981. K-Ar ages of metamorphic rocks at the base of the Samail ophiolite, Oman. J. Geophys. Res. 86, 2777-2782. LANPHERE, M.A., COLEMAN, R.G., KARAMATA, S. & PAMIC, J. 1975. Age of amphibolites associated with alpine peridotites in the Dinaride ophiolite zone, Yugoslavia. Earth Plan. Sci. Lett. 26, 271-276. LANPHERE, M.A.. COLEMAN, R.G. & HOPSON, C.A. 1981. Sr isotopic tracer study of the Samail ophiolite, Oman. J. Geophys. Res. 86, 2709-2720. LAPIERRE. H. & ROCCI, G. 1976. Le volcanisme alcalin du Sud-Ouest de Chypre et le probleme de l'ouverture des regions tethysiennes au Trias. Tectonophysics 30, 299-313. LAUBSCHER, H.P. 1975. Plate boundaries and microplates in Alpine history. Am. J. Sei. 275,865-876. LAURANT, R., TANER, M.F. & BERTRAND, J. 1984. Mise en place et petrologie du granite associe au complexe ophiolitique de Thetford Mines, Quebec. Can. J. Earth Sci. 21, 1114-1125. LEBAS, M.J. 1962. The role of aluminium in igneous clinopyroxene with relation to their parentage. Am. J. Sci. 260, 267-288. LEES, G.M. 1928a. The geology and tectonics of Oman and parts of S.E. Arabia. Quart. J. geol. Soc. Lond. ll0, 121-152. LEES, G.M. 1928b. The physical geography of SE Arabia. Geograph. Jour. 81,441-470. LEMOINE, M. 1973. About gravity sliding in the Western Alps. In: Gravity and Tectonics. DEJONG, K.A. & SCHOLTEN, R. (eds), John Wiley, New York. LENSCH, G. 1980. Major element geochemistry of the ophiolites in northeastern Iran. In: PANAYIOTOUA. (ed.) Ophiolites, Proceedings of the International Ophiolite Symposium, Cyprus 1979, 398-401. LENSCH, G., MIHM, A. & ALAVI-TEHRANI,N. 1977. Petrography and geology of the ophiolite belt north of Sabzevar/Khorassan (Iran). Nenes. Jahrb. Mineral. Abh. 131, 156-178. LENSCH, G., MIHM, A. & ALAVI-TEHRANI, N. 1979. Major element geochemistry of the ophiolites north of Sabzevar (Iran). Nenes. Jahrb. Palaontol. Monatsh, 415-447. LEWIS, A.D. & SMEWING, J.D. 1980. The Montgenevre ophiolite (Hautes Alpes, France): its structure, metamorphism and trace element geochemistry of its volcanic sequence. Chem. Geol. 28, 291-306. Llou, J.G. 1971. P-T stabilities of laumontite, waikirite, lawsonite and related minerals in the system CaAl~Si,Os - SiO~ - H20. J. Petrol. 12. 379-411. LIou, J.G., KUNIYOSHI,S. & ITO, K. 1974. Experimental studies of the phase relations between greenschist and amphibolite facies in a basaltic system. Am. J. Sci. 274. 613-632. LIPPARD, S.J. (ed), 1980. Open University Oman Project Map Sheet 2 - The Wadi Jizzi Area. Directorate of Overseas Surveys. LIPPARD, S.J. 1983. Cretaceous high pressure metamorphism in NE Oman and its relationship to subduction and ophiolite nappe emplacement. J. Geol. Soc. London 140, 97-104. LIPPARD, S.J. 1984. Petrology of alkali wehrlite sills in the Oman Mountains. Min. Mag. 48, 13-20. LIPPARD, S.J. & REX, D.C. 1982. K-Ar ages of alkaline igneous rocks in the northern Oman mountains, NE Arabia, and their relations to rifting, passive margin development and destruction of the Oman Tethys. Geol. Mag. 119,497-503. LIPPARD, S.J. & ROTHERY, D.A. (LOS) 1981. Open University Oman Project Map Sheet 3 - The Wadi Ahin - Yanqui Area. Directorate of Overseas Surveys. LIPPARD, S.J., SMEWING,J.D., ROTHERY, D.A. & BROWNING, P. 1982. The geology of the Dibba Zone, northern Oman mountains; a preliminary study. J. geol. Soc. Lond. 139, 59-66. LIPPARD, S.J., GRAHAM, G.M., SMEWING, J.D. & SEARLE, M.P. 1983. Melanges associated with the Semail ophiolite in the Northern Oman Mountains allochthon, southwest Arabia. In: MCCALL, G.J.H. (ed) Ophiolitic Melanges, Benchmark papers in Geology, 300-308. LISTER, C.R.B. 1972. On the thermal balance of a mid-ocean ridge. Geophys. J. R. astr. Soc. 26, 515-535. LOVELOCK, P.E.R., POTTER,T.L., WALSWORTH-BELL,E.B. & WEIMER, W.M. 1981. Ordovician rocks in the Oman mountains: the Amdeh Formation. Geol. Mijn 60, 487-495. LOWRIE, W., CHANNELL, J.E.T. & ALVAREZ, W. 1980. A review

