Alkaline Igneous Rocks
Geological Society Special Publications Series Editor
K. C 0 E
GEOLOGICAL
SOCIETY SPECIAL PUBLICATION
Alkaline Igneous Rocks E D I T E D BY
J. G. F I T T O N & B. G. J. U P T O N Grant Institute of Geology University of Edinburgh Edinburgh EH9 3JW
1987 Published for The Geological Society by Blackwell Scientific Publications OXFORD LONDON
EDINBURGH
BOSTON PALO ALTO MELBOURNE
N O 30
Published 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
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British Library Cataloguing in Publication Data
First published 1987
ISBN0-632-01616-7
Typeset, printed and bound in Great Britain by William Clowes Limited, Beccles and London
Library of Congress Cataloging-in-Publication Data
Alkaline igneous rocks.--(Geological Society special publications, ISSN 0305-8719) 1. Alkalic igneous rocks I. Fitton, J.G. II. Upton, B. G. J. III. Series 552'. 1 QE462.A4
Alkaline igneous rocks. (Geological Society special publication; no. 30) Bibliography: p. Includes index. 1. Alkalic igneous rocks. I. Fitton, J. G. II. Upton, B. G.J. III. Geological Society of London. IV. Series. QE462.A4A43 1987 552'. 1 86-26364 ISBN0-632-01616-7
Contents Preface Introduction: FITTON, J. G. & UPTON, B. G . J . BAILEY, D. K. Mantle metasomatism--perspective and prospect
vii ix 1
MENZIES, M. Alkaline rocks and their inclusions: a window on the Earth's interior
15
EDGAR,A. D. The genesis of alkaline magmas with emphasis on their source regions: inferences from experimental studies
29
LE Bhs, M. J. Nephelinites and carbonatites
53
TWYMAN, J. D. & GITTINS, J. Alkalic carbonatite magmas: parental or derivative ?
85
DAWSON, J. B. The kimberlite clan: relationship with olivine and leucite lamproites, and inferences for upper-mantle metasomatism
95
BERGMAN, S. C. Lamproites and other potassium-rich igneous rocks: a review of their occurrence, mineralogy and geochemistry
103
ROCK, N. M. S. The nature and origin of lamprophyres: an overview
191
CLAGUE, D. A. Hawaiian alkaline volcanism
227
WEAVER, B. L., WOOD, D. A., TARNEY, J. & JORON, J. L. Geochemistry of ocean island basalts from the South Atlantic: Ascension, Bouvet, St. Helena, Gough and Tristan da Cunha
253
HARRIS, C. & SHEPPARD, S. M. F. Magma and fluid evolution in the lavas and associated granite xenoliths of Ascension Island
269
FITTON, J. G. The Cameroon line, West Africa: a comparison between oceanic and continental alkaline volcanism
273
BAKER, B. H. Outline of the petroIogy of the Kenya rift alkaline province
293
MACDONALD,R. Quaternary peralkaline silicic rocks and caldera volcanoes of Kenya
313
WOOLLEY, A. R. & JONES, G. C. The petrochemistry of the northern part of the Chilwa alkaline province, Malawi
335
BOWDEN, P., BLACK, R., MARTIN, R. F., IKEI E. C. KINNAIRD, J. A. & BATCHELOR, R . A . Niger-Nigerian alkaline ring complexes: a classic example of African Phanerozoic anorogenic mid-plate magmatism
357
LI/~GEOIS, J. P. & BLACK, R. Alkaline magmatism subsequent to collision in the Pan-African belt of the Adrar des Iforas (Mali)
381
FLETCHER, C. J. N. & BEDDOE-STEPHENS, B. The petrology, chemistry and crystallization history of the Velasco alkaline province, eastern Bolivia
403
BARKER, D. S. Tertiary alkaline magmatism in Trans-Pecos Texas
415
EBY, G. N. The Monteregian Hills and White Mountain alkaline igneous provinces, eastern North America
433
vi
Contents
UPTON, B. G. J. & EMELEUS, C. H. Mid-Proterozoic alkaline magmatism in southern Greenland: the Gardar province
449
LARSEN, L. M. & SORENSEN, H. The Ilimaussaq intrusion--progressive crystallization and formation of layering in an agpaitic magma
473
NIELSEN,T. F. D. Tertiary alkaline magmatism in East Greenland: a review
489
DOWNES, H. Tertiary and Quaternary volcanism in the Massif Central, France
517
KOGARKO, L. N. Alkaline rocks of the eastern part of the Baltic Shield (Kola Peninsula)
531
Index
545
Preface The papers contained in this volume were presented at a symposium held in Edinburgh in September 1984, which marked the passage of ten years since the publication of The Alkaline Rocks edited by Henning Sorensen. In organizing the symposium and compiling this volume we aimed to review recent developments in the petrology and geochemistry of alkaline igneous rocks. We have, for example, paid particular attention to work on lamprophyres and carbonatites which are rock associations of current interest not covered in Sorensen's book. Reviews of recent work on some of the classic alkaline provinces, such as East Africa, southern Greenland and the Kola Peninsula, are included together with reviews of less wellknown areas. Other papers discuss the impact of experimental, geochemical and isotopic studies on our understanding of the generation and evolution of alkaline magmas. We are indebted to the contributors for their collaboration in producing this volume and it is with sadness that we note the death, on 14 February 1986, of Brian Baker, whose pioneering field studies formed the basis for much of our knowledge of the tectonic and volcanic evolution of the East African Rift. An obituary and appreciation of his work is published in the Journal of Volcanology and Geothermal Research (28, v-vii). We are grateful to The Geological Society and The Royal Society of Edinburgh for their generous assistance with the symposium costs, to Lucian Begg and Dodie James for their help with organizing the symposium and producing this volume, and to our colleagues for the care and enthusiasm with which they reviewed the manuscripts. The efforts of Edward Wates and his staff at Blackwell Scientific Publications are also gratefully acknowledged. To all these we offer our sincere thanks. J.G.F. B.G.J.U.
vii
Introduction J. G. Fitton & B. G. J. Upton Alkaline igneous rocks may be defined as those which have higher concentrations of alkalis than can be accommodated in feldspars alone, the excess appearing as feldspathoids, sodic pyroxenes, sodic amphiboles and other alkali-rich phases. These rocks are, therefore, deficient in silica and/or alumina with respect to alkalis and will have nepheline and/or acmite in their norms. In practice the term 'alkaline' is used to encompass a wide range of igneous rocks, not all of which conform to this rigid definition. Carbonatites, for example, are certainly silica-deficient but are rarely alkali-rich. True (nepheline-normative) alkali basalts grade into hypersthene-normative transitional basalts without any obvious change in mineralogy. Since transitional basalts are often closely associated with alkali basalts in the field, they are traditionally regarded as alkaline. It is now usual practice to define alkaline igneous rocks simply in terms of their alkali ( N a 2 0 + K20) and silica contents (see, for example, Le Bas et al. 1986). We have not attempted to review the classification of alkaline igneous rocks in this volume as this has been dealt with elsewhere (e.g. Sorensen 1974; Streckeisen 1967, 1980). The only alkaline rocks not covered in previous reviews are those hydrous mafic to ultramafic hypabyssal rocks known as the lamprophyres. The present volume includes three papers on this group. A comprehensive overview of lamprophyres is given by Rock and of the sub-group of lamproites by Bergman. The relationship between lamproites and kimberlites (which arguably belong to the lamprophyres) is discussed by Dawson. Volumetrically, alkaline rocks account for less than one per cent of all igneous rocks. Despite this, their remarkable mineralogical diversity has brought them repeatedly to the attention of petrologists and mineralogists, with the result that alkaline rocks account for about half of all igneous rock names. Sorensen (1974) lists no fewer than 400 alkaline rock types. This diversity springs largely from an abundance of alkalis and deficiency in silica which together generate a large number of mineral species not stable in more silica-rich, alkali-poor magmas. However, a large part of the attention given to alkaline rocks is due to their characteristic high concentrations of incompatible or large-ion lithophile elements (LILE). These are often of more than academic interest as most of the world's resources of niobium, tantalum and the rare-earth elements are found in or around alkaline igneous rock bodies. The economic importance of alkaline igneous rocks is further enhanced by their association with economic deposits of apatite (Kogarko) and with diamonds (Dawson;
Bergman). Evidence for continental alkaline magmatism can be found as far back as the late Archaean. For example, biotites from the Poohbah Lake syenite in north-western Ontario have been dated at around 2.7 Ga (Mitchell 1976) and a similar age has been reported by Larsen et al. (1983) for the Tupertalik carbonatite in western Greenland. At the present time, alkaline magmas are erupted in all tectonic environments with the possible exception of mid-ocean ridges. Even here, though, mildly alkaline lavas are sometimes erupted from off-axis volcanoes, as in the Vestmann Islands of Iceland. Alkaline igneous rocks are found on all the continents and on islands in all the ocean basins. Their occurrences may be classified on the basis of tectonic setting into three categories; continental rift valley From: FITTON, J. G. & UPTON, B. G. J. (eds), 1987, Alkaline Igneous Rocks, Geological Society Special Publication No. 30, pp. ix-xiv.
ix
x
Introduction
magmatism, oceanic and continental intraplate magmatism without clear tectonic control, and alkaline magmatism related to subduction processes. In practice, however, this classification is not always easy to apply. Continental rift valleys provide, volumetrically, the most important occurrences of alkaline igneous rocks although continental rifting is not always accompanied by magmatism. The best known example (arguably the type example) is the East African Rift which, in the course of its long history, has yielded almost the entire spectrum of alkaline magmas. Three papers in this volume are devoted to various aspects of magmatism in the East African Rift. Baker reviews its magmatic associations with respect to tectonic development and discusses the origin of the magmas, particularly in relation to the problems presented by the strongly bimodal distribution of basic and salic lava compositions. He concludes that the salic magmas evolved from basic parental magmas by processes of crystal fractionation (cf. Bailey). Macdonald focusses attention on the peralkaline silicic central volcanoes of Kenya and also favours an origin by crystal fractionation for most of the evolved magmas. There is little evidence for contamination of the evolving magmas with ancient continental crust except in the case of the Naivasha comendites. It is not always possible to demonstrate a genetic link between basic and evolved magmas in the East African Rift, however. The Chilwa alkaline province in Malawi, at the southern end of the rift, is an essentially intrusive province in which salic rocks predominate. The scarcity of basic rocks in this part of the rift has led Woolley & Jones to suggest that the evolved magmas were produced directly by melting of metasomatised mantle and lower crust. Insight into the processes occurring at depth beneath rift valleys may be gained by studying ancient and deeply eroded examples. The Proterozoic Gardar province in SouthWest Greenland is probably the best studied of these and is reviewed by Upton & Emeleus. One of the most striking features of the province is the presence of giant dykes, up to 800 metres wide. These are dominantly basic but in places show in situ differentiation into more salic rocks. Salic magma generated in the wider portions of these dykes migrated upwards and may ultimately have accumulated to produce central complexes in which basic magma was subordinate or absent (e.g. the Ilimaussaq intrusion, Larsen & Serensen). The giant dykes, therefore, play a crucial role in understanding the relationship between basic and salic magmas in this and possibly other rift systems. The separation of continents to form ocean basins must always be preceded by a phase of continental rifting, leaving volcanic and intrusive complexes stranded along passive continental margins. The vigorous magmatism which accompanies continental separation is generally tholeiitic in character as, for example, in the Karoo and Deccan flood basalt provinces. Alkaline magmas, however, may be emplaced along the trailing continental margin during the waning phase of magmatism, long after the spreading centre is established off-shore. The Tertiary volcanic rocks exposed along the east coast of Greenland (described by Nielsen) provide an excellent example of such an alkaline province. A second major occurrence of alkaline igneous rocks is provided by intraplate magmatic provinces whose activity and siting are not subject to any obvious tectonic control. In the ocean basins such magmatism manifests itself as ocean islands which are sometimes aligned in chains with ages increasing away from the active centres, as in the Hawaiian islands. In these cases it is possible to relate the magmatism to convective plumes within the asthenosphere. The Hawaiian islands (reviewed by Clague) show clearly defined
Introduction
xi
magmatic cycles starting with alkaline magmas (represented by Loihi seamount), passing through a voluminous tholeiitic shield-building stage and returning to alkaline magmatism during the waning phases of individual volcanic centres. These cycles seem to result from movement of the oceanic plate over a rising plume of partially molten asthenosphere in which the degree of melting increases towards the centre. They may be typical of ocean islands in general and are broadly analogous to similar cycles seen in flood basalt provinces preserved on passive continental margins. Continental intraplate magmatism may also show age progressions as in the NigerNigeria province (Bowden et al.). These progressions are, however, very rare and not so clearly defined as in ocean island chains, probably because the continental lithosphere is thicker and less easily penetrated than oceanic lithosphere. The Monteregian Hills and White Mountain provinces of eastern North America (Eby), for example, show no obvious progression but their seaward extension, the New England seamounts, show a regular decrease in age eastwards. Reviews of other continental intraplate alkaline provinces are given by Fletcher & Beddoe-Stevens (Velasco province, Bolivia) and Kogarko (Kola Peninsula). The Cameroon line in West Africa (Fitton) includes both continental and oceanic alkaline volcanic centres. None of these examples shows any clear progression of ages. Some continental provinces undergo repeated alkaline magmatism in one place over long periods. For example, the Kola Peninsula (Kogarko) was the site of alkaline magmatism in the mid Proterozoic and again in the Devonian. Such examples could be the result of coincidence but are more likely the result of the repeated exploitation of zones of weakness in the lithosphere. Destructive plate boundaries provide the third tectonic setting in which alkaline igneous rocks may occur. During the life of a subduction zone the characteristic calc-alkaline magmas tend to become more potassic with time and may give way to volcanic rocks of the shoshonitic association, some members of which may contain leucite. A discussion of subduction zone processes is beyond the scope of this volume and the reader is referred to the reviews of Gill (1981) and Ewart (1982). There are, however, two circumstances under which subduction processes can lead to the generation of more 'normal' alkaline magmas. Once the descending slab has become dehydrated at depth it loses its capacity to stimulate the generation of calc-alkaline magmas but it can still cause melting in the overlying asthenosphere. This can lead to the production of alkaline magmas from the mantle above the deepest parts of subduction zones. One such example of alkaline magmas erupted under a compressive regime is provided by the Trans-Pecos province of west Texas (Barker). The alkaline rocks in this area grade south-westwards into the calc-alkaline rocks of the Sierra Madre Occidental in Mexico and Barker relates both suites to subduction of the Farallon Plate. After the cessation of subduction, relaxation of the former compressive regime often results in extension and the generation of alkaline magmas. The resulting switch from subduction-related calc-alkaline to extensional alkaline magmatism appears to be a common phenomenon. It occurred, for example, in the western U.S.A. about 17 Ma ago and in parts of Africa and Arabia at the end of the Pan-African metamorphic episode during the late Precambrian. An example from the Pan-African belt in Mali is discussed by Liegeois & Black. The origin of alkaline magmas has attracted a great deal of interest among igneous petrologists over the last ten years or so. This interest has been stimulated by two important
xii
Introduction
and characteristic features of alkaline rocks. Firstly, they contain high concentrations of LILE and yet isotopic evidence suggests that their parental magmas had a mantle source which had been depleted in these elements for a long time. Thus alkaline igneous rocks may provide useful information about enrichment and/or melting processes in the mantle. Secondly, many mafic alkaline volcanic rocks contain xenoliths inferred to have originated within the mantle (Menzies). These are often enriched in LILE when compared with concentrations expected of chondritic mantle material and sometimes contain amphiboles and micas of metasomatic origin (Bailey). Such clear evidence for the existence of metasomatically enriched mantle, coupled with the problem of extracting LILE-rich magmas from LILE-poor mantle, has led to hypotheses involving mantle metasomatism as a precursor to alkaline magmatism. These hypotheses, reviewed by Bailey, have gained popularity over the past fifteen years and are invoked by several contributors to this volume. Mantle metasomatism neatly explains many of the features of alkaline magmatism. For example, the frequent association of alkaline magmatism with areas of large-scale regional uplift is consistent with the relatively low density of metasomatized mantle. An essential feature of all models involving a metasomatized mantle source for alkaline magmas is that this source must lie in the lithosphere. This is the only part of the mantle where enriched material can remain in one place for long periods without being swept away by convection. The lithospheric mantle beneath the continents is likely to be chemically and isotopically different from that beneath the oceans. Continental lithospheric mantle is old and will have had as complex a metamorphic and magmatic history as the overlying crust. Oceanic lithosphere, on the other hand, is relatively young and probably depleted in LILE. These differences should be reflected in the compositions of continental and oceanic alkaline rocks. However, alkali basalts erupted in continental and oceanic settings are generally identical both chemically (Fitton) and isotopically (Menzies). Since enriched mantle xenoliths are only commonly found in continental regions it follows that the enriched lithospheric mantle represented by these xenoliths is not the source of most continental and oceanic alkaline magmas. An asthenospheric source is therefore implied. This is not to say that enriched continental lithospheric mantle is never involved in the generation of alkaline magmas. There is good evidence (e.g. Edgar) that pockets of ancient enriched mantle beneath cratonic regions provide the source for LILE-rich mafic and ultramafic alkaline rocks such as micaceous kimberlite (Dawson), and lamproite and other potassic igneous rocks (Bergman). It is significant that these rock types are exclusively continental. More extensive melting may involve the continental lithosphere mantle in the production of less exotic rock types such as flood tholeiite and mildly alkaline basalt. Upton & Emeleus, for example, argue for a lithospheric mantle source for the Gardar alkaline magmas. If most alkaline magmas have an ultimate source in the asthenosphere then they must share this source with unequivocally asthenosphere-derived rocks such as mid-ocean ridge basalt (MORB). The consistent isotopic differences between MORB and alkali basalts (and indeed all intraplate basalts) requires that the asthenosphere be heterogeneous. This heterogeneity may result from the entrainment of lower mantle material in deep mantle plumes as suggested for the Hawaiian island chain (Clague). The entire convecting upper mantle may also be heterogeneous on a small scale and alkaline magmas may be generated by the selective melting of LILE-enriched streaks while more extensive melting produces MORB (Fitton). Geochemical studies on ocean island basalts from the South Atlantic
Introduction
xiii
(Weaver et a/.) suggest that one component of these enriched streaks is provided by subducted ocean-floor sediment. Derivation of LILE-rich magmas from an asthenospheric source depleted in these elements requires either very small degress of partial melting or extensive crystal fractionation. There can be no doubt that many alkaline rocks are the products of extensive low pressure crystal fractionation but this cannot be true of those alkaline rocks which host mantle xenoliths. Even the most magnesian alkaline rocks, which must represent nearprimary magmas, are rich in LILE. Small-degree partial melting ( < l ~ ) is therefore required to produce such magmas. McKenzie (1985) has recently shown that the extraction of melt fractions as small as 0.2~ is not only physically possible but inevitable where the melt viscosity is low, as it probably is in the case of alkaline magmas. Experimental studies on alkaline rocks and synthetic analogues (reviewed by Edgar) provide useful constraints on the feasibility of fractional crystallization and partial melting models and on the temperatures and pressures involved. Many lines of evidence suggest that volatile components form a significant part of alkaline magmas. The development of extensive zones of metasomatised country rock (fenite) around alkaline plutons, the abundance of chlorine and fluorine in some alkaline igneous rocks, and the frequently explosive eruption of alkaline magma all point to high concentrations of volatiles. These volatile components play an important role in the evolution of alkaline magmas and yet relatively little is known about them. Constraints on their composition have been provided by fluid inclusion studies (Harris & Sheppard) and by thermodynamic considerations (Kogarko). The effects of volatile components on the evolution of alkaline magmas can be seen clearly in both intrusive and extrusive rocks. Larsen & Serensen, for example, discuss the crystallization history of the Ilimaussaq intrusion in South-West Greenland and show how the upward migration of low-density, low-viscosity volatile-rich magma delayed crystallization under the roof of the intrusion. Silicic alkaline pyroclastic deposits around central volcanoes in Kenya often show striking variations in the abundance of some incompatible elements within a single vertical section, implying compositional zonation in the magma chamber before eruption. Macdonald shows that these variations are too large to be accounted for by crystal fractionation alone and suggests that some elements have been transported to the magma chamber roof zones as complex ions in a volatile phase. Carbonatites provide perhaps the best illustration of all of the influence of volatile components on the origin and evolution of alkaline rocks. There is now a consensus that their parental magmas originate by the separation of an immiscible carbonate liquid phase from a CO2-saturated nephelinite or phonolite magma. There is, however, some disagreement over the nature and subsequent evolution of this parental carbonate magma. Le Bas argues that the parental magma is rich in alkalis and similar in composition to the natrocarbonatite lavas erupted from Oldoinyo Lengai. This magma evolves at low pressure towards the more common calcite carbonatite (s6vite) by loss of alkalis to the surrounding country rocks which are metasomatized (fenitized) as a result. Twyman & Gittins offer an alternative scheme in which s6vite magmas are parental and natrocarbonatite magmas are derived from them by crystal fractionation. Most petrologists now believe that evolved alkaline magmas are produced by the fractional crystallization of basic magma. The more highly undersaturated parental magmas represented by basanite and nephelinite will produce undersaturated derivatives
xiv
Introduction
such as phonolite and foyaite. Mildly alkaline and transitional basalt m a g m a is likely to produce trachyte and, with extreme fractionation, alkali rhyolite. The production of peralkaline acid rocks by crystal fractionation alone is likely to be an inefficient process. The operation of this process, however, is clearly demonstrated by their association with transitional basalt on some ocean islands, such as Ascension (Harris & Sheppard). Peralkaline acid rocks only occur in a b u n d a n c e in continental environments, however, and here there is often good evidence that crustal contamination has a c c o m p a n i e d crystal fractionation. Well-documented examples of the operation of crustal c o n t a m i n a t i o n in the evolution of alkaline m a g m a s are presented by several of the contributors to this volume. Downes, for example, shows that the assimilation of lower crustal granulite has affected the evolution of alkaline m a g m a s in the F r e n c h Massif Central and uses isotope data to estimate the extent of this contamination. Other examples are presented by Bowden et al. ( N i g e r - N i g e r i a granite ring complexes), Eby (Monteregian Hills and White Mountain provinces, N o r t h America), Fitton (Cameroon line, West Africa) and Fletcher & BeddoeStevens (Velasco province, Bolivia). Other authors propose the derivation of evolved alkaline m a g m a s directly from m e t a s o m a t i z e d mantle or lower crust (Bailey; Woolley &
Jones). Our understanding of the origin and evolution of alkaline m a g m a s has come a long way since the publication of Sorensen's book in 1974, largely through the acquisition of a far larger geochemical and isotopic data base. The contributions to this volume review the current state of this understanding. Emphasis has shifted from crustal to mantle processes with the recognition of mantle m e t a s o m a t i s m and its possible role as a precursor to alkaline magmatism. More recently, though, there has been a swing towards the opposite view, that mantle m e t a s o m a t i s m is c a u s e d b y alkaline m a g m a t i s m . Theoretical and experimental studies on the migration and segregation of small-degree melts seem destined to accelerate this swing. Despite these advances, however, m a n y mysteries remain unsolved and alkaline rocks will still provide a fruitful field of research for m a n y years to come, yielding further insights into the nature of mantle processes and the evolution of magmas.
References EWART, A. 1982. The mineralogy and petrology of Tertiary-Recent orogenic volcanic rocks: with special reference to the andesitic-basaltic compositional range. In Thorpe, R. S. (ed.) Andesites. Pp. 25-87. John Wiley & Sons, London. GILL,J. B. 1981. Orogenic Andesites and Plate Tectonics, 390 pp. Springer-Verlag, Berlin. LARSEN,L. M., REX,D. C. & SECHER,K. 1983. The age of carbonatites, kimberlites and lamprophyres from southern west Greenland: recurrent alkaline magmatism during 2500 million years. Lithos 16, 215-21. LE BAS,M. J., LE MAITRE, R. W., STRECKEISEN,A. & ZANETTIN, B. 1986. A chemical classification of
volcanic rocks based on the total alkali--silica diagram. J. Petrol. 27, 745-50. MCKENZIE, D. 1985. The extraction of magma from the crust and the mantle. Earth planet. Sci. Lett. 74, 81-91. MITCHELL,R. H. 1976. Potassium-argon geochronology of the Poohbah Lake alkaline complex, northwestern Ontario. Can. J. Earth Sci. 13, 1456-9. SORENSEN,H. (ed.) 1974. The Alkaline Rocks. 622 pp. John Wiley & Sons, London. STRECKEISEN,A. 1967. Classification and nomenclature of igneous rocks. N. Jb. Miner. Abh. 107, 144-240. - - , 1980. Classificationand nomenclature of volcanic rocks, lamprophyres, carbonatites and melilitic rocks. Geol. Rundschau, 69, 194-207.
J. G. FITTON& B. G. J. UPTON,Grant Institute of Geology, University of Edinburgh, West Mains Road, Edinburgh EH9 3JW, U.K.
Introduction J. G. Fitton & B. G. J. Upton Alkaline igneous rocks may be defined as those which have higher concentrations of alkalis than can be accommodated in feldspars alone, the excess appearing as feldspathoids, sodic pyroxenes, sodic amphiboles and other alkali-rich phases. These rocks are, therefore, deficient in silica and/or alumina with respect to alkalis and will have nepheline and/or acmite in their norms. In practice the term 'alkaline' is used to encompass a wide range of igneous rocks, not all of which conform to this rigid definition. Carbonatites, for example, are certainly silica-deficient but are rarely alkali-rich. True (nepheline-normative) alkali basalts grade into hypersthene-normative transitional basalts without any obvious change in mineralogy. Since transitional basalts are often closely associated with alkali basalts in the field, they are traditionally regarded as alkaline. It is now usual practice to define alkaline igneous rocks simply in terms of their alkali ( N a 2 0 + K20) and silica contents (see, for example, Le Bas et al. 1986). We have not attempted to review the classification of alkaline igneous rocks in this volume as this has been dealt with elsewhere (e.g. Sorensen 1974; Streckeisen 1967, 1980). The only alkaline rocks not covered in previous reviews are those hydrous mafic to ultramafic hypabyssal rocks known as the lamprophyres. The present volume includes three papers on this group. A comprehensive overview of lamprophyres is given by Rock and of the sub-group of lamproites by Bergman. The relationship between lamproites and kimberlites (which arguably belong to the lamprophyres) is discussed by Dawson. Volumetrically, alkaline rocks account for less than one per cent of all igneous rocks. Despite this, their remarkable mineralogical diversity has brought them repeatedly to the attention of petrologists and mineralogists, with the result that alkaline rocks account for about half of all igneous rock names. Sorensen (1974) lists no fewer than 400 alkaline rock types. This diversity springs largely from an abundance of alkalis and deficiency in silica which together generate a large number of mineral species not stable in more silica-rich, alkali-poor magmas. However, a large part of the attention given to alkaline rocks is due to their characteristic high concentrations of incompatible or large-ion lithophile elements (LILE). These are often of more than academic interest as most of the world's resources of niobium, tantalum and the rare-earth elements are found in or around alkaline igneous rock bodies. The economic importance of alkaline igneous rocks is further enhanced by their association with economic deposits of apatite (Kogarko) and with diamonds (Dawson;
Bergman). Evidence for continental alkaline magmatism can be found as far back as the late Archaean. For example, biotites from the Poohbah Lake syenite in north-western Ontario have been dated at around 2.7 Ga (Mitchell 1976) and a similar age has been reported by Larsen et al. (1983) for the Tupertalik carbonatite in western Greenland. At the present time, alkaline magmas are erupted in all tectonic environments with the possible exception of mid-ocean ridges. Even here, though, mildly alkaline lavas are sometimes erupted from off-axis volcanoes, as in the Vestmann Islands of Iceland. Alkaline igneous rocks are found on all the continents and on islands in all the ocean basins. Their occurrences may be classified on the basis of tectonic setting into three categories; continental rift valley From: FITTON, J. G. & UPTON, B. G. J. (eds), 1987, Alkaline Igneous Rocks, Geological Society Special Publication No. 30, pp. ix-xiv.
ix
x
Introduction
magmatism, oceanic and continental intraplate magmatism without clear tectonic control, and alkaline magmatism related to subduction processes. In practice, however, this classification is not always easy to apply. Continental rift valleys provide, volumetrically, the most important occurrences of alkaline igneous rocks although continental rifting is not always accompanied by magmatism. The best known example (arguably the type example) is the East African Rift which, in the course of its long history, has yielded almost the entire spectrum of alkaline magmas. Three papers in this volume are devoted to various aspects of magmatism in the East African Rift. Baker reviews its magmatic associations with respect to tectonic development and discusses the origin of the magmas, particularly in relation to the problems presented by the strongly bimodal distribution of basic and salic lava compositions. He concludes that the salic magmas evolved from basic parental magmas by processes of crystal fractionation (cf. Bailey). Macdonald focusses attention on the peralkaline silicic central volcanoes of Kenya and also favours an origin by crystal fractionation for most of the evolved magmas. There is little evidence for contamination of the evolving magmas with ancient continental crust except in the case of the Naivasha comendites. It is not always possible to demonstrate a genetic link between basic and evolved magmas in the East African Rift, however. The Chilwa alkaline province in Malawi, at the southern end of the rift, is an essentially intrusive province in which salic rocks predominate. The scarcity of basic rocks in this part of the rift has led Woolley & Jones to suggest that the evolved magmas were produced directly by melting of metasomatised mantle and lower crust. Insight into the processes occurring at depth beneath rift valleys may be gained by studying ancient and deeply eroded examples. The Proterozoic Gardar province in SouthWest Greenland is probably the best studied of these and is reviewed by Upton & Emeleus. One of the most striking features of the province is the presence of giant dykes, up to 800 metres wide. These are dominantly basic but in places show in situ differentiation into more salic rocks. Salic magma generated in the wider portions of these dykes migrated upwards and may ultimately have accumulated to produce central complexes in which basic magma was subordinate or absent (e.g. the Ilimaussaq intrusion, Larsen & Serensen). The giant dykes, therefore, play a crucial role in understanding the relationship between basic and salic magmas in this and possibly other rift systems. The separation of continents to form ocean basins must always be preceded by a phase of continental rifting, leaving volcanic and intrusive complexes stranded along passive continental margins. The vigorous magmatism which accompanies continental separation is generally tholeiitic in character as, for example, in the Karoo and Deccan flood basalt provinces. Alkaline magmas, however, may be emplaced along the trailing continental margin during the waning phase of magmatism, long after the spreading centre is established off-shore. The Tertiary volcanic rocks exposed along the east coast of Greenland (described by Nielsen) provide an excellent example of such an alkaline province. A second major occurrence of alkaline igneous rocks is provided by intraplate magmatic provinces whose activity and siting are not subject to any obvious tectonic control. In the ocean basins such magmatism manifests itself as ocean islands which are sometimes aligned in chains with ages increasing away from the active centres, as in the Hawaiian islands. In these cases it is possible to relate the magmatism to convective plumes within the asthenosphere. The Hawaiian islands (reviewed by Clague) show clearly defined
Introduction
xi
magmatic cycles starting with alkaline magmas (represented by Loihi seamount), passing through a voluminous tholeiitic shield-building stage and returning to alkaline magmatism during the waning phases of individual volcanic centres. These cycles seem to result from movement of the oceanic plate over a rising plume of partially molten asthenosphere in which the degree of melting increases towards the centre. They may be typical of ocean islands in general and are broadly analogous to similar cycles seen in flood basalt provinces preserved on passive continental margins. Continental intraplate magmatism may also show age progressions as in the NigerNigeria province (Bowden et al.). These progressions are, however, very rare and not so clearly defined as in ocean island chains, probably because the continental lithosphere is thicker and less easily penetrated than oceanic lithosphere. The Monteregian Hills and White Mountain provinces of eastern North America (Eby), for example, show no obvious progression but their seaward extension, the New England seamounts, show a regular decrease in age eastwards. Reviews of other continental intraplate alkaline provinces are given by Fletcher & Beddoe-Stevens (Velasco province, Bolivia) and Kogarko (Kola Peninsula). The Cameroon line in West Africa (Fitton) includes both continental and oceanic alkaline volcanic centres. None of these examples shows any clear progression of ages. Some continental provinces undergo repeated alkaline magmatism in one place over long periods. For example, the Kola Peninsula (Kogarko) was the site of alkaline magmatism in the mid Proterozoic and again in the Devonian. Such examples could be the result of coincidence but are more likely the result of the repeated exploitation of zones of weakness in the lithosphere. Destructive plate boundaries provide the third tectonic setting in which alkaline igneous rocks may occur. During the life of a subduction zone the characteristic calc-alkaline magmas tend to become more potassic with time and may give way to volcanic rocks of the shoshonitic association, some members of which may contain leucite. A discussion of subduction zone processes is beyond the scope of this volume and the reader is referred to the reviews of Gill (1981) and Ewart (1982). There are, however, two circumstances under which subduction processes can lead to the generation of more 'normal' alkaline magmas. Once the descending slab has become dehydrated at depth it loses its capacity to stimulate the generation of calc-alkaline magmas but it can still cause melting in the overlying asthenosphere. This can lead to the production of alkaline magmas from the mantle above the deepest parts of subduction zones. One such example of alkaline magmas erupted under a compressive regime is provided by the Trans-Pecos province of west Texas (Barker). The alkaline rocks in this area grade south-westwards into the calc-alkaline rocks of the Sierra Madre Occidental in Mexico and Barker relates both suites to subduction of the Farallon Plate. After the cessation of subduction, relaxation of the former compressive regime often results in extension and the generation of alkaline magmas. The resulting switch from subduction-related calc-alkaline to extensional alkaline magmatism appears to be a common phenomenon. It occurred, for example, in the western U.S.A. about 17 Ma ago and in parts of Africa and Arabia at the end of the Pan-African metamorphic episode during the late Precambrian. An example from the Pan-African belt in Mali is discussed by Liegeois & Black. The origin of alkaline magmas has attracted a great deal of interest among igneous petrologists over the last ten years or so. This interest has been stimulated by two important
xii
Introduction
and characteristic features of alkaline rocks. Firstly, they contain high concentrations of LILE and yet isotopic evidence suggests that their parental magmas had a mantle source which had been depleted in these elements for a long time. Thus alkaline igneous rocks may provide useful information about enrichment and/or melting processes in the mantle. Secondly, many mafic alkaline volcanic rocks contain xenoliths inferred to have originated within the mantle (Menzies). These are often enriched in LILE when compared with concentrations expected of chondritic mantle material and sometimes contain amphiboles and micas of metasomatic origin (Bailey). Such clear evidence for the existence of metasomatically enriched mantle, coupled with the problem of extracting LILE-rich magmas from LILE-poor mantle, has led to hypotheses involving mantle metasomatism as a precursor to alkaline magmatism. These hypotheses, reviewed by Bailey, have gained popularity over the past fifteen years and are invoked by several contributors to this volume. Mantle metasomatism neatly explains many of the features of alkaline magmatism. For example, the frequent association of alkaline magmatism with areas of large-scale regional uplift is consistent with the relatively low density of metasomatized mantle. An essential feature of all models involving a metasomatized mantle source for alkaline magmas is that this source must lie in the lithosphere. This is the only part of the mantle where enriched material can remain in one place for long periods without being swept away by convection. The lithospheric mantle beneath the continents is likely to be chemically and isotopically different from that beneath the oceans. Continental lithospheric mantle is old and will have had as complex a metamorphic and magmatic history as the overlying crust. Oceanic lithosphere, on the other hand, is relatively young and probably depleted in LILE. These differences should be reflected in the compositions of continental and oceanic alkaline rocks. However, alkali basalts erupted in continental and oceanic settings are generally identical both chemically (Fitton) and isotopically (Menzies). Since enriched mantle xenoliths are only commonly found in continental regions it follows that the enriched lithospheric mantle represented by these xenoliths is not the source of most continental and oceanic alkaline magmas. An asthenospheric source is therefore implied. This is not to say that enriched continental lithospheric mantle is never involved in the generation of alkaline magmas. There is good evidence (e.g. Edgar) that pockets of ancient enriched mantle beneath cratonic regions provide the source for LILE-rich mafic and ultramafic alkaline rocks such as micaceous kimberlite (Dawson), and lamproite and other potassic igneous rocks (Bergman). It is significant that these rock types are exclusively continental. More extensive melting may involve the continental lithosphere mantle in the production of less exotic rock types such as flood tholeiite and mildly alkaline basalt. Upton & Emeleus, for example, argue for a lithospheric mantle source for the Gardar alkaline magmas. If most alkaline magmas have an ultimate source in the asthenosphere then they must share this source with unequivocally asthenosphere-derived rocks such as mid-ocean ridge basalt (MORB). The consistent isotopic differences between MORB and alkali basalts (and indeed all intraplate basalts) requires that the asthenosphere be heterogeneous. This heterogeneity may result from the entrainment of lower mantle material in deep mantle plumes as suggested for the Hawaiian island chain (Clague). The entire convecting upper mantle may also be heterogeneous on a small scale and alkaline magmas may be generated by the selective melting of LILE-enriched streaks while more extensive melting produces MORB (Fitton). Geochemical studies on ocean island basalts from the South Atlantic
Introduction
xiii
(Weaver et a/.) suggest that one component of these enriched streaks is provided by subducted ocean-floor sediment. Derivation of LILE-rich magmas from an asthenospheric source depleted in these elements requires either very small degress of partial melting or extensive crystal fractionation. There can be no doubt that many alkaline rocks are the products of extensive low pressure crystal fractionation but this cannot be true of those alkaline rocks which host mantle xenoliths. Even the most magnesian alkaline rocks, which must represent nearprimary magmas, are rich in LILE. Small-degree partial melting ( < l ~ ) is therefore required to produce such magmas. McKenzie (1985) has recently shown that the extraction of melt fractions as small as 0.2~ is not only physically possible but inevitable where the melt viscosity is low, as it probably is in the case of alkaline magmas. Experimental studies on alkaline rocks and synthetic analogues (reviewed by Edgar) provide useful constraints on the feasibility of fractional crystallization and partial melting models and on the temperatures and pressures involved. Many lines of evidence suggest that volatile components form a significant part of alkaline magmas. The development of extensive zones of metasomatised country rock (fenite) around alkaline plutons, the abundance of chlorine and fluorine in some alkaline igneous rocks, and the frequently explosive eruption of alkaline magma all point to high concentrations of volatiles. These volatile components play an important role in the evolution of alkaline magmas and yet relatively little is known about them. Constraints on their composition have been provided by fluid inclusion studies (Harris & Sheppard) and by thermodynamic considerations (Kogarko). The effects of volatile components on the evolution of alkaline magmas can be seen clearly in both intrusive and extrusive rocks. Larsen & Serensen, for example, discuss the crystallization history of the Ilimaussaq intrusion in South-West Greenland and show how the upward migration of low-density, low-viscosity volatile-rich magma delayed crystallization under the roof of the intrusion. Silicic alkaline pyroclastic deposits around central volcanoes in Kenya often show striking variations in the abundance of some incompatible elements within a single vertical section, implying compositional zonation in the magma chamber before eruption. Macdonald shows that these variations are too large to be accounted for by crystal fractionation alone and suggests that some elements have been transported to the magma chamber roof zones as complex ions in a volatile phase. Carbonatites provide perhaps the best illustration of all of the influence of volatile components on the origin and evolution of alkaline rocks. There is now a consensus that their parental magmas originate by the separation of an immiscible carbonate liquid phase from a CO2-saturated nephelinite or phonolite magma. There is, however, some disagreement over the nature and subsequent evolution of this parental carbonate magma. Le Bas argues that the parental magma is rich in alkalis and similar in composition to the natrocarbonatite lavas erupted from Oldoinyo Lengai. This magma evolves at low pressure towards the more common calcite carbonatite (s6vite) by loss of alkalis to the surrounding country rocks which are metasomatized (fenitized) as a result. Twyman & Gittins offer an alternative scheme in which s6vite magmas are parental and natrocarbonatite magmas are derived from them by crystal fractionation. Most petrologists now believe that evolved alkaline magmas are produced by the fractional crystallization of basic magma. The more highly undersaturated parental magmas represented by basanite and nephelinite will produce undersaturated derivatives
xiv
Introduction
such as phonolite and foyaite. Mildly alkaline and transitional basalt m a g m a is likely to produce trachyte and, with extreme fractionation, alkali rhyolite. The production of peralkaline acid rocks by crystal fractionation alone is likely to be an inefficient process. The operation of this process, however, is clearly demonstrated by their association with transitional basalt on some ocean islands, such as Ascension (Harris & Sheppard). Peralkaline acid rocks only occur in a b u n d a n c e in continental environments, however, and here there is often good evidence that crustal contamination has a c c o m p a n i e d crystal fractionation. Well-documented examples of the operation of crustal c o n t a m i n a t i o n in the evolution of alkaline m a g m a s are presented by several of the contributors to this volume. Downes, for example, shows that the assimilation of lower crustal granulite has affected the evolution of alkaline m a g m a s in the F r e n c h Massif Central and uses isotope data to estimate the extent of this contamination. Other examples are presented by Bowden et al. ( N i g e r - N i g e r i a granite ring complexes), Eby (Monteregian Hills and White Mountain provinces, N o r t h America), Fitton (Cameroon line, West Africa) and Fletcher & BeddoeStevens (Velasco province, Bolivia). Other authors propose the derivation of evolved alkaline m a g m a s directly from m e t a s o m a t i z e d mantle or lower crust (Bailey; Woolley &
Jones). Our understanding of the origin and evolution of alkaline m a g m a s has come a long way since the publication of Sorensen's book in 1974, largely through the acquisition of a far larger geochemical and isotopic data base. The contributions to this volume review the current state of this understanding. Emphasis has shifted from crustal to mantle processes with the recognition of mantle m e t a s o m a t i s m and its possible role as a precursor to alkaline magmatism. More recently, though, there has been a swing towards the opposite view, that mantle m e t a s o m a t i s m is c a u s e d b y alkaline m a g m a t i s m . Theoretical and experimental studies on the migration and segregation of small-degree melts seem destined to accelerate this swing. Despite these advances, however, m a n y mysteries remain unsolved and alkaline rocks will still provide a fruitful field of research for m a n y years to come, yielding further insights into the nature of mantle processes and the evolution of magmas.
References EWART, A. 1982. The mineralogy and petrology of Tertiary-Recent orogenic volcanic rocks: with special reference to the andesitic-basaltic compositional range. In Thorpe, R. S. (ed.) Andesites. Pp. 25-87. John Wiley & Sons, London. GILL,J. B. 1981. Orogenic Andesites and Plate Tectonics, 390 pp. Springer-Verlag, Berlin. LARSEN,L. M., REX,D. C. & SECHER,K. 1983. The age of carbonatites, kimberlites and lamprophyres from southern west Greenland: recurrent alkaline magmatism during 2500 million years. Lithos 16, 215-21. LE BAS,M. J., LE MAITRE, R. W., STRECKEISEN,A. & ZANETTIN, B. 1986. A chemical classification of
volcanic rocks based on the total alkali--silica diagram. J. Petrol. 27, 745-50. MCKENZIE, D. 1985. The extraction of magma from the crust and the mantle. Earth planet. Sci. Lett. 74, 81-91. MITCHELL,R. H. 1976. Potassium-argon geochronology of the Poohbah Lake alkaline complex, northwestern Ontario. Can. J. Earth Sci. 13, 1456-9. SORENSEN,H. (ed.) 1974. The Alkaline Rocks. 622 pp. John Wiley & Sons, London. STRECKEISEN,A. 1967. Classification and nomenclature of igneous rocks. N. Jb. Miner. Abh. 107, 144-240. - - , 1980. Classificationand nomenclature of volcanic rocks, lamprophyres, carbonatites and melilitic rocks. Geol. Rundschau, 69, 194-207.
J. G. FITTON& B. G. J. UPTON,Grant Institute of Geology, University of Edinburgh, West Mains Road, Edinburgh EH9 3JW, U.K.
Mantle metasomatism
perspective and prospect
D. K. Bailey S U M M A R Y : Mantle replenishment in lithophile elements has been discerned in the patterns of trace elements and isotopes in lavas. One replenishment process is identified as metasomatic replacement, seen in ultramafic xenoliths brought up in high-velocity alkaline eruptions. Thus alkaline magmatism provides the best primafacie evidence of metasomatism and open-system conditions in the upper mantle. The list of added lithophile elements includes the following: H, C, F, Na, A1, P, S, C1, K, Ca, Ti, Fe, Rb, Y, Zr, Nb, Ba and rare earths. Some metasomatism may be due to wall-rock alteration near magma bodies, but the evidence for metasomatism prior to melting opens the possibility that the process is a precursor to alkaline magmatism, giving the necessary source enrichment in lithophile elements. In some igneous provinces the metasomafism is widespread, intensive and pervasive; in others it appears as veining of variable intensity. Metasomatism as a largescale process is best indicated by the widespread distribution of alkaline magmatism in space and time: volatile flux through the lithosphere would then be the necessary precursor of metasomatism and magmatism. Volatile activity, metasomatism and melt enrichment clearly widen the scope for mafic magma generation in the mantle, but some long-standing problems of the alkaline associations (and indeed the calc-alkaline) also call for re-examination in terms of volatile activity in the lithosphere mantle. These include the diversity of magmas, the generation of large felsic volumes, and composition gaps in magma series. Experiments show that felsic minerals are stable to 30 kb, indicating the possibility of felsic-melt generation in the upper part of the mantle. A combination of volatile flux and melt percolation along geotherms that intersect the solidus at depths of less than 80 km would lead to enrichmentand metasomatism, providing distinct mantle sources for felsic magmas. Initial (or residual) melts from such a region, as distinct from those from greater depths, would be constrained by equilibria involving felsic minerals. Thus an igneous cycle could generate a bimodal association, with felsic melts forming in the upper regime and the mafic melts originating at depths below the range of felsic-mineral stabilities. Such a magma system is consistent with the observed eruptive characteristics, explains the typical ultramafic nodule and megacryst suites in the alkali olivine-basalt association and is free of difficulties with relative volumes of melts, with eruption timing and with rapid changes in erupted compositions.
Perspective From its Greek origins the term metasomatism should imply a 'change of body' or a change of substance. It must indicate a chemical change in a pre-existing rock or mineral, and traditionally has been used to make the distinction from metamorphism ('change of form') indicating reconstitution without chemical introduction. Unfortunately, early definitions of metasomatism are not rigorous, but common usage applies to cases where material has been transferred through a vapour or fluid without melting. More recently, in reference to the mantle, the usage has widened considerably, sometimes referring to simple melt infiltration, and even unspecified enrichment processes of a pre-existing mantle composition. Obviously the latter usages are not justified but there is a boundary problem. W h e n melt is introduced into a rock it may react with its surroundings and metasomatism may correctly describe the alteration process in the wall-rocks;
this borderline condition is perhaps part of the reason for some of the current ambiguity. There is at present a need for general agreement about terminology and it would certainly help to retain scientific precision if 'metasomatism' could be restricted to those cases where there is petrographic evidence of replacement of a pre-existing rock. Certainly the term should not be used to describe simple melt injection, producing a hybrid mixed rock, nor when there is no petrographic evidence that there has been replacement of an earlier mineralogy. In other words, the term should be applied only when there is unequivocal evidence of the previous substance--otherwise, to say that a rock has changed its chemistry is supposition. If there is a case for chemical introduction but the process is unclear or unknown, it would be better, and more straightforward, to use the term enrichment. Evidence for metasomatism in the mantle has
From: FITTON, J. G. & UPTON, B. G. J. (eds), 1987, Alkaline Igneous Rocks, Geological Society Special Publication No. 30, pp. 1-13.
2
D.K.
been steadily accumulating during the past 10 years (Harte et al. 1975; Lloyd & Bailey 1975) and the topic has provoked such interest that some surveys and reviews have already appeared (e.g. Boettcher & O'Neil 1980, Introduction and tabulation; Bailey 1982). The cited examples are relatively succinct and in some ways complementary, and may be taken as constituting a comprehensive starting point for the following discussion which aims to sketch an overall perspective to the subject. Occurrence and distribution Mantle samples showing evidence of metasomatism are provided in ultramafic nodules carried to the surface by high-speed volcanic eruptions. Some peridotite massifs (e.g. Lherz) contain veins of less refractory minerals, while others also contain evenly distributed minerals such as phlogopite and amphibole (e.g. Finero). These are sometimes cited as examples of 'enriched' and even metasomatized mantle but the evidence that such minerals have been introduced from an external source while the peridotite was part of the mantle has still to be established. Even in volcanically derived mantle fragments there may still be problems in distinguishing reaction products produced after incorporation in the magma; various criteria can be applied, however, and it is clear that in many cases minerals were introduced into the peridotite prior to eruption (Bailey 1982). Mantle nodules showing signs of metasomatism are brought to the surface exclusively by alkaline activity, especially in ultramafic and mafic eruptions. These range from kimberlites, which are characteristically fragmental, through to magmas in the alkali olivine-basalt association (basanites through to mugearites) which may occasionally carry ultramafic nodules. Hence, in terms of distribution, samples may be found in any parts of the stable plates, both oceanic and continental, where there has been alkaline igneous activity; such activity is usually associated with uplift and dislocation, in other words with some kind of lesion in the lithosphere. Compositions Metasomatism is recognized when there is replacement of a pre-existing peridotite mineralogy, new minerals being characterized by their content of mobile and volatile elements, as listed in Table 1. Introduction of the listed minerals into garnet or spinel peridotite (generally taken as representing average mantle composition) signifies mantle enrichment in lithophile ele-
Bailey TABLE 1. Lithophile-element-bearing m&erals in peridotite xenoliths which either may show metasomatic replacement of pre-existing peridotite mineralogy or are associated with metasomatism a (Major) b (Minor)
biotite, amphibole, clinopyroxene, carbonate phosphate, titanates, oxides, sulphides
Possible candidates for higher-level metasomatic minerals c (< 25 kb) feldspathoids, alkali feldspars ments. Most descriptions of mantle metasomatism deduce the introduction of some or all of the following elements: H, C, F, Na, A1, P, S, C1, K, Ca, Ti, Fe, Rb, Y, Zr, Nb, Ba, rare earths. If metasomatism were wholly a replacement process then there should be subtraction of equivalent material, balancing the introduction of the above elements. As yet this aspect has received little attention and this may be because the fabric of the rocks showing metasomatism is characteristically veined, with metasomatism proceeding along grain boundaries and fissures: hence the introduced material could be largely accommodated by an increase in rock volume, with redistribution rather than removal of the displaced elements. A parallel case may be seen in alkaline metasomatism in the crust (fenitization) which is characteristically marked by the introduction of new minerals along a close network of cracks. Timing In many cases the indications are that the metasomatic introduction of new minerals, although it must have preceded eruption, is part of the cycle of igneous activity, and there is isotopic harmony between the nodules and the magmas. In some instances there is evidence of an earlier metasomatic event (Erlank & Shimizu 1977; Menzies & Murthy 1980a) and there are cases of possible complex metasomatism and/or enrichment dating as far back as 3 Ga (Erlank et al. 1980; Menzies & Murthy 1980b). No doubt the picture will become still more complex as new data become available; this may be expected because alkaline activity has often been repeated through the same segment of lithosphere (Bailey 1977) and even if metasomatism were nothing more than a minor part of igneous activity there should be samples showing complex histories. Conditions of metasomatism Using a combination of experimentally determined mineral stabilities and solidi it is possible
Mantle metasomatism
3
to put limits on the pressures and temperatures of metasomatism, and hence on the conditions in the source mantle. Figs 1 and 2 summarize the conditions for stability of phlogopite, amphibole, clinopyroxene and felsic minerals in the mantle, and by reference to Fig. 4 the relative stability of carbonates can readily be envisaged. Obviously the diagrams must be generalizations because the individual mineral stabilities will depend on bulk composition (mineral and rock), coexistence with other metasomatic minerals, and the presence or absence of vapour, fluid or melt. In spite of these reservations the diagrams offer useful limits and indicate the following relationships. 1 Amphibole is stable near the peridotite solidus only in the upper part of the mantle, its lower boundary corresponding approximately to that of spinel and showing broad equivalence to the lower stability boundaries of feldspars and feldspathoids and the upper stability boundary of carbonate (see Figs 1, 2 and 4). 2 Phlogopite stability extends to higher temperatures and greater depths than amphibole, its
FIG. 2. P-T relationships of mineral stabilities and solidi, and geotherms. Solidi: as in Fig. 1 with the addition of WE, the solidus in the system KA1SiO4MgO-SiO2-H20-CO2 (Wendlandt & Eggler 1980). WE is used to provide a measure of phlogopite stability in peridotite mantle; its extrapolation to intersect KS at point P indicates a depth limit for phlogopite in the presence of melt (this is similar to the boundary given by Wyllie (1979)). Geotherms: as in Fig. 1. Mineral stabilities: C, coesite: Ks, kalsilite; AB, albite + nepheline; (AB), albite (jadeite + quartz); OR, K-feldspar; SC, solvus crest for alkali feldspars. stability limit at the solidus approximately coinciding with the inflexion region in some kimberlite geotherms (about 180km). Phlogopitebearing nodules or lavas from greater depths would therefore seem to be ruled out (Bailey 1986). 3 Clinopyroxene stability can extend up to the vapour-absent solidus and this mineral is arguably the most important but least appreciated metasomatic mineral.
FIG. 1. P-Trelationships of solidi, and geotherms. Solidi : PSD, peridotite vapour-absent solidus; KS, kimberlite solidus (Eggler & Wendlandt 1978); OE, peridotite solidus in the presence of H20 and CO2 (limited) (Olafsson & Eggler 1983). For convenience the stability region of spinel peridotite, with or without amphibole, is shown here rather than in Fig. 2. D is the diamond stability boundary (Kennedy & Kennedy 1976). Geotherms: S and O, shield and ocean (Clark & Ringwood 1964); 180, oceanic lithosphere 180 Ma old (Sclater et al. 1980); 30, oceanic lithosphere 30 Ma old (Oldenburg 1981). If geotherms should converge more rapidly in the depth range around 200 km as proposed by Tozer (1967) the geometry of the melt systems would change but the principles would remain the same.
From the above it may be deduced that carbonate and clinopyroxene may appear in any metasomatic regime, but the geological and experimental evidence suggests that high-pressure, low-temperature conditions (e.g. kimberlitic) favour carbonates while low-pressure hightemperature conditions (basanitic) favour clinopyroxene (see Fig. 5). Consideration of the magmas and their nodule suites also points to the fact that ultramafic alkaline magmas such as kimberlites, melilitites and nephelinites are characterized by phlogopite (in melts and xenoliths) while other magmas such as some nephelinites, but especially basanites, are characterized by xenoliths containing amphibole. Basanites in particular are likely to carry ultramafic nodules containing both spinel and amphibole, which
4
D.K. Bailey
may be an indicator of a generally shallower source depth for the nodules compared with ultramafic melts.
Outstanding issues While the reality of mantle metasomatism seems to have found wide acceptance, some differences may be perceived concerning such questions as the following: 1 Which comes first, the magmatism or the metasomatism? 2 Is metasomatism always a local phenomenon (in the vicinity of alkaline intrusions) or can it develop regionally? 3 What is the source of materials erupted at the surface, e.g. are they from the lithosphere or the asthenosphere ? The last question seems irresolvable until the nature of the asthenosphere can be unequivocally defined. At the moment, the ascription of a rock to an asthenosphere source (on the basis of selected aspects of its trace-element or isotope chemistry) is a convoluted way of saying that it has some attributes of mid-ocean ridge basalt; it would be wholly inappropriate to digress into this minefield. The first two questions are related, and controversy on these issues seems likely to be fruitless. Obviously, samples showing recognizable metasomatism can do so only on a local scale, and because volcanically transported fragments of metasomatized mantle must be small the evidence for metasomatism on a large scale must be sought in other ways. Appropriate lines of evidence, indicating large-scale metasomatism, have been described elsewhere (Bailey 1982) but may be summarized as follows: (a) an ultramafic nodule population dominated by metasomatized fragments throughout an igneous province; (b) a range of samples showing various degrees of metasomatism through to completely transformed nodules (alkali clinopyroxenites); and (c) the chemical equivalence of extensively metasomatized nodules and erupted melt compositions. Much of the cogency of the case for large-scale metasomatism is lost if discussion is allowed to focus just on laboratory data: it is then easy to lose sight of the wider geological perspective, where alkali- and volatile-rich magmatism has to be seen in the context of the geothermal and tectonic conditions of stable plates (Bailey 1983). This larger-scale (geological) evidence must be taken together with that of metasomatism on a small scale if we are to retain a balanced view. Both kinds of evidence exist (O'Reilly & Griffin 1984; Wilshire 1984) and they are not mutually exclusive. No amount of evidence or
argument in favour of one can, of itself, falsify the case for the other. Good petrographic evidence of local metasomatism is obviously immune to rebuttal, and no evidence has yet emerged to reject the case for regional development of the process; until such evidence emerges it is important that the argument should not become polarized (as in some older geological disputes) lest we end up with a controversy without roots in nature. The dilemma of the relationships between melting and metasomatism has been highlighted by Hawkesworth et al. (1984), who distinguish two enrichment processes, one typified by subsolidus metasomatism in kimberlite nodules and the other typified by injection of small-volume melts (typically associated with basanitic magmatism). These are essentially expressions of the effects of pressure and temperature, and it will be shown that they may be explained by the interplay of geothermal gradient, mantle solidus and mineral stabilities. It will be seen there that subsolidus metasomatism can also play a vital role in basanitic magmatism. Although it is valuable to identify the chemical characteristics of the two types of enrichment (Hawkesworth et al. 1984) it should be remembered that there is a spectrum of alkaline magmatism between kimberlites and basanites, and a spectrum of enrichment processes may be expected (see Fig. 5 and Bailey 1986).
Geological factors If all the observed metasomatism were a localized consequence of alkaline igneous activity, then the quantities of metasomatized mantle and the petrological significance of the process would be essentially trivial. It is when the geological picture is seen in total that such a conclusion becomes most suspect. Association with alkaline magmatism signifies the high activities of alkalis, mobile elements and volatiles, and the high-velocity eruptions that bring the samples to the surface emphasize the role of gases in this type of magmatism. When the tectonic framework, especially the connection between alkaline magmatism and lesions in stable lithosphere (and the repetition through time), is taken into account, it is difficult to escape the conclusion that we are observing a process in which volatiles must play a vital role. Alkali rich means volatile rich and, although there may be room for debate about whether the metasomatism is a precursor to magmatism in any particular instance, there can be little doubt that volatile migration may have far reaching effects without the necessary intervention of magma. One of the most graphic examples is in kimberlite peridotite
Mantle metasomatism nodules containing phlogopite: these were sampled from points on the geotherm well below the kimberlite solidus (Bailey 1982, Fig. 3) and hence well outside conditions where any melt could exist. Certainly there are strong grounds for supposing that volatile influx is a natural precursor to both metasomatism and alkaline magmatism.
Prospect Given an outline of present concepts of mantle metasomatism it becomes fruitful to look beyond and to consider other consequences of volatile movement through the mantle. One aspect, of immediate relevance to a volume on Alkaline Rocks, has been largely neglected in discussion of mantle metasomatism: this concerns the broader spectrum embracing not just mafic and ultramafic compositions but felsic alkaline magmas. At what has been described as the limiting case of cratonic magmatism (Bailey 1980a) there is kimberlite with no felsic associates, and closely related must be the lamproites and melilitites with very limited, if not uncertain, felsic connections. These and the ultramafic lamprophyres share other common features, such as style of eruption, that require special consideration in terms of petrogenesis; they have been discussed elsewhere (Bailey 1986). The more voluminous expressions of alkaline mafic magmatism are essentially basanitic (having some connections with nephelinites) grading through to transitional basalts, and these typically form major volcano complexes and have associated felsic magmas. Such alkali basalt magmas show evidence of mantle enrichment and characteristically may carry mantle nodules, which themselves show signs of metasomatism and/or enrichment. Their mantle source is hardly in doubt, and the possibility that this source may have undergone metasomatism raises the question whether the felsic magmas could be part of the same process. Some of the problems of alkaline felsic magmatism, such as regional magma development and comparative volumes, could be resolved in the context of regional metasomatic processes (Bailey 1972, 1974) and there have been no data on relevant mineral stabilities to exclude the possibility of sources in the upper mantle (Bailey 1976); more recent data lend credence to this possibility and allow a more penetrating look at the question. Before embarking on an examination of the more recent experimental information it is useful to look over the case in geological terms. The popular view for a long time has been that
5
felsic alkaline rocks are products of differentiation from mafic parents at relatively shallow (usually crustal) depths. Processes such as fractional crystallization are still currently in vogue in spite of many difficulties (see for instance Yoder 1979; Bailey 1981) and discussions of felsic magmas in particular seem impervious to contrary evidence. When the various pieces of evidence are assembled together, however, the case for low-pressure differentiation as a universal mechanism for producing felsic magmas can be seen to be without foundation. Essentially, evidence contrary to continuous differentiation at low pressure takes two forms: (1) felsic and intermediate volcanics carry ultramafic nodules; and (2) there are substantial composition and volume discrepancies in the igneous products of a given complex or province. As will be seen later, the first could be considered as the evidence from the high-energy ranges of the magmatic system and the second from the lower-energy ranges. Nodules
In addition to the dramatic examples of felsic lavas carrying ultramafic nodules (Wright 1966, 1969), mugearites and hawaiites are commonly reported as the hosts in ultramafic-nodule localities (see Boettcher & O'Neil (1980) for an informative listing). High-speed eruption is essential, and these are unequivocal examples of the existence of felsic and intermediate melts in the mantle: it is clear that low pressure differentiation is not essential for the formation of magmas in the range hawaiite-trachyte-phonolite. In fact, the composition of these melts may be a direct indication of physical and chemical conditions in the mantle source. Composition gaps
Daly (1925, 1927) first drew attention to the basalt-trachyte bimodality of oceanic lavas, a finding later supported by Barth (Barth et al. 1939) and statistically verified by Chayes (1963) who suggested that this raised doubts about the evolution of trachyte from basalt by continuous differentiation. Quite naturally, this suggestion provoked spirited opposition from advocates of fractional crystallization; the case was reviewed by Yoder (1973) who concluded that the gap was real and that this was not a realistic outcome of continuous fractional crystallization. More recent studies (e.g. Zielinski 1975) still invoke this mechanism, however, without apparent question, and so it is permissible to cite some additional geological facts.
6
D.K. Bailey
1 In deeply exposed continental sections revealing syenite plutons, comparable alkaline mafic and intermediate plutons are rare or absent. 2 Zoned plutons are rare, and where known are relatively small and generally considered to be composite (e.g. Monteregian Hills). 3 In keeping with 1 and 2, the common nodules in felsic volcanics are typically cognate, e.g. syenite in trachyte, peralkaline granite in comendite. To the above it may be added that the typical nodules in basanites are peridotite and gabbro. Hence it is hardly reasonable to argue that the basalt-trachyte bimodality in the ocean basins is an artefact of exposure level, eruption characteristics or sampling bias: the continental plutonic bodies and the nodule suites confirm the scarcity of intermediate magmas. Perhaps the most vivid evidence of magmatic bimodality lies on Graciosa in the Azores. Here there have been alternating eruptions of felsic and mafic pyroclastics, without detectable time breaks, requiring the simultaneous coexistence of contrasting magmas within the one volcanic system (Maund & Bailey 1982; Maund, 1985). This is simply a graphic example of the contemporaneous eruption of trachyte and basalt on oceanic islands generally, to which Daly called attention so long ago. Other composition gaps are present in other magma associations, such as basanite-phonolite and nephelinite-nephelinitic phonolite, but it may be relevant that even among the felsic rocks themselves there seem to be marked distribution maxima, as between phonolites, trachytes and rhyolites (and the plutonic equivalents). These gaps, too, militate against continuous differentiation, tending to favour partial melting or at least multiprocess origins (see Bailey (1976) for a full discussion).
Volumediscrepancies Another major problem for continuous differentiation at low pressure lies in the volume relations of the magmas in many provinces. The usual image of a large alkali basalt centre with small spines and flows of trachyte may be reasonable for some oceanic volcanoes, but it is a dangerous generalization, even for the oceans, e.g. Azores and Canaries. On the continents it is probably the exception rather than the rule: here again we have deep sections providing plutonic evidence. In major syenite provinces, such as Kola and Malawi, the basic and intermediate rocks required by continuous differentiation are not merely insufficient--they are lacking altogether.
Large syenite plutons (and large monotonic trachyte and phonolite volcanoes, as in the Kenya rift) are further evidence against low-pressure evolution of felsic magmas from a basaltic parent. Indeed, to suppose such differentiation has occurred when there is no sign of the earlier stages is to assume that syenite can form in only one way. In the volcanic regime the volume discrepancy can be seen in the Miocene-Recent activity of the East African rift zone where the calculated volumes of felsic and mafic volcanics are approximately equal (Williams 1972) and there have been regional floods of phonolite and trachyte. The latter are difficult to provide for in any process of magma generation but pose an acute problem for high-level differentiation (Bailey 1974, 1978). It should be said, too, that detailed studies of particular trachyte and pantellerite suites have not only failed to find any links with basalt (Bailey 1978) but have even revealed distinctive differences between felsic volcanoes erupting contemporaneously in the same province. In the area around Lake Naivasha in Kenya, there are Holocene phonolite, trachyte, pantellerite, comendite and basalt eruptions from overlapping centres that provide not only a series of major composition gaps but also distinctive major and trace-element patterns that are inexplicable in terms of continuous differentiation (Bailey & Macdonald, 1987). Clearly, there is commonality in that the magmas are part of a cycle of igneous activity, but the need is for a multisource/ multiprocess system of magma genesis. Yoder's solution (1973) to the dilemma of the Daly gap was the production of two contrasting magmas by fractional melting, along lines indicated by Presnall (1969). This envisages a source composition with two invariant points which by continuous melt extraction during an igneous cycle yields two contrasting magmas. Melt generation from two distinct invariant points would clearly alleviate some of the problems of alkaline magmatism and would explain composition gaps, but the generation of large volumes of felsic magma from a peridotite mantle source remains a difficulty, and a new problem is introduced of irreversibility and timing. Either the source volume would have to keep changing or the liquid from the lower-temperature invariant point could be erupted only at the start of a melting cycle. These difficulties could be solved through replenishment of felsic components by metasomatism or other forms of enrichment. At the very least therefore Yoder's solution requires an open system with melting pulses acting on periodically enriched sources (cycles of melting and enrichment). Even so the volume problem remains, and
Mantle metasomatism an additional factor is introduced by alternating mafic-felsic volcanism, exemplified by Graciosa, where there is effectively simultaneous eruption of felsic-mafic magmas poor in phenocrysts. Such activity implies the need for two (or more) distinctly different source compositions to be melted and tapped during the same igneous cycle. Geological evidence indicates that felsic sources richer than normal peridotite are feasible in the mantle and recent experimental evidence confirms that appropriate mineral stabilities extend into the upper part of the mantle. When these are taken together with eruption characteristics, depth indicators in the nodule suites and evidence of metasomatism, it becomes possible to propose a scheme of magma genesis that can reconcile all the seemingly conflicting observations and give an integrated pattern to the activity.
Metasomatism and the development of separate felsic sources Peridotites containing minor amounts of plagioclase, trachytes containing peridotite nodules, and sanidine-coesite eclogite nodules are some of the evidence indicating that felsic melts and sub-solidus felsic mineralogies are possible in the mantle, and it has been clear for some years that there is no a priori reason to suppose otherwise (Bailey 1976). In view of the problems with felsic magmatism, it is appropriate to update the evidence on mineral stabilities and then to enquire whether and how two sources could develop, and how magmas could be generated and erupted. In Figs 1 and 2 some of the relevant stability fields are plotted. Obviously the bulk mantle composition must influence some of these, but broadly speaking felsic minerals could exist to depths of about 100 kin, the exact limit depending also on the prevailing temperature. The special importance of the stability ranges shown in Figs 1 and 2 lies in the fact that the upper mantle within the felsic stability zone has the potential to yield melts of felsic character at a relatively low-temperature invariant point. No such invariant melting point is possible at greater depths, so that the first melts (from a phase assemblage containing no felsic minerals) must of necessity be different, and in all probability distinctly mafic. Thus the possibility of two distinct sources with completely different first melts is evident. Could their potential for producing contrasting magmas be established and enhanced by processes of enrichment and/or metasomatism ? Starting from the critical observation that high
7
volatile activity is a hall-mark of alkaline magnatism, it has been possible to develop a hypothesis relating magmatic variations to melting and metasomatism by volatile flux along different geothermal gradients (Bailey 1970, 1980a). Further consideration of the effects of the processes along low geothermal gradients has been given elsewhere (Bailey 1984, 1986). For the generation of felsic melts, it is necessary to look to activity along steeper geotherms more appropriate to oceanic conditions (or those away from continental craton nuclei). On steeper geotherms, such as those that would cross the solidus at high levels in the mantle (in the felsic stability region), volatiles could migrate along the geotherm only if the channelways were lined (Olafsson & Eggler 1983); melting is then more likely. When the possibility of flux melting along steeper geotherms was discussed previously (Bailey 1980a, 1983) attention was focussed on the eruption of melt through the overlying lithosphere with increasing departure of melt temperature from that of the wall-rocks, i.e. an intrusion, or melt diapir, rising as a detached thermal anomaly. In recent years, however, the concept of migration of very small percentage melts has found increasing favour (Walker et al. 1978; Waft 1980; Stolper et al. 1982). Essentially the melt is envisaged as percolating through the mantle, in which case it should maintain thermal equilibrium with its path, and in the simplest case could be envisaged as migrating along the geothermal gradient. Where the melt is generated by volatile flux, it would effectively extend the path of flux migration. Gradual percolation over a long distance would mean that the melt must also maintain phase equilibrium with the mineralogy through which it moves: most alkaline ultramafic melts are of alkali clinopyroxenite composition (Lloyd & Bailey 1975; Lloyd 1981) and could percolate only along channelways akin to the source mineralogy. Either there must be pervasive distribution of all the required phases or the movement must be restricted to channelways lined with these phases. As the initial melting is seen here as resulting from volatile influx, the pathways are more likely to be enriched by interaction with the melt, which may be contrasted with the scavenging process of melt percolation envisaged by Fitton & Dunlop (1985). In the higher section of the path the percolating melt will, in any case, start to reapproach the solidus and will then progressively precipitate solids such as clinopyroxene and, closer to the solidus, minerals such as amphibole, phlogopite, carbonate, feldspar and feldspathoids. Consequently the melt channels will be further lined or plated with minerals containing lithophile ele-
8
D.K. Bailey
ments. This may lead to occlusion of the percolation holes, requiring changes of course for subsequent percolating melt and thus extending and intensifying an enriched zone in the mantle. At the same time the temperature will be raised by the release of latent heat of fusion so that continuing percolation will act as an effective means of heat transfer and gradually steepen the geothermal gradient. Under conditions of 'equilibrium' percolation, melt will be used up where the geotherm passes back through the solidus, leaving only residual volatiles to continue along the sub-solidus path. These may be expected to produce an intense zone of metasomatism in the immediately overlying lithosphere segment indicated in Fig. 3. Melts or volatiles migrating along steep geotherms will not encounter the stability regions of either phlogopite or amphibole (see Fig. 2) so that most of the trapping of lithophile elements would be in clinopyroxene and/or carbonate until the solidus is approached. Thus, most of the mobile constituents such as Na, K, OH, F and C1 will be concentrated into any surviving lowtemperature melts and fluids, leading to enrichment in felsic minerals and amphibole at nearsolidus conditions. During the early phase of a new igneous cycle along steeper geotherms, therefore, there would be enrichment and meta-
800
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FIG. 3. The consequences of flux migration and melt percolation along an initial geotherm G1 (similar to O in Fig. 1) so that continued flux would develop alkaline ultramafic and/or mafic melts. As melt (arrows) percolates back towards the solidus (OE, as in Fig. 1) it precipitates lithophile-element-bearing minerals and releases heat, 'instantaneously' steepening the channel geotherm (G2). Vapour released, on solidification of melt, heats and metasomatizes the sub-solidus sector. With time, G2 will be smoothed by exchange with the greater mass of the surrounding mantle, and ultimately will be restored to a steady-state geotherm when activity ceases. (See text for further discussion.)
somatism in the amphibole-felsic zone of the upper mantle. Furthermore, the conditions would locally favour pegmatitic formation, with scope for the development of large megacrysts of amphibole, alkali felspars and soda-rich plagioclase; near-solidus volatile accumulation could then lead to small, explosive high-velocity eruptions, bringing such megacrysts to the surface. Melt percolating along grain boundaries obviously cannot penetrate below the solidus boundary, but continued percolation towards the solidus may lead to melt accumulation in this sector of the path, through a combination of steepening geothermal gradient and channel occlusion. The solidus boundary may thus act as a melt dam allowing the development of near-solidus melt pockets, capable of diatremic eruption, of which a typical manifestation would be lamprophyric breccias (see Bailey (1986) for further details). Accumulation of felsic melts would also be possible at temperatures below the peridotite solidus shown in Fig. 3 whenever the mantle becomes extensively modified so that the peridotite solidus no longer applies and other lowertemperature solidi become appropriate. In the strict sense this should take effect as soon as felsic minerals are present, but in the initial stages any melt would be volumetrically insignificant; these interstitial melts may themselves percolate, however, and accumulate into eruptible volumes at temperatures below the peridotite solidus. The preceding discussion provides the framework for generation of felsic melts in the alkaline magma association; it remains now to examine the conditions under which the mafic melts are generated, and in particular how some of these melts achieve the conditions necessary for highvelocity eruption.
Mafic melt diatresis Under the peridotite near-solidus conditions described above, the distinctive products are felsic magmas, high-velocity diatremic eruptions (felsic to mafic) carrying megacrysts and nodules, and lamprophyric breccias. In contrast, it is implicit in the discussion that many associated mafic volcanics originate below the felsic stability zone and in general will be at higher temperatures (the latter is also an essential element of Yoder's solution). Alkaline ultramafic and mafic rocks are indeed sometimes erupted as liquid, examples being available from lamproite through melilitite and nephelinite and extending to basalts; such melt eruptions present a fundamentally different facet of alkaline igneous activity. In addition to nearsolidus melts, the magma system has to permit
Mantle metasomatism melt accumulation at a temperature well above the vapour-saturated peridotite solidus, such that the eruption trajectory in P - T space intersects the Earth's surface above the solidus, as seen in Fig. 4. Some ultramafic eruptions may take the form of lapillae sprays in which the lapillae are lava droplets (Lloyd 1985); the associated pyroclastic deposits may also contain ultramafic nodules testifying to high-velocity eruption from a great depth. This must represent the case of a fluidized system originating in the mantle with a high liquid content and presumably driven by rapidly exsolved gas. Elsewhere ultramafic nodules are also to be found in lava flows, which must indicate more coherent high-velocity eruptions of melt plus fragments, and this style of eruption seems to apply most commonly in the case of nodule-bearing basanitic eruptions. The last factor, along with other basanite-eruption characteristics, is consistent with generally lower gas contents compared with nephelinite-melilitite magmas. High-velocity eruption requires a mechanism or agent of acceleration, and the only plausible 60(
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FIG. 4. P-Trelations of detached melt bodies rising from a region of accumulation on a geotherm (G) along an adiabatic path (A, an adiabat line with its origin at 1100~ from the Nyiragongo lava lake). Carbonate peridotite solidus is CSQR, with TQ the decarbonation boundary of Wyllie (1979); the equivalent boundary of Olafsson & Eggler (1983) is shown by the parallel (broken) line originating in OE (as in Fig. 1). Basanitic melt B accumulates and starts buoyant ascent from a point on the geotherm around 1200~ (approximating the asthenosphere boundary). High-velocity eruption ('melt diatresis') is triggered by massive gas exsolution as the magma passes through the melt structure transition where CO32- becomes unstable (taken as approximating the decarbonation boundary TQ). The pressure limit for amphibole and spinel peridotite stability (at near-solidus temperatures) is indicated at the right margin.
9
agent seems to be gas, either injected into an existing melt chamber or produced by rapid exsolution from the melt itself; at the moment the latter mechanism seems the logical candidate. If vapour-saturated melt is rising through the lithosphere it is to be expected that depressurization will cause progressive gas release; this would tend to quicken the rate of ascent but may be partly offset by increasing viscosity due to melt structure changes resulting from falling temperature and loss of volatiles. The effectiveness of viscosity changes must depend on initial melt composition and temperature. Speeding u p by gas expansion will be a function of release rate and gas density. High-speed liquid eruptions carrying mantle nodules are not the norm, however, and it would be useful if some additional 'firing' or 'lift-off mechanism could be identified in addition to simple decompression vesiculation. One potential trigger for the rapid and massive emission of gas would be when a rising body of melt passes through a phase boundary or melt structure transition. In the case ofbasanitic melts, the most common nodule mineralogies are spinel lherzolite, frequently including amphibole, to which experimental evidence can be applied to give a clear indication of the typical sampling depths of nodule-bearing basanitic eruptions (Fig. 4). This characteristic association with spinel lherzolite strongly suggests that nodulebearing basanitic eruptions get a sudden acceleration (or lift-off) within or close to the spinel and amphibole stability fields, i.e. massive gas exsolution occurs at depths around 80 km. As shown in Fig. 4, there is a ready explanation for this particular case, because around this level carbonate becomes unstable, and the passage of basanitic melt containing dissolved CO2 (in the form of CO32 -) through the carbonate stability boundary could provide the necessary trigger for highvelocity eruption. For liquid diatresis, it is necessary to have melt accumulation to the point where it starts to rise independently through cooler lithosphere, i.e. detached from the ambient mantle geotherm. It is important to realize, however, that although the melt and wall-rocks may be thermally separated they may still be chemically similar. This is because melt piercement of the lithosphere is the probable culmination of a period of volatile flux and melt percolation. Thus, a rising mass of melt is likely to encounter earlier-enriched wall-rocks in its upward passage. Olafsson & Eggler (1983) and Brey et al. (1983) have pointed out that amphibole, phlogopite and carbonate disappear within a small temperature interval above the solidus, so an intrusive melt may be expected to consume more of these phases in any interaction
I0
D.K. Bailey
it may have with the wall-rocks. An ascending melt body would thus tend to develop and preserve similar characteristics to its enriched path zone, in keeping with the observation that the magmas in a given province are chemically similar to the metasomatized nodules that they bring to the surface (Lloyd & Bailey 1975). The conclusion drawn then, that the magmas were formed by direct melting of the nodule compositions, may be too simple. The eruption characteristics require that at the level from which the nodules were picked up the carrier melt was at a temperature well above the vapour-saturated peridotite solidus. Hence the initial melt must have a deeper source than the bulk of the ultramafic nodule population. Depending on the source depth, its initial chemistry might even have been alkaline ultramafic, but it attained its erupted chemistry by interactions with metasomatized wall-rocks up to the point of lift-off. Most of the nodules were derived from the wallrocks around and in front of the point of rapid gas exsolution, which provided the impetus for their high-velocity transport to the surface. In areas where the ambient geothermal gradient is steep, an alkaline magma system can now be envisaged as comprising two distinct source regions, distinguished on the basis of felsic mineral stability. One region, below the maximum depths of felsic mineral stability (approximately coincident with spinel and amphibole stability) is the potential source for mafic hypersolidus melts whenever there is volatile flux through the lithosphere. The overlying zone of felsic stability is the locus of intense melt enrichment and metasomatism at near- and subsolidus conditions during volatile fluxing: this forms the source region for the felsic and nearsolidus portions of the activity. Mafic and felsic magmas may thus both be primary products of the same igneous cycle: the first forms by volatile influx leading to melt diapirism (mafic) well above the peridotite solidus (Fig. 4), the second by melt percolation leading to melt diapirism (felsic) near or below the peridotite solidus (Fig. 3).
Conclusion By combining the geological and experimental evidence it is possible to produce a unified scheme for the generation and eruption of mafic and felsic alkaline magmas, constituting the range sometimes described as the alkali olivine basalttrachyte association. A combination of volatile and melt percolation along geothermal gradients appropriate to thicker oceanic and younger
continental lithosphere leads to enrichment and metasomatism in the P-Tregion of the amphibole peridotite solidus, providing the major source region for felsic magmas: the main source of mafic magmas is at depths below the stability range of felsic minerals, eruption resulting from melt accumulation and buoyant uprise. These conditions are summarized in sections 0 and 30 in Figs 5 and 6. Melt generation from two distinct sources can account for the following characteristics of this magmatic association: 1 Composition gaps. Intermediate compositions may be formed but are not essential to petrogenesis. 2 Volume discrepancies. The relative volume of mafic-to-felsic magma results from the interplay of melt percolation and mafic melt diapirism. 3 Timing. Characteristic magmatic diversity results simply from the fluctuating interplay of volatile flux and melt generation (irreversibility is not an issue). 4 Synchroneity. Both felsic and mafic magmas can be available for eruption at all times. 5 Megacrysts. These are typically erupted in explosive diatremic activity and show significant differences from megacrysts associated with ultramafic magmatism. They are commonly amphibole and feldspar in the alkali basalt association and are explicable in terms of near-solidus formation. 6 Basanites may be erupted in a highly liquid state, carrying nodules of amphibole and spinel peridotite which are indicative of higher-temperature generation from a deeper source than felsic melts. High-velocity eruption is 'fired' by gas exsolution when CO32- becomes unstable in the melt. 7 Characteristic abundance of lithophile elements. Volatile fluxing, metasomatism and enrichment provide the required concentration mechanism. All the above features are also consistent with the observation that alkaline magmatism is controlled by lithosphere structures acting as focussing zones for volatile escape to the Earth's surface (Bailey 1977, 1980b, 1983). In earlier discussions attention has been concentrated on continental examples where, in addition to the advantages of many different erosion levels, there is frequently evidence of repetition of the activity through the same lesion in the lithosphere. Since a crucial observation for the present paper has been the alternating mafic-felsic volcanism in the Azores, it is appropriate to conclude by pointing out that the Azores constitute an excellent oceanic case for lithosphere control.
Mantle metasomatism GEOTHERM : NEAR S (SHIELD)
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MANTLE }
IN M O D I F I E D
WITH
[
MELTS
MANTLE
MANTLE
1
MELTS
MANTLE
I MELTS
AND MELTS
CONTINUOUS
FLUX
THROUGH
LESION
FIG. 5. Sample lithosphere sections along four selected geothermal gradients, derived directly from Figs 1 and 2. In each case it is assumed that the initial mantle composition is peridotite. On the left of each column are indicated the mantle changes induced by volatile flux along the geotherm, and on the right the kinds of melt developed in this regime. 'Melt' is not mentioned on the left in column 1 to emphasize the limited melting expected when the geotherm S is in grazing incidence with the kimberlite solidus (KS in Fig. 1). The base of columns 3 and 4 has been set at 1200~ on the geotherms, as an approximation to the asthenosphere boundary. As the geotherm becomes steeper (columns 1-4) the melts become more siliceous, with increasing separation from the solidus, in the manner prescribed by Olafsson & Eggler (1983). Volatile fluxing along geotherm O could not produce strong phlogopite enrichment but instead gives rise to assemblages dominated by clinopyroxenes, felsics and (at higher levels) amphibole, i.e. directly related to basanitic melts.
Felsic volcanoes, especially trachytes, form a major component of the activity on some of the islands, notably those lying along the Terceira rift: Sao Miguel, Terceira and Graciosa. On these three islands not only are the composition gaps pronounced but the large volumes of felsic
_
JL
Z
Z THOSP
ERE
,!:i;i
MOBILE ELEMENTS FLOW TO LESION
Fio. 6. Sketch to illustrate the operation of volatile focussing through lesions in the lithosphere of different thicknesses and geothermal gradients. Columns correspond to those depicted in Fig. 5. Discussion of the relevant applications of channelling volatiles through a narrow zone in the lithosphere ('pie-funnel effect') can be found in Bailey (1980b; 1983).
magmas constitute major embarrassments for continuous differentiation. Explanations for individual volcanoes have been attempted, but the case for continuous differentiation is lost when the activity is seen in regional context. It is hardly conceivable that the volcanoes strung along the Terceira rift should all be spectacularly successful products of high-level differentiation while the volcanoes on the other islands outside the rift have effectively failed. It is more reasonable to assume that the Terceira rift volcanoes are the result of structural control that has permitted the formation, below the rift, of mantle source regions capable of providing felsic melts. Such a conclusion is consistent with the scenario whereby the rift zone acts as a focus for the release of volatiles through the Azores lithosphere. ACKNOWLEDGMENTS: My appreciation of magmatic composition gaps has been sharpened by discussions with Felix Chayes and Ray Macdonald, and by Julian Maund's systematic observations on the 'zebra-striped' tephra deposits of Graciosa. Constructive reviews of the manuscript were provided by Godfrey Fitton and Martin Menzies.