References of magnetic stratigraphy investigations in Cretaceous pelagic carbonate rocks. J. Geophys. Res. 85(B7), 3597-3605. LUDDEN, T.N. & THOMPSON, G. 1979. An evaluation of the behaviour of the rare earth elements during the weathering of sea floor basalts. Earth Plan. Sci. Lett. 43, 85-92. LUYENDYK, B.P. & DAY, R. 1982. Palaeomagnetism of the Samail ophiolite, Oman. 2, The Wadi Kadir gabbro section. J. Geophys. Res. 87 (B13) 10903-10917. LUYENDYK, B.P., LAWS, B.R., DAY, R. & COLLINSON, T.B. 1982. Palaeomagnetism of the Samail ophiolite, Oman. 1. The sheeted dike complex at Ibra. J. Geophys. Res. 87 (B13), 10883-10902. MCBIRNEY, A.R. & NOYES, R.M. 1979. Crystallisation and layering of the Skaergaard Intrusion. J. Petrology 20, 487-554. MCCALL, G.J.H. & KIDD, R.G.W. 1982. The Makran, southeastern Iran: the anatomy of a convergent plate margin active from Cretaceous to present. In: LEGGETr, J.K. (ed) Trench - Forearc Geology. Spec. Publ. Geol. Soc. London 10, 387-397. MCCLAY, K.R. 1981. What is a thrust? What is a nappe? In: MCCLAY, K.R. & PRICE, N.J. (eds.). Thrust and Nappe Tectonics. Spec. Pubi. Geol. Soc. London 9, 7-9. MCCULLOCH, M.T., GREGORY, R.T., WASSERBURG, G.J. & TAYLOR, H.P. 1980. A neodymium, strontium and oxygen isotope study of the Cretaceous Samail ophiolite and implications for the petrogenesis and seawater-hydrothermal alteration of oceanic crust. Earth Plan. Sci. Lett. 46,201-211. MCCULLOCH, M.T., GREGORY, R.T., WASSERBURG, G.J. & TAYLOR, H.P. 1981. Sm-Nd, Rb-Sr and lSo-l~O isotopic systematics in an oceanic crustal section, evidence from the Samail ophiolite. J. Geophys. Res. 86, 2721-2736. MACDONALD, G.A. & KATSURA, Z. 1964. Chemical composition of Hawaiian lavas. J. Petrol. 5, 82-133. MCELHINNEY, M.W., BR1DEN, J.C., JONES, D.L. & BROCK, A. 1968. Geological and geophysical implications of palaeomagnetic results from Africa. Key. Geophys. 6, 201-238. MCKENZIE, D.P. 1967. Some remarks on heat flow and gravity anomalies. J. Geophys. Res. 72, 6261-6273. MCKENZIE, D.P. 1969. Speculations on the consequences and causes of plati~ motion. Geophys. J. R. astr. Soc. 18, 1-18. MAGARITZ, M. & TAYLOR, H.P.JR. 1974. Oxygen and hydrogen isotope studies of serpentinization in the Troodos ophiolite complex, Cyprus. Earth Plan. Sci. Lett. 23, 8-14. MALPAS, J. 1976. The petrology and petrogenesis of the Bay of Islands ophiolite, Newfoundland. Unpubl. PhD thesis, Memorial University, Newfoundland. MALPAS, J. 1977. Petrology and tectonic significance of Newfoundland ophiolites with examples from the Bay of Islands. North Am. ophiolites Bull. 95, 13-23. MALPAS, J. 1978. Magma generation in the upper mantle, field evidence from ophiolite suites and application to the generation of oceanic lithosphere. Phil. Trans. Roy. Soc. Lond. 288,527-546. MALPAS, J. 1979. The dynamothermal aureole of the Bay of Islands ophiolite suite. Can. J. Earth Sci. 16, 2086-2101. MANGHNANI, M.H. & COLEMAN, R.G. 1981. Gravity profiles across the Samail ophiolite Oman. J. Geophys. Res. 86 (B4), 2509-2525. MARCOUX, J. 1970. Age carnien des termes effusifs du cortege ophiolitique des nappes d'Antalya (Taurus Lycien oriental, Turquie). C.R. Acad. Sci. Paris 271,285-287. MARSH, N.G., SAUNDERS, A.D., TARNEY, J. & DICK, H.J.B. 1980. Geochemistry of basalts from the Shikoku and Daito Basins, DSDP Leg 58. In: KLEIN, G. DE V., KOBAYASHI,K. et al. (eds) Initial Reports of the Deep Sea Drilling Project 58, US Govt. Printing Office, Washington, 753-800. MATTEY, D.P. MARSH, N.G. & TARNEY, J. 1980. The geochemistry, mineralogy and petrology of igneous rocks from the West Philippine and Parece Vela Basins, and from the Kyushu-Palau and West Mariana Ridges, IPOD Leg 59. Initial Reports of the Deep Sea Drilling Project, 753-797. MENZIES, M.A. 1976. Rare earth geochemistry of fused ophiolitic and alpine lherzolites- 1 0 t h r i s , Lanzo and Troodos. Geochim. Cosmochim. Acta 40, 645-656. MENZIES, M. & ALLEN, C. 1974. Plagioclase lherzolite-residual mantle relationships within two eastern Mediterranean ophiolites. Contrib. Min. Pet. 45, 197-213.

I73

MENZIES, M., BLANCHARD,D. & JACOBS, J. 1977. Rare earth and trace element geochemistry of metabasalts from the Point Sal ophiolite, California. Earth Plan. Sci. Lett. 37,203. MENZIES, M., SEYFRIED, W., & BLANCHARD, D. 1979. Experimental evidence of rare-earth element immobility in greenstones. Nature 282,398-399. MERCIER, J-C. 1980. Single pyroxene thermobarometry. Tectonophysics 70, 1-37. MESORIAN, H., (et al.) 1973. Idees actuelles sur la constitution, l'origine et I'evolution des assemblage ophiolitiques mesogenes. Bull. Geol. Soc. Fr. 15,478-493. MEVEL, C. 1981. Occurrence of pumpellyite in hydrothermally altered basalts from the Vema Fracture Zone (MAR). Contrib. Min. Pet. 76, 386-393. MICHARD, A. 1983. Contribution a la conaissance de la marge nord de Gondwana, une chaine plisee Paleozoique, vraisenblement Hercynienne en Oman. C.R. Acad. Sci. Paris, 22. MICHARD, A., BOUCHEZ, J.L. & OUAZZANI-TOUHAMI, M. 1984. Obduction-related planar and linear fabrics in Oman. J. struct. Geol. 6, 39-49. MITCHELL, A.H.G. 1984. Post-Permian events in the Zangpo 'suture' zone, Tibet. J. geol. Soc. Lond. 141, 129-136. MIYASHIRO, A. 1975. Classification, characteristics, and origin of ophiolites. J. Geol. 83, 249-281. MIYASHIRO, A., SHIDO, F. & EWING, M. 1971. Metamorphism in the Mid-Atlantic Ridge near 24~ and 30~ Phil. Trans. Roy. Soc. Lond. A268, 589-603. MOLNAR, P. & GRAY, D. 1979. Subduction of continental lithosphere: Some constraints and uncertainties. Geology 7, 58-62. MONTADART, L., ROBERTS, D.G., DE CHARPEL, O. & GUENNOC, P. 1979. Rifting and subsidence of the northern continental margin of the Bay of Biscay. In: Initial Reports of the Deep Sea Drilling Project 48, U.S. Govt. Printing Office, Washington, 1025-1060. MONTIGNY, R., BOUGAULT, H., BOTHNGA, Y. & ALLEGRE, C.J. 1973. Trace element geochemistry and genesis of the Pindos ophiolite suite. Geochim. Cosrnochim. Acta 37, 2135-2147. MOODY, J.D. 1974. Late Cretaceous nappes in the Oman Mountains and their geologic evolution: discussion. Bull. Am. Assoc. Petrol. Geol. 58,889-895. MOODY, J.B. 1979. Serpentinites, spilites and ophiolitic metamorphism. Can. Min. 17, 871-887. MOORES, E.M. 1969. Petrology and structure of the Vourinos ophiolitic complex of northern Greece. Geol. Soc. Am. Spec. Paper ll8, 74pp. MOORES, E.M. 1982. Origin and emplacement of ophiolites. Rev. geophys. SpacePhys. 20, 735-760. MOORES, E.M. & VINE, F.J. 1971. Troodos Massif, Cyprus and other ophiolites as oceanic crust: evaluation and implications. Phil. Trans. Roy. Soc. Lond. A268, 443-466. MOORES, E.M. & JACKSON, E.D. 1974. Ophiolites and oceanic crust. Nature 250, 136-139. MOORES, E.M., ROEDER, D.H., ABBAS, S.G. & AHMAD, Z. 1980. Geology and emplacement of the Muslim Bagh ophiolite complex. In: PANAYIOTOU A. (ed.) Ophiolites, Proceedings of the International Ophiolite Symposium, Cyprus 1979, 424-429. MOORES, E.M., ROBINSON, P.T., MALPAS, J. & XENOPHONTOS, C. 1984. Model for the origin of the Troodos massif, Cyprus and other mideast ophiolites. Geology 12,500-503. MORELLI, C., GANTER, C., HONKASALO, T., MCCONNELL, R.K., SZABO, B., TANNER, J.G., UBITILA, U. & WHALEN, C.Z. 1971. The International Gravity Standardization Net 1971. Report to IUGG, Moscow 1971. MORTON, D.M. 1959. The geology of Oman. 5th World Petrol. Cong. New York Proc., Section 1,277-294. MOSELEY, F. 1969. The Upper Cretaceous ophiolite complex of Masirah Island, Oman. J. Geol. 6, 293-306. MOSELEY, F. & ABBOTrS, I.L. 1979. The ophiolitic melange of Masirah Oman. J. geol. Soc. Lond. 136, 713-724. MOTTL, M.J. 1983. Metabasalts, axial hotsprings, and the structure of hydrothermal systems at mid-ocean ridges. Bull. geol. Soc. Am. 94, 161-180. MOTrL, M.J. & HOLLAND, H.D. 1978. Chemical exchange during hydrothermal alteration of basalt by seawater - 1. Experimental

I74

References

results for major and minor components of seawater. Geochim. Cosmochim. Acta 42, 1103-1117. MOyrL, M.J. & SEYFRIED, W.E. 1980. Subseafloor hydrothermal systems: rock vs seawater dominated. In: RONA, P. (ed) Benchmark papers in Geology. MOI-rL, M.J., HOLLAND, H.D. & CANN, R.F. 1979. Experimental basalt-seawater interation - hydrothermal alteration. Geochim. Cosmochim. Acta 43,869-884. MOUVEAUX, et al. 1981. Ofioliti suppl. 6 (31-2 abs). MULLINS, H.T. & LYNTS, G.W. 1977. On the origin of the NW Bahama Platform, review and reinterpretation. Bull. geol. Soc. Am. 88, 1447-1461. MtJRRIS, R.J. 1980. Middle East: stratigraphic evolution and oil habitat. Bull. Am. Assoc. Petrol. Geol. 64, 597-618. NEAL, C. & STANGER, G. 1984. Calcium and magnesium hydroxide precipitation from alkaline groundwaters in Oman. and their significance to the process of serpentinization. Min. Mag. 48, 237-241. NEARY, C.R. & BROWN, M.A. 1979. Chromites from the Al Ays Complex, Saudi Arabia and the Semail Complex, Oman. Institute of Applied Geology, King Abdulaziz University, Jeddah, Kingdom of Saudi Arabia, Bull. 3. Evolution and Mineralisation of the Arabian Shield 193-205. Pergamon Press. NESBITT, H.W. & BRICKER, O.P. 1978. Low temperature alteration processes affecting ultramafic bodies. Geochim. Cosmochim. Acta 42. 403-409 NICOLAS, A. & JACKSON, E.D. 1972. Repartitions en deux provinces des peridotites des chaines alpines longeant la Mediterranee, implications geotectoniques. Bull. Swiss. Min. Petr. 52,479-95. NICOLAS, A. & POIRIER, J.D. 1976. Crystalline plasticitv attd solid state flow in metamorphic rocks, John Wiley & Sons, 444pp. NICOLAS, A. & PRINZHOFER, A. 1982. Cumulative or residual origin for the transition zone in ophiolites : structural evidence. J. Petrol. 24 (2), 188-206. NICOLAS, A. & VIOLEYrE, J.F. 1982. Mantle flow at oceanic spreading centres : models derived from ophiolites, Tectonophysics 81, 319-339. NICOLAS, A. & RABINOWICZ, M. 1984. Mantle flow pattern at oceanic spreading centres : relation with ophiolitic and oceanic structures. In: GAss, I.G., LIPPARD, S.J. & SHELTON, A.W. (eds.) Ophiolites and Oceanic Lithosphere. Spec. Publ. Geol. Soc. London 14, Blackwell Scientific Publications, Oxford, 147-151. NICOLAS, A., BOUDIER, F. & BOULLIER, A.M. 1973. Mechanisms of flow in naturally and experimentally deformed peridotites. Am. J. Sci. 272,853-876. NICOLAS, A., BOUDIER, F. & BOUCHEZ, J-L. 1980. Interpretation of peridotite structures from ophiolitic and oceanic environments. Am. J. Sci. 280--A, 192-210. NICOLAS, A., GIRARDEAU, J., MARCOUX, J,, DUPRE, B., XIBIN, W., YOUNGONG, C., HAIXIANG, Z. & XUCHANG, X. 1981. The Xigaze ophiolite (Tibet): a peculiar oceanic lithosphere. Nature 294, 414-417. NISBET, E.G. ,~ PEARCE, J.A. 1977. Clinopyroxene composition in mafic lavas from different tectonic settings. Contrib. Mitt. Pet. 63, 149-160. NITSCH, K-H. 1971. Stabilitatsbeziehungen yon prehnit- und pumpellyit-haltigen paragenesen. Contrib. Min. Pet. 30,240-260. NOIRET, G., MONTIGNY, R. & ALLEGRE, C.J. 1981. Is the Vourinos complex an island arc ophiolite? Earth Plan. Sci. Lett. 56,375-86. O'CONNOR, J.T. 1965. A classification of quartz-rich igneous rocks based on feldspar ratios. US Geol. Surv. prof. paper 525-B. O'HARA, M.J. 1965. Primary magmas and the origin of basalts. Scott. J. Geol. 1, 19-40. O'HARA, M.J. 1968. Are ocean floor basalts primary magma? Nature 220, 683-686. O'HARA, M.J. 1977. Geochemical evolution during fractional crystallization of a periodically refilled magma chamber. Nature 266, 503-507. O'HARA, M.J. & MATTHEWS, R.E. 1981. Geochemical evolution in an advancing, periodically replenished, periodically tapped, continuously fractionated magma chamber. J. geol. Soc. Lond. 138, 237-278. OHNENSTE'VrER, M., OHNENSTETrER, D., VIDAL, P.L., CORNICHET, J., HERNUVrE, D. & MACE, J. 1981. Crystallisation and age of zircon