I2
D. K. Bailey
References BAILEY, D. K. 1970. Volatile flux, heat focussing and the generation of magma. Geol. J. Spec. Issue No. 2, 177-86. - 1972. Uplift, rifting and magmatism in continental plates. J. Earth Sciences (Leeds), 8, 225-39. - - 1 9 7 4 . Continental rifting and alkaline magmatism. In: SORENSON, H. (ed.) The Alkaline Rocks, pp. 148-59. Wiley, New York. - - 1 9 7 6 . Applications of experiments to alkaline rocks. In: BAILEY, D. K. & MACDONALD, R. (eds) The Evolution of the Crystalline Rocks, pp. 416-69. Academic Press, London. --1977. Lithosphere control of continental rift magmatism. J. geol. Soc. London, 133, 103-6. - 1978. Continental rifting and mantle degassing. In: NEUMANN, E.-R. & RAMBERG, I. B. (eds) Petrology and Geochemistry of Continental Rifts, pp. 1-13. Reidel, Dordrecht. - 1980a. Volatile flux, geotherms, and the generation of the kimberlite-carbonatite-alkaline magma spectrum. Mineralog. Mag. 43, 695-9. - 1980b. Volcanism, Earth degassing and replenished lithosphere mantle. Phil. Trans. R. Soc. A 297, 309-22. - - 1 9 8 1 . Bowen and the Post-Impressionists. Geol. Mag. 118, 311-8. - 1982. Mantle metasomatism--continuing chemical change within the Earth. Nature, Lond. 296, 525-30. 1983. The chemical and thermal evolution of rifts. Tectonophysics, 94, 585-97. - 1984. Kimberlite: 'the mantle sample' formed by ultrametasomatism. In: KORNPROBST, J. (ed.) Kimberlites. 1 : Kimberlites and Related Rocks, pp. 323-33. Developments in Petrology 11 A, Elsevier, Amsterdam. - 1986. Fluids, melts, flowage and styles of eruption in alkaline ultramafic magmatism. Trans. geol. Soc. S.Afr. 88 (2), 449-57. -& MACDONALD, R. 1987. Dry peralkaline felsic liquids and carbon dioxide flux through the Kenya rift zone. In: MYSEN, B. O. (ed.) Magmatic Processes: Physiochemical Principles. Geochem. Soc., Spec. Pub. No. 1, in press. BARTH, T. F. W., CORRENS, C. W. & ESKOLA, P. 1939. Die Entstehung der Gesteine, 422 pp. Springer, Berlin. BOETTCHER,A. L. & O'NEILL, J. R. 1980. Stable isotope, chemical, and petrographic studies of high-pressure amphiboles and micas: evidence for metasomatism in the mantle source regions of alkali basalts and kimberlites. Am. J. Sci. 280A, 594621. BREY, G., BRICE, W. R., ELLIS, D. J., GREEN, D. H., HARRIS, K. L. & RYABCHIKOV, I. D. 1983. Pyroxene-carbonate reactions in the upper mantle. Earth planet. Sci. Lett. 62, 63-74. CHAYES, F.1963. Relative abundance of intermediate members of the oceanic basalt-trachyte association. J. geophys. Res. 68 (5), 1519-34. CLARK, S. P. & RINGWOOD, A. E. 1964. Density
distribution and constitution of the mantle. Rev. geophys. 2, 35. DALY, R. A. 1925. The geology of Ascension Island. Proc. Am. Acad. Arts Sci 60, 1-80. -1927. The geology of St. Helena Island. Proc. Am. Acad. Arts Sci. 62, 31-92. EGGLER, D. H. & WENDLANDT, R. F. 1978. Phase relations of a kimberlite composition. Cam. Inst. Wash. Yr Bk, 77, 751-6. ERLANK, A. J. & SHIMIZU, N. 1977. Strontium and strontium isotope distributions in some kimberlite nodules and minerals. Extended Abstracts, 2nd Int. Kimberlite Conference. - - A L L S O P P , H. L., DUNCAN, A. R. & BRISTOW, J. W. 1980. Mantle heterogeneity beneath southern Africa: evidence from the volcanic record. Phil. Trans. R. Soc. A297, 295-307. FITTON, J. G. & DUNLOP, H M. 1985. The Cameroon line, West Africa, and its bearing on the origin of oceanic and continental alkali basalt. Earth planet. Sci. Lett. 72, 23-38. HARTE, B., COX, K. G. & GURNEY, J. J. 1975. Petrography and geological history of upper mantle xenoliths from Matsoku kimberlite pipe. Phys. Chem. Earth, 9, 389-416. HAWKESWORTH, C. J., ROGERS, N. W., VAN GALASTEREN, P. W. C. & MENZIES, M. A. 1984. Mantle enrichment processes. Nature, Lond. 311 (5984) 331-5. KENNEDY, C. S. & KENNEDY, G. C. 1976. The equilibrium boundary between graphite and diamond. J. geophys. Res. 81, 2467-70. LLOYD, F. E. 1981. Upper-mantle metasomatism beneath a continental rift: clinopyroxenes in alkali mafic lavas and nodules from S. W. Uganda. Mineralog. Mag. 44, 315-23. - - 1 9 8 5 . Melting experiments on glassy olivine melilitites from strongly potassic mafic volcanics ofS. W. Uganda. Contrib. Mineral. Petrol. 90, 23643. - & BAILEY, D. K. 1975. Light element metasomatism of the continental mantle: the evidence and the consequences. Phys. Chem. Earth, 9, 389-416. MAUND, J. G. 1985. The volcanic geology, petrology and geochemistry of Caldeira volcano, Graciosa, Azores, and its bearing on contemporaneous felsicmafic oceanic island volcanism. Ph.D. Thesis, University of Reading (unpublished). - - , J. G. & BAILEY, D. K. 1982. Caldeira volcano, Graciosa, Azores: Progress Report. J. geol. Soc. Lond, 139, 658. MENZIES, M. & MURTHY, V. R. 1980a. Mantle metasomatism as a precursor to the genesis of alkaline magmas--isotopic evidence. Am. J. Sci. 280A, 622-38. - & 1980b. Nd and Sr isotope geochemistry of hydrous mantle nodules and their host alkali basalts: implications for local heterogeneities in metasomatically veined mantle. Earth planet. Sci. Lett. 46, 323-34. OLAFSSON, M. & EGGLER, D. H. 1983. Phase relations of amphibole, amphibole-carbonate, and phlogo-
Mantle metasomatism pite-carbonate peridotite: petrologic constraints on the asthenosphere. Earth planet. Sci. Lett. 64, 305-15. OLDENBURG, D. W. 1981. Conductivity structure of oceanic upper mantle beneath the Pacific plate. Geophys. J. R. astron. Soc. 65, 359-94. O'REILLY, S. Y. & GRIFFIN, W. L. 1984. Sr isotopic heterogeneity in primitive basaltic rocks, southeastern Australia: correlations with mantle metasomatism. Contrib. Mineral. Petrol. 87, 220-30. PRESNALL, D. C. 1969. The geometrical analysis of partial fusion. Am. J. Sci. 267, 1178-94. SCLATER, J. G., JAUPART, C. & GALSON, D. 1980. The heat flow through oceanic and continental crust and the heat loss of the Earth. Rev. Geophys. Space Phys. 18, 269-311. STOLPER, E., WALKER,D., HAGER, B. H. & JAYS, J. F. 1982. Melt segregation from partially molten source regions: the importance of melt density and source region size. J. geophys. Res. 86, 6261-71. TOZER, D. C. 1967. Towards a theory of mantle convection. In: GASKELL, T. F. (ed.) The Earth's Mantle, pp. 325-53. Academic Press, New York. WAFt, H. S. 1980. Effects of the gravitational field on liquid distribution in partial melts within the upper mantle. J. geophys. Res. 85, 1815-25. WALKER, D., STOLPER, E. & HAYS, J. F. 1978. A numerical treatment of melt/solid segregation: size of the eucrite parent body and stability of the
13
terrestrial low-velocity zone. J. geophys. Res. 83, 6005-13. WENDLANDT,R. F. & EGGLER, O. H. 1980. The origins of potassic magmas: 2. Stability of phlogopite in natural spinel lherzolite and in the system KA1SiO4-MgO-SiOE-H20-CO 2 at high pressures and high temperatures. Am. J. Sci. 280, 421-58. WILLIAMS,L. A. J. 1972. The Kenya Rift Volcanics: A note on volumes and chemical composition. Tectonophysics, 15, 83-96. WILSHIRE, H. G. 1984. Mantle metasomatism: The REE story. Geology, 12, 395-8. WRIGrtT, J. B. 1966. Olivine nodules in phonolite of the East Otago Alkaline Province, New Zealand. Nature, Lond. 210, 519-20. --1969. Olivine nodules and related inclusions in trachyte from the Jos Plateau, Nigeria. Mineral. Mag. 37, 370-4. WYLLIE, P. J. 1979. Petrogenesis and the physics of the Earth. In: YODER, n . S., Jr (ed.) The Evolution of the Igneous Rocks, pp. 481-520. Princeton University Press, Princeton, NJ. YODER, H. S. Jr. 1973. Contemporaneous basaltic and rhyolitic magmas. Am. Mineral. 58, 153-71. --(Ed) 1979. The Evolution of the Igneous Rocks. Fiftieth Anniversary Perspectives, 588 pp. Princeton University Press, Princeton, N.J. ZIELINSKI, R. A. 1975. Trace element evaluation of a suite of rocks from Reunion Island, Indian Ocean. Geochim cosmochim. Acta, 39, 713-34.
D. K. BAILEY, Department of Geology, University of Reading, Whiteknights, Reading, RG6 2AB, U.K.
Alkaline rocks and their inclusions: a window on the Earth's interior Martin Menzies S U M M A R Y : The majority of inclusion-bearing alkaline rocks have a source region identical with that of ocean island basalts regardless of whether they were erupted through oceanic or continental crust. Models have been proposed where oceanic alkaline rocks are derived from the asthenosphere and during passage to the surface there have occurred variable amounts of interaction with lithosphere of the mid-ocean ridge basalt type. A similar scenario would account for most inclusion-bearing continental alkaline rocks. Genesis of micaceous kimberlites and lamproites, however, may require that asthenospheric plumes have impinged on 'nuggets' of aged heterogeneous enriched sub-continental lithosphere. This would lead to mixing of isotopicallydistinct asthenospheric and lithospheric melts.
Introduction Inclusion-bearing alkaline rocks are erupted at the Earth's surface in the ocean basins and on the continental crust. They include basanites, nephelinites, hawaiites, alkali olivine basalts, phonolites, Group I and Group II kimberlites and lamproites. All these magmas or melts are known to transport spinel- or garnet-bearing peridotites and/or high-pressure megacrysts to the surface (e.g. diamond, sapphire, garnet etc.). The fact that they have disrupted and entrained mafic and ultramafic fragments stable at mantle pressures and temperatures argues very strongly for an ultimate mantle source. Important information about the geochemistry of the upper mantle is retained in the host magmas. Moreover, the inclusions provide added insight into the nature of the mantle occurring at shallower depths than the source of the host magma. Early research into the possible origins of alkaline rocks stressed the 'mantle-type' 87Sr/ S6Sr < 0.706) and 'crustal-type' (s 7Sr/S6Sr > 0.710) isotope ratios as a means of identification of possible source regions (Powell & Bell 1974; Rock 1976). Any similarity between isotope ratios measured in alkaline rocks and mid-ocean ridge basalts (MORBs) or ocean island basalts (OIBs) was taken as unequivocal proof that the particular rock in question was mantle derived (e.g. East African rift). Similarly, any alkaline rock with an isotopic composition similar to crustal values was assumed to be a crustal derivative or to have experienced crustal contamination (e.g. Kimberley, Australia). For example, Rock (1976) commented that in certain alkaline rocks the Sr isotopic ratios were 'so high (>0.710) that a significant contribution from crustal material was inescapable'. It is implicit in these models that interaction with a 'crustal' component occurred after magma
formation. More recently, alkaline rocks with radiogenic Sr (and non-radiogenic Nd) isotopic ratios have been interpreted as extracts from mantle source regions where addition of a crustal component to the mantle occurred prior to magma production (Chase 1981; Hofmann & White 1982; Zindler & Jagoutz 1987) by recycling processes involving subduction and/or delamination of aged sub-continental mantle (McKenzie & O'Nions 1983). Alkaline rocks may include components from the following:
1 Lower mantle." below the boundary layer (about 670 kin). 2 Upper mantle." oceanic lithosphere; continental lithosphere; convecting asthenosphere. 3 Recycled components." oceanic and continental crust and mantle. In this review isotopic data pertinent to inclusion-bearing alkaline rocks and their cargo of mantle inclusions erupted in ocean basins, continental rift valleys and stable cratonic regions will be used to construct an Earth model. Earth models will be considered where the structural elements that form the upper mantle below both oceans and continents are lithosphere (crust and mantle) and asthenosphere. These in turn are underlain by the lower mantle below the 670 km boundary layer.
Alkaline rocks and their mantle inclusions Random fragments of mantle are transported to the surface in alkaline rocks. The inclusions entrained in alkaline rocks range from garnet
From: FITTON,J. G. & UPTON,B. G. J. (eds), 1987, Alkaline Igneous Rocks, Geological Society Special Publication No. 30, pp. 15-27.
15
M. Menzies
[6
and spinel peridotites to mica clinopyroxenites, glimmerites, apatite-amphibole pyroxenites and mica-rutile-ilmenite-diopside (MARID) rocks. It is useful to think of the variety of inclusions as disrupted mantle pegmatites viz. pieces of mantle wall-rock (i.e. peridotites) and fragmented conduits of silicate melts or aqueous fluids (i.e. pyroxenites, glimmerites etc.). Analogues can be found in inclusion suites for the various melts or fluids that have migrated through the mantle. A m p h i b o l e - a p a t i t e - p y r o x e n ites found as veins in spinel lherzolite inclusions are believed to be a product of silicate-melt migration (Hawkesworth et al. 1984; Menzies et al. 1985c). A by-product of this magmatism is the Fe-Ti metasomatism observed in wall-rock peridotites. In other words, upwelling basanitic or nephelinitic melts may have crystallized in the mantle as these vein networks. It can be shown that their isotopic and trace-element characteristics are consistent with such an interpretation (Irving and Frey 1984; Roden et al. 1984). M A R I D rocks (Kramers et al. 1983; Smith 1983) found in kimberlite inclusion suites are believed to be portions of the conduit lining left after passage of a hydrous fluid enriched with incompatible elements. Glimmerites and micaceous pyroxenites (Erlank et al. 1987; Menzies et al. 1987) may also represent disrupted vein systems formed by crystallization of LIL- (LIL stands for large
05132 ~ N I V A K ~ " 1 0
o.51~L__~ 0.702
~
Ocean basins Before any discussion of the isotopic variability of alkaline basaltic and kimberlitic rocks erupted in continental regions, it would seem logical to review the data on alkaline rocks from ocean basins (Fig. 1). Although not strictly oceanic, the Nunivak data are included in Fig. 1. Alkaline rocks associated with disruption and entrainment of fragments of oceanic mantle have a very narrow range of Sr and Nd isotopic ratios (i.e. STSr/86Sr < 0.706 and 143Nd/144Nd > 0.51270 (Fig. 1, Table 1)). These magmas are clearly derived from mantle depleted in the light REE for a considerable period of time. A concomitant depletion in Rb is also apparent, but in the case of Tahiti and Malaita the magmas have slightly higher SVSr/86Sr ratios. This could result from (a) mixing of MORB mantle with a recycled radiogenic component or (b) influx of a high Rb/Sr fluid.
0.513211
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|
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I
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Host Magmas
0.513211_
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0.706
0.704
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ion lithophile elements) and LREE-enriched hydrous fluids in the mantle (LREE stands for light rare-earth elements). Mantle inclusions provide a very important record, not only of ancient depletion events (more than 103 Ma ago) related to crust formation, but more importantly of the passage of silicate melts and hydrous fluids from the mantle to the surface.
3 0.5132~,inlcl "-
~
i
HAWAllli
+10
-5 "O 0.5128 z
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0.5128
_
9
0.5124 I 0.702 87Sr/86S
I
I 0.704
I
I 0.706
o
0.5124
-I 0.702
1
I 0704
I
I 0.706
-5
-o Z
uo
r
FIG. 1. Nd versus Sr isotopic variation (present day) in mantle ultramafic and mafic inclusions and their host basaltic rocks from oceanic basins. Note the following: (i) the host magmas are derived from mantle depleted in Rb and light REE; (ii) the inclusions exhibit a greater range of Nd and Sr isotopes than the host magmas: (iii) the high 8VSr/a6Srratios in the Malaita and Tahiti inclusions. The vertical and horizontal reference lines in this and other diagrams define the generally accepted values for undifferentiated or chondritic mantle or bulk earth (87Sr/S6Sr= 0.7045 and 143Nd/144Nd=0.51264). Data are taken from several sources. Nunivak : Menzies & Murthy (1980a); Roden et al. (1984a). Malaita: Bielski-Zyskind et al. (1984). Tahiti: Vidal et al. (1984); Menzies & Hawkesworth (1987). Hawaii: Chert & Frey (1983).
Alkaline rocks and their inclusions
I7
TABLE 1.87Sr/S6Sr, 143Nd/144 Nd, R b / S r a n d S m / N d ratios f o r inclusion-bearing alkaline, kimberlitic a n d lamproitic m a g m a s
Locality
87Rb/86Sr
esra
/;Ndb
147Sm/
Ref
144Nd Afar (Assab) - 17 to - 10.0 -Africa (Tanzania) - 14.2 -Alaska (Nunivak) - 2 8 . 4 to - 17 0.049-0.121 Antarctica Foster Crater - 13.8 to - 11.4 - Ross Island - 19.2 0.089-0.165 Arizona Geronimo - 23.4 to - 17 0.118-0.240 San Carlos - 22.7 0.091 Australia Kiama - 7.1 0.051-0.164 Kimberley + 85 to + 218.6 0.593-9.855 Victoria - 28.4 to + 4.3 -California (Sierran Province) c + 4.3 to + 34.1 0.056-1.28 Fiji - 13.4 -France (Massif Central) - 14.2 to 0 0.153-1.225 French Polynesia (Tahiti) d + 7.1 to + 34.1 -Hawaii (Honolulu Series) - 18.5 to - 17.03 0.015-0.103 Scotland (Midland Valley and N W High- - 4.3 to + 25.6 0.077-0.203 lands) Solomon Islands (Malaita) + 1.4 to + 2.8 0.059-0.183 South Africa (various localities) Group I kimberlites - 11.4 to + 14.2 0.038-1.52 Group II kimberlites + 41.2 to + 90.8 0.140-0.786 South Yemen (Ataq) - 17.0 to - 2.8 0.028-0.692 Mantle xenoliths Keel e - 3 5 . 5 t o +1867.0-Asthenosphere -35.5to+21.3 --
+4.7 to +7.2 + 1.4 + 4 . 9 t o +7.8
-0.096 0.119-0.133
+3.3
--
to
+4.7
1 2 3
--
--
4
+ 4 . 9 t o +8.0 +7.6
0.101-0.126 0.130
5 6
- 0 . 6 t o +0.4 - 16.2 to - 8 . 0 - 4 . 7 to +5.0 - 9 . 6 to +2.9
0.116-0.120 0.065-0.082 0.112-0.119 0.103-0.140
+3.7
--
+ 7 . 0 t o +7.6 - 6 . 6 t o +3.1
0.120-0.168 0.068-0.108
7 8 9 10 11 12 13 14 15
+ l . 9 t o +2.9
0.102-0.121
16
- 4 . 1 to +1.0 - 12.9 to - 8.8 - 2 . 3 to +8.0
0.085-0.098 0.062-0.087 0.100-0.130
17 18
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---
19
+2.1 to +4.1
0.087-0.119
-
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"es~ represents the present-day 87Sr/86Sr ratio assuming a bulk earth s 7Sr/86Sr of 0.7045. beyo represents the present-day 143Nd/144Nd ratio assuming a bulk earth 143Nd/144Nd of 0.51264. cRange in Rb/Sr and Sm/Nd given for the Sierran Province lavas are for basaltic rocks that are not necessarily host to the xenoliths. The isotope data and trace-element data are for different samples. aTahitian samples are not the host magmas to the xenoliths. eThe range given for the 'keel' inclusions excludes inclusions found in diamonds (Richardson et al. 1984). 1 2 3 4 5 6 7 8 9 10
Betton & Civetta 1984; Cohen et al. 1984; Menzies & Murthy 1980a; Roden et al. 1984a; Stuckless & Ericksen 1976; Menzies et al 1985b; Menzies et al. 1985a; Evans & Nash unpublished data; Zindler & Jagoutz 1987; Wass & Rogers 1980; Menzies & Wass 1983; McCulloch et al. 1983 ; Frey & Green 1974; McDonough & McCulloch 1985; Van Kooten etal 1985;
T h e spinel a n d g a r n e t p e r i d o t i t e s a n d p y r o x e n i t e s e n t r a i n e d b y t h e s e m a g m a s (Fig. 1) h a v e a s i m i l a r r a n g e o f Sr a n d N d isotopes e x c e p t t h a t t h e i n c l u s i o n s do r e c o r d s o m e h i g h e r 87Sr/S6Sr ratios (e.g. M a l a i t a a n d Tahiti). It is v e r y difficult to state u n e q u i v o c a l l y t h a t t h e c h e m i c a l c h a r a c t e r istics o f t h e inclusions are n o t a n a d j u n c t o f
11 12 13 14 15 16 17 18 19
Gill 1984; Chauvel & Jahn 1984; Downes 1984; White & Hofmann 1982; Clague & Frey 1982; Chen & Frey 1983; Stille et al. 1983; Thirlwall 1982; Menzies & Halliday 1984; Bielski-Zyskind et al. 1984; Smith 1983; Menzies & Murthy 1980a; Menzies et al. 1987, and references cited therein; Menzies & Hawkesworth 1987.
m a g m a t i c processes, i.e. r e p e a t e d cycles o f m a g m a t i s m in t h e m a n t l e . R e g a r d l e s s o f this d i l e m m a , the basalt and inclusion data reveal mantle r e g i o n s d e p l e t e d in t h e light R E E a n d v a r i a b l y d e p l e t e d or e n r i c h e d in Rb. T h e latter e n r i c h m e n t m u s t be a r e l a t i v e l y r e c e n t p h e n o m e n o n (less t h a n 0.2 x 103 M a ago) w i t h i n t h e l i f e t i m e o f t h e
18
M. Menzies I
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0.5128 0.5124 1 0.702
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,
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TANZANIA
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0.750
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I ~ I 0704
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Magmas
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87Sr/86Sr
FIG. 2. Nd versus Sr isotopic variation (present day) in mantle ultramafic and mafic inclusions and their host basalts from rifted continental regions. Note the following : (i) the continental host magmas are derived from source regions not unlike the oceanic magmas in Fig. 1 ; (ii) in the majority of cases the host magmas and inclusions are derived from different mantle source regions; (iii) in the case of the African rift valley, fragments of heterogeneous mantle have been entrained. Data are taken from several sources. Middle East-Ethiopia: Menzies & Murthy (1980a); Betton & Civetta (1984); Menzies & Hawkesworth (1987). France and Germany: Stosch et al. (1980); Chauvel & Jahn (1983); Downes (1984). Eastern Australia: Menzies & Wass (1983). SW U.S.A. : Jagoutz et al. (1980); Stosch et al. (1980); Menzies & Hawkesworth (1987); Menzies et al. (1985a); Van Kooten et al. (1985); Zindler & Jagoutz (1987). ocean basin. It can be deduced from these data that the source regions of alkaline magmas and of the overlying mantle as represented by the inclusions are isotopically quite similar. Exposures of peridotite in the ocean basins (Roden et al. 1984b) exhibit isotopic variability that falls within the limits of the data in Fig. 1. Roden et al. (1984b) believe that the St Paul's rocks contain evidence of a recent (less than 0.2 • 103 Ma ago) trace-element enrichment related to alkaline magmatism.
Continental rift valleys Inclusion-bearing alkaline rocks erupted in the major rift valleys of the world have present-day Sr and Nd isotopic ratios not unlike those of the inclusion-bearing alkaline rocks from ocean basins. 87Sr/S6Sr ranges from 0.7027 to 0.7055, and 143Nd/144Nd ranges from 0.51305 to 0.51245 (Table 1). Interestingly, however, the marked disposition of the oceanic data towards high 87Sr/ S6Sr ratios is not evident in the continental
alkaline rocks. In fact, these data are more inclined to plot towards lower 143Nd/144Nd ratios (Fig. 2). Again, the alkaline host magmas are derived from a mantle source depleted in Rb and the light REE for a considerable period of time. Modifications to this mantle source have clearly occurred and account for the scatter of the data. For example, alkaline rocks from the Middle East (Betton & Civetta 1984; Menzies & Hawkesworth 1987) and SW U.S.A. (Van Kooten et al. 1985) plot towards more radiogenic 87Sr/86Sr ratios, a fact that indicates input of Rb into the source region or mixing of end-members with diverse 8VSr/86Sr ratios. The variety of mafic and ultramafic inclusions entrained in alkaline magmas erupted in rift valleys exposes a complex mantle structure that is at first glance extremely heterogenous for both Sr and Nd isotopes and is thus distinct from oceanic mantle. It is apparent, however, that the majority of inclusions have 87Sr/86Sr ratios ranging from 0.702 to 0.7065 and t43Nd/t44Nd ratios ranging from 0.51330 to 0.51240. Considering the paucity of data from ocean basins and the prolific number of isotopic
Alkaline
rocks and their inclusions
analyses from continental rift valleys, these data are only marginally different from oceanic-basin data. However, isotopic data from Tanzania appears to complicate matters (Fig. 2) since the inclusions display an extreme range in Sr and Nd isotopic composition. Nothing of this type has been seen in the ocean basins or in most of the continental rift valleys (compare Figs 1 and 2). These extremely radiogenic Sr and non-radiogenic Nd isotopic ratios record ancient (more than 103 Ma ago) trace-element enrichments involving addition of the light R E E and/or Rb to the mantle. With the passage of time, the isotopic ratios have adjusted so as to record the modified
0.5120
~
/
19
trace-element ratios. It is unlikely that such old trace-element enrichment events would be preserved in actively convecting asthenosphere, A mantle layer with fundamentally different isotopic characteristics clearly exists beneath the continents and not beneath the ocean basins. Furthermore, the similarity between oceanic and continental rift valley alkaline basalts points to a similar mantle source at depth for the host magmas. While most of the inclusions record an old depletion in both Rb and the light REE not unlike the oceanic mantle, certain fragments of 'anomalous' enriched mantle are occasionally entrained within continental rift valley magmas.
~-10 0'5125I
0,110
'
1o
0.702 Q704 0.706 0.708 0.710
0.700 0.710 0.720
( •Host Magmas 0.5132 0.5130 0.5128 0.5126
!
~
I
~
0.5124
0.740
( ~ Inclusions
I
~ t
0.730
I
...
-,10
I
SOUTH
AFRICA
-+~
o ~
' ,NOLUS,O,S
/
o.5122 o.5120 t Z ,r,-~., 0.5118 ~- I "O z
OY02
-
GROUP II~ K,MBERL,TES garnet inclusions in diamonds I
I
I
I
0.706
0.710
0.714
0.718
-lO
-15 "t~ Z
87Sr / 86 Sr
FIG. 3. Nd and Sr isotopic variation (present day) in mantle ultramafic and mafic inclusions and their host basalts and kimberlites from other continental regions. Note the following: (i) the host alkaline magmas from Scotland and Antarctica and the basaltic kimberlites (Group 1) from S Africa have Sr and Nd isotopic compositions not unlike OIBs; (ii) the mantle inclusions display, in most instances, a greater range of Sr and Nd isotopic ratios than the host magmas; (iii) a greater proportion of the inclusions than were entrained in oceanic basalts are from regions of the mantle enriched in Rb and light REE for a considerable time period; (iv) micaceous kimberlites (Group II) have Sr and Nd isotopic ratios distinctly different from those of OIB and similar to some of the inclusions. Data are taken from several sources. Scotland: Menzies & Halliday (1984); Halliday et al. (1985). Antarctica: Stuckless & Ericksen (1976); Kyle et al. (1985); Menzies et al. (1985b). South Africa: Menzies & Murthy (1980b); Kramers et al. (1981 ; 1983); Erlank et al. (1982, 1987); Smith (1983); Menzies et al. (1987).
M. Menzies
20
Although the source regions of the host magmas may be similar, clearly the potential exists in rift valleys for magmas to entrain fragments of mantle that does not exist under ocean basins.
istries. The range in isotopic composition is extreme, particularly with respect to Sr isotopic ratios (Fig. 3). These isotopic data record a timeintegrated response to (a) major partial melting or depletion events (similar to that observed in the alkaline volcanic rocks erupted in the oceans and rift valleys) and (b) ancient enrichment or metasomatic events caused by infiltration of silicate melts and aqueous fluids that have not as yet been reported from the ocean basins.
Other continental regions Alkaline basalts, lamproites, lamprophyres and kimberlites intrude regions of old crust. The alkaline basalts have Sr and Nd isotopic ratios consistent with derivation from either enriched (e.g. Scotland) or depleted (e.g. Antarctica) portions of the Earth's mantle (Fig. 3). Moreover, most of these continental alkaline basalts and basaltic kimberlites (Group I) are isotopically indistinguishable from ocean island alkaline basalts or rift valley basalts (Table 1). In contrast, micaceous kimberlites (Group II) and lamproites have more radiogenic Sr and less radiogenic Nd isotopic ratios than anything erupted in the ocean basins or rift valleys. In the case of the Group II kimberlites and lamproites we must invoke a source region distinctly different from that of alkaline rocks erupted in the ocean basins or continental rift valleys. It is perhaps to be expected that the petrographically and chemically diverse mafic and ultramafic inclusions entrained from below regions of old crust would display heterogeneous isotopic chemI
I
Asthenosphere-lithosphere The available host magma and inclusion data can now be integrated into some kind of Earth model. To assist us, the Sr and Nd isotopic data for continental and oceanic alkaline basaltic rocks (excluding Group II kimberlites) are compared with available data from several ocean islands and MORB (Fig. 4). The first point to be made is that the source is identical (Fig. 4) regardless of whether the rocks are erupted in continents or ocean basins. Their source must therefore lie not in the lithosphere but in or below the convecting asthenosphere since (a) the oceanic lithosphere is relatively sterile and not capable of producing large volumes of alkaline magma and (b) the I
I
I
+15
0.5134 MORB 0.5132
+1o
~Nd
0.5130 4"
A
+5
V
0.5128
<>
erguelen It
0.5126 9
A9 ~9
9
"-~
,O-
0.5124 I
0.702
0.703
-5
I
0.704
0.705
0.706
87Sr/86Sr
FIG. 4. Nd versus Sr isotopic variation in alkaline volcanic rocks and basaltic kimberlites. Comparative basalt data are taken from Dosso & Murthy (1980), Cohen & O'Nions (1982), Chen & Frey (1983) and Stille et al. (1983). Other sources of data are: o, San Quintin, Mexico (De Paolo 1978); , San Carlos, Arizona (Zindler & Jagoutz 1987); o, Geronimo, Arizona (Menzies et al. 1985a); O, Pisgah Crater, California (De Paolo 1978); ~ Nunivak, Alaska (Menzies & Murthy 1980a); ,t, Ataq, South Yemen (Menzies & Murthy 1980a); 4, Kiama, New South Wales (Menzies & Wass 1983); ~, Afar, Ethiopia (Betton & Civetta 1984); in, Massif Central, France (Chauvel & J ahn 1984); ~ , Malaita, Solomon Islands (Bielski-Zyskind et al. 1984); 4,, Victoria, Australia (McDonough & McCulloch 1985); ~, Sierran Province, California (Van Kooten et al. 1985); ~, Ross Island and vicinity, Antarctica (Menzies et al. 1985b); V, Tanzania (Cohen et al. 1984); v, Tahiti (White & Hofmann 1982); v, Hawaii (Stille et al. 1983); , , Scotland (Menzies & Halliday 1984; Halliday et al. 1985); *, South Africa (Smith 1983).
Alkaline rocks and their inclusions continental lithosphere is heterogeneous and differs markedly from the oceanic lithosphere. The second point to note concerns the similarity between mantle sampled by volcanic rocks erupted in ocean basins and continental rift valleys (summarized in Fig. 5). Essentially this fragmented and entrained lithosphere has a prehistory not unlike the source regions of alkaline magmas ultimately responsible for the transport of inclusions to the surface. This can be interpreted to mean that lithosphere has been rehomogenized by impingement of asthenospheric melts. Since rift valleys are relatively young (less than 0.2 x 10a Ma) any trace-element enrichment locked into the lithosphere would have had insufficient time to evolve isotopically or would have been obliterated by recent rehomogenization of the isotope systems. The third point has relevance to the presence
+1!
+10 I
+5 I E~
r ANTARCTICA
<~ISCOTLAND
0 I
-5 I
21
of aged isotopically heterogeneous lithosphere below the stable continental cratons. This is sampled during eruption of magmas through stable cratons or in some rift valleys (e.g. Tanzania) (Fig. 5). As a group, ultramafic and mafic inclusions from alkaline and kimberlite host rocks erupted through cratonic regions have a considerable range in Nd and Sr isotopic composition (Fig. 3). These values encompass mantle both enriched and depleted in light REE for a prolonged period of time (Menzies & Murthy 1980b; Erlank et al. 1982, 1987; Kramers et al. 1983; Cohen et al. 1984). Inclusions with the least radiogenic Nd isotopic composition were entrained by kimberlitic and alkaline magmas erupted through old stable cratonic regions (Menzies & Murthy 1980b; Cohen et al. 1984; Menzies & Halliday 1984; Richardson et al. 1984). Diopsides sepa-
-10 I
-15 I
-20 I
-25 I
s
-30 I
Nd
] ] ULTRAMAFIC
O~LSOUTH AFRICA
XENOLITHS
(Lithosphere)
F TANZANIA
1
CALIFORNIA-
i
] COL::A:~N-A NEW MEXICO
I
......
I
El AUSTRALIA ~:
[ EUROPE k S.YEMEN RED SEA GUADALUPE HAWAII
~
~ ~
HOST
MAGMAS
(Asthenosphere)
TAHITI MALAITA NEW ZEALAND
t
i
' ULTRAMAFIC HOST
ALASKA ALPINE
PERIDOTITE 0.5135
I 0.5130
I 0.5125
I 0.5120
1 0.5115
XENOLITHS
MAGMA
I 0.5110
84
I 0.5105
0.5100
143Nd / 144Nd
FIG. 5. Nd isotopic variation in host magmas and mantle inclusions erupted through stable cratons, in rift valleys and ocean basins. Note that, regardless of the tectonic setting, the alkaline host magmas are derived from a mantle source region with a narrow end range of + 8 to -- 7, similar to that observed in oceanic basalts. The inclusions exhibit a considerable range of eNdrelative to the host magmas. Mantle entrained from below stable cratons (e.g. Antarctica, Scotland and South Africa) is heterogeneous, whereas mantle sampled in rift valleys and ocean basins is not unlike MORB-OIB. Rift valleys formed through regions of old crust (e.g. Tanzania), however, appear to sample both types of mantle. Arguments can be put forward for the presence of old heterogeneous lithospheric mantle below stable cratons, which tends to be disrupted and replaced by asthenospheric processes during rift formation. Well established rift valleys would eventually rupture this carapace of aged lithosphere by upwelling of asthenospheric melts. Data are taken from the sources given in the captions to Figs 1-3 and data for Guadalupe and New Zealand (Menzies, unpublished data) and alpine peridotites (Menzies 1984, and references cited therein).
M. Menzies
22
formed by asthenospheric processes. Similarly, in the ocean basins the relatively young age of the crust makes the widespread survival of aged cold stable lithosphere somewhat unlikely.
rated from garnet peridotites have model Nd ages of less than 3.0 x 103 Ma (Menzies et al. 1985b; Erlank et al. 1987). Such ages indicate stabilization and isolation of the continental lithosphere for a considerable portion of the Earth's history. Furthermore, the isotopic variability of the lithospheric mantle overlaps that of the crust. Inclusions with an isotopic signature equivalent to M O R B or OIB have been found in oceanic regions and in continental regions that have experienced recent extension and presumably upwelling (i.e. rift valleys). Inclusions with isotopic ratios distinct from MORB and OIB are sampled by kimberlite or alkaline magmas erupted through old cratonic regions. In the case of stable cratons, the host magma is isotopically equivalent to OIB while the inclusions display a distinct isotopic signature unlike anything recovered from regions of young crust or the ocean basins (Fig. 5). It seems logical to infer that the difference must be that an aged heterogeneous component exists below regions of old crust. During rift valley formation this lithospheric carapace was presumably ruptured and trans-
Models for genesis of alkaline rocks The identification and characterization of the continental lithosphere, tectosphere or keel (Jordan 1975; Brooks et al 1976; Oxburgh & Parmentier 1978) has added a new dimension to our interpretation of the isotopic and traceelement characteristics of volcanic rocks erupted through continental crust. The continental lithosphere is defined as that portion of the upper mantle that moves as a coherent mechanical unit with a velocity similar to the surface crustal plate. Thermal and/or diapiric accretion may be responsible for stabilization of this region of the mantle (Jordan 1978; Oxburgh & Parmentier 1978). The thickness of the lithosphere below continental regions remains the subject of some debate (Jordan 1978; Oxburgh 1981). It is generally
TABLE 2. Source regions of inclusion-bearing alkaline and kimberlitic magmas Locality Afar (Assab) Africa (Tanzania) Alaska (Nunivak) Antarctica Foster Crater Ross Island Arizona Geronimo San Carlos Australia Kiama Kimberley Victoria California (Sierran Province)
Ref.*
Source region
1 Mantle mixing sources MORB and subducted lithosphere 2 Depleted mantle source 3 Mantle sourceOIB 4 Asthenospheric processes=OIB source 5 OIB source 6 OIB source 7 8 9 10
Mantle source modified by CO, MORB-enrichedmantle mixing Sub-continental heterogeneous mantle Subduction component in mantle-micaceous source
Fiji France (Massif Central) French Polynesia (Tahiti) Hawaii (Honolulu Series) Nigeria (Bokkos) Scotland (N Highlands) Solomon Islands (Malaita) South Africa (Various localities) South Yemen (Ataq)
11 OIB pods or dykes in MORB 12 Metasomatizedmantle source 13 Mantle source with recycled component 14 Multistage history of lithosphere-asthenosphere mixing 19 Fractionation in mantle, crustal contamination of evolved rocks 15 Old enriched sources: lithosphere-asthenosphere mix 16 Young depleted mantle plus recycled component 17 Lithosphere (enriched)--asthenosphere (depleted) sources 18 Mantle heterogeneity
* References 1-18 as in Table 1 ; 19, Irving & Price 1981.
Alkaline rocks and their inclusions accepted, however, that the sub-continental lithospheric mantle is considerably thicker than the crust and is of the order of 100 km or more compared with tens of kilometres beneath ocean basins. While the lithosphere below oceans is relatively young (less than 0.2x 103 Ma) and geochemically uncomplicated, the continental lithosphere retains a complex history that has been preserved in the stable non-convecting environment below the cratons. It has been suggested (Oxburgh & Parmentier 1978) that the lithosphere records the last major thermal event in the overlying crust. If so, aged heterogeneous mantle should be restricted to regions of Proterozoic and Archean crust. The prolonged stability of this region throughout geological time has important implications for the geochemical evolution of the sub-continental mantle. Brooks e t al. (1976) proposed that the keel soaked up fluids and melts from the asthenosphere and as a consequence the subcontinental lithosphere may contain a significant fraction of the incompatible-element inventory of the Earth. In theory, upwelling of silicate melts and aqueous fluids from the asthenosphere will modify the trace-element inventory of the lithospheric mantle. These enhanced trace-element
23
ratios will eventually register as modified isotopic ratios in regions of the mantle left undisturbed for considerable periods of time (more than 103 Ma). Extremely heterogeneous mantle will be produced by such processes. Any study of the evolution of continental alkaline (and tholeiitic) magmas must acknowledge the presence of aged heterogeneous lithosphere that may attain thicknesses in excess of 200 km below continental regions. Any magma believed to have been generated in the asthenosphere or lower mantle must first rupture this carapace of lithospheric mantle and then the crust, and establish by magma fracture a network of conduits to the surface. Since lithospheric mantle will be considerably thicker than crust, it presents a potential source of 'mantle contamination' for the asthenospheric melts. Mixing of asthenospheric and lithospheric melts may be a more significant process than crustal interaction in determining the isotopic and chemical characteristics of alkaline magmas (Table 2). Group II kimberlites and lamproites have Sr and Nd isotopic ratios distinctly different from those of OIB and more akin to those of crustal rocks or ultramafic xenoliths disrupted from the mantle below regions of old crust (Fig. 6).
I
~Nd
'
....MORB ~---~ASTHENOSPHERE 0.5130
~
R
-+10
BASALTS
m
0
OUP i K]-MBERLITES GROUP II KIMBERLITES
"O
z
"-~. 0 . 5 1 2 0
-
-10
-
-20
-
-30
LAMPROITES ~ "" ~
"O
z ,r-
'
[LITHOSPHERE I MANTLE INCLUSIONS
\MANTLE INCLUSIONS 0.5110 I,
I
I
I
0700
0750
0~00
0.850
87Sr/86Sr FIG. 6. t43Nd/~44Ndversus 878r/S6Sr variation in mantle inclusions and inclusion-bearing alkaline rocks. The extreme range in Nd and Sr isotopic composition defined by minerals separated from ultramafic inclusions encompasses the variation in Nd and Sr isotopic composition observed in alkaline rocks. Alkaline basalts and Group I kimberlites have isotopic characteristics identical with those of OIB derived from asthenosphere. In contrast, lamproites and Group II kimberlites may have a lithospheric component in their source regions. Data are taken from Menzies & Murthy (1980a, b), Erlank et al. (1987), Menzies (1983, and references cited therein), Menzies & Wass (1983), Cohen et al. (1984) and Menzies & Halliday (1984). Other data are taken from the references given in the caption to Fig. 4.
z4
M. Menzies
Contamination with significant amounts of crust is unlikely because of the rapidity with which these melts must reach the surface in order to maintain their load of high-density inclusions and diamonds. Furthermore, the extremely high concentrations of Sr (678-1845 ppm) and Nd (48-336 ppm) would require mixing with a highly radiogenic crustal component or stoping and assimilation of considerable amounts of crust. This is believed to be very unlikely. Now the problem arises of exactly which structural elements in the sub-continental mantle contribute to the genesis of these rocks. The isotopically heterogeneous nature of lamproites compares well with the heterogeneities recorded in continental lithospheric mantle inclusions (Fig. 6). Arguments could therefore be presented for a source region solely within the lithospheric mantle. This is rather unlikely as an asthenospheric component or influence may be needed to provide heat and volatiles and to trigger anatexis. Impingement of an asthenospheric plume on the continental lithosphere is believed to elevate the isotherms and induce melting in the lithospheric mantle immediately above the plume. The column of mantle above the plume may display a considerable range of Sr, Nd and Pb isotopic composition. To illustrate this it is
worth noting that ultramafic inclusions from the Bultfontein pipes, S Africa (Fig. 3), have an extreme range in Sr and Nd isotopes. If we accept that these inclusions represent a random vertical profile through the lithosphere, melts extracted from such a heterogeneous mantle would tend to be distinctly different from OIBs and the bulk of continental basalts particularly with regard to their isotopic ratios. If lamproites and micaceous kimberlites contain a major contribution from the lithospheric mantle it follows that their isotopic variability could theoretically be as great as that observed in lithospheric inclusions (Fig.
6). As more data become available, the range of Sr, Nd and Pb isotope ratios observed in mantle rocks is continually revised. Therefore a unique solution to the origin of lamproites and Group II kimberlites is impossible. It is perhaps prudent to limit comments on their origin to interaction between relatively homogeneous asthenosphere and extremely heterogeneous aged lithospheric mantle. The observed isotopic ratio in erupted lamproite or micaceous kimberlite will depend on (1) the mixing proportions of asthenosphere and lithosphere, (2) the Sr/Nd ratio in the asthenospheric and lithospheric end-members and (3) the isotopic ratio of these end-members.
FIG. 7. Earth models pertinent to the origin of inclusion-bearing alkaline rocks (IBAR). The primordial hotmantle plume (PHMP) and depleted low velocity layer (DLVL) model of Schilling (1973) and the model of Tatsumoto (1978) are based on REE and Pb isotopes. The third model is modified after Wasserburg & De Paolo (1979) and incorporates ideas from McCulloch et al. (1983), Smith (1983) and Chen& Frey (1983). The most significant aspect of this model is the involvement of the continental lithospheric mantle (CM) in the genesis of lamproites and Group II kimberlites. Wasserburg & De Paolo (1979) believe the CM to be 'sterile', but this is not compatible with geochemical data on inclusions. Continental flood basalts may have a contribution from CM early in their evolution while upwelling of asthenosphere triggers melting in the lithosphere and encourages mixing of the melts. Eventually, once the volcanic conduits are well established, continental flood basalts will essentially have the isotopic signature of OIBs. (CC, continental crust; OC, oceanic crust; IBAR, inclusion bearing alkaline rocks.)
Alkaline rocks and their inclusions McCuUoch et aL (1983) pursued a somewhat similar line of reasoning in their interpretation of the isotopic and trace-element variation observed in Australian lamproites. Mixing of a depleted end-member, not unlike OIB in isotopic composition, and an enriched mantle end-member, not unlike average crust in isotopic composition, produced the required range of St and Nd isotopes observed in the diamond-bearing lamproites. Similarly, in a detailed study of Group I (basaltic) and Group II (micaceous) kimberlites, Smith (1983) outlined a model of asthenosphere-lithosphere interaction. Since Group I kimberlites are isotopically identical with OIBs, their evolutionary history was believed to involve derivation from asthenospheric sources with little or no lithospheric mantle input. In contrast, Group II kimberlites require a more dominant lithospheric mantle component because of their radiogenic Sr and non-radiogenic Nd isotopic composition. McCulloch et al. (1983) defined a unique composition for the enriched mantle end-member. As more and more data become available, it is obvious that the sub-continental lithospheric mantle is extremely heterogeneous and could provide a variety of end-members for mixing models. To use the terminology of Powell and Bell (1974), several suites of garnet peridotites and pyroxenites stable at mantle pressures and temperatures have 'crustal-type' St, Nd and Pb isotopic compositions. As a consequence of this, the isotopic variability in the continental mantle is almost as diverse as in the overlying crust,
25
perhaps owing to incorporation of subducted melts or fluids over billions of years. Continental lithospheric mantle is in part transitional between M O R B - O I B type mantle and the crust. In its trace-element geochemistry and ultimately in its isotope geochemistry certain portions of the lithospheric mantle have more similarity to the crust than to M O R B - O I B mantle. It remains to be seen whether this is due to diapiric upwelling related to subduction of continental material or to other recycling processes.
Summarizing comments Impingement of asthenospheric or lower mantle plumes on oceanic or continental lithosphere is believed to initiate anatexis and promote mixing of isotopically distinct melts. Interaction with MORB-like lithosphere in ocean basins and continental regions gives rise to inclusion-bearing alkaline magmas and basaltic kimberlites (Group I) with isotopic signatures not unlike M O R B and OIB ( e N d = 8 to -- 9 ; eSr = -- 28 to + 34) (Table 2). Interaction with aged enriched continental lithospheric mantle may produce a more diverse suite of melts with isotopic characteristics unlike M O R B - O I B and more like that of ultramafic inclusions entrained in alkaline and kimberlitic melts (eNd= -- 24 to + 15 ;/3Sr= - - 35.5. to + 1867). Magmas believed to be produced by this process are micaceous kimberlites (Group II) and certain lamproites (Table 2).
References BETTON,P. J. & CIVETTA,L. 1984. Sr and Nd isotopic evidence for the heterogeneous nature and development of the mantle beneath Afar (Ethiopia). Earth planet. Sci. Lett. 71, 59-70. BIELSKI-ZYSKIND,M., WASSERBURG,G. J. & NIXON,P. H. 1984. Sm-Nd and Rb-Sr systematics in volcanics and ultramafic xenoliths from Malaita and Solomon Islands and the nature of the Ontong Java Plateau. J. geophys. Res. 89, 2415-24. BROOKS,C., JAMES,D. E. & HART,S. R. 1976. Ancient lithosphere: its role in young continental volcanism. Science, 193, 1086-994. CHASE, C. G. 1981 Oceanic island Pb: two-stage histories and mantle evolution. Earth planet Sci. Lett. 52, 277-84. CHAUVEL,C. & JAHN, B.-M. 1984. Nd-Sr isotope and REE geochemistry of alkali basalts from the Massif Central, France. Geochim. cosmochim. Acta, 48, 269-78. CHEN, C.-Y. & Frey, F. A. 1983. Origin of Hawaiian tholeiite and alkalic basalt. Nature, Lond. 302, 785-9.
CLAGUE, D. & FREY, F. A. 1982. Petrology and trace element geochemistry of the Honolulu volcanics, Oahu: implications for the oceanic mantle below Hawaii. J. Petrol. 23, 447-504. COHEN, R. S. & O'NIONS, R. K. 1982. The lead, neodymium and strontium isotopic structure of ocean ridge basalts. J. Petrol. 23, 299-324. , & DAWSON,J. B. 1984. Isotope geochemistry of xenoliths from East Africa: implications for development of mantle reservoirs and their interaction. Earth planet. Sci. Lett. 68, 209-20. DE PAOLO,D. 1978. Nd and Sr isotopic systematics of young continental igneous rocks. Short Papers of the 4th ICGCIG. Geological Survey Open File Report 78-701, pp. 91-3. United States Department of the Interior, Geological Survey, U.S. Printing Office. Dosso, L. & MURTHY, V. R. 1980. An Nd isotopic study of the Kerguelen islands: inferences on enriched oceanic mantle sources. Earth planet Sci. Lett. 48 268-76. DOWNES, H. 1984. Sr and Nd isotopic geochemistry of co-existing alkaline magma series, Cantal massif
26
M. Menzies
Central, France. Earth planet Sci. Lett. 69, 221334. ERLANK, A. E., ALLSOPP, H., HAWKESWORTH, C. J. & MENZIES, M. A. 1982. Chemical and isotopic characterization of upper mantle metasomatism in peridotite nodules from the Bultfontein kimberlite. Terra Cognita 2, 261-3. --, WATERS, F. G., HAWKESWORTH, C. J., HAGGERTY, S. E., ALLSOPP, H. L., RICKARD, R. S. & MENZIES, M. A. 1987. Evidence for mantle metasomatism in peridotite nodules from the Bultfontein floors, Kimberley, S. Africa. In : MENZIES, M. A. & HAWKESWORTH, C. J. (eds) Mantle Metasomatism, Academic Press, New York, 221311. FREY, F. A. & GREEN, D. H. 1974. The mineralogy, geochemistry and origin of lherzolite inclusions in Victorian basanites. Geochim. cosmochim. Acta, 38, 1023-59. GILL, J. B. 1984. Sr-Pb-Nd isotopic evidence that both MORB and OIB sources contribute to ocean island arc magmas in Fiji. Earth planet. Sci. Lett. 68, 44358. HALLIDAY,A. N., MENZIES,M. A. PALACZ,Z., HUNTER, R. H. & HAWKESWORTH, C. J. 1985. Hebridean mantle geochemistry. Trans. Am. Geophys. Union, 66, 414. HAWKESWORTH, C. J., ROGERS, N., VAN CALSTEREN, P. W. C. & MENZIES, M. A. 1984. Mantle enrichment processes. Nature, Lond. 311, 331-55. HOFMANN, A. & WrotE, W. M. 1982. Mantle plumes from ancient oceanic crust. Earth planet Sci. Lett. 57, 421-36. IRVING, A. J. & FREY, F. A. 1984. Trace element abundances in megacrysts and their host basalts: constraints on partition coefficients and megacryst genesis. Geochim. cosmochim. Acta, 45, 1201-21. & PRICE, G. 1981. Geochemistry and evolution of Iherzolite-bearing phonolitic lavas from Nigeria, Australia, East Germany and New Zealand. Geochim. cosmochim. Acta, 45, 1309-20. JAGOUTZ, E., CARLSON, R. W. & LUGMAIR, G. 1980. Equilibrated Nd, unequilibrated Sr isotopes in mantle xenoliths. Nature, Lond. 286, 708-10. JORDAN, T. H. 1975. The continental tectosphere. Rev. Geophys. space Phys. 13, 1-12. - 1978. Composition and development of the continental tectosphere. Nature, Lond. 274, 544-8. KRAMERS, J., RODDICK, J. & DAWSON, J. 1983. Trace element and isotope studies on veined, metasomatic and 'MARID' xenoliths from Bultfontein, South Africa. Earth planet. Sci. Lett. 65, 90-106. - - , SMITH, C. B., LOCK, N. P., HARMON, R. S. & BOYD, F. R. 1981. Can kimberlites be generated from ordinary mantle? Nature, Lond. 291, 53-6. KYLE, P., KIRSCH, I., GAMBLE, J. & MENZIES, M. 1985. Metasomatised ultramafic xenoliths from Foster Crater, Antarctica. Trans. Amer. geophys. Union 66, 409. MCCULLOCH, M. T., JAQUES, A. L., NELSON, D. R. & Lewis, J. D. 1983. Nd and Sr isotopes in kimberlites and lamproites from Western Australia: an enriched mantle origin. Nature, Lond. 302, 400-3. MCDONOUGH, W. F. & MCCULLOCI-I, M. T. 1985.
Isotopic and geochemical systematics in TertiaryRecent basalts from southeastern Australia and implications for the evolution of the sub-continental lithosphere. Geochim. cosmochim. Acta 49, 205168. MCKENZlE, D. P. & O'Nions, R. K. 1983. Mantle reservoirs and ocean island basalts. Nature, Lond. 301, 229-31. MENZlES, M. A. 1983. Mantle ultramafic xenoliths in alkaline magmas: evidence for mantle heterogeneity modified by magmatic activity. In: HAWKESWORTH, C. J. & NORRY, M. J. (eds) Continental Basalts and Mantle Xenoliths, pp. 111-38. Shiva, U.K. - - 1 9 8 4 . Chemical and isotopic heterogeneities in orogenic lherzolite massifs. In: GASS, I. G., LIPPARD, S. J. & SrIELTON, A. W. (eds), Ophiolites and Oceanic Lithosphere, Geol. Soc. Spec. Pubi. 13, pp. 231-40. - & HALLIDAY,A. N. 1984. Enriched mantle below the Scottish Highlands: Sr and Nd isotope and rare earth elements in peridotite and pyroxenite xenoliths from Loch Roag, Streap Comlaidh and Kilchatten, Scotland. Abstracts of 27th Int. Geological Cong., Moscow, p. 348. - - & HAWKESWORTH, C. J. 1987. Mantle xenoliths and upper mantle structure. In: NIXON, P. (ed.) Mantle Xenoliths, Wiley, New York, in press. --& MURTHY, V. R. 1980a. Nd and Sr isotope geochemistry of hydrous mantle nodules and their host alkali basalts: implications for local heterogeneities in metasomatically veined mantle. Earth planet Sci. Lett. 46, 323-34. - - & 1980b. Enriched mantle: Nd and Sr isotopes in diopsides from kimberlite nodules. Nature, Lond. 283, 634-6. --& WASS, S. Y. 1983. COz rich mantle below eastern Australia: REE, Sr and Nd isotopic study of Cenozoic alkaline magmas and apatite-rich xenoliths, Southern Highlands Province, New South Wales, Australia. Earth planet. Sci. Lett. 65, 287-302. - - , KEMPTON, P. & DUNGAN, M. 1985a. Interaction of continental lithosphere and asthenospheric melt below the Geronimo Volcanic Field, Arizona, U.S.A.J. Petrol., 26, 663-93. - - , KYLE, P. & GAMBLE,J. 1985b. Fragments of aged enriched sub-continental lithosphere entrained in asthenospheric magmas erupted at Foster Crater, Antarctica. Trans. amer. geophys. Union, 66, 409. - - - , ROGERS, N., TINDLE, m. & HAWKESWORTH,C. J. 1987. Metasomatic and enrichment processes in lithospheric peridotites, an effect of asthenosphere-lithosphere interaction. In: MENZlES, M. A. & HAWKESWORTI.I,C. J. (eds) Mantle Metasomatism, Academic Press, New York, 313-61. OXBURGH, E. R. 1981. Heat flow and differences in lithospheric thickness. Phil. Trans. R. Soc. Lond., Ser. A, 310, 337-46. - & PARMENTIER,E. M. 1978. Thermal processes in the formation of continental lithosphere. Phil. Trans. R. Soc. Lond., Set. A, 288, 415-29. POWELL, J. L. & BELL, K. 1974. Isotopic composition of strontium in alkalic rocks. In: SORENSEN, H.
Alkaline rocks and their inclusions (ed.) The Alkaline Rocks, pp. 412-420. Wiley, New York. RICHARDSON, S. H., GURNEY, J. J., ERLANK, A. J. & HARRIS, J. W. 1984. Origins of diamonds in old enriched mantle. Nature, Lond. 310, 198-202. ROCK, N. M. S. 1976. The comparative strontium isotopic composition of alkaline rocks: new data from southern Portugal and East Africa. Contrib. Mineral. Petrol. 56, 205-28. RODEN, M. F., FREY, F. A. & FRANCIS, D. i . 1984a. An example of consequent mantle metasomatism in peridotite inclusions from Nunivak Island, Alaska. J. Petrol. 25, 546-77. - - , HART, S. R., FREY, F. A. & MELSON, W. G. 1984b. Sr, Nd and Pb isotopic and REE geochemistry of St. Paul's Rocks: the metamorphic and metasomatic development of an alkali basalt mantle source. Contrib. Mineral. Petrol. 85, 379400. SCHILLING,J-G. 1973. Iceland mantle plume: geochemical evidence along the Reykjanes Ridge. Nature, Lond. 242, 565-571. SMITH, C. B. 1983. Pb, Sr and Nd isotopic evidence for sources of southern African Cretaceous kimberlites. Nature, Lond. 304, 51-4. STILLE, P., UNRUH, D. & TATSUMOTO,M. 1983. Pb, Sr, Nd and Hf isotopic evidence of multiple sources for Oahu, Hawaii basalts. Nature, Lond. 304, 259. STOSCH, H. G., CARLSON, R. W. & LUGMAIR, G. W. 1980. Episodic mantle differentiation: Nd and Sr isotopic evidence. Earth planet. Sci Lett. 47, 26371.
27
STUCKLESS, J. S. & ERICKSEN, R. L. 1976. Strontium isotope geochemistry of the volcanic rocks and associated megacrysts and inclusions from Ross Island and vicinity, Antarctica. Contrib. Mineral. Petrol. 58, 111-26. TATSUMOTO, M. 1978. Isotopic composition of lead in oceanic basalts and its implication to mantle evolution. Earth planet. Sci. Lett. 38, 63-87. THIRLWALL,i . 1982. Systematic variation in chemistry and Nd-Sr isotopes across a Caledonian calcalkaline volcanic arc: implications for source materials. Earth planet. Sci. Lett. 58, 27-50. VAN KOOTEN, G., LEEMAN, W. P. & MENZIES, M. A. 1985. Lithospheric peridotites and pyroxenites in alkaline basaltic rock from the Sierra Nevada, California, U.S.A. Trans. Amer. Geophys. Union, 66, 414. VIDAL, P., CHAUVEL, C. & BROUSSE, R. 1984. Large Mantle Heterogeneity beneath French Polynesia. Nature, Lond. 307, 536-8. WASS, S. Y. & ROGERS, N. K. 1980. Mantle metasomatism--precursor to continental alkaline volcanism. Geochim. cosmochim. Acta, 44, 1811-23. WASSERBURG, G. J. & DE PAOLO,D. J. 1979. Models of earth structure inferred from neodymium and strontium isotopic abundances. Proc. Natl Acad Sci. 76, 3594-8. WHITE, W. M. & HOFMANN, A. W. 1982. Sr and Nd isotope geochemistry of oceanic basalts and mantle evolution. Nature, Lond. 296, 821-5. ZINDLER, A. & JAGOUTZ, E. 1987. Trace element and Nd and Sr isotope systematics ofperidotite nodules from Peridot Mesa San Carlos, Arizona. Geochim. cosmochim. Acta, in press.
MARTIN MENZIES, Department of Geology, University of London, Royal Holloway and Bedford New College, Egham, Surrey, TW20 0EX, U.K.
The genesis of alkaline magmas with emphasis on their source regions: inferences from experimental studies A. D. Edgar SUMMARY: Experimental studies bearing on the genesis of alkaline magmas are reviewed, with emphasis on alkaline mafic-ultramafic magmas and the nature of their mantle sources. In contrast with previous experiments on simplified systems mainly at low pressures, more recent experiments under mantle pressures have resolved some of the problems of the earlier work. Inadequacies of petrogenetic schemes based on the model pyrolite mantle source for the generation of alkaline magma are also partly resolved by recent studies which suggest that with increasing alkalinity the possibility of deriving mafic-ultramafic magmas from a dry lherzolitic mantle source, or one with H20 as the only volatile, becomes increasingly difficult. The addition of CO2 and H20 to the source regions allows generation of magmas such as olivine melilitites, but few magmas of the olivine leucitite kindred can be generated from lherzolitic sources. Such magmas appear to be partial melts of clinopyroxenite or orthopyroxenite sources containing both volatiles. Preliminary experimental modelling of mantle metasomatism suggest that this process may readily occur. Experiments on felsic alkaline rocks with emphasis on liquid immiscibility and halogens and experiments aimed at solving specific problems are briefly considered.
Introduction Experiments on systems pertinent to alkaline rocks have undoubtedly made an immense contribution to our understanding of the genesis of these rocks, particularly in problems such as the attainment of peralkalinity, the depth of derivation of alkaline magmas and, most recently, the variation in the source materials from which alkaline magmas are derived relative to nonalkaline varieties. Some of the major experimental studies carried out since the publication of The Alkaline Rocks by Sorensen (1974) are reviewed in this paper. The term 'alkaline rocks' will be used to denote rocks containing feldspathoidal minerals (modal or normative) or rocks not necessarily containing a feldspathoid but with low SiOz and high alkali contents. Such a definition clearly precludes alkali granites but could include certain kimberlites and carbonatites which are not discussed in any detail. Within the past decade a number of reviews of the role of experimental petrology in our understanding of the genesis of alkaline rocks have been published. Edgar (1974) included data on alkaline systems published up to the early 1970s, while Bailey (1974) reviewed the application of experimental petrology to oversaturated peralkaline volcanic rocks and later (Bailey 1976) discussed experiments applicable to a wide spectrum of alkaline rocks ranging from strongly SiO2-undersaturated mafic varieties to undersaturated and oversaturated felsic types. More recently, the Basaltic Volcanism Study Project
(1981) (see particularly pp. 529-30, 561-2 and 564--7) has reviewed experiments on alkaline and highly undersaturated potassic basalts. Zyryanov (1981) discussed phase relations in feldspars and feldspathoids with some emphasis on petrological aspects, and Kogarko & Romanchev (1982) discussed phase equilbria in alkaline melts. Both of these papers include important references to the Russian literature. Writing from a petrologist's viewpoint, Edgar (1984) discussed the chemistry, occurrence and paragenesis of feldspathoids, including experiments bearing on the origin of pseudoleucites, the formation of melilite in nephelinites, melilitites and alkaline lavas of comparable compositions, geothermometry and the question of the origins of primary analcite. In the past 10-15 years the emphasis of experimental studies bearing on alkaline rock genesis has changed from low-pressure experiments in relatively simple synthetic systems to studies using rock compositions under mantle conditions in excess of 100 km depth and controlled partial pressures of volatiles. As a result of these high-pressure experiments the importance of H20, CO 2 and other volatiles on the genesis and evolution of alkaline magmas is much better known than it was a decade ago. The hypothesis that alkaline magmas as well as kimberlites, carbonatites and lamproites are partial melts of slightly to vastly modified lherzolite mantle sources has been strongly supported by experimental studies and has led to experiments in which processes of mantle metasomatism are being simulated. Investigations
From : FITTON,J. G. & UPTON, B. G. J. (eds), 1987, Alkaline Igneous Rocks, Geological Society Special Publication No. 30, pp. 29-52.
29
30
A. D. Edgar
have been undertaken to determine the compositions of liquids produced from partial melting of metasomatized mantle nodules. Results of such experiments, combined with major-element, trace-element and isotope geochemistry, have changed many of the models based on a pyrolitetype mantle source proposed in the late 1960s and early 1970s (e.g. O'Hara 1968; Green 1971) for basalt genesis, and particularly for the origin of alkali basalt magmas. Many experiments have also been carried out on the structures of silicate melts, many of alkaline affinities, which have important implications on the physical properties of magmas. Although these studies are beyond the scope of this paper, they are now being incorporated into petrology texts (e.g. Barker 1983). Other important advances in experimental petrology pertinent to alkaline rocks include studies on liquid immiscibility, the role of halogens and traceelement partitioning between crystals and melts. This review concentrates on experimental studies pertinent to the genesis of alkaline basalts and mafic rocks of extreme alkaline affinities. The implications of these studies on the mantle source regions of alkaline magmas are emphasized. Experiments at lower pressures pertinent to felsic alkaline magmatism are briefly considered.
Alkali mafic and ultramafic magmas--background studies Synthetic systems at low pressures Alkali basalts, in which potassium is a minor element, can be represented in the expanded basalt system (CaSiO3 (Wo)-NaA1SiO4 (Ne)Mg2SiO4 (Fo)-SiO2) by recalculating the rock composition into the CIPW normative minerals representing this system. In this system (Fig. 1) mildly alkaline basalts and associated rocks plot within the D i - N e - F o - A b volume and are separated from the more alkali-rich nephelinites and melilitites, plotting within the A k - D i - F o - N e volume, by the F o - D i - N e pseudoternary join. Experiments at low pressure within these two sub-systems of the expanded basalt tetrahedron and in systems with additional components such as CaA12Si20 8 (An) resulted in discrepancies in the paragenesis of the minerals, based on the interpretations of the experimental results and on textural observations on mildly alkali-enriched to highly alkali-enriched rocks. Part of this problem was due to treating the expanded basalt system on flow sheets as a quaternary system
Lo
Ne~
Fo
Mol per cent
FIG. 1. The expanded basalt system. Abbreviations for the CIPW normative minerals are as follows: La, larnite; Ne, nepheline; Fo, forsterite; Qz, quartz; Ab, Albite; En, enstatite; Di, diopside ; Ak, akermanite; Wo, wollastonite; Sm, soda-melilite; Mo, monticellite; Mer, merwinite; Ra, rankinite. Shaded joins represent planes of critical SiO2 undersaturation. when it is at least quinary (Schairer et al. 1969; Schairer & Yoder 1970). For similar reasons, the CaO-MgO-AI203-SiO 2 system (O'Hara & Biggar 1969; Schairer et al. 1969; Schairer & Yoder 1970) cannot be used to model alkaline mafic magmas. The obvious chemical deficiencies of the CaO-MgO-A1203-SiO2 system relative to the compositions of nephelinites and melilitites, and the now accepted hypothesis that nephelinites and melilitites form from the mantle-derived magmas at considerable depth make experiments at 1 atm in either the expanded basalt tetrahedron or the CaO-MgO-AI203-SiO 2 system quite inadequate for investigating the genesis of these rocks ( c f Thompson 1972). Although low-pressure experiments are largely inappropriate in explaining the genesis of alkaline magmas, some features such as the incompatibility of melilite and plagioclase in the same lava can be accounted for by using the schematic representation of liquidus volumes in the WoD i - A b - N e sub-system of the expanded basalt tetrahedron (Fig. 2). Data for this diagram are given by Yoder (1979). Along the Di-Ne join in this system diopside and nepheline do not coexist as liquidus phases but are separated by a field of
Genesis o f alkaline magmas olivine (Bowen 1922: Schairer et al. 1962; Yoder & Kushiro 1972). At sub-liquidus temperatures melilites with compositions comparable with those of melilite in melilitites and melilite nephelinites (E1-Goresy & Yoder 1974) occur. Extension of the primary olivine field into the D i - N e - W o - A b volume (Fig. 2) separates the primary fields of melilite and plagioclase and accounts for the incompatibilityof these minerals in nature. The stability of melilite molecules in various sub-systems of the expanded basalt tetrahedron and in other systems in which Ne and Di are components has been used by Yoder (1979) to explain the genesis of melilite-bearing rocks and associated lamprophyres. The limitations of the expanded basalt and CaO-MgO--A1203-SiO2 systems to explain the genesis of Na-rich mafic magmas become more pronounced for K-rich varieties of these magmas. Although not as common, K-rich mafic and ultramafic lavas may be more important than Na-rich varieties in holding the key to heterogeneities in the mantle source regions. Yoder (1975) has modelled these K-rich magmas using the system Ca2SiO 4 (La)-KA1SiO4 (Ks)Mg2SiO4 (Fo)-SiO2_H20. Figure 3(a) shows the relationship of leucite and akermanite to other major constituents of these rocks, whereas in Fig. 3(b) the system with H20 as an additional
31 Ks
o
LO
Met
Mo
Fo
Mol per cent
Ks
(b) Ne
Sm
Lo-
, Pl ~'"
wo wo
Ab
D. FIG. 2. The CaSiO3 (Wo)-CaMgSi206 (Di)NaA1Si308 (Ab)-NaA1SiO4(Ne)system. Abbreviations are as in Fig. 1 and as follows: Mel, melilite; O1, olivine; Pl, plagioclase; SS, solid solution. (After Yoder 1979.)
Mb
Mo Mol per cent
Fo
FIG. 3. The system Ca2SiO4 (La)-KA1SiO4 (Ks)MgzSiO4(Fo)-SiO2(Qz) (a) without H20 and (b) with H20. Abbreviations are as in Fig. 1 and as follows: Lc, leucite; Sa, sanidine; Phlog, phlogopite. (After Yoder 1979.) component has phlogopite, a common constituent of K-rich mafic lavas, as an additional phase. Many joins in this system have been investigated up to high pressures and, despite the obvious lack of Na20, the results provide some explanations for the genesis of K-rich alkaline mafic rocks. Using this system, Yoder (1975) related ugandite (olivine melilite leucitite) to mafurite (olivine kalsilite pyroxenite) by the reaction
A. D. E d g a r
32
2[Ca2MgSi207] (Ak) 4- 3[KA1Si206] (Lc) 4- (x 4-1)[MgzSiO4] (Fo) ugandite
4[CaMgSi206] (Di) + 3[KA1SiO4] (Ks) + x[MgzSiO4] (Fo) mafurite
As shown in Fig. 3(a)), the leucite + akermanite assemblage (bold lines) reacts to give the kalsilite + diopside assemblage (broken line). Both ugandite and mafurite contain olivine. If the reaction based on relationship in Fig. 3(a) is correct the simplified ugandite assemblage must be richer in olivine than the mafurite assemblage as indicated by the unknown value x. Highpressure experiments on these rocks (Edgar et al. 1980; Arima & Edgar 1983a) suggest that these magmas are not interrelated but may be derived independently from different metasomatized mantle sources. Using the relationship in Fig. 3(b) and adding COz, Yoder (1975) indicated a possible relationship between mafurite magmas and kimberlite magmas by the reaction CaMgSi206 (Di) + 3[KAISiO4] (Ks) 4- (x + 4) [Mg2SiO4] (Fo) 4- CO2 4- 3H20 rnafurite
3[KMg3A1Si3Olo(OH)2] (Phlog) + CaCO3(Ct) + x[Mg2SiO4] (Fo) kimberhte
In this reaction x represents the amount of olivine required to balance the equation and indicates that a kimberlitic groundmass will have less olivine than a simplified mafurite. On the basis of these equations, Yoder (1975) proposed that melilite-bearing magmas should be transformed to kimberlitic magmas. Melilite-bearing magma might be produced at low pressures under dry conditions (Fig. 3(a)) caused by volatile loss during ascent or by low-pressure crystallization. With the addition of H20 and CO2, such a melilite magma could be transformed to kimberlite. Although the kimberlite-melilitite association is not uncommon, kimberlites are only rarely associated with highly K-rich mafic volcanics such as mafurite. However, kimberlites are associated with lamproites sensu stricto (Bergman 1987) but probably not by the mechanism proposed by Yoder (1975). Many mafurites contain both kalsilite and primary phlogopite (Holmes 1950; Edgar 1979) which cannot be explained by this mechanism.
Rock systems--the pyrolite model As an alternative to the chemically simplified synthetic systems, experiments on rocks of alkali
olivine basalt, basanite, nephelinite and melilitite compositions (el Green & Ringwood 1967; Bultitude & Green 1968, 1971; Green 1969, 1970a, b, 1973a; Green & Hibberson 1970; Brey & Green 1977) under conditions ranging from crustal to upper-mantle depths clearly indicated the importance of mantle processes in the genesis of alkaline mafic magmas. Based on these studies and assuming a model pyrolite mantle source for these magmas (Ringwood 1975), Green (1970a, b) developed a petrogenetic grid for basaltic magmas (Fig. 4). For a model pyrolite with 0.1 H20 and no other volatile component, Fig. 4 shows that progressively more alkali-rich mafic to ultramafic magmas, ranging from alkali olivine basalts to olivine melilitites, might be produced at progressively greater depths with decreasing degrees of partial melting and increasing amounts of H20 in the melt. The grid shown in Fig. 4 can be used to divide fractionation trends, based on the liquidus and near-liquidus minerals, according to depth of generation of the magma. Green & Ringwood (1976), Green & Hibberson (1970) and Merrill & Irving (1977) proposed that mildly alkali basalts could be formed at approximately 45-60 km by fractionation of aluminous enstatite 4- aluminous subcalcic clinopyroxene from an olivine basalt magma to yield alkali olivine basalt magma with about 25~ normative olivine and 1 ~ - 2 ~ normative nepheline (Fig. 4). Green & Ringwood (1976) extended this fractionation scheme by suggesting that further fractionation of aluminous clinopyroxene_ aluminous enstatite from this magma can produce an olivine basanite with about 5~ normative nepheline. The absence of orthopyroxene as a liquidus phase at any pressure under dry conditions in alkali olivine basalt compositions (Ito & Kennedy 1968; Bultitude & Green 1971; Arculus 1975) makes this scheme highly unlikely. At greater depths (60-120 km) no fractionation trends linking the more extreme alkali magmas of nephelinitic and basanitic affinities could be found on the basis of results of experiments in the absence of H20. However, hydrous experiments on nephelinitic compositions with an estimated 5-7~ H20 in the melt (Bultitude & Green 1968, 1971; Green 1969, 1970a, b, 1973b; Green & Hibberson 1970) indicated that between 18 and 27 kb the primary field of orthopyroxene crystallization was extended, in contrast with anhydrous experiments in which liquidus phases were olivine, clinopyroxene and garnet. These experiments suggested a fractionation scheme whereby less extreme alkali basaltic magmas could be linked to more alkali-rich compositions, provided that H20 was available in their source regions. This scheme has
Genes& of alkaline magmas
33
% MELTING 5
I0
15
20
25
Qz.' Tholeiite Tholeiite/ ~ -----.__ 0-5 / i '
olivine
/,
,
~ ~
~
C\~e'/Oliv in 9 Th ole iit 9
\
5-10
?\~High AI203
.
20
Or tho'~vr ~
9OXene~
High AI2I~3
Hawc(iite"/"Alkaline Olivine
35
Tholeiite
5-15
High-AI203 e~ ~
30
Olivine
Olivine
Olivine Tholeiite
Tholeiite
~" 40
Olivine-rich ~-/~<~Y
.13
So.sanite
/
/
Alkaline Olivine Bo.satt 20 - 30
~s-25 / / ~O~ine Olivine-rich NL:phelinite Bosonite 2o-3o
@ L
D20
//2o-2s
U~ @
Otiv/ne Basalt 2~-25
5O
Otivine Thoteiite 20-25
Alkc{t'ine Picrite ~ 3o-35
Thoteiitic Picrite -3o
L
n
-~344v].Qe ~ M e t i t i t e ~ -~ Olivine Nepheti~te i . . o,~ ~30-35 / 3O Nephetln~te %, 25-35
I O,iv,ne
J~lMel.ilitit e ~ /
I~
.~35-~0
10 4
2
-O'[ivine N~.|i rite Nephelin'i, te ~35-40\
i
Picritic Basonite 30-35
Atkotine Picrite 30-35
E
60 v,
Thoteiitic Picrite -35
t70 .~ n 80 ~J
90 100
110 120
0%
0',3
%H20 in melt for simple m e l t i n g of p y r o l i t e c o n t a i n i n g
0.1~ H20
FIG 4. Green's (1970a, b) petrogenetic grid based on a model pyrolite mantle containing 0.1% H20 (the numbers below each rock represent normative olivine): - - - , stability limits of minerals indicated.
now been abandoned for highly alkaline magmas in favour of their being primary melts. The petrogenetic scheme for alkali basalt genesis outlined above differs considerably from that of O'Hara (1968) who suggested that alkaline magmas were generated at a greater depth (at least 100 km) by higher degrees of partial melting (up to 30%) and by a more complex fractionation scheme in which as much as 40% garnet+ clinopyroxene and 40% olivine are fractionated from a tholeiitic picrite at various stages of ascent. Ringwood (1975, pp. 166-71) gave a (biased?) review of the differences between O'Hara's scheme and that of Ringwood, Green and coworkers. His principal criticisms of O'Hara's hypotheses are the oversimplification of O'Hara's four-component system relative to natural conditions, the inconsistency in the Mg/ (Mg+Fe), Ni, Cr and incompatible elements required by the large amount of fractionation involved in O'Hara's hypothesis and those observed in alkaline olivine basalts, and O'Hara's lack of quantitative assessment of his petrogenetic scheme. The genesis of alkaline mafic-ultramafic magmas based on Green's petrogenetic grid (Fig. 4)
presents a number of difficulties. None of the compositions represents rocks with K20 > Na20 and those with high Na20 contents, e.g. olivine melilite nephelinites, require very small degrees of partial melting. The genesis of these highly alkaline SiO2-undersaturated magmas is much more likely to be the result of direct partial melting than of fractionation processes as these magmas have a primitive chemistry which precludes even moderate fractionation and they are in equilibrium with a pyrolitic mineral assemblage (cf Ringwood 1975). Because these magmas have much greater abundances of incompatible, major and trace elements than are present in less alkaline types, very low degrees of partial melting seem to be required (cf Kay & Gast 1973). Nevertheless the problem of segregation of these very small amounts of partial melt and their subsequent ascent from 100 km or greater is formidable (see Yoder 1976). This difficulty is even greater if, as Ringwood (1975, p. 153) suggests, less than 1% partial melting of a hydrous source may be required for the generation of olivine nephelinite, olivine melilitite and kimberlite magmas. More recently McKenzie (1984) has developed models based on equations involving
34
A. D. Edgar
the movement of melt and matrix in partialmelts. Based on his calculations, very low degrees of partial melting (less than 3~) at depths commonly accepted for alkali basalts can readily allow melts to ascend to the surface. For highly alkaline magmas derived by direct partial melting of a mantle source the grid shown in Fig. 4 may not be as valuable as for less alkaline types. The grid is based on a pyrolite (lherzolite) source in which H20 is the only volatile component. Recently the concept that alkaline magmas are partial melts of mantle sources metasomatized to produce regions of quite different compositions from that of pyrolite and the importance of volatiles other than H20 in such source regions has been widely discussed ( c f Bailey 1982). Much of the remainder of this review is devoted to experiments pertinent to these concepts.
Recent experimental studies on alkaline mafic-ultramafic magmas Alkaline olivine basalts
The Basaltic Volcanism Study Project (1981, chapter 3) defines alkali olivine basalts as those with not more than 5~ normative nepheline and lists (Appendix 3.1, pp. 614-15) 11 experimental studies carried out on natural or synthetic samples up to 1978. The pressures varied from 1 atm to 33 kb under dry to H20-excess conditions, many with controlled or reasonably estimated oxygen fugacity (fO2) and temperatures from nearliquidus to solidus. Experiments performed at low pressures are not pertinent to the problems of the generation of alkali magmas under mantle conditions, but are important in establishing the sequence of mineral crystallization particularly near the liquidus and possible parental magma compositions (see Thompson 1973). These experiments have also been important in determining the effects offO2 on the stabilities of minerals in alkali olivine basalts (see Helz 1973, 1976). Unfortunately, fewer data are available for the effects offO2 on alkali magmas than are available for other basalts. Thompson (1974) studied a primitive alkali olivine basalt composition from Skye, Scotland (Mg number, 65; ne, 2.84~), between 8 and 31 kb at fOz>iron-wustite (IW) buffer under dry conditions. Because of its high Mg value it is unlikely that this alkali olivine basalt could be derived from a tholeiitic parental magma by fractionation; it is more likely to represent a pristine melt as suggested by Green (1970a) for similar alkali olivine basalts. Near-liquidus rela-
tionships in these experiments are very complicated at about 17kb (see Thompson 1974) suggesting that liquids must be saturated in olivine + clinopyroxene_ pigeonite. On the basis of these results Thompson proposed that this alkali olivine basalt was either a direct partial melt of an Fe-rich lherzolite at about 50km depth or a partial melt of a more Mg-rich lherzolite from deeper sources which underwent some olivine fractionation. Frey et al. (1978) estimated that 11~-15~ partial melting of garnet-free lherzolite can produce an alkali olivine basalt magma. Experiments without volatiles give no information on the roles of amphiboles and micas on the genesis of alkali olivine basalt. Allen et al. (1975) studied a Hualalai, Hawaii, alkali olivine basalt under pressure conditions similar to those used by Thompson (1974) but with excess H20 and f O 2 approximating those of the hematitemagnetite (HM), nickel-nickel oxide (NNO) and magnetite-wustite (MW) buffers. This composition has a comparable Mg number (67) but a lower ne (0.20~) than the sample used by Thompson (1974). Their experiments are shown in Fig. 5 which indicates that under HzO-excess conditions amphibole + olivine + clinopyroxene are near-liquidus phases around 13 kb, olivine is a liquidus phase up to 25 kb and garnet is present between 18 and 23 kb (Fig. 5(a)). The main differences between the results of the dry experiments (Thompson 1974) and the H20-excess conditions used by Allen et al. (1975) are the presence of amphibole and absence of pigeonite in the HzO-excess experiments (Fig. 5(a)). Allen et al. (1975) found that a n f O 2 corresponding to the N NO buffer produced amphibole at higher temperatures relative to the HM and MW buffer conditions (Fig. 5(b)). The absence oforthopyroxene in the experiments reported by Allen et al. places some constraints on the potential of orthopyroxene fractionation in the genesis of these basalts under H20-saturated conditions. The experiments support the concept that andesitic magmas are derived from basaltic magmas by amphibole-liquid fractionation processes as amphibole is stable at temperatures greater than the H20-saturated liquidus of andesite. Thus fractionation of amphibole from basaltic liquids during partial melting or fractional crystallization processes would lead to SiO 2 enrichment. Melting at depths corresponding to pressures greater than 23 kb, where amphibole is no longer stable (Fig. 5), would enrich liquids in Na20 and H20. Figure 6 shows the results of dry experiments performed by Takahashi (1980) on a primitive alkali olivine basalt from Oki-Dogo Island, Japan (Mg number, 63; ne, 3.40~), under low-fO2
Genesis o f a l k a l i n e m a g m a s
[--,
| I"
O I + C p x + G a + Ru + O p
25 ~-rA'mphibo'0"~ "
~
9 f, . / LOlivine Out] "4="
35
, ~
,
,
Y+
x 20
13_
10
,, ,,'
800
900
,
+ Cpx
o
__ J
1000 1100 800
900
1000
1100
Temperature, ~ (b)
(o)
FIG. 5. Phase relations of an alkali olivine basalt from Hawaii : (a) under H20-saturated conditions; (b) effects of variousfO2 on amphibole stability. Abbreviations are as in Fig. 2 and as follows : Cpx, clinopyroxene; Am, amphibole; Ga, garnet; Ru, rutile; Op, opaque; N-NO, nickel-nickeloxide buffer assemblage; M-H, magnetite-hematite buffer assemblage ; M-W, magnetite-wustite buffer assemblage. Liquid(s) coexist with all assemblages. (Modified from Allen et al. 1975.) conditions within the wustite stability field. This basalt, which contains spinel lherzolite nodules, gives results similar to those obtained by Thompson (1974) with orthopyroxene+clinopyroxene on the liquidus at 14-15kb and olivine+orthopyroxene (pigeonite) + clinopyroxene + liquid between 13.5 and 14 kb. This suggests that the
i
I
I
magma could have coexisted with the latter assemblage at depths corresponding to this pressure. This assemblage, along with spinel, also occurs at an isobaric invariant point within the D i - A n - F o tetrahedron between 7 and 20 kb (Presnall et al. 1978, 1979). Kushiro (1973), Presnall et al. (1978, 1979) and Mysen & Kushiro
i
i
I
i
1500
ou 1 4 0 0
Ga§
.
OL+Opx
~
+ Cpx+L t
d
OL+Cpx+L
\Opx+Cpx+L
/
_ /
r 1300 n
E
r
1200
1100
j
~
OL§
0
n 5
n 10
n 15 Pressure,
n 20 kb
~ 25
I 30
35
FIG. 6. Phase relations in an alkali olivine basalt from Japan (abbreviations as in Figs 2 and 5). (Modified from Takahashi 1980.)
36
A. D. E d g a r
(1977) suggest that liquids in more complex natural systems become nepheline normative between 10 and 20 kb. All these results suggest that the alkali olivine basalt lay on the SiOz-poor side of the A b - F o - D i plane (Fig. 1) at 14 kb. Based on these experiments Takahashi (1980) proposed that alkali olivine basalt magmas can be partial melts oflherzolite at depths corresponding to 14 kb provided that H20 and CO2 are unimportant in the melting relations, or that they were generated at greater depths but remained in equilibrium with lherzolite in the upper mantle. Alkali olivine basalts and systems of corresponding composition have also been investigated experimentally at crustal pressures under a variety of conditions. Duke (1974) studied an alkali olivine basalt (Mg number, 57.3; he, 1.63~) from the Azores at 1 atm and withfO2 determined using CO2 and C O + C O 2 mixtures. He found that oxidation increased the stability of spinel and clinopyroxene but decreased that of olivine and plagioclase, promoted crystallization of magnesioferrite-rich spinel and acmitic pyroxene, and caused nephetine-normative liquids to become SiO2 saturated or oversaturated. Using the system CaO-MgO-AlzO3-SiO2-Na20-H2 O at 5 kb to model various basalt types, Cawthorn (1976) found that nepheline-normative liquids could crystallize amphibole___ olivine + Ca-rich clinopyroxene to produce quartz-normative residual liquids of andesitic composition.
Green (1973a) investigated a lherzolite-bearing olivine basanite composition from Mount Laura, Victoria, with 10% added olivine (Mg number, 70; ne, 10.5%) between 5 and 36 kb and over a range of H20-undersaturated liquid conditions. As shown in Fig. 7, liquids from this composition are saturated with olivine + orthopyroxene + clinopyroxene+garnet only under very limited conditions corresponding to 25-30kb, 12001300~ and 2 - 7 ~ H20. Based on these experiments, Green (1973a) concluded that basanites could be generated from a lherzolite mantle with 0.2-0;4% H20 by 6 ~ partial melting at a depth of around 100 km. He suggested that the excess olivine fractionated from the liquid prior to the incorporation of lherzolitic mantle xenoliths, presumably during ascent. Using a kaersutite eclogite nodule from Kakanui, Hawaii, of composition corresponding to olivine basanite (Mg number, 63; ne, 6.4~), Merrill & Wyllie (1975) showed that orthopyroxene+ clinopyroxene + garnet coexisted at 20 kb under HzO-excess conditions, and at greater than 30 kb with progressively lower H=O contents. With the addition of normative olivine to their bulk composition, they found a more extensive field of olivine + orthopyroxene + clinopyroxene+ garnet in equilibrium with liquid than that reported by Green (1973a) and they also found that kaersutite was stable on the H20-undersaturated liquidus between 12 and 21 kb for liquids 40
Basanites and nephelinites Basic rocks more alkaline than alkali olivine basalts are generally defined as those with more than 5% normative nepheline. They may also contain normative leucite and a few have normative kalsilite. These magmas have primitive Mg numbers and other chemical features indicative of an origin by direct partial melting from a mantle source. Some are extremely rich in K and other incompatible and large-ion-lithophile (LIL) elements. Because of the many problems involved in the genesis of these rocks, including the fact that many cannot apparently be derived by partial melting of a (normal) lherzolitic mantle source, a large number of experiments on these magma compositions have been carried out in recent years. Those experiments particularly applicable to the nature of the source regions are concentrated on in this section. Over a decade ago it was recognized that highly alkaline magmas could not be derived by partial melting of volatilefree mantle peridotite but that some of these magmas could represent partial melts of an HzOand CO2-bearing peridotitic mantle (see Basaltic Volcanism Study Project 1981, pp. 529-30).