from Corsican ophiolitic albitites : consequences for oceanic expansion in Jurassic times. Earth Plan. Sci. Lett. 54, 397--408. ORCUTT, J.A., MCCLAIN, J,S. & BURNETr, M. 1984. Evolution of the ocean crust: results from recent seismic experiments. In: GASS, I.G., LIPPARD, S.J. & SHELI'ON, A.W. (eds.) Ophiolites and Oceanic Lithosphere. Spec. Publ. Geol. Soc. London 14, Blackwell Scientific Publications, Oxford, 7-16. OUDIN, E., PICOR, P. & POUIT, G. 1981, Comparison of sulphide deposits from the East Pacific Rise and Cyprus. Nature 291, 404-407. OXBURGn. E.R. & TURCONE, D.L. 1968. Problems of high heat flow and volcanism associated with zones of descending mantle convective flow. Nature 216, 1041-1043. PAGE, N.J., PAl,LISTER, J.S., BROWN, M.A., SMEWING, J.D. & HAFFTY, J. 1982. Palladium, platinum, rhodium, iridium and ruthenium in chromite-rich rocks from the Samail ophiolite, Oman. Can. Mineralogist. 20, 537-548. PAI,LISTER, J.S. 1981. Structure of the sheeted dyke complex of the Samail ophiolite near Ibra, Oman. J. Geophys. Res. 86, 2661-2672. PALLISTER, J.S. 1984. Parent magmas of the Semail ophiolite, Oman. In: GASS, I.G., LIPPARD, S.J. & SHELTON, A.W. (eds.) Ophiolites and Oceanic Lithosphere. Spec. Publ. Geol. Soc. London 14, Blackwell Scientific Publications. Oxford, 63-70. PALLISTER, J.S. & HOPSON, C.A. 1980. Samail ophiolite magma chamber II. Evidence for cryptic variation and mineral chemistry. In: PANAYIOTOU A. (ed.) Ophiolites, Proceedings of the International Ophiolite Symposium, Cyprus 1979, 405-406. PALLISTER, J.S. & HOPSON, C.A. 1981. Samail ophiolite plutonic suite: field relations, phase variations, cryptic variation and layering, and a model of a spreading ridge magma chamber. J. Geophys. Res. 86, 2593-2644. PALLISTER~ J.S. & KNIGHT, R.J. 1981. Rare-earth element geochemistry of the Samail ophiolite near Ibra, Oman. J. Geophys. Res. 86, 2673-2698. PAMIC, J. & MAJER, V. 1977. Ultramafic rocks of the Dinaride central zone in Yugoslavia. J. Geol. 85,553-568. PAMIC, J., SESTINI, G. & ADIB, D. 1979. Alpine magmatic and metamorphic processes and plate tectonics in the Zagros range, Iran. Bull. geol. Soc. Am. 90, 569-576. PARROT, J.F. 1973. Petrologie de la coupe du Djebel Monssa, massif basique - ultrabasique du Kizil Dag (Hatay, Turquie). Sci. Terre. 18, 143-172. PARROT, J.F. 1977a. Ophiolites du nord-ouest Syrien et evolution de la croute oceanique tethysienne au cours du Mesozoique. Tectonophysics 41, 251-269. PARROT, J.F. 1977b. Assemblage ophiolitique du Baer-Bassit et termes effusifs du volcano-sedimentaire. Petrologie d'un fragment de ia croute oceanique tethysienne charriee sur la platform Syrienne. Tray, et Doc. ORSTOM, 72 Paris, 333pp. PARROT, J.F. 1980. The Baer-Bassit (Northwestern Syria) Ophiolitic Area. In: Roccl, G. (ed) Ofioliti Spec. Issue. Tethyan Ophiolites, Voi. 2 Eastern Area, 279. PEARCE, J.A. 1975. Basalt geochemistry used to investigate post tectonic environments on Cyprus. Tectonophysics 25, 41-67. PEARCE, J.A. 1980. Geochemical evidence for the genesis and eruptive setting of lavas from Tethyan ophiolites. In: PANAYIOTOU,A (ed). Ophiolites, Proceedings of the International Ophiolite Symposium, Cyprus 1979, 261-272. PEARCE, J.A. & CANN, J.R. 1971. Ophiolite origin investigated by discriminant analysis using Ti, Zr and Y. Earth Plan. Sci. Lett. 12, 33%349. PEARCE, J.A. & CANN, J.R. 1973. Tectonic setting of basic volcanic rocks determined using trace element analysis. Earth Plan. Sci. Lett. 12, 339-349. PEARCE, J.A., ALABASTER, T., SHELTON, A,W. & SEARLE, M.P. 1981. The Oman ophiolite as a Cretaceous arc-basin complex: evidence and implications. Phil. Trans. Roy. Soc. Lond. A3, 299-317. PEARCE, J.A., HARRIS, N.B.W. & TINDLE, A.G. 1984. Trace element discrimination diagrams for the tectonic interpretation of granitic rocks. J. Petrol. 25, 956--983. PEARCE, J.A., LIPPARD, S.J. & ROBERTS, S. 1984. Characteristics and tectonic significance of supra-subduction zone ophiolites. In:

References KOKELAAR, B.P. & HOWELLS, M.F. (eds) Marginal Basin Geology Spec. Publ. Geol. Soc. London, Blackwell Scientific Publications, Oxford, 77-94. PE-PIPER, G. 1982. Geochemistry, tectonic setting and metamorphism of mid-Triassic volcanic rocks of Greece. Tectonophysics 85, 253-272. PEDERSEN, R.B. & MALPAS, J. 1984. The origin of oceanic plagiogranites from the Karmoy ophiolite, western Norway. Contrib. Min. Pet. 88, 36-52. PERINCEK, D. 1980. Volcanics of Triassic age in Bitlis metamorphic rocks. Bull. Geol. Soc. Turkey, 23, 201-203. PENROSE CONFERENCE PARTICIPANTS 1972. Penrose field conference on ophiolites. Geotimes 17, 24-25. PETERS, T. & KRAMERS, J.D. 1974. Chromite deposits in the ophiolite complex of North Oman. Min. Deposita 9,253-259. PILGRIM, G. 1908. Geology of the Persian Gulf and adjoining portions of Persia and Arabia. Ind. Geol. Surv. Mere. 34, 1-77. POIRIER, J.P. 1975. On the slip systems of olivine. J. Geophys. Res. 80, 4059-4061. POIRIER, J.P. & NICOLAS, A. 1975. Deformation induced recrystallisation due to progressive misorientation of subgrains, with special reference to mantle peridotites. J. Geol. 83,707-720. POWELL, R. 1978. The thermodynamics of the pyroxenes geotherms. Phil. Trans. Roy. Soc. Lond. A288, 457-469. PRICE, R.A. & MOUNTJOY, E.W. 1970. Geologic structure of the Canadian Rocky Mountains between Bow and Athabasca Rivers - Progress report. Geol. Soc. Can. Spec. Paper 6, 7-25. PRINZHOFER, A. & ALLEGRE C.J. 1985. Residual peridotites and the mechanisms of partial melting. Earth Plan. Sci. Len. 74,251-265. RAASE, P. 1974. A1 and Ti contents of hornblendes, indicators of temperature and pressure of regional metamorphism. Contrib. Min. Pet. 45,231-236. RABINOWICZ, M., NICOLAS, A. & VIGNERESSE,J-L. 1984. A rolling-mill effect in asthenosphere beneath spreading centres. Earth Plan. Sci. Lett. 67, 97-108. RAHEIM, A. & GREEN, D.H. 1974. Experimental determination of the temperature and pressure dependence of the Fe-Mg partition coefficient for co-existing garnet and clinopyroxene. Contrib. Min. Pet. 48, 179-203. RAI~, R.W. 1963. The crustal rocks. In: HILL, M.N. (ed) The Sea, Vol 3, Wiley, New York, 85-102. READMAN, P.W. & O'REILLV, W. 1972. Magnetic properties of oxidised (cation deficient) titanomagnetites ((Fe, Ti, D)304). J. geomagn. Geoelect. 24, 69-90. REINHARDT, B.M. 1969. On the genesis and emplacement of ophiolites in the Oman mountains geosyncline. Schweiz. Min. Pet. Mitt. 49, 1-30. REINHARDT, B.M. 1970. On the genesis and emplacement of ophiolites in the Oman mountains geosyncline. Geol. Mijnbouw. 49, 161-163. REINHARDT, B.M. 1974. The relationship between spilites and other members of the Oman mountain's ophiolitic suite. In: AMSrUTZ. G.C. (ed) Spilites and Spilitic Rocks, Springer-Verlag, N.Y., 207-227. REX, D.C. & DODSON, M.H. 1970. Improved resolution and precision of argon analysis using an MS 1 mass spectrometer. Eclog. ,geol. Helv. 63, 275-280. RICATEAU, R. & RICHLY,H. 1980. Geology of the Musandam Peninsula (Sultanate of Oman) and its surroundings. J. Petrol. Geol. 3, 139-152. RICHTER, F.M. 1973. Finite amplitude convection through a phase boundary. Geophys. J. R. astr. Soc. 35,265-276. RICHTER, F.M. & MCKENZIE, D.P. 1978. Simple plate models of mantle convection, J. Geophys. 44,441-71. RIcou, L.E. 1968. Sur la raise en place au Cretace superieur cl'importantes nappes a radiolarites et ophiolites darts les Monts Zagros (Iran). C.R. Acad. Sci. Paris, 267, 2272-2275. RIcou, L.E. 1971a. Le croissant ophiolitique peri-Arabe, une ceinture des nappes mises en place au Cretace Superieure. Rev. Geograph. Phys. Geol. Dyn. Paris (2) 13, 327-350. Rlcou, L.E. 1971b. Le metamorphisme au contact des peridotites de Neyriz (Zagros interne, Iran): developpement de skarns a pyroxene. Bull. Soc. Geol. Fr. 13, 146-155. Rlcou, L.E. 1974. L'etude geolgique de la region de Neyriz, Zagros