Ol
/
,/ go+'~px,/ I
I
~1 t 'IX
l +wx/ l + opx /
I go
~',,~ + ol l
I
go
/o,/// /op,+.l
:~ 20
!
,/ I g ~
/
/
/
/
/
/
J
/
/ /
l,,
II
I/ II
I jq.~
/
/
~'~
~ ~
Olivine/
/
,~
\1
!9
/ O........ livine
\//
o
L 1100
,
J
"-.I 1300
/ ,
. 1500
TEMPERATURE, ~ FIG. 7. Phase relations in an olivine basanite + 10~ olivine under dry to HzO-saturated conditions: - - -, percentage HzO added. Abbreviations are as given in Figs 2 and 5. (After Green 1973b).
Genesis of alkaline magmas containing at least 2-3% H20. From this Merrill & Wyllie (1975) inferred that basanites with kaersutite megacrysts may contain 2-3% H20 at depths corresponding to these pressures. Arculus (1975) investigated basanites from Mount Shadwell, Victoria (Mg number 68; ne, 11.92%), and from Grenada, Lesser Antilles (Mg number, 74; ne, 7.84%), under dry conditions between 10 and 35 kb. Owing to the absence of liquidus assemblages compatible with a dry peridotitic source (notably orthopyroxene), he concluded that basanites could not be derived from such a source unless H20 and CO2 were present; the former stabilizes olivine relative to pyroxene, and the latter stabilizes orthopyroxene relative to olivine and clinopyroxene (Eggler 1974). Arculus (1975) also concluded that oxides of polyvalent cations such as TiO2 and P205, commonly abundant in basanites, might lower the olivine stability field by 2-3 kb. Thus the absence of volatiles and such polyvalent cations in previously proposed models (e.g. O'Hara & Yoder 1967; Kushiro & Yoder 1974) suggests that such models might be inapplicable for natural systems of basanitic affinities. Evolved alkali basalts Evolved alkali basalts with lower Mg numbers and less normative nepheline than the primitive basanites and nephelinites may be produced from these more primitive magmas by various fractionation mechanisms. Bultitude & Green (1971) suggested that many basanites and nephelinites might be products of olivine fractionation of lowSiO2 picritic magmas. Irving & Green (1972) demonstrated that kaersutite was a near-liquidus mineral at 14 kb from a nepheline mugearite composition with about 3.5% H20 which contained lherzolite nodules. These studies indicate that fractionation of olivine+kaersutite from basanitic magma could produce nepheline hawaiites and nepheline mugearites at depths around 40-45 km. Although experimental data are lacking, it seems likely that further fractionation of clinopyroxene+biotite could produce nepheline benmoreites and phonolites (cf Irving & Price 1981). Olivine-rich nephelinites and melilitites In experiments with olivine nephelinites and comparable compositions up to 36 kb and under dry to H20-excess conditions (Bultitude & Green 1968, 1971; Allen et al. 1975; Merrill & Wyllie 1975) it was shown that such magmas could not be in equilibrium with a lherzolitic mantle source. In all cases clinopyroxene rather than orthopyrox-
37
ene appeared on the liquidus. Bultitude & Green (1971) indicated the improbability that tholeiitic picrite or alkaline picritic magmas are related to olivine and olivine melilite nephelinites by fractionation processes, thus implying that these magmas are primary. Brey & Green (1975, 1976, 1977) and Brey (1976, 1978) conducted a very detailed study of olivine melilitite compositions to investigate the effects of H20:CO2 ratio a n d f O 2 up to 40 kb on the genesis of these magmas. These authors used an olivine melilitite from Laughing Jack Marsh, Tasmania (Mg number, 74; ne, 17.2%; lc, 6.6%), and a synthetic Ca- and Mg-rich olivine melilitite (Mg number, 75; ks, 4.4%). In addition to analyses of the solid phases, the H20 and CO2 contents of the liquids were also determined. Using the Laughing Jack Marsh sample Brey & Green (1975) found no combination of pressure, temperature and percentage H20 conditions under which this olivine melilitite could be in equilibrium with a pyrolite mantle assemblage. In the presence of H20 the solubility of CO2 is high in the olivine melilitite liquid, and it suppresses olivine and clinopyroxene and promotes orthopyroxene and garnet as liquidus or near-liquidus phases. These results, using an external HM buffer assemblage, are summarized in the 30 kb section shown in Fig. 8. With increasing COz (decreasing HzO/(H20 + CO2)) the liquidus temperature rises almost 200~ between 0.5 H20/ !
!
! 1400
? e
\
13oo
0
ox
§
0
9
~ 9
0.'25 0.~5
- 1200
0
0175
C02
38wt7.
T*C
H20 H20 * C02 mole fraction in c h a r g e
1100 H20 20wt7.
Fla. 8. Phase relations in an olivine melilitite showing variations in the liquidus phases at 30 kb and at f O 2 = HM buffer with varying HzO/(HzO + C02). Liquidus and near-liquidus phases are indicated above the liquidus: O ; garnet; +, orthopyroxene; 0, clinopyroxene; o, all liquid. (Modified from Brey & Green 1975.)
38
A.D.
(H20 + CO2) and CO2-only experiments. Based on these results, Brey & Green (1975) inferred that olivine melilitite magmas could be derived at 30kb and 1150-1200~ by less than 5~ equilibrium partial melting of pyrolite containing 0.5-1.0 H20/(H:O+CO2). They suggested low fO2 values (less than the N N O buffer) as most appropriate. Brey & Green (1977) refined the results of their earlier study by systematic investigation of the liquidus and near-liquidus relations with varying H 2 0 / ( H 2 0 + C O 2 ) and f O 2 conditions. The resuits shown in Fig. 9 represent near-liquidus phases at 30 kb andfO2 = HM buffer. Of particular importance is the narrow field of clinopyroxene separating the garnet+orthopyroxene+ clinopyroxene field from the olivine + clinopyroxene field for varying CO2 and H20 contents. Experiments at 27 kb with a composition containing 0.92 H20/(H20 + CO2) closely approached a near-liquidus olivine + orthopyroxene + clinopyroxene+garnet assemblage, compatible with equilibration of the olivine melilitite magma with a pyrolite source. Brey & Green (1977) proposed that the additon of 5% olivine (implying fractionation of a similar amount of olivine from the primary magma) would produce this four-phase assemblage. They also determined that liquidus and near-liquidus assemblages were not appreciably different at more realistic mantlefO2 values corresponding to those less than the N N O buffer assemblage. Based on these results, Brey & Green (1977) proposed that olivine melilitite magmas could be derived at depths corresponding to 27 kb at 1160~ by partial melting of a pvrolite mantle source containing both H20 and CO2 equivalent to 6 - 7 ~ CO2 and 7-8~o H20 dissolved in the melt. Brey (1978) proposed that olivine melilitites which were more SiO2-undersaturated than the ,
/ O~
4Q
O~. /"'f I"j1~
20
Laughing Jack Marsh sample might be derived at greater depths and with higher CO2 in the source region. Experiments at 30-35 kb with higher CO2 contents using a very SiO2-undersaturated synthetic olivine melilitite composition produced olivine+clinopyroxene as liquidus phases. At 30 kb this assemblage is clearly separated from the garnet + orthopyroxene field by an orthopyroxene field. He found that garnet stability increased at lower H20/(H20+CO2) and suggested that highly SiO2-undersaturated olivine melilitites could be derived at deeper levels (corresponding to 35 kb) by partial melting of a more CO2-rich pyrolite source. Although Brey & Green (1975, 1977) and Brey (1978) demonstrated the role of CO2 in addition to H20 for the genesis of olivine melilitite magmas, Eggler (1974, 1978), Huang & Wyllie (1974), Wyllie & Huang (1975, 1976) and Wendlandt & Eggler (1980a, b) proposed that CO2 was an essential component in the source regions for mafic SiO2-undersaturated magmas. Reviews of the experimental evidence for the role of CO2 are given by Eggler & Holloway (1977) and Wyllie (1977, 1979). Eggler (1978) studied joins in the system Na20-CaO-A1203-MgO-SiO2CO2 up to 30 kb and showed that orthopyroxene stability increased relative to that of olivine, thus producing liquids containing less SiO 2 than those formed in the absence of CO2. He suggested that nephelinites and melilite nephelinites could be partial melts of a peridotitic source at 50-90 km provided that CO2 was the predominant volatile component. If the total volatile contents (H20 + CO2) of the mantle source region are greater than 0.5 wt.% all variations between H20-saturated and CO2-saturated melts can occur, whereas with low total vapour concentrations in the source regions the compositions of the vapour a r e restricted and hence solidus temperatures and near-solidus melt compositions are also restricted. As an example of the role the amount of H20 and CO2 plays in the composition of melts. Eggler (1978) proposed that partial melts of an amphibole-peridotite in the presence of both H20 and CO2 (not exceeding 0.37 wt.%) will produce a nephelinitic melt composition at 15 kb.
/"
~i~ 20
Olivine Melilitite
,,
Edgar
Kimberlites "
ga + opx
i 40
ga + opx + cpx C pX ol § c p x ol 60 H20
FIG. 9. Phase relations in the system olivine melilititeH20-CO 2 at 30 kb andfO 2= HM buffer. Abbreviations are as in Fig. 5. (Modified from Brey & Green 1977).
The role of CO2 in the genesis of magmas is particularly relevant to kimberlites and carbonatites. Eggler & Wendlandt (1979) determined phase relations in a kimberlite with variable H20 and CO2 contents at 30 and 55 kb. At the higher pressure they found that kimberlite-like liquids coexisted with a lherzolite assemblage in experiments with 5% H20 and 5% CO2, but kimberlites
Genesis of alkaline magmas
could not be derived from a lherzolite source at the lower pressure. At 30 kb phlogopite and dolomite were sub-solidus phases in experiments with HzO-rich buffered vapour, whereas magnesite solid solution was a sub-solidus mineral at 55 kb. On the basis of phase relationships in the peridotite-COz-H20 system, particularly those determined by Eggler (1978), Wyllie (1977, 1978, 1979) and Ellis & Wyllie (1980), Wyllie (1980) suggested that kimberlite originated by partial melting of peridotitic mantle near 260 km in which solidus temperatures were reduced by upward migration of volatiles. This melting causes a density inversion and results in a diapiric uprise under adiabatic conditions with subsequent crystallization of the partially melted diapiric material beginning at depths of 10080 kin. Leucitites and olivine leucitites With the exception of kimberlites, all the alkali magmas so far described have N a z O > K 2 0 . Mafic-ultramafic rocks with K z O > N a 2 0 are essentially leucitites and olivine leucitites, but with a plethora of varietal names. Most of these rocks fall within the broad category of lamproites (cf. Bergman 1987). They commonly have normative leucite but may be SiO2-saturated to oversaturated and contain normative quartz. Problems involved in the genesis of these often highly LIL-element-rich rocks have been summarized by Bell & Powell (1969). Edgar et al. (1976) have pointed out some of the improbabilities of the older theories given by Bell & Powell. These theories fall into three categories. Firstly, differentiation of basaltic or peridotitic magma with fractionation of olivine (Wade & Prider 1940), eclogite and olivine (Holmes 1932), or eclogite (O'Hara & Yoder 1967). Secondly, assimilation of magmas of such diverse compositions as granite, carbonatite, nephelinite and monchiquite by a wide variety of crustal materials (cf Daly 1910; Bowen 1928; Williams 1936; Holmes & Harwood 1937; Larsen 1940; Holmes 1950, 1965; Turner & Verhoogen 1960). And thirdly, partial melting of biotites and hornblendes (Waters 1955) and kimberlite (Harris, quoted by Bell & Powell 1969). Any hypothesis involving basaltic magma seems unlikely, as in most occurrences of rocks of the leucititic kindred basalts are sparse or absent. Even if basaltic material was involved, very extreme degrees of crustal contamination or assimilation would be required to account for their high incompatibleelement content and to maintain their high MgO and CaO values. The theories involving partial melting of some types of mantle material are now widely accepted by most workers. Recent wide-
39
spread interest in these rocks has been prompted by the possibility of greater heterogeneity in the mantle source regions than previously proposed and by the current popularity of the process of mantle metasomatism (cf Bailey et al. 1980; Bailey 1982). In general, the similarities in high Mg number, low SiO2 content and high volatile contents of olivine leucitites, olivine basanites, olivine melilitites, olivine nephelinites and kimberlites suggest that these magmas are primary rather than derivative and hence can only be evolved from a mantle source containing H20 and CO2. Cundari & O'Hara (1976) conducted dry experiments on a mafic leucitite from New South Wales, Australia (Mg number, 79; lc, 0H; ne, 9.08%). They found clinopyroxene as the only liquidus phase between 20 and 40 kb and olivine at lower pressures. Absence of the assemblage olivine + orthopyroxene+ clinopyroxene _+garnet_+ spinel as near-liquidus assemblages precludes this magma's being derived by partial melting of dry lherzolite. Cundari & O'Hara (1976) concluded that this rock may have formed by partial melting of an olivine pyroxenite source containing essential phlogopite. Absence of normative leucite or quartz in this composition sets it apart from other compositions. Thompson (1977) conducted experiments under similar conditions on an apparently less-primitive leucitite from the Alban Hills, Italy (Mg number, 63; lc, 32.5%; ne 13.6%), with the same results. High-pressure experiments using H20 and H20 + CO2 have been carried out for compositions ranging from potassic intermediate to ultramafic from SW Uganda (Edgar et al. 1976, 1980; Ryabchikov & Green 1978; Arima & Edgar 1983a), from the Leucite Hills, Wyoming (Barton & Hamilton 1979, 1982), and from the W Kimberley area, W Australia (Arima & Edgar 1983b). All these compositions have primitive Mg numbers (70-78) but widely varying degrees of K enrichment (K/(K + Na) mol = 0.41-0.86). In Fig. 10 these rocks are plotted on the Mg2SiO4 (Fo)-KA1SiO~ (Ks)-SiO2 system which indicates the variable degrees of SiO2 saturation. The three compositions from SW Uganda (olivine ugandite, katungite and biotite mafurite) and the madupite from the Leucite Hills are all SiO2 undersaturated, whereas the orendite from the same locality and the wolgidite from W Kimberley are SiO2 saturated and slightly oversaturated respectively. The SiO2-undersaturated group shows distinctly different liquidus to near-liquidus phases from the group which is SiO 2 saturated to oversaturated. The primary phase fields at 3 kb PH2o (Luth 1967) and 28 kb Pmo (Green et al. 1984) in this system are also included in Fig. 10. With
40
A . D. E d g a r Oi
Sil wt.% FIG. 10. K-rich mafic-ultramafic rocks (O), for which experimental data are available, plotted as CIPW normative minerals in the system Mg2SiO4-KA1SiO4-SiO2 together with phase relations at PHzo of 3 and 28 kb (Luth 1967; Green et al. 1984)" HI, phlogopite composition. Abbreviations are as in Figs t, 2 and 3" Hy, hypersthene. Ks
Lc
So
increasing pressure the fields of phlogopite and enstatite expand relative to olivine. Thus partial melting of a mantle source with H20 as the only volatile phase results in liquids which may crystallize these minerals. With H20 as the only volatile phase neither orthopyroxene nor garnet crystallize from any of the SiO2-undersaturated compositions. Thus these magmas cannot have equilibrated with a normal lherzolitic mantle source. Figure 11 is a schematic diagram of near-liquidus phase assemblages for these compositions with H20 contents between 3 and 15 wt.~, representing vapourabsent conditions above about 10 kb. Clinopyroxene___phlogopite are the commonest near-liquidus minerals at high pressures and olivine+ clinopyroxene are the commonest at lower pressures. Based on these near-liquidus phase assemblages, possible source regions for these magmas range from a slightly K-enriched wehrlite for the least K-rich olivine ugandite magma to a phlogopite-rich clinopyroxenite for the highly K-rich mafurite magma (Fig. 11). The appearance of
phlogopite as a near-liquidus assemblage is directly correlated with the degree of K enrichment (Edgar & Arima 1983). Figure 12 shows near-liquidus assemblages for the SiO2-saturated (orendite) and SiO2-oversaturated (wolgidite) compositions in which only H20 has been added. Orthopyroxene is a common near-liquidus mineral for these compositions. In the orendite the near-liquidus assemblage of olivine + orthopyroxene + clinopyroxene + garnet + phlogopite (trace) at 30 kb suggests that this magma could be derived from a phlogopitebearing garnet lherzolite containing small amounts of H20. A phlogopite orthopyroxenite and a phlogopite harzburgite seem the likeliest sources for the wolgidite composition at high and low pressures respectively. The nature of the source materials and the depth of generation depend on the amount of H20 in the source. Ryabchikov & Green (1978) studied the effects of CO2 and H20 at 30kb on the mafurite composition used by Edgar et al. (1976). Figure 13 shows that, with total volatiles equivalent to 5
Genesis of alkaline magmas
40-
Madupite OI-ugandite Katungite Mafurite (Barton 8~ (Edgar e/o/. (Arima 8~ (Edgar elo/. Hamilton 1980) 1980) Edgar 1983a) 1976)
41
OI-ugandite (Edgareta~. 1980)
Katungite Mafurite (Arima & (Edgar et o/. Edgar 1983a) 1976) -140
T
T
35
t
30-
s_HlightlyK-enriched wehrlite (ol+cpx)
K-bearing clinopyroxenite (cpx) a_ 2 5 -
I
3%
-
RI---#-~ h g~ phlog_o_pite clinopyroxenite clinopyroxenite (cpx+phl) (cpx+ phi)
p_hloggRitewehrlite (cpx + ol + phi)
phlogopitewehrlite (cpx +ol+phl)
20-
, ]
P_h I._9_og_opit e___: clin.~_opyroxenite (cpx+phl)
t
120
slightly K-enriched wehrlite (ol+ cpx)
(ol only) p_hl og_op_it_eewehrlite (ol+cpx+phl)
I
~
(ol only)
l
5%
4
H20
p_hlogo!it.__eewehrlite (ol.cpxephl)
I
added
l
p_hl og_opit_.ee wehrlile (cpx + ol + phi) 15%
I00
80
I
60 I
FIG. 11. Schematic diagram showing near-liquidus phases for SiO2-undersaturated K-rich magmas in which H20 is the only added volatile. The underlined rock names are approximately equivalent to the near-liquidus assemblages in brackets. wt.% H20 , olivine+clinopyroxene are liquidus phases at CO2/(CO 2 Jr H 2 0 ) mol < 0.50, whereas orthopyroxene stability increases above this value. For CO2/(CO2 + H20) mol between 0 and 0.50 phlogopite is stable but clinopyroxene rapidly becomes unstable at higher ratios. For CO2/(CO2+H20 ) equivalent to 15 wt.% H20 , phlogopite and clinopyroxene stability decrease at low ratios, with orthopyroxene as the liquidus phase over most of the range of CO2/(C0 2 Jr H20) values. These relations and the absence of garnet suggest that a lherzolitic source may be possible for this highly K-rich composition. Figure 14 is a schematic diagram of three SiO 2undersaturated SW Ugandan lavas at CO2/ (CO2+H20) mo1=0.75 (equivalent to 15 wt.% H20 added). Only the low-K ugandite with orthopyroxene + clinopyroxene_+ garnet 4- spinel as near-liquidus phases from 22 to 40 kb can be derived from a lherzolitic source under these conditions. The katungite liquidus relationships suggest a garnet clinopyroxenite or a clinopyroxenite as the likeliest source compositions. Although phlogopite does not occur in these highCO2 experiments, the source regions must be K enriched in order to produce the K enrichment in the magmas provided that these are direct mantle melts. No experiments with high CO2 have been carried out on the SiO2-saturated to SiO 2oversaturated orendite or wolgidite.
Wendlandt & Eggler (1980a, b) modelled the genesis of potassic magmas based on melting relations in the systems KAISiO4-Mg2SiO4SiO2, KA1SiO~-MgO-SiO2-CO2 and KA1SiO4MgO-SiO2-H20-CO2 as simple mantle sources at pressures up to 50 kb. In these systems they were able to determine a large number of reactions producing a wide variety of possible liquids ranging from simplified tholeiites to carbonatites. On the basis of previous studies they suggested that the absence of Ca (as CaMgSi206) and Na (as NaA!SiO4 and NaA1Si3Os) would not appreciably alter their conclusions. The expansion of the primary phase fields of enstatite and sanidine in the KA1SiO~-Mg2SiO,~SiO2 system with increasing pressure is illustrated by the schematic diagrams given by Wendlandt & Eggler (1980a, Fig. 6). This results in liquids becoming progressively enriched in K20 and depleted in SiO2 at pressures ranging from 5 kb to greater than 30 kb. In the KA1SiO4-MgOSIO2-CO2 system, phase relations are similar but occur at lower pressures than in the CO2-free system. The trends of liquids with increasing pressure are schematically shown in Fig. 15 which is based on the variation in the extent of the primary solid phase fields with pressure in the systems KA1SiO4-MgO-SiO2-CO2 and KA1SiO4-Mg2SiO4-SiO 2 (Wendlandt & Eggler 1980a). The approximate pressures at which the
A. D. Edgar
42 Orendite {Barton 8 Hamilton 1982)
Wolgidite (Arima 8 Edgar !983b)
40 -f
-140
35
120 ph Iogopite
orthopyroxenite (opx+phl)
p_h I__qgo_pite ?garnet- Iherzolite {ol+opx+cpx+gt+ph)
30
Io0 T s s
25
ph I__0_ogo_pit3 orthgpy_roxenite (opx+phl)
p hl_#_ogop~t__e harzbuEgit__e (oi+ opx + phl)
K3
H. p_ ogoplte harzburqit._ee (opx +ol + phi)
80
20
(ol Only) I te p.hi ogopj harzburglte (o1+ opx+ phi) 5%
1'2%
60 15%
H20 added FIG. 12. Schematic diagram showing near-liquidus phases for SiO2-saturated and SiO2-oversaturated K-rich magmas in which H20 is the only added volatile. The underlined rock names are approximately equivalent to the near-liquidus assemblages in brackets. 1400
m
L
Iq
o
1400 1 30 kb (equiv to 15 wt% H20)
.+v/
J+FTa 1300
,,'
(bq~//
1300 -
,,
r
OI
,p
/
-....--
/ /
L O
/
~ ~ +
Xtols
/ /
1200
_
/
1200
-
,., C~4\\
-
L+Xtals
-% \\ \ \ \,o. \ \ \-o \%
30 kb {equiv. to 5 wf% H20) I100 H20 5 wt~
I
0.25
I
0-5
I
0.75
C02/(C%+ H20} mole fraction in charge
CO2 12.2 wt~
I100 H20 15 w1%
I
0.25
\
I
0"5
I
0.75
C02/(C02-1- H20) mole fraction in charge
C% 36.7 wt%
F]G. 13. Phase relations in the biotite mafurite system with varying C02/(H20 --kC O 2 ) contents equivalent to 5 and 15 wt.% H20 at 30 kb. Abbreviations are as in Figs 2-4. (Modified from Ryabchikov & Green 1978.)
Genesis of alkaline magmas Katungite (Arima a Edgar 1983a)
OI-ugandite (Edgar et 0/ 1980
40
t
:55 -
Mafurite (Ryabchikova Green 1978)
I
garnetIherzolite ? (cpx + opx + ga)
43
i
' ' garnet' clinopyroxenite (cpx + ga)
-
140
'
120
,r
E
v (D
50
Iherzolite (cpx + opx + ol)
clinopyroxenite (cpx)
r
ioo
d3
13._
25
_
garnet or spinel Iherzolite (opx + cpx + ol + ga + sp) 80
wehrlite (ol +cpx)
20 COz H201
27.4 3-5
27.5 3.8 % CO2 and H20 added
27.5 5"8
FIG. 14. Schematic diagram showing near-liquidus phases for SiO2-undersaturated K-rich magmas with CO2/(CO2 + H20) mol = 0.75 (equivalent to 15 wt.% HzO). The underlined rock names are approximately equivalent to the near-liquidus assemblages in brackets.
Mag Pressure (kb) of appearance of liquids Vol absent 19 34-5 --
CO2 satd. 14 27.5 > 29
Liquids Lc normative Ks normative Mag normative Carbonatitic
/
/
/
/
~.// 7
/
7
// "
~
/
/
/ 9
/
/
/~hl
/
Olivine tholeiitic
~/
/
/
/ //
/ /
/
/
Quartz.. tholeiitic
\
/
7
/.. Ks
/ Lc [ Sa (Ks normative] Alkali basaltic / Lc normative}
Qz
FIG. 15. Trends of liquids with increasing pressure in the systems KAISiO4-MgzSiO4-SiO2 and KA1SiO,-MgOSIO2-H20-CO2 (modified from Wendlandt & Eggler 1980a). Abbreviations are as in Figs 1 and 3 ; Mag, normative magnesite. The inset shows the normative nature of liquids at increasing pressures (as indicated by the line on the diagram) under volatile-free and CO2-saturated conditions.
44
A. D. Edgar
compositions of the residual liquids on partial melting become leucite, kalsilite and magnesite normative are shown in the inset to Fig. 15. At approximately 14 kb for CO2-saturated conditions and 19 kb for volatile-free conditions, liquids in equilibrium with sanidine, enstatite and forsterite are leucite normative and change from tholeiitic to alkali basaltic with increasing pressure. At greater than 27.5 kb under CO2saturated conditions and greater than 34.5 kb in the volatile-free system, sanidine is no longer stable and melts become extremely K20 rich and SiOz poor, yielding kalsilite-normative liquids. At pressures above 29 kb in the KA1SiO4-MgOSIO2-CO2 system melts are actually magnesite normative. Based on these results Wendlandt & Eggler (1980a) concluded that quartz tholeiitic liquids cannot fractionate to alkali basaltic ones. Liquids generated at high pressures with some fractionation could result in increasingly SiO2poor residual liquids. In contrast, fractionation at lower pressures could result in SiO2 enrichment. Wendlandt & Eggler's (1980a) experiments have been used by Kuehner et al. (1981) to model the Leucite Hills, Wyoming, magmas which show this trend of SiO2 enrichment. Although it is theoretically possible that leucite-normative rocks could be products of fractionation rather than direct partial melting, Wendlandt & Eggler (1980a) consider this to be unlikely. Using the KA1SiO4-MgO-SiO2-CO2-H20 system in which phlogopite is an additional phase, Wendlandt & Eggler (1980b) modelled the melting behaviour of phlogopite peridotite and its role in the genesis of K-rich magmas, kimberlites and carbonatites. Peridotite with small amounts of H20 + CO2 undergoes univariant melting in which H20 and CO2 vapour compositions are buffered by phlogopite and magnesite respectively. Such buffering is pressure dependent and hence the compositions of liquids produced by various reactions depend on the pressure of melting. With increasing pressure up to 30 kb, melting of a phlogopite peridotite with insufficient CO2 and H20 to hydrate all the potential phlogopite involves four reactions between combinations of phlogopite, enstatite, sanidine (or kalsilite), forsterite, liquid and vapour. As the pressure increases up to 30 kb the liquids change in composition progressively from quartz to enstatite to leucite to kalsilite normative. These relations are comparable with those for the KA1SiO4-Mg2SiO4-SiO2 system (Wendlandt & Eggler 1980a). Above 30 kb magnesite becomes stable and melting occurs by two reactions involving combinations of phlogopite, enstatite, magnesite, forsterite, kalsilite, liquid and vapour. The liquids produced are carbonatitic. With
increasing pressure phlogopite has decreased hypersolidus stability and at 50 kb ceases to be a solidus phase. Thus between 30 and 50kb, melting of phlogopite peridotite will produce increasingly K20-rich SiO2-poor liquids.
Mantle metasomatism Experimental studies and other data indicate that as the alkalinity of mafic to ultramafic magmas increases there is a decreasing possibility that these magmas are partial melts of a 'normal' pyrolitic model mantle composition. Although there are a number of mechanisms whereby local heterogeneities can occur in the mantle source, the concept of heterogeneities being caused by mantle metasomatism (cf. Lloyd & Bailey 1975; Bailey 1982) is now widely accepted. Experimental modelling of metasomatic processes is extremely difficult, particularly at mantle depths where very little is known of either the composition or the speciation of the fluids involved. Among the many elements considered to be responsible for metasomatism (Bailey 1982), K and Na are very important. Ryabchikov & Boettcher (1980) showed that the solubility of potassium (as K/O) in aqueous fluids increased about sixfold between 11 and 30 kb in the range 1050-1100~ Their values, ranging from 4 to 25 g K20 per 100 g H20, indicated that the solubility of K:O is sufficiently high to produce radical changes in the mantle. Ryabchikov et al. (1982) determined that the amount of sodium silicate leached by hydrous fluids at constant temperature from omphacitic pyroxenes increases with decreasing pressure. On the basis of both sets of experiments, Ryabchikov et al. (1982) proposed that K/Na fractionation in the upper mantle might produce K-rich metasomatism at deeper levels and Na-rich metasomatism at shallower levels which, on partial melting, would result in K-rich magmas generated at greater depths than Na-rich magmas. Edgar & Arima (1984; unpublished data) have carried out preliminary metasomatism-modelling experiments involving the system pyroliteK2Oaqueou s and pyrolite-(K20+Na20)aq ..... at 2 0 - 3 0 kb and between 850 and 950~ In these experiments K20 , as dried K2CO3, was dissolved in deionized H20. Na20 was added as Na2CO 3. This resulted in an aqueous solution containing negligible amounts of dissolved CO2. Figure 16(a) shows that at 30 kb an increased concentration of K20 in solution produces an initial increase in the proportions of phlogopite relative to orthopyroxene, clinopyroxene and olivine. Although
Genesis of alkaline magmas not in sufficient quantity to be observed in the Xray diffraction patterns (Fig. 16) garnet (at 30 kb) and spinel (at 20 kb) decrease with increasing K20 as observed optically. At about 3.8 g K20 per 10 g of solution, liquid appears at both 850 and 950~ indicating a decrease in the amount of phlogopite with K partitioning preferentially into the liquid. These results are in accordance with those of Takahashi & Kushiro (1983). Edgar & Arima (1984) were unable to analyse their run products using the electron microprobe because of the fine grain size, but they speculated that at 30 kb the formation of phlogopite involved a reaction between olivine, orthopyroxene and garnet in the approximate proportions ofpyrolitetype mantle at 30 kb (Wyllie 1975) with the K20 in solution to produce a product with less garnet and olivine and more orthopyroxene. Clinopyroxene is probably not involved in the formation of phlogopite. These results are in accordance with the observed decrease in the olivine/orthopyroxene and olivine/clinopyroxene intensities (Fig. 16(a)). In contrast with the results at 30 kb, at 20 kb an amphibole of pargasitic composition as analysed by microprobe shows an inverse relationship to phlogopite which remains the dominant 'metasomatic' product (Fig. 16(b)). On the basis of the relative proportions of phlogopite to the other mineral phases (Fig. 16(b)) and the ratios of olivine to orthopyroxene, olivine to clinopyroxene and orthopyroxene to clinopyroxene (Fig. 16(c)), the formation of phlogopite at 20 kb and at K20 concentrations where amphibole is no longer present (more than 2.0 g K20 per 10 g of solution) may involve a complex reaction between olivine, orthopyroxene, clinopyroxene, amphibole and spinel, together with K20 in solution, to form a product containing a higher amount of olivine and clinopyroxene and lower amounts of orthopyroxene relative to the reactants. A few preliminary experiments at 20 kb and 950~ in which both Na and K have been added in various proportions equivalent to 3.8 g total alkali per 10 g of solution show that, with increasing Na relative to K, the field of amphibole stability increases as phlogopite decreases. Although vastly simplified, these experiments indicate that alkali metasomatism can readily occur. Nodules with mantle mineralogy found in alkali basalts, kimberlites and related rocks often contain phlogopite and amphibole and range in overall composition from lherzolite to pyroxenite. Lloyd & Bailey (1975) and Lloyd (1981) studied nodules mainly of phlogopite clinopyroxenite composition from the K-rich volcanic province of SW Uganda and proposed that they might represent metasomatized mantle sources for these
45
lavas. Alternatively these nodules could be depleted residual source materials. Lloyd et al. (1985) determined the compositions of liquids produced from varying degrees of partial melting at 20 and 30 kb of an average nodule with no added volatiles (based on analyses of 84 nodules from this area). Figure 17 shows the compositions of liquids produced by various degrees of partial melting at 30 kb together with the coexisting minerals. At approximately 25% partial melting of this nodule, the liquid corresponds closely with that of an average katungite composition found in SW Uganda. Although this is a higher degree of partial melting than generally considered likely for the production of such K-rich magmas, it seems feasible in the light of the upwarped mantle isotherms, postulated in rifted areas such as SW Uganda (Bailey 1970), caused by mantle degassing (Bailey 1980), and in view of the high proportion of radioactive elements associated with these magmas.
Felsic alkaline rocks Although experiments pertinent to felsic alkaline rocks have been less numerous than those carried out on mafic varieties in the past decade, studies have been undertaken on liquid immiscibility, on the role of C1, F and P20 5 in alkaline magmas and on the problems of the genesis of pseudoleucites and primary analcites.
Liquid immiscibility Roedder (1979) reviewed the role of experimental work on the concept of liquid immiscibility as a viable process in the genesis of igneous rocks. Only studies applicable to alkaline rocks are considered here, most of which concern silicatecarbonatitic magmas. Freestone (1978) showed that the liquid immiscibility field in the fayalite-leucite-silica system was extended to more K-rich compositions by the addition of 1 mol ~ P205 and 3 mol ~ TiO2. He suggested that this expansion supports the liquidunmixing origin suggested by ocelli textures in syenites and sheets within alkali gabbro intrusions. The role of liquid immiscibility on the genesis of carbonatites was considered by Hamilton et al. (1979) and Freestone & Hamilton (1980). These workers used lavas ranging in composition from felsic nephelinites to phonolites comparable in composition with those in the Oldoinyo Lengai volcano, Tanzania. These were mixed with synthetic Na- and Ca-rich carbonatites. Experiments at 0.7-7.6 kb and 900-1250~ showed that
46
A. D. Edgar 8
_- q~p
/
6
"~ ~
o
o=~=I=ua 2 0
-
@
o
k
,~ ~
~ H
o
1.5~ I'0
o
u H
4
\ Iamph 2
o
oI
0
1"5~~I'0
9
-~---I
H
0.5 8
4
q,t-fl~p
4
6 n
"~
3 2 I
0
v
I'0
2.0
3-0
4-0
I
~
6.0
5.0
1
i0
2. 0
50
l
4-0
g K20/IO g sOIn.
g K 2 0 / I O g soln.
(a)
(b)
I 5.0
I 6,0
3-0 -
2o' I.O
o
u
-6
1.5
0"5
2"0 I-0 I ~ - -
0
a-
1
I-O
o
@ I
2.0
I
3.0
I
4-0
g K20/IO g soln. (C)
I
5-0
I
6-0
FIG. 16. Results of experiments in the pyrolite-K20 aqueous solution system plotted as ratios of X-ray diffraction intensities (I) of minerals against K 20 concentrations: (a) runs at 30 kb, 850~ ( 0 ) , and at 30 kb, 950~ (O); (b) runs at 20 kb, 950~ showing variations in the ratios of phlogopite and amphibole relative to other minerals; (c) runs at 20 kb, 950~ showing variations in the ratios of olivine, orthopyroxene and clinopyroxene. Abbreviations are as in Figs 2 and 3 ; q, quench phlogopite; p, primary phlogopite. (After Edgar & Arima 1984.)
Genesis of alkaline magmas
~_ApI
(not used in Freestone & Hamilton's experiments) liquid immiscibility could account for the carbonate-rich groundmass in kimberlites and the latestage associated carbonatite dykes and diapirs.
Phi C p x ~
Im. . . .
37
Halogen-bearing
I0
~
8-
u
/
-
e
-
systems
Hards (1976) determined the distribution of certain elements between the melt and vapour phases in granite and nepheline syenite compositions at 1 kb between 680 and 850~ in the presence of C1 and F. He found that C1 was more strongly concentrated in the vapour phase in granitic melts than it was in nepheline syenite melts, whereas F was concentrated preferentially in the liquid phase for both compositions and only entered the vapour phase after crystallization
O
~
6-
NazO
0 ~- ~ 9
e
O~
xcr..-"
if
0
47
J
w,o,o
II
L
I
i
I
~
I
20 40 60 80 Degree of melting (wt.%)
i
I
I00
FIG. 17. Compositions of liquids with varying degrees of partial melting of an average clinopyroxenite nodule from SW Uganda at 30 kb : - - , determined compositions; - - - , calculated compositions based on analyses of minerals. The phases present are given at the top of the figure. Cpx, clinopyroxene: Phi, phlogopite: Ilm, ilmenite: Ap, apatite. (After Lloyd et al. 1985)
Si% + AiaO3
CaO (a)
Na20
the miscibility gap expands with increasing Pco2 and decreasing temperature (Fig. 18). Results at 0.7 kb and l l00~ are plotted on the N a 2 0 (SiO2+AlzO3)-CaO system (Fig. 18(a)). Under the miscibility gap inincreasing P t o t a l = P c o 2 , creases and there is a rotation of tie-lines resulting in more CaO-rich liquids (Fig. 18(b)). From the compositions of the melts, Freestone & Hamilton (1980) concluded that the natrocarbonatites from Oldoinyo Lengai separated from a phonolitic rather than a nephelinitic magma. Because the miscibility gap closes away from the Na20-rich compositions (Fig. 18), exsolution of a carbonatitic melt is more likely in salic than in mafic silicate magmas. Freestone & Hamilton's (1980) results do not support the idea that kimberlite sills and dykes form by an immiscibility process at crustal pressures. Bedson & Hamilton (1981), however, showed that in the presence of H 2 0
SiOa + Ala03
CaO
(b) FIG. 18. Liquid immiscibility in the system NazOCaO-(A1203 + SiO2) at (a) 0.7 kb and 1100~ and (b) 1100~ and varying pressures Pc% = Ptotal. (After Freestone & Hamilton 1980.)
A. D. E d g a r
48
was completed. Sodium was the most common element in the vapour phase of both compositions. These experiments may explain the presence of Cl-bearing sodalite in SiO2-undersaturated rocks but the absence of Cl-bearing minerals from granitic rocks. On the basis of Hard's experiments the presence of F-bearing minerals in both types of rocks is to be expected. On the basis of experiments at 1 kb in the system NaA1SiO4 (Ne)-SiO2-NaC1-H20, Barker (1976) found that in the SiO2-undersaturated portion of the system NaC1 stabilized sodalite whereas in the SiO2-oversaturated part it has no effect on phase relations because of immiscibility between NaCl-rich liquid and SiO2 melt. These results suggest that the high C1 content of feldspathoidal rocks is not a cause of their SiO 2 undersaturation and that such rocks are unlikely to be produced by anatexis or assimilation of halite-bearing sediments (Appleyard 1974; Ayrton 1974). For the SiO2-undersaturated part of the same system, Binsted (1981) showed that sodalite might exist as a primary mineral between about 25% and 60% Ne in the pseudobinary ( A b + Ne)gs(NaC1)5 join (Binsted 1981, Fig. 13). He found that melts crystallizing sodalite follow a non-linear trend (Binsted 1981, Fig. 14) owing to partial replenishment of NaCI in the melt from the fluid reservoir. With increasing NaC1, nepheline will crystallize as the second phase following albite. These relationships may be important in the genesis of agpaitic rocks such as the lujavrites in Ilimaussaq, Greenland, where sodalite is a primary mineral.
sial for many years. Taylor & Mackenzie (1975) studied differences in leucite morphologies with different cooling methods at 2 kb Pnz o in the system NaA1SiO4-KAISiO4-SiOz-H20. As shown in Fig. 19 the compositions of the leucites produced appeared to depend on their morphologies. From the common occurrence of natural pseudoleucites along the leucite-analcite join, Taylor & MacKenzie (1975) suggested that metastable Na-rich leucites are produced by extremely rapid alkali-ion exchange rather than by the breakdown of high-temperature leucites or by leucite breaking down to nepheline+ feldspar at the reaction point in this system.
Primary versus secondary origin of analcites Experimental studies bearing on the problem of the origin of analcite phenocrysts in a few rare volcanic rocks have been carried out by Roux & Hamilton (1976) in the system NaA1SiO4KA1SiO4-SiO2. They found that analcite coexisted with liquid above about 5 kb PH2o and 600~ supporting the concept of primary analcites. In contrast, Gupta & Fyfe (1975) found a very rapid reaction between leucite and saturated NaCI solutions at 1 kb Pn2 o at low temperatures. This supported the idea that these natural analcite phenocrysts could be transformed from leucites in a low-temperature hydrothermal environment. The plausibility of both of these mechanisms occurring in nature has been discussed by Edgar (1984).
Other experiments on felsic alkaline rocks Pseudoleucites The origin of pseudoleucites has been controver-
Dolfi et al. (1978) and Ruddock & Hamilton (1978) investigated the KA1Si206 (Lc)-
Zoned
c Ne
"x,~~Ks Greatly exsolved
leucite
Slightly ( I11r-='_lib IIIII
exsolved I--'l
leucite
~
FIG. 19. Recalculated microprobe analyses and morphologies of leucites crystallized from a bulk composition (+) crystallized in part of the system NaA1SiO4-KA1SiO4-SiO2-H20. Lines join spots probed on leucite crystals to compositions in the system. (After Taylor & Mackenzie 1975.)
Genesis of alkaline magmas CaMgSi20 6 (Di) and KA1Si206-CaMgSi20 6S i O 2 - H 2 0 - C O 2 systems respectively. In the former system, Dolfi et al. showed that increasing pressure up to 12 kb increased the primary diopside field at the expense of leucite. In experiments at 4 kb in the system KA1Si20 6CaMgSizO6-H20, Ruddock & Hamilton (1978, Fig. 19) indicated that with increasing pressure the fields of leucite and quartz contract while those of diopside and sanidine expand. Phlogopite is restricted to compositions with low diopside contents. On the basis of these results they explain the common phenocryst assemblage of diopside and phlogopite along with groundmass quartz and potash feldspar in many lamprophyres. High-pressure experiments up to 30 kb on the composition Lc60DisQ35 with H 2 0 and CO2 (Ruddock & Hamilton 1978, Fig. 19) confirmed these conclusions. These experiments also suggested that minettes are derived by partial melting of the mantle and that primary carbonate minerals in these rocks could have formed at mantle pressures. In order to determine the effect of CaA12Si2Os
49
(An) on phase relations in the system NaA1SiO4KA1SiO4-SiOz-H20, Norris & MacKenzie (1976) and Whiteley (1981) have determined phase relations at 1 kb PH20 with 3, 5 and 10 wt.% CaA12SizOs. With increasing An contents, the leucite and alkali feldspar fields are reduced as the plagioclase field expands. For the 3% An plane, Norris & MacKenzie (1976, Fig. 30) located a eutectic between nepheline, K-rich feldspar and Na-rich feldspar. For the 5% and 10% An planes (Norris & MacKenzie 1976, Figs 31 and 32) the projected phase volumes of Kfeldspar and Na-feldspar cut the leucite phase volume at progressively more K-rich compositions as the K-rich feldspar field contracts. In the absence of a Ca-bearing mafic mineral, such as diopside, these systems more closely approximate felsic magmas.
ACKNOWLEDGMENTS: I am grateful to M. Arima and A. Kolisnik for their assistance in the preparation of the diagrams for this paper. The Natural Science and Engineering Research Council of Canada provided financial support.
References ALLEN,J. C., BOETTCHER,A. L. & MARLAND,G. 1975. Amphiboles in andesite and basalt: I. Stability as a function of P - T - f 0 2 . Am. Mineral. 60, 1069-85. APPLEYARD,E. C. 1974. Syn-orogenic igneous alkaline rocks of eastern Ontario and northern Norway. Lithos, 7, 147-60. ARCULUS, R. J. 1975. Melting behaviour of two basanites in the range 10-35 kbar and the effect of TiO 2 on the olivine-diopside reactions at high pressures. Yearb. Carnegie Inst. Washington, 512-15. ARIMA, M. & EDGAR, A. D. 1983a. High-pressure experimental studies on a katungite and their bearing on the genesis of some potassium-rich magmas of the west branch of the African rift. J. Petrol. 24, 166-87. & - - 1983b. A high pressure experimental study of a magnesian-rich leucite-lamproite from the West Kimberley area, Australia: petrogenetic implications. Contrib. Mineral. Petrol. 84, 228-34. AYRTON, R. 1974. Rifts, evaporites and the origin of certain alkaline rocks. Geol. Rundsch. 63, 430-50. BAILEY, D. K. 1970. Volatile flux, heat focussing and the generation of magma. Geol. J. Spec. Issue 2, 177-86. 1974. Experimental petrology relating to oversaturated peralkaline volcanics, a review. Bull. Volcanol. 38, 637-52. Applications of experiments to alkaline rocks. In: BAILEY,D. K. & MACDONALD,R. (eds) The Evolution of the Crystalline Rocks, pp. 419-69. Academic Press, New York. 1980. Volatile flux, geotherms, and the generation 74,
-
-
-
-
- - 1 9 7 6 .
-
-
of the kimberlite-carbonatite-alkaline magma spectrum. Mineral. Mag. 43, 695-9. - 1982. Mantle metasomatism--continuing chemical change within the earth. Nature, Lond. 296, 525-30. --,TARNEY, J. & DUNHAM, K. C. (eds) 1980. Evidence for chemical heterogeneity in the earth's mantle. Phil Trans. R. Soc. Lond. Ser. A, 297, 134493. BARKER, D. S. 1976. Phase relations in the system NaA1SiO4-SiO2-NaC1-H20 at 400-800~ and 1 kilobar, and petrologic implications. J. Geol. 84, 77-106. - - 1983. Igneous Rocks, 417 pp. Prentice-Hall, Englewood Cliffs, NJ. BARTON, M. & HAMILTON,D. L. 1979. The melting relationships of a madupite from the Leucite Hills, Wyoming, to 30 kb. Contrib. Mineral. Petrol. 69, 133-42. & - - 1 9 8 2 . Water undersaturated melting experiments bearing upon the origin of potassiumrich magmas. Mineral. Mag. 45, 267-78. BASALTIC VOLCANISMSTUDY PROJECT, 1981. Basaltic Volcanism on the Terrestrial Planets, 1286 pp. Pergamon Press, Oxford. BEDSON, P & HAMILTON, n. L. 1981. Kimberlites, carbonatites and liquid immiscibility. Prog. Exp. Petrol., NERC Publ. Set. D, 29-32. BELL, K. & POWELL, J. L. 1969. Strontium isotopic studies of the Burunga and Toro-Ankole regions, east and central equatorial Africa. J. Petrol. 10, 536-72. BERGMAN,S. C. 1987. Lamproites and other potassium-
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rich igneous rocks: a review of their occurrence, mineralogy and geochemistry. In: FITTON, G. J. & UPTON, B. G. J. (eds) Alkaline Igneous Rocks, Geol. Soc. Spec. Publ. 30, pp. 103-90. BINSTED, N. 1981. The system Ab-Ne-NaC1-HzO. Prog. Exp. Petrol., NERC Publ. Ser. D, 34-35. BOWEN, N. L. 1922. The reaction principle in petrogenesis. J. Geol. 30, 177-98. --1928. Evolution of the Igneous Rocks, 251 pp. Dover, New York. BREY, G. 1976. CO2 solubility and solubility mechanisms in silicate melts at high pressures. Contrib. Mineral. Petrol. 57, 217-21. 1978. Origin of olivine melilitites--chemical and experimental constraints. J. Volcanol. geotherm. Res. 3, 61-88. & GREEN, D. H. 1975. The role of CO2 in the genesis of olivine melilitite. Contrib. Mineral. Petrol. 49, 93-103. & -1976. Solubility of COz in olivine melilitite at high pressures and the role of CO2 in the earth's upper mantle. Contrib. Mineral. Petrol. 55, 217-30. & -1977. Systematic study of liquidus phase relations in olivine melilitite + H20 + CO 2 at high pressures and petrogenesis of an olivine melilitite magma. Contrib. Mineral. Petrol. 61,141-62. BULTITUDE, R. J. & GREEN, D. H. 1968. Experimental study at high pressures on the origin of olivine nephelinite and olivine melilite nephelinite magmas. Earth planet. Sci. Lett. 3, 325-37. & -1971. Experimental study of crystal-liquid relationships at high pressures in olivine nephelinite and basanite compositions. J. Petrol. 12, 12147. CAWTI-IORN, R. G. 1976. Melting relations in part of the system CaO-MgO-A1203-SiO2-Na20-H20 under 5 kb pressure. J. Petrol. 17, 44-72. CUNDARI, A. & O'HARA, M. J. 1976. Experimental study at atmospheric and high pressures of a mafic leucitite from New South Wales, Australia. Prog. Exp. Petrol., N.E.R.C. Publ. Ser. D, 260-2. DALY, R. A. 1910. Origin of the alkaline rocks. Geol. Soc. Am. Bull. 21, 87-115. DOLFI, O., HAMILTON,D. H. & TREGILA, R. 1978. The system KA1Si206-CaMgSi206 at 4 and 12 kilobars. Prog. Exp. Petrol., NERC Publ. Ser. D, 21-2. DUKE, J. M. 1974. The effect of oxidation on the crystallization of an alkali basalt from the Azores. J. Geol. 82, 524-8. EDGAR, A. D. 1974. Experimental studies. In: SORENSEN, H. (ed.) The Alkaline Rocks, pp. 355-89. Wiley, New York. 1979. Mineral chemistry and petrogenesi: of an ultrapotassic-ultramafic volcanic rock. Contrib. Mineral. Petrol. 71, 171-5. 1984. Chemistry, occurrence and paragenesis of feldspathoids: a review. In: BROWN, W. L. (ed.) Feldspars and Feldspathoids, pp. 501-32. Reidel, Dordrecht. & ARIMA, M. 1983. Conditions of phlogopite crystallization in ultrapotassic volcanic rocks. Mineral. Mag. 47, 11-19. -& -1984. Experimental studies on K-metasomatism of a model pyrolite mantle and their -
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bearing on the genesis of ultrapotassic magmas. Proc. 27th Int. Geological Cong. on Petrology (Igneous and Metamorphic Rocks), Vol. 9, 509-41. V N U Science Press, Utrecht. CONDLIFFE,E., BARNETT,R. L. & SHIRRAN, R. J. 1980. An experimental study of an olivine ugandite magma and mechanisms for the formation of its K-enriched derivatives. J. Petrol. 21,475-97. - - , GREEN, D. H. & HIBBERSON, W. O. 1976. Experimental petrology of a highly potassic magma. J. Petrol. 17, 339-56. EGGLER, D. H. 1974. Effects of CO2 on the melting of peridotite. Yearb. Carnegie Inst. Washington, 73, 215-24. 1978. The effect of CO2 upon partial melting of peridotite in the system Na20-CaO-A1203-MgOSIO2-CO2 to 35 kb, with an analysis of melting in a peridotite-H20-CO2 system. Am. J. Sci. 278, 305-43. --& HOLLOWAY, J. R. 1977. Partial melting of peridotite in the presence of H20 and CO2: principles and review. Oreg. Dep. Geol. mineral Industries Bull. 96, 15-36. • WENDLANDT, R. F. 1979. Experimental studies of the relationship between kimberlite magmas and partial melting of peridotite. In: BOYD, F. R. & MEYER, H. O. A. (eds) Kimberlites, Diatremes and Diamonds: Their Geology, Petrology and Geochemistry, pp. 330-8. American Geophysical Union, Washington, DC. EL-GORESY, A. & YODER, H. S. 1974. Natural and synthetic melilite compositions. Yearb. Carnegie Inst. Washington, 73, 359-71. ELLIS, D. E. & WYLLIE, P. J. 1980. Phase relations and their petrological implications in the system MgOSIO2-H20-CO2 at pressures up to 100 kbar. Am. Mineral. 65, 540-56. FREESTONE, I. C. 1978. Liquid immiscibility in alkalirich magmas. Chem. Geol. 23, 115-23. --& HAMILTON, D. L. 1980. The role of liquid immisciblity in the genesis of carbonatites--an experimental study. Contrib. Mineral. Petrol. 73, 105-17. FREY, F. A., GREEN, D. H. & ROY, S. D. 1978. Integrated models of basalt petrogenesis: a study of quartz tholeiites to olivine melilitites from southeastern Australia utilizing geochemical and experimental petrological data. J. Petrol. 19, 463-513. GREEN, D. H. 1969, The origin of basaltic and nephelinitic magmas in the earth's mantle. Tectonophysics, 7, 409-22. 1970a. A review of experimental evidence on the origin of basaltic and nephelinitic magmas. Phys. Earth planet. Inter. 3, 221-35. The origin of basaltic and nephelinitic magmas. Trans. Leicester Lit. Phil. Soc. 44, 26-54. 1971. Composition of basaltic magmas as indicators of conditions of origin: applications to oceanic volcanism. Phil. Trans. R. Soc. Lond. Ser. A, 268, 707-25. 1973a. Conditions of melting of basanite magma from garnet peridotite. Earth planet. Sci. Lett. 17, 456-65. Experimental melting studies on model -
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Genesis of alkaline magmas upper mantle compositions at high pressure under both water-saturated and water-undersaturated conditions. Earth planet. Sci. Lett. 19, 37-53. - - , GUPTA, A. K. & TAYLOR, W. R. 1984. Experimental studies on the effects of H20 and CO2 on liquidus relations in Fo-Ks-SiO2 and F o - N e SiO2 at 28 kb. Abstracts, Workshop on Experimental Geochemistry, Monash University, Melbourne. HmBERSON, W. O. 1970. Experimental duplication of conditions of precipitation of high pressure phenocrysts in a basaltic magma. Phys. Earth planet. Inter. 3, 247-54. & RINGWOOD, A. E. 1967. The genesis of basaltic magmas. Contrib. Mineral. Petrol. 15, 103-90. GUPTA, A. K. & FYFE, W. S. 1975. Leucite survival: the alteration to analcime. Can. Mineral. 13, 3613. HAMILTON, D. L., FREESTONE, I. C., DAWSON, J. B. & DONALDSON, O. H. 1979. Origin of carbonatites by liquid immiscibility. Nature, Lond. 279, 52-4. HARDS, N. 1976. Distribution of elements between the fluid phase and silicate melt phase of granites and nepheline syenites. Prog. Exp. Petrol., NERC Publ. Ser. D, 88-90. HELZ, R. T. 1973. Phase relations of basalts in their melting range at PH2o=5 kb as a function of oxygen fugacity. Part I. Mafic phases. J. Petrol. 14, 249-302. 1976. Phase relations of basalts in their melting ranges at PH2o = 5 kb. Part II. Melt compositions. J. Petrol. 17, 139-93. HOLMES, A. 1932. The origin of igneous rocks. Geol. Mag. 69, 543-58. 1950. Petrogenesis of katungite and its associates. Am. Mineral. 35, 772-92. -1965. Principles of Physical Geology (2nd edn), 623 pp. Ronald Press, New York. --& HARWOOD, F. 1937. The petrology of the volcanic rocks of Bufumbira. Mere. Geol. Surv. Uganda, 3 (2), 1-300. HUANG, W. L. & WYLLIE, P. J. 1974. Eutectic between wollastonite I1 and calcite contrasted with thermal barrier in i g O - S i O 2 - C O 2 at 30 kilobars with emphasis to kimberlite-carbonatite petrogenesis. Earth planet. Sci. Lett. 24, 305-10. IRVING, A. J. & GREEN, D. H. 1972. Experimental study of phase relationships in a high pressure mugearitic basalt as a function of water content. Geol. Soc. Am. Meet. Abstracts with Programs 4, 550-1. - - - & PRICE, R. C. 1981. Geochemistry and evolution of high pressure phonolitic rocks from Nigeria, Australia, Eastern Germany and New Zealand. Geochim. cosmochim. Acta, 45, 1309-20. ITO, K. & KENNEDY, G. C. 1968. Melting and phase relations in the plane tholeiite-lherzolite-nepheline basanite to 40 kilobars with geological implications. Contrib. Mineral. Petrol. 19, 177-211. KAY, R. W. & GAST,P. W. 1973. The rare earth content and origin of alkali rich basalts. J. Geol. 81, 65382. KOGARKO, L. N. & ROMANCHEV, B. P. 1982. Phase equilibria in alkaline melts. Zap. vses. mineral. Ova. 8, 167-182 (in Russian). -
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KUEHNER, S. M., EDGAR, A. D. & ARIMA, M. 1981. Petrogenesis of the ultrapotassic rocks from the Leucite Hills, Wyoming. Am. Mineral. 66, 663-77. KUSHIRO, I. 1973. Origin of some magmas in oceanic and circum-oceanic regions. Tectonophysics, 17, 211-22. -& YODER, n . S. 1974. Formation of eclogite from garnet lherzolite: liquidus relations in a portion of the system MgSiO3-CaSiO3-A1203 at high pressures. Yearb. Carnegie Inst. Washington, 73, 266-9. LARSEN, E. S. 1940. Petrographic province of central Montana. Geol. Soc. Am. Bull. 51,887-948. LLOYD, F. E. 1981. Upper mantle metasomatism beneath a continental rift: Clinopyroxenes in alkalic mafic lava and nodules from South West Uganda. Mineral. Mag. 44, 315-24. & BAILEY, D. K. 1975. Light element metasomatism of the continental mantle: the evidence and the consequences. Phys. Chem. Earth, 9, 389-416. , ARIMA,M. & EDGAR, A. D. 1985. Partial melting of a phlogopite-clinopyroxenite nodule from south-west Uganda: an experimental study bearing on the origin of highly potassic continental rift volcanics. Contrib. Mineral. Petrol., 91, 321-9. LUTH, W. C. 1967. Studies in the system of KA1SiO4Mg2SiO4-SiO2-H20. I. Inferred phase relations and petrologic applications. J. Petrol. 8, 372-416. MCKENZlE, D. P. 1984. The generation and compaction of partially molten rock. J. Petrol. 25, 713-65. MERRILL, R. B. & IRVING, A. J. 1977. Chemistry and phase relations of an orthopyroxene-bearing transitional alkalic basalt (abstract). Los, 58, 526. --& WYLLIE, P. J. 1975. Kaersutite and kaersutite eclogite from Kakanui, New Zealand--water excess and water deficient melting relations to 30 kilobars. Geol. Soc. Am. Bull. 86, 555-70. MYSEN, B. O. & KUSHIRO, I. 1977. Compositional variations of coexisting phases with degree of melting of peridotite in the upper mantle. Am. Mineral. 62, 843-65. NORRIS, G. & MACKENZIE,W. S. 1976. Phase relations in the system NaA1SiO4-KAISiO4-CaA12Si2OsSiO2 at PH2o = 1 kb. Prog. Exp. Petrol., NERC Publ. Set. D, 79-81. O'HARA, M. J. 1968. The bearing of phase equilibria studies on synthetic and natural systems on the origin and evolution of basic and ultrabasic rocks. Earth Sci. Rev. 4, 69-133. & BIGGAR, G. M. 1969. Diopside-spinel equilibria, anorthite and forsterite reaction relationships in silica-poor liquids in the system C a - M g O A1203-SiO 2 at atmospheric pressure and their bearing on the genesis of melilitites and nephelinites. Am. J. Sci. 267A, 364-90. YODER, H. S. 1967. Formation and fractionation of basic magmas at high pressures. Scott. J. Geol. 3, 67-117. PRESNALL, D. C., DIXON, J. R., O'DONNELL, T. H. & DIXON, S. A. 1979. Generation of mid-ocean ridge tholeiites. J. Petrol. 20, 3-35. , DIXON, S. A., DIXON, J. R., O'DONNELL, T. H., BRENNER, N. J., SCHROCK, R. L. & DYCUS, D. W. 1978. Liquidus phase relations on the join diopside-forsterite-anorthite from 1 atm to 20 kbar; -
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their bearing on the generation and crystallization of basaltic magma. Contrib. Mineral. Petrol. 66, 203-220. RINGWOOD, A. E. 1975. Composition and Petrology of the Earth's Mantle, 618 pp. McGraw-Hill, New York. ROEDDER, E. 1979. Silicate liquid immiscibility in magmas. In: YODER, H. S. (ed.) The Evolution of the Igneous Rocks: Fiftieth Anniversary Perspectives, pp. 15-58. Princeton University Press, Princeton, NJ. Roux, J. & HAMILTON, D. L. 1976. Primary igneous analcite--an experimental study. J. Petrol. 17, 244-57. RUDDOCK, D. I. & HAMILTON, D. L. 1978. The system KA1SizO6-CaMgSi206-SiO2-H20 at 4 kilobars. Prog. Exp. Petrol., NERC Publ. Ser. D, 25-31. RYABCHIKOV, I. D. & BOETTCHER,A. L. 1980. Experimental evidence at high pressure for potassic metasomatism in the mantle of the earth. Am. Mineral. 65, 915-19. & GREEN, D. H. 1978. The role of carbon dioxide in the petrogenesis of highly potassic magmas. In: Problems of Earth's Crust and Upper Mantle, Trudy Instituta Geologii GeofizikL So An SSR 403, Nauka, Novosibirsk (in Russian). , SCHREYER, W. & ABRAHAM, K. 1982. Compositions of aqueous fluids in equilibrium with pyroxenes and olivines at mantle pressures and temperatures. Contrib. Mineral. Petrol. 79, 80-4. SCHAIRER, J. F. & YODER, H. S. 1970. Critical planes and flow sheet for a portion of the system CaOA1203-SiO 2. Yearb. Carnegie lnst. Washington, 68, 202-14. , TILLEY, C. E. & BROWN, G. M. 1969. The join nepheline-diopside-anorthite and its relation to alkali basalt fractionation. Yearb. Carnegie Inst. Washington, 66, 467-71. , YAGI, K. & YODER, H. S. 1962. The system nepheline-diopside. Yearb. Carnegie lnst. Washington, 61, 96-8. SORENSON, H. (ed.). 1974. The Alkaline Rocks, 622 pp. Wiley, New York. TAKAHASHI, E. 1980. Melting relations of an alkali olivine basalt to 30 kbar, and their bearing on the origin of alkali basalt magmas. Yearb. Carnegie Inst. Washington, 79, 271-6. - & KUSHIRO, I. 1983. Melting of a dry peridotite at high pressures and basalt magma genesis. Am. Mineral. 68, 859-79. TAYLOR, D. & MACKENZIE,W. S. 1975. A contribution to the pseudoleucite problem. Contrib. Mineral. Petrol. 49, 321-33. THOMPSON, R. N. 1972. The one atmosphere melting patterns of some basaltic volcanic series. Am. J. Sci. 272, 901-32. - - 1 9 7 3 . One atmosphere melting behaviour and nomenclature of terrestrial lavas. Contrib. Mineral. Petrol. 41, 197-204. - 1974. Primary basalts and magma genesis. I. Skye, North-West Scotland. Contrib. Mineral. Petrol. 45, 317-41.
Primary basalts and magma genesis. III. Alban Hills, Roman comagmatic province, central Italy. Contrib. Mineral. Petrol. 60, 91-108. TURNER, F. J. & VERHOOGEN, J. 1960. lgneous and Metamorphic Petrology, 694 pp. McGraw-Hill, New York. WADE, A. & PRIDER, R. T. 1940. The leucite-bearing rocks of the West Kimberley area, Western Australia. Q.J. geol. Soc., Lond. 96, 39-98. WATERS, A. C. 1955. Volcanic rocks and the tectonic cycle. Geol. Soc. Am. Spec. Pap. 63, 703-722. WENDLANDT, R. F. & EGGLER, D. H. 1980a. The origins of potassic magmas. 1. Melting relations in the systems KA1SiO~-Mg2SiO4-SiO2 and KA1SiO4-MgO-SiO2-CO2 to 30 kilobars. Am. J. Sci. 280, 385-420. - & 1980b. The origins of potassic magmas. 2. Stability of phlogopite in natural spinel lherzolite and in the system KA1SiO4-MgO-SiOz-H20CO2 at high pressures and temperatures. Am. J. Sci. 280, 421-58. WIalTELEY, C. 1981. Phase relations in the system NaA1SiO4-KA1SiO4-CaA12Si2Os-SiO2-HzO at 1 kb. Prog. Exp. Petrol., NERC Publ. Ser. D, 33-4. WILLIAMS, H. 1936. Pliocene volcanoes of the NavajoHopi county. Geol. Soc. Am. Bull. 47, 111-72. WYLLIE, P. J. 1975. The earth's mantle. Sci. Am. March 1975. - - 1 9 7 7 . Mantle fluid compositions buffered by carbonates in peridotite-CO2-H20. J. Geol. 8 5 , 187-207. - - 1 9 7 8 . Mantle fluid compositions buffered in peridotite-H20--CO2 by carbonates, amphibole and phlogopite. J. Geol. 86, 687-713. - - 1 9 7 9 . Magmas and volatile components. Am. Mineral. 64, 469-500. - 1980. The origin of kimberlite. J. geophys. Res. 8 5 , 6902-10. - & HUANG, W. L. 1975. Peridotite, kimberlite, and carbonatite relations in the system CaO-MgOSiO2-COz. Geology, 3, 621-4. - - & 1976. Carbonation and melting reactions in the system CaO-MgO-SiO2-CO2 at mantle pressures with geophysical and petrological applications. Contrib. Mineral. Petrol. 54, 79-107. YODER, H. S. 1975. Relationship of melilite-bearing rocks to kimberlite: a preliminary report on the system akermanite-CO2. Phys. Chem. Earth, 9, 883-94. -1976. Generation of Basaltic Magmas, 265 pp. National Academy of Sciences. Washington, DC. - 1979. Melilite-bearing rocks and related lamprophyres. In." YODER, H. S. (ed.) The Evolution of Igneous Rocks. Fiftieth Anniversary Perspectives, pp. 391-411. Princeton University Press, Princeton, NJ. - & KUSrtIRO, I. 1972. Composition of residual liquids in the nepheline-diopside system. Yearb. Carnegie Inst. Washington, 71,413-16. ZYRYANOV, V. N. 1981. Phase Correspondence in the Systems of Feldspars and Feldspathoids, 217 pp. Nauka, Moscow (in Russian). - - 1 9 7 7 .
A. D. EDGAR, Department of Geology, University of Western Ontario, London, Ontario, N6A 5B7, Canada.
Nephelinites and carbonatites M. J. Le Bas S U M M A R Y : Carbonatites found in strongly alkaline intra-plate petrographic volcanic provinces are associated with olivine-poor nephelinites and with phonolites. Olivine-rich nephelinites occur in basanitic and alkali basalt provinces, normally without carbonatites. These igneous provinces are marked by epeirogenic crustal uplift. Nephelinites, ijolites and carbonatites form discrete magmatic events within individual complexes and correspond to the silicate-carbonate conjugate immiscible liquids observed in the laboratory. Carbonate liquids of variable alkali content can separate from both nephelinitic and phonolitic liquids. The silicate liquids give rise to pyroxenites, ijolites, nepheline syenites and nephelinitic pyroclast-rich strato-volcanoes. The carbonate liquids lose alkalis and fractionate to s6vite, alvikite and ferrocarbonatite, each with increasing incompatible-element content. Further fractionation of carbonatite magma can produce mineralizing fluids rich in rare-earth elements, F, Ba, U and Th. Dolomite carbonatite forms only in the deeper parts of carbonatitic complexes, perhaps at depths greater than 2 km. Explosive carbonatite volcanism can occur giving widespread carbonate tufts, rarely with lavas. Fenitization characterizes ijolite-carbonatite intrusive complexes. Syenitic fenites containing alkali feldspars, sometimes perthitic, are formed as aureoles 500 m wide around ijolites in granitic terranes, with nepheline syenite commonly formed as a contact-reaction rock. The feldspar-rich syenitic fenite aureoles which develop around the early carbonatites usually contain pure orthoclase or pure albite.
Carbonatites achieved geological respectability only about 30 years ago, largely through the impact of work in Scandinavia and Africa. Br6gger's work (1921), at Fen in S Norway, which was strongly disputed by Bowen (1924) but later re-established by Saether (1957), demonstrated that calcite-rich rocks could have all the characteristics of intrusive magmatic bodies and that they alkali metasomatize (fenitize) their host rocks. In Sweden, von Eckermann (1948) conducted a meticulous study of the Aln6 complex. He maintained that alkali carbonatite magma was the prime intrusive material at Aln6, that the associated alkaline silicate rocks were all metasomatic products of fenitization and that the calcite-rich carbonatite was the residue of the intrusive carbonate magma. Dixey et al. (1935) were the first to identify carbonatites in Africa. Garson re-investigated the Malawi complexes of Chilwa, Tundulu, Kangankunde, Songwe and others, and wrote a series of memoirs culminating in a valuable summary account (Garson 1965). The association of carbonatites with nephelinitic volcanism was first recognized by King (1949) at the Napak complex in Uganda. Tuttle & Gittins (1966) and Heinrich (1966) wrote comprehensive accounts of intrusive carbonatites, but extrusive carbonatites have only recently become widely appreciated although they were first identified in 1960 by Dawson (1962).
Strongly alkaline igneous rocks are notorious for their diversity of rock type and abundance of names. Some names, such as ijolite and urtite, are necessary, but many can be discarded. The names used here are based on the work of the IUGS Subcommission on the Systematics of Igneous Rocks (Streckeisen 1976, 1978). Carbonatites are defined as igneous carbonate rocks with more than 50% modal carbonate minerals. If the carbonate is calcite, the rock is a calcite carbonatite which is called a s6vite if it is coarse grained with adcumulate texture or a micro-s6vite if it is medium or fine grained. Beforsite is equivalent to dolomite carbonatite. A less common variety of carbonatite is alvikite, which typically has a texture of small calcite rhombs packed together and is enriched in incompatible elements such as rare-earth elements (REE), Ba, Mn, Zn and commonly Nb. The boundary between s6vite and alvikite is still uncertain but may be set at 0.4% MnO, 1500 ppm Ba and 2000 ppm REE. The calcite of alvikite is slightly ferroan. Ferrocarbonatite is a ferroan calcite to ankeritic calcite carbonatite even more strongly enriched in some or all of REE, Ba, Mn, Fe, Zn, F and U, with possible lower limits of 1.0% MnO, 5000 ppm REE and 5000 ppm Ba. This rock type is very rare, but it is the main source for mineralization of these elements in carbonatites. Natrocarbonatite is carbonatite lava. It is fine grained, extrusive and composed of nyerereite ((Na,K)2Ca(CO3)2) and gregoryite
From: FITTON,J. G. & UPTON, B. G. J. (eds), 1987, Alkaline Igneous Rocks, Geological Society Special Publication No. 30, pp. 53-83.
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((Na,K)2CO3). Some natrocarbonatites are richer in Ca and have calcite in place of gregoryite. The purpose of this contribution is to give a petrological synthesis of the nephelinite-carbonatite complexes, at both plutonic and volcanic levels. It begins with the rise of the primary magma through the uppermost mantle, then considers the processes suffered by the magma when it stops in reservoirs at or near the base of the crust and finally traces the passage of the different magmas, carbonate and silicate, through the crust to the ultimate products at the surface. Kimberlitic carbonatites are not included because they have recently been reviewed by Dawson (1980) and because they are different from nephelinitic carbonatites in their primary lack of alkalis, particularly Na. This lack is expressed by the absence of fenitization around the kimberlitic carbonatites. However, the zonation of the E African and E Siberian ultraalkaline provinces, recently compared by Le Bas (1986), indicates that these two types of carbonatite are genetically related in their mantle sources.
Distribution Although nephelinite-carbonatite complexes are rare, being only a fraction of 1~ of all igneous rocks, their distribution is worldwide (Fig. 1) with ages back to the early Proterozoic. Nephelinite-carbonatite complexes occur in intra-plate environments, both continental and, rarely, oceanic. For instance they occur in the Canary 150 r
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90 ~
89
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30 T
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Islands and the Cape Verde Islands, where there is a clear regional association with ocean island basalts. Petrographically there is no distinction between the complexes that occur singly, for example as at Kaiserstuhl in S Germany, and those in groups such as the complexes composing the 20 Cenozoic carbonatitic volcanic centres of the inner portion of the E African province (Le Bas 1980). The E African province (Fig. 2) is zoned, with kimberlites in a central core about 300 km across. Carbonatites, olivine-poor nephelinites and ijolites, but not basalts, comprise the next inner concentric zone 100-300 km wide, and at the edge is a carbonatite-free outer zone ranging in width from 100 to 300 km which is composed of basanites, alkali basalts and some olivine-rich nephelinites (Fig. 3). The large Palaeozoic-Mesozoic E Siberian province is similarly zoned (Dawson 1980; Le Bas 1986). In the Angola province, or half-province if S Brazil is rejoined to it, the distribution is different. There the alkali basaltic centres cluster near the coast line, with carbonatitic complexes further inland and kimberlite diatremes furthest inland. Namibia is similar (Prins 1981). The scale of these distributions suggests deep-seated mantle processes. They are not related to plate motion or plume traces. These alkaline and ultra-alkaline provinces are also marked by epeirogenic uplift (Le Bas 1971, 1980; Baker 1987) which produces broad swells and plateau uplift both continental (e.g. the Rhine upwarp and the East African plateau) and oceanic (e.g. Cape Verde Rise). These uplifts
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FIG. ]. World distribution of nephelinite-carbonatRe complexes: I I , r215
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FIG. 2. Map of the E African Alkaline Province (or Lake Victoria Province) showing the zonation ( - - - ) : 1, central zone of kimberlite diatremes (K); 2, inner zone of ultra-alkaline complexes (U); 3, outer zone of alkali basaltic volcanism (A). Radial lines indicate the edge of the E African uplift. Bold lines indicate rift faults. The faults show no apparent relation to the zonation. The inset square marks the area shown in Fig. 3. (RW, Rwanda; BU, Burundi.)
evidently relate to the same deep-mantle processes as those giving rise to the ultra-alkaline magmatism (Bailey 1987). Such up-arching produces tension in the upper crust which is quite independent of the motion of the whole tectonic plate. The two motions, the first vertical, covering an equidimensionat area and emanating from the mantle associated with ultra-alkaline magmatism, and the second essentially horizontal and related to plate motion, are interpreted to be distinct and not necessarily related (Le Bas 1986). It is argued that ultra-alkaline magmatism and regional uplift are not necessary initial stages of the Wilson cycle which describes the break-up of continents followed by sea-floor spreading.
Nephelinite magma There are nephelinites and nephelinites, the two varieties being distinguished by the properties given in Tables 1-3. Those defined as Group I
nephelinites are olivine rich, and are typically associated with alkali basalt and basanite volcanic provinces. Group II nephelinites are rarer and comprise olivine-poor nephelinites which occur with carbonatitic and ijolitic complexes (Fig. 3). The olivine-poor nephelinites commonly have abundant euhedral clinopyroxene phenocrysts, whereas the few olivine phenocrysts often show resorbed crystal margins, sometimes even reaction rims, indicating that olivine was no longer in equilibrium. The abundance of pyroxene usually means that these rocks are metanephelinites rather than nephelinites, and many could be better described as clinopyroxenephyric melanephelinites. Aphanitic varieties are rare. Olivine-poor nephelinite can fractionate to olivine-free nephelinite and to phonolite (Table 1). These phonolites are often glassy, commonly making abundant dykes and plugs, and Lippard (1973) distinguished them (his Gwasi-type phonolites) from the phonolites in and around the
56
M. J. Le Bas 134o +
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FIG. 3. Map of Kenya and adjacent areas showing the present distribution of the Group II olivine-poor nephelinitic (and carbonatitic) volcanic centres corresponding to a portion of the inner zone shown in Fig. 2, and the Group I olivine-rich nephelinitic (non-carbonatitic) volcanic centres corresponding to the Kenyan region of the outer zone shown in Fig. 2. Between these two zones lies a zone of mixed nephelinitic types resulting from an overlap or mixing of the inner- and outer-zone magmas.
Kenya rift valley (his Plateau and Kenya types). The latter are derived from basanitic and basaltic sources, and form extensive lavas evidently of low viscosity. The phonolites derived from the nephelinites are rich in Sr (more than 1000 ppm) and Ba (more than 100 ppm, and some more than 1000 ppm), and are richer in total alkalis (about 14%) relative to silica (about 53%) than those derived from plagioclase-bearing m a g m a s w h e n both Sr and Ba contents are less than 100 ppm.
The Miocene lavas of the Kisingiri volcano in W Kenya provide a good example of the normal fractionation trend from olivine-poor nephelinites and they are not dissimilar from those at other centres nearby except in detail. The early stages of fractionation are dominated by the precipitation of m u c h clinopyroxene plus a few oxides and some olivine at first (Fig. 4), producing only a m i n i m a l rise of Si in the m a g m a with fractionation. On some occasions Si falls, partic-
Nephelinites and carbonatites
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FIG. 4. CaO-MgO plot of fields for olivine-rich nephelinites (Group I) and olivine-poor nephelinites or melanephelinites (Group II), and their mutual fractionation product of ordinary nephelinites and phonolites. • is the composition of olivine fractionated; + - - + are successive compositions of clinopyroxenes fractionated from melanephelinites, nephelinites and phonolites. Dotted lines are paths of fractionation, and to the right of the fields are crystal extracts with the pyroxene-to-olivine ratio indicated. (Data from Le Bas (1978), Varne (1968), Saggerson (1970), Spencer (1969), Strong (1972) and EAGRU data file (unpublished.)
ularly if no oxides are precipitated, and melilite nephelinite magma can form. Some data are given in Table 1, and further data can be obtained from the U K - I G B A data base which can be accessed via BGS (London) or the World Data Center-A (Colorado). They show that the production of phonolite from olivine-poor nephelinite is accompanied by a fall in Ca, Mg, total Fe, P and sometimes Ti, with a rise in Si, A1, Na, K and the ratios Fe/Mg and K/Na. The fractionation of trace elements from olivine-poor nephelinites can be shown on twoelement plots, but is best seen on primordial mantle-normalized plots (Figs 5 and 6). Figure 5 is a plot of four samples from Kisingiri (data in Table 1) and shows a marked relative increase in the incompatible elements Th, Nb and Zr, with depletion in Sr, P, Ti and V; the latter is pulled out by the precipitation of pyroxene, apatite and any titanomagnetite. The increase in the ratio Zr/Ti appears to be a sensitive index of fractionation, and a similar rotation of the Zr-Ti tie-line (i.e. Zr/Ti increases) is also seen from the trace data from other complexes. Figure 6 compares melanephelinites of carbon-
atitic complexes of different ages in E Africa and reveals that, while the high-field-strength elements Nd, P, Zr, Ti and Y remain closely similar, the others vary slightly, particularly the largeion-lithophile (LIL) elements. The younger complexes of Homa Mountain and Elgon both show a slight enrichment in Ba and Th, and Ba mineralization is seen at Homa (Le Bas 1977, p. 237). If data ranges were marked on Fig. 6(a), the values of Rb, Sr, Nb and REE would show overlap, except in the case of Homa Mountain where all the data consistently show higher REE and lower K values. This is not instrument bias since all the data were obtained on the X-ray fluorescence (XRF) spectrometers at both Edinburgh and Leicester University. If the Kisingiri and N a p a k averages plotted in Fig. 6(a) are taken as normal for nephelinitic rocks associated with carbonatites (but whether these are nephelinites from before or after possible liquid-immiscible separation from carbonate melt is not yet certain), they should be chemically comparable with other carbonatitic nephelinites. Appropriate data are sparse, but they do compare closely with the Triassic nephel-
58
M. J. Le Bas
TABLE 1. Selected nephelinitic lavas from carbonatitic and other provinces Elgon, WKenya
Kisingiri, W Kenya
Napak, EUganda
RR659 RR585 RR457 U1041
SUN6 SUN89 SUN21
ElK8
ElK6 ElK7
54.73 0.47 21.37 5.13 0.23 0.21 2.70 9.74 5.72 0.16
43.50 2.03 10.30 15.03 0.18 9.18 16.94 1.26 1.68 0.58
42.84 2.77 11.47 15.46 0.29 7.16 15.33 2.76 1.27 0.99
42.20 44.12 2.48 2.61 12.04 13.89 1 3 . 9 7 14.11 0.27 0.25 6.22 5.03 1 5 . 1 7 11.42 3.93 5.52 2.54 2.78 0.88 0.95
101.27 100.46
100.68
97.76
97.23
100.35
99.68
100.68
82 92 333 31 6 749 1098 46 49 88 38 135 18
19 121 291 54 7 1251 1167 67 69 132 50 184 25
12 133 260 55 8 1212 1267 73 74 133 56 206 26
29 117 286 60 15 1857 1946 163 96 154 57 271 36
29 114 211 45 12 1211 910 111 92 158 61 193 25
30 121 218 58 12 1128 684 123 94 157 57 214 28
Major elements (wt.%) SiO 2 TiO 2 AI20 3 Fe203 T MnO MgO CaO Na20 K20 P205
Total
41.99 3.08 12.94 15.22 0.24 7.43 13.65 3.46 1.58 0.89
45.42 49.73 2.56 1.95 1 4 . 8 8 16.72 13.46 10.62 0.23 0.22 4.39 2.36 9.17 6.75 7.26 8.35 2.80 3.82 0.93 0.75
100.48 101.10
40.38 42.56 2.52 2.48 1 1 . 6 8 13.62 1 6 . 2 7 14.48 0.25 0.23 6.27 5.52 1 2 . 5 3 10.76 4.42 4.58 2.33 2.50 0.81 0.79
Trace elements (ppm) Ni Zn V Rb Th Ba Sr Nb La Ce Nd Zr Y
47 114 323 40 12 1862 3135 119 105 194 79 284 32
25 119 238 63 11 1161 1342 111 92 174 70 317 29
9 129 143 81 15 1401 1448 132 95 161 132 362 30
4 168 52 141 56 1139 808 336 74 135 40 816 24
RR659, aphanitic melanephelinite lava, top of Kaniamwia, Kisingiri (Leicester XRF); RR585, nephelinite lava, Rangwe, Kisingiri (Leicester XRF); RR457, nephelinite dyke, with katophorite, Kiawindu, Kisingiri (Leicester XRF); U 1041, phonolite lava, trachytic texture, Nyamaji, Homa Bay (Leicester XRF); SUN6, pyroxene-phyric melanephelinite lava, Napak (wet analysis by R. C. Tyler 1966; traces Edinburgh XRF); SUN89, melanephelinite lava, Napak (Leicester XRF); SUN21, aphanitic melanephelinite lava, Okathim (Edinburgh XRF); ELK8, melanephelinite lava above Kitum Cave, E flanks Mount Elgon (Leicester XRF); ELK6, melanephelinite lava, Kitum Cave, E flanks Mount Elgon (Leicester XRF); ELK7, nephelinite lava, Kitum Cave, E flanks Mount Elgon TABLE 2. Worm average nephelinites
SiO2 TiO2 A1203 Fe203 FeO MnO MgO CaO Na20 K20 P205 H20
Group I (olivine rich) (wt.% (SD))
Group II (olivine poor) (wt.~(SD))
39.8 (1.5) 3.0 (0.3) 11.3 (0.9) 5.1 (1.9) 7.9 (1.2) 0.2 (0.01) 12.3 (1.8) 12.8 (1.3) 3.4 (0.3) 1.1 (0.3) 0.9 (0.6) 2.1 (0.6)
41.5 (2.0) 2.8 (0.3) 11.4 (1.6) 7.4 (1.5) 6.0 (1.0) 0.2 (0.01) 8.1 (2.0) 13.2 (3.2) 3.7 (1.9) 1.8 (1.5) 0.8 (0.6) 3.0 (1.1)
99.9 n=81
99.9 n--97
After Le Bas 1978.
inites of the Karoo province with w h i c h carbonatites are associated (Bristow & Saggerson 1983). The Miocene nephelinites of the southern portion (south of the equator) of the K e n y a rift valley, particularly those in the Kishalduga region, are almost identical (cf. Figs 6(a) and 6(b), and have olivine contents m i d w a y between the rich and poor types defined above (Kishalduga data from R. Tarzey, personal communication). Miocene nephelinites of N K e n y a are m a r k e d l y poorer in K and the more incompatible trace elements. N o olivine-free nephelinites have yet been observed in the northern rift. The few data that are available from the olivine-rich nephelinites of the Massif Central in France and of Greenland, Hawaii and the Comores Islands are all closely similar to the carbonatitic nephelinites, apart from variable K / T h ratios and Rb and K contents.
Nephelinites and carbonatites
Homa Mt., W. Kenya
N Kenya rift
S Kenya rift
59
HF566
HC42
11.108
8.428
15.591
17.548
Massif central 42466
Hawaii
Comores
65KAPAAI 1
Moh24
45.17 2.94 11.16 12.06 0.13 6.48 11.81 2.84 1.75 0.79
53.28 0.59 21.14 4.23 0.16 0.48 3.33 9.97 4.82 0.18
41.71 2.23 11.52 12.22 0.18 12.72 12.64 3.06 0.77 0.41
42.84 2.13 12.53 11.74 0.18 10.64 12.01 3.99 1.27 0.49
42.78 4.05 10.09 15.13 0.20 9.13 15.79 1.62 1.16 0.73
43.46 3.73 9.82 15.49 0.21 8.54 14.67 2.81 1.36 0.68
43.61 3.62 8.80 11.08 0.15 13.44 11.45 2.38 3.66 0.82
40.15 2.55 11.45 14.40 0.22 12.77 12.10 3.51 0.77 0.80
41.37 2.71 11.49 13.89 0.24 11.03 12.10 3.39 1.56 1.00
95.12
98.18
97.45
97.81
100.69
100.76
99.01
98.67
98.78
37 113 350 46 10 4668 1213 105 92 170 71 249 25
4 120 62 104 33 1224 1153 143 103 151 37 490 26
214 78 317 9 7 551 604 69 56 108 39 147 23
180 83 297 17 10 697 685 80 73 135 49 166 24
135 109 357 45 13 623 915 128 105 202 85 345 33
121 114 344 47 13 651 1075 116 96 183 74 335 29
309 80 -78 9 824 764 93 79 174 73 409 23
340 107 290 20 6 780 1050 -53 110 56 171 --
168 127 -47 -1250 880 ----257 --
(Leicester XRF); HF566, pyroxene-phyric melanephelinite plug, N flanks Homa Mountain (Edinburgh XRF); HC42, feldspar-phyric phonolite plug, S flanks Homa Mountain (Edinburgh and Durham XRF); 11.108, olivine nephelinite lava, Kasorogol, S Turkana (Edinburgh XRF); 8.428, olivine nephelinite, Nasaken, S. Turkana (Edinburgh XRF); 15.591, olivine nephelinite, S. Nakuru (Edinburgh XRF); 17.548, olivine nephelinite, Kishalduga, Narok (Leicester XRF); 42466, olivine nephelinite, Cantal, France (Downes 1984); 65APAAll, olivine nephelinite, Honolulu (Clague & Frey 1982); MOH24, olivine nephelinite, Moheli Island, Comores (Strong 1972). T h e similarities of the trace-element signatures in Fig. 6 and Table 1 indicate broadly similar chemical sources for the m a g m a s in the mantle, w h e t h e r or not carbonatites are associated. A c h a r a c t e r c o m m o n to all nephelinites is the e n r i c h m e n t in K , Th, N b and the light R E E w h i c h points to a m a n t l e source rich in the heatproducing elements, although the K / T h ratio does vary, indicating a variable m a n t l e source for these components. A further study of the detailed variation in nephelinitic and other primitive m a g m a s across the zoned alkaline igneous province of E Africa, c o m p a r i n g early Cenozoic with later Cenozoic m a n t l e m a g m a t i c products, is being carried out at present by R. Tarzey. A similar investigation of the Cape V e r d e m a g m a tism is being u n d e r t a k e n by N. Hodgson. The olivine-rich and olivine-poor nephelinites, in terms of their m a j o r - e l e m e n t chemistry, could
be derived from e a c h other, but their geographical separation by some h u n d r e d s of kilometres into the c o n t e m p o r a n e o u s outer and inner zones of the E A f r i c a n province shows that distinct processes occurred in the two cases.
Magma plumbing These two types of nephelinites are distinct chemically and physically (Table 3). The olivinerich nephelinites have high Mg and N i contents a n d few or no fractionation products. T h e y build small volcanic cones with an occasional abund a n c e of lherzolite nodules, w h i c h m a y be t a k e n to indicate that these m a g m a s c a m e directly from the m a n t l e with few or no intervening crustal reservoirs. In strong contrast, the olivine-poor nephelin-
M. J. Le Bas
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FIG. 5. Primordial mantle (PM) normalized plot of four selected lavas from the Kisingiri strato-volcano (Table 1). Note the decreasing coherence of distribution, particularly of the high-field-strength elements during fractionation. The PM values of Wood et al. (1979) were used; V was taken as 84 ppm.
ites make large central strato-volcanoes and almost never carry mantle-type xenoliths. They exhibit strong fractionation to phonolite and ijolite, which suggests that magma produced in the mantle must have stopped in a magma reservoir during its upward passage to the surface. This is corroborated by the low Mg number and Ni content of magmas that reach the surface. The question arises: why should one type of nephelinite consistently stop on its rise through the continental lithosphere, and the other not? Trace elements do not support the argument of mantle chemical heterogeneity as the cause. Therefore, either the physical conditions of passage of the two differ, i.e. one portion of the lithosphere is easier to penetrate than another, or the physical properties of the two types of magma must differ. No lateral variation of the lithosphere is apparent in any of the provinces and therefore the second alternative must be examined. The density of primitive nephelinite magma would be about 2.9-3.0 g cm- 3, whilst the density contrast across Moho from mantle to continental crust is about 3.4-2.8 g cm-3. Thus a nephelinite magma would be expected to stop rising in the lower continental crust, but perhaps not in the lower oceanic crust (density about 3.0 g cm-3). Why then do nodule-bearing nephelinites pass rapidly up through continental crust without stopping ? The answer may lie in the volatile content. Nephelinites associated with carbonatites by liquid immiscibility must be carbonated nephelinite magmas and hence have a high C O 2 / ( H 2 0 -kCO2) volatile fraction. If, however, the COz/ ( H 2 0 + C O 2 ) ratio is lower, the nephelinite magma would be more hydrous and less viscous
500
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FIG. 6. (a) Normalized plot of average trace-element abundances for melanephelinites from four Tertiary carbonatitic volcanic complexes in E Africa (Fig. 3); (b) normalized plot of averaged values of trace elements in olivine nephelinites from the northern and southern sectors of the Kenya rift valley.
Nephelinites and carbonatites
6i
TABLE 3. Contrasts between nephelinites of Groups I and H Character
Principal phenocrysts Clinopyroxene MgO content of rock Mg no. (100Mg/(Mg+Fe2+)) at 100 K/(Na + K) at Associated basalts Chemical fractionation Volcanic structure Percentage pyroclastic Plutonic equivalents Mantle-type xenoliths Examples
Group I (olivine-rich nephelinites) Olivine and clinopyroxene Lilac and Ti rich > 10 wt.% 74 18 Abundant Very slight Small parasitic Moderate Rare Common Kenya Rift Comores Islands Hawaiian Islands Texas Massif Central
and would perhaps be able to rise through the lithosphere without stopping. Some dynamic fractionation during ascent may also have taken place. Now that the style of nephelinitic magmatism associated with carbonatites has been distinguished from other nephelinitic developments, the term nephelinitic volcanism in the remainder of this contribution will refer to the olivine-poor nephelinites and their derivatives. During or before retention in a magma reservoir, olivine was precipitated and hence largely lost to the remaining nephelinitic magma. Resorption and reaction rims on remaining olivine phenocrysts indicate that they were no longer stable in the magma. Augite was also being precipitated in abundance. It seems that the initial precipitation of olivine increased the already high CO2 content in the magma, which eventually destabilized any remaining olivine and widened the pyroxene stability field. Cumulate bodies of dunite within carbonatite complexes are rare (e.g. Kovdor, Kola, and Shawa, Zimbabwe), whilst alkali pyroxenites occur in most carbonatitic complexes.
Pyroxene fractionation Well over half the extrusive products of nephelinite volcanism are pyroclastic, indicating a volatile-rich source. Among the products in W Kenya are numerous megacrysts of clinopyroxene up to 3 cm long. The chemistry of these and other clinopyroxenes in nephelinitic rocks is revealing
Group II (olivine-poornephelinites (melanephelinites)) Clinopyroxene and rarely olivine Pale brown or greenish and Ti poor < 10 wt.% 67 24 None Strong Large central High Frequent Very rare W Kenya E Uganda Malawi Siberia Amba Dongar, India
in the deciphering of the upward passage of the magma. Table 4 gives the analyses of a range of pyroxenes as xenocrysts, phenocrysts and groundmass phases from melanephelinites, nephelinites and phonolites. Most of them have low Ti and A1 contents which, if they came from groundmass phases, would be at variance with the higher Ti and A1 contents normally encountered in strongly alkaline rocks (Thompson 1974). Instead, as pointed out by Thompson, the low Ti and AI appear to indicate a high temperature of formation. The high Ca content of these magmas would also favour partition of the Ti into perovskite or sphene, which are both quite abundant as microphenocrysts in these rocks. Until analysed by microprobe, the presence of Cr-diopside among the xenocrysts and phenocrysts in the melanephelinites was not appreciated (Table 4, column 1). Cr-diopside is typical of pyroxenes from mantle-type lherzolites, and sometimes forms whole xenocrysts. These xenocrysts can also be present as colourless to palegreen cores to more normal pyroxene compositions (Table 4, columns 4a and 4b) and fall in the mantle pressure field for chromiferous pyroxenes described by Onuma & Tohara (1984) where Na > A1vI. The presence of Cr-diopside is thought to indicate that the nephelinitic magma does record a history of passage through upper-mantle lherzolite. Olivine, with more than 90% Fo, has not been recognized among the xenocrysts or phenocrysts in the lavas, and the few olivines present range from 86% to 81% Fo. Olivine with Fo > 90% does occur in the cumulates, however.
62 TABLE 4.
M. J. Le Bas
Analyses of clinopyroxenes from melanephelinites, nephelinites and phonolites 2a
2b
53.08 0.62 0.58 . 3.19 0.11 16.16 24.31 0.39 0.00 0.14 98.58
51.67 0.87 1.61 . 6.73 0.13 13.58 23.80 0.66 0.00 0.15 99.20
51.25(0.55) 1.06 (0.19) 1.84 (0.26) . . 4.63 (0.28) 0.05 (0.02) 14.83 (0.34) 24.08 (0.09) 0.41 (0.03) 0.02 (0.00) 0.28 (0.07) 98.45
51.94 49.95 0.71 1.95 1.53 3.09
Total
53.20(0.45) 0.45 (0.07) 1.00 (0.14) . . 2.93 (0.13) 0.06 (0.01) 16.85 (0.18) 23.43 (0.37) 0.48 (0.07) 0.00 0.86 (0.27) 99.33
48.51 1.56 4.49 3.19 4.34 5.44 4.81 0.05 0.05 0.14 15.12 14.02 13.58 23.81 24.55 22.40 0.47 0.40 0.96 0.01 0.02 0.00 0.44 0.01 . 98.42 99.47 99.64
46.55 47.80 50.45 50.41 2.19 1.37 0.80 0.82 5.04 4.82 2.30 2.27 4.01 -12.80 13.51 7.19 17.00 8.45 8.27 0.30 0.65 0.55 0.52 10.29 6.05 4.87 4.53 22.07 19.20 13.98 13.43 1.22 2.79 5.08 5.89 0.18 -0.00 0.34 . . . . 99.04 99.68 99.28 99.99
Si Ti A1 Cr Fe 3+ Fe 2+ Mn Mg Ca Na K Ca Mg Fe* Na Mg Fe z+ + M n Mg/(Mg + Fe*)
1.956 0.013 0.046 0.025 . 0.090 0.002 0.924 0.923 0.034 . 47.6 47.7 4.7 3.4 91.2 5.4 91
1.971 0.017 0.025 0.004 . 0.099 0.003 0.894 0.968 0.028 . 49.3 45.5 5.2 2.8 90.0 7.2 90
1.940 0.025 0.071 0.005 . 0.211 0.004 0.760 0.958 0.048 . 49.6 39.3 11.1 5.0 79.1 15.9 80
1.923 0.030 0.081 0.008 . . 0.145 0.002 0.830 0.968 0.030 . . 49.8 42.7 7.5 3.1 86.9 10.0 85
1.943 1.867 1.817 0.020 0.055 0.044 0.068 0.136 0.198 0.013 . . . 0.090 0.136 0.170 0.151 0.002 0.002 0.004 0.843 0.781 0.758 0.954 0.983 0.899 0.034 0.029 0.070 . 49.3 50.8 47.3 43.6 40.3 39.9 7.1 8.9 12.9 3.5 3.2 7.8 86.5 85.9 84.9 10.0 10.9 7.3 86 82 76
1.786 0.063 0.228 . 0.116 0.231 0.010 0.589 0.907 0.091 0.009 48.9 31.8 19.3 10.9 70.5 18.6 62
1
SiO2 TiO2 A1203 Fe203 FeO MnO MgO CaO Na20 K20
Cr203
.
.
3
4a
4b
5
6
7
8
9
1.871 1.949 1.942 0.040 0.023 0.024 0.222 0.105 0.103 . . -0.372 0.392 0.557 0.273 0.266 0.022 0.018 0.017 0.353 0.280 0.260 0.805 0.579 0.554 0.212 0.381 0.440 --0.017 46.4 38.0 37.2 20.3 18.4 17.5 33.3 43.6 45.3 23.6 40.0 46.8 39.4 29.4 27.7 37.0 30.6 25.5 38 30 28
1, Cr-diopside xenocryst from strongly porphyritic olivine melanephelinite lava, Kisingiri volcano, W Kenya (RR710) (mean of six wavelength-dispersive microprobe analyses (WDMAs); standard deviation to 2tr given in parentheses); 2a, core of pale-green zoned pyroxene in block of pyroxenite ejected from Oldoinyo Lengai volcano, N Tanzania (24407) (energy-dispersive microprobe analysis (EDMA)); 2b rim of 2a; 3, euhedral megacryst (2 cm) of black diopside in agglomerate between melanephelinite lavas of Kisingiri volcano, Kenya (K79) (mean of four WDMAs; standard deviation given); 4a, core of colourless zoned Cr-diopside in phenocryst in normaltype melanephelinite lava, lower N flanks of Kisingiri volcano, Kenya (U16) (WDMA); 4b rim of 4a; 5, bulk pyroxene from fine-grained melanephelinite lava, S Ruri, W Kenya (S197) (wet analysis by H. Lloyd); 6, bulk pyroxene from nephelinite lava, S. Ruri ($33) (wet analysis by H. Lloyd); 7, Aegirine-augite phenocryst in feldspar-phyric phonolite lava of Kisingiri volcano, Wasaki Peninsula, W Kenya (U410) (EDMA); 8, Aegirineaugite microphenocrysts in glassy phonolite, volcanic plug, S Ruri (S 17) (wet analysis by H. Lloyd); 9, Aegirineaugite microphenocrysts in porphyritic phonolite, volcanic plug, S Ruri ($29/1) (wet analysis by H. Lloyd). Fe*, total F e + Mn. Titaniferous lilac-coloured Al-augite phenocrysts are exceedingly rare in these nephelinites, but are a b u n d a n t in the olivine-rich nephelinites and have been interpreted (Fodor et al. 1982) as being precipitated at pressures in the vicinity of 1520 kb during passage t h r o u g h the u p p e r mantle. Since Cr-diopside occurs but Ti-Al-augite is rare in the olivine-poor nephelinitic rocks, it is inferred that the latter were not precipitated during the passage of the p a r e n t m a g m a t h r o u g h the mantle. Similar chromiferous diopsides in pyroxenite (e.g. Table 4, c o l u m n 2a) are inter-
preted as cumulates, w h i c h further suggests that a c c u m u l a t i o n in a deep reservoir did occur. T h e more n o r m a l pyroxene phenocrysts in the m e l a n e p h e l i n i t e are a b u n d a n t low-Ti low-A1 augites (Table 4, columns 2b and 3) w h i c h zone out to more aluminous augites (Table 4, columns 4b and 5). They also show m o d e r a t e Fe enrichment, indicative of fractionation in a m a g m a chamber. Large euhedral black glistening megacrysts showing signs of resorption on the crystal faces also occur. T h e y have the same composition as
Nephelinites and carbonatites the phenocrysts (Table 4, column 3) and hence presumably have the same source. Evidently they were transported directly to the surface without stopping. The rare presence of pargasitic amphibole in some E African pyroclastic rocks (Varne 1968; Le Bas 1977) may be related to fractionation at mantle pressures as proposed by Varne, but they may also be the product of strongly hydrous conditions obtaining at only slightly lower pressures in a deep crustal magma chamber and the consequent explosive eruption of this magma to the surface. If this interpretation of the magmatic plumbing is correct, then the subvolcanic coarse-grained bodies of ijolite exposed by erosion at carbonatitic centres are products of nephelinitic magma rising from the lower crustal reservoir which stopped and crystallized. Therefore, the subvolcanic bodies would not mark the position of the main magma chambers feeding the volcanic structures above. Crystal fractionation of the melanephelinitic magma in the lower crustal reservoir produced nephelinitic and finally phonolitic liquids with increasingly Fe-rich sodic pyroxenes (Table 4, columns 6-9) as plotted in Fig. 7. Compositions richer in acmite than shown in
63
Fig. 7 do occur, but only rarely, and are confined to needles growing in late-stage fractionates. Also significant, with a bearing on the plumbing, is that some pyroxenes have green cores and show reverse zoning. They are quite common in melanephelinites and nephelinites. Table 5 lists analyses of three such zoned pyroxenes which have cores richer in Fe and Na than their colourless rims. The green cores correspond to pyroxenes from phonolites, while the rims correspond to the nephelinitic and melanephelinitic hosts. These give evidence for magma mixing, whereby pyroxene phenocrysts from phonolite are mixed into incoming magma to produce the reverse-zoned pyroxenes which can then be carried to the surface. Such a process necessitates a magma reservoir as proposed above. Analysis 3 in Table 5 further shows that mixing did not always lead to eruption, but that cumulative pyroxenites could be produced. Similar magma mixing revealed by green cores with reverse zoning is well known in many parts of the world. Brooks & Prinzlau (1978) give an excellent summary covering such mixed products from Antarctica, Africa, Europe, America and Greenland.
Crustal intrusive processes
\
~/
\
ee/
...:,.'.. Fe 2+ + Mn
FIG. 7. Na-(Fe 2+ + Mn)-Mg at.% plot of pyroxenes taken from Table 4 together with unpublished energy dispersive microprobe data on further E African samples: II, Cr-diopside from RR710 (Table 4, column 1); A, megacryst from K79 (Table 4, column 3); + - - +, zoned core-to-rim pale-green pyroxene in pyroxenite (Table 4, columns 2a and 2b); O, phenocryst and groundmass pyroxenes (the tie-line with the arrow indicates zoning); - - - , pyroxene trend for the sequence pyroxenite to ijolite to nepheline syenite in the sub-volcanic complexes of E Uganda (Tyler & King 1967) and W Kenya (Le Bas 1977, and unpublished data). The path from melanephelinite to nephelinite to phonolite corresponds to fractionation beginning with lowfO2.
Recent geochemical modelling by the author using major-element, trace-element and isotopic data shows that mixing of phonolite with melanephelinite magma can also produce lavas of mugearitic and trachyandesitic compositions. Such lavas are rare, but have been recognized in the nephelinitic volcanoes of E Africa (King 1949; Le Bas 1977). In the absence of any basalts in these volcanoes, they had been thought to be products of crustal contamination of melanephelinite magma. Another feature indicative of magma residing in a continental crustal chamber for long periods is that, the more extreme is the phonolitic fractionate from a nephelinite parent, the greater is the STSr/86Sr ratio. In W Kenya the 8781"/8651" initial ratios of melanephelinites are usually between 0.7032 and 0.7038 (Norry et al. 1980), whilst phonolites and trachyphonolites have values of 0.7046 and 0.7053 respectively (Le Bas 1977, p. 140). These higher values in the fractionates are interpreted as the product of crustal contamination. The density-contrast conditions cited above for arresting the upward passage of nephelinitic magmas near the mantle-crust boundary hold true only for continental crust. They cannot be applied to oceanic islands such as the Canary and
M . J. Le Bas
64
TABLE 5. Analyses of clinopyroxenes with green cores and reverse zoning 1
Core
2
Rim
SiOz TiOz A1203 FeO T MnO MgO CaO NazO CrzO 3 Total
51.12 0.52 1.96 15.50 0.36 8.47 19.44 2.21 . 99.58
50.80 1.50 2.19 6.50 0.12 14.36 23.42 0.45 . . 99.34
Si Ti A1 Cr Fe r Mn Mg Ca Na Na Mg Fe 2+ +Mn Mg/(Mg + Fe*)
1.972 1.903 0.015 0.042 0.089 0.098 . . . 0.500 0.204 0.012 0.004 0.487 0.802 0.804 0.940 0.165 0.033 16.2 3.4 47.8 82.3 36.0 14.3 57 85
3
Core
Rim
Core
Rim
47.56 1.53 5.12 17.90 0.39 5.32 19.98 1.88 . 99.68
48.32 2.15 4.67 7.58 0.12 12.53 22.58 1.02
51.26 0.37 1.07 15.69 0.35 8.31 20.36 2.09 0.16 99.66
50.43 0.87 2.19 10.11 0.22 11.45 22.88 0.76 0.19 99.10
1.987 0.011 0.049 0.005 0.509 0.011 0.480 0.845 0.157 15.8 48.2 36.0 57
1.924 0.025 0.098 0.006 0.323 0.007 0.651 0.935 0.056 5.9 67.9 26.2 72
98.97
1.866 0.045 0.237
1.830 0.061 0.208
0.587 0.013 0.311 0.840 0.143 15.4 33.5 51.1 40
0.240 0.004 0.707 0.916 0.075 8.3 78.5 13.2 86
.
1, Fine-grained nephelinite lava with occasional phenocrysts of pyroxene, mostly with green cores and colourless rims, showing reversed zoning, Kisingiri volcano, Kenya (K61) (EDMA); 2, hybrid fine-grained lava with scattered large phenocrysts of phonolitic pyroxene rimmed by groundmass composition pyroxene of melanephelinite host, NE flanks of Kisingiri volcano (U926) (EDMA); 3, apatite pyroxenite block from Oldoinyo Lengai volcano, N Tanzania (24420) containing subhedral pyroxenes with green cores and pale-green margins (EDMA). Cape Verde Islands, where nephelinites and carbonatites occur and where the distinction between the olivine-rich and olivine-poor nephelinites is not so clear. Green pyroxene cores exist, and so do mantle nodules. If a deep crustal reservoir is lacking, then fractionation must take place during ascent or in high-level chambers such as those now occupied by the subvolcanic plutonic complexes where fractionation and segregation can be seen to have occurred. A final question may be asked concerning the stage and depth at which magma chambers are formed in continental or oceanic crust. What relation do these feeders have to the onset of liquid immiscibility which, as will be shown below, affected the carbonated nephelinite magma and permitted the formation of conjugate silicate and carbonate magmas? If the unmixing of liquids is an exothermic process, its onset together with loss of carbonate fraction from the silicate liquid would temporarily raise the liquidus temperatures of the magmas until each had precipitated its more refractory phases, mainly
olivine and/or pyroxene +oxides and a p a t i t e + oxides respectively, until the remaining liquids were again buoyant. In summary, the following magma plumbing is envisaged. A carbonated olivine-rich magma is derived by partial melting of the asthenosphere. This rises through the mantle, gathering lherzolite fragments from the walls and perhaps precipitating some olivine. On passing into the base of the crust, the magma stops temporarily in a reservoir where olivine is precipitated, C02 builds up and the pyroxene stability field expands. The limit for liquid immiscibility is reached and carbonatitic magma separates. Pyroxene crystallizes to yield pyroxenite cumulates and the magma differentiates to phonolite. Influxes of new primitive magma mix with the phonolite producing reversezoned pyroxenes and, in extreme cases, new melts of mugearitic composition. All these magmas can rise directly to the surface and erupt as lavas, but some stop at subvolcanic levels and crystallize there, forming small plutons of ijolite and nepheline syenite.
Nephelinites and carbonatites
However, kimberlites do seem to produce a residual carbonatitic fraction in some cases, such as at Benfontein, S Africa (Dawson & Hawthorne 1973). Gittins (1978) likewise describes some carbonatitic and ultramafic rocks from the Cargill complex, N Ontario, as related by simple magmatic fractionation. The third possibility is that carbonatite and nephelinite magmas can be derived as independent melts in the asthenosphere, a view upheld by Woolley & Jones (1987), but, as Treiman & Essene (1983) point out, carbonatites emanating from partial melting of mantle peridotite would be Na-poor. Derivation of a carbonate melt from the mantle has been shown by the experimental work of Eggler (1978) and Wyllie (1978, 1980). They show that at about 30 kb in a phlogopitebearing mantle with a high CO2/H20 ratio the initial melt could be carbonatitic and associated with kimberlitic or melilitic magmas. Wyllie further showed that at lower CO2/H20 ratios the upper asthenosphere can produce an alkaline ultrabasic silicate partial melt from which an immiscible alkali-rich carbonate liquid would probably separate.
Liquid immiscibility The process of the separation of two immiscible liquids is considered to be central to the interpretation of the evolution of carbonatitic magmatism. Therefore liquid immiscibility must be examined to see whether it best accords with the petrological evidence. First, however, the alternatives for the production of carbonatite magma should be briefly considered, with the assumption that carbonatites are magmatic and not hydrothermal metasomatic deposits (Bulakh & IskozDolinina 1978). There are three possible styles of origin. First, the carbonatite magma is a product of melting of limestone in the crust. This was upheld by Shand (1949) among others, but experimental and geochemical work on synthetic systems, mainly by Wyllie (well summarized in his 1978 paper), has demonstrated the igneous nature of carbonatites. The second possibility is that carbonatite is the product by fractional crystallization of nephelinitic and ijolitic magma (King 1965). The discreteness of many ijolite and carbonatite intrusions and the almost complete lack of mineralogical and chemical continuity between the silicate and carbonate magmas is not in keeping with carbonatite as a residual magmatic fraction from nephelinite magma, a point discussed by Prins (1978). Volcanic rocks midway between nephelinite and carbonatite are unknown except for a few pyroclastic examples. Plutonic rocks midway between ijolite and carbonatite are uncommon, and bimodality can be demonstrated in most complexes. In the few complexes where transitional rock types occur, the structural and textural relations of these rocks in both the field and the laboratory usually show that ijolitic rocks have been invaded by carbonatite, producing hybrids. 900
L1
65
Experimental evidence In their study of possible liquid immiscibility, Koster van Groos & Wyllie (1966, 1968, 1973) observed liquid immiscibility along the join NaA1Si3Os-Na2CO3 above 870~ and at 1 kb or more, but not at atmospheric pressure (Fig. 8). Introducing water into the system (Koster van Groos & Wyllie 1968) Caused the field of liquid immiscibility to shrink. At lower temperatures (about 700-750~ albite is joined by cancrinite on the liquidus. The accompanying H20-COz vapour contained Na20 and SiO2, and such vapours are considered to be the synthetic
[
L 2----~ LI+L2
T~ 800
NO+L2
NC+LI+L 2
_
~
NC+L~
700 Albite
I
20
40 WT %
I 6o
I 8o
Na2CO3
FIG. 8. Highly simplified phase diagram of the join Ab-NC with 10% H20 at 1 kb pressure. The fields of cancrinite and vapour have been omitted. L1 and L2 are immiscibleliquids. (After Koster van Groos & WyUie 1973.)
66
M. J. Le Bas
equivalents of N a - K fenitizing fluids known to emanate from carbonatite magmas. In 1973 Koster van Groos & Wyllie added plagioclase (Abs0 and Abs0) to the runs and observed liquid immiscibility along both joins between the plagioclase and Na2CO3. Cancrinite lay on the solidus and the carbonate liquid was richer in Ca than the silicate liquid which was dominantly albite and nepheline normative. Watkinson & Wyllie (1971) studied the comparable system NaA1SiO4-CaCO3-H20 at 1 kb but found no liquid immiscibility along the join nephelinecalcite. Had higher pressures been used, the experiments of Freestone & Hamilton (1980) indicate that liquid immiscibility would have been observed. Besides revealing the extent of liquid immiscibility in certain silicate-carbonate systems, these experiments showed (Fig. 8) that the silicate melt could have moderate quantities of dissolved carbonates, the proportion varying with water and Ca contents. In contrast, the carbonate melt usually had only about 5~ dissolved silicates. The proportion of silicate in carbonate melts and carbonate in silicate melts decreases as the temperature reduces. These distributions of one component dissolved in another correspond closely to those seen in natural ijolites and carbonatites. Calcite is sometimes observed as a late-stage trace constituent in ijolites, urtites and nepheline syenites, and fluid inclusions trapped in apatites in ijolites also reveal the presence of carbonates early in the history of crystallization of those melts (Rankin & Le Bas 1974a). Introducing K into the phase system likewise permits liquid immiscibility (Wendlandt & Harrison 1979). P and Ti also enhance the occurrence of liquid immiscibilityby increasing the polymerization of the melt (Freestone 1978; Mysen e t al. 1981). P is particularly relevant to carbonatites which are often rich in apatite. Because previous experiments did not contain a pyroxene component, Verwoerd (1978) carried out melting experiments at 2 kb on compositions along the join synthetic ijolite (5acmite +3 nepheline + 2 Na-disilicate) and an Na-Ca-natrocarbonatite end-member in order to obtain synethtic melts close to those of natural compositions. The results confirmed the liquid immiscibility shown by Koster van Groos & Wyllie. Freestone & Hamilton (1980) furthered the study of liquid immiscibility by conducting experiments over the range 0.7-7.6 kb and 9001250~ Using natural nephelinite and phonolite starting materials from Oldoinyo Lengai, together with a simplified synthetic natrocarbonatite, they showed that liquid immiscibility occurs between relatively polymerized and depolymerized liquids
and that both nephelinite and phonolite magmas are immiscible with alkali-rich carbonate melts (Fig. 9). Increasing the pressure widened the immiscibilityfield and increased the partitioning of Mg into the carbonate fraction but decreased that of K. Decreasing the temperature also widened the immiscibilitygap. Most nephelinites and phonolites of carbonatite complexes plot near the solvus for 7.6 kb at 1100~ (Fig. 9). These conditions are interpreted as those closest to the natural ones. Bedson (1983) took the experiments further, extending the pressure range to 25 kb. Amongst various distribution coefficients measured, he found that the rare earths were approximately equally distributed between the silicate and carbonate liquids, but that Ta, and more particularly Hf, were enriched in the silicate liquid. Hence immiscible silicate melts coexisting with carbonate melts should have high Hf/Ta ratios, a feature which would not be consistent with crystal fractionation of carbonate melt from silicate (Bedson 1984). Not enough data exist to test this hypothesis at present. Some early experiments by Brey & Green (1976) suggested the possibility that liquid immiscibility might occur between olivine melilitite and carbonatite, but later work (Brey 1978) showed that this was not the case. Similarly, no evidence has been found that kimberlite and carbonatite are related by liquid immiscibility. Natural evidence
At Fen and Aln6 the ijolitic and carbonatitic components do not crop out as discrete intrusions. Instead, the early ijolites and associated fenites are intruded and refenitized by carbonatites. Since the intrusion of carbonatite is frequently A
NazO + K20
IX iquids
~/ /
II
0.7/ ~.> / / /,i L . ~ "
Si02 + AI205
r~o~_
~ i
wt. %
/SOl
X
\\ \
CaO
FIG. 9. Liquid immiscibilityplot showing silicate fields I-VIII (described in the text p. 68) and the natrocarbonatite field IX around the two-liquid boundaries for 0.7, 3.0 and 7.6 kb at 1100~
67
Nephelinites and carbonatites accompanied by extensive brecciation, the ijolites are fragmented and at Fen, where the original extent of the ijolite intrusion was at least 2 km across, relatively unaltered ijolitic and melteigitic rocks are now present only in a small area (less than 1 km across) around Melteig farm in the SW corner of the complex. The rest are permeated and veined by carbonatite, giving the impression that a complete spectrum of compositions from pure silicate to pure carbonate rocks exists with the carbonate as a late-stage component. Tveitasite, hollaite, kasenite and ringite are local names given to these mixed rocks, but since they are hybrid rocks the names are of little value. In some complexes the ijolite and carbonatite intrusions do not penetrate each other but form discrete intrusions. The type locality of ijolite at Iivaara in NE Finland (Lehijarvi 1960) is one such discrete 2 km wide plug, with no signs of carbonatite at the surface. Carbonatites can also form discrete intrusions, as in the 5 km diameter pipe at Sokli on the west flank of the Kola ultraalkaline province (Vartiainen 1980). The field relations in E Africa show distinct bimodality of the intrusions (Le Bas 1977). Most intrusions combine discrete ijolitic and carbonatitic components with few or no hybrids. The clearest example is the Usaki complex (Le Bas 1977) which is an ijolite ring complex 3 km across with micro-ijolite marginal facies but with no development of carbonatite or calcite ijolite (i.e. an ijolite with more than 10% calcite). A few samples do show late-stage interstitial calcite (about 1%) as well as a few calcite veins, and this calcite is interpreted as being the residue of the few per cent of carbonate remaining dissolved in the silicate melt after the immiscible separation of the carbonate melt. The carbonatite partner of this ijolite is the Wasaki subvolcanic complex which occurs 5 km away to the north. The bimodality is particularly well marked on the rare occasions when extrusive volcanic products are preserved. At Oldoinyo Lengai the lavas are either silicate (nephelinite to phonolite) or carbonate. Parallels of this bimodal behaviour are seen in the extrusive products of all other known volcanic structures, even the newly discovered structure on the Cape Verde Islands (Silva et al. 1981). Fluid inclusions provide another line of evidence that liquid immiscibility can separate carbonate and silicate magmas. During heating and homogenization experiments on fluid inclusions contained in the apatite crystals of E African ijolites, a sequence was observed showing liquid immiscibility actually taking place in the heating stage under the microscope (Rankin & Le Bas 1974b). At 960~ a homogeneous liquid
filled the negative-crystal-shaped cavity of the fluid inclusion, but at 950~ small globules appeared and by 800~ the globules had solidified to silicate glass. At 575~ the remaining liquid solidified and instantaneously crystallized to a strongly birefringent mesh of elongated carbonate crystals. Further evidence of immiscibility comes from fluid inclusions in apatites within ijolite, which have bulk compositions close to natrocarbonatite (Le Bas & Aspden 1981). The ijolite-inclusion tie-line (Fig. 10, Ij-IC) attests that the relation is one of liquid immiscibility. It is interpreted that nephelinite magma was constrained in its crystallization path by being on the solvus and that, as it progressively crystallized, so it lost by immiscibility more of its small content of dissolved carbonate. Since the change in bulk chemical composition of the silicate melt as it migrated along the solvus was governed by loss of Ca (precipitation of diopsidic aegirine-augite and apatite), it moved towards urtite (Fig. 10, Ur). During this process the carbonate liquid became richer in Ca (Fig. 10, UC). Hence it is recognized that not only do droplets ofcarbonatite liquid separate from crystallizing nephelinite magma but also the droplets change composition with fractionation of that magma. One criterion for liquid immiscibility is that, not only must the two liquids be in equilibrium with each other, but so also must any additional phase. Several s6vitic carbonatites contain olivine, e.g. Aln6, Jacupiranga, Kaiserstuhl, Mbeya, Oka and Palaborwa, and, where measured, they lie in the range Fo95_ss. This range is similar to those in co-magmatic olivine ijolites, olivinites and olivine-bearing nephelinites. Such composiA
84
Na20 + K20
t, NCb i
~iIc
/-.// / S[O2+ AI203
'JIj
/ ,liquid
\ CoO
FIG. 10. Plot as Fig. 9 showing joins of silicate host and carbonate fluid inclusions for ijolite (Ij-IC) and urtite (Ur-UC) compared with the solvus for 3 kb at 1100~ NCbl and NCb2 are the span of natrocarbonatite analyses from Oldoinyo Lengai.
68
M. J. Le Bas
tions are in keeping with the proposition that the carbonate and silicate magmas could be related by immiscibility. Liquid immiscibility is also demonstrated by cryptic variation in the apatites in carbonatites and ijolitic rocks. Apatites precipitating from ensuing fractions of silicate magmas take different paths from those in fractionating carbonatite magmas, although all commence at a composition common to both silicate and carbonate hosts (Le Bas & Handley 1979). If the ijolites and carbonatites were related by fractional crystallization, a solid-solution phase crystallizing throughout would be expected to show a continuous variation throughout. Lapin & Vartiainen (1983), in their study of apatites from Finnish orbicular carbonatites, also reached this conclusion.
Relation between magmas and the system (Si + AI)-(Na + K)-Ca The carbonate-saturated system (Si + A1)-(Na + K)-Ca reported by Freestone & Hamilton (1980) accounts for 75%-85~ of the components of ultra-alkaline rocks and is therefore suitable for discussion here. When the compositions of common igneous rocks as well as ultra-alkaline rocks are plotted on this phase system using solvus relations at 1100~ (Fig. 9), a significant distribution is seen. This temperature is chosen because melting experiments by Piotrowski & Edgar (1970) on W Kenyan nephelinites, ijolites and nepheline syenites showed that liquidus temperatures lay in the range 1100-1200~ at 1 arm pressure, with solidus temperatures of about 950-1000~ and liquid immiscibility operates only above solidus temperatures. Figure 9 shows that, apart from some unusual rock types (fields IIa and IIb), most compositions are confined to the (Si + A1) corner with fields II and III lining up along an apparent boundary parallel but lower down in the diagram than the 7.6 kb solvus. Assuming an approximately regular interval between the solvus isobars, the apparent boundary can be taken at about 10 kb. The first significant feature is that it is the nephelinites and ijolites (field II) and phonolites and nepheline syenites (field III) which abut the solvus, and that all the other common rock compositions plot away from the solvus. This distribution supports the proposition that, if experimentally determined tie-lines are used (Freestone & Hamilton 1980), only these strongly alkaline compositions are possible candidates for liquids which could be conjugate with C a - N a - K carbonate liquids (field IX).
A second feature is that olivine nephelinites, melanephelinites and melilite nephelinites (field I) do not apparently lie on the solvus, although they could do so at higher pressures or lower temperatures, both of which enlarge the twoliquid field. The latter case of a lower temperature is petrologically unrealistic, but having mafic magmas of field I on the solvus at higher pressure is quite possible in the conditions envisaged for magma chambers deep in the crust. Even if these nephelinitic compositions were not on the solvus, they could become so after slight fractionation to field II compositions. This distinction perhaps explains why melanephelinites are associated with carbonatites but olivine nephelinites, which seldom fractionate, do not have accompanying carbonatites, it being assumed that the nephelinites were carbonated in the first place. It is further possible that if an olivine nephelinite rose high into the crust and then did fractionate, the magma would not impinge on the solvus and carbonatite could be produced by fractional crystallization. Syenites (field IV) also plot away from the proposed solvus whilst nepheline syenites plot on it, and this relation corresponds to the known absence and presence respectively of carbonatites with these rocks in the field. A third feature is that melilitites (field V) and kimberlites (field VI) lie well away from the proposed solvus, in keeping with the known geology of melilitite and kimberlite occurrences. Likewise basalts, andesites and rhyolites (fields VII and VIII) cannot produce carbonatites by liquid immiscibility. Fields IIa (alnoite) and IIb (urtite) are interesting exceptions. Alnoite is a frequent accessory magma in carbonatitic complexes and is itself commonly strongly carbonated. It is suggested that alnoite could be the fractionated product of olivine-melilite nephelinite which had penetrated high into the crust and from which the late-stage carbonate-rich residue had not separated. Urtite (field IIb) is rare even when compared with the ijolite in which it occurs, usually as minor intrusions at relatively high levels in the ijolitic plutons. It is proposed that urtite can result from liquid immiscibility taking place in an ijolite pluton at mid-crustal levels. The low Ca content of urtite compared with ijolite (Fig. 9) is consistent with loss of a C a - N a rich carbonate fraction by liquid immiscibility. This is further corroborated by the composition of the carbonate fluid inclusions trapped in urtite which are relatively Ca-rich compared with those trapped in ijolite, as shown in Fig. 10. Many of these possibilities must remain speculative until the important constituents Fe, Ti, P, F and oxygen fugacity are incorporated into the system.
69
Nephelinites and carbonatites
TABLE 6. Analyses of typical carbonatites (wt.~)
Carbonatites Despite the fact that nearly all the carbonatites of the world are calcite or dolomite-rich, the primary and parental carbonatitic magmas produced by liquid immiscibility are considered to be strongly alkaline, akin to natrocarbonatite. Twyman & Gittins (1987), however, do not consider natrocarbonatite to be either primary or produced by liquid immiscibility. The close agreement between the experimental work described above leaves no doubt that liquid immiscibility must be significant and that the primary composition is rich in Na + K. The fact that fluid inclusions trapped in early apatites in carbonatites are alkali-rich (Rankin 1975, 1977) lends further support. Most carbonatitic complexes display a wide variety of carbonatites. Field work invariably shows that the earliest intrusions are s6vites, with later alvikites and then ferrocarbonatites. N a - K fenitization is usually confined to the s6vite. Alvikites rarely show fenitization and ferrocarbonatites never show it. S6vites occasionally form diatremes 1-3 km in diameter. More usually they are tens or hundreds of metres across, while alvikites and ferrocarbonatites occur in swarms of dykes and veins.
Natrocarbonatite
1
SiO2 TiO2
AI203 Fe203 MnO MgO CaO Na20 K20 P205 CO2 F C1 SO3 SrO BaO REE
2
3
4
0.05 0.88 0 . 1 6 6 . 1 2 0.01 0.18 0 . 0 7 0 . 6 8 0.11 0.37 0 . 1 7 1.31 0.41 2.62 4.04 7.55 0.48 0.39 0.41 0 . 7 5 0.48 0.31 0.67 12.75 14.43 53.60 51.20 29.03 33.89 0.09 0 . 2 5 0.14 8.39 0.03 0.01 0.79 0.93 3.18 1.52 2 . 6 6 30.53 38.38 39.50 37.03 0.09 2.71 0.06 - . . 3.81 Trace . -0.89 2.88 - 1.35 0.23 0 . 1 0 0.01 0.08 0 . 1 7 0.11 1.26 0.05 0.3 ND c.0.1
5
6
3.24 0.83 0 . 0 0 0.07 0.20 0.65 11.50 11.00 5.18 5.53 10.74 0.36 25.85 43.60 -0.05 -0.06 1.27 0.42 32.62 30.42 --. 0.49 - 0 . 7 3 0.07 2.48 >4.0 2.82 c. 1.5
All analyses sum to 100 __+1.7 (LOI included). 1, Natrocarbonatite lava, 1960 eruption, Oldoinyo Lengai, N Tanzania (Gittins & McKie 1980); 2, s6vite dyke, Tundulu, Malawi (Garson 1965, p. 24); 3, alvikite cone sheet, Homa Mountain, W Kenya (Le Bas 1977, p. 317 (HC629) traces by C. Barber, unpublished); 4, dolomite carbonatite (beforsite dyke), Aln6, Sweden (von Eckermann 1948, p. 122); 5, ferrocarbonatite (high Mg), Kangankunde, Malawi (Garson 1965, p. 54); 6, ferrocarbonatite (low Mg), Homa Mountain, W Kenya (Le Bas 1977, p. 318 (HF15) traces by C. Barber, unpublished).
magma
The chemical composition of parental natrocarbonatite is still uncertain. Table 6, column 1, gives the composition of natrocarbonatite extruded from Oldoinyo Lengai in 1960. This is the only known instance of fresh carbonatite lava, and its composition (CC, 30 wt.~; NC, 58 wt.%; KC, 12 wt.~o) when plotted on Fig. 9 (field IX) indicates that it is the conjugate melt with phonolite magma, as interpreted by Freestone & Hamilton (1980). The conjugate carbonate melt with nephelinite would, according to them, be more calcic, but no natural examples are known except perhaps at Kaiserstuhl (Keller 1981). O and C isotopes confirm the igneous nature of natrocarbonatite and indicate a temperature of crystallization in the range 400-800~ (Suwa et aL 1975). Natrocarbonatite has been described in detail by Dawson (1962), Du Bois et al. (1963), Cooper et al. (1975), McKie & Frankis (1977) and Gittins & McKie (1980). It is porphyritic with large clear platy crystals of nyerereite ((Na,K)2Ca(CO3) 2 solid solution) and large brownish rounded crystals of gregoryite (Na2CO3 with some Ca + K substitution for Na) in a glassy carbonate matrix containing quench needles of nyerereite.
Table 7 gives a modal analysis of a sample collected by N. J. Guest from the 1960 eruption. Older examples of carbonatitic lavas or their pyroclastic equivalents have recently been discovered. Deans & Roberts (1984) described carbonatite tufts and lava clasts from the Miocene volcano at Tinderet in Kenya in which flowaligned calcified nyerereite crystals were identified. At one time they were thought to be pseudomorphs after melilite, and although relict nyerereite has never been found (nor will it be since it is water soluble) further textural study leaves little doubt that the platy cavernous pseudomorphs, largely composed of myriads of small calcite crystals, are after nyerereite. Such pseudomorphs are now also known in lavas at Kerimasi in N Tanzania, and suspected at Homa Mountain and the Ruri Hills, W Kenya, at Kruidfontein, S Africa, and in the Cape Verde Islands. The platy pseudomorphs at Kerimasi volcano described by Mariano & Roeder (1983) were identified on textural features by Hay (1983) as former nyerereite crystals. They coexist with fresh tablet-shaped calcite phenocrysts. Although both pseudomorphs and phenocrysts are com-
7o
M. J. Le Bas
TABLE 7. Modes ofnatrocarbonatites (volume ~ ) Oldoinyo Kerimasi, Lengai, Tanzania Tanzania 82-8-15A2
Kerimasi, Tanzania 82-8-14C6
Kerimasi, Tanzania 83-7-26A
Tinderet, Kenya L. KIP 1
Tinderet, Kenya U. KIP 3
Kaiserstuhl, S Germany, and Fort Portal, W Uganda
A few Abundant
9.3 2.7
Very few Abundant
Abundant None
Phenocrysts Calcite Nyerereite Gregoryite
Nil 38.1 48.0
29.7 67.0 1
7.6 46.0 __
Groundmass Calcite Nyerereite Opaques Apatite Perovskite Pyrochlore
-6.02 1.1 ----
2.8 -0.3 0.2 Trace --
46.33 Trace Trace ---
Interstitial glassa
6.8
--
--
c.50 m
7.1 79.3 0.7 0.4
m
0.5 m
1, not identified in thin section; 2, may include groundmass gregoryite; 3, Cc" N y ~ 1 1 4, includes quench nyerereite. posed of calcite, electron diffraction microprobe analysis (Table 8) distinguishes the typical igneous calcite with high Sr from that in the pseudomorphs which have low Sr but detectable Mg and Na. This distinction, the interpretation that the pseudomorphs cannot be after calcite, the relations known from the C a - N a - K carbonate system (Fig. 11) and the probable liquid immiscibility relations shown in Fig. 9 all confirm the identification. The modal variation of natrocarbonatites (Table 7) shows that these lavas can have a wide range of compositions, assuming that the phenocryst phases are complete and cognate. By using the phase data of the synthetic C a - N a - K carbonate system at I kb (Fig. 11) determined by Cooper et al. (1975), it can be seen that the natrocarbonatite of Oldoinyo Lengai plots in field A on Fig. 11, but the Kerimasi and Tinderet
lavas are more calcic, since both nyerereite and calcite are precipitated, and plot in the field B. Clasts from the Cape Verde Islands also appear to plot in field B. The occurrences of carbonatite lava at Kaiserstuhl in S Germany (Keller 1981) and at Fort Portal in W Uganda (in samples kindly provided by P. H. Nixon) show phenocrysts of calcite alone. This cannot easily be reconciled with the two-liquid field system shown in Fig. 9 and the liquidus surfaces of Fig. 11 unless the lavas were relatively N a poor. It is concluded that a wide variety of natrocarbonatite magmas can exist. Some occur with gregoryite and nyerereite phenocrysts, indicating compositions appropriate to immiscible liquids conjugate with phonolite magma, as at Oldoinyo Lengai. Some have nyerereite phenocrysts alone or with calcite (Table 7), appropriate to immiscible liquids conjugate with nephelinite magma.
TABLE 8. Microprobe analytical data on Kerismasi natrocarbonatites Brown pseudomorphs
Clear calcite crystals 1
MgO CaO S~ Na20
BD 55.1• 0.7• BD
2
BD 54.8• 0.7• BD
3
BD 51.5• 0.6• BD
4
0.8• 48.3• BD 0.5•
5
0.4• 50.6• BD BD
1 and 2, Large phenocrysts in lava (samples 82-8-14C6 and 83-7-26A respectively); 3, groundmass quench prismatic crystal (sample 83-7-26A); 4 and 5, two large pseudomorphs replaced partly by granules of calcite and partly by voids (with 1 and 2 respectively). Each analysis is the mean of four points with standard deviation to 2tr. The energy dispersive microprobe analyses were performed using the Cambridge Microscan Mk.5 at the University of Leicester. BD = below detection limit.
Nephelinites and carbonatites
/
/
NazCOs
/ K2Ca(CO3)2
Gry/(~A"XNyr/~ \\ / ~ ) ~ ~/ .) CaCOs No2Co(C03)2 Na2C%(C05)3
FIG. 11. CC-NC-KC carbonatite system showing field of natrocarbonatites from Oldoinyo Lengai (A) and from Kerimasi and Tinderet (B). Gry and Nyr are plotted at the natural compositions of gregoryite and nyerereite minerals. NY is the end-member nyerereite Na2Ca(CO3)2. FC is fairchildite K2Ca(CO3)2. (After Cooper et al. 1975).
Natrocarbonatites contain few minerals apart from carbonates. The few include oxide and phosphate minerals, but no silicates apart from rare biotite. Deans & Roberts (1984) report 15 wt.~o SiO2 and up to 5~o modal silicates, often biotite, in the Tinderet lavas. Bedson (1983) too usually found 1-5 wt.~ SiO z in his experimental carbonate charges. A120 3 was also present but was even more weakly partitioned into the carbonate melt. Biotite is the most frequent silicate mineral encountered, but only in accessory proportions. In some s6vitic carbonatites, mica and other silicates, particularly mildly sodic augite sometimes rimmed by aegirine, become abundant, and these rocks are the metacarbonatites. Vartiainen (1980) described a collar 2 km wide of metacarbonatite (average SiO2 content, 25 wt.~) around the s6vite core 2 km in diameter at Sokli. Armbrustmacher (1979) described petrographically similar rocks in Colorado, and Robbins & Tysseland (1983) have described silicocarbonatites (metacarbonatites) derived from gabbros at Pollen in N Norway. Metacarbonatites are interpreted to be 'reaction rocks' between magmatic carbonatite and country rock; the resulting 'mixed rock' is often up to 70~o calcite. Similar reactions, but on a smaller scale, are known in nearly all s6vites, and the products are often incorporated as xenoliths in s6vite intrusions. The experimentally determined distribution coefficients of elements between immiscible car-
71
bonate and silicate melts (Freestone & Hamilton 1980; Bedson 1983) help explain some of the observed variability. P partitions strongly into the carbonate melt, and hence nephelinites are often poor in apatite while the early s6vites are rich in apatite. Carbonatite lavas, however, always have less apatite than predicted by the above. Apatite is a dense mineral (specific gravity, 3.1) whereas carbonatite magma has a very low density of 2.2 g cm- 3 (Nesbitt & Kelly 1977), and therefore it is not surprising that apatite attains eruption less commonly. The partition of Mg between silicate and carbonate melts depends mainly on Pco2. At 7.6 kb Mg appears to partition preferentially into the carbonate (the data are poor), but at low pressure (1-3 kb) it partitions into the silicate. The distribution coefficient of K is never far from unity, unlike Na and Ca which both partition strongly into the carbonate melt. This causes a fall in the Na/K ratio of the silicate magma after the separation of carbonatite. Sr also partitions preferentially into the carbonate melt, particularly in the lower-temperature range 600-700~ but the Ca/Sr ratio remains unaltered in all melts (Koster van Groos 1975). Greater pressures increase the partition of REE into carbonate melt, with more light REE than heavy REE. This gives the steep REE pattern typical of carbonatites, but the steepness can vary considerably depending on the temperature at the time of separation (Bedson 1984). Intrusive carbonatites
The intrusion of carbonatite in pipe-like structures usually follows that of the silicate rocks (Fig. 12). The commonest carbonatite is s6vite showing adcumulate texture, and it is normally composed of 90~-100}/o calcite (Table 6). Dolomite may occur instead, particularly in older and more deeply eroded complexes, in keeping with the fact that dolomite is not stable at depths less than about 2 km at carbonatite magmatic temperatures (600-800~ Periclase occurs in the Amba Dongar and Kerimasi shallow-seated carbonatites. Common accessory minerals in carbonatites are fluorapatite, magnetite, biotite, pyrochlore and olivine, and some less-common minerals are aegirine-augite, arfvedsonite, K-feldspar, pyrite, zircon, baddeleyite, perovskite, sphene, quartz, fluorite, baryte, bastnaesite, parisite and monazite. The last six minerals occur mainly in the late-stage carbonatites, the ferrocarbonatites, while the remainder can usually be found in the earlier and more abundant carbonatites. There is a noticeable lack of sodic minerals
M. J. Le Bas
72 0
~,
.
4 0 0 0 t_ |
L
3 0 0 0 [I
NAPAK
,
5 [
-...--~--J~--~
+ + + +'+
metres above sea-level
/ ~ o~\~
~\c 9~J . ~'G'
, ,~o\~~ ~ \ \ ~
.~,~'~ /I --oba b~e-~ ~ '
2ooo t - _ ~ 1000~+
, l .Km
+ -r+
\
~o.~;' . . . . . . . . . . 9
Vole~
I "] -
. . . . . . .
~
UI~ A / t-IIM
~..
LOKUPOI
~eg~,,z,~" ~~ .,e o4 s u b -
-~ .
.-.n,~ ~o~
"oo @
_I
~"
J ""
~
"--.. " - ~
J
- q
+ + + + + /
Fenitized
~.
1/
basement
/ Ijolite
~ Carbonatite
~, Basement
FIG. 12. Cross-section of Napak nephelinite-carbonatite volcanic complex in E Uganda. The structure is typical of such complexes. (After King 1949.) and few potassic ones, considering the abundance of Na and K in parental natrocarbonatite. The Na and K are lost by metasomatic reaction of the early carbonatites with country rock to form fenites. Therefore the carbonatites seen cropping out are not products of the complete magma less any volatiles but have also lost their original complement of alkalis (Table 6). In the case of adcumulate s6vites the loss could have accompanied the escape of the intercumulus material. The loss of alkalis unfortunately masks the original alkali content of the parental natrocarbonatite, and therefore it is not possible to decide whether the carbonate immiscibly separated from a nephelinite or a phonolite melt. However, in some complexes (e.g. Ruri Hills, W Kenya) there is a close temporal association of phonolite and carbonatite intrusions, which could thus be pairs of conjugate immiscible liquids. The immiscible separation of carbonate melt from silicate melt takes place at l l00-1000~ This is well above the liquidus temperature for the carbonate component of the melt, but not above that for olivine and apatite. Both olivine and prismatic apatite therefore precipitate from carbonatite magma and under these conditions form cumulates at the base of the intrusions such as those occurring in Kola. Not only are olivine and apatite much denser (specific gravity of 3.3 and 3.1 respectively) than carbonatite magma (specific gravity, 2.2), but carbonatite magma has an unusually low viscosity of 5x 10-2 P (Treiman & Schedl 1983) permitting efficient gravity settling. If oxidizing conditions prevail, magnetite may also precipitate and accumulate. The carbonate liquidus is encountered at 500600~ This temperature is gauged from homogenization studies on fluid inclusions trapped in apatites (Rankin 1975, 1977). These apatites are not the prismatic crystals which accumulated from the early phase of cooling but are the large ovoid apatite crystals which characterize s6vite. This temperature is in general agreement with the liquidus temperatures determined by Wyllie and others for carbonatite magma. The low
temperatures are also marked by fenitization of the thermal aureole around the carbonatite, in which low albite, orthoclase, microcline, phlogopite and arfvedsonite commonly occur. With cooling of a carbonatite intrusion, fractionation of s6vite to alvikite to ferrocarbonatite takes place, with a final small-scale development of calcite carbonatite which lithologically looks like s6vite but is devoid of the usual accessory mineralogy characteristic of normal s6vite. This sequence is observed in E Africa (Le Bas 1977), in the U.S.S.R (Kapustin 1981), the Chilwa Province of Malawi (Garson 1965) and elsewhere. A variant on this occurs when the carbonatite is dolomitic. In this case the sequence is s6vite to dolomite carbonatite to ferrocarbonatite. These sequences accord with Wyllie's (1965) experiments. Olivine (Foss_90), rarely monticellite and very rarely melilite are restricted to the earlier and deeper calcite carbonatites (s6vites). Apatite and pyrochlore can crystallize throughout the sequence of carbonatites but are most abundant in the early s6vites, with apatite sometimes being so abundant that it forms apatite rock. Such apatite deposits have been mined. Apatite rock also occurs as disrupted schlieren in s6vite and dolomite carbonatite where they are dragged and folded by the magmatic flow. These apatites are almost completely deficient in Sr and REE and can be matched with microfragments of similar textured apatite rock caught up in the nephelinitic lavas. The main apatite in the s6vites occurs as large slightly resorbed crystals which have much the same contents of Sr and La as the apatite phenocrysts in nephelinitic lavas, as is appropriate to the postulated liquid-immiscible relation. Lapin & Vartiainen (1983) similarly describe olivine, apatite and magnetite which have fractionated from carbonatite. Biotite can occur in all intrusive carbonatites, as can pyroxenes and amphiboles, but they are more frequent in the early s6vites and dolomitic carbonatiteso Biotite flakes, often with margins oxidized to opaque oxides, together with magne-
Nephelinites and carbonatites tite phenocrysts can also occur in carbonatite lavas such as those near Fort Portal in SW Uganda. Feldspar, usually potassic but sometimes sodic, occurs in the margins of s6vitic and dolomitic carbonatites and is derived from the enveloping feldspathic fenite. Many but not all s6vites carry 1%-2~ of euhedral magnetite microphenocrysts which often coexist with biotite, apatite and pyrochlore. The magnetite is Ti poor with usually only 1-2 wt.% TiO2, and at Jacupiranga, Brazil, the magnetites in the s6vites are zoned with decreasing contents of Ti towards the rims. MnO and MgO, both of which usually show contents of less than 1 wt.~ and less than 4 wt.~ respectively, also decrease with the zoning. Magnetites in the later intrusive phases of carbonatite have compositions comparable with these outer zones and are also pure magnetite (Prins 1972; Gaspar & Wyllie 1983a). Ilmenite is rare in carbonatites and has very variable Mg, Mn and Nb contents. At Jacupiranga, Brazil, where it coexists with magnetite in the reaction zone betweenjacupirangite (titaniferous pyroxenite) and later s6vite, Gaspar & Wyllie (1983b) have calculated an equilibration temperature of 570-595~ with oxygen fugacities of 10- lS'5-10-19.5 atm, which are similar to the values estimated by Prins (1972) for magnetites from a variety of African carbonatites (540-575~ 10- 23-10- 24 atm). Alvikites are medium grained compared with the coarse-grained s6vites and are less voluminous, although in some complexes they appear to be very abundant because they form numerous but thin minor dykes and cone-sheets which comprise the majority of the outcrop. The mineralogy of alvikites is similar to that of s6vite, with a tendency for alvikite to have fewer accessory oxide and silicate minerals. The principal difference lies in the calcite which, in alvikite, was slightly ferroan when it crystallized. Probe analysis of the calcite often fails to reveal this ferroan character, evidently because the iron has unmixed from the calcite structure as a result of late-stage processes to which all carbonatites are prone. In such cases the iron is visible only as minute black specks scattered throughout the carbonatite quite separate from any magnetite microphenocrysts, and as brown staining along the crystal boundaries of the calcites. Alvikites are also marked by slightly greater contents of Mn, Ba and REE (Table 6). These elements appear to be in solid solution in the carbonate minerals. The apatites in alvikites are also rich in REE as well as Sr. True alvikites usually have a pale brownish colour, reflecting the Mn and Fe contents. Some, however, are white, like s6vite, and these commonly do not
73
contain as much Mn, Fe, Ba and REE as true alvikites and are in fact micros~vites. The distinction in the field is not easy as the brown coloration is often obscured by secondary processes, but trace-element analysis of the rock or probe analysis of the apatite can reveal the distinction. Much of the economic wealth of carbonatites lies in the ferrocarbonatites. The primary carbonate mineral of ferrocarbonatite is ankeritic, but with weathering much of it breaks down to secondary calcite or dolomite rimmed by iron oxides. Ferrocarbonatite occurs in only minor quantities in most complexes, and appears as thin brown dykes and veins with chilled margins and cross-cutting relations against earlier carbonatites. There are many good examples in E Africa, Malawi, Aln6 and the Cape Verde Islands. In other cases ferrocarbonatites form diffuse veins and patches apparently replacing s6vites, and are considered to be a late-stage development of the magmatic ferrocarbonatite. This residue of carbonatite magma appears to have become so rich in volatiles that it no longer behaved as a melt but as a volatile-rich fluid. Brown ferrocarbonatitic material can be seen progressively penetrating coarse calcite and dolomite carbonatites, migrating along crystal boundaries and gradually replacing the calcite and dolomite. The replacement process can go to completion, and if it begins along a fracture, as it usually does, the end-product can be a parallel-sided zone a metre or so wide which can look rather like a crosscutting dyke. This process can be seen replacing s/Svite in Kenya, Malawi and Fen. At Newania, India, and Sarfartoq, W Greenland, it replaces dolomite carbonatite. The apatites in the ferrocarbonatites also provide evidence for the replacement process. The cryptic variation shown by apatite is usually recorded by increases in REE, which reach a maximum in the ferrocarbonatites, but often the apatite in ferrocarbonatite lies compositionally in the alvikite and sometimes s6vite fields. This is consistent with the interpretation that many ferrocarbonatites were formerly alvikites subsequently transformed and mineralized. Mineralization
Four stages of mineralization can be recognized, all following on the magmatic Fe-enrichment (and Mn-enrichment) stage of ferrocarbonatite development. The first stage of mineralization is the production of REE-rich carbonatites from residual Frich magmas (Moiler et al. 1980; Samoylov & Smirnova 1980; Viladkar & Dulski 1986). The
74
M. J. Le Bas
REE minerals segregate in pockets and fractures along sheet-like ferrocarbonatite bodies. The minerals, principally bastnaesite ((Ce,La)FCO3), appear as pale-brown streaks within the darkerbrown to black ferrocarbonatite, and show strong light REE enrichment with La/Yb ratios of 1001000. U - T h radioactivity and associated fluorite or baryte are usually indications of the presence of REE minerals. Nearly all carbonatite complexes show this mineralization to some extent. The largest is at Mountain Pass, California, where Sr-rich carbonatites have more than 10% bastnaesite and other REE fluorcarbonate minerals cut by barytebearing and silicified radioactive veins (Olsen et al. 1954). Experiments by Jones & Wyllie (1983) suggest that the bastnaesite could have been precipitated from a low-temperature carbonatite magma (perhaps about 500~ or less). The Kangankunde carbonatite in Malawi has a similar REE mineralization but with Sr > Ba. Following the formation of the replacement ferrocarbonatites is the second-stage mineralization associated with late-stage fluids and the precipitation of fluorite at lower temperatures (100-200~ (Roedder 1973). Massive fluorite usually forms along fractures, particularly in the roof zones of carbonatitic complexes. It is variously purple, yellow or colourless, and can also occur as scattered crystals in the early carbonatites adjacent to the fractures along which the fluids passed. Evidently, F-rich fluids penetrated into the wall-rocks. At Kruidfontein near Johannesburg and at Amba Dongar, India, the F-rich fluids rose to the top of the carbonatite complex and entered and mineralized the fractured rocks capping the complex. At Kruidfontein the fluids penetrated up into bedded carbonatitic tufts sitting in a volcanic caldera structure and formed fluorite-rich horizons interbedded with the tufts. The third stage is the formation of baryte. It is present in nearly all carbonatite complexes and follows closely on the formation of fluorite with which it commonly occurs but in much smaller quantities. Local concentrations do occur and have been mined (e.g. at Aln6, Sweden). The Ba, F and REE are all mantle derived. They are not products of assimilation of continental crust, as it is sometimes considered for the Curich mineralization at the Palaborwa carbonatite in S Africa, because exactly similar Ba, F and REE mineralization is known in the Cape Verde Islands which are seated directly on oceanic crust (Le Bas 1984). The fourth stage of mineralization is marked by the formation of U - T h minerals together with pervasive silicification. This mineralization is
never very extensive and is characterized by low U/Th ratios which makes some deposits uneconomic (e.g. Loe-Shilman, N W Pakistan (Jan et al. 1981)). This is often followed by late-stage calcite veining and reprecipitation of Fe oxides, usually hematite. The Fe is derived from ferrocarbonatites penetrated by hydrothermal fluids. If the fluids become locally all-pervasive, assisted by incursions of groundwaters, then large hematite deposits can form. The Rodberg iron deposit at Fen, S Norway, is an example of this process (Andersen 1984). These Fe deposits are distinct from the magnetite deposits such as that at Kovdor, Karelia, U.S.S.R., which is a product of primary magmatic precipitation.
Fenitization The products of fenitization are mineralogically syenitic, with the result that they are sometimes misidentified as igneous syenites. The chief points of distinction lie in the texture and structure. Syenitic fenites are often heterogeneous and show low-temperature deformation or shattering. With increasing metamorphic temperatures, schlieren and mechanical flow textures develop, and next to the igneous contact the fenites become coarse grained. The effects of fenitization are more extreme than those of ordinary contact metamorphism because the metasomatism causes strong chemical reactions which can totally obscure the original composition of the rock. As a result there are many theories of fenitic processes (King & Sutherland 1960; Woolley 1982). The key lies in the realization that fenitization is not a single process but is often multiple. The variables involved are (a) the geochemistry of the magma causing the metasomatism, (b) the nature of the fluids leaving the magma, (c) the depth profile of the magma (different reactions develop at different depths), (d) the porosity and structure of the country rocks being fenitized, (e) the mineralogy and geochemistry of the rocks being fenitized and (0 any water-rock reaction.
Syenitic fenites generated by ijolite Fenites are not seen around nephelinitic volcanic pipes and dykes, nor around xenoliths of country rock trapped in nephelinitic magma. Evidently time is insufficient for identifiable metasomatism to take place. However, fenites develop strongly around larger intrusive bodies of nephelinite magma that fractionate and crystallize to ijolite. Fenitic zones can often be identified around ijolite plutons, and Fig. 13 gives the typical
Nephelinites and carbonatites ORIGINAL MINE RA~.,~400
300
METRES 200
I
!
75 0
100 |
CONTACT
GRANODIORITE ZONE IT
ZONE I
=
ZONE 111"
ZONE
I
Quartz ~
Ae~i,rine- ou~lite Mognetite ]lmenite Oligoclose
'
Atbite (An3-s) i
Microcline
Orthoctose
_.
Biotite ,,, ~ Alk(:Ui FeldSpQr
Mesoperthlte & . . . . . o~tipe'rthite iNephetine . . . . .
.........
,
!
~
Pate brown biotite ,,I
; .
.
.
Phlogopite . . . .
.
W
I
I
i_j
tO Hornblende
~, s.0dic a..~hibote (,~Y,)
Sodic ;Qmphibote ' WoliQstonite ~"
~ SHATTERIN~L. w-
STRONG MORTAR
FENIT/ZEO
GRANULATION TEXTURES
GRANOOtoRITE
COMPLETE
& "-
GRANULATION
$YENtTh
FIG. 13. The mineralogical changes across the fenite zones surrounding ijolite in granodiorite. (After Rubie, in Le Bas 1977, p. 72.)
mineralogical changes that take place across the zones. The pyroxene grades from a weakly sodic aegirine-augite near the contact to almost pure aegirine in zone 1, and can form 5%-50~ of the rock. The feldspar near the contact is new hypersolvus alkali feldspar commonly in the range Orgo_6oAb60_40, but in the outer zones the original sub-solvus feldspars of the granitic basement survive. The hypersolvus character of the feldspars and sub-solidus texture of the fenites suggest fenitization temperatures of about 800~ Ijolite intrusion causes uplift of the country rocks. At one time it was thought that the uplift was related to the metasomatic transfer of material from magma to fenite, but Rubie (1982) has shown that in W Kenya the transfer of material, mainly SiOz, was from country rock to magma, and that granodiorite suffered a volume loss of up to 20~ near the contact. This results in the formation of feldspathized ijolite (petrographically equivalent to nepheline syenite) at the margin of ijolite intrusions (Fig. 14). Aluminium appears to have been almost immobile during fenitization, and the K / N a ratio
decreases slightly away from the contact with falling temperature (Rubie & Gunter 1983). Other studies have been based on the assumption that oxygen is immobile, a condition which approximates to a constant-volume fenitization (McKie 1966), but Appleyard & Woolley (1979) based their study of fenitization on a constant-specificgravity model which suggested, at the Borrolan syenite in N W Scotland and the Sokli carbonatite in N Finland, that modest volume increases had taken place. Carbonate 9minerals are rare in fenites around ijolites unless later carbonatite intrusions are present. Evidently the metasomatic solutions causing the fenitization were not carbonate bearing.
Syenitic fenites generated by carbonatite N a - K fenites around carbonatites are much more varied than those around ijolites. Only s6vites, dolomitic carbonatites and rarely micros6vites show fenitization. Alvikites and ferrocarbonatites almost never show fenitization, and this is
76
M. J. Le Bas +
+ +
+
BASEMENT IttLIIII IIIIIIIII1,, Nepheline-poor Nep~;~i;;-po, o:
_jr_
.....
+
--- =
,,
11111IT,iF-_--
x
lllrlll
-
--
•
IIII rr ;i-~-
Nepheline-ll
,,
x
•
'I[II[EIIJV-s'
+
" "
x "
inward d ffus on of Si during fenitization
Ii'/
~X_--~lllrelative
-.~
from less ldesilicified
~- '~1, basement
X
~/N,,~
/,, I
;^~J I
]l~lll(
~
~o.-o,,,o,,e
~'1
x
+
~
+
screen of chilled m c r o - j o te prevents diffusion of Si
X
-
-Jr-
Ix|/i!/
x
aegirine-rich
to
--~-_(!1 'jmore d e s i l i c i f i e d
rich fenite /
+ Jlil~
downward
-', =lllt~'"""~176o, ~,
lJ
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from b a s e m e n t
FIG. 14. Schematic cross-section of ijolite intrusion with chilled margin and influx of silica during fenitization, and the development of feldspathized ijolite (mineralogicallyequivalent to nepheline syenite but texturally distinct) particularly along the upper parts of the margins. taken to indicate that early carbonatite magmas are alkali-rich but that these alkalis are lost during the early stages of carbonatite fractiona1 ~
0 + K20
/ -.6"D~ /
i' ;,,!gO + Fe20T
~k
~
A A
A
9
x X CaO
wt% FIG. 15. Carbonatite plot showing the path of
fractionation of carbonatite magma from natrocarbonatite (#) to s0vite (O) and alvikites ( x ) to beforsites and ferrocarbonatites (A): P, Fort Portal lava; T, estimated Tinderet lava; K, estimated Kerimasi lava (Table 7). (After Le Bas 1981.)
tion (Fig. 15). In some cases even the final phases of s0vite intrusion fail to cause fenitization. Evidently by that stage the s6vite magma has lost its alkalis, and instead could only brecciate the fenites formed earlier. This is contrary to the proposal that the concentration of alkalis increases during carbonatite fractionation (Twyman & Gittins 1987). Depth profiles are important. Fenitization around the topmost parts of a s6vite is strongly potassic, usually with the production of almost pure K-feldspar rock. Early records of this came from Malawi (Dixey et al. 1935), Songwe, Tanzania (Brown 1964), Kaiserstuhl, S Germany (Sutherland 1967) and Zambia (Bailey 1966). At Koga in Swat, N Pakistan, K-feldspar rock (plus rare aegirine) about 50 m thick completely caps a s0vite plug 500m across (Mian, personal communication). The K-feldspar is usually orthoclase Or87_9s, whereupon the rocks are termed orthoclasites (Sutherland 1965). Sometimes the K-rich feldspar is microcline, and Heinrich & Moore (1970) named such cases 'burnt rock' as they had that appearance. Commonly, the only other phase present is 1%-2~ of limonitic iron oxides, with calcite absent. The high K content is character-
Nephelinites and carbonatites istic of these fenites, whereas the K-feldspar developed in feldspathized ijolites and ijolitic fenites is less potassic (Or<87) (Le Bas 1984). When s6vites intrude mafic rocks, the fenites formed are phlogopite rich and not feldspar rich (Gittins et al. 1975), and it has been observed that, when s6vite cuts ijolite containing pyroxenite, the pyroxenite is phlogopitized and the ijolite is feldspathized, indicating that the Mg for the formation of phlogopite comes from the rock being fenitized and not from the carbonatite. Oxygen-isotope analyses on mineral pairs to determine temperature have yet to be made successfully, but X-ray diffraction of the K-rich feldspar indicates a low temperature of formation with obliquities up to 0.6 (Le Bas 1984). Haggerty & Mariano (1983) calculated a temperature of 500-550~ for the fenites around the Parana carbonatite, which agrees with that generally estimated from the mineralogy of carbonatitic fenites. In older and more deeply eroded complexes, the feldspathic fenites developed around so~ites are albite rich, sometimes coexisting with Mgarfvedsonite, aegirine and iron oxides. Such albitites occur around carbonatites in the fenitized quartzites at Turyi and in migmatites at Vuoriyarvi, both on the Kola Peninsula, and in gneisses at Novopoltavskoye between the Azov and Caspian Seas, U.S.S.R (Kapustin 1981). Albitite and aegirine-albite fenites are also formed from grits, granites and migmatites around the Sangu carbonatite in W Tanzania (Coetzee 1963). These albitite and albite-aegirine schists are lithologically similar to the syenitic fenites around the silicate complexes, but they can be distinguished mineralogically. The feldspar of fenites around carbonatites is almost pure albite (Ab>90) or almost pure orthoclase (Or>85), whereas the feldspar in the syenitic fenites around ijolite and syenite is an alkali feldspar Ab60_40 Or4o_60. The compositions reflect the low temperature of carbonatite fenitization (below 550~ in contrast with the higher temperatures of the silicate fenitization (more than 700~ Albitite also occurs as fragments brought up by carbonatitic volcanic breccias, e.g. Amba Dongar, India (Deans et al. 1972), and NyamajiWasaki, W Kenya (Le Bas 1977), indicating albite-rich fenites at depth whilst K-rich fenites are exposed at the surface. A similar zonation occurs around the Novopoltavskoye carbonatite, Azov, where inner zone fenites are albitic but the outer ones are microcline-rich (Kapustin 1982). The N a - K zonation is in accord with Orville's (1963) experiments on alkali-ion exchange between fluid and feldspar at falling low tempera-
77
tures, and finds support from Woolley's worldwide investigations (1982). S6vite and alvikite carry very little F, but the natrocarbonatitic magmas from which they were derived were probably F rich and C1 rich. Natrocarbonatite lava has 1.49-2.93 wt.% F and 0.77-4.15 wt.% C1 (Gittins & McKie 1980). Some minerals in carbonatites also show that they crystallized from halogen-rich magma, and Dawson & Fuge (1980) estimate that F > C1 was usual. In addition to apatite, biotite and Na-amphibole, which are the most important F-bearing minerals, clinohumite (Mg(F,OH)2.4MgzSiO4) has been reported at Cargill in Ontario, Jacupiranga in Brazil, Palaborwa in S Africa, Sokli in Finland and Gardiner in E Greenland (Gittins 1978). Even the residual nephelinitic glasses from Oldoinyo Lengai volcano have appreciable halogen contents (0.25-0.69 wt.% C1), and the F/C1 ratio for the whole rock is 1.6 (Donaldson & Dawson 1978). The F and C1 are evidently mantle derived, probably from phlogopite rather than apatite (Edgar & Arima 1985). However, at Oka, Quebec, Treiman & Essene (1984) calculate that the fluorine fugacity was low (10-44). If carbonatite magma were rich in halogens, particularly F, then all fenitic fluids emanating from it would also be F rich as would the residual fluids. Evidence for the former is seen in the fluorapatites and fluoramphiboles developed in fenites, and for the latter in the fluorite late-stage mineralization. Yardley (1985) also finds evidence from apatites that F might be important in metamorphic fluids. It is considered probable that a fluoride species is the main solvent of the fenitizing fluids.
Conclusions In this assessment of the main features of carbonatite-nephelinite volcanism, the following salient points can be made. The starting point is carbonated nephelinitic magma generated from an enriched mantle source such as that described by Cohen et al. (1984). Such magma rises through the lithosphere, zone refining as it rises, until it stops in the lower crust. Fractionation at that stage would permit olivine and pyroxene cumulates to form. At the same time the melt would fractionate to phonolite, and liquid immiscibility could take place forming a series of conjugate carbonatite magmas (Fig. 16). Nephelinitic magma is dense (about 2.7 g cm-3) and hence is not seen where the crust is composed of thick low-density unaltered sediments. Where nephelinitic magma has the buoyancy to penetrate the crust it can make
78
M. J. Le Bas
MELANEPHELINITE NEPHELINITEPHONOLITE LAVAS&PYROCLASTICS
PSEUDO-
DYKES
Rap,d
t /
IJOLITE
,
Ropid ascent
" "
I --.~
K.FEN!TE~A~ vK~TEE/
FenilizatiOn
IPER,DOTITE
/ /
FERROC,~BONATIT~
/I.SYE
I Slow / aster, I
I
LATE- STAGE CARB~NATITE
/
I:k/1 t oo.o
oscenl
o, .....
"~ "
TRACHYTE
URTITE & WOLL-URTITE LATE- STAGE ~11 ALKALINE FLUIDS/' CALCITE 1~ / IJOLITE -
SYE.,TES \
CARBONATED BRECCIASTUFFS & AGGLOMERATES
CARBONAIIELAVA & PYROCLASTC IS
.a
PQchonohon Low volahle ~ .
silicote melt
Volobie-rJch fluid
I lower crustal magma chamber
HELINITE I LIQUID FRACTIONATES I TO PHONOLITE I IMMISCIBILITY I
CUMULATES
? BASE OF CRUST
CARBONATED NEPHELINITIC MAGMA
PARTIAL MELT OF ENRICHED MANTLE SOURCE
FIG. 16. Schematic flow chart for the evolution of nephelinitic and carbonatitic magmas. (After Le Bas 1977, Fig. 24.2.) pyroxenite-ijolite-urtite plutons, with the ijolite perhaps developing an outer zone and capping of feldspathic ijolite (Fig. 14) as a result of the inward diffusion of silica from the country rocks undergoing fenitization. The country rocks are metasomatized to syenitic fenites with alkali feldspar and Na-pyriboles (Fig. 17). Although natrocarbonatite magmas of various (Na + K)/Ca ratios can be produced by the liquid
immiscibility, depending on whether the conjugate silicate melt is nephelinite or phonolite, the alkali loss to fenitization obscures the differences. All carbonatites follow similar paths producing s6vites, dolomite carbonatites, alvikites and ferrocarbonatites with REE enrichment, with the final production of residual mineralizing fluids which precipitate fluorite, baryte and U - T h minerals. F-rich alkali-bearing fluids cause feni-
Nephelinites and carbonatites
79
J
Cone-sheets~ plones of brecciation end K |enitizetion
'.?3
:"
+ +
/ ~_
.
I/0tite~, ~ + 1 ~ - I * +N'. ;, § i..I/."~ 4-
O,L
__.~.___,
- IN Cl i . r'." .I
FIG. 17. Schematic cross-section across an idealized nephelinite-carbonatite volcanic complex showing the normal structure and sequence of intrusion, the zonal fenitization around the ijolite, the deeper-level albitization (Na), the nearer-surface potassic feldspathization (K) accompanied by brecciation (A) and the emplacement of s6vite (CI), the cone-sheet fracture system of the alvikites (C2) and ferrocarbonatites (C3), and the final development of late-stage carbonate veins (C4) and a mineralized capping above the intrusive centre. (After Le Bas 1977, Fig. 23.1.) tization. Fluids e m a n a t i n g from early carbonatites cause albitization with albite-rich fenites +_ Na-pyriboles being developed at deeper levels around the carbonatite pluton. K-rich feldspathic fenites form towards the top and cap the s6vitic intrusions. They are often a c c o m p a n i e d by
brecciation and are usually followed by an F, Ba and U - T h mineralization. ACKNOWLEDGMENTS" I acknowledge most sincerely the many colleagues who have helped me reach these conclusions.
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BAILEY, D. K. (1966) Carbonatite volcanoes and shallow intrusions in Zambia. In: TUTTLE,O. F. & GITTINS, J. (eds) Carbonatites, pp. 127-154. Wiley, New York. --1987. Mantle metasomatism--perspective and prospect. In: FtTTON, J. G. & UPTON, B. G. J. (eds) Alkaline Igneous Rocks, Geol. Soc. Spec. Publ. 30, pp. 1-13. BAKER, B. Iq. 1987. Outline of the petrology of the Kenya rift alkaline province. In: FITTON, J. G. & UPTON, B. G. J. (eds) Alkaline Igneous Rocks, Geol. Soc. Spec. Publ. 30, pp. 293-311.
8o
M. J. Le Bas
BEDSON, P. 1983. The origin of the carbonatites and their relation to other rocks by liquid immiscibility. Ph.D Thesis, University of Manchester (unpublished). - 1984. Rare earth element distributions between immiscible silicate and carbonate liquids. Prog. Exp. Petrol. NERC 6th Rep., pp. 12-19. BOWEN, N. L. 1924. The Fen Area in Telemark, Norway. Am. J. Sci. 8, 1-11. BREY, G. 1978. Origin of olivine melilitites--chemical and experimental constraints. J. Volcanol. geotherm. Res. 3, 61-88. - - & GREEN, D. H. 1976. Solubility of CO2 in olivine melilitite at high pressures and role of COz in the earth's upper mantle. Contrib. Mineral. Petrol. 5 5 , 217-30. BRISTOW, J. W. & SAGGERSON, E. P. 1983. A general account of Karoo vulcanicity in southern Africa. Geol. Rundsch. 72, 1015-60. BROGGER, W. G. 1921. Die Eruptivgesteine des Kristianiagebietes. IV. Das Fengebeit in Telemark, Norwegen. Vidensk. Skr. 1. Mat-Naturv. Klasse, 1920 (9). BROOKS, C. K. & PRINZLAU,I. 1978. Magma mixing in mafic alkaline volcanic rocks: the evidence from relict phenocryst phases and other inclusions. J. Volcanol. geotherm. Res. 4, 315-31. BROWN, P. E. 1964. The Songwe scarp carbonatite and associated feldspathization in the Mbeya Range, Tanganyika. Q.J. geol. Soc. 120, 223-40. BULAKH, A . G . & ISKOZ-DOLININA, I. P. 1978. Origin of carbonatites in the Central'nyy massif, Turig peninsula. Int. geol. Rev. 20, 822-30. CLAGUE, D. A. & FREY, F. A. 1982. Petrology and trace element geochemistry of the Honolulu Volcanics, Oahu: implications for the oceanic mantle below Hawaii. J. Petrol. 23, 447-504. COETZEE, G. L. 1963. Carbonatites of the Karema depression, western Tanganyika. Trans. geol. Soc. S. Afr. 66, 283-333. COHEN, R. S., O'NIONS, R. K. & DAWSON, J. B. 1984. Isotope geochemistry of xenoliths from East Africa: implications for development of mantle reservoirs and their interaction. Earth planet. Sci. Lett. 68, 209-20. COOPER, A, F., GxYrINS, J. & TUYrLE, O. F. 1975. The system Na2CO2-K2CO3-CaCO 3 at 1 kilobar and its significance in carbonatite petrogenesis. Am. J. Sci. 275, 534-60. DAWSON, J. B. 1962. The geology of Oldoinyo Lengai. Bull. Volcanol. 24, 349-87. --1980. Kimberlites and their Xenoliths, Springer, Berlin. --& FUGE, R. 1980. Halogen content of some African primary carbonatites. Lithos, 13, 139-43. -& HAWTHORNE,J. B. 1973. Magmatic sedimentation and carbonatitic differentiation in kimberlite sills at Benfontein, South Africa. J. geol. Soc. Lond. 12, 61-85. DEANS, T. & ROBERTS, B. 1984. Carbonatite tufts and lava clasts of the Tinderet foothills, Western Kenya: a study of calcified natrocarbonatites. J. geol. Soe. Lond. 141,563-80. - - , SUKHESWALA,R. N., SETHNA,S. F. & VILADKAR,
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82
M. J. Le Bas
igneous Province, South West Africa. Annen. Univ. Stellenbosch, Ser. A1 (Geol.) 3, 145-278. RANKIN, A. H. 1975. Fluid inclusion studies in apatite from carbonatites of the Wasaki area of western Kenya. Lithos, 8, 123-36. -1977. Fluid-inclusion evidence for the formation conditions of apatite from the Tororo carbonatite complex of eastern Uganda. Mineral. Mag. 41, 155-64.
LE BAS, M. J. 1974a. Nahcolite (NaHCO3) in inclusions in apatites from some E. African ijolites and carbonatites. Mineral. Mag. 39, 563-70. - & -1974b. Liquid immiscibility between silicate and carbonate melts in naturally-occurring ijolite magma. Nature, Lond. 2 5 0 , 206-9. ROBBINS, B. &. TYSSELAND, M. 1983. The geology, geochemistry and origin of ultrabasic fenites associated with the Pollen carbonatite (Finnmark, Norway). Chem. Geol. 40, 65-95. ROEDDER, E. 1973. Fluid inclusions from the fluorite deposits associated with carbonatite at Amba Dongar, India and Okorusu, South West Africa. Trans. Inst. Min. Metall., Sect. B, 82, 35-9. RUBIE, D. C. 1982. Mass transfer and volume change during alkali metasomatism at Kisingiri, western Kenya. Lithos, 15, 99-109. - & GUNTER, W. D. 1983. The role of speciation in alkaline igneous fluids during fenite metasomatism. Contrib. Mineral. Petrol. 82, 165-75. SAErHER, E. 1957. The alkaline rock province of the Fen area in southern Norway. Kgl. norske Videns. Selsk. Skr. 1957 (1). SAGGERSON, E. P. 1970. The structural control and genesis of alkaline rocks in central Kenya. Bull. Volcanol. 34, 38-76. SAMOVLOV,V. S. & SMIRNOVA,Ye. V. 1980. Rare earth behaviour in carbonatite formation and the origin of carbonatites. Geochem. Int. 17, 140-52. SnAND, S. J. 1949. Eruptive Rocks. Thomas Murby, London. SILVA, L. C., LE BAS, M. J. & ROBERTSON, A. H. F. 1981. An oceanic carbonatite volcano on Santiago. Cape Verde Islands. Nature, Lond. 294, 664-5. SPENCER, A. B. 1969. Alkalic igneous rocks of the Balcones Province, Texas. J. Petrol. 10, 272-306. STRECKEISEN,A. 1976. To each plutonic rock its proper name. Earth Sci. Rev. 12, 1-33. -1978. IUGS Subcommission on the systematics of igneous rocks; classification and nomenclature of volcanic rocks, lamprophyres, carbonatites and melilitic rocks. Neues Jb. Mineral. Abh. 134 (i), 114. STRONG, D. F. 1972. Petrology of the island of Moheli, western Indian Ocean. Geol. Soc. Am. Bull. 83, 389-406. SUTHERLAND,D. S. 1965. Nomenclature of the potassicfeldspathic rocks associated with carbonatites. Geol. Soc. Am. Bull. 76, 1409-12. -1967. A note on the occurrence of potassium-rich trachytes in the Kaiserstuhl carbonatite complex, West Germany. Mineral. Mag. 36, 334-41. SUWA, K., OANA, S., WADA, H. & OSAKI, S. 1975. - - &
Isotope geochemistry and petrology of African carbonatites. Phys. Chem. Earth, 9, 735-45. THOMPSON,R. N. 1974. Some high-pressure pyroxenes. Mineral. Mag. 39, 768-87. TREIMAN, A. H. & ESSENE, E. J. 1983. Mantle eclogite and carbonate as sources of sodic carbonatites and alkalic magmas. Nature, Lond. 302, 700-3. --&--1984. A periclase-dolomite-calcite carbonatite from the Oka complex, Quebec, and its calculated volatile composition. Contrib. Mineral. Petrol. 85, 149-57. - - & SCHEOL, A. 1983. Properties of carbonatite magma and processes in carbonatite magma chambers. J. geol. 91,437-47. TUTTLE, O. F. & GITTINS, J. 1966. Carbonatites. Wiley, New York. TWYMAN, J. D. & GITTINS,J. 1987. Alkalic carbonatite magmas: parental or derivative. In: FITI'ON, J. G. & UPTON, B. G. J. (eds) Alkaline Igneous Rocks, Geol. Soc. Spec. Publ. 30, pp. 85-101. TYLER, R. C. & KING, B. C. 1967. The pyroxenes of the alkaline igneous complexes of eastern Uganda. Mineral. Mag. 36, 5-21. VARNE, R. 1968. The petrology of Moroto Mountain, eastern Uganda, and the origin of nephelinites. J. Petrol. 9, 169-90. VARTIAINEN, H. 1980. The petrology, mineralogy and petrochemistry of the Sokli carbonatite massif, northern Finland. Geol. Surv. Finland, Bull. 313. VERWOERD, W. J. 1978. Liquid immiscibility and the carbonatite-ijolite relationship: preliminary data on the join NaFea+Si206-CaCO3 and related compositions. Ann. Rep. geophys. Lab. Wash. 77, 767-74. VILADKAR,S. G. & DULSKI,P. 1986. Rare earth element abundances in carbonatites, alkaline rocks and fenites of the Amba Dongar complex, Gujarat, India. Neues Jb. Mineral. 1986, H.1, 37-48. WATKINSON,D. H. & WYLLIE, P. J. 1971. Experimental study of the composition join NaAISiO4-CaCO3H20 and the genesis of alkalic rock-carbonatite complexes. J. Petrol. 12, 357-78. WENDLANDT, R. F. & HARRISON, W. J. 1979. Rare earth partitioning between immiscible carbonate and silicate liquids and CO2 vapour: results and implications for the formation of light rare earthenriched rocks. Contrib. Mineral. Petrol. 69, 40919. WOOD, D . A . , JORON, J.-L., TREUIL, M., NORRY, M. J. & TARN~Y, J. 1979. Elemental and Sr isotope variations in basic lavas from Iceland and surrounding ocean floor. Contrib. Mineral. Petrol. 70, 319-39. WOOLLEY, m. R. 1982. A discussion of carbonatite evolution and nomenclature, and the generation of sodic and potassic fenites. Mineral. Mag. 46, 1317. - & JONES, G. C. 1987. The petrochemistry of the northern part of the Chilwa Alkaline Province, Malawi. In: FITTON, J. G. & UPTON, B. G. J. (eds) Alkaline lgneous Rocks, Geol. Soc. Spec. Publ. 30, pp. 335-55.