I75

Iranien, et l'evolution structurale des Zagrides. Universite de Paris-Sud, Orsay, Series A 126A, 321pp. RIcou, L.E. 1976. Evolution structurale des Zagrides. Le region clef de Neyriz (Zagros iranien). Mere. Soc. Geol. Fr. 55, 125, 140pp. Rlcou, L.E., BRAUD,J. & BRUNN, J.H. 1977. Le Zagros. Mem. h. ser. Soc. Geol. Fr.(7), 17, 1024-1044. ROBERTS, S. 1985. The role of igneous processes in the formation of ophiolitic chromitite. Unpubl. PhD Thesis, Open University. ROBERTSON, A.H.F. & HUDSON, J.D. 1973. Cyprus umbers: chemical precipitates on a Tethyan ocean ridge. Earth Plan. Sci. Len. 18, 93-101. ROBERTSON, A.H.F. & STILLMAN, C.J. 1979. Submarine volcanism and associated sedimentary rocks of the Fuerteventura Basal Complex, Canary Islands. Geol. Mag. 116, 203-214. ROBERTSON, A.H.F. & WOODCOCK, N.H. 1979. Mamonia complex, southwest Cyprus, evolution and emplacement of a Mesozoic continental margin. Bull. geol. Soc. Am. 90, 651-665. ROBERTSON, A.H.F. & WOODCOCK,N.H. 1981. Godene zone, Antalya Complex: volcanism and sedimentation along a Mesozoic continental margin, SW Turkey. Geol. Rundschau 70, 1177-1214. ROBERTSON, A.H.F. & WOODCOCK, N.H. 1983a. Genesis of the Batinah Melange above the Semail ophiolite, Oman. J. Struct. Geol. 5, 1-17. ROBERTSON, A.H.F. & WOODCOCK, N.H. 1983b. Zabyat Formation, Semail Nappe, O m a n Sedimentation onto an emplacing ophiolite. Sedimentology 30, 105-116. ROBINSON, P.T., MELSON, W.G., O'HEARN, T. & SCHMINCKE, H-U. 1983. Volcanic glass compositions of the Troodos ophiolite, Cyprus. Geology 11,400-404. Roccl, G. & LAPIERRE, H. 1969. Etude compartiv des diverse manifestations du voicanisme preorogenique au sud du Chypre. Schweiz. mineral, petrol. Mitt. 49, 31-46. RoccI, G., OHNENSTETYER, D. & OHNENSTEITER. M. 1975. La dualite des ophiolites tethysiennes. Petrologie 1, 172-174. ROEDER, P.L., CAMPBELL, I.H. & JAMIESON, H.E. 1979. A re-evaluation of the olivine spinel geothermometer. Contrib. Min. Pet. 68,325-335. RONA, P.A. 1978. Criteria for recognition of hydrothermal mineral deposits in oceanic crust. Econ. Geol. 73, 135-160. RONA, P.A. 1984. Hydrothermal mineralization at sea-floor spreading centres. Earth Sci. Rev. 20, 1-104. ROSENCRANTZ, E. 1983. The structure of sheeted dikes and associated rocks in North Arm massif, Bay of Islands ophiolite complex and the intrusive processes at oceanic spreading centres. Can. J. Earth Sci. 20,787-801. ROSENDAHL, B.R., RAIFI-, R.W., DORMAN, L.M., BIBEE, L.D., HUSSONG, D.M. & SuttoN, G.H. 1976. Evolution of oceanic crust. A physical model of the East Pacific Rise crest derived from seismic refraction data. J. Geophys. Res. 81, 5294-5304. Ross, J.W., MERCIER, J.F., AVE LALLEMENT, H.G., KARTER, H.L. & ZIMMERMAN, J. 198(/. The Vourinos Ophiolite Complex, Greece: the tectonite suite. Tectonophysics 70, 63-83. ROTHERY, D.A. 1982. The evolution of the Wuqbah block and the applications of remote sensing in the Oman ophiolite. Unpubl. PhD Thesis, Open University, 414pp. ROTHERY, D.A. 1983. The base of the sheeted dyke complex, Oman Ophiolite: implications for magma chambers at oceanic spreading axes. J. geol. Soc. Lond. 140, 287-296. ROTHERY, D.A. 1984. The role of Landsat Muitispectral Scanner (MSS) imagery in mapping the Oman ophiolite. In: GASS, I.G., LIPPARD, S.J. & SHELTON, A.W. (eds.) Ophiolites and Oceanic Lithosphere. Spec. Publ. Geol. Soc. London 14, Biackwell Scientific Publications, Oxford, 405-413. SABZEHEI, M. 1974. Les melanges ophiolitique de la region d'Esfandagheh (Iran meridienal). Etude petrologique et structurale. Interpretation dans le cadre lranien. Doct III ~m~ cycle Thesis, University of Grenoble, 205pp. SADRED1NI, E. 1974. Geologie und Petrographie im Mittelteil des Ophiolithzuges nordlich SabzevarKhorassan (Iran). Dissertation Univ. Saarbrucken, 120pp. SAUNDERS, A.D. & TARNEY, J. 1979. The geochemistry of basalts from a back-arc spreading centre in the East Scotia Sea. Geochim. Cosmochim. Acta 43, 555-572.