Nephelinites and carbonatites WYLLIE, P. J. 1965. Melting relationships in the system CaO-MgO-CO2-HzO, with petrological applications. J. Petrol. 6, 101-23. -1978. Silicate-carbonate systems with bearing on the origin and crystallization of carb0natites. Proc. 1st Int. Syrup. on Carbonatites, Ministerio das Minas e Energia, Departamento Nacional da
83
Produqao Mineral, Pogos de Caldas, Minas Gerais, Brazil, pp. 61-78. - - - 1980. The origin of kimberlites. J. geophys. Res. 85, 6902-10. YARDLEY,B. 1985. Apatite composition and fugacities of HF and HC1 in metamorphic fluids. Mineral. Mag. 49, 77-9.
M. J. LE Bks, Department of Geology, University of Leicester, Leicester, U,K.
Alkalic carbonatite magmas- parental or derivative? James D. Twyman & John Gittins S U M M A R Y : During the 1950s and 1960s the Tanzanian volcano Oldoinyo Lengai erupted a highly alkalic lava composed almost exclusively of the N a - K - C a carbonate minerals nyerereite and gregoryite and containing about 32~ Na20 and 7% K20. For many years the lava was considered a petrological curiosity but gradually the idea developed that alkalic carbonatite magmas might develop generally during the evolution of commoner carbonatite magmas. In 1981 Le Bas, followed in 1982 by Woolley, ascribed to the Oldoinyo Lengai magma a parental status and erected a scheme whereby other commoner carbonatite rock types are derived from it; the alkalic carbonatite magma is one of two derived by immiscible separation from a nephelinitic magma, the other being ijolitic. The carbonatite liquid has 500~ of superheat and loses alkalis progressively as the magma cools to its liquidus temperature of 400~176 by which stage it has become a calcitic-dolomitic liquid. This liquid then differentiates to produce the commoner types of carbonatite rocks. In rejecting such a scheme we argue as follows: such a scheme would require the magma to be water saturated at the moment of immiscibility, and progressive loss of water and alkalis would induce crystallization thus preventing the formation of a calcite-dolomite magma; the 500~ of superheat is a consequence of the design of the Freestone & Hamilton experiments and does not exist; carbonatite magmas generally begin to crystallize at temperatures greater than 900~ rare-earth-element compositions do not permit derivation of alkalic carbonatite magma from nephelinite magmas by liquid immiscibility; the mafic silicate mineralogy of carbonatites demonstrates alkali enrichment rather than depletion. We propose that the commonest parental magma of carbonatites is a mildly alkalic olivine s6vite composition and that most carbonatite rock types are derived from it by fractional crystallization, allowing the development of cumulates that are well lubricated by intercumulus liquid and capable of much subsequent movement, deformation and re-intrusion as a crystal mush. As this parental magma crystallizes, continued fractionation of alkalideficient and anhydrous minerals increases both its alkali and water contents until water saturation is reached. This in turn causes the development of an aqueous fluid which limits the alkali content of the magma. The composition of the magma at which water saturation develops is controlled by the rate of rise of the magma, and therefore aqueous fluids with various N a : K ratios can develop and control the type of fenitization that occurs. Under plutonic conditions alkalis in excess of the amount that can be fixed as silicate minerals are carried away as fenitizing fluids. Alkali loss does therefore occur, and through the medium of an aqueous fluid, but calcite and/or dolomite crystallize throughout the magmatic cooling history rather than simply at the low-temperature end of this history. The alkalic carbonatite magma of Oldoinyo Lengai type also develops through fractional crystallization of the same alkali-poor olivine s6vite magma, but under essentially dry conditions where the magma is kept liquid by alkalis and halogens, and alkali loss is prevented by the absence of an aqueous phase. Alkalic carbonatite magmas are therefore late differentiates of a more normal mildly-alkalic olivine s6vite magma developed under low water fugacity. They are of very small volume and require long quiescence to develop. They are not to be considered parental to other carbonatite rock types.
Introduction The problem of the c h a r a c t e r and origin of carbonatite m a g m a s is at once tantalizing and enigmatic. For m a n y years few geologists believed in the existence of m a g m a t i c c a r b o n a t e rocks at all, largely because of the high melting t e m p e r a t u r e of CaCO3 and failure to appreciate the effect of H 2 0 , F a n d alkali carbonates in lowering that temperature. T h e melting studies p e r f o r m e d by Wyllie and Tuttle a n d their coworkers b e t w e e n 1957 and the 1970s gradually o v e r c a m e these conceptual difficulties and the
petrological world was won over to the idea of carbonatite m a g m a s . Petrologists in the U. S.S.R., w h e r e there has always been a great propensity for m e t a s o m a t i s m , r e m a i n e d resistant to the idea of carbonatite m a g m a s , but they have n o w j o i n e d the igneous fold. Once the initial barrier of carbonate melting at reasonable t e m p e r a t u r e and pressure was overcome, the course of petrogenetic thought was g o v e r n e d largely by the melting studies of Wyllie and his students at the P e n n s y l v a n i a State U n i v e r s i t y and later at the U n i v e r s i t y of Chicago. T h e d e v e l o p m e n t of calcitic and dolomitic liquids
From: FITTON,J. G. & UPTON,B. G. J. (eds), 1987, Alkaline Igneous Rocks, Geological Society Special Publication No. 30, pp. 85-94.
85
86
J. D . T w y m a n & J. G i t t i n s
from various ultramafic ultrabasic compositions over a wide range of crustal and mantle pressures was pursued in this work. It must be remembered that this experimental programme was not designed solely to elucidate carbonatite genesis, but was at least as much concerned with the development of kimberlite magmas and magmas that might be parental to the basalts of the Earth's surface and upper crust. Consequently, Na and K did not figure prominently in these experimental systems. The general pattern emerged that calcitic and dolomitic liquids, given adequate H20 and CO2 fugacities, could develop from certain SiO2undersaturated peridotitic compositions such as might commonly be expected in the Earth's mantle. Consequently, they were thought to offer a probable explanation of how carbonatite magmas are formed. However, at the same time some petrologists were suggesting that these experimentally derived liquids were more applicable to the rather special type of carbonatite that is associated with kimberlites and the carbonate that so frequently pervades kimberlite intrusions, and that they might have relatively little relevance to the carbonatites of the classic carbonatitealkalic igneous rock association. Almost at the beginning of the melting studies an interesting turn of events was introduced when in 1960 the Tanzanian volcano Oldoinyo Lengai erupted its unusual N a - C a - K carbonate lava at temperatures too low to be incandescent. Earlier eruptions had, of course, been noted in the 1950s by Guest (1956), but were so unusual that the significance was not widely recognized. The initial tendency, with a few exceptions, was to consider it as a petrological curiosity of little application to the broader problems of carbonatite genesis, but it gradually assumed greater importance in the minds of certain petrologists. Dawson (1964) had originally suggested its importance and he was later supported by Cooper et al. (1975) and Gittins et al. (1975). In each of these papers it was argued that alkalic carbonatite magmas probably develop during the evolution of a carbonatite-igneous rock complex but that in most cases the alkalis are lost to aqueous or aqueous-halide fluids that may be involved in fenitization. The residues were thought to be normal calcitic and dolomitic carbonatites. In general it was felt that there was overwhelming evidence that most carbonatite magmas could not just be calcitic or dolomitic because of the frequent presence within the rocks of sodic pyroxenes, phlogopite-biotite and sodi-potassic amphibole, as well as fenitization adjacent to dykes and other intrusions; extreme examples were the presence of several centimetres of
phlogopite-rich rock in pyroxenite adjacent to carbonatite dykes. The concept emerged, and has now been accepted more widely, that carbonatite rocks do not necessarily represent the composition of the magma from which they crystallized. Le Bas (1981) argued the case that a highly alkalic magma of Oldoinyo Lengai type is parental to the full range of carbonatite rock types. His thesis was that carbonated nephelinite magma separates immiscibly into an ijolitic silicate liquid and an alkalic carbonate liquid at a temperature of 1000~176 The silicate liquid is said to be at its liquidus but the carbonate liquid is at least 500~ above its liquidus, with crystallization not commencing until 400~ 600~ The pressure at which the immiscible separation occurs was not stated but the context seems to imply lower-crustal rather than mantle conditions. The scheme envisaged by Le Bas involves the formation of a progressively lesssodic carbonate magma as alkalis are lost, presumably, although not clearly stated, during cooling through the 500~ superheat range, followed by crystallization differentiation that forms calcitic carbonatites at first and then increasingly more dolomitic and ankeritic carbonatites as the magma cools from its 400~ 600~ liquidus. Le Bas further suggested that the proportion of Ca: Na: K in the alkalic carbonatite magma may be a function of the temperature and pressure at which immiscible separation occurred, a conclusion drawn from the experimental work of Freestone & Hamilton (1980) and which he referred to as 'dynamic un-mixing'. The variable proportion of Na-rich to K-rich fenites was attributed to possible temperature,controlled variations in relative Na and K mobility, which Gittins et al. (1975) had also invoked. The following year Woolley (1982), while supporting the evolutionary concept of the formation of calcitic, dolomitic and ankeritic carbonatites from alkalic magma through alkali loss and fractional crystallization, introduced a modification by suggesting that the aqueous fluid becomes progressively more potassic. In this way he sought to explain the superposition of potassic fenitization on early sodic fenitization observed in some carbonatites. Thus the highly alkalic carbonatite magma of Oldoinyo Lengai was promoted from a stage in the evolution of carbonatite complexes (Cooper et al. 1975) to a parental magma derived immiscibly from a carbonated nephelinite magma. Purportedly, it gives rise, through progressive loss of alkalis followed by fractional crystallization, to the full diversity of carbonatite rock types. It is this parental status and proposed differentiation scheme with which we disagree. We
87
Alkalic carbonatite magmas argue our case by examining progressively the evidence of an immiscible origin, the evidence for or against alkali depletion, rare-earth-element (REE) data, the problems of generating and then abstracting an aqueous fluid phase and problems of magma temperatures.
Critique of the parental-status scheme The Freestone & Hamilton experiments Freestone & Hamilton (1980) established that molten nephelinite and phonolite are each immiscible with molten alkalic carbonatite lava. What was done was the following: a nephelinite and an alkalic carbonatite lava, both of which had erupted from the volcano Oldoinyo Lengai, were mixed together; the mixture was heated to a temperature above the liquidus of both rocks and then quenched in order to establish texturally that two immiscible liquids had coexisted. The experiments were repeated with phonolite and produced similar results. It was concluded that the silicate and carbonate liquids separated in the volcano from a common parent, for which phonolite was preferred to nephelinite. However, the experiments do not prove this. They establish that the two liquids are immiscible at 1 kb but not that they came into being together. There is a weakness in the design of the experiments. If the alkalic carbonatite is a highly evolved rock with a low liquidus temperature (about 500~ as we shall attempt to show, then melting it with a silicate rock whose liquidus is about 1000~ is bound to produce liquids with sharply divergent liquidi. The superheat is a consequence of the experimental design, and the argument is dangerously circular. Two rocks from the same volcano, but which may have been derived by fractional crystallization of two entirely different magmas, are melted together and found to produce immiscible liquids with liquidi 500~ apart. This is held to prove that the two rocks represent the compositions of magmas that were derived immiscibly from a common parent. The experiments prove what is possible but not that the possible happened. It is analogous to calculations that show a reaction to be thermodynamically feasible but not whether it is kinetically possible. We agree that alkalic carbonatite and silicate magmas are mutually immiscible, but we believe that the immiscibility occurred a long way back in the petrogenetic history, and in the mantle.
Alkali depletion or alkali enrichment? The arguments surrounding the alkalic carbonatite question have become rather theoretical with insufficient attention being paid to real rocks. Natural carbonatites can settle the question of whether alkali depletion occurs, for if it does the associated silicate minerals ought to reflect this depletion. Amphiboles from three calcitic and dolomitic carbonatite complexes are shown in Fig. 1. The direction of amphibole evolution is established by compositional zonation and is, from core to rim, an increase in Na + Fe accompanied by a decrease in Ca + Mg. In short, it is the reverse of what is required by the parentalstatus scheme. Rare-earth-element data It is possible to calculate the REE distribution in a carbonate liquid that would be in equilibrium with the Oldoinyo Lengai phonolite by using the partition coefficients published by Wendlandt & Harrison (1979). The result is shown in Fig. 2 where it is compared with the measured patterns of the alkalic carbonatite lava and both the nephelinite and the phonolite. Even after allowing for the fact that the nephelinite and phonolite analyses are not of the highest quality there is a considerable discrepancy between the calculated carbonatite pattern and the measured pattern, 0"7 [ " OLIVINE CARBONATITES l + CLINOHUM,CARBONATITES/
t( ,
~
l-Z+. ~ \
0.5
Co
9 ~%t
0.4
No+K 0.5 0.2 0.1 ol
"'s
~
,
I
l I0
,
I , 20 30
Mg/Fe FIG. 1. Compositions of amphiboles from the Argor and Goldray carbonatites, Ontario, Canada (Twyman 1983) and the Safartoq carbonatite, Greenland (Secher & Larsen 1980). The arrows indicate the direction of compositional change from core to rim. The petrographic varieties represented are those with amphibole alone, amphibole with titanian clinohumite and amphibole with olivine. The general trend of amphibole evolution is enrichment in Na, K and Fe with depletion in Ca and Mg.
88
J. D. Twyman & J. Gittins IO3
\.
,o'
-
--__-~ I I LO Ce
I Nd
I I 1 SmEu Tb
II YbLu
FIG. 2. Chondrite-normalized REE distribution for Oldoinyo Lengai lavas. The calculated composition of the alkalic carbonatite ( - - - - - - ) is based on the distribution coefficients of Wendtlandt & Harrison (1979) for a carbonatite in equilibrium with nephelinite and a phonolite whose measured values are taken from Gerasimovsky et al. (1972). The measured values of the alkalic carbonatite are taken from Twyman (1983) ( - - . . - - ) and Philpotts et al. (1972) (----). and the divergence increases towards the heavy REE. It has been argued already that the melting experiments performed by Freestone & Hamilton (1980) offer no proof that alkalic carbonatite magma was derived from nephelinite magma. The REE data provide further support for this view. The problem of generating an aqueous fluid from the alkalic carbonatite m a g m a
Progressive alkali loss is supposed to have occurred through 'exchange processes [with the] enveloping country rock' and 'by reaction' (Le Bas, 1981, pp. 138, 139). The aqueous fluid that has been invoked by other workers (cited earlier) seems to be inherent in the proposed alkali loss, and so such a fluid must have been generated from the carbonate liquid in some manner. That, in turn, requires the liquid to have reached water saturation. The available possibilities are that the carbonate liquid was saturated at the moment of immiscible separation from its alleged nephelinitic parent, or that it reached saturation during cooling. It is unlikely that the carbonate liquid reached saturation during cooling. The commonest process by which an aqueous fluid is developed from a magma is the progressive crystallization of anhydrous minerals which increases the concentration of water in the remaining liquid. Although the minerals of the Oldoinyo Lengai carbonatite lava are anhydrous, this process is not available
because we are asked to believe that the magma cooled through 500~ of superheat. It was therefore above its liquidus and so could not have been crystallizing. It is possible for a magma to reach water saturation as it rises to a regime of lower pressure but only if the solubility of water in it is relatively low. Solubility data for alkalic carbonatite magmas are incomplete, but it is probable that the solubility is very high and that saturation would not be reached in this way. The remaining alternative, that the carbonate liquid was water saturated at the moment of immiscibility, is also fraught with difficulty. This would require the ijolitic liquid as well as the alkalic carbonatite liquid to be saturated. Yet ijolites, being anhydrous rocks, appear to have crystallized from fairly dry magmas. Furthermore, nephelinitic magmas have fairly low water contents and so, for the alkalic carbonate magma to reach saturation, the available water would have to partition almost exclusively into the carbonate liquid. However, available data on water solubility in nepheline-bearing liquids suggest that the solubility of water in ijolitic magma would not be inconsiderable. The problem of generating an aqueous fluid within the constraints of the parental-status scheme appears to be insuperable, yet without such a fluid the scheme does not work. The effect of continuous removal of water and alkalis from a m a g m a
Suppose that an aqueous fluid could be generated in some manner. What would be the effect of its continued removal from the alkalic carbonatite magma? It has been well established experimentally, through the studies of Wyllie, Gittins and their coworkers, that calcitic and dolomitic liquids are possible primarily through the grace of water and alkalis. If either, or both, are removed, crystallization is the inevitable result. This is especially so since the solubility of water in calcitic and dolomitic liquids is markedly less than in sodi-potassic carbonate liquids. In short the proposed method by which the magma composition changes would be fatal to the magma's continued existence. In addition, under the proposed magmatic evolutionary scheme there is no reason why the Oldoinyo Lengai alkalic carbonatite lava should have erupted at all. And if it did, it ought to have been a vigorous eruption as its aqueous phase escaped and the remaining water in the saturated magma boiled offexplosively. In contrast, the 1960 eruption was quiescent and composed of anhydrous lavas. If the proposed alkali loss was not through the medium of an aqueous fluid but rather by wall-
89
Alkalic carbonatite magmas rock reaction, the same arguments apply. Reaction of such a magma with the enclosing silicate rocks must, inevitably, lead rapidly to crystallization.
Magma temperatures and crystallization temperatures The parental-status scheme has the crystallization of calcitic and dolomitic magmas commencing at 400~176 and is in accord with much popular thinking about the temperature ranges over which carbonatite magmas crystallize. Unfortunately, there is a good deal of misleading thought on this subject. Low temperatures (below 500~ are overemphasized, largely through misuse of the calcite-dolomite geothermometer. The fact that temperatures obtained from measurements of the amount of Mg substituting Ca in calcite that coexists with dolomite are minimum temperatures is often overlooked. If dolomite has exsolved from the calcite electron microprobe analysis cannot generally be used, and even if the calcite is homogeneous it cannot be assumed that there has been no diffusion of Ca and Mg between calcite and dolomite after crystallization, particularly if a fluid phase is present. Published crystallization temperatures are rarely over 550~ and many are as low as 200~ Yet it can be shown that in at least one carbonatite complex (Argor, N Ontario, Canada) a carbonate mineral began to crystallize at a temperature higher than 885~ The bulk composition of coarsely exsolved single crystals (74 CACO322 MgCO3-4 FeCO3) gives an equilibrium temperature of 885~ indicating that the liquidus temperature of the carbonatite magma was at least that high. Recrystallization and re-equilibration during cooling render accurate geothermometry in carbonatites very difficult and have often caused overemphasis on the 500~ range (Gittins 1979). Argor offers proof of the existence of a carbonatite magma at not less than 900~ and we suggest that this is an approximate liquidus temperature for many carbonatite magmas. It is, of course, in sharp contrast with the 400~ 600~ liquidus temperature required by Le Bas for the magma from which calcitic and dolomitic carbonatites crystallized.
The volume requirements of the alkali-depletion scheme A very general idea of the volume of material that would have to be removed from an alkalic carbonatite magma to produce a calcitic-dolomitic carbonatite can be obtained from a simple
subtraction diagram. This does not require knowledge of the fluid phase compositions since we can assume that everything in the fluid came originally from the magma. In Fig. 3 the data are the compositions of the Oldoinyo Lengai alkalic carbonatite lava, the solid composition of the subtracted phase and the average carbonatite of Gold (1966) which, despite the limited value of averages from that period, suffices for the purpose. The composition of the subtracted phase is taken at FeO = 0, and the diagram indicates that 95~ of the alkalic carbonatite magma must be removed to convert it into normal carbonatite. If the calculations are repeated for a subtracted phase with SiO 2 -- 0, the proportion is still 90%. If we are to assume that all normal calciticdolomitic carbonatites are the residue remaining after the removal of 9 0 ~ - 9 5 ~ of the mass of a parental alkalic carbonatite magma the volume requirements become unrealistically enormous, even if it were possible to have a magma still in existence after abstracting so much water and alkalis. Again, if wall-rock reaction rather than fluid involvement is responsible for the alkali loss the proportion of carbonatite to metasomatic encircling rock would have to be vastly different from what is generally observed.
Origin of alkalic carbonatite magma We have attempted to show that alkalic carbonatite magma is not parental to the more normal w t~ < -r
laJ
w
R.~o
o
----Z
Z
4 0 " ""~:
~"
OXIDE IO
0
I 2
5
4
5
6
S iO 2
FIG. 3. Subtraction diagram showing the volume and composition of a fluid phase that must be removed from an alkalic carbonatite to produce a normal carbonatite. See text for further discussion.
J.D. Twyman & J. Gittins
90
calcitic-dolomitic carbonatites. Yet it clearly exists. Where, then, does it come from? We believe that it can be generated by fractional crystallization of a mildly-alkalic olivine s6vite magma which we propose as a fairly common carbonatite parental magma. The arguments in support of such a composition are to be published separately but are based on the magnesian composition of carbonatite olivines (most commonly F095_85), which implies generation of the magma from a very mafic source, and upon the partition coefficients derived by Freestone & Hamilton (1980) from their melting studies. Although we do not believe that the results of their experiments demonstrate what they set out to prove, the data can be used to estimate the N a 2 0 and K 2 0 content of carbonatite liquids in equilibrium with immiscible silicate liquids of
various compositions. We believe that the olivine s6vite magma is the result of immiscible separation from an olivine nephelinite magma at a pressure of approximately 27 kb. The partition coefficients for nephelinite give N a 2 0 + K 2 0 18~, but a more primitive melanephelinite gives a value of about 8%. Our conclusion is that the commonest parent magma that generates the diversity of carbonatite rocks is an olivine s6vite with about 8% N a 2 0 + K20. A computer-modelling scheme has been employed to test whether an Oldoinyo Lengai alkalic carbonatite lava can be produced by fractional crystallization of this mildly-alkalic olivine s6vite magma. The compositions used are an olivine s6vite from the Argor intrusion (Ontario, Canada) with additional alkalis to bring N a 2 0 + K 2 0 up to 8.1~, and the mineral compositions from
TABLE 1. Results of the computer modelling of the differentiation of carbonatite magmas
SiO2 A12Oa TiOz FeO MgO MnO CaO Na20 K20 PEO5 CO2 Total
Weighting factor
Parent olivine s6vite
5.0 1.0 1.0 20.0 5.0 0.1 10.0 50.0 10.0 50.0 5.0
6.56 10.66 1.09 10.66 3.68 0.25 34.50 5.74 2.41 5.92 28.17 100.00
Target alkalitic carbonatite 0.66 0.33 0.11 0.33 1.34 0.16 17.80 33.86 8.18 1.09 36.34 100.00
Calculated target 0.66 0.33 0.12 0.33 1.34 0.14 17.85 33.84 8.30 1.09 36.66 100.44
Difference (%) 0.0 0.0 1.0 0.0 0.2 - 14.6 0.3 - 0.1 1.5 0.0 0.9
Minerals removed ct SiO2 A1203 TiO 2 FeO MgO MnO CaO NazO
1.77 0.68 0.36 51.07
ilm
51.70 48.30
mag
apt
3.05 96.84 0.11 57.10
dt
58.39
42.69 9.93 0.45 16.50 20.05 0.08 0.09
9.52 14.00 1.65 29.34
2.04 8.16 1.74
P2Os Total Solution
bt
0.03 12.57 17.07
K20 CO 2
amph
10.21
42.90 46.13
45.50
100.00
100.00
100.00
100.00
-44.67
-1.57
-6.80
-13.40
100.00 -3.78
100.00 -9.95
ct, calcite; ilm, ilmenite; mag, magnetite; apt, apatite; amph, amphibole; bt, biotite; dt, dolomite. 83.87~ crystallization of parent olivine s6vite will generate a residual alkalic carbonatite.
100.00 -3.70
Alkalic carbonatite magmas the Argot carbonatites. The computer program is a modification by Dr. M. P. Gorton (University of Toronto) of the linear-programming and leastsquares method P E T M I X due to Wright & Doherty (1970). It is essentially a reiterative regression analysis that minimizes the squares of the residuals. The results of the calculation are given in Table 1, and in summary it can be said that 84~ crystallization can generate a lava of Oldoinyo Lengai type by removal of minerals that would form a rock composed of 53% calcite, 4 ~ dolomite, 16~ apatite, 8 ~ magnetite, 2 ~ ilmenite, 12~ biotite and 5 ~ amphibole. This will be recognized as a common type of s6vite. We do not claim a definitive status for the calculation but we believe that it does demonstrate that alkalic liquid can be derived through crystal fractionation of mildly-alkalic olivine s6vite magma. Further support for this concept comes from the chemical compositions of the amphiboles and phlogopite-biotites from the Argor intrusion. Throughout most of the crystallization of the magma there is very little change of A1 relative to Mg/Fe in amphibole (Fig. 4) or of Ti relative to Mg/Fe in biotite (Fig. 5), but in a small number of highly evolved rocks there is a substantial increase amounting to tenfold for A1 and thirteenfold for Ti. Both elements can be treated as trace elements and the curves are strikingly similar to trace-element enrichment curves for liquid predicted by Rayleigh fractionation with bulk crystal-liquid partition coefficients of less than unity. These enrichments can be used to estimate the amount of crystallization required to produce them: about 9 0 ~ - 9 2 % They offer further support for the concept that alkalic carbonatite magmas are produced by fractional crystallization.
An alternative petrogenetic scheme We propose that the alkalic carbonatite magmas known to have erupted from Oldoinyo Lengai, and possibly from other volcanoes (Dawson 1964; Hay & O'Neil 1983, Deans & Roberts 1984; Hay 1984), are of small volume and are derived from a more primitive parental olivine s6vite magma (with about 8 ~ alkalis) by fractional crystallization and gravity separation. Differentiation of this magma can follow more than one course. Development
of normal
carbonatite
The general course of differentiation, whether wet or dry, is toward increasingly alkalic compositions, and this is portrayed by the silicate minerals (principally the amphiboles). Initial magma temperatures are high (close to 1000~
91 -I
0.5
-!0 9 8 7
0.4 AI 0.:5
~
6
5 2-64 CAI
0.2 0.1 0 0
i'. "A~
5 2 " .,,v~- I J t t I I t I J 0 I 2:54 5 6 78 9
Mg/Fe FIG. 4. Compositions of amphiboles from the Argor carbonatite showing variation of A1 with Mg/Fe. Note the dramatic increase in AI as Mg/Fe falls below a value of 2. 15
0.30 0.25 0.20
7.5
CTi o
5
CTi
"
Ti 0.15 0.10 ~o
0.05 0
9
i
i
I
2
9
~ oo
i 5
I .
i
i
4
5
:o
i 6
9
~
.25
I
7
Mg/Fe
FIG. 5. Compositions of biotites from the Argor carbonatite showing variation of Ti with Mg/Fe. Note the dramatic increase in Ti as Mg/Fe fails below 2. and are similar to those of the associated silicate magmas. The alkali content of about 8 ~ and the fairly low water content are sufficient to keep it liquid at high pressure. (A study of the system calcite-dolomite with sodium and potassium carbonate at 1 kb (Beckett 1986) has established that at 8 ~ N a 2 0 the system would have extensive liquid at 900 ~ C.) Crystallization of dominantly anhydrous minerals and, initially, minerals of low alkali content (calcite, dolomite, apatite and Fe-Ti oxide) drives up both the alkali and the water contents which become increasingly concentrated in a diminishing volume of magma and progressively reduce its viscosity. At some stage the magma will become water saturated and an aqueous fluid will separate carrying with it most of the alkalis that have not already been fixed as silicates. The stage at which this occurs will be strongly dependent on the pressure and hence upon the rate at which
92
J. D. Twyman & J. Gittins
the magma has risen through the crust. Thus it is possible for separation of the aqueous phase to occur at various degrees of alkali enrichment according to the rate of magma rise. For example, a rapidly rising magma will not have sufficient time for crystal fractionation to produce a high degree of alkali enrichment before water saturation occurs and so the escaping fluid will carry off very little N a 2 0 + K 2 0 . In contrast, a magma that has risen slowly will have fractionated to an advanced degree, generating a high water content, and so its aqueous fluid will be very alkalic and a powerful fenitizing agent. The wide variations in the extent of fenitization associated with carbonatite complexes can be explained in this way. Woolley's (1982) suggestion that the relative proportion of sodic to potassic fenitization is depth dependent also fits into this scheme. In this scheme magmas of varied alkalic enrichment are possible as a result of fractional crystallization, the degree of enrichment being limited only by the stage at which water saturation develops and an aqueous fluid phase separates carrying off alkalis in excess of the amount that can be fixed as silicates. We suggest that, under plutonic conditions, the process is prevented, by continuous alkali loss, from producing a magma of the composition of the Oldoinyo Lengai lava.
Development of alkalic carbonatites How then does the extreme alkalic carbonatite magma develop? The discussion so far has focussed on two factors that serve to reduce the liquidus temperature of our proposed parental carbonatite magma: increasing water content and increasing alkali content. However, water has the further effect of removing alkalis once water saturation is reached, and so it seems unlikely that the extreme alkalic composition can be reached in a hydrous magma. We pointed out earlier that the aqueous fluid phase will control the alkali content of the magma. Furthermore, the mineralogy of the Oldoinyo Lengai lavas shows that the magma was very dry when erupted. The answer to the dilemma seems to be that the extreme alkalic magma is produced by fractional crystallization under fairly dry conditions. We have already noted that an olivine s6vite with about 8% alkalis will be largely liquid at 900~ but factors other than alkali content can be invoked to reduce the liquidus in such a magma. The Oldoinyo Lengai lava contains as much as 8% F and as much as 4% C1. F is not at all uncommon in carbonatite complexes as indicated by fluorite (see yon Eckermann (1948) on Aln6) and the extensive substitution of F for OH in many carbonatite amphiboles and micas. F
and C1, together with the alkalis, would maintain the parental carbonatite magma in a liquid state throughout its crystallization history. Increase in alkali content, through continued fractionation of alkali-free minerals and a very small percentage of alkali-bearing minerals, is able to continue uninterrupted by alkali loss since an aqueous fluid phase does not develop. A continuum of magmatic compositions from the initial 8% alkalis to the Oldoinyo-Lengai-type extreme of about 37% is possible, but the latter can be reached only if the magma is dry. In this scheme the majority of carbonatite rock types would be formed by cumulus processes as crystals sink through a very fluid magma. They are not cumulates in the passive sense of, for example, gabbro cumulates. The magmatic plumbing system in a carbonatite complex is probably much more dynamic so that the cumulates, with their lubricating interstitial liquid, are readily disturbed, re-intruded and deformed by pulses of new magma. Much of the chaotic intrusive structure of carbonatite complexes can probably be explained in this way.
Conclusions Alkalic carbonatite magma of the type erupted as lava in the Tanzanian volcano Oldoinyo Lengai is not a parental liquid from which other carbonatite magmas and rock types are derived. Rather, it is a late-stage liquid derived through fractional crystallization of an olivine s6vite magma containing about 8% N a z O + K 2 0 . In the parental-status scheme nephelinite magma divides into two immiscible liquids, of which one is ijolitic and the other is an alkalic carbonatite composition with about 500~ of superheat. During cooling, the latter continuously loses alkalis through the medium of an escaping aqueous fluid phase or by wall-rock reaction, and it changes composition until it becomes a calciticdolomitic magma which differentiates further to give rise to a wide variety of residual products. We have argued that, although the details are not brought out in the original paper, the scheme requires that the alkalic carbonatite liquid must be water saturated from the moment of its creation and that progressive loss of water and alkalis will induce crystallization rather than development of a calcitic-dolomitic magma. We have argued that the scheme is not supported by the geothermometry of carbonatite rocks or by REE studies and that mineral zonation in carbonatite amphiboles implies alkali enrichment in the magma rather than alkali depletion.
Alkalic carbonatite magmas The 500~ of superheat is a consequence of experimental design rather than a logical deduction from unconstrained experiments, and does not exist. We propose an alternative petrogenetic scheme in which the parent magma, from which carbonatite rock types are derived, is a mildly-alkalic olivine sSvite composition with about 8 ~ total alkalis. This magma probably originates by immiscible separation from an olivine nephelinite magma deep within the mantle, a proposal that makes less extreme demands on element and component partitioning than does the parentalstatus scheme. During its subsequent rise into and through the crust it undergoes fractional crystallization in which anhydrous and alkalipoor minerals sink, leaving the residual liquid progressively more enriched in alkalis and water. Water saturation is eventually reached but this can occur over a wide range of compositions and pressures governed by the rate at which the magma rises and the time available for differentiation. Rapid rise with little time for fractionation will generate saturation at a relatively low alkali content. Slow rise will generate water saturation at a much later stage and a higher alkali content because of the much higher solubility of water in alkalic carbonatite liquids than in calcitic-dolomitic liquids. The alkali content of the rocks is limited by the amount that can be fixed as silicates and this in turn is limited by the activities of Si and AI, both of which are
93
very low. Separation of an aqueous fluid phase when water saturation is reached controls the alkali content of the remaining magma, and is itself controlled by the solubility of alkali carbonates and halides in supercritical water and by the rate at which anhydrous alkali-poor minerals continue to crystallize. Thus in most plutonic settings the crystallizing magma will lose alkalis at some stage when water saturation is reached; this may be a longcontinued process providing a source of fenitizing fluids. The pre- and post-saturation history of the magma will probably control the N a : K ratio of the fluid and have a profound bearing on the type of fenitization that ensues. A major difference from the parental-status scheme is that the water content of the magma gradually increases through fractional crystallization until saturation is reached, rather than escaping continuously from an already alkalic liquid that is merely cooling but not crystallizing. If the parental olivine s6vite magma has exceptionally low water and high halogen contents, alkali enrichment might reach a very advanced stage and even develop the extreme composition of the Oldoinyo Lengai lava. This will be possible because the magma does not reach water saturation and so alkalis are able to accumulate continuously without being abstracted by an aqueous fluid. Alkalic carbonatite magmas are late derivatives and not primary liquids.
References BECKETT, M. F. 1986. Phase relations in alkali-bearing dolomite carbonatites; effect of alkalinity and fluorine content on the solubility of pyrochlore in carbonatite magma. M.Sc Thesis, University of Toronto, Canada (unpublished). COOPER, A. F., GITTINS,J. & TUTTLE,O. F. 1975. The system NazCO3-KzCO3-CaCO 3 at 1 kilobar and its significance in carbonatite petrogenesis. Am. J. Sci. 275, 534-60. DawsoN, J. B. 1964. Reactivity of the cations in carbonatite magmas. Geol. Assoc. Can. Proc. 15, 103-13. DEANS, T. & ROBERTS,B. 1984. Carbonatite tufts and lava clasts of the Tinderet foothills, western Kenya: a study of calcified natrocarbonatites. J. geol. Soc. Lond. 141, 563-80. VONECKERMANN,H. 1948. The alkaline district of Aln5 Island. Sver. geol. Unders., Ser. C, 36. FREESTONE,I. C. & HAMILTON,D. L. 1980. The role of liquid immiscibility in the genesis of carbonatites--an experimental study. Contrib. Mineral. Petrol. 73, 105-17. GERASIMOVSKY,V. I., ALASHOW,Yu. A. & KARPUSHINA, V. A. 1972. Geochemistry of the rare earth
elements in the extrusive rocks of the East African rift zones. Geochem. Int. 1972, 305-19. GITTINS, J. 1979. Problems inherent in the application of calcite-dolomite geothermometry to carbonatites. Contrib. Mineral. Petrol. 69, 1-4. - - , ALLEN, C. R. & COOPER,A. F. 1975. Phlogopitization of pyroxenite; its bearing on the composition of carbonatite magmas. Geol. Mag. 112, 5037. GOLD, D. P. 1966. The average and typical chemical composition ofcarbonatites. Proc. 4th Gen. Meeting Inter. Mineral. Assoc., 1964, India, 83-9. GUEST, N. J. 1956. The volcanic activity of Oldoinyo L'Engai, 1954. Tanganyika geol. Surv. Rec. 1954, 4, 58-9. HAY, R. L. 1984. Natrocarbonatite tephra of the volcano Kerimasi, northern Tanzania. Geology, 11, 599-602. -& O'NEm, J. R. 1983. Carbonatite tufts in the Laetolil Beds of Tanzania and the Kaiserstuhl in Germany. Contrib. Mineral. Petrol. 82, 403-6. LE BAS, M. J. 1981. Carbonatite magmas. Mineral. Mag. 44, 133-40.
94
J. D. Twyman & J. Gittins
PHILPOTTS, J. A., SCHNETZLER,C. C. & THOMAS,H. H. 1972. Petrogenetic implications of some new geochemical data on eclogitic and ultrabasic inclusions. Geochim. cosmochim. Acta, 36, 113166. SECHER, K. & LARSEN, L. M. 1980. Geology and mineralogy of the Sarfartoq Carbonatite Complex, southern West Greenland. Lithos, 13, 199-212. TWVMAN, J. 1983. The generation, crystallization, and differentiation of carbonatite magmas: evidence from the Argor and Cargill complexes, Ontario. PhD Thesis, University of Toronto (unpublished).
WENDLANDT, R. F. & HARRISON, W. J. 1979. Rare earth partitioning between immiscible carbonate and silicate liquids and CO2 vapor: results and implications for the formation of light rare earthenriched rocks. Contrib. Mineral. Petrol. 69, 40419. WOOLLEY, A. R. 1982. A discussion of carbonatite evolution and nomenclature, and the generation of sodic and potassic fenites. Mineral. Mag. 46, 13-7. WRIGHT, T. L. & DOHERTY, P. C. 1970. A linear programming and least squares computer method for solving petrologic mixing problems. Bull. Geol. Soc. Am. 81, 1995-2008.
JAMES D. TWYMAN* & JOHN GITTINS, Department of Geology, University of Toronto, Toronto M5S 1A1, Canada. * Present address: Arco Oil and Gas Company, Lafayette, LA, U.S.A.
The kimberlite clan: relationship with olivine and leucite lamproites, and inferences for upper-mantle metasomatism J. B. Dawson S U M M A R Y : Kimberlites are rare ultrabasic potassic low-volume melts that originate in the diamond stability field of the upper mantle. They are petrographically complex, and wide
mineralogical and chemical variations suggest that they should be regarded as a group or clan rather than a single narrowly-defined rock type. In S African kimberlites there is a major division between highly-micaceous (Group II) kimberlites that are mineralogically, chemically and isotopically distinct from poorly-micaceous (Group I) kimberlites. Nd and Sr isotope studies of diamondiferous olivine and leucite lamproites from NW Australia indicate that they are isotopically similar to Group II kimberlites but are distinct from Group I. New bulk rock analyses of Group II kimberlites indicate closer chemical similarities with olivine lamproites than previously supposed. and carbonatisation. Kimberlite commonly contains inclusions of upper mantle derived ultramafic rocks. Variable quantities of crusKimberlites are rare ultrabasic potassic igneous tal xenoliths and xenocrysts may also be rocks occupying small vents, sills and dykes. present. Kimberlite may contain diamond but They are petrographically complex because of only as a very rare constituent.' variations in texture (diatreme facies versus The number of qualifications (my italics) in this hypabyssal facies (Dawson & Hawthorne 1970)), definition serves to illustrate the petrographic variability and particularly the range of mineral variations in xenolith and xenocryst content, variable content and proportions of macrocrysts, species developed in the matrix. The mineralogy and variations in the mineralogy of the fine- is also complicated by the fact that some species grained matrix, all of which combine to give a (e.g. calcite and apatite) that are relatively wide range of chemical composition (Dawson common in hypabyssal-facies kimberlites have 1967, 1980). These variations, particularly in not developed to the same extent in diatrememineralogy, have created difficulties in defining facies kimberlites, presumably owing to volatile kimberlite, as exemplified by the recent definition l o s s during diatreme formation; also, certain of kimberlite by Clement et al. (1984): species such as monticellite are extremely suscep'Kimberlite is a volatile-rich potassic, ultratible to calcitization under high Pco2 conditions. basic igneous rock which occurs in small Despite these complications it can be seen that volcanic pipes, dykes and sills. It has a individual kimberlites, although not necessarily distinctively inequigranular texture resulting containing identical matrix minerals, do possess from the presence of macrocrysts set in a some permutation of the minerals listed by finer-grained matrix. This matrix contains, as Clement et al. (1984), and their similar and prominent primary phenocrystal and/or overlapping mineralogies suggest that they should groundmass constituents, olivine and several be regarded as a group or clan of closely-related of the following minerals: phlogopite, carbonrock types rather than a single tightly-defined ate (commonly calcite), serpentine, clinopyrock type. roxene (commonly diopside), monticellite, There is nonetheless a fundamental division apatite, spinels, perovskite and ilmenite. The within the kimberlite clan, recognized as such by macrocrysts are anhedral, mantle-derived, Wagner (1914), between 'lamprophyric' kimberferromagnesian minerals which include olilites having a highly micaceous matrix, and the vine, phlogopite, picroilmenite, chromian spiso-called 'basaltic' variety containing little or no nel, magnesian garnet, clinopyroxene mica in the matrix. (The term 'basaltic' is quite (commonly chromium diopside), and orthopyinappropriate and Dawson (1980) has proposed roxene (commonly enstatite). Olivine is exthat it should be dropped.) In addition to the tremely abundant relative to the other major difference in the phlogopite content of the macrocysts, all of which are not necessarily matrices of these two types of kimberlite, there present. The macrocrysts and relatively-earlyare certain other differences. formed matrix minerals are commonly altered 1 In the case of S Africa, the micaceous by deuteric processes, mainly serpentinisation
Varieties of kimberlite
From:FITTON, J. G. & UPTON,B. G. J. (eds), 1987, Alkaline Igneous Rocks, Geological Society Special Publication No. 30, pp. 95-101.
95
J. B. Dawson
96
kimberlites were intruded 120-200 Ma ago which contrasts with the non-micaceous varieties intruded 80-90 Ma ago. 2 The micaceous kimberlites tend to contain considerable amounts of matrix calcite and apatite. 3 The micaceous kimberlites generally also contain groundmass diopside, which is quite rare in the non-micaceous types (Dawson et al. 1977). 4 Although macrocrystal picro-ilmenite is common in the non-micaceous kimberlite, it is relatively rare in micaceous varieties. 5 The micaceous kimberlites are occasionally accompanied by other high-K minor intrusives, as at the Helam Mine in the Swartruggens area of the western Transvaal where an E-W-trending dyke-swarm contains both diamondiferous kimberlite and non-diamondiferous sanidine-nepheline-lamprophyre (the so-called 'Male' dyke) (Skinner & Scott Smith 1979). Another major difference between the micaceous and non-micaceous kimberlites in S Africa has recently been recognized on isotopic grounds (Smith 1983). The two groups are quite distinct by reason of their U, Pb, Sm, Nd and Sr contents
and isotopic ratios, which are interpreted as reflecting substantial chemical differences in the mantle source rocks for the two groups. The nonmicaceous kimberlites (denoted Group I by Smith) are derived from undifferentiated to slightly depleted sources relative to bulk earth (Sr and Nd systematics) but with high U/Pb ratios. In contrast, the micaceous kimberlites (referred to as Group II by Smith) are derived from an enriched source with high Rb/Sr and Nd/Sm ratios but with low U/Pb ratios; it was suggested by Smith that this enrichment event took place in excess of 103 Ma ago. The isotopic differences between the Group I and Group II kimberlites are well illustrated on a Sr versus Nd isotope plot (Fig. 1) on which the Helam Mine lamprophyre dyke, referred to above, falls close to the micaceous kimberlites. In summary the term 'kimberlite' in S Africa has been used to cover rocks that are distinct petrographically and is 9 and which, furthermore, were intruded at different times. Their common factors are that they are both products of low-volume ultrabasic upper-mantle melts originating in the diamond stability field of the Earth's upper mantle in the late-Jurassic~Sr
-50
0
50
100
150
200
I
]
I
I
i
i
I~MORB I
[] Ohvme " ' lamproltes "
\ ' ~ \ DEPLETED MANTLE (DM) '~'~
9 Leucite lamproites
~ ,
~" Helam lamprophyre
" ..... . a/ ,J,h oc~' K~.~ ~3". ~ o ~"N~. 9'BULTFONTEIN / \ "
/'%
-4
o~J .... /OCEAN ISLANDBASALTS
Gpl ~"-~ \ KIMBERLITES ~ ,, / /
9 Phases in Bultfontein metasomites
~'\
"~
~/(Sr/Nd)gM
_ 2
~'~~Sr/~d)EM
!_Sr/Nd!DM _ 10 ~ (Sr/Nd)EM ~ \ . -12
?-
-
~:r-
~.~AUSTRALIA 9
-'---ZD
Gpll KIMBERLITES
-~.
[] "'0---_.
~ . ~
_
.
//
-16
0.705
I 0.710
I 0.715
ENRICHED MANTLE (EM) I 0.720
87Sr/86Sr (I) FIG. 1. Nd versus Sr isotope plot for kimberlites, lamproites and Bultfontein metasomites. Data for kimberlites and the Helam lamprophyre are from Smith (1983) and those for the Bultfontein metasomites are from Kramers et al. (1983). The lamproite data, the points for depleted and enriched mantle and the depleted/enriched mantle value curves are from McCulloch et aL (1983).
The kimberlite clan early-Cretaceous period. The proposal to regard kimberlite as a clan is particularly appropriate now that it has been shown that differing members have been derived from isotopically distinct areas within the upper mantle.
Relationship between kimberlites and lamproites Although diamond has been mined in the C6te d'Ivoire from dykes described as 'fitzroyites' for some years (Knopf 1970), the discovery of diamonds in olivine- and leucite-bearing lamproites in W Australia has stimulated a number of studies into the relationship between kimberlites and lamproites (Mitchell 1981 ; Atkinson et al. 1984; Jaques et al. 1984; Nixon et al. 1984). In addition, a re-investigation of the diamondiferous 'kimberlite' at Prairie Creek, Arkansas, U.S.A., has shown more mineralogical affinities with W Australian lamproites than with kimberlite (Scott Smith & Skinner 1984). Together, all these studies are unanimous in recognizing distinctions between kimberlites and lamproites on the following grounds. 1 The lamproites contain glass. 2 The lamproite groundmass contains potassic richterite. 3 Compared with typical groundmass micas in kimberlites (Smith et al. 1978) the lamproite micas are richer in Ti, Fe and Na, but poorer in A1. 4 The groundmass diopsides in the lamproites have higher Ti contents than those in micaceous kimberlite (Dawson et al. 1977) (diopside from the Helam Mine lamprophyre resembles that in the lamproites). 5 Compared with groundmass spinels in kimberlites, those in Prairie Creek lamproite have lower A1203 and MgO contents and are relatively restricted in overall compositional range. 6 There is a paucity of garnet and magnesian ilmenite in the W Australian rocks compared with the S African kimberlites. 7 It is claimed that, compared with the average kimberlite analyses by Dawson (1967), the olivine lamproites of Kimberley are characterized by higher SiO2, TiO2, P2Os and particularly K20 contents, but by lower A1203 and FeO and considerably lower CaO and CO2 contents. 8 The mode of emplacement of the lamproites is different from the kimberlites in that some of the lamproite magma intruded to high level and apparently consolidated as lava lakes within craters. The mineralogical differences between kimberlites in general and leucite lamproites are indis-
97
putable, although it should be appreciated that phlogopite (high pressure) is chemically equivalent to the low-pressure combination of leucite and forsterite, and hence the difference could be due to pressure differences at the time of crystallization. Calcite is virtually absent in the lamproites where the main Ca-bearing species are richterite and diopside; diopside is present in Group II kimberlites but contains less TiO2 (typically less than 1~) (Dawson et al. 1977) compared with values generally in excess of 1~ in olivine lamproites (Jaques et al. 1984). The presence of relatively high Ti contents within the silicates (it is also high in the richterite and in groundmass phlogopite) and the relative lack of variety of ilmenite and rutile in the lamproites contrasts with kimberlites where the reverse is true. There are, however, some mineralogical similarities between olivine lamproites and Group II kimberlites. The presence of diopside (although with TiO2 differences) and olivine in two generations is common to both groups, as is abundant perovskite and apatite; moreover, although the rims of lamproite groundmass micas are higher in TiO2 and Fe 3 + but lower in A1203, the cores of these micas are compositionally very similar to the kimberlite groundmass Group II phlogopites (Smith et al. 1978). The main differences are the absence of calcite in the lamproites and the presence in them of glass and richterite; the latter features (together with the relatively quiet emplacement of the lamproite intrusions) are perhaps due to low COz contents in the original magma. However, superseding all these differences is the presence of diamond which gives this group of K-rich, low-volume melts the distinctive fingerprint of having originated in the diamond stability field of the upper mantle. Since the mineralogy is directly related to the bulk chemistry, in addition to the effects of changing oxygen fugacity, temperature and pressure during consolidation, it is also pertinent to examine the suggested differences in bulk composition between kimberlites and lamproites; these differences were recognized by Jaques et al. (1984) from a comparison of lamproite analyses with the average kimberlite analyses published by Dawson (1967). It is proposed here to compare the chemistry of several new analyses of highlymicaceous Group II kimberlites with that of the olivine lamproites since even superficial comparison indicates no overlap with the more siliceous leucite lamproites; it should be noted that the new analyses presented in Table 1 all show higher K20 contents than that of the average micaceous kimberlite (Dawson 1967) used by Jaques et al. (1984) in their comparisons. Table 1 shows the chemistry of Group II kimberlites, the sanidine-
98
J . B . Dawson
leucite lamprophyre from the Helam Mine and W Australian olivine lamproites. The Group II kimberlites are poorer overall in most oxides but are considerably enhanced in CaO and CO2 relative to the Australian rocks, and it is apparent that these other oxides are mainly 'diluted' by modal calcite. Table 2 shows the Group II kimberlites recalculated without CaCO 3 and H 2 0 - compared with an average olivine lamproite and a range of lamproite compositions recalculated on the same basis; oxide ratios of the type used by Jaques et al. (1984) in their comparative study are also given in Table 2 and it should be noted that only the A1203/CaO ratio is affected by the recalculation. In the case of SiO2 the kimberlites have only slightly lower absolute values than the lamproites and have very similar recalculated values. Similarly total iron and MgO are similar in both groups of rock, as are K 2 0 and P205. Furthermore MgO/SiO 2 and KEO/A1203, which were used by Jaques et al. (1984) to distinguish between the two groups, are very similar, and the Group II kimberlites have even higher K 2 0 / N a 2 0 ratios than the lamproites. As observed by Jaques et al. (1984) the lamproites have lower TiO2, Al203 (slightly), CaO and CO2 contents, but in other respects are similar.
Of the trace elements, the lamproites are considerably richer in Ba, Rb and Zr, richer in N b and Ni, and similar in Pb, Zn and Cu compared with the kimberlites; however, they have lower V, Cr, Sr and Y contents (Table 3). These values give rise to relatively high Rb/Sr, T i / N b and Zr/Nb ratios and relatively low K/ Rb, K/Ba and Ca/Sr ratios in olivine lamproite compared with Group II kimberlites; Ti/Zr ratios overlap for the two groups (Table 3). Scott Smith and Skinner (1984) used a N b / Z r plot to illustrate differences between the Prairie Creek and Australian lamproites and the S African kimberlites. Figure 2 is a similar plot incorporating new data from Jaques et al. (1984) and Atkinson et al. (1984) for lamproites, together with the Group II kimberlite data from Table 3. The olivine lamproites lie between the kimberlites and leucite lamproites on the Zr axis. The Group II kimberlites lie within the general field and trend of S African kimberlites in general ( N b / Z r = 2 ) but also lie very close to the position of the diamondiferous Argyle olivine lamproite. In the case of the rare-earth elements (REE) the pattern (light R E E enriched) and abundances for olivine lamproites overlap those for S African Group II kimberlites from Swartruggens (Mitchell & Brunfelt 1975) and the Bellsbank dykes (Fesq et al.
TABLE 1. Analyses o f Group H S African kimberlites compared with N W Australian olivine lamproites (wt,%) I
2
3
4
5
6
7
8
SiO 2 TiO2 A1203 Fe203 FeO MnO MgO CaO Na20 K20 H20 + H20P20 5 CO2 F SO3
36.53 1.58 4.72 1.94 3.86 0.61 11.81 17.40 0.25 3.98 0.91 0.38 0.48 13.3 n.d. 0.74
36.28 1.42 4.16 5.85 3.11 0.21 22.89 7.99 0.14 4.98 4.40 1.52 1.37 3.73 n.d. 0.03
39.26 1.75 4.88 5.59 3.27 0.42 19.81 6.30 0.14 4.37 4.58 2.54 0.71 4.18 0.01 0.09
38.66 1.72 4.89 4.16 4.58 0.16 29.82 5.49 0.17 5.24 4.00 3.18 1.61 2.77 0.03 0.60
37.53 1.50 4.10 3.23 4.46 0.13 23.74 6.33 0.16 4.27 5.46 2.43 1.22 3.42 0.02 0.37
42.82 1.46 5.23 6.42 3.62 0.15 18.70 8.34 1.10 1.58 4.59 4.60 0.75 0.34 NA NA
42.6 3.43 3.96 -8.30* 0.14 25.0 5.05 0.52 4.45 4.96 -1.27 0.18 ---
45.0 3.32 4.84 3.00 4.66 0.12 21.21 4.88 0.46 5.50 3.01 0.67 1.58 0.50 NA NA
Total
98.49
98.08
97.93
98.18
98.37
99.70
--
98.83
* Total iron expressed as FeO 1, BD 1088, Zout en Zuur dyke, Boshof District; 2, BD 1089, New Elands dyke, Boshof District; 3, BD 1102, Gordonia dyke, Theron District; 4, BD 1268, Main dyke, Helam Mine, Swartruggens, Transvaal; 5, BD 1267, Changehouse dyke, Helam Mine; 6, BD 1269, sanidine-leucite lamprophyre dyke ('Male' dyke), Helam Mine; 7, average of 21 Ellendale, W Australia, olivine lamproites (Jaques et al. 1984); 8, magmatic olivine lamproite, AK1 Argyle diamond mine, W Australia (Atkinson et al. 1984) (total includes 0.09 wt.~ BaO). Analyses 1-5 by X-ray fluorescence (analyst, R. Kanaris-Sotiriou) except FeO, H20, CO2 and F (wet methods; analyst, A. Saxby); analysis 6 by wet methods (analyst, J. R. Baldwin).
The kimberlite clan
99
TABLE 2. Analyses of Group H kimberlites and olivine lamproites calculated free of CaC03 and H20-
(wt.% ) 1
SiO2 TiO2 A1203 Fe203 FeO MnO MgO CaO Na20 KzO HzO § P205 SO3 KzO/A1203 AI203/CaO K20/Na20 MgO/K20 MgO/SiO2
2
53.80 2.32 6.95 2.86 5.68 0.90 17.40 0.71 0.37 5.86 1.34 0.70 1.09 0.8 9.8 15.8 3.0 0.32
3
41.20 1.61 4.72 6.64 3.53 0.24 25.99 3.68 0.16 5.66 4.98 1.55 0.03 1.2 1.3 35.4 4.6 0.63
4
43.86 1.95 5.45 6.24 3.65 0.47 22.13 1.12 0.16 4.88 5.12 0.80 0.10 0.9 4.9 30.5 4.5 0.50
5
43.59 1.94 5.51 4.69 5.16 0.18 23.48 2.22 0.30 5.90 4.51 1.81 0.67 1.1 2.5 19.7 4.0 0.54
42.57 1.70 4.65 3.66 5.06 0.16 26.93 2.25 0.18 4.84 6.19 1.38 0.42 1.0 2.1 26.9 5.6 0.63
6
7
42.8 3.44 3.97 -8.46 0.14 25.10 4.83 0.52 4.47 4.98 1.28 -1.12 0.82 8.5 5.6 0.6
40.71-42.61 2.71-5.77 3.30-4.51 3.80-4.94 2.37-4.98 0.11-0.14 19.04-26.90 4.06-5.61 0.31-0.63 3.46-5.11 3.74-4.93 0.60-1.48 0.9-1.3 0.6-1.1 8.1-12.2 3.7-7.5 0.4-0.7
Analyses 1-5 from Table 1 recalculated on a CaCO 3- and H20-free basis; analysis 6 is the average of 21 Ellendale (W Australia) olivine lamproites (recalculated free of CaCO3 and H 2 0 - and with all Fe as FeO); analysis 7 is a range of six olivine lamproite analyses (Jaques et al. 1984). 1975). H o w e v e r , the closest t r a c e - e l e m e n t similarities b e t w e e n G r o u p II k i m b e r l i t e s a n d olivine l a m p r o i t e s c a n be seen on an N d versus Sr isotope plot (Fig. 1). T h e i m p l i c a t i o n s f r o m this plot are that, like the l a m p r o i t e s , G r o u p II k i m b e r l i t e s are d e r i v e d f r o m a n c i e n t e n r i c h e d m a n t l e and, in this p a r t i c u l a r respect, h a v e m o r e affinities w i t h l a m p r o i t e s t h a n w i t h the G r o u p I kimberlites.
Relevance to upper-mantle metasomatism B o t h the k i m b e r l i t e s a n d the d i a m o n d i f e r o u s l a m p r o i t e s o r i g i n a t e at d e p t h s in excess o f 150 k m in the u p p e r m a n t l e a n d their differences in c h e m i s t r y reflect differences in t h e source regions.
TABLE 3. Trace-element analyses of micaceous kimberlites and N W Australian lamproites (ppm) 1
Cr Ni Ba Pb Zn Cu Rb Sr Y Zr Nb
144 1212 799 1445 47 115 30 142 647 25 189 100
K/Ba K/Rb Rb/Sr Ca/Sr Ti/Nb Ti/Zr Zr/Nb
22.9 223 0.219 192 94 50 1.9
V
2
118 1401 945 2192 26 82 50 185 1618 26 347 96 18.8 162 0.114 35 43 12 3.6
3
167 1631 812 3459 23 65 35 224 865 21 218 110 10.5 176 0.258 52 95 48 2.0
4
5
6
7
151 1266 768 3456 47 87 46 257 836 35 400 156
103 1265 857 1770 40 76 38 198 1133 27 283 107
72 1126 1024 9714 32 75 58 452 1102 16 784 175
19-142 528-1703 673-1500 3867-18281 20-50 61-90 52-93 300-611 959-1245 10-20 564-1214 118-244
13.1 179 0.307 47 57 22 2.6
20.0 159 0.175 40 82 31 2.6
3.4 0.41 31 116 26 4.5
1, BD 1088; 2, BD 1089; 3, BD 1102; 4, BD 1268; 5, BD 1267; 6, 7, mean and range of NW Australian olivine lamproites (Jaques et al. 1984; Table 5, analyses 1-6). The analyst for 1-5 was R. Kanaris-Sotiriou.
J. B. Dawson
I00 6O0
500
9
High-K kimberlites
[]
Kimberley kimberlites
z~pc Prairie Creek lamproites
400
o
Olivine lamproite
9
Leucite lamproite I N ' W
IA
] Australia
Argyle lamproite
E O. O. 300 ..Q Z 200
DD/
DDC'/
~176176I
t~
0I
~
/
O
Oo
OA
,...~,
~
f~pc
5 0I0
I
m
/
I 1000
1500
20100
Zr ppm
FIG. 2. Nb versus Zr plot for kimberlites and lamproites. Data sources: Kimberley kimberlites, Clement (1982) high-K kimberlites, new data (this paper); Argyle lamproite, Atkinson et al. (1984); Prairie Creek, Arkansas, lamproites, Scott Smith & Skinner (1984); NW Australian lamproites, Jaques et al. (1984).
Jaques et al. (1984) envisage a source region of harzburgitic composition (to account for the low AI and Ca in the lamproites) that has been subsequently subjected to K, Ti, Ba, Rb and Zr enrichment. In the case of the Group II kimberlites their source region must have contained some garnet (to account for the higher A1 and Cr relative to lamproites) and carbonate; although the K, Nb, Pb and Sr enrichment was similar to that of the lamproites, the kimberlite source region was not so enriched in Ti, Ba, Rb and Zr. Conversely, the relatively young enrichment event that is reflected in the Group I kimberlites resulted in the addition of less K, but higher Ti, to the Group I source region. All these differences highlight the fact that, although the same groups of elements are added during upper-mantle enrichment, there are specific differences from place to place in the upper mantle and, in the case of S Africa, at different times. The case in S Africa is particularly complex because, in addition to the old Group II and younger Group I enrichment events, the enrichment event that produced the veined and metasomatized xenoliths from the Bultfontein Mine (Kramers et al. 1983; Dawson 1987), although producing an isotopic signature identical with that of the Group I enrichment event (Fig. 1), took place in the amphibole stability field, i.e. at a considerably
shallower depth in the upper mantle than the kimberlite source region.
Conclusions
Recent bulk chemical and isotope data have reinforced the older petrographic distinctions between mica-rich (Group II) and mica-poor (Group I) kimberlites. The Group II kimberlites have closer chemical similarities to olivine lamproites than to Group I kimberlites. Nevertheless, all kimberlites and diamondiferous lamproites originate within the diamond stability field of the upper mantle. Their chemical and isotopic differences point to different degrees of initial melting followed by differing amounts and type of enrichment, at different geologic times, of the source mantle from which the different types of kimberlite and lamproite were derived. As such, they offer further convincing evidence of the temporal and spatial inhomogeneity of the upper mantle. ACKNOWLEDGMENTS: The kimberlite analyses were made by R. Kanaris-Sotiriou, A. Saxby and J. R. Baldwin. The figures were drawn by M. Cooper and the manuscript was typed by P. Mellor. I extend my thanks to all of them.
The kimberlite clan
I 01
References ATKINSON, W. J., HUGHES, F. E. & SMITH, C. B. 1984. A review of the kimberlitic rocks of Western Australia. In: KORNPROBST, J. (ed.) Kimberlites I." Kimberlites and Related Rocks, pp. 195-224. Elsevier, Amsterdam. CLEMENT, C. R. 1982. A comparative geological study of some major kimberlite pipes in the Northern Cape and Orange Free State. PhD Thesis, University of Cape Town (unpublished). , SKINNER, E. M. W. & SCOTT SMITH, B. H. 1984. Kimberlite re-defined. J. Geol. 92, 223-8. DAWSON, J. B. 1967. Geochemistry and origin of kimberlite. In: WYLLIE, P. J. (ed.) Ultramafic and Related Rocks, pp. 269-78. Wiley, New York. -1980. Kimberlites and their Xenoliths, 252 pp. Springer, Berlin. 1987. Metasomatized harzburgites in kimberlite and alkaline magmas: enriched restites and 'flushed' lherzolites. In: MENZIES, M. A. & HAWKESWORTH, C. J. (eds) Mantle Metasomatism. pp. 125-144 Academic Press, London. & HAWTHORNE, J. B. 1970. Intrusion features of some hypabyssal South African kimberlites. Bull Volcanol. 34, 740-57. , SMITH, J. V. & HERVIG, R. L. 1977. Late-stage diopside in kimberlite groundmass. Neues Jb. Mineral. Abh. 1977, 529-43. FESQ, H. W., KABLE, E. J. D. & GURNEY, J. J. 1975. Aspects of the geochemistry of kimberlites from the Premier Mine and other selected South African occurrences with particular reference to the rare earth elements. Phys. Chem. Earth, 9, 687-707. JAQUES, A. U, LEWIS, J. D., SMITH, C. B., GREGORY, G. P., FERGUSON, J., CHAPPELL, B. W. & MCCULLOCH, M. T. 1984. The diamond-bearing ultrapotassic (lamproitic) rocks of the West Kimberley region, Western Australia. In : KORNPROBST, J. (ed.) Kimberlites I: Kimberlites and Related Rocks. pp. 225-54. Elsevier, Amsterdam. KNOPF, D. 1970. Les Kimberlites et les Roches Apparentbes de C6te d'Ivoire, 202 pp. SODEMI, Abidjan. -
-
-
-
KRAMERS, J. D., RODDICK, J. C. M. & DAWSON, J. B. 1983. Trace element and isotope studies on veined, metasomatic and 'MARID' xenoliths from Bultfontein, South Africa. Earth planet. Sci. Lett. 65, 90-106. MCCULLOCH, M. T., JAQUES, A. L., NELSON, D. R. & LEWIS, J. D. 1983. Nd and Sr isotopes in kimberlites and lamproites from Western Australia: an enriched mantle origin. Nature, Lond. 302, 400-3. MITCHELL, R. H., 1981. Titaniferous phlogopites from the leucite lamproites of the West Kimberley area, Western Australia. Contrib. Mineral. Petrol. 76, 243-51. & BRUNFELT, A. O., 1975. Rare earth element geochemistry of kimberlite. Phys. Chem. Earth, 9, 671-86. NIXON, P. H., THIRLWALL, M. F., BUCKLEY, R. & DAVIES, C. J. 1984. Spanish and West Australian lamproites: aspects of whole rock geochemistry. In: KORNPROBST, J. (ed.) Kimberlite I: Kimberlites and Related Rocks, pp. 285-96. Elsevier, Amsterdam. SCOTTSMITH, B. H. & SKINNER, E. M. W. 1984. A new look at Prairie Creek, Arkansas. In : KORNPROBST, J. (ed.) Kimberlites I: Kimberlites and Related Rocks, pp. 255-83. Elsevier, Amsterdam. SKINNER, E. M. W. & SCOTT SMITH, B. H. 1979. Petrography, mineralogy and geochemistry of kimberlite and associated lamprophyre dykes near Swartruggens, western Transvaal, R.S.A. Extended Abstracts, De Beers Kimberlite Symp. 11, Cambridge. De Beers, Cambridge. SMITH, C. B. 1983. Pb, Sr, and Nd isotopic evidence for sources of southern African Cretaceous kimberlites. Nature, Lond. 304, 51-4. SMITH, J. V., BRENNESHOLTZ,R. & DAWSON, J. B. 1978. Chemistry of micas from kimberlites and xenoliths. I. Micaceous kimberlites. Geochim. cosmochim. Acta, 42, 959-71. WAGNER, P. A. 1914. The Diamond Fields of Southern Africa. 355 pp. Transvaal Leader, Johannesburg. -
-
J. B. DAWSON, Department of Geology, University of Sheffield, Mappin Street, Sheffield S1 3JD, U.K.
Lamproites and other potassium-rich igneous rocks: a review of their occurrence, mineralogy and geochemistry* Steven C. Bergman S U M M A RY : In this paper the geological occurrence, geochemistry and mineralogy of ultrapotassic (KzO/Na20 > 3 (molar ratio)) and perpotassic (K20/A1203 > 1 (molar ratio)) igneous rocks, especially lamproites, are reviewed and discussed in the context of compositionally-similar mantle-derived melts. Lamproites are K- and Mg-rich igneous rocks (typically K20 > 5 wt.%, MgO > 5 wt.%) which possess an exotic and diagnostic mineralogy and geochemistry. Known lamproites occur in 21 major suites or localities in continental regions with a variety of geological and tectonic environments; they range in age from the early Proterozoic dykes at Holsteinsborg, W Greenland, and Chelima, India, and Precambrian pipe at Argyle, W Australia, to the Middle Pleistocene flows and the Recent volcanics of the Leucite Hills, Wyoming, and Gaussberg, Antarctica, respectively. Intrusive and extrusive forms of lamproites include flows, a variety of pyroclastics (welded tufts, piperno, air-faU tufts, volcanic breccias etc.), cinder cones, dykes, sills and diatremes. Whereas kimberlite diatremes tend to be carrot shaped, the shape of olivine lamproite diatremes approximates a sherbet-glass. The recent discovery of diamondiferous lamproites of large volumetric proportion in the E and W Kimberleys, NW Australia, and the reclassification of the diamondiferous micaceous peridotite at Prairie Creek, Arkansas, as a lamproite substantiate their economic importance. The 21 lamproite suites considered here tend to be localized marginal to continental craton Cores in areas that overlie fossil Benioff zones, in contrast with the general occurrence of kimberlites more interior to continental cratons. The petrographic diversity of lamproites has historically hindered the development of a concise and universal classification and nomenclature. Lamproites are distinguished from kimberlites and alkali basalts (and lamprophyres) in terms of mineralogy, mineral chemistry, geochemistry and volcanic extrusive character. Relative to kimberlites, lamproites are enriched in K, Si, Ti, AI, Rb, Sr, Zr and Ba and depleted in CO2, Ca, Mg, Fe, Ni, Co and Cr. Lamproites are characterized by the general presence of phlogopite, diopside, leucite and K-richterite, occasional glass, olivine, sanidine, priderite, perovskite, wadeite, apatite and chrome spinel, and very rare ilmenite. Lamproite amphiboles, diopsides and phlogopites are distinctly depleted in A1~O3 relative to those of nearly all other igneous rocks. Lamproite magmas are produced by the partial melting of old refractory mantle peridotite (approaching a dunite or harzburgite in mineralogy) that was enriched in K-bearing and other incompatibleelement-enriched phases, such as phlogopite and apatite, most probably as a result of some metasomatic event which occurred prior to melting. In contrast with alkali basalt and kimberlite melts which are apparently produced from the partial melting of a CO2-enriched mantle peridotite (i.e. a source with a relatively high CO~/H20 ratio), water is the key volatile species involved with lamproite petrogenesis (source with a low COz/H20 ratio).
Introduction 'Lamproite' designates a K-rich mafic to ultramafic alkaline igneous rock as originally defined by Paul Niggli and P. J. Beger in 1923 in the context of Niggli's detailed chemical system of igneous rock classification. Relative to nearly all other igneous rock types, they possess high K / N a and K/A1 ratios. In addition, lamproite mineralogy and geochemistry are unparalleled. Niggli and Beger recognized two subtypes of the lamproite clan, based on the distinctive compo* This paper is dedicated to the memory of the late Th. G. Sahama who contributed much to the infant stages of the lamproite revolution.
sitions of rocks from the Leucite Hills, W y o m i n g , and SE Spain, a n d included fortunite, verite, orendite, prowersite, wyomingite a n d j u m i l l i t e in the clan. Tr6ger (1935) redefined the term 'lamproite' as the extrusive equivalent of lamprophyres that are rich in K and Mg. In the ensuing five decades, over 150 papers have been published on the subject of lamproites and an exceedingly complex terminology has been generated. Whereas the earliest studies e m p h a s i z e d the two classic localities in Spain and W y o m i n g , it is the purpose of this paper to summarize, compare and contrast the geological, geochemical and mineralogical characteristics of 21 lamproite occurrences on six continents and place these lamproite features in the context of compositionally allied
From: FITTON, J. G. & UPTON, B. G. J. (eds), 1987, Alkaline Igneous Rocks, Geological Society Special Publication No. 30, pp. 103-190.
IO3
Io4
S. C. Bergman
alkaline igneous rocks, including other ultrapotassic rocks, kimberlites, lamprophyres and alkali basalts. While not all K-rich rocks will be reviewed, several of the more important suites, geographically significant localities (because of close proximity to lamproites) or localities which have been previously described as lamproites but are not considered so in the present discussion will be included in this discussion. This paper is not meant to be an exhaustive review of K-rich rocks, but rather a springboard to the K-rich rock literature. Prior to the late 1970s lamproites were thought to represent petrological oddities of limited extent and importance with extreme and exotic mineral assemblages and chemical compositions (e.g. Turner & Verhoogen 1960). However, the discovery of diamond-bearing lamproites at Ellendale and Argyle, on the SW and E margins of the Kimberley craton in N W Australia (Atkinson et al. 1984a, b; Jaques et al. 1984), and the reclassification of the diamond-bearing micaceous peridotite at Prairie Creek, Murfreesboro, Arkansas, from a kimberlite (albeit anomalous) (Miser & Ross 1923; Bolivar 1977; Lewis 1977; Meyer et al. 1977) to a lamproite (Bolivar 1984; Scott-Smith & Skinner 1984a, b) have placed the lamproite clan in an important position in terms of both petrogenesis and economics. Not only must diamond-exploration models be revised to include lamproites but, in addition to kimberlites, certain members of a second group of igneous rocks must originate from within the diamond stability field in the Earth's upper mantle (depths of greater than 150 km) and be explosively transported to the surface in order to preserve diamonds that are unstable in these oxidized magmas. The diamondiferous sandy tuff lamproite at Argyle, W Australia, is among the largest (125 acres and more than 100 million tons) and is the highest grade (about 600 ct per 100 ton) igneous diamond deposit thus far discovered, dwarfing many of the S African kimberlite pipes (typically 20-80 ct per 100 tons in grade and 2060 acres in area) in the quantity of diamonds contained. This paper is organized as follows. After a brief consideration of the igneous geochemistry of K, Na and A1, previous reviews and benchmark papers on lamproites and other ultrapotassic igneous rocks will be briefly summarized. The nomenclature and classification oflamproites will then be discussed, followed by a review of the geological occurrence oflamproite and significant compositionally related suites. The major-, traceelement and isotope geochemistry of lamproites and their mineral chemistry will then be summarized and compared with those of compositionally
related rock types. Finally, after a review of experimental data involving lamproites, current theories on the petrogenesis of lamproites will be presented. It should be noted that, while the literature is presented in an objective forum, many of the author's opinions, which are not necessarily shared by all workers in this constantly debated arena of alkaline rocks, have infiltrated the review, especially in the discussion section.
K, Na and AI in igneous systems Important rock-forming minerals containing substantial proportions of both K and A1 include sanidine, biotite, leucite and rare kalsilite; the molar ratio K/A1 is nearly always less than unity for these minerals. Relative to both A1 and Na, K displays a remarkable regularity in its general distribution and abundance in igneous rocks (Fig. 1). In a bulk-rock sense, K is more incompatible than A1 and Na in basaltic and more differentiated igneous systems subjected to processes involving either partial melting or fractional crystallization. For example, both K/A1 and K/ Na increase in progressing from the 'bulk Earth' (carbonaceous chondrites) or the 'primitive upper mantle' to basalts (either alkaline or tholeiitic) as the result of partial melting, possibly combined with a small amount of fractionation (Fig. 1). These ratios both increase in the generalized sequence basalt, andesite, dacite, each progressively increasing in the degree of fractional crystallization that the given magma has experienced (Fig. 1). For the vast majority of igneous rocks, KzO/ N a 2 0 molar ratios are less than 1.0. As has been suggested by Johannsen (1931, 1933, 1935, 1939), those rocks with KzO/Na20 (molar) ratios of 13 are best termed 'potassic'. Relatively common rocks with K 2 0 / N a 2 0 ratios of about 1-3 include some granites, rhyolites, syenites, trachytes, latites, leucite tephrites, leucite basanites and minettes. Rarer rock types with a 'potassic' character include shonkinites and absarokites. Those rocks with KzO/Na20 ratios in excess of 3 should be termed 'ultrapotassic' and include some minettes, leucite phonolites, rare alkaline rocks (e.g. juvites), kimberlites and lamproites. To describe the relative amounts of the alkalis and A1, Shand (1927) used the term 'peralkaline' for those rocks with the sum of molecular K 2 0 and Na20 in excess of molecular A1203. The numerical value of the ratio (K20 + NazO)/A1203 has been named the agpaitic index or coefficient by Ussing (1912) and has been discussed by Polanski (1949) and Sorensen (1960). Peralkaline
Lamproites and K-rich igneous rocks ........
~g~[,:rE[ r " " 1
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!
K20 ! Na20 (~rr RATIO)
FIG. 1. (a) Plot of the covariation of K20/A1203 and K 2 0 / N a 2 0 for average igneous rocks and carbonaceous chondrites (Nockolds 1954; Goles 1967; Mason 1967, 1979; LeMaitre 1976) and a few sediments (Turekian & Wedepoh11961) compared with the estimated compositions of various portions of the Earth (continental crust (Clark & Washington 1924; Goldschmidt 1933; Poldervaart 1955; Ronov & Yaroshevsky 1969), primitive mantle (Taylor 1979), and lower, upper and Archaean crust (Taylor & McLennan 1981)). Note that Earth materials show a regular and positive correlation between the degree of'perpotassic' and 'ultrapotassic' character. (b) Same plo t as (a) except that the scale is reduced to include the average compositions of lamproites and related rocks (see text); the shaded region corresponds to the field of Earth materials in (a). Three sedimentary rock compositions are also shown. Lamproites are unparalleled in their degree of K enrichment relative to A1 and Na. See Appendix 3 for rock suite abbreviations. rock types include some rhyolites, trachytes, syenites, granites, minettes, lamproites and kimberlites. Most lamproites and some kimberlites (e.g. the Group II kimberlites of S Africa (Smith 1983a, b) and other m i c a c e o u s kimberlites) are a m o n g the only rock types that are so enriched in K 2 0 that it is present in excess of A1203 (atomic basis). J o h a n n s e n (1931, 1933, 1935, 1939) sug-
gested that the term 'perpotassic' be used to describe these exotic rocks. Therefore most lamproites and some mica-rich kimberlites and minettes share the unique quality of being a m o n g the only rock types that are both ultrapotassic and perpotassic. Those rocks whose elevated K contents are due to near-surface or late-stage m a g m a t i c alteration (such as the potassic altera-
IO6
S . C. B e r g m a n
and contrasts lamproites with lamprophyres and Dawson (1987) compares olivine-rich lamproites with micaceous kimberlites. Foley (1985) considers the oxidation state of lamproite magmas. 'Shoshonites' (i.e. those andesitic intrusive and Previous reviews extrusive rocks with K 2 0 / N a 2 0 near or greater Lamproites and other K-rich alkaline rocks have than 1.0, with slight ~enrichments in CaO and MgO and depletions in FeO and TiO 2 relative to been the subject of many important reviews and contributions, although no-one has attempted a NazO-rich alkali basaltic rocks) and related comprehensive integrated synthesis of tarn~- potassic_ro~s,_~vhich nre minor but ubiquitous components of many island arcs and continental proites. Various aspects of selected potassic and ultrapotassic rock suites have been discussed in margins, are much more aluminous than lammany benchmark papers and books on alkaline proites and have received attention in review and other igneous rocks (Jensen 1908; Daly 1910, papers by Joplin (1965, 1968) and Morrison 1914, 1933; Niggli 1923; Shand 1927; Smyth (1980). Recent contributions to the petrogenesis 1927; Terzaghi 1935; Saether 1950; Turner & of shoshonite magmas include Gest & McBirney Verhoogen 1960; Sobolev 1970; Bardet 1973, (1977), Boccaletti et al. (1978), Pagel & Leterrier 1974, 1977; Carmichaeletal. 1974; Stewart 1979; (1980), Ewart (1982), Keller (1983) and Friend & Janardhan (1984), and references cited therein. Smith 1984b; Venturelli et al. 1985; Foley et al., K-rich rocks occur as local variations within in review). Niggli (1923, pp. 182-5) and Beger (1923, pp. 448-51) were the first to synthesize the various portions of alkaline pyroxenite-bearing ultramafic plutons and gneiss complexes in the early 1920s understanding of ultrapotassic rocks. Their two subtypes of lamproites were character- Oman Mountains, Arabia (Searle 1984), at ized by the following Niggli values (Niggli values Ampipal in the central Nepal Himalayas (Lasare the ratio of the weight per cent oxide to the serre 1976), at Shaki and Okeho in the basement gram molecular weight, normalized to one cation complex of W Nigeria (Oyawoye 1976), in the Batbjerg complex in E Greenland (Gittins et al. • 1000; k = k/(na + k), mg = mg/(mg +fe)). 1980; Brooks et al. 1981), in the Blue Mountain Subtype si al fm c alk k mg c/jm complex, New Zealand (Grapes 1975), in S Murcia 140 16 55 13 16 0.80 0.80 0.23 Greenland (Upton & Thomas 1973) and in the Wyoming 165 18 41 14 27 0.85 0.75 0.35 Mordor complex in central Australia (LangworLacroix (1926) summarized the systematics of thy & Black 1978) among others. The petrogenetic syenitic leucite-bearing rocks including those of relationship between these rocks and lamproites the Leucite Hills and Gaussberg. On the basis of must await more detailed study. Many alkalic-carbonatitic complexes possess petrography and geochemistry, Sahama (1974) differentiated orenditic from kamafugitic rocks fenitized zones which are characterized by rocks in a review of K-rich igneous rocks; the former composed almost entirely of potassium feldspar (Heinrich 1966; McKie 1966; Verwoerd 1966; group corresponds to the lamproites discussed Heinrich & Moore 1970). Whereas these rocks here. A recent book reviewing mineralogical, petrological, geochemical and experimental as- possess some geochemical and mineralogical features common to lamproites, they tend to be pects of rocks that contain modal leucite (Gupta & Yagi 1980) includes a discussion of many of extremely depleted in MgO and it is clear that the lamproite suites synthesized and updated in their petrogenesis and paragenesis are distinct the present work. Wendlandt & Eggler (1980a) from those of lamproites. performed a statistical study of 835 leucitenormative Cenozoic volcanic rocks and recognized two compositionally-distinct sub-groups, one enriched in K20 and A1203 but depleted in Lamproite nomenclature and MgO and another enriched in MgO but depleted classification in K20 and A120 3. Interestingly, lamproites apparently did not fall in either group and were Lamproite terminology has historically been one evidently not included in the original data set of the most complex and disjointed of that for any group of igneous rocks (Table 1). In the first (RKNFSYS; Chayes 1976). Mitchell (1985) half of this century, a nomenclature was develcontributed a comprehensive review of the oped that centred around the mineralogy of mineralogy and mineral chemistry of lamproites individual lamproite suites with little regard to and Wagner & Velde (1986a) discussed the lamproite suites elsewhere in the world. Table 1 mineral chemistry of 10 K-richterite-bearing summarizes the sometimes redundant terms that lamproites from six suites. Rock (1987) compares
tion zones in some fenites and porphyry systems) are excluded from the present discussion.
Lamproites and K-rich igneous rocks
Io7
TABLE 1. Historic summary of lamproite terminology Rock term
Principal mineral and/or phenocryst GGM assemblage opx ol amph cpx phi leuc san
Wyomingite Orendite Madupite Cedricite Mamilite Wolgidite Fitzroyite Verite Jumillite Cancarixite Fortunite Cocite Kajanite Gaussbergite
x
x x x x
x x
x
x x x x x
x x x x x x x x
x x x x
x
x x
x)
x
Rock suite
Leucite Hills, U.S.A.
Original reference
Cross 1897
•
x x x x x x
x
Fitzroy Basin, W. Australia Wade & Prider 1940
•
x ~ Murcia-Almeria, SE Spain ) Murcia-Almeria, SE Spain Murcia-Almeria, SE Spain Coc-Pia, N Vietnam Oele Kajan, Borneo
x
x
x
x x x x
x
x
x
Gaussberg, Antarctica
Osann 1899 Parga-Ponda11935 De Yarza 1893 Lacroix 1933b Brouwer 1909; Lacroix 1926 Lacroix 1926
opx, orthopyroxene; ol, olivine; amph, amphibole; cpx, clinopyroxene; phi, phlogopite; leuc, leucite; san, sanidine; GGM, glassy groundmass.
have been introduced for mineral assemblages characteristic of individual or several lamproite suites. Note, for example, that fitzroyite and wyomingite are equivalent, and cocite and jumilite are mineralogically identical. Unfortunately, the most recent report of the IUGS Subcommission on the Systematics of Igneous Rocks, which specifically emphasized the classification and nomenclature of lamprophyres, carbonatites and other alkaline rocks (Streckeisen 1980), neglected to include lamproites in its discussion. Under the earlier IUGS scheme (Streckeisen 1967), lamproites would fall in the compositional fields of alkali trachyte, phonolite, phonolite foidite or leucitite. Tr6ger (1935) and Johannsen (1931, 1933, 1935, 1939) describe selagites from Italy which some workers (D. Velde, personal communication, 1985) place in the lamproite clan. However, selagites contain about 10~oplagioclase (oligoclase) and therefore do not meet the mineralogical criteria for lamproites discussed in more detail below. Although Middlemost (1975) included lamproites in the basalt clan, lamproites should not be considered as basalts (despite the appearance of basaltic cinder cones and lava flows in many lamproite occurrences such as the Leucite Hills). Lamproites do not contain plagioclase, a phase which is generally regarded as an essential component of basalts. Note, however, that Zirkel (1870) and Rosenbusch (1877) used the term 'leucite basalt' for rocks composed of leucite, augite and olivine which were devoid of plagio-
clase. If olivine lamproites are to be regarded as basalts, kimberlites must also be included in the basalt clan. Lamproites have historically been split up and placed in a variety of sometimes distinct rock groups; these classifications, furthermore, have hindered the development of a clear understanding of the lamproite clan and its petrogenesis. Rosenbusch (1887) classed the Prairie Creek, Arkansas, lamproite with biotite-peridotites but the Leucite Hills rocks with leucite basalts. Niggli (1923) introduced the term 'Mediterranean province type' for potassic alkaline rocks of syenitic to shonkinitic composition; the type locality was the Roman co-magmatic province whose rocks are not lamproites because of their metaluminous to peraluminous character. Nevertheless, some lamproite suites have been included in the Mediterranean magma type (e.g. Turner & Verhoogen 1960). Lacroix (1933) put lamproites in his Division II, Family A: missourites, leucitites and albanites. Tr6ger (1935) placed the Prairie Creek, Arkansas, lamproite in his glimmerite Family (no. 721, olivine-glimmerite). Other lamproites fall into Tr6ger's nepheline syenite (orendite) and shonkinite (cocite, wyomingite, gaussbergite, jumillite) families. Johannsen (1931, 1933, 1935, 1939) detailed the mineralogy and texture of many of the lamproites and ultrapotassic rocks known at that time and placed lamproites and other ultrapotassic rocks in six different families within three classes of his rock-classification scheme. These families
IO8
S. C. B e r g m a n
spanned a wide range of compositions, including feldspathoidal rocks, peridotites, trachytes, perknites and alkali syenites. Williams et al. (1954) included Leucite Hills lamproites in the leucite phonolite petrographic group. Rittmann (1951), who considered both mineralogy and geochemistry in devising his nomenclature for volcanic rocks, placed lamproites in the 'lamproitic phonolite', 'lamproitic leucitite' or'lamproite trachyte' groups. He additionally suggested different terms for established lamproite rock names, i.e. madupite~--pheno-phlogopite-mafitite~-~-lamproite leucitite,verite~pheno-mafitite(phlogopite)~lamproitic trachyte, and wyomingite~pheno-leucitemafitite----lamproitic phonolite. Sahama (1974) classified K-rich alkaline rocks into two groups, lamproitic (or orenditic) and kamafugitic, on the basis of chemistry and petrography. Barton (1979) recognized three sub-groups of K-rich alkaline rocks on the basis of bulk rock and mineral chemistry and petrography: the Leucite Hills type, the Toro-Ankole type and the Roman province type. The 21 lamproite suites considered below fall in both the Leucite Hills and Roman province fields in Barton's scheme; none, however, overlaps the Toro-Ankole field. On the basis of petrographic modal and textural variability, Mitchell (1985) suggested that two broad subdivisions of lamproites be made: phlogopitesanidine-leucite-diopside-lamproites and madupitic lamproites. The former group possesses resorbed phenocrystal phlogopite and includes wyomingites, orendites, fitzroyites, cedricites, verites etc. The latter group is characterized by poikilitic groundmass phlogopite and includes madupites, wolgidites and jumillites. As many lamproites possess an extremely finegrained or glassy groundmass, classification based on mineral abundances alone is doomed to failure because of the uncertainty in the mineral equivalent of the groundmass. Since bulk-rock composition reflects mineral chemistry and abundance of component phases (and vice versa), whole-rock geochemistry can also be used in establishing a classification system for lamproites. This is extremely useful for fine-grained rocks and is essential for glass-rich rocks. Nonethe-less, it is often difficult to pigeon-hole rocks solely on the basis of their geochemistry. ScottSmith & Skinner (1984b) suggested classifying lamproites on the basis of the modal abundance of principal primary minerals, following an analogous scheme proposed by Skinner & Clement (1979) and Clement et al. (1977) for kimberlites. Their six major divisions were phlogopite, K-richterite, olivine, diopside, sanidine and glassy lamproites. Mitchell (1985) has followed their suggestion but has criticized the elimination
of leucite from the list of major phases and the inclusion of glass in a mineralogical classification, two criticisms that are taken into consideration below. Therefore leucite is substituted for glass in the six major sub-divisions proposed by ScottSmith & Skinner (1984b). Minerals of secondary importance can be used as modifiers to the most abundant phases and permit further sub-divisions to these six major classes. Mitchell (1985) also recommended recognizing the texture of phlogopite in the classification, and suggested the following sub-divisions: phlogopite lamproites (in which phlogopite occurs as phenocrysts) and madupitic lamproites (in which phlogopite is poikilitic in the groundmass). Mitchell's mineralogical classification removes many of the redundancies of the classical terminology, and tends to over-simplify the mineralogy of a given rock. Nevertheless, until more is known about the genetic controls on the petrography of lamproites, this classfication scheme is the best alternative.
Lamproite defined The definition of lamproite proposed and used here is based on those of Scott-Smith & Skinner (1984b) and Mitchell (1985) but includes several other parameters. It is suggested that the term 'lamproite' be applied to those intrusive and/or (more commonly) extrusive rocks that fulfil the following mineralogical and/or geochemical constraints. Mineralogical conditions 1 Dominant (more than 5~-30%), although not essential, primary phases (groundmass or phenocryst) include titanian phlogopite (Al-poor), clinopyroxene (Al-poor diopside, more rarely diopsidic augite), Al-poor alkali amphibole (commonly K-Ti-richterite, more rarely K-riebeckite, K-arfvedsonite or other alkali amphiboles), olivine (often in aggregated grains but also euhedral), leucite (Na-poor, commonly Fe-rich) and sanidine (commonly Fe-rich). 2 Accessory (less than 5%-10%) but often characteristic primary phases include spinel (intermediate Cr-Mg-Fe-A1-Ti compositions), priderite ((K,Ba)(Ti,Fe 3+ )8016 ), wadeite (K,~Zr2Si6018), shcherbakovite ((Na,K) (Ba,K) Ti 2Si4014 ), jeppeite ((K,Ba)2 (Ti,Fe)6013 ), apatite, perovskite, sphene, armalcolite ((Mg,Fe)Ti4010), enstatite (rare) and ilmenite (very rare). 3 Alteration or secondary phases comprise analcime, chlorite, quartz, TiO 2 polymorphs,
Lamproites and K-rich igneous rocks carbonate minerals, nontronite (and other zeolites), chlorite, serpentine, barite (Sr-rich), albite and a variety of clay minerals. 4 If a rock carries any of the following, it falls outside the lamproite sensu stricto group: (a) primary plagioclase (late- to post-magmatic secondary albite can occur); (b) melilite or monticellite (places a rock in the more calcic kamafugitic, ultramafic, lamprophyre, kimberlite or basalt clans); (c) kalsilite (requires a SiO2 activity much lower than that observed in lamproites); (d) nepheline, sodalite, nosean and hafiyne (these sodic phases are not compatible with an ultrapotassic assemblage); (e) melanite (despite the TiOz-rich and Al-poor composition oflamproites, melanite contains too much Fe to be stable). 5 Xenoliths and xenocrysts are neither necessary nor sufficient to the definition of lamproites but are mentioned to permit further distinction of lamproites from kimberlites and ultramafic to mafic lamprophyres. Cr-rich pyrope xenocrysts, which are ubiquitous components of kimberlites, are rare components of olivine lamproites (e.g. Ellendale, Argyle, Prairie Creek). 'Dogtooth olivine' aggregated monomineralic xenoliths are a characteristic feature of many olivine lamproites (e.g. Ellendale, Prairie Creek (Scott-Smith and Skinner 1984a, b)) and are distinct from the ubiquitous rounded olivine macrocysts of kimberlites. Spinel and garnet lherzolite xenoliths are extremely rare in lamproites but common in kimberlites. The most common xenoliths found in lamproites (apart from ubiquitous shallowcrustal country-rock fragments) are biotitite, biotite-clinopyroxenite and pyroxenite xenoliths that are most probably cognate. 6 Mineral chemistry (see below) can also be used to distinguish between lamproites and other rocks.
10 9
Occurrences of lamproites and other potassic to ultrapotassic rocks This section summarizes geological and petrological data and literature pertinent to those localities which contain rocks that conform to the constraints discussed above and hence can be called lamproite sensu stricto. Localities containing potassic-ultrapotassic rocks which do not meet the requirements for classification as lamproites will also be considered, and the rationale for not including them in the lamproite group will be given. Tables 2 and 3 abstract the geological parameters of each locality and most of these are plotted in Fig. 2. Note that the abbreviations used to reference each locality will be used throughout this paper and, in particular, that those localities not recognized as containing lamproites sensu stricto are set in square brackets. An alphabetical listing of the abbreviations is given in Appendix 3. Figure 3 illustrates the variety of rock textures observed in the suites from Australia and N America, and Fig. 4 shows representative thin-section photomicrographs from 11 lamproite localities. N America and Greenland
Kimberlites, lamproites, carbonatites and allied rocks are widespread in N America (Fig. 5), however, in contrast with kimberlites and carbonatites, lamproites do not occur in the interior of the craton but are generally located near its margin (Meyer 1976; Turtle & Gittins 1966). The estimated age of the crustal basement to these lamproite localities varies from 1.5 x 103 to 2.5 • 103 Ma. Enoree Vermiculite district, S Carolina ( E V D ) (34 ~ 30' N, 82 ~ W)
Geochemical conditions
When a rock possesses the following general compositional features, it is suggested that it be included in the lamproite clan: K20/A1203 > 0.8 (molar), K 2 0 / N a 2 0 > 4 (molar), Mg number> 70, with general compositional ranges of 45-55 wt.~ SiO2, 4-10 wt.~ A12Oa, 1-5 wt.~ TiO2, 210 wt.% CaO, 5-10 wt.~ K20, 0.2-1.5 wt.~ Na20, 0.5-2.0 wt.% P205 and 1.0-3.0 wt.% BaO. It is recommended that both mineralogy and geochemistry be used to distinguish lamproites from other rocks. However, where mineral data are lacking or impossible to acquire (glassy rocks), these compositions can be used with caution.
Late Proterozoic to Early Palaeozoic vermiculite bodies occur throughout the Appalachians and Piedmont and are best exposed and studied near Enoree, SC, just W of the Kings Mountain Belt (Stewart 1949; McClure 1963; Libby 1975). Palaeozoic metamorphism produced a mineral assemblage of phlogopite, diopside, tremolite (or talc), K-feldspar, apatite, sphene, monazite and zircon. Metamorphism has clearly modified the primary mineralogy and only a few of the major lamproite phases are present. The geochemical character of the EVD rocks, provided that it has not been modified by the metamorphism, is diagnostically lamproitic and nearly identical with the geochemistry of the average lamproite (Tables 4 and 5).
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FIG. 2. World map showing the locations of lamproite sensu stricto suites and some of the compositionally-similar potassic-ultrapotassic rock suites. For a more complete listing, see Tables 2 and 3 and text. Hills Pond, Kansas ( H P ) (37 ~ 40' N, 95 ~ 45' W)
Cretaceous (88 Ma (Zartman et al. 1967)) lamproite plutons, sills and dykes intrude Precambrian granites of the Rose Dome and nearly flatlying Pennsylvanian shales and limestones near Hills Pond, Woodson County, SE Kansas (Twenhofel & Bremer 1928; Knight & Landes 1932) (Fig. 6). About 13 km SW of the Rose Dome, a 'mica peridotite' (madupitic lamproite (P. Berendsen, personal communication, 1984)) of approximately the same age (90-91 Ma (Zartman et al. 1967)) intrudes the Silver City Dome (Bickford et al. 1971; Franks 1966). The HP rocks possess all the earmarks of olivine lamproites and consist of Ti-phlogopite, serpentinized olivine, K-richterite and Ti-diopside in a fine-grained groundmass of serpentine, perovskite, apatite and chrome spinel (Merrill et al. 1977) (Fig. 4). Other petrological and geological contributions concerning the HP intrusives and their xenoliths have been made by Wagner (1954), Franks (1959, 1966), Franks et al. (1971) and Cullers et al. (1985). Fresh rocks have been recovered by drilling; the surface exposure of the HP body has been intensely altered to form a vermiculite body. The petrography of madupitic (K-richterite-diopside) lamproites from the Silver City Dome area are discussed in Berendsen et al. (1985). Other lamproite diatremes occur in the area (H. Coopersmith, personal communication, 1985).
The geochemistry of the HP olivine lamproite is nearly identical to those of the diamondiferous olivine lamproites from Ellendale, Argyle and Prairie Creek. Evidently, the quiescent intrusive character of the Hills Pond magmas, manifested by the preponderance of thin sills, was not conducive to the preservation of diamonds, assuming the magmas were derived from a similar depth and source rock as the three diamondiferous suites mentioned. The HP area is one of the rare occurrences in which lamproite forms welldeveloped sills and the only known occurrence of olivine-lamproite in well-developed sills. Holsteinsborg, W Greenland ( H O L ) (66 ~ 45' N, 53 ~ IV)
Proterozoic lamproite (1200 Ma) and post-tectonic kimberlite (585 Ma) dykes (less than 1-2 m wide) intrude metamorphic rocks 1650-1740 Ma old in the southern part of the Nagssugtoquidian mobile belt on the coast of central W Greenland (Noe-Nygaard & Rambert 1961; Escher & Watterson 1973; Scott 1977, 1979, 1981; Brooks et al. 1978; Thy 1985) (Fig. 7). The lamproites contain phenocrysts of olivine, diopside, Krichterite, Ti-phlogopite and pseudoleucite in a groundmass of Ti-phlogopite, diopside, K-richterite, sanidine, pyrite, a carbonate mineral and priderite (Fig. 4), whereas the kimberlites contain macrocrysts of olivine, phlogopite, picroilmenite and rare pyrope in a groundmass of olivine, Ti-
Lamproites and K-rich igneous rocks
I 15
(a)
(c)
(b)
(e)
(d) FIG. 3. Photographs of lamproite slabs showing typical rock textures (all listed clockwise from top left): (a) olivine lamproites from the Prairie Creek area (PRA) (American Mine, Kimberlite mine, Crater of Diamonds pipe (three samples) and Black Lick); (b) phlogopite lamproites from the Prairie Creek, Crater of Diamonds diatreme (PRA), all from the West Hill, showing tufts and magmatic--autolithic breccias; (c) Fitzroy Basin (WKB) xenolithic/autolithic lamproite breccias (Mount Abbot, Big Spring, 81 mile vent, Mount Ibis, Calwynyardah Pipe and Fishery Hill; center; Ellendale pipe and Wolgidee Hill); (d) Fitzroy Basin (WKB) magmatic and volcanic rocks (Mount North phlogopite lamproite, vesicular 81 mile vent phlogopite lamproite, Ellendale pipe B olivine lamproite, vesicular 81 Mile Vent glassy phlogopite lamproite, altered near-surface Ellendale pipe B; center glassy Oscars Plug phlogopite lamproite); (e) Fitzroy Basin lamproites; (f) Leucite Hills lamproites.
(f)
(Scale: photograph a= 17 • 29 cm)
I I6
S . C . Bergman
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p (q) r) (s) t) FIG. 4. Photomicrographs (plane light) of various lamproites. Massive magmatic lamproites: (a) Wolgidee Hills wolgidite with perovskite, priderite, shcherbakovite, diopside, K-richterite in view; (b) Machells Pyramid Hill phlogopite lamproite; (c) Oscar Plug (WKB) glassy phlogopite lamproite; (d) Priestly Peak (PP) phlogopite lamproite; (e) Gaussberg (GSB) leucite lamproite; (f) Smoky Butte (SB) phlogopite lamproite; (g) Kamas (KAM) phlogopite lamproite; (h) Spring Butte (LH) phlogopite lamproite; (i) Holsteinsborg (5646) (HOL); (j) Chelima (CHE) phlogopite lamproite. Lamproite breccias: (k) Oscar Plug (WKB) lamproite-sandy tuff-breccia; (1) Boars Tusk (LH) phlogopite lamproite; (m) Argyle AK1 (ARG) sandy tuff; (n) Calwynyardah (WHB) olivine lamproite. Olivine lamproites: (o) Hills Pond (HP); (p) Kimberlite Mine (PRA); (q) Black Lick (PRA); (r) Crater of Diamonds (PRA); (s) Ellendale pipe B with dunite xenolith; (t) Ellendale pipe B with garnet xenocrysts. Scale: each photograph is approximately 4 x 5 mm. phlogopite, apatite, diopside, serpentine, a carbonate mineral, perovskite and spinel (Scott 1977, 1979, 1981; Brooks et al. 1978). So-called 'anomalous' lamprophyre dykes that are extremely rich in diopside with phlogopite, kaersutite and titanomagnetite also occur (Scott 1977, 1979, 1981). Thy (1982) reported an assemblage
of K-richterite-arfvedsonite-riebeckite-actinolite from related dykes. Lherzolite, harzburgite and granulite xenoliths occur in the kimberlite dykes (Scott 1979, 1981). The emplacement of the HOL lamproites is broadly contemporaneous with the emplacement of the Gardar igneous province ((1.1-1.3) x 10 3
Lamproites and K-rich igneous rocks
I 17
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9 ULTRAMAFIC LAMPROPHYRE o KIMBERLITE LOCALITY / PROVINCE t, LAMPROITE LOCALITY / PROVINCE 9 K-RICH BASALT/LAMPROPHYRE ULTRAPOTASSIC ROCKS
& OTHER
9 LAMPROPHYRE 9 CARBONATITE
FIG. 5. Map of N America showing the locations of lamproite s e n s u s t r i c t o suites, other K-rich rock suites, kimberlites, carbonatites and several notable lamprophyres. Principal geological and tectonic features, i.e. Pacific oceanic crust fracture zones, the New England seamount chain, the mid-continental rift zone, the Colorado lineament zone, the inner boundary of Phanerozoic orogenic belts, the limit of areas presumed to be underlain by Archaean cratonic material and major crustal boundaries (with the approximate basement age after Condie 1976) are also shown. For lamproite and K-rich rock suite abbreviations see Tables 2 and 3 and text. Precambrian crustal provinces: Wy, Wyoming; CN, Central; GR, Grenville; SU, Superior: KA, Kaminak; SL, Slave; CH, Churchill. Pacific fracture zones: ME, Mendicino; PI, Pioneer; MU, Murray; MO, Molokai; CL, Clarion. Other features: NES, New England seamounts, CL, Colorado lineament: MCRS, Mid-continent rift system; OA, Ouachita fold belt.
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I20
TABLW5. Comparison of the average major-element compositions (wt.~) of diamondiferous lamproites and Kimberlites with barren varieties Known diamondiferous
SiO2 TiO2 Al,,O3 FeO* MnO MgO CaO Na20 K20 P205 BaO COE H20 + Mg number n
Known non-diamondiferous
Lamproites
Kimberlites
Lamproites
Kimberlites
47___7 2.9_+ 1 4.8+1 8_+2 0.12_+0.04 23_+6 8+8 0.4-+ 0.3 3.5_+ 1 1.3_+0.7 0.9+0.8 4_+4 3.3_+2 82_+5 46
40+7 2.4_+ 1.7 5_+2 11 _+3 0.18_+0.1 28_+7 10_+8 0.6_+ 0.9 1.5_+ 1.1 0.9_+ 0.8 0.19_+0.2 4.0_+5 7_+4 81 _+6 325
53___6 3.1 ___2 10___2 6_+2 0.10_+0.05 10_+4 6-+3 1.6_+ 1 7.5_+3 1.3 _+0.7 0.6_+0.6 1.7_+2 3_+2 72 +_8 270
36+7 3.0-t-2 5_+4 12+2 0.20_+0.1 27_+8 13_+7 0.6_+ 1.2 1.4_+ 1 0.9_+0.8 0.16_+0.1 8.0_+5 5_+3 78 +_12 100
All oxides, except CO., and H20, are normalized to 100~o on a volatile-free basis to facilitate comparison (see also Dawson 1986). n, number of analyses. Ma (Blaxland 1976a, b)) 700-800 km to the S of HOL as well as a variety of alkaline complexes in Canada (Gittins et al. 1967; Currie et al. 1975), some of which contain pseudoleucite (Watkinson & Chao 1973), and the Kali 'kimberlitic' dyke in N Sweden (Kresten et al. 1977). Proterozoic ((1.01.2) x 103 Ma) potassic ultramafics in S Greenland (Upton & Thomas 1973) are similar in age and general geochemical character to the HOL rocks, but do not fit the definition of lamproite. It should be noted that diamonds have been recovered in kimberlites from the area between Ivigtut and Frederikshab (Andrews & Emeleus 1971) and in heavy mineral concentrates from stream sediments in the Fiskenaesset area (Nielsen 1976) about 500-700 km S of HOL. Over a dozen alkaline rock suites (carbonatites, lamprophyres and kimberlites) occur in a coastal belt S of HOL and range in age from 2.65 x 103 to 0.12x 103 Ma, indicating mantle-derived alkaline magmatic activity recurring in the same region for a time span in excess of half of Earth history (Larsen et al. 1983). Kamas and Moon Canyon, Utah ( K A M ) (40 ~ 40' N, 111 ~ 12' IV)
Middle Tertiary dykes, flows and small plugs of phlogopite-olivine lamproites are spatially associated with more syenitic dykes in Moon Canyon, near Kamas, Scott County, Utah (Fig. 8). They occur near the junction of the W margin of the E - W trending Uinta Arch and the N-S trending Wasatch Mountains (Morris 1953 ; Larsen 1954; Best et al. 1968; Henage 1972) just W of the
Colorado Plateau. Morris (1953) called the rocks melteigites and Best et al. (1968) termed them wyomingites and orendites. Phlogopite and olivine phenocrysts are set in a glassy groundmass of analcime, sanidine, phlogopite, clinopyroxene, K-richterite, ilmenite, pseudobrookite, chromite, a roedderite-like phase and apatite (Wagner & Velde 1986a) (Fig. 4). The dykes range in age from 13 to 40 Ma (Crittenden & Kistler 1966; Best et al. 1968). Henage (1972), on the basis of trace-element data on mineral separates and whole-rock dyke samples, established that the distribution of trace elements was heterogeneous as a result of both crystal separation and volatilephase transfer. According to Boutwell (1912), mica peridotite dykes were also intersected by underground mine workings in the nearby Park City mining district. The K A M dykes and flows are located 250 km SW of the LH lamproites. Leucite Hills, Wyoming ( L H ) (41 ~ 50' N, 109 ~ 10'
w) The 14 Leucite Hills (22 major exposures) consist of cinder cones, lava flows, volcanic necks, dykes, intrusive sheets and pyroclastic rocks which were intruded through and extruded over Tertiary shales, Cretaceous limestones and clastic sedimentary rocks on the northern nose of the Rock Springs uplift in the Green River Basin, Sweetwater County, Wyoming (Fig. 9). The LH are among the youngest lamproites (about 1 Ma old) (Bradley 1964; MacDowell 1966), second only to Gaussberg, and occur over an area of 2000-
L a m p r o i t e s a n d K - r i c h igneous r o c k s
x 21
FIG. 6. (a) Generalized geologicalmap of the Rose Dome Area, Kansas, showing the location and setting of the Hills Pond lamproite (after Jewett 1964); (b) geological map of the Hills Pond lamproite with cross sections AA' and BB' based on drilling (after Wagner 1954). 2500 km 2. Cross (1897), Kemp (1897) and Kemp & Knight (1903) generated the nomenclature for LH that has since been applied to several other lamproite suites (Exeter area dykes, U.K., Kamas, Utah, and Fitzroy Basin, W Australia). More recent geochemical and petrological work includes that of Johnston (1959), Smithson (1959), Yagi & Matsumoto (1966), Carmichael (1967),
Kay & Gast (1973), Ogden et al. (1978), Barton (1979), Ogden (1979), Kuehner (1980), Ogden et al. (1980), Barton & van Bergen (1981), Kuehner et al. (1981), Vollmer et al. (1981, 1984), and Salters & Barton (1985). Petrographic types (Cross 1897) include wyomingite (leucite-phlogopite-diopside lamproiteolivine_+ K-richterite), orendite (leucite-sani-
I22
S . C. [ ] PROTEROZOIC GNEISSES & METASEDIMENTARY ROCKS []
ARCHAEAN GNEISSES PALAEOZOIC KIMBERLITE 9 PROTEROZOIC LAMPROITE FAULTS DOME 0 100 km
Bergman Crater o f Diamonds State Park and associated intrusives, Prairie Creek, Murfreesboro, Arkansas ( P R A ) (34 ~ 2' N, 93 ~ 40' W)
Although the diamondiferous Prairie Creek intrusions in the vicinity of the Crater of Diamonds L (COD) Park, Pike County, Arkansas, were previously thought to possess affinities with 56* 52* kimberlites (Miser & Ross 1923; Scull 1959; -" ' 0' ::::::::::::::::::::::: 48~ Meyer 1976; Bolivar 1977, 1982; Lewis 1977; Meyer et al. 1977; Gogineni et al. 1978; Bolivar EGEOESMIN !!iiiiiiiii~. 69* ~, =================================& Brookins 1979) it is now apparent that they , ~ :iii:::::::i:::iiiii have more lamproite than kimberlite affinities (Scott-Smith & Skinner 1982, 1984a, b, c; Mitchell & Lewis 1983; Bolivar 1984). Note that Carmichael et al. (1974, p. 517) were arguably the HOLSTEINSBOR i ii!iiii)iii! first to suggest this feature. The COD lamproite is a 73 acre diatreme with three different lithological units: a massive hypabyssal olivine.... ~i iiiiiiiii+iiiiii phlogopite lamproite, an intrusive/extrusive volcanic breccia and a phlogopite lamproite tuff ~ ======================================== i!i!iiiiiiiiii (Figs 3, 4 and 10). Only the breccia is presently S U K K E R T O P P E j,.~ ~ i i ::::,o.~ i : : : :::i!!!!! i i ...... ::::!!!~:::::::: believed to contain an economic concentration of diamonds, although the three lithologies have not been tested using modern techniques. At least four other associated lamproite intrusions occur 1-3 km to the NE of COD: the Kimberlite mine, the American mine, Twin Knobs and Black Lick FIG. 7 Generalized geologicalmap of W Greenland prospect. All are diamondiferous except for the showing the relationship between lamproite and last. kimberlite intrusives and structural elements. (After The PRA intrusion was first mentioned by Noe-Nygaard & Rambert 1961 ; Scott 1977, 1981). Powell (1842; cited by Miser & Ross 1923) and received early attention as a micaceous peridotite (Branner & Brackett 1889; Williams 1891). Miser dine-diopside-phlogopite lamproite -t- K-richter- & Ross (1923) were perhaps the first to put the ite) and madupite (diopside-phlogopite lam- PRA intrusive into the geological context of the proite) (see Figs 3 and 4). The last mentioned is better-studied diamondiferous peridotites of typically SiO2-undersaturated whereas the first southern Africa, even though diamonds were two are often SiO2 saturated or oversaturated. initially discovered at COD in 1906. Additional Accessory phases include priderite, wadeite, studies of the geology, geochemistry and minerapatite, perovskite, magnetite (rare), pseudo- alogy of the PRA rocks include those by Miser & brookite, spinel and Sr-barite (Carmichael 1967; Purdue (1929), Ross et al. (1929), Moody (1949), Kuehner et al. 1981). Thoenen et al. (1949), Stone & Sterling (1964), Experimental studies of LH rocks (Yagi & Gregory (1969), Gregory & Tooms (1969), Lewis Matsumoto 1966; Sobolev et al. 1975; Barton et al. (1976), Meyer (1976), Brookins et al. (1979), 1976; Barton & Hamilton 1978, 1979, 1982) are Steele & Wagner (1979) and Waldman et al. summarized below. (1985). PRA diamonds and their mineral incluLherzolite (rare), harzburgite (rare), pyroxenite sions have been the subject of many studies and mica pyroxenite xenoliths occur at nearly all (Langford 1973; Giardini et al. 1974; Giardini & lamproite exposures (Ogden et al. 1978; Kuehner Melton 1975a, b, c) as have been their occluded 1980; Barton & van Bergen 1981) and particularly and included gases (Melton et al. 1972; Melton & in the volcanic necks. Crustal xenoliths occur in Giardini 1976, 1980). Roedder (1984, pp. 509the LH lavas (Kay et al. 1978). Detrital Cr10) disagreed with the interpretations made in pyrope, Cr-diopside and enstatite of ultimate the last three studies. mantle origin but unknown magmatic source (not The PRA lamproites intrude the nearly horithe LH) have been found in ant hills W of the LH zontal Lower Cretaceous Trinity Formation (clay, in the Green River Basin (McCandless 1982, silt, sand, gravels and limestones); peridotite 1984). pebbles occur at the base of the Upper Cretaceous
z:.
:ii
i if!
=======================
L a m p r o i t e s a n d K-rich igneous rocks
123
FIG. 8. Generalized geological map of the W Uinta Mountain area, Utah, showing the geologicalposition of the Kamas and Moon Canyon lamproites. (After Best et al. 1968; Henage 1972; Hintze 1980.)
Tokio formation (Miser & Purdue 1929). Zartman (1977) and Gogineni et al. (1978) suggest an age of intrusion of 97-106 Ma on the basis of K-Ar dating of phlogopite separates. The PRA intrusions occur on the margin of the Ouachita fold belt on the northern edge of the Gulf coastal plain (Wickham et al. 1976). The hypabyssal and breccia phases of the PRA intrusions consist of olivine phenocrysts (often serpentinized) in an extremely fine-grained groundmass of phlogopite, serpentine, perovskite, K-richterite, Cr-spinel, magnetite, diopside, wadeite and priderite (Figs 3 and 4). The epiclastic tuff unit (Scott-Smith & Skinner 1984a, b, c) contains abundant phlogopite, sanidine and quartz in an extremely altered clay-chloritecarbonate matrix. Both the breccia and tuff units contain abundant autolithic and xenolithic (sedimentary rock) fragments. Garnet (pyrope and almandine) is extremely rare and typically only encountered in heavy-mineral concentrates; picroilmenite only occurs as inclusions within garnets and as extremely rare xenocrysts. Coarsegrained lamproite xenoliths consisting of Krichterite, phlogopite, diopside and priderite from the massive hypabyssal facies have been described by Mitchell & Lewis (1983).
Leucite- and pseudoleucite-bearing intrusions (syenites and tinguaites) cut nepheline syenites at Magnet Cove, Arkansas, 90 km NE of PRA (Williams 1891); lamprophyres are widespread in the Ouachita Mountains (Robinson 1976; Mullen & Murphy 1985; Mullen & Petty 1985). Smoky Butte, Montana (SB) (47 ~19'N, 107~3' W)
N-NE-trending dykes and small plugs of lamproite intrude the Tullock member of the Palaeocene Fort Union formation (sandstone and shale) along the axis of a very broad gentlydipping regional syncline over a distance of 3 km just W of Jordan, Montana (Matson 1960) (Fig. 11). The three major rock types at SB contain the following mineral assemblages in addition to glass: (i) sanidine, K-richterite, K-riebeckite and armalcolite; (ii) Ti-phlogopite, diopside, olivine, sanidine and armalcolite; (iii) Ti-phlogopite, diopside, olivine, analcime and armalcolite (Velde 1975; Matson 1960). Velde (1975)noted the presence of priderite and Wagner & Velde (1986a) additionally recognized pseudobrookite and chromite. D. Velde (personal communication, 1985) noted the presence of a new mineral, davanite (K2TiSi601 s), in the SB rocks.
I24
S. C. B e r g m a n
FIG. 9. Generalized geologicalmap of the northern Rock Springs Arch of the Green River Basin showing the location of the Leucite Hills lamproites : BT, Boars Tusk; MAT, Matthew; SM, Steamboat Mountain; BR, Black Rock; OR, Orenda; BAD, Badgers Teeth; ZIR, Zirkel Mesa; EM, Emmons Mesa; HAT, Hatcher Mesa; HAG, Hague; HAL, Hallock; END, Endlich; NP, North Pilot Butte; NT, North Table Butte; ST, South Table Butte; CR, Cross; OS, Osborn; P, Pilot. (After Cross 1897; Kemp & Knight 1903; Love et al. 1955.) The dykes are generally extremely fine grained to glassy in texture and often contain zeolite or carbonate-filled vesicles, but massive coarsegrained (1-2 mm) type (ii) lamproite with residual groundmass glass occurs at the Smoky Butte plug. Bedded welded autolithic lamproite lapilli and piperno (welded agglutinate tufts) in clay-rich matrices overlie and are marginal to the intrusions. Contact intrusive breccias composed of vesicular deformed magmatic lamproite fragments in a clay-rich matrix also occur at the plug and are almost identical with the breccias observed at LH, WKB, MAP and in other lamproite suites (S. C. Bergman, unpublished data).
Mitchell & Hawkesworth (1984) discussed the trace-element and Nd-Sr isotopic geochemistry of the SB intrusions. On the basis of K-Ar dating (Marvin et al. 1980), the SB intrusives are 27 Ma old, relative to the 47-53 Ma ages of the Winnett alnoite (100 km WSW of SB), the Missouri Breaks kimberlites and ultramafic lamprophyres (160 km WNW of SB) (Hearn 1968, 1979; Hearn & McGee 1982, 1984), and the Highwood Mountain shonkinites and minettes (300 km W of SB). Sapphire-bearing minette-like lamprophyres occur at Yogo Gulch in central Montana (e.g. Meyer & Mitchell 1985). Although only obscurely described in the literature, dykes whose occurrence and mineralogy are similar to that of SB
Lamproites and K-rich igneous rocks
125
FIG. 10 (a) Generalized geological map of the Murfreesboro area, Arkansas, showing the positions of lamproite intrusives at Prairie Creek, Twin Knobs, Kimberlite Mine, Black Lick and American Mines; (b) generalized geological map of the Prairie Creek lamproite diatreme; (c) E-W cross-sections of the Prairie Creek lamproite diatreme showing two alternative interpretations (left half of diagram after Bolivar 1977).
I26
S. C.
Bergman
FIG. 11 (a) Map of Montana showing the locations of the Montana petrographic province alkaline rock suites (after Larsen 1940; Hearn 1979; Marvin et al. 1980) (LBM, Little Belt Mountains; HB, Haystack Butte; HM, Highwood Mountains; CAM, Castle Mountains; CZM, Crazy Mountains; BBM, Big Belt Mountains; MR, Madison Range; GM, Gallatin Mountains; ABR, Absaroka Range; SH, Sweet Grass Hills, Haystack Butte, East Butte; BPM, Bearpaw Mountains; RB, Rattlesnake Butte; TB, Twin Butte; LRM, Little Rock Mountains; MBD, Missouri Breaks Diatremes; SB, Smoky Butte; ID, Ingomar Dome Diatreme (Froze-to-Death Butte); PD, Porcupine Dome; WA, Winnett Alnoite; JM, Judith Mountains; MM, Mocassin Mountains); (b) generalized geological map of NE Montana showing the geological environment of the Smoky Butte lamproite intrusives and extrusives (after Ross et al. 1955); (c) geological map of the Smoky Butte intrusives and extrusives (after Matson 1960).
Lamproites and K-rich igneous rocks occur in Rosebud County, Montana, 90-100 km S of SB at T10N, R36E, near the Ingomar Dome at Froze-to-Death Butte (Heald 1926, pp. 26-7; Marvin et al. 1980) and at T10N, R38E, on the Porcupine Dome (Bowen 1915, p. 66). The welldescribed igneous rocks geographically closest to SB are those of the Winnett alnoite sill (Ross 1926a; Marvin et al. 1980). [Central Sierra Nevada, California ( C S N ) (37 ~ 29'N, 119 ~ 19' W ) ]
Ultrapotassic basanitic lavas and plugs (free of plagioclase) and associated potassic olivine and alkali basalts (plagioclase bearing) were erupted from 3.3 to 3.8 Ma ago through the core of the Sierra Nevada batholith (Van Kooten & Peck 1977; Van Kooten 1980; Dodge & Moore 1981). These rocks possess atomic K/Na and K/A1 ratios, MgO contents and Rb contents which are too low and CaO and Na20 contents which are too high to be lamproites. The phlogopites and diopsides of CSN lavas are richer in AI than are those of lamproites. Although the BaO contents of CSN lavas are extreme (0.3-0.6 wt.%) and similar to those of lamproites (0.6 wt.~), the Rb levels are extremely low (less than 100 ppm) for lamproites (typically 200-300 ppm). [Deep Springs Valley, California ( D S V ) (37 ~ 21' N, 118 ~ 03' W ) ]
Four leucite-bearing mafic, olivine melaleucitite, olivine-leucite trachybasalt and potassic basalt plugs and flows (Dodge & Moore 1981) contain plagioclase and have whole-rock K 2 0 / N a 2 0 and K20/A1203 atomic ratios too low to be lamproites. [Highwood Mountains, Montana ( H M ) (47 ~ 31' N, 110 ~ 48' IV)]
I z7
[Navajo-Hopi Buttes, Arizona ( N H B ) (35 ~ 25' N,
11o~ w ) I Dykes, flows, volcanic necks and pyroclastics of potassic and ultrapotassic mafic melts were erupted through Triassic and Cretaceous sediments in two intervals (2-5 and 25-30 Ma ago (Roden et al. 1979)) in the Four Corners area on the SW part of the Colorado plateau (Williams 1936; Anderson et al. 1970; Roden 1981; Rogers et al. 1982). The Navajo rocks are potassic whereas the Hopi Butte localities tend to be sodic. These rocks are not lamproites sensu stricto because they have atomic K 2 0 / N a 2 0 = 1.3, K20/ AlzO3=0.3 and their phlogopites contain 1215.5 wt.~ A1203; however, the presence of analcime-bearing minettes suggests that lamproites may occur as end-members of the suite.
[Seward Peninsula, Alaska ( S E W ) (65 ~ N, 116 ~ IO' W ) ]
A belt of alkaline rocks 350 km long trends NE from the Darby Mountains of the eastern Seward Peninsula to beyond the Kobuk-Selawik lowlands 200 km to the NE (Miller 1970, 1972). The complex is middle Cretaceous in age (97-110 Ma) and consists dominantly of peraluminous potassic nepheline syenites and related rocks with a minor component of peraluminous ultrapotassic rocks. As shown by Miller (1972), the SEW rocks are nearly identical, in terms of both mineralogy and geochemistry, to the HM suite, Montana. Some of the SEW rocks are extremely ultrapotassic, and contain among the highest K20 contents of igneous rocks recorded in the world because of the abundance of kalsilite (atomic K z O / N a 2 0 = 1 3 , 16.6 wt.~ K20: Kobuk Selawik Lowland juvite) (Miller 1972).
The Highwood Mountains sub-province (Larsen et al. 1941 ; O'Brien et al. 1985) displays the best-
developed suite of potassic-ultrapotassic mafic alkaline rocks (see Fig. 11) of any suite within the alkaline petrographic province of central Montana (Weed & Pirsson 1895; Buie 1941; Larsen 1940). The HM rocks occur as flows, plugs, dykes, plutons, sills, lopoliths and laccoliths and were emplaced 50-53 Ma ago (Marvin et al. 1980) about 160 km E of the Rocky Mountain front. Petrographic types of interest include nepheline-normative shonkinite, missourite and minette (mafic phonolite). The HM rocks are not as perpotassic (KzO/A120 3 = 0.4) or ultrapotassic (KzO/NazO,~ 1.5) as typical lamproites (Tables 4 and 5).
[Fortification dyke, Colorado ( F O R ) (40 ~ 48' N, 107 ~ 40' W ) I
Ross (1926b) described a dyke 16 km long (3-6 m wide) trending N60~ about 30 km N of Craig, Colorado, that contains a majority of oligoclase and orthoclase with minor mafics. The FOR dyke, though originally termed 'verite' by Ross (1926b), is not a lamproite because it contains sodic plagioclase and consequently has low K20/ NazO and K20/A1203 ratios. The present author classifies the Fortification dyke as a lamprophyre, but modern diagnostic mineral and chemical studies are needed to say anything further.
S. C. Bergman
128 [ Two Buttes, Colorado ( TB ) (37~
102~
' W) ]
Cenozoic laccoliths and associated dykes of prowersite at Two Buttes in the Arkansas Valley (Gilbert 1896; Cross 1906) are minettes even though they were originally included in Niggli's lamproite clan. Despite an ultrapotassic geochemistry (atomic KzO/Na2O=5), A120 3 contents (12.3~) are slightly high for lamproites. Although the analysis may be suspect, the AI20 3 content of an augite separate is 3.1 wt.~ (Cross 1906), which is much too high for lamproites (see below). [Knox County, Maine ( K N X ) (44 ~1 I' N, 69 ~10' W)]
Prowersites from Appleton Maine (Bastin 1906) contain large phenocrysts of perthitic feldspars. Despite a low A120 3 content (10.6 wt.~) and an ultrapotassic character (atomic K20 / Na20 = 2.5), the KNX rocks are not lamproites or even lamprophyres. [Galore Creek, British Columbia (GAL) (57 ~ 14' N, 131 ~ 30' W ) ]
Ultrapotassic pseudoleucite-bearing syenite porphyries and phonolites occur in the upper Galore Creek Valley, NW British Columbia (Allen et al. 1975). Although altered with abundant evidence for orthoclase replacement, these peraluminous rocks contain elevated K20 contents (9-13 wt.~) and atomic K20/Na20 ratios (about 3). [Spotted Fawn Creek, Yukon (SFC) (64 ~ 22' N, 133 ~ 42' W)]
Pseudoleucite-bearing tinguaites (coarse-grained phonolites occur in the Spotted Fawn Creek area in W central Yukon (Templeman-Kluit 1969). The rock is ultrapotassic, yet metaluminous with atomic K20/Na20 = 3.1 and KzO/A1203 = 0.7. [A illikBay, Labrador (ALL) (59 ~15' W, 55 ~15'N)]
Minette dykes occur as part of the Mesozoic mafic-ultramafic lamprophyre-carbonatite suite of Aillik Bay, Nain Province, Canada (Hawkins 1976; Foley 1982, 1984). These dykes intrude Archaean gneisses and Proterozoic metavolcanics and metasediments and include the following petrographic types: minette, monchiquite, alnoite, aillikite and carbonatite. Although the heretofore understood geochemistry and petrography (Kranck 1953; Hawkins 1976) of the AIL 'minettes' (6~ olivine, 359/00augite, 28~ biotite, 25~ orthoclase and 4 ~ oxide; 44~ SiO2, 4.0 wt.~ TiO2, 9.1~ A1203, 11.0~ FeO*, 10.5~ MgO, 7.5~ CaO, 6.3~ K20, 1.8~/o Na20, 1.1~ P205) displayed characteristics typical of minettes as a
group (cf Rock 1986), recent work has demonstrated the existence of priderite in some of the 'minettes' (K. Collerson, personal communication, 1984). More detailed work is therefore required to establish the true affinities of the AIL dykes, although it is possible that true minettes may carry priderite. [Chino Valley, Arizona ( C H I N ) (35 ~ N, 112 ~ 30'
w)] A 25 Ma old eclogite- and peridotite-xenolithbearing potassic latite (55-60 wt.~ SiO2, atomic K20/Na20 = 1-3 and 4-5 wt.~ MgO) and potassic felsic latites occur at Sullivan Buttes, near Chino Valley, Arizona (Arculus & Smith 1979; Roden et al. 1979; Schulze & Helmstaedt 1979). [Colima Graben, Mexico (COL) (19 ~ 40' N, 103 ~ 40' W) ]
Late-Quaternary K-rich basalts, andesites and related rocks occur in the Colima Graben of the E-W-trending Mexican volcanic belt (Luhr & Carmichael 1980, 1981). These workers have described phlogopite-bearing K-rich basalts which are petrographically similar to minettes. [Other K-rich rock localities]
Brief descriptions of other occurrences of N American K-rich rocks can be found in Glazner & Stork (1976), Shafiqullah et al. (1976), Glazner (1979), Rowell & Edgar (1983), and references cited therein.
Australia
In contrast with the widespread occurrence of diamonds and kimberlites in Australia (Fig. 12), lamproites sensu stricto only occur along the SW and SE flanks of the Kimberley craton in NW Australia. The leucitites of New South Wales are not included in the lamproite clan in this paper because of their anomalous character. The Terowie, S Australia, 'lamproites' reported by Colchester (1982, 1983) are also interpreted as olivine nephelinites (Carr & Olliver 1980; S. C. Bergman, unpublished data). W Kimberleys, W Australia ( W K B ) " ~ 8 ~ N, 124 ~ 45'E)
Over 100 individual occurrences of many petrographic types of lamproite sensu stricto flows, plugs, dykes, plutons, sills, diatremes, cryptovolcanic structures, tufts and other pyroclastics occur on the margin of the Proterozoic King Leopold mobile zone, the Fitzroy Trough and the
Lamproites
and K-rich igneous rocks
I29 [ ] PROTEROZOIC FOLD BELTS []ARCHAEAN
CRATON
RIFT VALLEY
~
CONTINENT-OCEAN
BOUNDARY ( M o ) -----" FRACTURE ZONE
-~- EXTRA-ARCH BASIN ::;:: TASMAN ::::::::::.~, i:i:i ===========BELT ========FOLD ===========:5:::::::::: =======================
22222224~222222;
O DIAMOND
ii!iiiiiii~ili!i!i!ili!i!i!iiiiiii!!i!i!i!iiiiilil
OCCURRENCE &, LAMPROITE ULTRAPOTASSIC ROCK 9 KIMBERLITE | ULTRAMAFIC LAMPROPHYRE 9 CARBONATITE
0 I
400 ~
km
I
FIG. 12. Generalized map of Australia showing the major Archaean cratons, Proterozoic mobile zones, continent-ocean boundary, oceanic crust fracture zones (and other features) and the locations of lamproites, kimberlites, carbonatites, ultramafic lamprophyres and diamond occurrences. (After Garlick 1979, 1983; Stracke etal. 1979; Veevers 1981 ; Atkinson et al. 1984a.)
Lennard Shelf (in three separate fields--Ellendale, Calwynyardah and Noonkanbah), all of which form the Canning Basin on the SW flank of the Kimberley Basin (Fig. 13). Diamonds occur in over 30 separate intrusions (Atkinson 1982; Atkinson et al. 1984b). These intrusions were discovered by Fitzgerald (1907), initially petrographically and chemically described by Simpson (1925), Farquharson (1920, 1922) and Skeats & Richards (1926), and have since received considerable attention in detailed work by Wade (1937), Prider (1939, 1960, 1965, 1982), Wade & Prider (1940), Prider & Cole (1942), Bell & Powell (1970), Derrick & Gellatly (1972), Mason (1977), Atkinson et al. (1982, 1984a, b), Jaques et al. (1982, 1983, 1984a, b), Nixon et al. (1982, 1984), McCulloch et al. (1983a, b), Smith (1984a, b) and Jaques et al. (1986). Although early K - A r and Rb-Sr work erroneously indicated ages of 35 Ma and 250 Ma respectively for WKB rocks (Prider 1960; Kaplan et al. 1967), more recent K - A r and Rb-Sr work on mineral separates limits the age of emplacement of 14 of the WKB lamproites to 17-25 Ma (Wellman 1973; Jaques et al. 1984a; Bergman & Onstott, unpublished data). This ultrapotassic and perpotassic suite consists of a petrological continuum from diamondiferous olivine lamproites to barren leucite lamproites (Atkinson et al. 1983, 1984a; Jaques et al. 1984b,
1986). The WKB olivine lamproites contain olivine megacrysts and aggregates in a finegrained to glassy groundmass of olivine, phlogopite, diopside, apatite, K-richterite, perovskite, wadeite and spinel, whereas the leucite lamproites are composed of phlogopite, diopside, K-richterite and leucite phenocrysts in a groundmass of several of the above phases with spinel, wadeite, priderite, jeppeite, shcherbakovite, perovskite, apatite and rare ilmenite (Figs 3 and 4). Kfeldspar is absent except where it replaces leucite. Wade & Prider's (1940) nomenclature for individual rock types is summarized in Table 1. The WKB lamproites occur on the southern margin of the Precambrian Kimberley block and intrude mostly Mesozoic and Palaeozoic sediments, although a few intrusions cut Proterozoic granitic rocks. The regional structure consists of dominantly NW-trending faults with E-W-trending anticlines and subordinate synclines oriented in an en echelon pattern. Mantle xenoliths are rare in WKB rocks but include dunites and harzburzites (Atkinson et al. 1984a; Jaques et al. 1984b, 1986). Many of the olivine-rich xenoliths display features similar to the aggregates with smaller olivine grains. Hall & Smith (1985), Jaques et al. (1986) and Smith (1984a) discussed aspects of the ARG and WKB diamonds.
S. C. Bergman
I3o 125
124
CRETACEOUS ~ ! CANNING BASIN SEDIMENTARY ROCKS TRIASSIC [ ] SEDIMENTARY ROCKS PERMIAN [ ] SEDIMENTARY ROCKS DEVONIAN [ ] REEF CARBONATES KIMBERLEY BASIN SEDIMENTARY ROCKS (,ARCHAEAN BASEMENT) L. PROTEROZOIC [ ] KING LEOPOLD MOBILE ZONE
126
9 LEUCITE LAMPROITE 9 OLIVINE LAMPROITE FAULT . . . . . ANTICLINE OR SYNCLINE
0 I
50 km i
1
(a)
[]MAGMATIC LAMPROITE COARSELY MICACEOUS [ ~ MAGMATIC LAMPROITE FINELY MICACEOUS []TUFF [ ] S A N D Y TUFF
0 L
I km J B
C /
A' A' ~ /
PLAN
CROSS-ECT,ON
:'it',
,,
0-200mCROSS-SECTIONS A
A'
B
B' C
C' .
.
.
.
.
.
.
.
.
.
i~ i,,:i-::.~i, -"
0 I
(b)
300~m J
40Ore•
[ ] MAGMATIC LAMPROITE DMUDSTON E [ ] TUFF
~oom
200m 300,m (c)
FIG. 13. (a) Generalized geological map of the Canning Basin and King Leopold Mobile Belt, W Australia, indicating the locations of olivine and leucite lamproite intrusives (after Bureau of Mineral Resources, Geology, Geophysics 1976; Atkinson et al. 1984; Jaques et al. 1984a); (b) geological map and cross-sections of the Ellendale olivine lamproite diatreme, Lennard shelf (after Atkinson et al. 1984a); (c) geological map and crosssection of the Calwynyardah olivine lamproite diatreme, Fitzroy Trough (after Atkinson et al. 1984a; Madigan 1983).
Lamproites and K-rich igneous rocks
Creek mobile zone at Lissadell Road and Bow Hill (Atkinson et al. 1984a, b, Jaques et al. 1986).
Argyle, E Kimberleys, W Australia ( A R G ) (16 ~ 40' S, 128 ~ 25' E)
The diamondiferous Argyle AK1 olivine lamproite pipe is located on the eastern flank of the Kimberley craton in the 1940-1800 Ma old Halls Creek mobile zone (Atkinson et al. 1982, 1984a, b; Madigan 1983; Jaques et al. 1986). This intrusive is 125 acres in exposed area and is the richest diamond deposit in the world in terms of grade, containing proven reserves of 61 million tons at 680 ct per 100 tons and additional probable reserves of 14 million tons at 610 ct per 100 tons (Atkinson et al. 1984b) (Fig. 14). An extensive alluvial diamond deposit also occurs adjacent to the Argyle pipe in Smoke Creek and Limestone Creek (Madigan 1983; Meakins 1983). The ARG intrusive is most probably Precambrian in age since it is overlain by Cambrian sediments; it consists predominantly of an olivine lamproite sandy tuff that is intruded by dykes of magmatic olivine lamproite. The massive magmatic phases consist of euhedral olivine phenocrysts (replaced by talc and carbonate) and ragged tetraferriphlogopite microphenocrysts in a fine-grained groundmass of phlogopite, anatase, sphene, perovskite, apatite, Mn-ilmenite, Ti-Mg chromite and sulphides (Fig. 4). The sandy tuff is massive or weakly bedded and contains magmatic lamproite clasts in a groundmass of olivine, ash and xenolithic rounded quartz grains (Fig. 4). Leucite has also been reported in the ARG rocks. Dykes of mica peridotite occur near ARG in the Halls
A
13 ]
[Lake Cargelligo Area, New South Wales ( N S W ) (33 ~ 20' S, 146 ~ 25' E ) ]
Leucite-bearing lava flows and associated scoria cones and pyroclastics were erupted from 10 to 15 Ma ago (Wellman et al. 1970; Cundari et al. 1978) through a peneplaned basement complex of Palaeozoic geosynclinal sediments and granites and an overlying thin veneer of Cenozoic sediments of the central and southern Highlands Fold Belt over a 450 km • 150 km belt with concentrations near Lake Cargelligo, New South Wales (Cundari 1973; Cundari & Ferguson 1982, and references cited therein). The NSW essential mineralogy consists of diopside, olivine, leucite, Fe-Ti oxides, Ti-phlogopite and Ti-richterite with accessory apatite, sanidine and rare nepheline. Although the essential minerals of lamproites are present in the NSW rocks, their compositions deviate slightly from lamproites with Ti-richterite intermediate between the Na and K endmembers (K20/Na20=0.8-3.0 (weight basis)) and high A1203 contents of 1.7-2.4 wt.~; diopsides additionally possess a wide range in AI203 contents (as high as 2.8 wt.%). However, compositions of some phlogopite and diopside phenocryst cores possess typical lamproite characteristics with extremely low A1203 contents (9.7 wt.~ and 0.09 wt.~ respectively).
At31o
A':
[ ] LAMPROITENON SANDY TUFF LAMPROITESANDY TUFF [ ] LAMPROITECONTACT ZONE [ ] PROTEROZOIC SEDIMENTARY ROCKS
w FAULT ~
t
400m
-C
200m
500m J
PLAN
I
|"J 100 C'
I
100 CROSS-SECTIONS
FIG. 14. Generalized geological map and cross-section of the Argyle AKI olivine lamproite diatreme, Halls Creek mobile zone. (After Madigan 1983; Atkinson et al. 1984a).
I32
S. C. Bergman
Although the low CaO, A1203 and Na20 contents of the NSW rocks are typical of lamproites (average of 28 analyses, 9.0 wt.~, 8.7 wt.~ and 1.8 wt.~ respectively), the average atomic KzO/Na20 ratio of 2.1 and atomic K20 / A1203 ratio of 0.6 are rather low for lamproites. TiO2 contents in the NSW rocks are extremely high (range of 3.4-8.0 wt.%; mean, 4.3 wt.~) for lamproites (mean, 2.8 wt.~). The average Mg number of the NSW rocks (66) is somewhat low for lamproites (average, 74). The Sr-Nd isotopic compositions of NSW rocks are anomalous in that they plot close to bulk Earth instead of well within the 'enriched-Sr, depleted-Nd mantle quadrant' as other lamproites on a 143Nd/144Nd versus 87Sr/86Sr plot (R. Mitchell, personal communication, 1984). The NSW rocks are therefore not considered lamproites sensu stricto because of their major-element geochemistry, isotope geochemistry and mineralogy. They are closely allied, however, to the lamproite clan. [Mordor complex, Central Australia ( M O R ) (23 ~ 32'S, 134 ~ 27' E ) ]
The ultramafic to felsic rocks of the Mordor complex in central Australia (Langworthy & Black 1978) are potassic to ultrapotassic (atomic K20/Na20 -- 2-5) and mildly peraluminous (molar (K20 + Na20)/AI203 = 0.3-0.6) and are therefore not lamproites.
Europe Post-tectonic lamproites sensu stricto occur in at least four districts in Europe: in the Hercynides of Cornwall at Pendennis Point, in the Appenides of SE Spain between Murcia and Almeria, in NW Italy and in an isolated occurrence on the island of Corsica. Potassic to ultrapotassic nonlamproite rock suites characterize the Mediterranean area, the most notable being the strongly peraluminous central Italian volcanic province, but also including the Permian Exeter volcanics and Channel Island dykes of the U.K., the Bohemian Massif in Czechoslovakia, the Anatolia province of Turkey, the Spednogorie Zone, Bulgaria, the Hellenic arc in southern Greece, the Aeolian arc, southern Italy, and isolated occurrences in Iran and Norway. Interestingly, both the lamproites and compositionally similar potassic to ultrapotassic rocks of Europe are generally depleted in Ti and Nb (about 1-2 wt.~ TiO2 and less than 40 ppm Nb) relative to these rock types from most other worldwide localities (2-5 wt.~ TiO2 and more than 50 ppm Nb). These compositional features of Ti and Nb depletion are shared by subduction-derived ba-
salts and andesites (Ewart & LeMaitre 1980; Gill 1981 ; Briqueu et al. 1984) and will be discussed in more detail below. Note that minettes are widespread throughout Europe (Velde 1969b). Murcia-Almeria province, S E Spain ( M A P ) (38 ~ 30' N, l ~ W)
Breccia-mantled pipes, plugs, sills, flows and tephra of lamproite were emplaced through Mesozoic and Tertiary sedimentary rocks from 6 to 8 Ma ago (Bellon & Letousey 1977; Bellon 1981 ; Bellon et al. 1981 ; Nobel et al. 1981) at the SE edge of the Betic and Sub-Betic orogenic belts of the Alpine nappes over an area of 15 000 km 2 (Fig. 15). Ironically, these rocks have historically received more attention than any other lamproite suite, and yet they are the least like the other lamproites under present discussion. Niggli (1923) used the MAP suite as a type locality for lamproites. They were initially described by Lewis (1887), Yarza (1895), Osann (1889, 1906) and Washington (1903), and more recently by Meseguer Pardo (1924), Jeremine & Fallot (1929), Fallot & Jeremine (1932), Hernandez-Pacheco (1935), Parga Pondal (1935), San Miguel de la Camara (1935, 1936), San Miguel de la Camara et al. (1952), Fuster & Pedro (1953), Fuster et al. (1954, 1967), Fuster (1956), Fuster & Gastesi (1964), Hernandez-Pacheco (1965), Marinelli & Mittempergher (1966), Borley (1967), Fermoso (1967a, b), Velde (1969b), Fernandez & Hernandez-Pacheco (1972), Pellicer (1973), Caraballo (1975), Lopez Ruiz & Rodriguez Badiola (1980), Nixon et al. (1982, 1984), Venturelli et al. (1984a), Hertogen et al. (1985) and Wagner & Velde (1986a). Petrographic varieties include jumillite (abundant diopside and sanidine with minor phlogopite, K-richterite and olivine), fortunite (abundant hypersthene, phlogopite and sanidine with minor olivine, diopside and K-richterite), verite (abundant phlogopite with minor sanidine, diopside and olivine in a glass-rich groundmass) and cancarixite (abundant sanidine and Krichterite with minor olivine, diopside and hypersthene). Accessory phases include leucite, Crspinel, rutile, apatite, analcime and ilmenite, and alteration phases include quartz, serpentine, carbonate, chlorite and zeolite. Wadeite, priderite or shcherbakovite have not been reported. Compared with other lamproite suites, the MAP rocks are the most enriched in SiO2 and depleted in TiO2 and Nb. They are also the only lamproites that contain apparently primary magmatic orthopyroxene. It is interesting that the MAP lamproites are least like the other lamproites sensu stricto in terms of rock and mineral
Lamproites and K-rich igneous rocks
chemistry, despite the fact that Niggli used them as the type locality for lamproites. The jumillites are the most lamproite-like of all the MAP rocks.
N W l t a l y ( N W I ) (45 ~ 35'N, 7~ 45'E) and Orciatico, Pisa, Italy (PIS) (44 ~ N, 10 ~ E)
Dykes and rare flows of post-Alpine ultrapotassic lamprophyre and associated shoshonitic calcalkaline rocks were intruded N of the Canavese tectonic line in the Sesia-Lanzo and Combin Units (ophiolites, schists) of the internal NW Alps of NW Italy from 29 to 33 Ma ago (De Marco 1959; Krummenacher & Evernden 1960; Carraro & Ferrara 1968; Dal Piaz et al. 1973, 1979; Hunziker 1974; Scheuring et al. 1974; Zingg et al. 1976; Venturelli et al. 1984b). The mineralogy of the NWI ultrapotassic rocks is characterized by an abundance of phlogopite, diopside and sanidine with minor to trace quantities of altered olivine, riebeckite-arfvedsonite, apatite, sphene, Fe-Ti oxide and carbonate minerals (Dal Piaz et al. 1979; Venturelli et al. 1984b). The diopsides and phlogopites from the more ultrapotassic rocks (atomic K20/ NazO = 5-8) are typical of those of lamproites in that they have low A1203 contents (less than 0.3 wt.% and about 12 wt.~ respectively). The lesspotassic rocks (atomic K20/Na20<3.7), however, contain as much as 1.5 wt.~ A1203 in diopsides and 15 wt.~ A1203 in phlogopites. The geochemistries of the more-potassic rocks are also diagnostically lamproitic, with atomic K20/ A1203,~0.9, atomic K20/Na20~4, high MgO contents (9-14 wt.%) and low CaO contents (4.17.7 wt.~) (see Tables 4 and 5). Refractory traceelement levels are lamproitic, with 315-460 ppm Ni and 600-800 ppm Cr; contents of large-ion lithophile (LIL) elements also display lamproite affinities with less than 850 ppm Zr, less than 1240 ppm Sr and less than 570 ppm Rb. The NWI ultrapotassic dykes therefore fulfil the lamproite constraints summarized above, although Velde and Venturelli (personal communication to Bachinski 1985) group many of these rocks in the minette group. Rocks related to lamproites (termed selagites) containing olivine, clinopyroxene, phlogopite, sanidine, ilmenite (Borsi et al. 1967), K-richterite, chromite and apatite were erupted 1-4 Ma ago at Orciatico, Pisa, Italy (Stefanini 1934; Barberi & Innocenti 1967; Wagner & Velde 1986a). However, these rocks apparently do not contain oligoclase, a phase which Tr6ger (1935) and Johannsen (1931, 1933, 1935, 1939) include in selagites.
133
Pendennis ( P E N ) (50 ~ 5' N, 5 ~ 2' W) and Holmeade Farm ( H L M ) (50 ~ 54' N, 3 ~ 28' IV) Cornwall, U.K.
Ultrapotassic dykes, here considered lamproites, occur on the coast of SW England at Pendennis Point where they intrude Palaeozoic slates (Hall 1974, 1982). Analcime-bearing lamprophyres have been reported further inland at Holmeade Farm near Tiverton (Tidmarsh 1932; Knil11969, 1982; Velde 1971; Cosgrove 1972). Knill (1969) likened the HLM rocks to wyomingite and orendites from Leucite Hills and both Cosgrove (1972) and Velde (1971) termed them lamproites despite their minette affinities. The PEN dykes are estimated to be Permo-Carboniferous in age (Hall 1982), whereas HLM has been dated at 281__+11 Ma (Miller & Mohr 1964). The HLM locality falls within the widespread Exeter volcanic series which is treated separately below as a suite of non-lamproite ultrapotassic rocks. The PEN petrography is characterized by sanidine, phlogopite, alkali amphiboles (riebeckite-arfvedsonite), diopside, apatite, barite, spinel, futile, hematite, ilmenite, zircon, quartz, calcite and sulphide. The mineral chemistry of the dykes (low A1203 in diopside, amphibole and phlogopite) and their extremely ultrapotassic and perpotassic character (Table 4) permits their classification as lamproite sensu stricto. The HLM rocks are more difficult to classify with confidence because of a lack of detailed petrographic and mineral chemical data. They fall on the margins of the lamproite chemistry range, with atomic KzO/A1203 =0.6, which is too low for lamproites, and A1203 = 13.6 wt.~, which is too high. Their constituent phases include olivine, phlogopite, sanidine, diopside, anatase, magnetite, apatite, perovskite and analcime that was originally identified as leucite (Tidmarsh 1932; Velde 1969b). However, recent work by Jones & Smith (1985) on the Loxbeare Farm minettes demonstrates that the chemistries of phlogopites (13-15 wt.~ A1203), diopsides (23 wt.~ A1203 ; Mg number, 84-87) and sanidines (Or64_97 ; less than 1.0 wt.~ Fe203) are more typical of the minerals of minettes than of lamproites.
Sisco, Corsica, France (COR) (42 ~ 50' N, 9 ~ 25'E)
A small lamproite sill 13.5-15.4 Ma old intrudes Mesozoic schists W of the hamlet of Sisco, N of Bastia on the northern part of the island of Corsica (Fig. 16) (Prime11963; Velde 1967, 1968; Feraud et al. 1977; Civetta et al. 1978; Bellon 1981). The lamproite (termed porphyry by Primel,
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S. C. Bergman
FIG. 15. (a) Map of Spain and NW Africa showing the generalized structure and geology and the distribution of lamproites and other ultrapotassic rocks. (After UNESCO 1971 ; Venturelli et al. 1984a; Vila et al. 1974); (b) location map of SE Spain showing the principal lamproite and other potassic volcanic rock occurrences (after Venturelli et al. 1984a).
FIG. 16. Map of Corsica showing the generalized geology and the location of the Sisco lamproite, the size of which has been grossly exaggerated for clarity. (After Velde 1967; and Geological Survey of Italy 1961.)
Lamproites and K-rich igneous rocks
and minette and lamproite by Velde) contains phenocrysts of phlogopite and altered olivine in a groundmass of sanidine, diopside, phlogopite and richterite-arfvedsonite with accessory sphene, anatase, rutile, ilmenite, chromite, priderite and calcite (Wagner & Velde 1985). Phlogopite (13.8-15 wt.% A1203) and richterite (2.3 wt.~ A1203; 4.7 wt.~ K20) compositions are atypical of lamproites, but the presence of priderite (Velde 1968) suggests lamproite affinities. The rock is SiO2 rich (56-57 wt.%) and MgO poor (6-7 wt.%) for lamproites in general but possesses a high K20 content (10~) and high atomic K20/Na20 and K,O/AI203 atomic ratios (5.4 and 0.9 respectively). [Tuscany ( T U S C ) (42 ~ 15' N, 11 ~ 32' E ) ]
Quaternary volcanic centres at San Venanzo in the Perugia province and at Cuppaello in the Rieti province (both about 50-60 km E of the central volcanic centres of the Latial province) contain kamafugitic lavas (Holm & Munksgaard 1982; Gallo 1984). The geochemistry of the TUSC kamafugitic rocks is very similar to that of the Toro-Ankole province with the exception of slightly lower TiO2, Nb and Sr contents in the TUSC rocks. The relatively high CaO and low SiO2 contents and the presence of kalsilite and melilite in the TUSC rocks distinguish them from lamproites. [Sunnfjord, W Norway ( S U N N ) (61 ~ 20' N, 7~ 20' E) ]
Two hydrothermally altered peralkaline ultrapotassic syenite dykes (Furnes et al. 1982) consist almost entirely of microcline (75~-80~) plus 159/0-209/o phlogopite, but the presence of the K Zr silicate dalyite, which is compositionally similar to wadeite, and the Ba-Ti silicate labuntsovite, which is similar to shcherbakovite, suggests that these rocks have a lamproite affinity. As it is not known to what extent the hydrothermal alteration has modified the primary chemistry and mineralogy, this occurrence is grouped with the lamproite allies. The presence of microcline is consistent with low-temperature reequilibration. [ E Srednogorie, Bulgaria ( S RE ) (42 ~40' N, 26 ~E) ]
Potassic rocks of the SRE include leucite basanites, limburgites, shoshonitic basalts, trachytes and latites (e.g. Boccaletti et al. 1978). Grozdanov (1979) and Stefanova & Boyadjieva (1975) discussed the mineralogy of rocks very close to lamproites in composition from Svidnya in the Sofia.
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[Bohemian Massif, Czechoslovakia ( B O H ) (49 ~ N, 14 ~ E ) ]
Although most BOH minettes (N~mec 1972, 1973, 1974) contain a typical assemblage of sanidine-phlogopite-augite, one particularly mafic and alkali-rich minette dyke also contains alkali amphibole (N6mec 1972). This rock contains 10.8 wt.~ AI20 3, 8.2 wt.% K20 and 1.0 wt.% Na20 and is geochemically and mineralogically similar to lamproites. Schulze et al. (1985) and N~mec (1985) presented mineral chemistry data on some of the BOH rocks which generally show that these alkaline minettes are distinct from lamproites sensu stricto. However, N~mec's (1985) report of priderite in the BOH rocks indicates the need for further study. Holub (1977) described durbachitic plutons (i.e. syenites) which contain apparently cognate xenoliths of ultrapotassic melasyenites that possess compositions similar to lamproites. These xenoliths contain plagioclase and are therefore not true lamproites. Radulescu (1966) discussed ultrapotassic rocks in the E Carpathians. [Exeter volcanics ( E X V ) (50 ~ 26' N, 3 ~ 19' W) and Channel Island lamprophyres ( C H N ) (49 ~ 13' N, 2~ ' W), U.K.]
The EXV (Hobson 1892; Tidmarsh 1932; Knill 1969, 1982; Jones & Smith 1985) comprise four geochemical groups of rocks (Cosgrove 1972): two are shoshonitic (K20/Na20,~I; K20/ A1203 ~ 0.2) and two are ultrapotassic but are not lamproites ( K 2 0 / N a 2 0 ~ 5 and K20/A1203 ~ 0.5). The HLM lamproite sensu latu discussed above falls in Cosgrove's latter group of EXV. The Channel Island dykes form a suite that extends from Guernsey and Jersey to the coast of Normandy (Lees 1974) and are mainly minettes. Wagner & Velde (1986a) described minette from St. Helier that is compositionally similar to minettes from NHB. They are only mentioned here because of their mildly-ultrapotassic character and spatial association with the EXV. [Other European K-rich rocks]
Although these localities are not related to lamproites, they are mentioned here because of their mildly potassic and alkalic character. Potassic volcanics occur in the Roman Province (RP) and Aeolian Arc, Italy (Appleton 1972; Rosi 1980; Civetta et al. 1981 ; Barberi et al. 1974, 1978), the Laacher See, F.R.G. (Duda & Schminke 1978; Wimmenauer 1974), the Hellenic Arc, Aegean Sea (Innocenti et al. 1982), Anatolia, Turkey (Keller 1983) and NW Iran (Riou et al. 1981).
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Africa
The most thoroughly studied African lamproites occur at Bobi and Sequela in the Ivory Coast but, in marked contrast with European or American lamproites, only a limited amount of petrological information is available in the literature. In this regard, it is ironic that Africa is the site of the largest number of petrogenetically related kimberlites for which a wealth of data is available. The kamafugitic ultrapotassic province at ToroAnkole on the E African rift is not a lamproite suite but is perhaps the most thoroughly studied suite of its type. The following lamproite summaries must therefore be disappointingly brief until more thorough studies are published.
Bobi-Sequela, Ivory Coast (BOB) (8 ~ 9' N, 6 ~ 33'
w) Sheared dykes, termed 'meta-kimberlites' by Bardet (1973), intrude Archaean crystalline rocks in the Bobi and Toubabouko regions, Seguela, central Ivory Coast (Knopf 1970; Bardet 1973; Mitchell 1985). These dykes are thought to be 1150-1430 Ma old and diamondiferous in character (Bardet & Vachette 1966; Bardet 1973). The BOB dykes possess a geochemical signature similar to that of lamproites (Table 4); however, it is not known to what extent metamorphism has modified their primary composition. Dykes called fitzroyite and wyomingite by Bardet (1973) occur in the BOB area; they lack pyrope and ilmenite but contain abundant chromite. Intimately associated Proterozoic kimberlites and alnoites followed the strike of older fracture systems during their emplacement. The mineralogy of the BOB 'meta-kimberlite' includes phlogopite, diopside, spinel and talc. Alluvial diamonds from the Seguela area have been traced upstream to dykes of altered fitzroyite and mica-peridotite (Dawson 1967a, b; Knopf 1970).
Luangwa Graben, Zambia ( L U A ) (11 ~ S, 33 ~ E)
Little has been published on the K-Ti-richteritebearing lamproites from the Luangwa Graben (Murray, personal communication to Dawson 1970; Dawson 1980, p. 8; Mitchell 1985). A recent review of diamonds (Wilson 1982) mentions the probability of their existence. Kimberlites in the area have been briefly discussed by Dawson (1970, 1980). The lamproites include dykes, pipes and craters which intrude Karoo sedimentary rocks and are probably Jurassic to Cretaceous in age (B. ScottSmith, personal communication, 1986). Crater
facies lapilli tufts and associated pyroclastic and epiclastic deposits occur in the LUA. All of the characteristic lamproite phases occur (B. ScottSmith, personal communication, 1986), yet wadeite or priderite have not been discovered. The LUA is an exceptional lamproite that occurs in a well-developed continental rift setting. Pniel, Postmasburg, Swartruggens etc. South Africa (PPS) (33 ~ 54' S, 18 ~ 58' E)
Dykes of phlogopite-rich and leucite-bearing 'lamprophyres' are found in association with diamondiferous micaceous kimberlites at the Hellam Mine at Swartruggens, S Africa (Skinner & Scott 1979; Hargraves & Onstott 1980). The Swartruggens 'leucite lamprophyre' contains phlogopite, olivine, diopside, leucite and spinel (Skinner & Scott 1979) and possesses a geochemical composition closer to that of olivine lamproites than that of leucite lamproites (Table 4). Lamproite dykes are also reputed to occur at Pniel (Baster's Mine), Postmasburg and Pielansburg in the Johannesburg and Kimberley areas of S Africa (Dawson 1970, 1980, p. 8; Erlank 1973; L. G. Krol, R. Mitchell and A. J. A. Janse, personal communication, 1982, 1985). Feldspathoid-bearing kimberlites (typically containing nepheline, often out of Sr istopic equilibrium with the host rock (Smith 1983b)) have been reported by De Beers (Skinner, personal communication to Smith 1983b) at Klipfontein (28 ~ 23' S, 24 ~ 08' E) and Poortjie (28 ~ 01' S, 24 ~ 17' E). Unfortunately little information is currently published on their geochemistry and mineralogy. The compositionally related (perhaps identical ?) Group II kimberlites of S Africa are discussed by Smith (1983b) and Dawson (1987). [ Toro-Ankole and Birunga volcanic field, Uganda ( T A N ) (1 ~ S, 29 ~ E ) ]
Potassic-ultrapotassic mafic-ultramafic rocks at Toro-Ankole and Birunga characterize the western branch of the E African rift valley system in SW Uganda. Tufts, lavas, agglomerates and explosion breccias consist of a variety of kamafugitic rock types, ranging from melilite- and kalsilite-bearing katungites through olivine leucitites to K-trachyandesites (Holmes 1950; Bell & Powell 1969; Edgar & Arima 1981 ; Ferguson & Cundari 1982). The more mafic rock types do not contain any plagioclase or K-feldspar. These rocks are not lamproites sensu stricto, despite the fact that some workers (e.g. Vollmer & Norry 1983b) refer to them as such. The TAN rocks are members of the kamafugitic suite of ultrapotassic rocks (Sahama 1974) that are mineralogically and geochemically distinguished from true lamproites
Lamproites and K-rich igneous rocks by their nepheline, melilite and kalsilite, as well as by their elevated A1203 and CaO contents and lower SiO 2 contents (Table 4). [Azzaba, Algeria ( A Z Z ) (36 ~ 43' N, 7~ 5' E) ]
Glassy dykes and extrusive rocks, termed 'potassic olivine trachytes' by Vila et al. (1974), occur near Azzaba, Algeria (see Fig. 15a). Olivine ranges in composition from Fo84 to Fo91 and occurs with augite, sanidine and chromite (Vila et al. 1974; Velde, personal communication, 1985). The AZZ rocks are ultrapotassic (K20/ Na20 = 3.5), but their A1203 contents are rather high for lamproites (12.9%) and the rocks are peraluminous with atomic K20/A1203=0.6. These rocks are therefore best described as lamproites sensu lato until more detailed mineral chemistry and petrographic data are available. [Shaki and Okeho, S W Nigeria ( S H O ) (8 ~ 39'N, 3~
Ultrapotassic, nearly perpotassic syenites and biotite pyroxenites occur in several isolated bodies in the basement complex of W Nigeria (Oyawoye 1976). They consist of microcline, albite, biotite, amphibole, pyroxene, quartz and sphene, and possess atomic KzO/A1203 ratios of 0.6 and atomic K 2 0 / N a 2 0 ratios of 3 with 8.7
wt.% K20. Antarctica
Figure 17 shows the location of the three known lamproite sensu stricto occurrences in Antarctica:
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Gaussberg, Mount Bayliss and Priestly Peak. These rocks range in age from Palaeozoic to Quaternary and the most recent rocks at Gaussberg occur on the E coast where an extension of the Kerguelen Plateau of the Indian Ocean intersects the continent. The 2000 km long Kerguelen-Gaussberg aseismic ridge has been the site of extensive hot-spot activity, involving the generation of mildly-potassic alkali basalts, for the last 27 Ma (Stephenson 1972; Nougier 1970). All the Antarctic lamproites are rich in TiO 2 (3-5.5 wt.~ TiO2; average lamproite, 2.8 wt.%) and LIL elements (400-1500 ppm Ba, 900-1800 ppm Zr, 1300-3000 ppm Sr, 40-150 ppm Nb and 270-350 ppm Ce). The Antarctic lamproites are also particularly rich in P205 with 1.5-3.3 wt.~ (4-10 modal ~ apatite); they share this trait with the Indian apatite-rich lamproites (see below). It should be noted that E India and E Antarctica were adjacent prior to the breaking of Gondwanaland, which suggests a unique character in the underlying mantle lithosphere. Gaussberg (GSB) (66 ~ 48' S, 89 ~ 19' E)
Leucite lamproite lavas and tephra were erupted from the Gaussberg volcano approximately 56 000 years ago on the coast of Wilhelm II land, E Antarctica (Drygalski 1904; Phillipi 1912; Reinisch 1912; Lacroix 1926; Nockolds 1940; Vyalov & Sobolev 1959; Sheraton & Cundari 1980; Cundari & Ferguson 1982; Tingey et al. 1983). The more recent K - A r work on leucite separates by Tingey et al. supersedes the whole-
FIG. 17. Generalized geology of E Antarctica and the southern Indian Ocean showing the distribution of lamproites and seamounts of the Kerguelen-Heard Island chain. (After UNESCO 1976; Sheraton & England 1980.)
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S. C. B e r g m a n
rock K-Ar work of Ravich &Krylov (1964) and Soloviev (1972) which suggested ages of 20 Ma and 9 Ma respectively. In addition, Sheraton & Cundari (1980) and Tingey et al. (1983) described as fission-track work by Gleadow which suggested an age of 25 000+12 000 years. Dort (1972) postulated a late Pleistocene to Recent age on the basis of geomorphology. The GSB lavas have been termed leucitites by most workers. The lavas are nearly aphyric with phenocrysts of olivine, diopside and leucite in a glassy matrix (50~-60~ of the rock) containing quench crystals of leucite, diopside, phlogopite, amphibole, ilmenite and chromite (Sheraton & Cundari 1980) (Fig. 4). The mineral chemistry of the silicate phases is diagnostic of lamproites with low A1203 contents in diopside and phlogopite (averaging 0.4 wt.~ and 6.9 wt.~ respectively). The GSB whole-rock geochemistry is also typical of lamproites (Table 4), and emphasizes its perpotassic and ultrapotassic character. Both cognate (olivine-diopside-leucite-phlogopite) and accidental (spinel lherzolite) xenoliths occur in the volcanics (Sheraton & Cundari 1980). Collerson & McCulloch (1982) discussed the Nd and Sr isotope and rare-earth element geochemistry of the GSB lavas. Mount Bayliss ( M B A Y) (73 ~ 26' S, 62 ~ 50' E)
A dyke 5 m thick, termed 'melasyenite' by Sheraton & England (1980), intruded Archaean felsic rocks during the Silurian (413-430 Ma) in the southern Prince Charles Mountains, MacRobertson Land, approximately 300 km SSW of the Radok Lake alnoites (Trail 1963; Walker & Mond 1971; Sheraton & England 1980; Tingey 1981 ; Sheraton 1983). Medium-grained rocks are composed of microcline, K-richterite, K-arfvedsonite, phlogopite, apatite, anatase, zircon, calcite and Fe-Ti oxide. Finer-grained samples have been recovered from MBAY and these differ slightly from the coarser-grained rocks in that they contain pseudoleucite, more phlogopite, less amphibole and ilmenite. The mineral chemistries ofMBAY phlogopites and amphiboles are similar to those from WKB and LH lamproites. The geochemistries of the MBAY rocks are also typical of lamproites in that they possess atomic ratios K 2 0 / N a 2 0 = 3 . 2 and K20/A1203= 1.1; however, they are more evolved than the average lamproite (Mg number, 74) with an Mg number of 56. Priestly Peak (PP) (67 ~ 11' S, 50 ~ 22' E)
A similar dyke (albeit undated) cuts a Proterozoic dolerite dyke at Priestly Peak in Enderbyland
near the coast (Sheraton & England 1980; Sheraton et al. 1980). The PP dyke is finer grained than that of MBAY and consists predominantly of phlogopite, microcline and K-arfvedsonite with minor amounts of apatite, quartz, rutile, sphene, zircon and barite (Fig. 4). The PP dyke (Mg number, 70) is similar in composition to the GSB lavas (Mg number, 70) in that it is more primitive than the MBAY dyke (Mg number, 56). Asia and Indonesia
The only known occurrences of lamproite sensu stricto in Asia and Indonesia are the Indian dykes at Chelima and the Gondwana Coalfields. Those at Coc Pia, North Vietnam are closely allied to lamproites. Phlogopite-leucite 'basalts' occur in central Kalimantan; however, the petrological data required to characterize them are not available. Potassic-ultrapotassic intrusive and extrusive rock suites have been reported for several areas in the U.S.S.R. In recent discussions with geologists at several Chinese geological institutes it was found that lamproites have not yet been recognized in China (C. Hearn & A. J. A. Janse, personal communication, 1985). Coc Pia and Sin-Cao, Upper Tonkin, N Vietnam (COC) (22 ~ 15' N, 103 ~ 30' E)
These lamproites, termed 'cocites' by Lacroix (1926, 1933a, b), occur near the N Vietnam-Laos border in Indochina. The dykes at Coc Pia cut Mesozoic alkaline syenites and granites and are grey fine-grained rocks consisting of olivine, diopside, phlogopite, magnetite, sanidine and leucite (Lacroix 1933a). According to Fromaget (1933), alkaline dykes also occur at Sin Cao, NE of Laid Chau, where they intrude Triassic sediments. The Sin Cao rocks consist primarily of augite and phlogopite phenocrysts with sanidine and analcime pseudomorphs after leucite in the groundmass. Lacroix (1933a) included the Sin Cao rocks in his cocite group, and noted their chemical and mineralogical similarity to the Spanish jumillites. The geochemistry of the COC rocks is typical of lamproites with the exception of a rather low K20 value (4~-5~) and low atomic K20/Na20 and KEO/A1203 ratios (about 1.0 and 0.4-0.6 respectively). However, these discrepancies can be easily reconciled if the possibility of late- to post-magmatic alteration of leucite to analcime is considered. Lacroix (1933a) and Fromaget (1933) also described shonkinites and nepheline syenites at Pin Chai, S of Pu To and N of Lai Chau in Upper
Lamproites and K-rich igneous rocks
Tonkin. Wagner & Velde (1986b) re-examined the two type specimens from Coc Pia and have noted the absence of Ti-rich oxides and Krichterite, and the presence of the Ti-rich phlogopites titanomagnetite and feldspars (Or80_87). These two type specimens fulfil the mineralogical and compositional requirements for inclusion in the lamproite sensu strieto group; however, feldspar compositions are different from those of true lamproites. Further field work is therefore justified to search for other potential lamproites in the Coc Pia area. Chelima (CHE) (15 ~ 27' N, 78 ~ 42'E) and Gondwana Coalfields (GD W) (24 ~ N, 86 ~ E), India
Proterozoic (1200 Ma old) (Crawford 1969; Crawford & Compston 1973) phlogopite-rich, mafic-ultramafic dykes intrude Proterozoic shales of the Cumbum formation over a 10 km x 6 km area near the Chelima Railway Station on the western margin of the Cuddapah Basin and the eastern margin of the Dharwar craton, Andhra Pradesh, India (Fig. 18) (Appavadhanulu 1962; Sen & Narasimha Rao 1970; Sarma 1983; Bergman & Baker 1984). The two early reports described the dykes as minettes and the rocks as transitional between kimberlites and carbonatites, but the last work suggested, on the basis of rock and mineral chemistry, that they are lamproites. The rocks occur with massive magmatic to glassy breccia phases and consist of abundant foliated phlogopite (with talc and chlorite secondary intergrowths), rutile, apatite, perovskite, ilmenite, serpentine pseudomorphs after olivine and a secondary ferroan dolomite. Phlogopite composition and zoning characteristics are typical of lamproites (see below and compare Mitchell (1985)) with (core --* rim): Mg number, 8 5 4 65; AlzO3, 11 ~ 8 wt.~; TiO2, 5.5--* 3.5 wt.~; BaO, 0.7 4--, 0.3 wt.~; CrzO3, 0.2 ---, 0.0 wt.%. The geochemistry of the dykes is also typical of lamproites when corrected for the dolomite alteration (compare with Table 4): 48 wt.~ SiO2; 5.0 wt.~ A1203; 7.4 wt.~ TiO2; 9.4 wt.~ FeO; 18.9 wt.~ MgO; 4.2 wt.% CaO; 0.2 wt.~ Na20; 3.3 wt.~ K20; 2.5 wt.% PzOs; 560 ppm Cr; 450 ppm Ni; 2500 ppm Ba; 1920 ppm Sr; 160 ppm Rb; 380 ppm La; 700 ppm Ce; 1020 ppm Zr. Sen & Narasimha Rao (1970) noted micro-diamonds in the Chelima dykes; ancient workings, presumably for diamonds, are also present. Alluvial diamonds occur in an adjacent drainage within 6 km of this area (Bergman & Baker, submitted for publication). A widespread suite of Mesozoic lamprophyres and mica peridotites, some of which contain
139
leucite, occur about 1000 km to the N N E of the CHE dykes in the Gondwana Coalfields at Jharia, Bakaro and Raniganj on the N margin of the Singhbum craton (Fox 1930; Ghose 1949; Banerjee 1953; Makherjee 1961 ; Sanyal 1964; Chatterjee 1974; Sathe & Oka 1975; Sarkar et al. 1980). The geochemistry of the leucite-bearing GDW dykes is typical of lamproites (Table 4) and is much more mafic than typical minettes (compare with Rock 1986). Gupta et al. (1983) discussed the petrology of carbonated apatite glimmerites from the Damodar Valley which contain phlogopite, apatite, ankerite and chrome spinel phenocrysts in a groundmass of the same minerals plus rutile, devitrified glass, priderite and pyrite. All these GDW dyke rocks are extremely enriched in P205 (3-6 wt.~), TiO2 (5-7 wt.~) and carbonate minerals (10-20 wt. ~ CO2) compared with typical lamproites and, as a group, probably represent a distinct subtype of the lamproite clan. Indian kimberlite diatremes (840-1200 Ma old) are known in three districts: at Wajrakarur (more than six pipes), in the core of the Dharwar craton (about 250 km W of the Chelima dykes), and at Panna (more than four pipes) and Jungel (more than three pipes) near the SE margin of the Aravalli craton (500-800km W of the Gondwana Coalfield dykes) (Satyanarayana Rao & Phatare 1966; Crawford & Compston 1970, 1973; Marthur & Singh 1971; Paul et al. 1975; Balasubrahmanyan et al. 1978; Paul 1979; Murty 1980; Murty et al. 1980). Alluvial diamonds are widespread throughout central and northern India. Although the lamproites discussed above are widely separated from these three kimberlite occurrences alluvial diamonds and lamproites do occur in the same areas. Further work is clearly required on the Indian lamproites and lamprophyres. Kajan River, Kalimantan ( KAJ) (1 ~ 30' N, 115 ~ 30' E)
Brouwer (1909) reported a phlogopite-leucite basalt on the Kajan River, central-E Kalimantan, containing resorbed phenocrysts of phlogopite with reverse pleochroism, typical of the tetraferriphlogopites from lamproites and kimberlites (Mitchell 1985); additional minerals include leucite, diopside, olivine and Fe-Ti oxides. The KAJ geochemistry, however, is not strictly lamproitic (K20/A1203=0.4; K20/ Na20=2.0), and until more detailed data are available the KAJ rock must be grouped outside the lamproite clan. Lacroix (1926) included the 'kajanite' in his systematic treatment of the syenitic leucitites, but Niggli (1923) was apparently unaware of the kajanite occurrence or
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S. C. B e r g m a n
FIG. 18. (a) Generalized geological map of India showing the three Archaean cratons and the distribution of lamproites, kimberlites, lamprophyres and diamonds (after Roy 1962; Roy Chaudhury 1973); (b) geological map of Andhra Pradesh State, India, illustrating the location of the Chelima lamproite, Wajrakrur kimberlite and alluvial diamond occurrences (after Roy 1962; Roy Chaudhury 1973); (c) map showing the distribution of lamproite dikes at Chelima, Andhra Pradesh (after Sarma 1983). neglected it in his treatise on lamproites. Bergman et al. (1985) discussed the mineral and rock chemistry of minettes of non-lamproite affinity in the nearby Karamu River area, central Kalimantan. Bucking (1899) noted a biotite-leucite basalt in the S. Celebes. Indonesian arc volcanics (IA V) (6 ~ S, 111 ~ E)
Leucite basalts, tephrites and potassic andesites characterize isolated occurrences from the active Sunda arc in Java, the Celebes and other islands, and are peraluminous with low atomic K20/ A1203 ratios (0.2-0.3) and elevated A120 3 contents (16-22 wt.~o) (Grubb 1965; Baren 1948; Foden 1983).