I76

References

SAXENA, S.K. 1979a. Garnet-clinopyroxene geothermometer. Contrib. Min. Pet. 70, 229-235. SCHMINCKE H-U., RAUTENSCHLEIN,M., ROBINSON, P.T. & MEHEGAN, J.M, 1983. Troodos extrusive series of Cyprus: A comparison with oceanic crust. Geology 11,405-409. SCLATER,J.G. & FRANCHETEAU,J. 1970. The implications of terrestrial heat flow observations on current tectonic and geochemical models of the crust and upper mantle. Geophys. J. R. astr. Soc. 20, 509-535. SEARLE, M.P. 1980. The metamorphic sheet and underlying volcanic rocks beneath the Semail ophiolite in the northern Oman mountains of Arabia. Unpubl. PhD Thesis, Open University, 213pp. SEARLE, M.P. 1984. Alkaline peridotite, pyroxenite, and gabbroic intrusions in the Oman Mountains, Arabia. Can. J. Earth Sci. 21, 396-446. SEARLE, M.P. & GRAHAM, G.M. 1982. The "Oman Exotics" - oceanic carbonate build-ups associated with the early stages of continental rifting. Geology 10, 43-49. SEARLE, M.P. & MALPAS, J. 1980. Structure and metamorphism of rocks beneath the Semail ophiolite of Oman and their significance in ophiolite obduction. Trans. Roy. Soc. Edinburgh, Earth Sciences 71,247-262. SEARLE, M.P. & MALPAS, J. 1982. Petrochemistry and origin of subophiolite metamorphic and related rocks in the Oman mountains. J. geol. Soc. Lond. 139, 5-248. SEARLE, M.P., LIPPARD, S.J., SMEWING, J.D. & REX, D.C. 1980. Volcanic rocks beneath the Semail ophiolite nappe in the northern Oman mountains and their significance in the Mesozoic evolution of Tethys. J. geol. Soc. Lond. 137,589-604. SEARLE, M.P., JAMES, N.P., CALON, T.J. & SMEWING, J.D. 1983. Sedimentological and structural evolution of the Arabian continental margin in the Musandam Mountains and Dibba zone, United Arab Emirates. Bull. geol. Soc. Am. 94, 1381-1400. SENGOR, A.M.C. 1985. Story of Tethys: How many wives did Okeanos have? Episodes 8, 3-12. SENGOR, A.M.C. & KIDD, W.S.F, 1979. Post-collisional tectonics of tlae Turkish-Iranian plateau and a comparison with Tibet. Tectonophysics 55,361-376. SEYFR1ED, W.E., BISCHOFF, J.L. & MorrL M.J. 1978. Seawater/basalt ratio effects on the chemistry and mineralogy of spilites from the ocean floor. Nature 275, 211-213. SEYFRIED, W.E. JR., MO'i'rL, M.J. & JANECKY, D.R. 1979. Origin of metalliferous deposits of the Troodos ophiolite: experimental seawater-basalt interaction in a sea-water dominated system. Abst. Prog. Geol. Soc. Amer. Ann. Meeting 11,514. SHEARMAN, D.J. 1976. The geological evolution of southern Iran. The Report of the Iran-Makran Expedition. Geograph. J. London 142,393-404. SHELTON, A.W. 1984. Geophysical studies on the Northern Oman Ophiolite. Unpubl. PhD thesis, Open University, 353pp. SIGVALDASON, G.E. 1962. Epidote and related minerals in two deep geothermal drill holes, Reykjavik and Hveragerdi, Iceland. U.S. Geol. Sur. Prof. Paper 450-E, 77-84. SIMONIAN, K.O. & GASS, I.G. 1978. The Arakapas fault belt, Cyprus : a fossil transform fault. Bull. geol. Soc. Am. 89, 1220-1230. SKELTON, P. & NOLAN, S. 1985. Rudist palaeoecological zonation in a Maastrichtian beach to shoreface environment, Jebel Faiyan, U.A.E. Abstracts of the Palaeont. Assoc. Conference, Aberystwyth 1985, 19. SMEWIN6, J.D. (ed), 1979. Open University Oman Project Map Sheet 1 - The Sumeini-Shinas Area. Directorate of Overseas Surveys. SMEWING, J.D. 1980a. Regional setting and petrological characteristics of the Oman ophiolite in North Oman. In: Roccl, G. (ed.) Ofioliti Special Issue on Tethyan Ophiolites Vol 2 Eastern Area, 335-378. SMEWIN6, J.D. 1980b. An Upper Cretaceous ridge-transform intersection in the Oman ophiolite. In: PANAYIOTOUA. (ed.) Ophiolites, Proceedings of the International Ophiolite Symposium, Cyprus 1979, 407-413. SUEWIN6, J.D. 1981. Mixing characteristics and compositional differences in mantle-derived melts beneath spreading axes: Evidence from cyclically layered rocks in the ophiolite of North Oman. J. Geophys. Res. 86, 2645-2660.

SMEWING, J.D., SIMONIAN, K.O., EL BOUSHI, I.M. & GASS, I.G. 1977. Mineralised fault zone parallel to the Oman ophiolite spreading axis. Geology 5,534--538. SMEWING, J.D., CHRISTENSEN, N.I., BARTHOLOMEW, I.D. & BROWNIN6, P. 1984. The structure of the oceanic upper mantle and lower crust as deduced from the northern section of the Oman ophiolite. In: GASS, I.G., LIPPARD, S.J. & SHELTON, A.W. (eds.) Ophiolites and Oceanic Lithosphere. Spec. Publ. Geol. Soc. London 14, Biackwell Scientific Publications, Oxford, 41-54. SMITH, A.G. 1971. Alpine deformation and the oceanic areas of the Tethys, Mediterranean and Atlantic. Bull. geol. Soc. Am. 82, 2039--2070. SMITH, A.G. 1981. Subduction and coeval thrust belts, with particular reference to North America. In: MCCLAY, K.R. & PRICE, N.J. (eds.) Thrust and nappe tectonics. Geol. Soc. Spec. Publ. 9, Blackwell Scientific Publications, Oxford, 111-124. SMITH, A.G. & BRIDEN, J.C. 1977. Mesozoic and Cenozoic Paleocontinental Maps, Cambridge University Press, Cambridge, 65pp. SMITH, A.G. & WOODCOCK,N.H. 1976. Emplacement model for some "Tethyan" ophiolites. Geology 4, 653-656. SPARKS, R.S.J., SIGURDSSON, H. & MEYER, P. 1979. Density variation amongst mid-ocean ridge basalts: implications for magma mixing and the scarcity of primitive lavas. LOS 60, 971 (abst). SPOONER, E.T.C. & FYFE, W.S. 1973. Sub-sea-floor metamorphism, heat and mass transfer. Contrib. Min. Pet. 42,287-304. SPRAY, J. 1982. Mafic segregations in ophiolite mantle sequences, Nature 299,524-528. SPRAY, J.G. I984. Possible causes and consequences of upper mantle decoupling and ophiolite displacement. In: GASS, I.G., LIPPARD, S.J. ,s, SHELTON, A.W. (eds.) Ophiolites and Oceanic Lithosphere. Spec. Publ. Geol. Soc. London 14, Blackweli Scientific Publications, Oxford 255-268. SPRAY, J.G. & RODDICK, J.C. 1980. Petrology and 4~ geochronology of some Hellenic sub-ophiolite metamorphic rocks Contrib. Min. Pet. 72, 43-55. SPRAY, J.G. & RODDICK, J.C. 1981. Evidence for Upper Cretaceous transform fault metamorphism in West Cyprus, Earth Plan. Sci. Lett. 55, 273-291. SPULBER, S.D. & RUTHERVORD,M.J. 1983. The origin of rhyolite and plagiogranite in oceanic crust: an experimental study. J. Petrol. 2 4 ( 1 ) , 1-25. STAKES, D.S. a O'NEIL, J.R. 1982. Mineralogy and stable isotope geochemistry of hydrothermally altered oceanic rocks. Earth Plan. Sci. Lett. 57,285-304. STAKES, D.S., TAYLOR, H.P. JR. & FISHER. R.L. 1984. Oxygen isotope and geochemical characterization of hydrothermal alteration in ophiolite complexes and modern oceanic crust. In: GAss, I.G., LIPPARD, S.J. & SHELTON, A.W. (eds.) Ophiolites and Oceanic Lithosphere. Spec. Publ. Geol. Soc. London 14, Blackwell Scientific Publications, Oxford, 199-214. STANGER, G. 1985. Silicified serpentinite in the Semail Nappe of Oman. Lithos 18, 13-22. STANGER, G. 1986. The hydrogeology of the Oman Mountains. Unpubl. Ph.D. Thesis, Open University, 572pp. STAUDIGEL, H., HART, S.R. & RICHARDSON, S.H. 1981. Alteration of the oceanic crust: processes and timing. Earth Plan. Sci. Lett. 52, 311-327. STHGER, R.H. a JAGER E. 1977. Subcommission on Geochronology: Convention of the use of decay constants in geo- and cosmochronology. Earth Plan. Sci. Lett. 36,359-362. STEINMANN, G. 1927. Die ophiolithischen Zonen in den mediterranen Kettengebirgen. Congres Geol. Internat, )(IV Sess. Madrid, 637-677. STERN, C., DEWIT, M.J. & LAWRENCE, J.R. 1976. Igneous and metamorphic processes associated with the formation of Chilean ophiolites and their implications for ocean floor metamorphism, seismic layering and magnetism. J. Geophys. Res. 81 (23), 4370-4380. STEWART, R.A., PILrEY, O.H. & NELSON, B.W. 1965. Sediments of the northern Arabian Sea. Marine Geol. 3, 411-427. STILLE, H. 1924. Grundfragen der vergleichenden Tektonik. Borntraeger, Berlin, 443pp. STILLMAN, C.J., FURNES, H., LEBAS, M.J., ROBERTSON, A.H.F. & ZIELONKA, J. 1982. The geological history of Maio, Cape Verde Islands. J. geol. Soc. Lond. 139, 347-361.