Maksimov 1982; Kostyuk 1983). The most ultrapotassic occurrences (those of the Synnyr, Murun and Inagli massifs) do not, apparently, include lamproites. They are petrographically and geochemically similar to the HM, SEW and BRA ultrapotassic suites. Lacroix (1926) made an obscure mention of several 'kajanites' (phlogopite-leucite-diopsideolivine basalts) from Vouyerska Kosa and Sveta Petka, Siberia (with a reference to Raoult and a personal communication from M. Koyitch).
South America Brazil (BRA) (20 ~ S, 48 ~ W)
Baikal-Aldan Belt (BAL) and Pamir, Mongolia (PAM), U.S.S.R.
Potassic-ultrapotassic rocks are widespread along the 1500 km Baikal-Aldan alkalic intrusive belt on the Siberian Platform (Mineyeva 1972;
Shoshonitic volcanics, shonkinite and kamafugitic rocks occur in isolated localities in Brazil (Ulbrich & Gomes 1981, and references cited therein). Associated rock types include ugandite, analcite basalts, limburgite and minette, and
Lamproites and K-rich igneous rocks
although kimberlites and other alkaline rocks are widespread in Brazil (Svisero et al. 1979a, b, 1984), lamproites sensu stricto have not yet been reported.
Age of emplacement Known lamproites range in age from Proterozoic (Argyle, Holsteinsborg and Chelima) to Quaternary (Leucite Hills and Gaussberg). Of the 21 lamproite suites or localities recognized here,
141
nine are Cenozoic, four are Mesozoic, four are Palaeozoic and four are Proterozoic. Therefore, in addition to xenolith-bearing alkali basalts, lamproites are apparently most common in the Cenozoic. This age relationship could result from selective preservation of the extrusive members of recent suites or alternatively from some mantle or tectonic process permitting their formation which has been most active in recent times. The fact that the intrusion of Indian-AntarcticAustralian lamproites (CHE, GDW, GSB,
I42
S. C.
Bergman
MBAY, PP, A R G and WKB) has taken place in the same large region (when corrected for continental drift) over more than a billion years suggests that the former hypothesis may be the most likely explanation. On the basis of the better-studied Cenozoic suites (e.g. WKB and MAP), intrusive-extrusive activity takes place over a time span of 2-10 Ma. In areas such as the Fitzroy Basin the emplacement of primitive diamondiferous olivine lamproite diatremes can occur contemporaneously with that of more differentiated leucite lamproite flows and plugs (Jaques et al. 1984a). Therefore it is possible that diamondiferous olivine lamproite diatremes may be erupted in the Leucite Hills area in the next 10 Ma! However, the WKB suite is the only one in which such contemporaneity has been noted.
Intrusive-Extrusive forms Lamproites have been reported to occur in nearly every igneous form possible, the most common being dykes and flows. The most recent lamproites from GSB and LH display all the volcanological features typical of alkali basaltic eruptives, with classic lava flows, pyroclastic ejecta, cinder cones etc. (Fig. 19(c)). Therefore there is sufficient evidence to suggest that the more eroded slightly older lamproites (e.g. WKB) that occur as dykes and necks also probably possessed extrusives similar to those present at LH and GSB. Lamproite sills occur at HP and COR, and a shallow coarse-grained (grain size of less than 2 cm) lamproite pluton occurs at the 2 km diameter Wolgidee Hills intrusive (WKB). Vesicular rocks are ubiquitous in all but the hypabyssal intrusives; extremely vesicular scoriaceous lavas occur at GSB, LH and MAP among others. Palagonite and agglutinate tufts (welded fall-out tufts equivalent to piperno) and/or bedded pyroclastics occur in the more recent suites (e.g. GSB, LH, MAP, SB, WKB and PRA). Their absence in older suites can be easily attributed to selective erosion of extrusive forms. These extrusive facies are evidence of base surge as well as other volcanic processes (Atkinson et al. 1984b). Autolithic breccias composed of fragments of magmatic lamproite that are often plastically deformed and set in a fine-grained groundmass consisting of both magmatic and xenolithic (pulverized country rock) material characterize the contact zones of nearly every lamproite intrusive (Figs 3 and 4). These autolithic breccias are diagnostic of lamproites and are texturally similar to pyroclastic breccias associated with silicic volcanic centres. Glass-rich rocks are
present in many lamproite sensu strieto occurrences (e.g. LH, GSB, MAP, WKB and SB). The size oflamproite occurrences varies widely from locality to locality, from dykes of limited aerial extent less than 1-2 m wide at HOL, N W I and MBAY, PP and CHE, through small plugs (10-100 m wide) at KAM, MAP and HP, to the volcanic cone 370 m high and 1400 m in diameter at GSB. The estimated total volume of the erupted material at LH, WKB and MAP exceeds 10-100 km 3 in each case. Therefore, whereas lamproites are volumetrically of limited size compared with mafic alkalic igneous rocks in general (Mauna Loa volcano, Hawaii, alone exceeds 100km3), a given eruptive suite can approach the size of a continental alkali basalt volcanic field. Lamproites also occur as diatremes or pipes, and these igneous forms have the greatest potential for possessing a diamondiferous olivine lamproite phase such as those at PRA, WKB and ARG. These lamproite diatremes differ in appearance from the typical kimberlite diatremes (Figs 19(a)and 19(b)). Whereas kimberlite pipes tend to be carrot-shaped and possess 2-3 km of vertical flaring (Hawthorne 1975; Harris 1984), lamproite diatremes are characterized by a funnel or sherbet-glass shape with less than 0.1-0.5 km of vertical flaring. This contrast has important diamond-exploration implications (i.e. volume calculations and magnetic modelling) and can be reconciled by a consideration of the volatile abundances and compositions of the respective intrusive systems (Harris 1984). The larger proportions of both CO2 and H20 (a factor of 23) in kimberlites relative to lamproites produces a relatively-deep (2-3 km) explosive boiling stage in the former intrusives. The volatile budgets of lamproite magmas are dominated by the extremely-soluble HzO so that they boil at shallower depths of less than 1 km. An alternative explanation for the difference in shape between kimberlite and lamproite diatremes involves the relatively-shallower erosion level displayed by the latter. Since the vast majority of kimberlites are Pre-Cenozoic, whereas most lamproites are Cenozoic, this erosion level can be easily rationalized. However, even the older (Palaeozoic) lamproites (e.g. Argyle) display tuffaceous facies.
Tectonic-geological environment Whereas some of the lamproite sensu lato or compositionally related shoshonite localities occur on the margins of continents or in oceanic margins (e.g. Indonesia and the Mediterranean), the 21 lamproite sensu stricto localities described above all occur in continental regions. This
Lamproites a n d K-rich igneous rocks KIMBERLITE
DIATREME
~-
MODEL
.....
.,r--Z
CRATERFACIES~'~,{~'~.-~.7'_~>~<~~ - ~ - ' - - - ' -
LAMPROITE
I43 PIPE M O D E L
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--'" "
FINE-GRAINED MASSIVE
- ~ - - / ' ~ , . ~'~,~.~ ..... ~;,..~ '~ , ' ~ EROSIONAL ...... / _ _ ~ ' ~ ~ ~ i";'.':i:-!~" ;!!!~~,,~,f - ~ SURFA C E
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~
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.
.
.
.
.
.
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ALKALI BASALT CONE AND FLOW MODEL -~
\,
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- ' ~ ~ ~~?~--" ~ - . ~ C I N D E R MULTIPLE FLOWS . ~ A S H ~ F, ',~Lj~\
FIG. 19. Idealized cross-sections showing the morphological features of (a) olivine lamproite diatremes, (b) kimberlite diatremes and (c) alkali basalt cinder cones. (After Hawthorne 1975.) continental feature is shared by alkalic rocks in general (Bailey 1974) as well as by leucite-bearing rocks, with the exception of the four oceanic islands discussed by Gupta & Yagi (1980): the Cape Verde Islands, the Kerguelen Islands, the Marquesas Islands and Tristan da Cunha. However, whereas alkaline basaltic magmas with N a > K (weight basis) characterize nearly all well-developed continental rift systems (Bailey 1974; Barberi et al. 1982), with the exception of LUA, lamproites sensu stricto are absent and only
"CONE --,i
"
(c) a few related ultrapotassic rock suites (e.g. LAC, TAN and BAL) occur in well-developed continental rift systems. Lamproites are, in general, the result of postorogenic magmatic phenomena in regions that have experienced collisional orogenesis (with underlying fossil Benioff zones) several tens to many hundred million years prior to their eruption. In contrast, shoshonitic rocks are characteristically more directly linked with subduction. Morrison (1980) found that, while
I44
S. C. Bergman
shonkinites are associated with calc-alkaline rocks of orogenic zones, they tend to be younger and occur over the deeper parts of Benioff zones or in association with block faulting and uplift accompanying the 'flipping' ofa subduction zone. Therefore, whereas shoshonites are temporally associated with the waning stages of subduction, lamproites are clearly post-subduction and apparently require the passage of time in addition to the stabilization of orogenic belts for their formation. It should be noted that Helmstaedt & Gurney (1984) related the kimberlites of S Africa to subduction processes. In contrast with the general restriction of kimberlites to craton interiors (e.g. Janse 1984; Dawson 1980), lamproites generally occur closer to craton margins. In contrast with ocean island chains which young in a certain direction, lamproites do not and they consequently do not owe their presence to hot spots. Note, however, that Edgar (1983) related ultrapotassic magmatism of the Western U.S.A. (including lamproites) to the Yellowstone mantle plume. Several lamproite suites are associated with broad anticlinal or synclinal structures (e.g. WKB, LH and KAM), whereas others intrude relatively undeformed sedimentary-platform sediments resting on Proterozoic basement rocks marginal to the core of the continental craton (e.g. HP, PRA, SB, WKB and LH). Nearly all lamproite suites are associated with surficial lineaments or fault lines which probably connect with well-established deep basement fractures or zones of weakness. Lamproites share this feature with kimberlites and many other mantle-derived extrusive and intrusive rocks. Rowell & Edgar (1983), speculating on the spatial relationships between 15 dated occurrences of Cenozoic K-rich volcanics ( K 2 0 > Na20; SiOz<60 wt.~) in the western U.S.A. and the outer limits of palaeo-arc-related magmatism, suggested that the K-rich volcanism was directly related to deep subduction rather than intra-plate tectonics. Although most of the nonlamproite potassic suites correlated well, none of the lamproites sensu stricto did: they were emplaced long after the ambient western U.S.A. subduction had ceased. Lamproites are therefore controlled by intra-plate tectonic processes, although an association with regions that have experienced collisional orogenesis is indicated by the more recent suites.
Diamondiferous character Except for the CHE lamproite in India which requires more study, there are five known diamondiferous lamproite suites: WKB, ARG,
PRA, LUA and BOB. As the first three have received detailed study, but the latter two have not, the following is based largely on the first three suites. Diamondiferous lamproites all share the following features. 1 They are generally olivine lamproites (mafic to ultramafic) with elevated refractory element contents: 14-30 wt.~ MgO, 35-56 wt.~o SiOz, 26 wt.~ CaO, 0.1-0.8 wt.% Na20, 3-7 wt.% A1203, 400-1300 ppm Ni and 300-1800 ppm Cr (Tables 5 and 6). 2 They are moderately to extremely enriched in incompatible elements: 1-6 wt.~ K20, 100-500 ppm La, 100-600 ppm Ce, 200-11 000 ppm Ba, 100-600 ppm Nb, 100-700 ppm Rb, 300-2000 ppm Sr and 200-1200 ppm Zr; many of these concentrations are in excess of those found in kimberlites. 3 They occur as diatremes with sherbet-glass or funnel-shaped cross-sections, are relatively large in surface area (75-210 acres) and possess multiple intrusive-extrusive phases including a massive magmatic phase, breccias and tufts. 4 They can occur throughout geological history, with the five diamondiferous magmas discovered heretofore erupting in Precambrian, Cretaceous and Miocene times. 5 They are distinctly different from non-diamondiferous varieties of lamproites. The average major- and trace-element compositions of diamondiferous lamproites are compared with those of barren lamproites below (see Tables 5 and 6). Although abundant associated leucite lamproite intrusions are only found in the WKB suite, further mapping is likely to reveal associated differentiated lamproites in the PRA and ARG suites. The eruption of diamondiferous olivine lamproites in the WKB suite is contemporaneous with that of more differentiated leucite lamproites.
Lamprophyre versus iamproite Although the term 'minette' in its present sense, has been used for over 150 years (Bachinski, personal communication, 1984), Gfimbel (1874) introduced the broad textural term 'lamprophyre' or 'shining rock' in reference to a dark dyke rock composed of mafic phenocrysts and groundmass phases. Rosenbusch (1887) gave the term its present significance by defining 'lamprophyres' as dark-coloured dyke rocks of porphyritic texture characterized by a panidiomorphic texture and consisting of a high percentage of mafic minerals as phenocrysts in a fine-grained groundmass composed of the same mafic minerals as well as feldspars and/or feldspathoids. Streckeisen (1980) emphasized the feature that feldspars and/
Lamproites and K-rich igneous rocks
~ =~ ~
~
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~
~
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~
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o
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I46
S. C. Bergman
or feldspathoids, when present, occur only in the groundmass. Rock (1977, 1984, 1985, 1986, 1987) and Wimmenauer (1973) have synthesized the lamprophyre literature and have contributed much to the clarification of many of the problems that have long plagued these mafic-ultramafic rocks. Rock (1986) has chosen to include lamproites as well as kimberlites in the lamprophyre clan. Although the term 'lamprophyre' is extremely broad and includes a host of intermediate, mafic and ultramafic rock types, it is best utilized as a field term, for which it was originally intended, and should not be used to encompass such widely separated rock types as kimberlites and lamproites which possess such distinct evolutions. Clearly, many basalts fulfil the definition of lamprophyre, yet basalts and lamprophyres are petrogenetically distinct (Rock 1987). However, Yagi et al. (1975) described a lamprophyre dike which represents the subvolcanic feeder zone of alkali basalt flows. The original definitions of 'lamprophyre' have emphasized the shallow intrusive dyke character of lamprophyres. Lamproites, as we have seen, characteristically occur as extrusives or shallow intrusives, and it is only the very old occurrences that expose deeper feeder-dyke levels. Clearly, every volcanic edifice can generally be traced to a feeder dyke. As lamproites often possess a feldspathoid (leucite) among the phenocryst population (e.g. SB, WKB etc.), they do not qualify to be called lamprophyres. Tr6ger's (1935) use of the term 'lamproite' as an extrusive equivalent of lamprophyres that are rich in K and Mg (presumably minettes) should be abandoned. As discussed above and below and by Rock (1984, 1987) minettes and lamproites are geochemically and mineralogically distinct, although there are some overlapping features. Therefore minettes and lamproites cannot be related by currently understood processes that could link an intrusive to a coexisting extrusive form. Several, albeit extremely rare, minettes occur as extrusives (e.g. NHB and COL) (see above), but these are not lamproites as we understand them at present. Nevertheless, it is possible that minettes may be nothing more than deep-seated equivalents of extrusive lamproite lavas. The fact that minettes are generally Palaeozoic in age (notable exceptions are the NHB, COL and KAJ minettes of Cenozoic age) and therefore have experienced more erosion than younger lamproites supports this view. More work is required to solve this problem. It is therefore suggested that lamproites as well as kimberlites be excluded from the lamprophyre clan if only to remove some of the inherent genetic biases introduced by associating them.
This opinion is shared by several other workers (J. B. Dawson, personal communication, 1984; R. Mitchell, personal communication, 1985).
Lamproite geochemistry Rather than reproducing individual geochemical analyses, the averages and ranges (represented by standard deviations) of geochemical parameters for various rock suites are discussed in this section, recognizing the inherent biases introduced by such an analysis. The literature sources for individual analyses are tabulated in Appendix 1. Of the 275 analyses available for lamproites, the lamproite group average is weighted toward the MAP, WKB, PRA and LH suites (i.e. 75% of the analyses come from 4 or 20% of the suites). The most apparent feature of the geochemical data is the extreme range in major-, trace-element and isotopic composition of individual lamproite suites and of lamproites as a group.
Major-element geochemistry The average major-element compositions of lamproite suites and other ultrapotassic rock suites are tabulated in Table 4. Univariate frequency-distribution bar charts including all the lamproite major- and minor-element data are illustrated in Fig. 20. These distributions are heterogeneous for all elements and ratios considered, emphasizing the compositional complexity and variability of lamproites. Some distributions are approximately normalor slightly skewed (e.g. K20, P205 and MgO); however, most elements possess frequency distributions that are bimodal or extremely skewed (e.g. SiO2, A1203, TiO2, FeO, CaO, Na20, BaO, H20 and ZrO2). Elemental ratios (e.g. Mg number, K20/ NazO and K20/A1203) display distributions characterized by a high kurtosis. With the exception of a few elements, lamproites overlap in major-element composition with kimberlites and lamprophyres (Table 4). The range in composition of lamproites (expressed as the arithmetic mean_+ 1 standard deviation) overlaps with that of kimberlites for all major elements except K, Si, Mg and A1. However, statistical comparisons of the two groups, lamproites and kimberlites, using a general linear-models procedure, show that they are different in terms of all major and minor elements. Lamproites can be statistically distinguished from lamprophyres, as a group, with respect to all major elements, but most significantly with potassium (general linear model F value, 840) and less so for CaO, FeO*, A1203 (F
Lamproites and K-rich igneous rocks
8oI 6o 5o 3o
FREQUENCY
40 20
I
FREQUENCY
FREQUENCY
40
6
5
io
ib
147
I0
20
_
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_
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20
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30
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20
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,
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'
5
i
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0
i
0.5 1.0 K20/AI20s
30 3O
zb
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ib
'
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~060~
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.
6
I .'0
210 SZ0
410
40
0
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I
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!
I
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I
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0
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o 35
45
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65
75
o'.s
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0 . 0 6 0.12 0.18 0.24 0 . 3 0 Zr02
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Na20 4- K20 AI20~
FIG. 20. Frequency-distribution diagrams for major- and minor-element contents and various elemental ratios in lamproites (volatile free; weight per cent basis).
S. C. Bergman
I48
values of 300-700). Of all the lamprophyres, minettes approach lamproites the closest; however, when non-diamondiferous lamproites are statistically compared with minettes (general linear models Student t test), the two groups differ in A1203, FeO*, CaO, MgO, Na20 , K20 and TiO2 at the 99.9~ confidence level. Figure 21 illustrates the compositions of all lamproites compared with kimberlites and lamprophyres in two ternary projections used for representing kimberlites and related rocks (Cornelissen & Verwoerd 1975). Although a great deal of overlap occurs in both plots, lamproites occupy a relatively K20-rich field in the K20-MgO-A1203 ternary which is not occupied by either kimberlites or lamprophyres. In the FeO-A1203-MgO projection, lamproites are intermediate between kimberlites and lamprophyres. In fact, with exception of Si, P, Ti, K, Fe and Ca, lamproites are compositionally intermediate between kimberlites and lamprophyres (Table 4). Lamproites AI203
i
',-~< ~",
( ;e
-'_;
9 FeO
\ MgO
(a) K20 A
MgO
~ LAMPROITES . . . . . . . . . LAMPROPHYRES ..... LITES
AI203
(b)
FIG. 21. Ternary diagram showing the compositions of lamproites, lamprophyres and kimberlites projected on (a) the A1203 MgO-FeO* and (b) the K20A1203-MgO plane (weight per cent).
have the highest average K20, TiO2 and SiO 2 and the lowest average CaO and FeO contents of any of these mafic-ultramafic alkalic rock types. Lamproites are distinguished from kamafugitic rocks by distinctly higher SiO 2 and lower A1203 and CaO contents as well as by a higher degree of peralkalinity. Figure 22 and Tables 4 and 5 emphasize this K20-SiO2 distinction between lamprophyres, lamproites and kimberlites as a group. Alkali basalt compositions greatly differ from lamproites in that they are significantly enriched in A12Oa, Na20 and CaO, and depleted in K20 , MgO, SiO2, P205, BaO and ZrO 2 (Tables 4 and 5). Figure 22 illustrates the covariance of selected major elements and their ratios for individual lamproite suites and compares the average compositions of kimberlites, alkali basalts, lamprophyres and several lamproite minerals. With few exceptions, the size of an individual lamproite suite field is positively correlated with the number of analyses of the given suite, a feature which emphasizes the extreme variability in lamproite geochemistry. It should be noted that in nearly every plot, diamondiferous suites overlap with each other in composition. In a plot of the degree ofultrapotassic versus perpotassic character (Fig. 22(a)) most lamproite suites individually show a positive correlation, as displayed by most rock types in Fig. l, although several (e.g. ARG) show a slight negative correlation. As observed in Fig. 22(b), some of the WKB lamproites are the most perpotassic of all lamproites whereas the MAP suites contain some of the least perpotassic lamproites. Figure 22(c) displays the extreme ranges in both CaO and K20 contents for given suites; diamondiferous suites are characterized by the lowest CaO contents. The nature of magmatic differentiation in lamproite suites is documented in a plot of Mg number versus SiO2 (Fig. 22(d)). In addition, the generally primitive character of lamproites is evident; lamproites possess extremely high Mg numbers for their relatively high SiO2 contents compared with nearly all other rock types. Although some suites display decreasing Mg numbers with increasing SiO2 (e.g. WKB, HOL and LH), a trend that is expected to result from the fractionation of mafic phases such as olivine, diopside and phlogopite, many lamproite suites possess trends that display very limited Mg number variation (less than 10) with extremely large changes in SiO2 (e.g. ARG, PRA, KAM and NWI). Therefore some mafic phase fractionation can explain the SiO2-Mg number variations displayed within some suites, but other types of differentiation (e.g. vapourphase transfer) or, alternatively, differences in source composition, pressure and/or temperature
Lamproites and K-rich igneous rocks are required to explain the chemical variations within other lamproite suites. Figure 22(e) illustrates the covariance of TiO2 and K20. Lamproite suites can be divided into two groups according to their TiO2 contents: one group with TiO2 < 2 tool.% includes COR, NWI, PEN, LH and MAP, and the other with TiO2 > 1.5 mol.% includes WKB, ARG, BOB, SB, MBAY, GSB, HP and HOL. Interestingly, the NWI, COR, PEN, LH and MAP suites also share the feature of occurring in a young palaeoorogenic zone and are most probably depleted in TiO2 (as well as Nb etc.) because of a TiOzdepleted source, a feature shared by orogenic andesites and related rocks. The extreme enrichments in TiO 2 displayed by the WKB, CHE, BOB and SB suites must represent TiOz-enriched source rocks.
Normative composition The average CIPW normative compositions for most lamproite suites are given in Table 7, bearing in mind the limitations of the CIPW normative calculations in describing phlogopiteand amphibole-rich rocks. The average lamproite consists of nearly equal amounts (by weight) of salic and femic normative minerals. Almost half (43%) of the 295 lamproites that have been included in this data set are quartz normative; in fact, all suites for which more than five rocks have been analysed contain at least one quartznormative rock. Only eight of the 21 suites contain rocks with normative leucite, and normative K-metasilicate is present in small amounts because of the Al-depleted character of lamproites. All suites but two contain at least a few rocks with normative feldspathoids. Lamproites as a group are characterized by significant normative orthoclase (16-59 wt.%, average 35 wt.%) and femic minerals (33-81 wt.%, average 51 wt.%), but relatively low normative albite (012 wt.%, average 5 wt.%) and anorthite (0-6 wt.%, average 1 wt.%) relative to nearly all other igneous rocks. As with most peralkaline rocks, almost all lamproites contain normative acmite (average 2.5 wt.%). Compared with minettes, lamproites contain similar average amounts of normative quartz, orthoclase and nepheline but significantly lower albite and anorthite, slightly lower clinopyroxene, and slightly higher orthopyroxene, acmite, olivine and leucite. Lamproites as a group are somewhat more enriched in femic minerals than minettes. Despite the common occurrence of m o d a l leucite in lamproites, many lamproite magmas tend to differentiate towards quartz-normative residua. This is based on the natural evidence of
I49
the normative compositions of interstitial glasses and whole-rock samples of glassy lamproites (e.g. GSB (Sheraton & Cundari 1980), LH (Kuehner et al. 1981) and WKB (Wade & Prider 1940; Prider 1982)). In fact many leucite-bearing lamproites are SiO 2 saturated or oversaturated. This differentiation trend can be explained by a small amount of phlogopite fractionation and additionally by the experimental work of Luth (1967) in the system kalsilite-forsterite-silicawater which demonstrated that the phlogopite liquidus surface bridges the forsterite-kalsilite thermal divide at PHzo > 1 kb.
Trace-element chemistry A summary of the trace-element geochemistry of lamproites, compared with that of kimberlites, lamprophyres and alkali basalts, is given in Table 6 and Fig. 23(a). All trace elements show an extreme degree of variance with coefficients of variation typically more than 50% and often more than 100%. If these extreme variations are ignored, some general conclusions can be drawn. On average, lamproites have the highest incompatible-element contents (e.g. Rb, Sr, Ba, U, Zr and rare-earth elements (REE)) and alkali basalts the lowest, whereas kimberlites have the highest compatible-element contents (e.g. Ni, Cr, Co, Sc, Zn and Cu) and alkali basalts generally the lowest. Lamproites are more enriched in both compatible and incompatible elements compared with other ultrapotassic rocks. With respect to the average K/Rb ratios, the various rock groups are ranked as follows: alkali basalts (350), nondiamondiferous lamproites (220), calc-alkaline and alkaline lamprophyres (180), kimberlites (160) and diamondiferous lamproites (85). The average REE contents of lamproites, lamprophyres and kimberlites all overlap to some degree (Table 6 and Fig. 23(a)), and all are markedly light REE enriched and heavy REE depleted compared with alkali basalts. Although a large degree of overlap exists, average lamproite REE contents generally exceed those of kimberlites. The average ratios of light REE to heavy REE (La/Lu)on are ranked as follows: lamproite (100), kimberlite (90), ultramafic lamprophyre (47), alkaline and calc-alkaline lamprophyre (27) and alkali basalt (11), showing that lamproites contain the smallest quantities of heavy REE (relative to light REE) compared with these other rock types. Since the degree of elemental variation for these rock types is so great, overlap is generally observed in comparing the trace-element contents of the various rock groups. Nevertheless, the
S. C. Bergman
15 ~
2.0]
t
DIAMONDIFEROUS SUITES 9 AVE, KIMBERLITE (550) 9 AVE. LAMPROPHYRE (800) ",9 AVE. ALKALI BASALT (4230)
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DIAMONDIFEROUS SUITES AVE. KIMBERLITE (550) AVE. LAMPROPHYRE (800) AVE. ALKALI BASALT (4230) AVE. LAMPROITE PHLOGOPITE (229) AVE. LAMPROITE AMPHIBOLE (97)
K20/AI203 = 1
=
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(b)
10 AI203 (mole %)
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DIAMONDIFEROUS SUITES 9 AVE. KIMBERLITE (550) 9 AVE. LAMPROPHYRE (800) AVE. ALKALI BASALT (4230) I-I AVE. LAMPROITE PHLOGOPITE (229) O AVE. LAMPROITE AMPHIBOLE (97)
.:..< .;-..-----
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Lamproites and K-rich igneous rocks gO-
,
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~ Z BOB
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o r,
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40-
9 9 ~, I-1 O
DIAMONDIFEROUS AVE. KIMBERLITE (550) AVE. LAMPROPHYRE (800) AVE, ALKALI BASALT (4230) AVE. LAMPROITE PHLOGOPITE (229) AVE. LAMPROITE AMPHIBOLE (97)
;o
3O
(d)
WKBf
"::. . . . . " .... . .....-'/ - ......... .--
~
;o
;o
SIO 2 (mole %)
109 9 zx 17
WKB
DIAMONDIFEROUS SUITES AVE. KIMBERLITE (550) AVE. LAMPROPHYRE (800) AVE. ALKALI BASALT (4230) AVE. LAMPROITE PHLOGOPITE (229) E. LAMPROITE AMPHIBOLE (97)
COR SB HP YAM GSB MBAY PEN NWI PPN MAP HOL LH ENO CHE NSW
~
~
..............
-----~ ~ - - - r ,'----r ---.-.-"~r
_o
'...
"~
;
,~ _
..-~---~
>-~--53
0
( e )
K20 (mole %)
FIG. 22. Plots illustrating the compositional space occupied by various lamproite suites, with average rock and mineral compositions plotted for comparison (see text and Tables 2 and 3 for abbreviations): (a) KzO/A1203 versus K20/Na20; (b) K20 versus A1203; (c) K20 versus CaO; (d) Mg number versus SiOz; (e) TiO2 versus K20.
trace elements that are most useful in distinguishing lamproites from kimberlites are as follows: (a) Rb, Sr, Ba, Zr and the K/Rb ratio; (b) Ni, Cr, Co and the Ni/Cr ratio. Kimberlites generally have Rb < 150 ppm, Sr < 1200 ppm, Ba < 3000 ppm, Z r < 8 0 0 ppm and K/Rb<150, whereas lamproites generally have R b > 150 ppm, Sr> 1200 ppm, Ba > 3000 ppm, Zr > 800 ppm and K/ R b > 150. With respect to the compatible elements, kimberlites generally have Ni > 500 ppm, Cr>800 ppm, Co>50 ppm and Ni/Cr>0.7, whereas lamproites generally have Ni < 500 ppm, Cr<800 ppm C o < 5 0 ppm and Ni/Cr<0.7. In general, the absolute concentrations of both compatible and incompatible elements are larger
in kimberlites and lamproites than in alkali basalts and lamprophyres. Individual lamproite suites show widely-ranging trace-element abundances. The Indian lamproites (CHE, GDW) are the most enriched in REE of all lamproites yet contain typical lamproite signatures with respect to the other trace elements. Lamproite suites possessing the lowest LIL-element contents include KAM, N W I and ARG. On the basis of the available data diamondiferous lamproites display no statistically-significant differences relative to barren lamproites in their trace-element systematics, with the exception of an enrichment in Nb and Ga and in the
S. C.
I52
Bergman
TABLE 7. Average CIP W normative compositions of various lamproite and related rock suites Suite
n
q
ARG 3 BOB 1 CHE 2 COC 2 COR 2 EVD 4 GSB 1 HLM 2 HOL 12 HP 4 KAM 9 LH 24 MAP 99 MBAY 6 NWI 9 PEN 1 PRA 24 SB 13 WKB 51 Lamproite 284 Minette 50 Lamprophyre728
6.1• 4.4 --1.1• ---6.3• -2.8• 0.6• 5.0• 4.3• 0.2• 4.4 5.8• 5.4• 5.5• 3.5• 2.7• 0.6•
n
ol
Suite
ARG 3 BOB 1 CHE 2 COC 2 COR 2 EVD 4 GSB 1 HLM 2 HOL 12 HP 4 KAM 9 LH 24 MAP 99 MBAY 6 NWI 9 1 PEN PRA 24 SB 13 WKB 51 Lamproite 284 Minette 50 Lamprophyre728
17.9• 3.5• 20.9• 1.3• 26.7• 11.3 14.3• 22.7• 25.2• 8.6• 6.5• 7.4• -11.1• -17.4• 0.7• 7.8• 9.9• 6.2• 10.4•
or
25.8• 4.6 15.8• 26.4• 59.1• 21.8• 54.2 47.4• 39.9• 31.6• 36.8• 41.9• 38.0• 49.6• 45.1• 56.1 16.5• 41.8• 38.1• 35.0• 34.5• 13.1•
ab
0.5• . 1.2• 11.3• 5.0• 1.0• -12.5• 0.1• 0.5• 7.0• 0.9• 11.0• . 3.0• 2.4• 7.8• 0.3• 5.4• 13.6• 12.9•
an
lc
0.6• .
-.
--
.
.
.
. 13.4• 0.4 0.1• -4.4• 10.0• -. --0.4• -0.9• 2.0• 0.3• 1.5•
. 2.7• -0.5• 1.5• 0.3• 1.3• 5.6• 14.6•
hm
sp
rt
3.1• -3.4• 3.9• 0.9• 2.8• -2.1• --
5.7•
-7.2 0.3• ---
--------
-5.9 --
1.5• 0.2• 1.0• 0.1• 1.2• 2.8 1.8• -0.6• 1,3• 3.6• 5.4•
2.8• --1.4• 1.8• 1.2• 0.6• --4.7• 0.6• 3.1• 1.7• 1.3• 1.4•
--1.5• 1.7• 0.7• 0.1•
-------0.02• ---
m
0.1• 2.3• 3.5• 1.1• 0.2• 0.03•
0.6• 0.1• 1.0• 0.3• 0.02• 0.1•
ap
2.2• 3.9• -7.4 -4.4• -4.9• 5.6• 2.4• 5.6• 3.1• 3.7 0.5• 2.3• 2.3• 2.5• 0.1• 0.4•
3.0• 4.1 4.9• 1.4• 1.5• 7.0• 3.6 6.6• 2.9• 1.8• 3,0• 3.6• 2.5• 6.5• 2.9• 4.3 2.5• 4.1• 2.6• 2.9• 2.7• 1.7•
cc
2.6• -13.0• 0.5• 2.9• -0.2 0.1• 10.0• -0.1• 0.5• 1.2• 0.8• -6.2 4.4• 3.9• 2.6• 2.2• 2.1• 2.2•
hy
di
1.2• --
-4.2 -1.7• 1.0• 0.3• 1.7• 0.1• 1.1• 1.0• 0.4 0,2• 0.4• 2.6• 0.8• ---
1.3• -0.2• -0.8• 1.5• 1.7• 6.7•
il
ac
0.6• 4.3 . --
. 5.7• . 3.2• ----1.2• 0.6• 0.5•
--
mt
11.3• 1.6• 2.4• 3.8• 6.6 3.5• 6.4• 5.8• 3.0• 2.6• 2.3• 7.3• 2.4• 3.2 3.7• 8.6• 5.0• 4.3• 3.3• 4.4•
ks
.
. 0.7• . 4.8• -6.1• --2.6• 0,6• 2.8• .
ne
5.2• --25.3• 6.4• 14.1• 10.3 0.6• 1.0• 8.7• 9.7• 13.1• 8.3• 3.7• 13.5• -2.4• 4.6• 6.5• 9.8• 14.9• 18.7•
26.9• 25.4 43.3• 15.1•
2.8• 8.5• 23.8• 8.7• 6.0• 14.2• 17.4• 11.4• 18.4 35.2• 15.3• 15.9• 13.7• 6.5• 3.7•
Z Foid
Z Salic
~ Femic
0.6 4.1 -5.7 0.3 16.6 5.7 -3.7 2.5 6.0 12.5 1.0 1.6 3.2 0.4 0.8 0.8 3.8 4.3 2.0 8.4
33.9 54.4 19.0 44.2 65.3 44.1 54.9 66.8 40.6 32.1 55.0 55.5 58.0 54.8 52.4 60.9 26.5 56.6 45.5 48.7 59.0 49.9
66.1 48.6 81.0 55.8 34.7 55.9 45.1 33.4 59.4 67.9 45.0 44.5 42.0 45.2 47.6 39.1 73.5 43.3 54.5 51.3 41.0 50.1
Means • standard deviations are given for n > 2 analyses; weight per cent normative minerals; only those minerals w h i c h occur in significant concentrations (i.e. more than 0.1 wt) are tabulated here. N o r m s are calculated using the reported FeO and Fe203 contents.
compatible elements (Cr, Ni, Co). However, diamondiferous lamproites possess slightly higher average light REE (La, Ce), Rb, Ba and Ta contents and have higher La/Lu (1400 compared with 800), Ba/Sr (6 compared with 3) and much lower K/Rb ratios (85 compared with 220) relative to barren varieties. SB magmas contain among the highest REE, Sr and Ba contents of all lamproite suites, but possess more typical lamproite concentrations of other LIL elements
(e.g. U, Th and Zr) and among the lowest Rb contents of any lamproite. WKB are likewise exceptionally enriched in LIL elements. Stable and radiogenic isotopic geochemistry of ultrapotassic magmas Sr and/or Nd isotopic ratios have been measured in rocks from eight lamproite suites (over 160 samples) and in eight potassic-ultrapotassic rock
Lamproites and K-rich igneous rocks
i
153
i
PN~
/ c~
=, E
...... :~.. . :<
I
.c
_~i,.. o
-
~
~.'.';/../~ ,.>/ /. X
!/
i~
~.i " ~ . . ~
/
~ ~
\
J
~,~/
o
.J
~
o
?-
d
~ ~ . / 9
.
~
.5
w. i
,
,
i
,
,
,
d
i
'
'
,
,
i
,
,
,
,
,
,
,
,
,
d
c~
PN~,~,LIPN~ L
~o_
/
,
G
8
0
/t
"-
Cl
~
~'" ~'o ~c
eg ~am E~
o~
~,
~
~
._~
o
~ , ~
o-~
~
~
~
.
~-
I
0
~.,
~oo~
S. C. Bergman
154
suites other than lamproites (over 200 samples); the results are summarized in Tables 8 and 9 and Fig. 23(b). Whereas ultrapotassic rocks other than lamproites generally fall in the ranges 87Sr/ 86Sr~=0.704-0.708 and 143Nd/144Ndt=0.5122 0.5126 (where t indicates present-day ratios), lamproites are characterized by much more and much less radiogenic compositions respectively, with 8VSr/S6Srt= 0.705-0.718 and 143Nd/ 144Ndt=0.5112-0.5123. Lamproite Nd-Sr isotopic data fall along two district trends on an end-eSr plot (Fig. 23(b)). One trend, which includes the SB and LH suites, is characterized by slightly radiogenic 87Sr/86Sr ratios with extremely nonradiogenic 143Nd/l~4Nd ratios, and another shallow trend, which includes GSB, MAP and WKB rocks, is characterized by extremely radiogenic 87Sr/86Sr ratios and non-radiogenic 143Nd/ 144Nd ratios (Table 8). These data indicate that the source regions of these mantle-derived magmas were characterized by very old rocks (more than (1-3) • 10 3 Ma old) with Rb/Sr and Nd/Sm ratios greater than those of bulk Earth, the former source material possessing slightly lower Rb/Sr ratios or being younger than the latter. Most important is the undeniable conclusion that the source regions of primary mantle-derived dia-
mondiferous melts (e.g. Ellendale) may be extremely old and enriched in Rb relative to Sr and Nd relative to Sm. In the past it has generally been assumed that basalts possessing relatively high 87Sr/86Sr ratios (above 0.7050) (Faure & Powell 1972; Faure 1977) must have been contaminated by radiogenic continental crustal material during their ascent. However, combined with the Sr-Nd isotopic data on some kimberlites and included megacrysts and xenoliths (Menzies & Murthy 1980a, b; Erlank et al. 1982; Hawkesworth et al. 1983; Kramers et al. 1983; Smith 1983a, b), the lamproite data indicate that contamination by continental crust is not required to explain their isotopic systematics and that extremely enriched portions of the sub-continental upper-mantle lithosphere exist. Oxygen isotopic studies have thus far been largely limited to ultrapotassic rocks other than lamproites, with the exception of the LH and GSB suites. Studies by Garlick (1966), Taylor (1968), Barberi et al. (1975, 1978), Taylor & Turi (1976), Turi & Taylor (1976), Taylor et al. (1979, 1984), Ferrara et al. (1980, 1985), Kuehner (1980) and Kyser et al. (1981) demonstrate that unaltered lamproites generally possess slightly higher 8180 ratios (7-12) compared with typical mantle-
TABL~ 8. S r a n d N d isotopic c o m p o s i t i o n s o f l a m p r o i t e s Locality*
87Sr/86Srt n
Yr
WKB
LH MAP PRA NWI GSB SB NSW WKB SB MAP
8
0.7169
13 6 17 53 6 34 1 4 7 4 ~ 15 6 8
0.7142 0.7133 0.7063 0.7057 0,7048 0.7096 O.7O79 0.7187 0.7096
0.7171 0.7061 0.7188
Range
t
n
4SN d / t 44N d t
Yc
Range
0.7047-0.7050 0.712-0.721 0.7104-0.7187 0.7114-0.7165 0.7056-0.7070 0.7053-0.7078 0.7036-0.7058 0.7064-0.7132 -0.7171-0.7221 0.7092 0.7098 0.7059 0.7063 0.7050-0.7057 0.7113-0.7209 0.7059-0.7063 0.7175-0.7208
8 13
0.51206 0.51119
0.5120-0.5123 0.5114-0.5111
19
0.51193
0.5118-0.5121
7 4 9 15 6 8
0.5111
0.51111-0.51117 0.5113-0.5115 0.51252-0.51264 0.5117-0-5121 0.5113-0.5115 0.5112-0.5113
0.51189 0.5114 0.5112
Reference
1 2 3 4 5 2 6 2 7 8 9 10 11 12 13 13 14
* See Tables 2 and 3 and text for abbreviations. 1, Kaplan et al. 1967; 2, Powell & Bell 1970; 3, Nixon et al. 1984; 4, McCulloch et al. 1983a, b; 5, Jaques et al. 1984a; 6, V o l l m e r e t a l . 1984; 7, Bolivar 1977; 8, Smith 1983b; 9, V e n t u r e l l i e t a l . 1984a, b; 10, Collerson & McCulloch 1982; 11, Mitchell & Hawkesworth 1984; 12, Mitchell, personal communication, 1984; 13, Fraser et al. 1985; 14, Nelson et al. 1986.
Lamproites and K-rich igneous rocks
[55
TABLE 9. Sr and N d isotopic compositions o f ultrapotassic-potassic rocks other than lamproites Locality*
Rock type
87Sr/S6Srt n
HM NHB TAN
Shonkiniteetc. Minette, monchiquite etc. Kamafugites, ugandite, kivite, shoshonite etc.
TUSC
Kamafugites
LAC
Basalts
RP
Basalts, intermediate and felsicrocks
CSN
Ultrapotassic andpotassic basalts DSV Leucitite NSW Leucitite S African kimberlites Group I Group II N American kimberlites Group I Finsch II
2
Range
143Nd/144Ndt n
:~
Reference
Range
10 0.7078 0.7072-0.7087 12 0.7062 0.7041-0.7081 104 8 37 6 1 1 9 9 22 13 14 5
0.7059 0.7036-0.7083 0.7047 0.7046-0.7048 0.7069 0.7046-0.7103 0.7035-0.7096 0.7057 0.7054-0.7059 0.7112 -0.7104 0.7054 0.7046-0.7062 __ 0.7044-0.7047 0.7092 0.7085 0.7068-0.7098 0.7097 0.7088-0.7111 0.7064 0.7061-0.7068
2 11
0.7119 0.7119-0.7124 0.7053 0.7050-0.7057
24 28
0.7057 0.7088
0.703-0.706 0.708-0.710
13 0.7948 0.704-0.705 7 0.7100 0.7095-0.7119
4 0.5127 0.51270-0.51272 12 0.51250 0.51211-0.51276
9
__
0.51260-0.51265
8 0.51212 0.51209-0.51216
11 0.5118
0.5116-0.5119
8 0.51276 0.51257-0.51282 6 0.51219 0.51213-0.51234
2 3 4 5 18 6 7 8,9 10 11 12 13 14 14 16 15 15
6 0.51219 0.51217-0.51224
* See Tables 2 and 3 and text for abbreviations. 1, Powell & Bell 1970; 2, Bell & Powell 1969; 3, Vollmer & Norry 1983a; 4, Vollmer & Norry 1983b; 5, Rock 1976; 6, Holm& Munksgaard 1982; 7, Vollmer 1976; 8, Hoefs &Wedepohl 1968; 9, Hurleyetal. 1966; 10, Staudigel & Zindler 1978; 11, Powell & Bell 1974; 12, Cox et al. 1976; 13, Rogers et al. 1985; 14, Van Kooten 1981; 15, Smith 1983b; 16, Fraser et al. 1985; 17, Nelson et al. 1986; 18, Aoki & Kurasawa 1984. derived material and support the following conclusions: (a) in regions such as the Roman Province (RP), where isotopically heavy crust exists (8180---+15 to +25), contamination by crustal material is required to explain the stable isotope compositions, and Sr and Nd isotopic compositions are also probably modified by crustal contamination (Taylor et al. 1984); (b) some lamproites (e.g. GSB) and other ultrapotassic rocks (e.g. TAN) display primary mantle 8180 signatures (about + 6.5) when corrected for nearsurface alteration effects (Taylor et al. 1984); and (c) the source regions of some ultrapotassic magmas do not substantially differ in their 8180 compositions from those of typical mid-ocean ridge basalts (MORBs) (about + 5.5) or kimberlites and their xenoliths ( + 5.5 to + 7.5) (Taylor 1968; Kyser et al. 1981, 1982). It is also possible that lamproite source regions are anomalous in their 8180 systematics compared with other mantle rocks.
Pb isotopic studies are similarly largely restricted to ultrapotassic rocks other than lamproites. Most Pb isotopic data exists for the TAN, RP, CSN and DSV suites (Vollmer 1976, 1977; Van Kooten 1981; Vollmer & Norry 1983a, b) and suggest that, in the first case, a P b - P b model age of 500 Ma dates a mantle-source homogenization event (metasomatism?). Pb isotopic systematics of the RP rocks, however, are most consistent with the mixing of two distinct reservoirs; this interpretation is also consistent with the oxygen isotope data noted above. Van Kooten's (1981) Pb isotope work suggests a rather recent mantle-homogenization event on the basis of large variations in 238U/2~ and 232Th/ :~ ratios with rather restricted ranges in 2~176 2~176 and Z~176 ratios, perhaps as recent as 10 Ma ago. Pb isotopic studies on lamproites from L H and SB are currently under-way at the Open University (e.g. Fraser et al. 1985) and on W K B rocks at the
156
S. C. Bergman
Australian National University (e.g. Nelson et al. 1986) and tentatively indicate model ages of (1.5-3.0) x 103 Ma for these mantle source regions. Further radiogenic and stable isotope studies are clearly required for a better understanding of the petrogenesis of lamproites. For example, are the two trends recognized thus far in terms of eSr-end systematics simply the result of too few samples, and will these two trends develop into one broad trend for more lamproite suites? Are all lamproites characterized by 'enriched' Sr and 'depleted' Nd isotopic signatures in contrast with the vast majority of isotopically 'depleted' signatures commonly observed in alkali basalts and MORB (Hofmann et al. 1979)?
Lamproite mineral chemistry The mineralogy and mineral chemistry of lamproites has been recently reviewed by Mitchell (1985) and Wagner & Velde (1985), to which the reader is referred for detailed accounts. Therefore only the most important features, abstracted from Mitchell's review, will be highlighted in this section, and the mineral chemistry of lamproites will be compared with that of associated rock types. An important feature of lamproite mineral compositions is that they reflect the exotic compositions of the magmas from which the minerals crystallized. A tabulation of literature containing lamproite-mineral chemical data can be found in Appendix 2. The average lamproite and related rock mineral compositions are provided in Table 10.
Olivine Olivine (commonly pseudomorphed by secondary phases) occurs in two generations (large anhedral macrocrysts (dimensions of more than 1 mm) and small strain-free crystals or aggregates (dimension of less than 1 ram)) in the more mafic members (and some felsic members, e.g. MAP) of nearly all lamproite suites. The compositional range observed for both types is Fo77-93 with most about Fo89. NiO contents are typically in the range 0.15-0.6 wt.~ whereas CaO is generally less than 0.2 wt.~. The Mg numbers and CaO and NiO contents of lamproite olivines overlap those observed in kimberlites and lamprophyres but are distinct from those in alkali basalts and related rock types.
Orthopyroxene Orthopyroxenes ranging in composition from En61 to En94 (Cr20 3 ~ 0.3 wt.~; CaO < 1.3 wt. %;
TIO2<0.4 wt.~) occur in the MAP, LH and WKB suites. Their extremely wide range in Mg number and texture (commonly rimmed by phlogopite) suggests that some may be xenocrystic; however, their extremely low A1203 contents (typically less than 1.0 wt.~), are much lower than those observed in orthopyroxenes from marie igneous rocks and suggest some type of lamproite affinity. Therefore many of the Fe-rich orthopyroxenes most probably originate by disaggregation of phlogopite-pyroxene or other cognate xenoliths.
Clinopyroxene Either diopside or salite has been observed in virtually every lamproite; most clinopyroxenes are nearly pure diopside-hedenbergite solid solutions (22-26 wt.~ CaO). The AI203 contents in lamproite clinopyroxenes are remarkably low (average 0.5 wt.~) relative to clinopyroxenes of other marie and ultramafic volcanic rocks (less than 5-10 wt.% A1203), except those in some kimberlites (1-2 wt.~); in fact, many lamproite clinopyroxenes have insufficient A1 to fill all the tetrahedral sites. Anomalously high A1203 contents in some lamproite diopsides (less than 1.0 wt.~) are interpreted as indicating a xenolithic relationship or contamination, based on the compositions of grains in both the groundmass and phenocryst populations of the better-studied localities (i.e. LH, WKB and HOL). Lamproite diopsides typically contain 0.4-3.5 wt.% TiO2 (HP with the highest content, PRA with intermediate content, and MAP and LH with the lowest content), 0.0-1.5 wt.~ Cr203 (LH and HP with the lowest content, and WKB and PRA with the highest content) and 0.0-1.5 wt.% Na20 (generally 0.3-0.4 wt.~). Of all the marie and ultramafic rock types, lamproites contain diopsides with the lowest Na20 contents. It is therefore remarkable that aegirine occurs on the rims of diopside grains, especially those within vesicles, in LH rocks (Kemp & Knight 1903; Carmichael 1967; Kuehner 1980) and other lamproites (Wagner & Velde 1985). Clinopyroxenes found in some kimberlites and some ultramarie lamprophyres are similar in composition to those found in lamproites.
Alkali amphiboles K-Ti-richterite is a common alkali amphibole observed as late-stage groundmass grains in most lamproite suites (e.g. HOL, WKB, LH, SB, HP, PRA and GSB), although K-Ti-Mg-arfvedsonites and K-riebeckites also occur (e.g. at PP,
Lamproites and K-rich igneous rocks TABLE 10.
I57
Summary of mineral chemistry of lamproitephases compared with those of other rock types* Olivines Groundmass, phenocryst phases
n SiO2 TiO2 A1203 Cr203 FeO* MnO NiO MgO CaO Total Mgno.
Xenocryst, xenolith phases
LAMP
LAMPH
KIMB
AOB
ALK
KIMB
AOB
LAMPH
58 40_+1 0.0 0.05 0.03 10_+3 0.17 0.33_+0.1 49 _+2 0.17_+0.2 99.9 89_+3
13 40-+1 0.03 0.03 0.04 12_+5 0.24 0.39_+0.7 47 _+5 0.t6_+0.1 100.1 87_+6
150 41+_1 0.02 0.03 0.06 9_+2 0.11 0.32_+0.1 50 _+2 0.07_+0.8 100.2 91_+3
78 38_+2 0.03 0.03 0.03 22_+9 0.35 0.13_+0.11 39 _+8 0.30_+0.1 99.8 75_+12
14 39 0.07 0.04 0.06 20 0.30 0.24 40 0.42 99.8 77
552 41+2 0.03 0.04 0.04 9_+3 0.11 0.35_+0.1 50_+ 3 0.5_+0.6 100.2 90+3
319 40_+1 0.02 0.17 0.04 12_+5 0.17 0.28_+0.1 47 _+4 0.13_+0.1 100.0 87_+6
51 40_+1 0.07 0.02 0.05 10_+2 0.15 0.34_+0.14 49_+ 2 0.07_+0.6 100.2 90+2
Orthopyroxenes Xenocryst, xenolith phases
Groundmass, phenocryst phases
n SiO2 TiO2 A120~ Cr203 FeO* MnO NiO MgO CaO Na20 Total Mgno.
LAMP
LAMPH
INT
LAMP
LAMPH
KIMB
AOB
7 55_+2 0.27_+0.2 0.9_+0.6 0.29 12_+6 0.27 -30 1.0_+0.4 0.11 99.7 81_+10
7 57_+4 0.31_+0.1 1.31_+0.5 0.9 5_+ 1 0.10 -33 2.2_+0.8 0.12_+0.1 99.6 92_+2
37 53 0.17 0.50 -21 0.56 -24 1.2 -100.1 67
6 56+_2 0.07 1.3_+0.9 0.32_+0.3 7_+6 0.24 0.12 34_+ 5 0.6_+0.3 0.10_+0.8 99.8 89_+10
95 56_+ 1 0.18_+0.1 2.7_+1 0.36_+0.2 7_+2 0.16 0.10 33 _+2 0.9_+0.4 0.15+_0.1 100.4 89+_4
657 57_+2 0.10_+0.1 1.4_+1 0.33_+0.2 6_+42 0.13 0.09 34_+ 3 0.64_+0.4 0.13_+0.1 100.1 90_+6
380 54+_2 0.15_+0.1 3.9_+1.5 0.36_+0.2 8_+4 0.17 0.08 32_+ 3 0.88_+0.7 0.10+0.8 99.8 88_+7
Amphiboles Xenocryst, xenolith phases
Groundmass, phenocryst phases
n SiO 2 TiO2 A1203 ChO3 FeO* MnO NiO MgO CaO Na:O K20 BaO Total Mgno.
LAMP
LAMPH
INT
AOB
ECL
ATP
ALK
KIMB
AOB
LAMPH
118 53+2 4.5+2 0.8_+1.7 0.03_+0.5 8__+7 0.39_+1 0.04_+0.4 17_+5 5-+2 4.9_+1.5 4.2-+1.6 0.17_+0.2 97.1 78_+20
53 40_+3 4.3+1 13_+3 0.02_+0.2 13_+4 0.41_+0.5 -12_+3 11+2 2.5_+1 1.4_+0.6 . . 97.6 62_+13
21 44 3.2 11 -15 0.30 -12 11 1.7 0.8 . 98.8 60
27 40 5.2 13 0.4 12 0.17 -12 11 2.8 1.4 . 97.9 64
91 50 0.2 9 0.3 12 0.12 0.11 13 8 3.6 0.29 . 97.0 65
31 49 1.2 9.2 1.0 3.3 0.07 0.10 20 12 2.4 0.31 . 97.2 91
19 41 2.7 12 0.8 18 0.7 -9 11 3.0 1.4 . 98.0 48
55 49_+7 0.9_+1 7.7_+7 0.6_+0.7 5.2_+4 0.10 0.12 19_+4 9___2 3.3_+0.7 2.7_+1.9 . 97.5 86_+11
219 41_+2 3.7_+1.9 14_+2 0.8-+0.9 9.4-+4 0.16 0.04 14_+3 11_+1 2.9-t-0.6 1.2_+0.6
28 40+2 4.1+1.6 14_+1.5 0.4-+0.6 10-+3 0.17 0.06 13_+3 11_.+0.7 2.4-+0.6 1.5+_0.8
97.9 72+_13
97.6 70+_10
.
S. C. Bergman
I58
TABLE 10. (cont.) Summary of mineral chemistry of lamproitephases compared with those of other rock
types* Clinopyroxenes Groundmass, phenocryst phases LAMP n SiO2 TiO2 A1203 Cr203 FeO* MnO NiO MgO CaO Na20 Total Mgno. wo en fs
LAMPH KIMB
158 128 53• 50+3 1.0+0.6 1.6• 0.51• 4.2_+2.7 0.19_+0.3 0.14• 3.5• 6.8• 0.11_+0.7 0.21• 0.06+0.4 0.12• 17+2 14• 24• 22• 0.46• 0.8• 99.7 99.8 90• 78• 47.5 46.2 47.0 42.4 5.5 11.4
188 54• 0.42• 3.5• 0.8-1-1 4.2+2 0.11• 0.03• 16• 19• 2.2• 99.7 87+8 42.0 50.4 7.5
Xenocryst, xenolith phases
AOB
ECL ATP INT ALK LAMP
KIMB
175 48• 2.4• 6.0• 0.2_+0.2 8.5-+3 0.19• 0.03• 12• 21• 1.0• 99.5 72+13 47.1 38.0 14.9
86 55 0.19 8.2 0.8 6.8 0.5 0.4 9.2 14 6.0 99.6 69 44.2 39.0 16.7
1057 724 54• 51• 0.28•177 4.6_+4.1 6.0• 1.0+0.9 0.7• 3.6• 5.0• 0.10 0.13 0.6 0.4 16• 15• 18• 20+2 2.6• 1.1• 99.9 99.7 88• 84• 42.4 44.3 50.9 47.1 6.7 8.6
115 52 0,39 4.8 0.72 2.8 0.8 0.4 16 22 1.0 99.8 91 46.5 48.8 4.7
31 52 0.57 2.6 0.12 8.5 0.25 0.3 15 21 0.37 99.6 77 42.2 44.2 13.6
135 49 1.1 4.8 0.10 10 0.37 0.6 12 22 1.3 99.7 67 47.2 35.5 17.3
24 54• 0.55• 1.2• 0.9• 4.8• 0.29 0.8 16• 23• 0.87• 100.0 85• 46.6 45.6 7.8
AOB
LAMPH 233 53• 6.0• 0.5+_0.6 5.2• 0.11 0.9 14_+3 18+3 2.5+__1.8 99.8 83• 42.9 47.2 9.9
Spinels Xenocryst, xenolith phases
Groundmass, phenocryst phases
n SiO2 TiO2 A1203 Cr203 FeO* MnO NiO MgO V205 Total
LAMP
AOB
KIMB
LAMPH
25 0.2 9• 3• 37• 41• 0.9• 0.1 • 8• -98
37 0.2 18• 5• 7• 61_+14 0.7+0.4 0.1 • 5• 0.4• 97
253 29 0.2 0.2 10• 10• 9_+12 8• 2 0 + _ 2 3 17• 48+23 59+21 0.6• 0.5• 0.1 • 0.2• 11• 7• 0.5• 0.2• 97 97
ATP
LAMP
AOB
KIMB
LAMPH
45 0.1 0.1 28• 39+10 19_+8 0.3_+0.1 0.1 • 13• -99
13 0.2 1.3• 17+11 45• 22+12 0.6_+0.3 0.2_+0.1 13• 0.3• 99
337 0.2 2.8• 42+20 17+15 21+18 0.3_+0.2 0.4• 17• 0.1• 99
503 0.2 4.3+7 16• 36• 29• 0.4_+0.3 0.2• 13• 0.6• 99
48 0.2 1.6• 39• 23• 19+15 0.2_+0.2 0.4• 17• 0.5 99
Ilrnenites Xenocryst, xenolith phases
Groundmass, phenocryst phases LAMP n TiOz A1203 Cr~O3 FeO* MnO NiO MgO CaO Total
KIMB
4 331 50 • 1 50 + 5 0.2• 0.3• 0.05___0.2 1.2• 44 • 2 37 + 9 1.3• 0.9• 0.09 0.1 _0.1 3.3• 9.2• 0.1 0.15• 99 98
LAMPH
AOB
KIMB
LAMPH
AOB
23 52 • 2 0.5• 0.06• 43 • 3 0.5+0.8 0.01 3.8___4 0.26• 99
11 52 0.4 0.05 40 0.57 0.06 5.1+2 0.38 99
582 49• 0.5• 1.3• 37• 0.6• 0.5• 10• 0.04• 98
23 52• 0.5• 1.I• 36• 0.3• 0.0 9.5• 0.05• 99
54 52• 0.5• 0.06• 37• 2.1• 0.04• 7.6• 0.04• 98
Lamproites and K-rich igneous rocks
159
TABLE 10. (cont.) Summary of mineral chemistry of lamproite phases compared with those of other rock
types* Phlogopites Groundmass, phenocryst phases
n
SiOz TiO 2 A1203 Cr203 FeO* MnO NiO MgO CaO Na20 K20 BaO Total Mg no.
Xenolith, xenocryst phases
LAMP
LAMPH KIMB
INT
ATP
ALK
KIMB
258 41 _+2 5.0_+2 10_+3 0.3_+0.4 6.8_+4 0.06 0.11 22_+3 0.07 0.30 10_+0.6 0.8_+0.6 95.2 85_+10
87 37_+2 4.3_+2 15_+2 0.2_+0.4 11_+4 0.14 0.02 18_+4 0.37 0.42 8.5_+2 0.7_+0.4 94.6 74_+14
10 36 2.9 16 -20 0.41 -12 0.08 0.14 9 0.14 96.0 52
5 39 1.3 15 1.2 2.6 0.12 -24 0.02 0.6 9 -93.2 94
48 37 4.6 14 0.02 15 0.31 ~ 14 0.2 0.6 9 1.3 95.5 62
258 25 40_+2 40-+2 1 . 8 _ + 2 3.3-+2 13_+2 14-+1 0.7_+0.8 0.7+0.6 4.8_+2 7_+4 0.03 0.04 0.15 0.04 24-1-2 21 _+3 0.08 0.14 0.33_+0.3 0.45_+0.3 10_+1 10_+0.5 0.2_+0.2 0.9_+0.5 94.6 95.8 90-+5 84_+10
147 39_+2 2.3_+2 13_+3 0.4_+0.5 7_+4 0.06 0.06 23__.4 0.16 0.20 10_+1 1.7_+1 94.2 85+8
Sanidines (or alkalifeldspar) LAMP n SiO2 AIEO3 FezO3 MgO MnO CaO Na20 K20 BaO
Leucites
LAMPH
28 16 63 _+1 65 _+2 17_+ 1 19_+ 1 2.3-+ 1 0.3_+0.1 0.5_+0.4 0.01 0.02 ,0.01 0.25_+0.4 0.24+_0.3 0.47+0.2 3.6_+5 1 6 _ + 0 . 5 11_+5 1 _+0.7 0.5 99 99
AOB
36 37_+2 4.8_+3 16_+1 0.8_+0.8 8.2+5 0.06 0.08 19_+4 0.07 0.9_+0.3 9_+0.7 1.0_+0.6 95.6 80-+13
Priderites
LAMP SiO2 AIEO3 Fe203 CaO Na20 K20
LAMPH
15 56.0 21.0 1.2 0.01 0.05 21.3 99.6
LAMP SiO2 TiO2 A1203 Cr203 FeO CaO Na:O KzO BaO
20 0.1 71.8-+2 0.5_+0.9 0.5_+0.1 9.1 _+3 0.04 0.38_+0.2 7.7_+1.6 9.8_+5 99.6
Garnets Xenocryst, xenolith phases
n SiO 2 TiO 2 A1203 Cr203 FeO* MnO NiO MgO CaO Na20 Total Mgno.
LAMP
KIMB
LAMPH
AOB
18 41 ___1.5 0.27_+0.2 21___2 3.3_+3 12_+9 1.1+1.3 0.01 18_+6 4.5+1.6 0.01 100.0 72+22
1354 41 +- 1 0.36+_0.6 21+3 3.1_+3 11+5 0.38-+0.2 0.04_+0.5 18_+5 6_+3 0.08_+0.2 100.1 74-+14
119 41 _+ 1 0.33_+0.2 22_+1 1.2_+2 12_+4 0.4-+0.2 0.01 17_+4 6.3_+3 0.6_+0.8 100.3 71+_11
31 41 ___1 0.20_+0.1 23_+1 0.8_+1 11+5 0.35-+0.1 0.0 18_+4 5.5_+1 0.05+_0.1 99.9 73-+14
* LAMP, lamproites; LAMPH, lamprophyres; AOB, alkali olivine basalts; KIMB, kimberlites; INT, intermediate volcanic rocks (e.g. andesites); ALK, alkaline plutonic rocks; ATP, alpine-type peridotites; ECL, eclogites. All major-element analyses in weight per cent. FeO*, total Fe as FeO.
I60
S. C. Bergman
MBAY, SB, WKB, HOL and NWI). The fact that many so-called K-richterites from lamproites possess K20/Na20 molar ratios of less than unity is consistent with this late-stage interpretation. Amphibole also occurs as late-stage euhedral crystals lining vesicles in some lamproite lavas. Lamproite amphibole compositions vary widely with the following typical ranges: Mg number, 40-94; A1203, 0.0-1.5 wt.%; TiO2, 1-7 wt.% (Fig. 24). As with lamproite clinopyroxenes and phlogopites, there is usually insufficient A1 to fill all the tetrahedral sites in lamproite amphiboles. Individual K-richterite grains are typically zoned with rims displaying enrichments in Ti, Na and Fe and depletions in K and Mg relative to the cores. Therefore many of the more evolved amphiboles (e.g. some of those in MAP, PEN, MBAY, WKB, PP, SB and HOL) display K20 / Na20 < 1 (molar). The most potassic richterites (K20/Na20 > 1.5 (molar) occur in the WKB, HP and PRA suites. In contrast with the extremely titaniferous phlogopites in the SB suite (more than 10 wt.% TiO2), SB amphiboles have TiO2 contents (3-4 wt.%) similar to those in the average lamproite. The MnO contents of lamproite amphiboles are similar to those of lamprophyres but are significantly higher than most compositionally-similar rock types. K-richterites are also found in some K-rich ultramafic rocks previously thought to be kimberlites (Baster's mine, Pniel, Barkly West (Erlank 1973); this rock is considered a lamproite in the present study) and in MARID-type xenoliths (Dawson & Smith 1977; Erlank & Richard 1977; Dawson 1979); however, these tend to be depleted in Ti relative to lamproite amphiboles. Relative to nearly all other related rock types, lamproite amphiboles are compositionally unique, being the most depleted in A1 and Ca and enriched in K, Ti and Si. Mason (1977) studied the geochemistry of Krichterites from Wolgidee Hills, Western Australia, and found that they are extremely depleted in REE relative to other amphiboles from alkalicmafic magmatic settings. This REE depletion is in agreement with the interpretation mentioned above, i.e. that K-richterite precipitated relatively late in the crystallization sequence following the crystallization of REE-enriched phases such as apatite.
Phlogopite Ti-rich Al-poor phlogopites and tetraferriphlogopites of widely-varying compositions occur as an essential phase in all lamproites, as both phenocryst and/or groundmass phases. Grains with Al-deficient tetraferriphlogopite margins,
often demonstrating a reverse pleochroic formula, are common in both phenocryst and poikilitic groundmass grains. The characteristics of lamproite phlogopites are generally as follows: Mg numbers, 37-94; TiO2, 2-11 wt.%; Cr203, 0.0-1.5 wt.%; A1203, 1-14 wt.%; SiO 2, 3843 wt.%. Extreme zoning in Ti, A1, Mg and Fe is often observed within a given phenocryst, with grain rims most typically enriched in Ti and Fe and depleted in Mg and A1 relative to the cores. The most AlzO3-depleted phlogopites (1-3 wt.%) occur in the WKB, PRA (Kimberlite mine only) and HP suites (see Fig. 25), whereas the MAP and LH suite phlogopites are relatively A1203 rich (11-14 wt.%). The SB and WKB lavas contain the most TiOz-rich phlogopites (7-12 wt.% and 5-11 wt.% respectively) of all lamproite localities. Lamproite phlogopite compositions are distinct from those in almost all igneous and metamorphic rocks (Fig. 25) with the exception of some minette phlogopites (e.g. Bachinski & Simpson 1984; Velde, personal communication, 1985). The most notable exceptions are some kimberlites and their peridotite xenoliths, especially the MARID-suite types (Dawson & Smith 1977; Erlank & Richard 1977; Dawson 1980). Velde (personal communication, 1985) described phlogopites intermediate between phlogopiteannite-tetraferriphlogopite and ferri-annite from a Jersey minette.
K-feldspar Sanidine or rare microcline (and/or orthoclase) are the only feldspars found in the more aluminous members of various lamproite suites (e.g. HOL, LH, KAM, GSB, SB and MAP). Sanidine is restricted to the groundmass and is a late-stage phase. With the exception of the MAP rocks, lamproite sanidines are extremely K rich (Or>95 ; Na20 = 0-2 wt.%; CaO <0.5 wt.%) with anomalously high Fe203* (typically 0.5-5 wt.%) and BaO (0.2-1.6 wt.%) contents. The extreme Fe-rich character of lamproite sanidines separates lamproites from nearly all other rock types; the sanidines from lamprophyres are usually in the range 0r60_95 and contain about 1.0 wt.% Fe203 (Bachinski & Simpson 1984; Rock 1986). The absence of plagioclase in lamproites provides an effective means for distinguishing lamproites from other ultrapotassic-potassic rocks.
Leucite and analeime Leucite is the only feldspathoid that occurs in lamproites. Euhedral and anhedral leucites are
Lamproites and K-rich igneous rocks
161
~
il
i! <~ ~-~C~ i
i
(% JaW OeO
,9.~
/,,"
..~. /,,0 / ..q',
I,,
l': 2:ili i /I/
o \
A
,
c
c
e~
~ v
~..!
.
~e
~X
9
~o
x
0~ .~=? ..~
K~
li",I I I
1
o (% J.~) CO'iV
~
(% ~
0
~0!I
.
~d~2
I62
S. C. Bergman 20..... MAP'
.
GSB
~ . . . .
10-
"...... ~ ~ : -
~"MBA~jl-"
~.
~ ,,..
~,., /1.4
.." VY-~ ,ZI I
0 Mg number
20~ - , . - ~ " ~ : ' ~ . " : r . . ~. . . . .
.
~
'
~
_
~
\
eo10-
%
Mg number
.... ....... . . . .
" LAMPROITES (229) LAMPROPHYRES (98) ALKALINE INTRUSIVES (32) FKIMBERLITE XENOLITHS (193) t-K G R O U N D M A S S - P H E N - X E N C (166)
12-
10-
A8-
Y 4-
'30
;0
~
do
7'0
do
~
40
MO number
FIG. 25. Phlogopite compositions from lamproites compared with other rock types: (a) At203 versus Mg number as a function of the lamproite suite; (b) A1203 and (c) TiO2 versus Mg number for lamproites and other rock types. (See Table 10 and Appendix 2.) Bold and light-face curves as in Fig. 24.
Lamproites and K-rich igneous rocks
typically weakly anisotropic and twinned (e.g. WKB) or isotropic and twin free (e.g. LH) and are most often altered and pseudomorphed by secondary phases (e.g. sanidine, analcime, quartz and zeolites). Lamproite leucites are commonly non-stoichiometric with a marked excess of Si, Fe 3+ and K. Leucite Fe20 3 contents are as high as 1-2 wt.%. These chemical features separate lamProite leucites from the leucites of other potassic-ultrapotassic rocks (e.g. RP and HM) which possess higher A1 and lower Si, K and Fe contents. Analcime has been recognized in the SB, KAM, LH and MAP lamproite suites and most probably originates by the secondary replacement of leucite, although this view is debatable. Whereas lamproites never display classic pseudoleucite textures (i.e. K-feldsparnepheline intergrowths (Shand 1927)), other potassic-ultrapotassic rocks (e.g. HM, TAN and NSW) commonly contain classic pseudoleucite.
Fe-Ti-Cr-Mg-AI oxides A wide variety of oxide phases in the rhombohedral, cubic and orthorhombic series have been noted in lamproites. Spinels (sensu lato) fall into four compositional groups and occur in rocks from WKB, LH, PRA, GSB, ARG, HP, HOL and MAP, among others: (a) Ti-poor Al-rich Mg chromites; (b) Ti-rich A1-Mg chromites; (c) A1poor Ti-Mg chromites; (d) Mg-Ti-magnetites. Whereas some of these spinel types are commonly found in rocks other than lamproites, lamproite spinels are more Fe rich than kimberlite spinels but are similar to those of other alkaline igneous rocks (Haggerty 1976). Spinel zoning, when present, displays Cr-depleted and Fe 3 +-enriched rims relative to the cores (Mitchell 1985). Lamproite spinels tend to be more Mn rich than those from other alkalic mafic rocks (Table 10). Armalcolite has only been found in SB lavas (Velde 1975), but has been produced experimentally in studies on WKB rocks by Arima & Edgar (1983) (see below). Ilmenite with compositions relatively similar to those of other rock types has been noted in the GSB lavas (4 wt.% MgO (Sheraton & England 1980)), at Sisco (COR) (Velde 1968), at Oscar's Plug (WKB) (S. C. Bergman, unpublished data), at Mount North (WKB) Wagner & Velde 1986a) and at Jumilla (MAP) (2.5-3.5 wt.~ MgO) (Mitchell 1985), but is generally absent. Lamproite ilmenite MnO contents are higher, on the average, than those from compositionally related rock types (except alkali basalt xenolith ilmenites) (see Table 10). Anatase occurs in the altered BOB rocks (Mitchell 1985) and rutile occurs in the CHE lamproites;
163
Van Kooten (1980) and Kuehner (1980) discussed ferro-pseudobrookites from the CSN ultrapotassic rocks and LH lamproites respectively, but these phases are all most probably secondary.
K-Fe-Ti-Ba oxides The natural occurrence of the exotic phases priderite ((K,Ba)(Yi,Fe3+)8016) and jeppeite ((K,Ba)z(Ti,Fe 3+)6013) is almost restricted to lamproites. Priderite has thus far been found in the WKB, COR, LH, KAM, HOL, SB and PRA lamproites, and will most probably be discovered in other suites as more detailed studies are performed. Prider (1939) originally mistook priderite for rutile in the WKB rocks, and Cross (1897), Smithson (1959) and Johnston (1959) mistook it for rutile in LH rocks; the more detailed study of the oxides in the WKB rocks by Norrish (1951) and in the LH rocks by Carmichael (1967) demonstrated its existence. Apart from the priderite-bearing apatite glimmerites in India (Gupta et al. 1983), which are probably lamproites, the only known extra-lamproite priderite occurrence (albeit Ba enriched relative to lamproites) was reported by Zhuravleva et al. (1978) in the olivinite-ijolite members of a carbonatite pluton in the Kola Peninsula, U . S . S . R . K . Collerson (personal communication, 1984) identified priderite in AIL minettes, Labrador. N6mec (1985) described priderite from the K-richer members of Czechoslovakian minettes which may be lamproites. Jeppeite (Bagshaw et al. 1977; Pryce et al. 1984) is apparently restricted to the Wolgidee Hills pluton (WKB) as well as the PRA priderite-bearing phlogopite pyroxenite xenoliths (Mitchell & Lewis 1983). Dubeau & Edgar (1985) presented experimental data on the stability of priderite as a function of pressure and temperature. Much more work on the distribution of priderite-jeppeite in lamproites and minettes is therefore required to understand its paragenesis.
Wadeite Wadeite (K4Zr2Si6018) is a hexagonal ringsilicate which was also originally discovered in the WKB lamproites (Prider 1939), where it occurs as rods up to 1 mm long. Carmichael (1967), Kuehner (1980), Kuehner et al. (1981) and Henage (1972) additionally noted its presence in LH and KAM rocks. B. H. Scott-Smith (personal communication, 1983) discovered wadeite in lamproites from PRA. Tikhonenkov et al. (1960) reported an extra-lamproite occurrence of
I64
S. C. Bergman
wadeite in nepheline-feldspar veins of the Khibina alkaline massif, U.S.S.R. Wadeite is compositionally related to dalyite (K2ZrSi6Ols) and other Zr silicates (zektzerite, sogdianite and darapiosite) which occur in rocks more enriched in SiO2 than lamproites (e.g. hyperalkaline granites (Robins et al. 1983; Thibaut et al. 1972)). Arima & Edgar (1980) studied the stability of pure wadeite and found that, by itself, it can occur under a wide range of uppermantle to crustal P - T conditions.
Shcherbakovite (noonkanbahite) Shcherbakovite ((Na,K)(Ba,K)TiaSi4014) is an unusual phase that was discovered in the WKB wolgidites by Prider (1965). Eskova & Kazakova (1955) initially reported its occurrence in alkalic pegmatites of the Khibina massif, U.S.S.R. Shcherbakovite is closely compositionally related to batisite (Na2BaTi2Si4014), which was initially reported by Kravchenko et al. (i960). Unfortunately, shcherbakovite has not received much attention since Prider's original work and only a limited amount of mineralogical data is available.
Other phases Important and common accessory phases of all lamproites include apatite and perovskite. Sphene (mostly secondary) and zircon are rare accessory minerals. Wagner & Velde (1986a) described rare roedderite-like phases (K, Fe, Mg silicates) in MAP and KAM lamproites. Lamproite secondary phases are widespread and include Sr-rich barite, chlorite, heulandite, nontronite, other zeolites, serpentine, carbonate minerals, quartz and rare albite. As mentioned above, xenocrysts of a wide variety of phases, including diamonds and Cr-rich pyrope, can occur in lamproites. However, lamproite pyropes are more enriched in MnO than are kimberlite and other pyropes, but are otherwise similar in chemistry.
Mineralogy and mineral chemistry: summarizing comments Lamproites contain a diverse suite of minerals, some of which are restricted to lamproites. The diagnostic mineral assemblages (i.e. containing phlogopite, K-richterite, priderite, shcherbakovite and wadeite) and unique mineral compositions (K-rich, Ca- and Al-poor amphiboles, Na- and Al-poor diopsides; Ti-rich, Al-poor phlogopites; Ba- and Fe-rich sanidines; Fe-, K- and Si-rich leucites; the mafic minerals largely possessing
primitive Mg numbers in the range 85-93, although more differentiated phases can occur) shown by lamproites reflect their peralkaline and primitive compositions and provide effective means for distinguishing lamproites from other alkalic mafic--ultramafic magmatic rocks. As discussed at length by Mitchell (1985), the zoning characteristics displayed by lamproite phases place important constraints on both the crystallization history of individual magmas and the petrogenetic relationships between several lamproite magmas. Interestingly, the elevated MnO contents of lamproite amphiboles, ilmenites, spinels and garnets are in marked contrast with the MnO-depleted character of the lamproite whole-rock samples (0.10 wt.% (Table 5)) compared with kimberlites, alkali basalts etc.
Experimental studies on K-rich rocks and their minerals Reviews of experimental studies dealing with Krich rocks and their minerals have been given by Gupta & Yagi (1980) and Gittins (1979). Therefore important and more recent papers on the subject are summarized in this section.
Natural lamproites The high-P-T phase relations of lamproite rock compositions have been determined by Barton (1976), Barton & Hamilton (1978, 1979, 1982) and Arima & Edgar (1983). Barton & Hamilton studied orendites, wyomingites and madupites from the LH and found that only the orendites and wyomingites could equilibrate with a garnet lherzolite assemblage at P > 2 6 kb and Pi_i2o= Ptotal" Madupites probably represent partial melts of phlogopite-pyroxenite or phlogopite-olivinepyroxenite assemblages (Barton & Hamilton 1979). Barton & Hamilton's experiments indicate that the dominant liquidus or near-liquidus phases are leucite, olivine, orthopyroxene, clinopyroxene and garnet at P < 3 0 kb. They also concluded that the peralkalinity of these ultrapotassic magmas could reflect either primary source rock compositions or the selective melting of phlogopite + pyroxene. They postulated that the association of low-SiO: madupites with highSiO2 orendites at LH results from variations in H20 and CO2 and local mineral assemblages in the upper mantle source region. It should be noted that A. Edgar (personal communication, 1985) questions the validity of the phase relationships determined by Barton because of Fe partitioning in platinum capsules. Sobolev et al.
L a m p r o i t e s a n d K-rich igneous rocks (1975) studied the low-P near-liquidus phase equilibria of a LH wyomingite and found that diopside is the sole liquidus phase at temperatures around 1320~ and P = 1 b despite the fact that phlogopite phenocrysts are common and that the 1 b solidus temperature is about 1000~ Arima & Edgar (1983) studied the hydrous (+ COO liquidus and sub-liquidus phase relations of a wolgidite (from the WKB, Mount North) to 40 kb. They found that olivine (at P < 15-24 kb) and orthopyroxene (P > 15-24 kb) occur on the liquidus; phlogopite, rutile, clinopyroxene, armalcolite and priderite follow at lower P and T. Rutile reacts with phlogopite and liquid to produce priderite at P < 15 kb and T < 1010~ and armalcolite reacts to form priderite at T< 1010~ and P < 15 kb; rutile is the high-pressure phase and armalcolite is the high-temperature phase. These phases (priderite, armalcolite and rutile) were only found in runs containing added H20 (13 wt.%) and were absent from runs with low H20 contents (3 wt.~). Their results indicate that wolgidite magma most probably represents the partial melt of a source mantle containing phlogopite, rutile, olivine and orthopyroxene. It is unlikely that a wolgidite-composition melt would be derived from the partial melting of a simple garnet- or spinel-bearing lherzolite mantle, in contrast with the derivation of typical alkali basalts from such a mantle. Other ultrapotassic rock compositions
Edgar et al. (1976, 1979), Edgar (1979), Edgar & Arima (1983), Ryabchikov & Green (1978), Arima & Edgar (1983) and Edgar & Condliffe (1978) investigated the phase behaviour of several TAN rocks (biotite ugandite, biotite mafurite and katungite) in the presence of H20 + CO2 to 40 kb. They found that these melts never equilibrate with an orthopyroxene- or garnet-bearing assemblage and that only diopside, olivine, ilmenite and phlogopite occur on the liquidus, even at high pressure. These three magmas cannot be related to each other by fractionation alone or by partial melting of a single mantle source region. The three magma compositions could be derived by partial melting, at various depths, of clinopyroxenite or peridotite sources with varying degrees of enrichment in K where low H 2 0 / C O 2 ratios produce low enrichments in K. Their experimental findings favour metasomatism of a peridotitic mantle source as a prerequisite for the production of K-rich magmas. Ryabchikov & Green (1978), however, concluded that biotite mafurite, olivine leucitite and ugandite liquids could be produced by the partial melting in the presence of H20 and CO2 of a
I65
lherzolite source mantle locally enriched in phlogopite. Phlogopite occurs on the liquidus at high H20/CO 2 ratios (Xco2=0-0.25), but becomes unstable at higher CO2 contents (Xco2 > 0.25). Cundari & O'Hara (1976) studied the phase equilibria of a leucitite from NSW under anhydrous conditions to 40 kb and determined that garnet did not exist with enstatite above 35 kb, indicating that orthopyroxene did not occur in an anhydrous source peridotite. Thompson (1977) experimentally studied a clinopyroxene leucitite from the RP and found that the phenocryst assemblage could be duplicated at 14 kb and 1260~ indicating a mantle derivation. He postulated that the experimental data were consistent with the view that the apparent crustal contamination (suggested by SrO isotope data) was due to the partial fusion of subducted ocean-floor sediments within the upper mantle. Other experimental work on potassic and ultrapotassic rocks includes that of Dolfi et al. (1976), Lloyd et al. (1985), and Esperanca & Holloway (1985). Compositions in the system N a 2 0 - K 2 0 - M g O AI203-SiO2-H20-CO 2
Schairer & Bowen (1938) studied the anhydrous system KA1Si206-CaMgSi206-SiO2 at 1 atm and established the phase equilibria pertinent to Fe-free K-rich alkaline magmas. They found an extremely large leucite field in the leucitediopside-silica ternary system. However, more recent work on the hydrous analogue of this system (Ruddock & Hamilton 1978a) demonstrates that the leucite field is extremely compressed, which explains the general absence of leucite in minettes. Luth (1967) determined the phase equilibria in the system MgzSiO 4KA1SiO4-SiOz-H20. Bravo & O'Hara (1975) studied the partial melting of a synthetic phlogopite-bearing garnet and spinel lherzolite (CO2 free) and found that partial melts of these solids at 15 and 30 kb had extremely low K20 contents (1-4 wt.~) and were quartz- and/or hypersthene-normative, These liquids share an SiO2-rich compositional feature with lamproites, although they are relatively depleted in K20 relative to lamproites. Modreski & Boettcher (1972, 1973) investigated the stability of phlogopite in a model system and found that it is unstable under oceanic geothermal conditions of P > 25 kb but may persist in sub-continental geothermal conditions to P > 50 kb. Wendlandt (1977a, b) and Wendlandt & Eggler (1980a, b, c) investigated the stability and melting behaviour of sanidine and phlogopite in olivine-
I66
S. C. Bergman
bearing assemblages. They found that sanidine + forsterite breaks down to kalsilite+enstatite along a P-Tlocus including 20 kb at 1000~ and 30 kb at 1300~ indicating that the latter assemblage will be stable along most geotherms (Wendlandt & Eggler 1980c). The melting behaviour of phlogopite in peridotite was found to be controlled by the equilibrium mineral assemblage as well as by the composition of the volatile phase (H20 versus CO2). At P < 1 4 kb the liquid compositions were quartz or enstatite normative, but became leucite to kalsilite normative at higher pressures in the presence of CO2 + H20 vapour. At pressures above 30 kb in the presence of magnesite, phlogopite peridotite melts to produce carbonatitic liquids which become progressively enriched in K20 and SiO2 as pressure increases. Phlogopite is not stable in a peridotite assemblage above 50 kb. Ryabchikov & Boettcher (1981) investigated the solubility of K in an aqueous fluid in equilibrium with phlogopite-pyrope-forsterite gels at 10-30 kb and 1050-1100~ and found that water is capable of dissolving 4-25% K20 and that solute KzO/AI203 ratios are close to unity. Ruddock & Hamilton (1978a) experimentally studied the system leucite-diopside-quartz-H20 to 4 kb and confirmed that, under hydrous conditions, diopside and phlogopite will be the first two liquidus phases followed by sanidine and quartz. This study nicely explains the petrography of minettes and other lamprophyres (however, see Rock (1984, Fig. 8) for an alternative interpretation). Ruddock & Hamilton (1978b) determined the stability of carbonate in an ultrapotassic mafic assemblage and found that a minette composition liquid is saturated in olivine, garnet, orthopyroxene and clinopyroxene at pressures above 15-20 kb and is saturated in carbonate at pressures above 20-30 kb. Schneider (1982) and Schneider & Eggler (1984) determined the high P - T (15-20 kb and 750-1100~ solubility of major-element oxides in H20 and H20-CO2 fluids in equilibrium with jadeite-, amphibole- and phlogopite-peridotite and several individual minerals. They found that the fluid's solute contents were in the range 0.518 wt.% and increased with decreasing CO2 content. H20 or H20 + CO2 fluids in equilibrium with phlogopite-peridotite were found to be peraluminous whereas H20 at- CO2 fluids in equilibrium with amphibole- or jadeite-peridotite ranged from mildly peraluminous to strongly peralkaline in their solute compositions. Solute K/Na ratios were found to be controlled by H20 / CO2 as well as by P, T and the bulk composition of the solid, whereby K/Na increases with increasing CO2/H20 ratio of the fluid.