References STOCKLIN, J. 1968. Structural history and tectonics in Iran, a review. Bull. Am. Assoc. Petrol. Geol. 52, 1229-1258. STOCKI,IN, J. 1974. Possible ancient continental margin in lran. In: BURKE, C.A. & DRAKE, C.L. (eds) The geology of continental margins. Springer-Verlag, New York, 873-887. STOLPER, E. 1980. A phase diagram for mid-ocean ridge basalts: Preliminary results and implications for petrogenesis. Contrib. Min. Pet. 74, 13-27. STONELEY, R. 1974. Evolution of the continental margins bounding a former Southern Tethys. ln: BURKE, C.A. & DRAKE, C.L. (eds) The geology or" continental margins. Springer-Verlag, New York, 889-903. SIONELEY, R. 1975. On the origin of ophiolite complexes in the Southern Tethys Region. Tectonophysics 25,303-322. STONELEY, R. 1980. The Geology of the Kuh-e-Dalneshin area of southern Iran, and its bearing on the evolution of the Southern Tethys. J. geol. Soc. Lond. 137,509-526. STURT, B.A. & THON, A. 1978. An ophiolite complex of probable early Caledonian age discovered on Karmoy. Nature 275,538-539. SUESS, E. 1909. Das Antlitz der Erde (Bd 31) Tempsky, Wien, Fr.eytag, Leipzig. TAKIN, M. 1972. Iranian geology and continental drift in the Middle East. Nature 235, 147-150. TARNEY, J., WOOD, D.A., SAUNDERS, A.D., CANN, J.R. & VARET, J. 1980. Nature of mantle heterogeneity in the North Atlantic: evidence from deep sea drilling. Phil. Trans. Roy. Soc. Lond. 297A, 179-202. TELEKI, O. 1981. Subduction complex of pre-Jurassic age, Northern Anatolia, Turkey. Geology 9, 68-72. TERRY, J. 1971. Sur l'age Triasique des laves associees a ia nappe ophiolitique du Pinde septentrional (Epire et Macedonie, Grece). C.R. somm. Soc. Geol. Ft. 7,384-385. THAKUR, V.C. 1981. Regional framework and geodynamic evolution of the Indus-Tsangpo suture zone in the Ladakh Himalayas. Trans. Roy. Soc. Edinburgh 72, 89-97. THOMPSON, G. 1973. A geochemical study of the low-temperature interaction of sea-water and oceanic igneous rocks. EOS, Trans. Am. geophys. Un. 54, 1015-1019. THUIZAT, R., WHITECHURCH, H., MONTIGNY, R. & JUTEAU, T. 1981. K-Ar dating of some infra-ophiolitic metamorphic soles from the Eastern Mediterranean: new evidence for oceanic thrustings before obduction. Earth Plan. Sci. Lett. 52,302-310. TILTON, G.R., HOPSON, C.A. & WRIGHT, J.E. 1981. Uranium-lead isotopic ages of the Samail ophiolite, Oman, with applications to Tethyan Sea ridge tectonics. J. Geophys. Res. 86, 2763-2776. TINKLER, C., WAGNER, J.J., DELALOYE, M. & SELCHLE, H. 1981. Tectonic history of the Hatay ophiolites (S. Turkey) and their relation to the Dead Sea Rift. Tectonophysics 72, 23-41. TIPPET, R.P., PESSAGNO, E.A. & SMEWING, J.D. 1981. The biostratigraphy of sediments in the volcanic unit of the Samail ophiolite. J. Geophys. Res. 86, 2756-2762. TOMASSON, J. & KRISTMANNSDOYfIR,H. 1972. High temperature alteration minerals and thermal brines, Reykjanes, Iceland. Contrib. Min. Pet. 36, 123-127. TREUIL, M. & JORON, J.L. 1974. Etude geochimique des elements entraces dans le magmatisme de l'Afar. Implications petrogenetiques et comparisons avec le magmatism de l'Islande et de la dorsale medio-Atlantique. In: PILGER, A. & ROSLER, A. (eds) Afar, between continental and oceanic rifting. Int. Union Comm. Ge0dyn. 16, 26-79. TRUMPY, R. 1960. Paleotectonic evolution of the central and western Alps. Bull. geol. Soc. Am. 71,843-908. TScHoPP, R.H. 1967. The general geology of Oman. 7th World Petrol. Cong. Mexico, Proc. vol. 2, 243-250. VAIL, P.R., MITCHUM, R.M. & THOMPSON, S. 1977. Global cycles of relative changes in sea level. In: PAYTON, C.E. (ed) Seismic Stratigraphy- Application to Hydrocarbon Exploration. Mere. Am. Assoc. Petrol. Geol., Spec. Publ. 26, 83-99. VAZIRI-TABAR, F. 1976. Geologie und Petrographie der Ophiolithe und ihrer vulkanose dimentaren Folge produkte im Osttcil des Bergzuges nordlich SabzevarKhorassan (lran). Dissertation Univ. Saarbrucken, 152pp. VENTURELLI, G., THORPE, R.S. & PoTrs, P.J. 1981. Rare earth and

177

trace element characteristics of ophiolitic metabasalts from the Alpine-Apennine belt. Earth Plan. Sci. Lett. 93, 109-123. VINE, F.J. 1966. Spreading of the ocean floor: New evidence. Science 154, (3755) 1405-1415. VINE, F.J. & MOORES, E.M. 1972. A model for the gross structure, petrology and magnetic properties of oceanic crust. Geol. Soc. Amer. Mere. 132, 195-205. VITA-F1NZI, C. 1973. Late Quaternary subsidence. In: The Musandam Expedition 1971-72, Scientific results: Part 1. Geograph. J. 139, 414-421. VITA-F1NZI, C. 1982. Recent coastal deformation near the Strait of Hormuz. Proc. Roy. Soc. London A 382,441-457. WAGER, L.R., BROWN, G.M. & WADSWORTH, W.J. 1960. Types of igneous cumulates. ]. Petrol. 1, 73-85. WALKER, R.G. & Mvrrl, E. 1973. Turbidite facies and facies associations. In: Turbidites and Deep water Sedimentation. Soc. Econ. Palaeont. Mineral. Pacific Section, Los Angeles, Short Course Lecture Notes. WEERTMAN, J. 1968. Dislocation climb theory of steady state creep. Trans. Amer. Soc. Metals 61,681-694. WELLEND, M.J.P. & MHCHELL, A.H.G. 1977. Emplacement of the Oman Ophiolite. A mechanism related to subduction and collision. Bull. geol. Soc. Am. 88, 1081-1088. WELLS, P.R.A. 1977. Pyroxene thermometry in simple and complex systems. Contrib. Min. Pet. 62, 129-139. WENNER, D.B. & TAYLOR, H.P.JR. 1971. Temperatures of serpentinization of ultramafic rocks based on ~sO/160 fractionation between coexisting serpentinite and magnetite. Contrib. Min. Pet. 32, 165-185. WENNER, D.B. & TAYLOR, H.P.JR. 1978. Oxygen and hydrogen isotope studies of the serpentinization of ultramafic rocks in oceanic environments and continental ophiolite complexes. Am. ]. Sci. 273,207-239. WHITE, R.S. 1977. Fold development in the Gulf of Oman. Earth Plan. Sci. Lett. 36, 85-91. WHITE, R.S. 1979. Gas hydrate layers trapping free gas in the Gulf of Oman. Earth Plan. Sci. Left. 42, 114-120. WHITE, R.S. 1982. Deformation of the Makran accretionary sediment prism in the Gulf of Oman (NW Indian Ocean). In: LEGGETr, J.K. (ed) Trench - Forearc Geology. Spec. Publ. Geol. Soc. London 10,357-372. WHITE, R.S. & KLITGORD, K.D. 1976. Sediment deformation and plate tectonics in the Gulf of Oman. Earth Plan. Sci. Lett. 32, 199-209. WHITE. R.S. & Ross, D.A. 1979. Tectonics of the Western Gulf of Oman. ]. Geophys. Res. 84, 3479-3489. WHITE, R.S. & Ross, D.A. 1982. Deformation of the Makran accretionary sediment prism in the Gulf of Oman (NW Indian Ocean). In: LEGGETT, J.K. (ed) Trench-Forearc Geology. Spec. Publ. Geol. Soc. London 10, 357-372. WHITEHEAD, J.A.. DICK, H.J.B. & SCHOUTEN, H. 1984. A mechanism for magmatic accretion under spreading centres. Nature 312, 146-148. WHITMARSH, R.B. 1979. The Owen Basin off the southwest margin of Arabia and the evolution of the Owen Fracture Zone. Geophys. J. R. astr. Soc. 58, 441-470. WICKS, F.J. & WHI-FrAKER, E.J.W. 1977. Serpentine textures and serpentinisation. Can. Mineral. 15,459-488. WILLIAMS, H. 1975. Structural succession, nomenclature and interpretation of transported rocks in W. Newfoundland. Can. J. Earth Sci. 12, 1874-1894. WILLIAMS, H. & MALPAS, J. 1972. Sheeted dikes and brecciated dike rocks within transported igneous complexes; Bay of Islands, Western Newfoundland. Can. J. Earth Sci. 9, 1216-1229. WILLIAMS, H. & SMYTH, W.R. 1973. Metamorphic aureoles beneath ophiolite suites and alpine peridotites; tectonic implications with West Newfoundland examples. Am. J. Sci. 273, 594-621. WILSON, H.H. 1969. Late Cretaceous eugeosynclinal sedimentation, gravity tectonics and ophiolite emplacement in the Oman Mountains, southeast Arabia. Bull. Amer. Assoc. Petrol. Geol. 53, 626-671. WILSON, H.H. 1972. Late Cretaceous nappes in the Oman mountains: a discussion. Bull. Am. Assoc. Petrol. Geol. 57, 2282-2287. WILSON, R.A.M. 1959. The geology of the Xeros-Troodos area. Cyprus Geol. Surv. Dept. Mem. 1, 135pp.

t78

References

WOLERY, T.J. & SLEEP, N.H. 1976. Hydrothermai circulation and

geochemical flux at mid-ocean ridges. J. Geol. 84,249-276. WOOD, D.A., GIBSON, I.L. & THOMPSON, R.N. 1976. Elemental mobility during zeolite facies metamorphism of the Tertiary basalts of Eastern Iceland. Contrib. Min. Pet. 55,241-254. WOODCOCK, N.H. & ROBERTSON, A.H.F. 1977. Origins of some ophiolite related metamorphic rocks of the Tethyan belt. Geology 5, 373-376. WOODCOCK, N.H. & ROBERTSON, A.H.F. 1982a. Stratigraphy of the Mesozoic rocks above the Semail ophiolite, Oman. Geol. Mag. 119, 67-76.

WOODCOCK, N.H. & ROBERTSON, A,H.F. 1982b. The Upper Batinah Complex, Oman: Ailochthonous sediment sheets above the Semail ophiolite. Can. Jour. Earth Sci. 19, 1635-1656. WORZEL, J.L. 1965. Deep structure of coastal margins and mid-ocean ridges. In: WHILLARD, W.F. & BRAI)SHAW, R. (eds) Submarine geology and geophysics. Butterworths, London, 335-361. WORZEL, J.L. 1974. Standard oceanic and continental structure. In: BURK, C.A. & DRAKE, C.L. (eds) The geology of continental margins. Springer-Verlag, New York, 59-66.

Vereinigte Arabische Emirate Oman

Read more

Oman A Country Study

Read more

Oman - The Islamic Democratic Tradition

Read more

Oman (World Bibliographical Series)

Read more

Child of the Northern Spring

Read more

Child of the Northern Spring

Read more

Child of the Northern Spring

Read more

The Gods of Northern Buddhism

Read more

The Army of Northern Virginia

Read more

Child Of The Northern Spri

Read more

Roles of the Northern Goddess

Read more

Child of the Northern Spring

Read more

Child of the Northern Spring

Read more

Child of the Northern Spring

Read more

John Oman and his Doctrine of God

Read more

The Northern Crusades

Read more

The Northern Devil

Read more

The Northern Crusades

Read more

The Northern Clemency

Read more

Under the Northern Lights

Read more

The Northern Clemency

Read more

The Northern Devil

Read more

The Northern Girl

Read more

The Northern Clemency

Read more

The Northern Girl

Read more

Under the Northern Lights

Read more

The Northern Clemency

Read more

Under the Northern Lights

Read more

Northern Exposure

Read more

Northern Renegade

Read more

Recommend Documents

Vereinigte Arabische Emirate Oman

Die Autorin Kirsten Baron, geboren 1954 in Kiel, studierte Agrarwissenschaften. Seit mehr als zehn Jahren ist sie als R...

Oman A Country Study

Oman, a country study Federal Research Division Kessinger Publishing, LLC ISBN-10: 1419138294 ISBN-13: 978-1419138294 F...

Oman - The Islamic Democratic Tradition

Oman – The Islamic Democratic Tradition Oman is the inheritor of a unique political tradition, the imama (imamate), an...

Oman (World Bibliographical Series)

Child of the Northern Spring

Child of the Northern Spring

Child of the Northern Spring

The Gods of Northern Buddhism

The Army of Northern Virginia

...

Child Of The Northern Spri