Petrogenesis A wide variety of models have been advocated for the evolution of K-rich rocks; the reader is referred to previous reviews by Bell & Powell (1969), Sahama (1974) and Gupta & Yagi (1980). Most hypotheses can be placed in one of three broad groups. 1 Partial melting of mantle material that was metasomatically (or otherwise, e.g. by plutonism) enriched in phlogopite and other LIL-elementenriched minor phases (e.g. apatite, zircon, sphene etc.); the melt may be subject to variable amounts of fractionation but does not undergo substantial assimilation of crustal material (Waters 1955; Yagi & Matsumoto 1966; Kay & Gast 1973; Boettcher et al. 1975; Beswick 1976; Gupta et al. 1976; Van Kooten 1980). 2 Assimilation of continental crustal material by 'normal' mantle-derived alkali-basaltic, carbonatitic or alkali-ultramafic melts (Daly 1910, 1933; Shand 1931; Rittman 1933; Larsen 1940; Williams 1936; Holmes 1950; Powell & Bell 1970; Taylor & Turi 1976; Turi & Taylor 1976; Taylor et al. 1984). 3 Evolution involving processes such as zone refining, fractional resorption, gaseous transport etc. in otherwise typical mantle-derived melts (Bowen 1928; Saether 1950; Kennedy 1955; Harris 1957; Prider 1960; Marinelli & Mittempergher 1966; Fuster et al. 1967; Kogarko 1980; Kogarko et al. 1968; Kushiro & Aoki 1968; Harris & Middlemost 1969; Stewart 1979; Wones 1979; Ryabchikov et al. 1982). Sahama (1974) observed that researchers tend to desire hypotheses that attempt to explain the petrogenesis of K-rich rocks in general; however, the geochemical distinctions between lamproites and kamafugitic as well as other K-rich rocks do not permit such a luxury. The tectonic-geological and geochemical distinctions between lamproites, kamafugites, shoshonites and other Krich rocks clearly support this view. The last statement in Sahama's (1974)review of K-rich rocks speculates on a relationship between kimberlites, lamproites and kamafug o ites. Gupta & Yagi (1980) additionally recognized the importance of this speculation. The last 10 years have witnessed major discoveries of diamondiferous lamproites and other developments which make this speculation a reasonable, if not undeniable, view. Although kimberlites and olivine lamproites evolve differently, they possess many similarities which have led several workers to suggest a petrogenetic model for lamproites
Lamproites and K-rich igneous rocks whereby kimberlite magma is differentiated to produce a more potassic SiO2-rich lamproitekamafugite magma (Holmes 1932; Wade & Prider 1940; Scott 1981). Although more data are required to establish an unambiguous petrogenetic picture for lamproites, the data now available make it possible to constrain their petrogenesis to a much greater degree than was possible as little as 10 years ago. In addition, many aspects of the petrogenesis of lamproites have important implications for a variety of aspects of mantle evolution and magma genesis. Lamproites are mantle-derived melts and do not simply result from the melting of recycled continental crust. This is supported by (a) the presence of diamonds in olivine lamproites, (b) the presence of mantle-derived peridotite xenoliths (albeit much rarer than their abundance in alkali basalts and kimberlites) in leucite, phlogopite and olivine lamproites, (c) a primitive (Mg number 74 + 9) mafic-ultramafic major-element composition, (d) an enrichment in the refractory trace elements Co, Cr, Ni and Sc, and (e) the mantle-type pressures and temperatures (P = 2030 kb and T= 1100-1300~ calculated by performing thermodynamic isoactivity calculations (Nicholls & Carmichael 1972; Carmichael et al. 1974, 1977; Barton & Wood 1976) in which lamproite magmas are equilibrated with various source-mantle assemblages. In contrast with being enriched in LIL elements, lamproite magmas possess major-element compositional features typical of partial melts of a refractory (i.e. Mg number, 88-92, diopsidedepleted, harzburgite ?) mantle : they are depleted in the incompatible elements Na, A1 and Ca compared with alkali basalts and other primary melts considered to represent low degrees (less than 10%) of partial melting of a lherzolitecomposition mantle. A depleted diopside-poor source is further supported by the high-P-T experimental studies on lamproite compositions which suggest a non-lherzolite mantle source for lamproite magmas (see above) and also by the characteristically refractory harzburgite and dunite xenolith population that occurs rarely in several lamproite suites (e.g. WKB, LH, MAP). However, there are other ways of interpreting this depletion in certain major elements. Lamproites are phlogopite-rich rocks and many whole-rock analyses of lamproites approach but are more calcic and siliceous than the bulk composition ofphlogopite. Therefore it is possible that lamproites represent the selective fractional fusion of mantle phlogopite and other LILelement enriched phases (e.g. apatite, zircon etc.) diluted by small amounts of peridotite compo-
167
nents (clinopyroxene, orthopyroxene and olivine). Since phlogopite is severely depleted in Na and Ca and moderately depleted in AI compared with basaltic melts, this model nicely explains the bulk chemistry of lamproites. It is further supported by the fact that phlogopite is not expected to survive extensive amounts of partial melting, most of it being consumed at extremely low degrees of melting. The ultrapotassic but variable K/Na ratios of lamproites can be interpreted in a number of ways. K can be enriched relative to Na as a result of (a) source-mantle characteristics, i.e. precipitation of phlogopite (Na absent) via metasomatism in a Na-poor refractory peridotite, (b) fractionation of high K/Na phases (e.g. phlogopite, leucite) and (c) selective volatile-phase transfer of K relative to Na. The experimental work by Schneider & Eggler (1984) and Schneider (1982) contributes important data with which to judge the effectiveness of the latter process, i.e. K/Na in the fluid varies with CO2/H20. However, much more work on it is required to understand vapour-phase transfer of alkalis properly. Since this feature is a local as well as a general feature of lamproites, it most probably results from a combination of all these processes (see discussion in Sahama (1974, p. 106).) Like alkali basalts and lamprophyres, the volatile budgets of leucite-lamproite and phlogopite-lamproite magmas are water-dominated (relative to CO2), whereas those of kimberlites (and olivine-lamproites) possess higher and subequal H20 and CO2 contents. The H20 contents of kimberlites indicated in Tables 4 and 5 are maximum values for the magmas, and the true values are most probably a factor of 2 lower owing to the ubiquitous and often secondary serpentinization that characterizes kimberlite groundmass and xenocrysts. As mentioned above this HzO-rich character provides an explanation for the sherbet-glass shape of lamproite pipes. It is also consistent with the SiO2-rich and variable composition of lamproites. Experimental work by Mysen & Boettcher (1975a, b) and Eggler (1977) has shown that partial melting of a peridotite with a high H20/(CO2 + H 2 0 ) ratio will produce a relatively SiO2-rich melt, eliminating the need to invoke fractionation of low-SiOz mafic minerals to produce SiO2-rich lamproites. Slight variations in the source H 2 0 / ( C O 2 -k-H 2 0 ) ratio would produce variations in the SiO2 contents of partial melts (e.g. Bachinski & Scott 1979). Therefore, whereas kimberlites and alkali olivine basalts are SiO 2 undersaturated and are most probably derived from a mantle source with a low H20/(CO 2 -k-H20) ratio, lamproite sources represent the complementary volatile situation
I68
S. C. B e r g m a n
(high H20/(CO 2+H20)) in addition to other contrasts, i.e. LIL-element enrichments. Of all the alkaline rocks, lamproites are consistently among the most enriched in LILelements. They are also characterized by enrichments in higher-series elements relative to lowerseries elements in the same group (e.g. lower K/ Rb and Sr/Ba) compared with virtually all other rock types. It is now clear that this feature is a mantle phenomenon and does not require crustal contamination for its explanation. The highly variable concentrations of LIL-elements in lamproites is easily explained by a heterogeneous distribution of LIL-element-enriched phases in the source mantle in addition to volatile-phase transfer processes acting during magma ascent. These minerals, most importantly phlogopite but also including other LIL-element-enriched phases such as apatite, monazite, sphene, rutile and zircon, originated from the process of mantle metasomatism, most recently reviewed by Bailey (1982) and Dawson (1984). These phases could be precipitated or otherwise produced from any number of sources, including (a) fluids ascending from a down-going slab which hybridize the overlying mantle wedge, (b) fluids emanating from an ascending plume, (c) 'background' deepmantle fluids ascending throughout the upper mantle and (d) silicate melts of variable origins. As mentioned above, this metasomatism must have been fairly old (10-103 Ma) to generate the 87Sr/86Sr and 143Nd/144Nd ratios of lamproites. The age of this lamproite-source metasomatism contrasts with that suggested for the source regions of alkali basalts and related melts. Since H20 is suggested as the most important 'lamproite volatile', the mantle metasomatism associated with lamproite source regions may be distinct from that associated with alkali basalts (CO2-rich?). If we accept the view that olivine lamproites and leucite lamproites (or phlogopite lamproites) form a cognate petrological continuum, a view that is supported by the time-space-composition relationships present within the WKB and PRA suites, several implications arise. The compositional variations in these suites indicate the existence of differentiation processes more complex and less understood than simple crystal fractionation. Also, the preservation of diamonds within the WKB and PRA suite diatremes indicates that these magmas explosively ascend to the surface from depths of greater than 150 km, i.e. within the diamond stability field in the upper mantle because lamproite melts, which are extremely hydrous, would otherwise oxidize the diamonds. Although this feature is shared by kimberlites, many lamproite diamonds display a
greater degree of graphitization and rounding than those from kimberlites; these features may indicate that lamproite magmas ascend more slowly than kimberlites. Much more work on the stability of diamond in lamproite magmas is required to acquire a better understanding of the ascent characteristics of lamproites. Another implication resulting from the mantle derivation of lamproites involves the radiogenic isotope systems Sm-Nd and Rb-Sr. Some researchers have recently interpreted the isotopic data of K-rich rocks as indicating the operation of continental crust contamination (e.g. TAN (Vollmer & Norry 1983b)) or the mixing of a depleted mantle component (similar to MORB: eNa = + 10 and eSr 40) with an enriched mantle component (eNd= -- 16 and ~Sr= + 240, e.g. WKB (McCulloch et al. 1983a, b)). However, the isotopic data can also be interpreted in the context of a single-mantle reservoir that is heterogeneous but enriched in its Rb/Sr and Nd/ Sm ratios relative to the bulk Earth. The heterogeneities could be related to partial melting or plutonic processes. Small differences in these ratios can produce the time-integrated range in 143Nd/l*4Nd and 87Sr/86Sr ratios observed for lamproites. Regardless of the specific model used to interpret the isotopic data (as all are nonunique), an enriched mantle source is required by all. This 'enriched' source is also represented by Group II kimberlites (Smith 1983a, b) and the diopsides from kimberlite xenoliths (Menzies & Murthy 1980a, b; Basu & Tatsumoto 1980). However, the major-element chemistry of lamproites is characterized by a depletion in some of the major elements that are the first to enter partial melts (Ca, AI and Na) and an enrichment in refractory elements (Mg and Ni) so that the source probably experienced a partial-melting event prior to the enrichment in Nd/Sm and Rb/ Sr (see above). When this is viewed in the light of the Pb isotope data obtained by Fraser et al. (1985) and Nelson et al. (1986) (see above), it is seen that lamproite mantle sources must have been subjected to depletion events followed by enrichment events a long time ago ((1-3)x 103 Ma). This model has interesting implications for Earth history because not only does it suggest the enriched complement to the depleted mantle source represented by MORB and nearly all alkali basalts (except suites such as Kerguelen and Tristan da Cunha), it also supports the existence of Precambrian partial-melting episodes in the mantle and suggests a three-stage model of petrogenesis. If the more recent lamproite suites that occur overlying fossil Benioff zones indicate the general tectonic environment of lamproites, then it is possible that lamproites represent =
- -
Lamproites and K-rich igneous rocks partial melts of sub-continental lithosphereasthenosphere that was (a) initially depleted by the formation of MORB at a spreading centre, (b) moved beneath a continent and (c) later hybridized perhaps by fluids and melts emanating from a down-going slab. All this must have taken place sufficiently long ago to permit the development of time-integrated high 87Sr/86Sr and low 143Nd/14aNd ratios. The most recent melting event that led to the surface emplacement of lamproite magmas could be triggered or enhanced by a number of mechanisms, including transient hot-spots (randomly contacting the unique lamproite-source hybridized mantle; unlike ocean island chains, however, lamproites do not generally form time-space-related linear trends), the build-up of radiogenic heat (due to sources enriched in K, Th and U) and the tectonic conditions of the overlying crust (involving deep fractures). It should be noted that variations of this three-stage model of depletion-enrichmentmelting have been suggested by various workers (Best 1975; Beswick 1976; Thompson et al. 1984). When the lamproite sensu stricto and other ultrapotassic rock localities are plotted on a map with the continents in the positions of 180-200 Ma ago (see Fig. 26), a zonal pattern of the distribution of these K-rich rocks emerges. Two broad belts are evident: one belt 3000 km long includes the Indian, W Australian and Antarctic occurrences, and the other more diffuse belt includes the N American, European and W African localities. Since the mantle source regions
9
~ 7
of lamproites are evidently very old, this pre-drift zonation is perhaps expected and not only permits the possible prediction of other lamproite localities but also has important implications for mantle evolution. Is it possible that the southern hemisphere lamproite belt (representing 1.2x 103 Ma of magmatic activity) delineates a Precambrian (Archaean?) zone of recycling of crustal material (fossil Benioff zone)? While much speculation has filtered into this discussion, it is hoped that a coherent picture of lamproite petrogenesis has emerged. The lamproite revolution is only in its infancy and future studies of these exotic rocks will surely clarify many of the issues only briefly discussed here.
Kimberlite--lamproite relationships Dawson (1987) summarizes the similarities in geochemistry and mineralogy between olivine lamproites and the more micaceous kimberlites (including those in Group II of Smith (1983a, b)) and suggests that olivine lamproites may, in fact, be members of the Group II kimberlites. Smith (1983b) demonstrated Nd-Sr-Pb isotopic similarity between Group II kimberlites and olivine lamproites from PRA and WKB and suggested further that they are one and the same, a view that is additionally held by the present author. Smith (1983b) also suggested that lamproites (as well as Group II kimberlites) were derived by the partial melting of sub-continental lithosphere,
~
~'/~'~ ~
169
LAMPROITES 8~ II ULTRKAsPOTASSIC~/
~'~"~7
9 KIMBERLITES j
EAnt;
~.
\\ \
4-
"')l.".
-
\
,,
4-
FIG. 26. Distribution oflamproites, ultrapotassic rocks and kirnberlites on a map showing a reconstruction of the continents 180-200 Ma ago (Owen 1983). Lamproites fall along two broad belts.
I7O
S. C. Bergman
whereas Group I kimberlites were derived by the partial melting of sub-continental lithosphere or underlying asthenosphere. The fact that MARID-suite xenoliths occur in kimberlites, indicating lamproite-type plutonic-metasomatic processes in sub-continental lithosphere encountered by the host ascending kimberlite magmas, is consistent with both lamproite and kimberlite melts existing in similar areas in the upper mantle.
Relationships between lamproites and MARID and other mica-rich xenoliths The similarity between the geochemistry and mineralogy of lamproites on the one hand and the M A R I D glimmerite and PKP (amphiboleperidotite) suites of xenoliths found in kimberlites on the other suggests some type of relationship. MARID xenoliths, which are also known as 'mantle pegmatites' consist of mica, amphibole, rutile, ilmenite and diopside (Dawson & Smith 1977); in addition, compositionally-similar phases (Ti-mica, K-richterite) are observed as metasomatic alteration products of many peridotite xenoliths from kimberlites (Erlank 1970, 1973; Erlank & Rickard 1977; Erlank et al. 1982; Jones & Smith 1985). Haggerty (1983) and Haggerty et al. (1983) discussed K - B a - C r titanates from metasomatized peridotite xenoliths from S Africa; these phases are not unlike the priderite-jeppeite phases from lamproites. Jones et al. (1985) presented isotopic data for xenoliths from Kimberley, S Africa, which indicated that glimmerite xenoliths were disrupted pegmatitic segregations of Group I kimberlites (terminology of Smith (1983a, b)), whereas M A R I D and PKP xenoliths formed by magmatic-metasomatic processes involving both Group I and Group II components. However, trace-element and isotopic studies on metasomatized peridotite and M A R I D xenoliths from the Bultfontein kimberlite mine, Kimberley, S Africa (Kramers et al. 1983) have demonstrated that the S r - N d - P b isotopic features are consistent with mixing of a metasomatic fluid from a slightly depleted mantle source with mantle material showing a timeintegrated enrichment in incompatible elements. Nevertheless, the dominant (albeit extremely rare) amphibole in kimberlites or in the replacement zones in their included-mantle xenoliths is the K-rich richterite (slightly depleted in Ti relative to lamproite K-richterites) which is the characteristic (and nearly ubiquitous) amphibole in lamproites. The subtle compositional differ-
ences between K-richterites of lamproite, M A R I D xenoliths and metasomatized peridotite xenoliths can be explained by a variable buffering capacity of the host mantle peridotite which may modify the original amphibole composition. The presence of K-richterite, combined with the Tiand phlogopite-rich nature of M A R I D suite xenoliths, further suggests that lamproite-like fluids were the agents of patent metasomatism of kimberlite xenoliths; lamproite melts may be responsible for the precipitation of M A R I D xenoliths. However, much remains to be explored in the M A R I D and metasomatized xenoliths before the extent of this genetic relationship is fully realized. Interestingly, the nature of metasomatic phases (mica and amphibole) in lherzolite and harzburgite xenoliths from alkali basalts (higher A1, and kaersutite or pargasite rather than richterite (Boettcher & O'Neil 1980; Menzies & Murthy 1980b)) substantiates a distinct nature to the mantle source regions of MARIDsuite xenoliths and metasomatized alkali basalt xenoliths.
Conclusions Lamproites are recognized as a coherent but variable and exotic petrographic and chemical group of K-rich igneous rocks which border on and share certain petrogenetic aspects with the alkali basalt, kimberlite, lamprophyre and other K-rich rock clans. Lamproites should not be included in the lamprophyre clan. Lamproites can be effectively distinguished from these other rock types on the basis of mineralogy, mineral chemistry and whole-rock major, trace-element and isotope chemistry. Lamproites are unparalleled by other rock types with respect to their compositions. Known lamproites occur in 21 suites on six continents. Lamproites occur closer to the margins of continents whereas kimberlites are most abundant nearer the craton cores; lamproites often intrude crust that overlies fossil Benioff zones. Lamproites are partial melts of a metasomatized (i.e. phlogopite-, apatite-bearing) but depleted (in Na, AI, Ca) source-mantle peridotite (harzburgite). A three-stage model (depletion-enrichment-melting), at least, is required to explain the evolution of the source regions of lamproite magmas. Lamproites differentiate by processes including crystal fractionation, fractional resorption and volatile-phase transfer, among others. ACKNOWLEDGMENTS: This research was supported by the Anaconda Minerals Company, a division of the Atlantic Richfield
Lamproites and K-rich igneous rocks Corporation, in conjunction with their diamond exploration programme in the U.S. and Kalimantan. I thank R. W. Knostman, I. Gemuts and L. G. Krol for their continued support. Preprints supplied by R. Mitchell, B. Scott-Smith, D. Colchester, J. V. Smith, P. H. Nixon, P. Gregory, D. Velde, R. Vollmer, C. Hawkesworth and M. Menzies are gratefully appreciated. These workers, as well as L. G. Krol, N. R. Baker, M. Skinner, D. Velde, S. Bachinski, P. Berendson, M. Bickford, K. Collerson, F. Albarede, K. Fraser and F. Dodge, participated in many fruitful discussions which served to improve the content of this report. Two days with D. and B. Velde in Ayron proved enlightening. Pat Bickford generously supplied unpublished mineral chemical data on the Hills Pond rocks, Peter Nixon provided samples of the Murcia-Almeria lamproites, Pieter Berendsen provided samples of the Hills Pond lamproites and B. Scott-Smith supplied samples of the Holsteinsborg lamproites. I thank R. Baker, D. Dunn, J. Crann and W. Turner for logistical support and assistance in the field in the U.S.A., Indonesia and Australia. S. Self greatly assisted in the interpretation of the piperno and other pyroclastic rock textures. L. Noodles and Maggie clarified many confusing concepts. An exceedingly constructive review by Sharon Bachinski was invaluable; D. Velde, L. Krol and J. G. Fitton additionally provided useful criticisms on an earlier draft. I thank J. G. Fitton and B. G. J. Upton for only requiring a 25% reduction in the length of the original manuscript. I thank G. W. DeArmond, G. McEntire, D. Schraeder, L. Auvermann and I. Martinez for performing many arduous literature searches and for obtaining many of the obscure references cited in the text. I thank J. Toney and D. J. Henry for much invaluable help in the microprobe work. The rock and mineral chemistry data bases were compiled with the able assistance of V. Mount, S. R. Yang, F. Stiffand E. Kinsel to whom I express sincere thanks. The manuscript was typed by S. Epperson and the graphics were professionally produced by N. Murray.
PRA
SB
(ALL) (CSN) (DSV) (FOR) (GAL) (HM)
(KNX) (NHB)
(SEW) (SFC) (TB) (CHIN)
171
Bolivar 1977; Gogineni et al. 1978; Williams 1891; Scott-Smith & Skinner 1984a, b; Bergman, unpublished data Matson 1960; Velde 1975; Bergman, unpublished data Hawkins 1976 Van Kooten 1980 Dodge & Moore 1981 Ross 1926b Allen et al. 1975 Weed & Pirsson 1896; Pirsson 1905; Osborne & Roberts 1931 ; Wolff 1938; Hurlbut & Griggs 1939; Buie 1941; Tappe 1966; Witkind 1969, 1973; Woods 1975; Bergman, unpublished data Bastin 1906 Williams 1936; Lewis 1973; Schmitt et al. 1974; Rogers et al. 1982 Miller 1972 Templeman-Kluitt 1969 Cross 1906 Arculus & Smith 1979; Schulze & Helmstaedt 1979
Australia WKB
ARG (NSW)
Wade & Proder 1940; Prider 1960, 1982; Atkinson et al. 1984a; Jaques et al. 1984b; Nixon etal. 1984; Bergman, unpublished data Atkinson et al. 1984a Cundari 1973
Europe Osann 1906; Washington 1917; Parga Pondal 1935; Fuster & Pedro 1953; Borley 1967; Fuster et al. 1967; Nixon etal. 1984; Venturelli etal. 1984a Dal Piaz et al. 1979; Venturelli et al. NWI 1984b Hall 1982 PEN Tidmarsh 1932; Knill 1969 HLM Velde 1967 COR (TUSC) Gallo 1984 (SUNN) Furnes et al. 1982 N6mec 1972 (BOH) Duda & Schminke 1978 (LAC) MAP
APPENDIX 1 Sources of whole-rock major and trace-element analytical data in Tables 4-7 and associated figs 20-22
N America EVD HOL HP KAM LH
Libby 1975 Scott 1979 Franks et al. 1971 ; Merrill et al. 1977 Best et al. 1968 ; Bergman, unpublished data Carmichael 1967; Kuehner 1980, Kuehner et al. 1981; Schultz & Cross 1912; Yagi & Matsumoto 1966; Bergman, unpublished data
Africa BOB PPS
Bardet 1973 (Swartruggens only) Skinner & Scott 1979; Bergman, unpublished data
172 (TAN) (AZZ)
S. C.
Bergman
Holmes & Harwood 1937; Holmes 1950; Higazy 1954; Sahama 1974; Mitchell & Bell 1976 Vila et al. 1974
Antarctica
GSB Sheraton & Cundari 1980 MBAY, PP Sheraton & England 1980 Asia and Indonesia
COC Lacroix 1933a, b CHE Bergman & Baker 1984 GDW S a r k a r e t a l . 1980; G u p t a e t a l . 1983 (BAL) Kostyuk 1983 (KAJ) Brouwer 1909 Abbreviations are given in Tables 2 and 3 and in the text. Average compositions
The major- and trace-element data for lamproite sensu stricto are cited above; the major-element data for lamprophyres, kimberlites and alkali basalts were compiled on a computerized data bank that is a combination of the Mutchler et al. (1973) PETROS data bank and a literature compilation by Bergman (File DIROC). Individual references can be obtained by writing to Bergman. Trace-element data for kimberlites and alkali basalts are taken from Wedepohl & Muramatsu (1979) and data for lamprophyres are taken from Rock (1984, 1986). The estimated composition of the primitive mantle comes from Taylor & McLennan (1981) or Mason (1979).
leucite, A analcime, J priderite and jeppeite, A amphibole, I ilmenite, G garnet (suite abbreviations are as given in appendix 3). ARG
Atkinson et al. 1984a (G, H, C); Scott-Smith & Skinner 1984b (P) GSB Sheraton & Cundari 1980 (I, S, H, O, P, C) HOL Scott 1981 (A, S, O, P, C) HP Merrill et al. 1977 (A, P, C); P. Bickford, personal communication 1984 (A, C, P, S); Mitchell 1985 (P, S) LH Carmichael 1967 (A, S, O, P, C, J, F, L); Barton 1979 (L, F, C, P); Kuehner 1980 (A, S, O, P, C); Barton & van Bergen 1981 (H, O, P, C); Kuehner et al. 1981 (A, S, O, P, C); Mitchell 1985 (S, J) MAP Borley 1967 (P, C); Carmichael 1967 (P, F, O, C, A); Fustor et al. 1967 (P); Lopez Ruiz & Rodriguez Badiola 1980 (A, O, P, C); Venturelli et al. 1984a (A, H, S, P, P, C); Mitchell 1985 (P, F, I) MBAY, PP Sheraton & England 1980 (A, P, I,
F) PEN PRA
SB
WKB
APPENDIX 2 Sources of mineral chemical data of Table 10 and Figs 23-25
The mineral chemical data used in these figures and table have been compiled in a computerized data file DIMIC that comprises over 137,000 analyses of phases from kimberlites, lamprophyres, lamproites, alkali basalts and their included xenoliths and megacrysts (S. C. Bergman, unpublished). Since it is unreasonable to tabulate all the data sources here, a copy of the sources can be obtained by writing to Bergman. Only sources containing lamproite mineral chemical data will be summarized. The abbreviations are as follows: P phlogopite, O olivine, C clinopyroxene, H orthopyroxene, S spinel, F sanidine, L
Hall 1982 (A, P, F) Lewis et al. 1976 (G, S, O, P, C); Gogineni et al. 1978 (A, G, S, O, P, C); Mitchell & Lewis 1983 (A, C, P, J); Scott-Smith & Skinner 1984a, c (A, S, O, P, C); Mitchell 1985 (J) Velde 1975 (A, S, P, C); Mitchell 1985 (P, O, N, F, S); Henry & Bergman, unpublished data (A, P, C,O, S) Carmichael 1967 (C, P, A, F, L, J); Mitchell 1981 (P); Mitchell & Lewis 1983 (A, P, C): Atkinson et al. 1984a (G, H, S, C); Jaques et al. 1984a, b (A, G, H, S, O, C); Pryce et al. 1984 (J); Scott-Smith & Skinner 1984b (A, P, S); Mitchell 1985 (P, A, C, L, S,J)
APPENDIX 3 Alphabetical listing of rock suites discussed in the paper
ARG (AZZ) (BAL) BOB (BOH) CHE
Argyle, W Australia Azzaba, Algeria, Africa Baikal Rift-Aldan Shield, U.S.S.R. Bobi and Seguela, Ivory Coast, Africa Bohemian Massif, Czechoslovakia Chelima, Andhra Pradesh, India
Lamproites and K-rich igneous rocks (CHN) (CJ)
(coc) (COL) COR (CSN) DSV EVD (FOR) (GDW) GSB HLM (HM)
HOL HP KAM (KNX) (LAC) LH LUA MAP MBAY (NHB)
(NSW) NWI
Channel Islands, U.K. Campos de Jordao, Brazil Coc Pia, N Vietnam Colima Graben, Mexico Sisco, Corsica, France Central Sierra Nevada, California Deep Spring Valley, California Enoree Vermiculite, S Carolina Fortification Dyke, Colorado Gondwana Coalfields, India Gaussberg, Antarctica Holmeade Farm, U.K. Highwood Mountains, Montana Holsteinsborg, W Greenland Hills Pond, Kansas Kamas, U t a h Knox County, Maine Laacher See Province, Germany Leucite Hills, Wyoming Luangwa Graben, Zambia Murcia-Almeria province, Spain Mount Bayliss, Antarctica N a v a j o - H o p i Buttes, Arizona Lake Cargelligo Area, NSW, Australia N W Italy
PAM PEN PIS PD PP PPS PRA (RP) (SAB) (SAC) SB (SEW) (SRE) (SUNN) (TAN) (TB) (TUSC) WKB
173
Pamir, U.S.S.R. Pendennis, U.K. Orciatico, Pisa, Italy Piedade, Brazil Priestly Peak, Antarctica Postmasburg, Pneil etc. S Africa Prairie Creek, Arkansas Roman Province, Italy Santo Antonia da Barra, Brazil Sacramento, Brazil Smoky Butte, Montana Seward Peninsula, Alaska Srednogorie, Bulgaria Sunnfjord, Norway Toro-Ankole, Birunga, U g a n d a Two Buttes, Colorado Tuscany, Italy W Kimberley, Fitzroy Basin, W Australia
Those suites in parentheses are not lamproites but merely K-rich rocks; those not in parentheses conform to the present definition of lamproite sensu stricto
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STEVENC. BERGMAN,ARCO Resources Technology, Exploration and Production Research Center, 2300 W. Piano Pkwy, Piano, Tx 75075, U.S.A.
The nature and origin of lamprophyres: an overview Nicholas M. S. Rock S U M MA R Y: Lamprophyres are 'alkaline rocks' because they carry high alkalis at a given percentage of SiOz, together with one or more of normative ne, lc or ac, modal foids, and NaK-Ti-rich amphiboles or pyroxenes. They reach higher whole-rock H20,CO2, Sr and Ba contents than other silicate igneous rocks. Contents of related element subsets in amphiboles (Ti, Ba), K-feldspars (Ba, Fe3), phlogopites (Ti, Ba, Fe 3) and pyroxenes (Ti, A1, Fe 3) include among the highest values known in nature for these minerals. 'Primitive' minerals (diopside, forsterite) commonly coexist with 'evolved' minerals (albite + orthoclase, quartz). Four welldefined 'branches' of the lamprophyre 'clan' have distinctive compositions: calc-alkaline (shoshonitic) lamprophyres (minettes etc.), alone among lamprophyres, have mixed alkalinecalc-alkaline affinities; alkaline lamprophyres (camptonites etc.) are basanitic to nephelinitic and, alone among lamprophyres, usually have Na >>K; ultramafic lamprophyres (aln/Sites etc.) are the most Si poor and Ca rich of silicate igneous rocks, and grade into carbonatites; lamproites (orendites etc.) are uniquely rich in K, Rb, Ba, Th, Mg, Cr and Ni at mainly 'andesitic' SiO2 contents, and carry a suite of diagnostic minerals (wadeite etc.). Each branch comprises at least four rock types which resemble each other much more than rock types of other branches; however, some rock types can be grouped into slightly distinct "families' within one branch (e.g. phlogopitic and madupitic lamproites). A case can be made for including kimberlites as a fifth branch of the lamprophyre clan. Synoptic plots and tables, based on some 5000 major, trace-element and mineral analyses, are presented, to aid identification and classification. Lamprophyres are far more common than generally stated, occurring worldwide in more tectonic settings than many other alkaline rocks and throughout the geological record. They may approximate intratelluric magma compositions. Nearly all represent primitive magmas, and many represent primary magmas. Some represent parental magmas to a wide range of hydrous alkaline intrusive suites: calc-alkaline lamprophyres to potassic pyroxenite-diorite-shonkinite-syenite-granite plutons (Cortlandt); alkaline lamprophyres to hornblendic gabbro-syenite plutons (Monteregian Hills); and ultramafic lamprophyres to ijolite-carbonatite complexes (Fen).
Introduction
Nomenclature
Lamprophyres are the only 'alkaline rocks' covered neither by Sorensen (1974) nor yet in a specialized monograph. Despite Streckeisen's rationalized nomenclature of 1979 and many petrologists' awakening recognition of their global significance, they remain to many others a curiosity (Seeliger 1975) and continue to be treated at best as an 'uncomfortable afterthought' (Hughes 1982) in textbooks. The term 'lamprophyric' often conceals meanings as vague but varied as 'altered', 'biotite-rich' and 'exotic' (Heinrich & Alexander 1975) or even 'not studied in detail'! Yet it is one of few rock names intrinsically conveying its meaning (Greek lampros, porphyros, glistening porphyry or purple rock)--compare locality-based names such as kimberlite, and terms such as granophyre and keratophyre which, sophistically, do not imply porphyritic texture. Furthermore, lamprophyres can often be identified in the field (unlike many fine-grained rocks), and they are among the most widespread of alkaline rocks. This overview is part of a continuing attempt to give them their rightful place in the literature.
Previous confusion perhaps attained its nadir with the Arrow Peak dyke, Montana, variously called 'orthoclase-camptonite' (Rosenbusch 1897), 'minette' (Pirsson 1905), 'leucite-monchiquite' (Beger 1923), 'diopside-lamprophyre' (Knopf 1936) and even 'mafic phonolite' (Buie 1941). A hierarchical classification, expanded in the following ways from Streckeisen (1979), is used here (Fig. 1). 1 As in Rock (1986), Streckeisen's 'melilitic lamprophyres' are renamed 'ultramafic lamprophyres'. 2 As in Rock (1977, 1981), lamproites (the leucite-lamproites of some authors) are regarded as a fourth 'branch' of the lamprophyre 'clan'; they share all the characteristics defined below, and have almost invariably been described as 'lamprophyric' (from Niggli (1923) to Jaques et al. (1984)); the name lamproite, of course, intrinsically implies close affinity with lamprophyres. 3 As in Hughes (1982), and as elaborated in Rock (1987) kimberlites are regarded as a fifth
From: FITTON,J. G. & UPTON,B. G. J. (eds), 1987, Alkaline lgneous Rocks, Geological Society Special Publication No. 30, pp. 191-226.
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Nature of lamprophyres branch of the lamprophyre clan (not as a single rock type). Kimberlites must be a subdivision of lamprophyres, and not vice versa, because (a) lamprophyre is a broader 'sack-term' than kimberlite, (b) kimberlites are frequently described as 'lamprophyric' but few lamprophyres are 'kimberlitic', (c) kimberlite is a locality-based name, theoretically with a type-rock, and lamprophyre is more suitable as a clan term, having no type-rock, and (d) lamprophyre has historical precedence. Careful use of this terminology should, in time, overcome much of the lamprophyre 'mystique', provided that (i) the very great differences between the four branches are no longer ignored and (ii) the term 'lamprophyre' is mainly restricted to field use and the rock names in Fig. 1 employed in petrological descriptions. Both changes are necessary because 'lamprophyre' alone covers an inordinate range of compositions from ultramafic to mesotype, ultrabasic to intermediate, Na rich to K rich, melilite to quartz bearing, and peraluminous to perpotassic (Table 1). It is thus even more petrologically vague than 'diabase'. The 'simple descriptive terminology' (biotite-pyroxene-lamprophyre etc.) advocated by Knopf (1936) and Currie (1976) is no less ambiguous for similar reasons.
Purpose of this overview Rock (1977, 1984, 1986) has specificallyreviewed alkaline (AL), calc-alkaline (CAL) and ultramafic (UML) lamprophyres. Gupta & Yagi (1980), Mitchell (1985), Bergman (1987) and Dawson (1987) have covered lamproites (LL). The aim of the present overview is to identify features which distinguish (1) all lamprophyres from 'common' igneous rocks and (2) the different branches from each other. It also indicates how lamprophyres may be fundamental for our understanding of other alkaline rocks. Heterogeneities do exist within each lamprophyre branch. For example, several chemically distinct groups of minettes can be distinguished (Rock 1980, 1984). However, differences within branches are much less than between branches (Metais & Chayes 1964). Moreover, many unimodal (approximately normal or log-normal) element distributions are revealed within each branch by, for example, (i) histogram shapes (Figs. 2 and 7) and (ii) various statistical tests such as kurtosis value, equality of mean and median, and Fillibens's r (Lister 1982). Because such distributions are characteristic of single rock types (LeMaitre 1982), each branch can legiti-
I93
mately be considered as a single petrological entity, at least for comparison with the other branches. Individual rock types will therefore only occasionally be considered hereafter.
Criteria for identification of lamprophyres as a clan A hierarchy of criteria, in roughly decreasing order of importance, is outlined below and explained in later sections; more specific criteria for each branch have been described by Rock (1977, 1984, 1986). Mineral assemblage
Essential phases should include amphiboles, biotites-phlogopites and, often, minerals rich in F, C1, SO3, CO2 and H20 such as halides, carbonates, sulphates and zeolites. Compositional fields of common minerals are delineated in Tables 4-7 and Figs 2-4. High Ba, Fe 3, A1 or Ti in K-feldspars, amphiboles, ph!ogopites and pyroxenes (as appropriate) are characteristic. Mg-rich mafic minerals commonly coexist with alkali feldspars and even with quartz. Cognate minerals absent from all lamprophyres include olivines with mg ~<75, Fe-Mg-Mn- and tremolitic amphiboles (sensu Leake 1978), muscovite, orthopyroxenes, pigeonite and wollastonite. Minerals additionally absent from each individual branch are detailed in Table 1. Texture
The rock should be mesotype to hypermelanocratic and porphyritic (preferably panidiomorphic) with no feldspar or quartz phenocrysts. Feldspathoid or melilite phenocrysts and groundmass olivines are almost always absent. Felsic globular structures and pseudohexagonal or 'castellated' (battlemented) biotites (often with deepbrown rims and pale cores) are characteristic. Most fine-grained mafic dykes with abundant macrocrysts or phenocrysts of biotite or amphibole are lamprophyres (though many lamprophyres only have groundmass amphibole or biotite). Mode of occurrence
The rock should form a dyke, sill, plug, stock, (sub)-volcanic vent or diatreme, or, exceptionally, a margin to a larger intrusion. Associated breccias are very common. Intrusions more than a few square kilometres in outcrop remain unconfirmed. Only a few examples of lamprophyre lavas and pyroclastics are at present known (for
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Nature of lamprophyres minettes and lamproites). Lamprophyre dykes are o f t e n c h a r a c t e r i z e d b y s e g m e n t e d f o r m , offsetting, a n d b a n d e d or z o n e d i n t e r n a l s t r u c t u r e s i n d i c a t i n g flow d i f f e r e n t i a t i o n , m u l t i p l e i n t r u s i o n etc. ( U p t o n 1965; C u r r i e & F e r g u s o n 1970; D e l a n e y & P o l l a n d 1981).
Whole-rock composition C o m p o s i t i o n a l fields are set o u t in T a b l e s 8 - 1 0 a n d F i g s 5 - 1 0 , a n d c h e m i c a l s c r e e n s are g i v e n b y R o c k (1977, 1984, 1986); n o o v e r a l l l a m p r o p h y r e s c r e e n is feasible. H i g h levels o f HEO, C O 2 , F, C1, alkalis, Sr, T h , P, B a a n d light r a r e - e a r t h e l e m e n t s ( R E E ) , b u t n e a r - b a s a l t i c levels o f Y, Ti, h e a v y R E E a n d Sc are c h a r a c t e r i s t i c . H i g h H 2 0 a n d COE d o n o t n e c e s s a r i l y i m p l y a l t e r a t i o n o w i n g to high m o d a l c o n t e n t s o f b i o t i t e , a m p h i bole and carbonates.
195
a s s o c i a t i o n s ( T a b l e 2). ' L a m p r o p h y r e s ' f r o m o u t s i d e t h e s e a s s o c i a t i o n s (e.g. D i e t r i c h s o n 1955) n e a r l y a l w a y s v i o l a t e t h e p r e s e n t definitions.
Miscellaneous features A l t e r a t i o n is c h a r a c t e r i s t i c a l l y s e l e c t i v e : if m a f i c ( p h e n o c r y s t ) m i n e r a l s a r e altered, felsic ( g r o u n d m a s s ) m i n e r a l s are o f t e n fresh, a n d v i c e versa. Lamprophyres often exfoliate characteristically, w e a t h e r i n g to a r e d m i c a - r i c h soil. A b u n d a n t a n d v a r i e d x e n o l i t h s a n d x e n o c r y s t s ( T a b l e 12) a r e characteristic (although not diagnostic).
Identification and nomenclature of rock types closely related to lamprophyres Volatile-poor relatives of lamprophyres
Association R o c k s s a t i s f y i n g t h e f o u r p r e v i o u s c r i t e r i a are r e s t r i c t e d to a f e w w e l l - d e f i n e d p e t r o l o g i c a l
TABLE 2.
S o m e A L g r a d e into b a s a l t i c r o c k s o f s i m i l a r b u l k c o m p o s i t i o n , v i a d e c r e a s e o f [ a m p h i b o l e + biotite], loss o f g l o b u l a r t e x t u r e and, o f t e n , a p p e a r -
Summary of known lamprophyre associations Calc-alkaline lamprophyres All rock types
Alkaline lamprophyres
Mainly Mainly spessar- minettes rites
All rock types
Ultramafic Lamproites lamprophyres
Mainly monchiquites
Associations Post-orogenic calc-alkaline granitoid plutons Shoshonitic suites Appinite-breccia pipecomplexes Kimberlites Alkaline rock-carbonatite complexes Alkaline syenite-gabbro complexes Regional dyke-swarms unrelated to central magmatism Localized isolated magmatism
x x x x
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---
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9
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X
Tectonicsettings Oceanic islands Island arcs Major transcurrent faults Orogenic belts Continentalcratons Rift valleys
Partly adapted and expanded from Wimmenauer (1973a). x x , widespread occurrence; x , minor occurrence; ( x ) , rare occurrence; - - , unknown.
?, unconfirmed occurrence;
196
N. M. S. Rock
ance of plagioclase phenocrysts or groundmass olivines. A comprehensive nomenclature covering such gradations, using terms such as 'camptonitic basalt' has been given by Rock (1983). This could be extrapolated to rock types transitional between UML and 'dry' nephelinitesmelilitites ('aln6itic melilitite' (Rock 1986)) and to rare volatile-poor CAL or LL (e.g. Vaniman et al. 1985).
Mafic and felsic-enriched relatives of lamprophyres Widespread mafic-enriched rock types in lamprophyric suites (Rock 1979, 1983) can be covered by terms such as 'picritic monchiquite' and 'melaspessartite'. Such modified terms are perhaps warranted for melanocratic CAL, AL or LL (cf 'madupite') and for UML with fewer light minerals than dark minerals. Unequivocal felsic relatives of LL and UML are not yet recognized. However, intermediate to felsic rock types are commonly associated with CAL (Wimmenauer & Miiller 1985) and have been termed porphyrites or odinites (microdioritic lamprophyric rocks with plagioclase phenocrysts), semi-lamprophyres or malchites (lamprophyres with differentiation index 20-33 (Wimmenauer 1973b)), acidporphyrites (porphyrites with quartz and K-feldspar) and porphyries. A felsic relative of the Navajo minettes (U.S.A.) has been termed both 'felsic minette' and 'sanidine trachyte' (Bachinski & Scott 1980). Mugearitic-trachytic-phonolitic rocks commonly associated with AL have been termed bostonites, hedrumites, heronites, maenites etc. (Scott & Middleton 1983; Laderoute et al. 1985). Although CAL --* porphyrite ~ porphyry and AL ~ bostonite rock series partly mirror calc-alkaline andesite ~ dacite ~ rhyolite and alkaline basalt ~ trachyte series respectively, textural and compositional peculiarities may justify separate terminology (Rock et al. 1986).
Plutonic and volcanic relatives of lamprophyres As recognized by Streckeisen (1979), lamprophyres are not simply textural variants of common plutonic or volcanic rocks, partly because of their exceptionally high contents of volatiles and incompatible elements (Figs 5 and 7). Only AL broadly correspond to any common igneous rocks, having volcanic relatives among alkali basalts, basanites and nephelinites (Rock 1977) and plutonic relatives among kaersutite-bearing gabbroic rocks (Grapes 1975). Of all silicate igneous rocks, UML reach the lowest SiO2 (down to 15%, where they grade into carbonatites) and highest CaO (up to 30%), while LL attain the highest
KzO (up to 13%) combined with very high MgO (up to 27%) but with 'andesitic' SiO 2 contents (Fig. 6). Neither have recognized volcanic or plutonic relatives. CAL only have close relatives among uncommon 'shoshonitic' rocks (Rock 1984).
Discrimination between the four lamprophyre branches Much confusion has been caused by the absence hitherto of discriminatory tables or plots showing the very real differences between the four branches. The tables and plots given here are based on a computerized data base of some 5000 major, trace-element and mineral analyses (Rock 1977, 1984, 1986); this only compiles analyses of rocks called 'lamprophyres' (or one of the rock names in Fig. 1) which actually comply with the present definitions. Through the kind offices of Dr. S. C. Bergman, data have been exchanged between the present data base and his own independently compiled data base. The paucity of analyses contained in only one data base prior to combination suggested that both now contain the vast majority of available lamprophyre data to mid-1986. Statistics presented here and by Bergman (1987) are based on essentially the same data sets, apart from minor differences in emphasis or definition. Following Metais & Chayes (1964), linear discriminant analysis (Davis 1973; LeMaitre 1982) was found to distinguish both whole-rock and mineral compositions efficiently (Table 3). Plots based directly on discriminant functions, however, have three main disadvantages: (1) to many, they are abstract and arcane; (2) ideally, they have to be recalculated for each significant addition to a data base; (3) multigroup discriminant analysis is best for lamprophyres but is least amenable to simple tests of its effectiveness. Twogroup discriminant analysis was therefore performed pairwise, and Figs 3-6 are based on the oxides found to provide the best overall discrimination. Such plots are necessarily compromises since different oxides are effective in different comparisons (Ti discriminates CAL from UML biotites but not, say, CAL from AL biotites). Nevertheless, many plots approach the maximum efficiencies achieved by formal discriminant functions in Table 3, and they provide a good practical aid to correct classification. Use of all the criteria and plots proffered here requires re-naming of many rocks described in the literature. For example 'central complex
Nature of lamprophyres
197
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198
N. M. S. Rock
kimberlites' become 'aillikites' (UML) (Rock 1986), while most 'peralkaline minettes' (Velde 1967; Hall 1982), 'leucite-lamprophyres' and associated 'mica-peridotites' (Ghose 1949; Mukherjee 1961; Sarkar et al. 1980) become lamproites. The present view approximates Rosenbusch's (1897) original lamprophyre concept and attributes much confusion to later reassessments (Knopf 1936). Most spatial associations reported between different lamprophyre branches either are coincidental (Kranck 1953), reflect incorrect naming (cf. Tyrrell & Neilson 1938; Baker 1953), or lack age data to confirm consanguinity (Basta et al. 1985). The only common genetic association and gradation between different lamprophyre branches (Table 2) occurs in carbonatite-alkaline provinces (sensu Tuttle & Gittins 1966). Here some or all of AL (mainly monchiquites), UML and, more rarely, CAL (minettes) and LL may be present (e.g. Arkansas, U.S.A.). Each type generally shows its own distinct composition ( e l Tables 8 and 9), although intermediate varieties do sometimes occur. Outside carbonatitic provinces, associations and petrological gradations between different branches are virtually unknown. A note on lamprophyre-kimberfite relations
Except for very rare wholly mica-free types, kimberlites share most characteristics of lamprophyres, and their average geochemistries are in many respects similar (Fig. 5(b)). Even micapoor types may be sufficiently rich in volatile components to be retained within the lamprophyre clan. Kimberlites and lamproites are the only two primary diamond sources yet confirmed, and they intergrade via olivine-lamproites (Scott Smith & Skinner 1984a, b; Bergman 1987; Dawson 1987). Nevertheless, many apparent associations between lamproites and kimberlites have proved fortuitous (Scott 1979, 1981 ; Scott Smith & Skinner 1984b), and lamproites differ from kimberlites in carrying amphibole, glass, leucite and wadeite or related minerals, with higher whole-rock SiO2 and lower Mg, Fe, Cr and Ni (Atkinson et al. 1984, figs 8-10; Jaques et al. 1984, fig. 10; Scott Smith & Skinner 1984a, figs 6-10; Nixon et al. 1984, figs 1-6). Again kimberlites are similar enough to UML in carbonatite complexes to have been repeatedly confused (Mitchell 1979), even though most examples are readily discriminated (Rock 1986). Kimberlites contrast most with CAL and AL in both mineralogy (feldspar content) and chemistry (Mg and Si contents).
Comparativemineralogy Only 'lamprophyres' which comply with the criteria outlined here and by Rock (1977, 1984, 1986) are considered in the remainder of this paper. Considerable space is given to mineralogy because, in view of the doubtful significance of whole-rock compositions (see below), mineral chemistry may provide the most effective way of identifying and distinguishing different lamprophyres. The aim of the present section is to identify both unusual mineral species and compositional fields of common minerals which are characteristic either of lamprophyres collectively or of individual branches. Features such as complex mineral zoning (Scott 1980) and multiple compositional trends within single rock-suites (Mitchell 1981, 1985) are mostly irrelevant for these purposes, but have been considered by Rock (1977, 1984, 1986) and are briefly cited in Table 1 where particularly significant.
Mineralogical features diagnostic of the whole lamprophyre clan
Lamprophyres carry one or more of the following minerals which link them with alkaline rocks: alkali or Ti-rich pyroxenes and amphiboles, BaTi-phlogopites, Ba-K-feldspars, feldspathoids, melanitic garnet, melilite, monticellite, perovskite and silicates of alkalis with Fe, Ti, Zr, Nb etc. Significant contents of minerals rich in F, C1 and SO3 (Ca-Sr-Ba sulphates (Mitchell 1985; N6mec 1985), fluorite and even scapolite (Laderoute et al. 1985)) may be characteristic. Alkali feldspar assemblages are unusual in three ways for fine-grained sub-volcanic rocks known to cool unusually rapidly (Mauger 1984): (1) all structural types of K-feldspar (from high-sanidine to adularia) can occur (Rock 1984); (2) pure K-feldspar not infrequently coexists with pure albite (Rock 1979; Mauger 1982, 1984); (3) alkali feldspars invariably coexist with MR-rich marie minerals. These peculiarities may reflect either the high volatile fugacities in lamprophyre melts, or the disequilibrium or sub-solidus origin of some mineral assemblages (see below). Other striking peculiarities include (1) highMg and/or Mn ilmenites, (2) high-Zn spinels, (3) high-Ti amphiboles, phlogopites and (often) pyroxenes, and (4) high-Ba K-feldspars, amphiboles and phlogopites. Maximum reported Mg, Mn, Ba and Ti contents are among the highest known for the appropriate mineral (Table 4) (cf Deer et al. 1962, 1978). It may eventually be possible to set limits (more than 0.5% BaO or
Nature of lamprophyres
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N. M. S. Rock
more than 6% TiO 2 in amphiboles; more than 3% BaO or more than 8% TiO2 in phlogopites; more than 1.5% BaO in K-feldspars) above which such contents are diagnostic of lamprophyres. These peculiarities are partly but not wholly linked. For example, high mineral Ti contents contrast strikingly with near-basaltic whole-rock Ti contents (Fig. 5(b)). This implies that most Ti in lamprophyres is carried in silicates, and may explain the unusual Mg-Mn-Zn-rich, relatively Ti-poor oxide phases as a corollary. However, high Ti in phlogopites is believed to reflect physical (high-temperature, high-fO2, low-pressure) conditions (Arima & Edgar 1981; Tronnes et al. 1985), whereas high Ti in pyroxenes and amphiboles is believed to reflect chemical (Narich, Si-undersaturated) conditions which apply only to AL and UML. This may explain why AL have high Ti in all mafic silicates, whereas CAL and LL have high-Ti phlogopites but lower-Ti pyroxenes and amphiboles (Table 4; see also below).
40
Calc-alkaline lamprophyres (I0) II I I ~ o II Alkaline Improphyres ~41)
2O
% 2o
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80
84
88
92
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Mineral combinations diagnostic of each lamprophyre branch (Table 1)
LL carry a unique suite of K - F e - B a - T i - Z r silicates, including biotite (N~mec 1985), jeppeite, priderite, shcherbakovite and wadeite (Mitchell 1985), which reflect their high K/A1 ratios. Leucite and potassium titanian richterite (sensu Leake 1978) together distinguish LL from other lamprophyres (and all other igneous rocks). Minerals almost confined to UML may include a suite of N a - K - F e sulphides such as bartonite, djerfisherite, erdite and rasvumite (Czamanske et al. 1981). Ti-Zr-garnet, ferri-diopside, melilite, Mg-Mn-spinels, and monticellite distinguish UML from other lamprophyres (Rock 1986). Neither CAL nor AL are known to carry any uniquely diagnostic minerals, but the assemblages [quartz + plagioclase] and [nepheline + plagioclase] respectively distinguish them from other lamprophyres. Differences between cognate minerals common to all lamprophyre branches Olivines (Fig. 2)
Over 200 analyses reveal Fo contents L L > UML > AL. Fo9o + olivines are practically con-
FIG. 2. Distribution of mg (100 Mg/(Mg + Fe 2) at.) ratios in lamprophyre olivines. The numbers of analyses are given in parentheses. The sources are as follows. CAL: Roden & Smith 1979; Allan & Carmichael 1984--10 individual values only, indicated by dots. AL: Rock 1979; McHone & Corneille 1980; Mitchell & Janse 1982; Foley & Malpas 1984; Furnes& Stillman 1987. UML: Walton & Arnold 1970; Kresten & Edelman 1975; Vartiainen et al. 1978; Nikishov et al. 1979; Nixon et al. 1980; Platt & Mitchell 1979, 1982; Foley & Malpas 1984; Meyer & Villar 1984; J. Craven, personal communication. LL: Carmichael 1967; Scott 1979; Kuehner et al. 1981 ; Prider 1982; Scott Smith & Skinner 1984a; Venturelli et al. 1984a; Mitchell 1985. A wider range (without actual data) has been reported by Jaques et al. (1984). The LL histograms are biased by 46 determinations in Scott Smith & Skinner (1984a), but the range is unaffected by excluding these. fined to LL. NiO and CaO contents are generally 0.2%-0.4%, high CaO being noteworthy (Table 4). Clinopyroxenes (Fig. 3(d) )
Nearly 400 analyses cluster around the diopsidesalite-augite boundaries in the pyroxene quadrilateral (Rock 1984, 1986), but A120 3 contents and mg numbers provide reasonable discrimination (Fig. 3(d)). A1203 (like TiO2) contents in AL
FIG. 3. Simple plots discriminating the compositional fields on cognate minerals common to the different lamprophyre branches: (a) biotites; (b) amphiboles; (c) alkali feldspars; (d) clinopyroxenes. Variables are chosen for their high contributions towards linear discriminant functions (Table 3). The sources, which are too numerous to list individually, are available on request. The numbers of analyses are given in Table 4 (total 1200).
Nature of lamprophyres
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The individual symbols may represent several overlapping analyses. Total Fe as FeO (except in (c)) owing to large numbers of probe data. Oxides and simple oxide ratios are in weight per cent, and m g numbers are in atomic per cent. Mineral zoning, phenocryst and groundmass variations, and distinct mineral populations are ignored for the purposes of this diagram. Outlying analyses are deliberately excluded from the field-defining lines. Megacrysts and xenocrysts are excluded. Note that the biotite-phlogopite boundary in (a) is only approximate owing to inclusion of Fe203 in the abscissa. Note the logarithmic scales in (b) and (d). In (c) note that alkali feldspars are only an occasional accessory mineral in UML. A few Na-rich pyroxenes are omitted in (d) to preserve scale. Symbols are as follows: O, calc-alkaline lamprophyres (CAL); • alkaline lamprophyres (AL); V , ultramafic lamprophyres (UML); r~, lamproites (LL). For further plots separating CAL and LL minerals, and separating UML from kimberlite minerals, see Rock (1984, 1986).
202
N. M . S. R o c k
and UML pyroxenes include some of the highest known values; some have even been termed 'fassaites' (Durazzo et al. 1984). TiO2 contents vary too widely to provide effective discrimination, but t-tests and M a n n - W h i t n e y tests indicate the order AL > U M L > L L > CAL in central tendency. Na-pyroxenes have been reported in all four branches (Table 4) mainly as groundmass grains or phenocryst rims but occasionally as cores (Mullen & Petty 1985). However, too few analyses are available to draw distinctions. Note that orthopyroxenes are relatively common in a few lamproites but, as in other lamprophyres, they are almost certainly xenocrystic (Mitchell 1985).
Amphiboles (Table 1 ; Fig 3(b)) Over 250 analyses define diverse fields, and the significance of certain gaps between remains unclear. Nevertheless, the following generalizations are warranted (all amphibole names are sensu Leake (1978) and Rock & Leake (1984)). 1 AL amphiboles cover a more restricted range than the other branches; more than 90% of available analyses are 'kaersutites' or 'titanian pargasites', which are also found in some UML. 2 Hastingsitic amphiboles are common t o m a n y CAL, some AL and a few UML. 3 Titanian potassium richterites are unique to LL. 4 Tschermakitic amphiboles are known only from CAL. 5 Ti contents decrease in the general order AL > U M L ~ LL > CAL, but the great majority of lamprophyre amphiboles are titanian. 6 Some amphiboles which have the same names ('arfvedsonite-eckermannite') form more than one compositional group (Fig. 3(b)). 7 Alkali amphiboles are common in LL and U M L and present in some CAL, but are rare in AL; magnesio-riebeckite-arfvedsonites in CAL, UML and LL (Fig. 3(b)) are unreported from AL or indeed from other igneous rocks, being otherwise confined to metasomatic rocks (fenites) and metamorphosed evaporites.
Biotites-phlogopites (Fig. 3(a) ) Over 500 analyses are very diverse. Phenocryst cores and rims, and one or more coexisting groundmass micas, may all have distinct compositions (Nixon et al. 1980, 1985). Synoptic studies of micas (Mitchell 1981, 1985; Bachinski & Simpson 1984) provide further details, but the
following are perhaps the most significant generalizations. 1 A few lamprophyre phlogopites can be matched with those from kimberlites, leucitebearing rocks and mantle-peridotite xenoliths. 2 Nevertheless, the trends of kimberlite and lamprophyre micas are different, especially with respect to A1 and Ti. 3 UML and LL phlogopites often carry extensive tetraferriphlogopite (KMg3Si3Fe3010 (OH)2); LL micas reach the lowest A1 contents known (Fig. 3(a)). 4 Ca-bearing micas have been reported from CAL, AL and U M L (Table 4), but information is insufficient to determine whether this reflects impurities. 5 Only AL micas are commonly biotite (as opposed to phlogopite). 6 Ti contents are in the general order L L > AL > U M L > CAL, although LL/AL and AL/ U M L mean/median values, and all maximum values, are statistically equal.
Feldspars Plagioclases are present only in CAL and AL. Probe and optical data for both cover the range Anl-80, but 'typical' plagioclases are andesineoligoclases. Zoning can reach 50~ An content. Alkali feldspars (Fig. 3(c) and Table 5) are absent from U M L (except in accessory amounts) and from some LL. Sparse data suggest that BaO contents are in the order LL > CAL > AL, with BaO exceeding 1~ for many LL and some CAL K-feldspars. Pure albites have been reported from CAL, AL and UML. The gap in the AL field between Or 1 0 ~ - 3 0 ~ (Fig. 3(c)) may be a sampling effect.
Feldspathoids (Table 5) Analcime can occur in all lamprophyres, ranging from a few modal per cent in some CAL (Roden & Smith 1979) to a dominant phase in some monchiquites (Evans 1901). Most reliable analyses show little deviation from the standard formula, but maximum reported CaO and K 2 0 contents reach 2.14~ and 2.65~ respectively (Furnes & Stillman 1987). Nepheline and sodalite minerals occur only in AL and UML. Analyses are available only for AL (Table 5), where they are richer in Na than the Morozewicz composition but have similar amounts of excess SiO2 (Rock 1978a, 1979). Leucite occurs only in holocrystalline non-ultramafic LL, although analcime and even feldspathic ocelli have been repeatedly misidentified as leucite in other lam-
Nature of lamprophyres
203
TABLE 5. Mean compositions (weight %except for ratios) of feldspathoids, melilite and alkali feldspars in
the different lamprophyre branches Analcime ~
Total analyses4, 5 SiO2 A1203 FeO, MgO CaO Na20 K20 TiO2 (Na + K)/A1 (at.) K/(Na+K) (at.)
Calc-alkaline
Alkaline
Lamproites
1 53.8 23.3 0.79 --10.8 0.45 0.06 0.78 0.026
14 54.1 + 1.68 23.4 ___1.5 0.12+0.08(10) 0.02(2) 0.67 • 0.69,(12) 11.9 +_1.7 0.4 +0.7(12) 0.02+_0.01(4) 0.89 +_0.12(12) 0.02 +-0.004(12)
2 58.74 21,31 0.86 -0.05 9.35 0.32 -? 0.027
Nepheline 2
Melilite3
Alkaline
Ultramafic
21 46.5 _ 1.9 32.4 +_1.0 0.5 +_0.1 0.01(11) 0.80 +_0.24 15.4 +0.4 4.3 +0.9 0.04+_0.06(16) 0.93 + 0.03 0.15+0.03
8 43.4 6.4 2.9 8.8 34.3 3.8 0.1 0.1
Alkali feldspars 6 Calc-alkaline Total analyses4, 5 FeO~ Na20 K20 K/(Na+K)(at.) BaO
Alkaline
26 29 0.72{0.13-2.0}(23) 0.24{0.06-0.71} 2.82{0.27-6.3}(22) 2.9{0.12-10.4} 12.6 {7.0-16.3} 11.8{0.35-16.5} 73.7 {42.5-97.4}(22) 72.7{2.2-98.9} 0.37{0.06-1.2}(8) --
Ultramafic 3 1.73{1.1-2.8} 0.50{0.39-0.57} 15.26{14.8-15.6} 95.2{94.6-96.1} --
Lamproites 47 1.71{0.05-4.3} 0.69{0.02-2.4} 15.2 {11.9-16.8} 93.6 {76.3-99.8} 0.92{0.07-1.9}(26)
1No data for ultramafic lamprophyres; some probe analyses appear to have lost alkalis. 2Absent from calc-alkaline lamprophyres and lamproites; no data for ultramafic lamprophyres. 3Absent from all other lamprophyres. 4 See notes 2 and 4 to Table 4. s Numbers in parentheses indicate the numbers of values where less than the total analyses. 6Figures are means {ranges}; standard deviations are inappropriate given the wide range of Na/K ratios, especially for alkaline lamprophyres.
prophyres (Rock 1984; Mitchell 1985). Pseudoleucite has been confirmed in Greenland lamproites, where it occurs both as macrocrysts and in the matrix (Scott 1981 ; Thy 1985).
doluminescence and isotopic studies are urgently needed before the exact status of lamprophyre carbonates can be assessed. Nevertheless numerous workers (Frenzel 1971) have provided good evidence for primary status.
Carbonates (Fig. 4) Lamprophyres in general carry more carbonates than other silicate igneous rocks; U M L have higher CO2 even than kimberlites (Rock 1986). Probe data have begun to reveal previously unsuspected compositional variety (Fig. 4) including not only the usually reported calcite but also Fe-dolomite and (so far only in U M L and LL) breunnerite. At least four possible origins must be contemplated for such minerals: (1) primary, (2) secondary by carbonation of silicate minerals, and secondary by dolomitization of either (3) primary or (4) secondary calcite (el Le Bas 1984). As primary and secondary carbonates are difficult to distinguish in thin section, catho-
Oxide minerals (Tables 6 and 7) These are rare in LL, where high K/A1 usually causes Fe and Ti to enter priderite and related minerals (priderite itself is not known to crystallize with spinels (Mitchell 1985)). Ilmenites in U M L are unusually rich not only in Mg but also consistently in Mn (Rock 1986). Either or both characteristics may apply to CAL, AL and LL ilmenites, but data are limited (Table 6). Spinels are extremely variable. Phenocryst cores and rims, groundmass spinels, xenocrysts, and primary and secondary spinels in one rock may all be distinct (Nixon et al. 1980; Mitchell
2o4
N. M. S. Rock Co
Dolomite
Mcj
Ankeri'ie
Breunnerite
Fe
FIG. 4. Compositions of carbonate minerals in lamprophyres (Mole per cent) (symbols as in Fig. 3). The sources are as follows. AL: Scott & Middleton 1983; Foley & Malpas 1984. UML: Walton & Arnold 1970; Platt & Mitchell 1982; Foley & Malpas 1984. LL: Hall 1982. No data are available for CAL. Only a few representative calcites are plotted, as most so far analysed are nearly pure CaCO3. 1985). As data are sparse for CAL and AL, only the following tentative generalizations are offered. 1 Some CAL and AL spinels, together with some groundmass spinels in UML, appear to be 'common' igneous types.
2 Some Cr-rich spinels in CAL are similar only to those in boninites among volcanic and subvolcanic rocks (Suzuki & Shiraki 1980). 3 Zn-Mg-chromites in CAL (Wagner & Velde 1985) are similar only to spinels associated with F e - N i sulphide ores; high-Zn spinels are also known from AL (Table 7) and U M L (B. G. J. Upton, personal communication). 4 Some Mg-Cr-rich spinels in U M L and LL overlap with peridotite and kimberlite spinels only. 5 Other U M L spinels are subtly different from kimberlite spinels in both compositional range and element correlations (Mitchell 1979; Scott Smith & Skinner 1984a). 6 As with ilmenites, spinels rich in both Mg and Mn have been found to date only in U M L (Platt & Mitchell 1979, 1982).
Comparative whole-rock geochemistry Major elements
(Tables 8 and 9; Figs 5, 6 and 7)
Figure 6(b) compares lamprophyres with the most recent I U G S alkali-silica classification of volcanic rocks; it also discriminates effectively between the four branches. Apart from a few CAL (mainly kersantites and spessartites), all lamprophyres lie on the alkaline side of the Hawaiian alkaline-tholeiitic divide. In normative composition (Table 8), LL show their unique
TABLE 6. Summary of available ilmenite compositional data for the different lamprophyre branches Calc-alkaline Analyses" 2 1 MgO (~) 0.75-2.6 0.03 MnO (%) -9.17 Source Roden & Wagner & Smith 1979 Velde 1985 Area
Navajo, U.S.A.
Jersey, U.K.
Alkaline 1 1 1.86 0.31 3.00 0.59 Furnes& Scott & Stillman Middleton 1987 1983 Cape Verde Gran, Islands Norway
Lamproites 1 2.80 2.24 Mitchell & Janse 1982 Wawa, Ontario
? <8
3 2.6-3.3
--
1.1-1.9
Jacques et Mitchell al. 1984 1985 W KimJumilla, SE berley, Spain Australia
Ultramafic I Analyses2 3 MgO (%) < 0.4 MnO (~) 0.4-3.2 Source Platt & Mitchell 1979 Area Marathon, Ontario
! 7.3 0.71 von Eckermann 1968 Aln6, Sweden
3 5.1-7.7 3.74-4.45 Vartiainen et al. 1978 Sokli, Finland
3 0.10-0.15 4.7-6.1 Griffin & Taylor 1975 Fen, Norway
1
0.35 2.0 Kresten et al. 1981 Kalix, Sweden
1For plots and compilation of full analyses of UML ilmenites see Rock (1986). 2Most data are for single analyses; ranges are given where more than one analysis is available.
Nature of lamprophyres
20 5
TABLE 7. S u m m a r y o f spinel compositions present in the different lamprophyre branches Area
Reference
No. of analyses
Compositions represented
Calc-alkaline
Colima, Mexico Navajo, U.S.A. Alps, Italy
Allan & Carmichael 1984 Roden & Smith 1979 Wagner & Velde 1985
4 2 3
Ti-magnetite; Mg-chromite Ti-magnetite-ulvospinel Zn-Mg-chromite
Mitchell & Janse 1979 Scott & Middleton 1983 Foley & Malpas 1984 Ohashi 1980 Furnes& Stillman 1987 Rucklidge et al. 1980
2 2 10 1 6 2
Ti-(Mg)-magnetite Ti-(AI-Zn)-magnetite Ti-magnetite Cr-Ti-magnetite Ti-(Mg)-magnetite Ti-magnetite
Malaita, Solomons
Nixon et al. 1980
6
Aland, Finland Aln6, Sweden Ailik, Labrador
Kresten & Edelman 1975 von Eckermann 1968 Foley & Malpas 1984
2 1 10
Fen, Norway
3
Kalix, Sweden Marathon, Canada
Griffin & Taylor 1975; Mitchell 1979 Kresten et al. 1981 Platt & Mitchell 1979
Mg-ulvospinel-magnetite; Cr-ceylonite; Mg-Al-chromite Ti-magnetite Magnetite Ti-magnetite; Mg-A1-Ti-magnetite, MgAl-chromite Ti-magnetite ; Ti-A1-Mg-chromite
Ile Bizard, Canada Tanakh, Siberia
Mitchell 1979 Nikishov et al. 1979
Alkaline
Wawa, Ontario Gran, Norway Aillik, Labrador Setogawa, Japan Cape Verde Islands E Greenland Ultramafic 1
4 17
Magnetite; Ti-magnetite; Ti-picotite Mn-Mg-ulvospinel-magnetite; Ti-AI-Mgchromites; Ti-Al-chromites; Tiulvospinel-magnetites Ti-ulvospinel-magnetites A1-Mg-Cr-Ti-magnetite
Lamproites 2
Leucite Hills, U.S.A. Prairie Creek, Ar, U.S.A. W Kimberley, Australia SE Spain
Carmichael 1967 Scott, Smith & Skinner 1984a; Mitchell 1985 Jaques et al. 1984; Mitchell 1985
Ti-magnetite Ti-Mg-chromite Ti-Mg-Al-chromites; Ti-Mg-chromites; Ti-Cr-magnetite Hercynite; Cr-hercynite; Mg-chromite
Venturelli et al. 1984a; Mitchell 1985
See Rock (1986) for listing of full analyses of UML spinels. 2See Mitchell (1985) for listing of full analyses of LL spinels.
composition by combining high ks (K2SiO3) with abundant ru, and either q or lc. Many U M L are too Si poor and CO2 rich to yield norms, but high la (CazSiO4) is characteristic. Most AL have basanitic norms (i.e. a b + o r + a n + n e + d i ) , although monchiquites often carry lc, ka (KA1SiO4) and la, indicating gradation into foidites and UML. CAL vary from q to ne and from an to ac normative, but this in itself manifests their compositional peculiarity.
Trace elements (Table 10; Figs 5(b), 7-9) Available data are regrettably inhomogeneous-not least in ranges of determined elements and
analytical methods. Basic statistics are compared in Table 10, and spidergrams, whose striking similarities clearly support the concept of a lamprophyre 'clan', are compared in Fig. 5(b). The following are the most significant generalizations. 1 Lamprophyres are linked together and with alkaline rocks by their high average contents of low-field-strength incompatible elements K, Rb, Ba and Sr, but near-basaltic levels of many highfield-strength elements (Ti, Y and heavy REE) and Sc. 2 Negative Nb anomalies are present in LL and especially CAL, but not in AL or UML.
2o6
N. M. S. Rock
TABLE 8. Mean major element composition and range in normative character o f the different lamprophyre branches Calc-alkaline lamprophyresl
Alkaline lamprophyres'
Ultramafic lamprophyres
Lamproites 2
Total analyses ~' 3 754 563 245 293 Major elements (wt. ~ ) (means _ standard deviations (numbers of values where less than the total number of analyses) 3 SiO2 51.5 +4.0 (706) 41.9 +3.6 (561) 30.9 +5.8 49.5 +7.4 A120 3 14.0 +2.3 (679) 13.7 +2.7 (561) 6.8 +3.0 8.5 +2.5 Fe20 3 3.7 +2.3 (723) 5.3 +2.8 (560) 7.9 +4.5 3.7 +2.0 (257) FeO 4.9 + 1.8 (640) 6.5 +2.0 (527) 7.1 +2.6 (229) 4.1 +2.3 (229) FeO, 7.4 +1.9 (751) 10.9 +2.2 13.9 +4.3 6.1 +1.3 MgO 6.9 +2.6 (751) 7.2 +3.0 15.1 __+4.7 11.4 -t-5.6 CaO 6.6 +2.2 (728) 10.6 +2.4 15.0 -+4.8 5.4 +3.0 NazO 2.7 _1.0 (727) 3.2 _1.3 1.1 +1.1 1.3 +0.9 KzO 3.8 +2.0 (731) 2.3 +__1.1 1.8 __+1.0 7.0 +2.6 H,O § 2.7 -t-2.0 (640) 3.6 __+2.3 (513) 3.8 +3.5 (185) 2.8 ___1.9 (241) TiO2 1.3 +0.6 (713) 3.0 +_1.1 (558) 3.6 +1.7 2.7 + l . 5 P2Os 0.71+0.52(622) 0.85+0.47(520) 1.2 -t-0.8 (239) 1.1 +0.5 MnO 0.15+0.14(630) 0.21 +0.09(499) 0.26+0.17(235) 0.09+0.04(284) CO2 2.2 +2.4 (311) 2.7 +2.4 (309) 8.2 +6.9 (162) 3.0 +3.0 (137) Atomic ratios (means + standard deviations (numbers of values where less than the total number of analyses)) (Na+K)/AI 0.62+0.19(678) 0.57+0.14(561) 0.60+0.35 1.17+0.29 K/(Na+K) 47.5 +18.9 (727) 32.6 +14.6 57.1 __+21.6 77.0 +14.7 M g / ( M g + F e 2) 69.7 +_11.1 (640) 64.5 +9.6 (527) 78.1 +8.5 (229) 83.8 +6.9 (229) Mg/(Mg+Fe2)s ~ 68.7 +10.0 (751) 60.9 +9.2 72.6 +8.3 80.5 +6.9 Normative composition4 (CO, ignored in calculations and not treated as C a C O J Complete analyses s 675 561 1996 Percentage of complete analyses with the following character Q> 0 42 * * Hy>0, Q=0 25 3 * OI>O, Ne=Q=O 0 0 * Ne > O, Lc = 0 30 75 82 Lc >O, Ne=O 0 0 0 Ne>0, Lc>0 3 21 14 Ka>O * 18 85 Ka > 0 * 10 79 Ge, Ak>O * * 15 Cm>O 10 * 1 An>O, Cm=O 86 I" 97 Ac>O, Ns=O 2 * 3 Ac>O, Ns>O < 1 * < 1 Ac>O, Ks>O * * 0 Ac >O, Ns> O, Ks >O * * 0 Wo>O <1 5 0 Hm>O 10 7 0 Ru > 0 3 6 0
293 43 21 <1 10 18 8 9 * * * 21 5 20 9 44 1 14 38
* Zero by definition. ~" 100% by definition. 1As an indication of a comparability of pre- and post-1930 analyses, calc-alkaline and alkaline lamprophyre averages were calculated for (a) all available data and (b) post-1930 data only (total analyses 683 and 487 respectively. These two sets of means differed at most by only 1-2 in the first decimal place and are not therefore quoted separately. -"Mean presented merely for completeness and should be treated with extreme caution owing to large compositional differences between the different lamproite families (Bergman 1987). 3 See notes 2 and 4 to Table 4. All analyses screened for conformity to definitions in Streckeisen (1979) and Rock (1977, 1984, 1986). No analyses of tufts, breccias or pyroclastic rocks are included in the statistics. Fe normalized to Fe/total Fe (cation)= 0.30 (Rock 1986). 5 Remainder of 'total analyses' lack a value for one or more of the major oxides essential for calculating a CIPW norm (e.g. SIO2); most are partial analyses with only one to two oxide values. 6The remaining 46 of the 245 complete analyses are too Si poor and CO: rich to calculate CIPW norms. With CO2 ignored, there is insufficient SiO, to form even gehlenite molecule, but COz is too high to be calculated as CaCO 3.
Nature o f lamprophyres
207
TABLE9. Major-element (wt. %) contrasts between contemporaneous lamprophyres of different branches
within the same alkaline province or complex Aillik Bay complex, Labrador, Canada (Foley & Malpas 1984) Lamprophyre branch SiO2 A1203 FeOt MgO CaO Na20 K20 TiO2
AL 2 (19)
UML 2 (7)
33.5+2.4 22.9+2.7 8.3___2.3 3.4-t-1.3 12.5___1.5 10.9___1.1 8.5+1.9 16.0+3.0 15.0+2.1 17.8+2.3 2.7__ 1.1 0.6+0.4 2.0___0.6 2.1+0.5 4.5+1.6 3.0+0.6
Damaraland province Namibia, SW Africa (Visser 1964)
Meimecha-Kotui province, Siberia (Egorov 1970)
CAL 3
AL 3 U M L 3
UML 3
AL 3
48.44 13.73 13.66 2.58 5.80 3.10 3.83 2.25
41.16 16.87 10.50 5.30 8.12 3.00 3.00 3.00
38.00 10.69 13.70 10.12 14.76 5.83 2.60 1.10
44.25 13.54 9.42 7.62 9.58 4.13 2.20 2.88
33.28 18.51 16.08 6.02 12.03 0.55 3.02 2.30
Central Arkansas province, U.S.A. (Steele & Wagner 1979; Scott Smith & Skinner 1984a) 1 AL 3 (14)
UML 3
LL 2 (24)
43.2-t-2.7 39.00 49.9___7.9 13.2-t-2.0 14.15 4.8-t-1.0 10.4-t-2.2 11.02 8.6___2.3 7.8-t-3.1 8.80 24.5+7.0 11.8-1-2.9 17.40 4.7___1.3 1.8+ 1.4 3.20 0.4+__0.3 3.0__ 1.3 1.40 3.1_ 1.2 3.4___0.8 3.26 2.6_+0.7
1Most lamprophyres recorded in the literature have been renamed according to the definitions in the present paper. The Prairie Creek 'kimberlite', for example, is a lamproite (Bergman 1987; Scott Srrtith & Skinner 1984a), and most 'ouachitites' are ultramafic lamprophyres and not alkaline lamprophyres. Means include data from several unpublished theses. 2 Mean ___standard deviation (number of analyses in brackets). 3 Single analysis of individual dyke.
tCorb~
y
~
* Kimberlites
400
-
?
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6
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oo-
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x ...... x o-----13 + ....... + 0-----o
/ !
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CAL AL UML LL
....
& 9'Common'bosic/.~^,/x kT.2.,-,L 0 (o)
igneous rocks 1 _- __...~....-~Averoge igneous rock " ', I'. i I ~ 2 4 H20
50
* I
6
,i':'//7-
2 o
~.,3
~,o FIG. 5. Comparison of lamprophyre whole-rock compositions with other igneous rocks. Fields of individual analyses would cover much of the plots, so only mean values are shown. The lamprophyre means are taken from Tables 8 and 10. (a) Volatile contents. Shaded field based on LeMaitre's (1976) means for various common basic igneous rocks. Symbols as Fig. 3. Various kimberlite means from Dawson (1980) and Bergman (1987). (b) MORB-normalized 'spidergrams" using the scheme of Pearce (I 982, 1983). Kimberlite (KIMB) and alkali basalt (AOB) means, from Bergman (1987).
*
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l
t%, I/ x--'~,7*, 7. 0"5 I
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208
N. M. S. Rock
!~
(b)
21 (a)
]
i'!'
-
-
r
20 (c)
MgO
40 Si02
60
o
(d) AI203
CoO 9
2b
40
SiOe
60
SiOJlO
TiO z • 4
FIG. 6. Simple plots discriminating the whole-rock compositional fields of lamprophyres and comparing them with recent chemical classifications o f ' c o m m o n ' igneous rocks. The numbers of analyses are given in Table 8 (total, 1850). Fields for each lamprophyre branch enclose 95~ of available analyses to avoid outlying and possibly inaccurate data. The means from Table 8 are shown using the symbols of Fig. 3. All oxides are in weight per cent. The andesite classification fields of Gill (1982) are also shown in (a), and the tholeiite-alkaline divide of Macdonald & Katsura (1964) and the recent IUGS proposal for a definitive total alkali-silica classification of volcanic rocks (LeMaitre 1984) are included in (b). Note that this proposal applies to analyses with original H20 and CO, contents less than 2%, recalculated volatile free, and its comparability with lamprophyre data may therefore be limited. The field labelled 'basanite' covers tephrite to tephriphonolite; 'trachandesite' covers hawaiite, benmoreite etc. (c) and (d) are plots providing better discrimination between branches not so efficiently separated by (a) and (b). CAL and LL cannot be efficiently separated by whole-rock composition alone unless five or more chemical variables are used. For further discriminatory plots see Rock (1984, 1986).
Nature of lamprophyres "OJ
209
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FIG. 8. Rare-earth profiles (total 150) normalized to the chondrites of Nakamura (1974). Shaded fields and vertical bars indicate ranges. The means are shown alone for invariable spectra. The methods of analysis (where stated) are as follows: neutron activation (INAA); induced coupled plasma (ICP); X-ray fluorescence (XRF): spark-source mass spectrography (SSMS). The sources for (a) calc-alkaline lamprophyres are given in fig. 3 of Rock (1984). The sources for (b) alkaline lamprophyres are as follows: <~, mean and range of six monchiquites/ camptonites, Monchique, Portugal (Rock 1978a, 1979) (SSMS); Q, mean of three camptonites, Spanish Peaks, U.S.A. (Jahn et al. 1979) (INAA); ]~, mean and range of 18 camptonites, Sunnhordland, Norway (Faerseth et al. 1976) (XRF) (La to Nd only); r-q, one monchiquite, Hopi, AZ, U.S.A. (Nicholls 1969) (XRF) (La to Sm only): ,, one camptonite, Girnar, India (Paul et al. 1977) (INAA); ~-, two monchiquites, Callander Bay, Canada (Cullers & Medaris 1977) (INAA); ~I,, monchiquite (misnamed 'aln6ite'), Oka, Quebec (Eby 1975) (XRF); x, camptonite, Eilean Donain, Scotland (Morrison 1980) (INAA). The sources for (c) lamproites are as follows: 9 'peralkaline minette' (lamproite), Pendennis, England (Hall 1982) (ICP); x, three Leucite Hills lavas, WY, U.S.A. (Kay & Gast 1973) (IN AA) ; ~, mean and range of 48 lamproites, Holsteinsborg, Greenland (Scott 1979) (XRF) (La and Ce only); ~, mean and range for four fitzroyite family lamproites, NW Australia (Nixon et al. 1984) (ICP); ornamented field, range for 15 jumillite family lamproites, SE Spain (N ixon et al. 1984) (ICP). The sources for (d) ultramafic lamprophyres are given in fig. 6 of Rock (1986). Other data (Thompson et al. 1984) are omitted, as they substantially overlap those shown.
3 Sr and Ba c o m m o n l y reach higher concentrations than in other silicate igneous rocks; the highest recorded Sr content (7275 ppm) exceeds Wedepohl's (1978) m a x i m a even for carbonatites.
4 Th, Zr and Rb also reach higher values than in other igneous rocks, except for exotic types of pegmatites and agpaites. 5 Although Y generally has levels characteristic
!
212
N. M. S. Rock 5 Minette and UML fields overlap substantially. 6 There is some overlap between lamprophyre (CAL, UML) and kimberlite fields. 7 On average, lamprophyres show steeper REE profiles than alkali basalts, but have sub-equal contents of the middle REE (around Gd-Er).
of mid-ocean ridge basalt (MORB), it also achieves a few high values (more than 150 ppm in AL and UML). 6 The mean AL spidergram has an almost identical shape to the alkali basalt spidergram, but at enrichments higher by a factor of 1.5-2; this confirms the view (Rock 1977) that AL are volatile- and incompatible-element-enriched alkali basaltic rocks. 7 The mean kimberlite spidergram shows greatest similarities with that of UML.
Isotopic data
Combined Sr and Nd isotopic data are only available for lamproites (Table 11). Sr ratios for AL and UML are consistently low, but CAL display some tendency towards higher values (Fig. 10). Virtually no Pb or stable isotope data are available. Integrated S r - P b - N d - O isotopic studies are urgently needed.
Rare earths (Figs 8 and 9)
Apart from one or two 'anomalous' examples, all lamprophyres show comparable REE spectra, with sub-equal contents of middle REE. Nevertheless the following differences are apparent from Fig. 9.
1 Plagioclase-bearing lamprophyres (spessartires and camptonites) have substantially lower Sm and La/Yb than plagioclase-free types. 2 The order of both light REE enrichment and slope in general is LL > UML > CAL > AL. 3 The light REE thus increase both with increasing K / N a and K-feldspar/plagioclase ratios (cf. Cullers & Graf 1984). 4 Different lamproite families define quite distinct fields.
t
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Miscellaneous petrological characteristics of the whole lamprophyre clan Xenolith and megacryst/macrocryst assemblages (Table 12)
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FIG. 9. Comparison of lamprophyre REE spectra with those of kimberlites and carbonatites. Symbols as in Fig 3. Note the log-log scales. Sources as in Fig. 8. The majority fields are labelled. Kimberlite field after Nixon et al. (1980).
Lamprophyres characteristically carry abundant crustal xenoliths, which sometimes appear to sample inferred subcrop faithfully (Read et al. 1926). Many lamprophyres, moreover, grade into (or are associated with) breccias, ~zarrying abundant country-rock material (Nixon et al. 1980; Rock 1983). All four lamprophyre branches have also furnished spectacular mantle xenolith suites: CAL, Navajo, U.S.A. (Ehrenberg 1979); AL, Scottish Highlands (Rock 1983); UML, Malaita, Solomons (Dawson et al. 1978); LL, W Kimberley, Australia (Atkinson et al. 1984). The number of recent reports suggests that lamprophyres may even carry proportionately more mantle material than basalts, taking into account relative hostrock abundances. For example, monchiquites have yielded some of the most varied and abundant mantle material among Scottish Permo-Carboniferous alkali basaltic intrusions (Rock 1983; Upton et al. 1983). Curiously,
TABLE 11. Combined Sr and N d &otopic data for lamproites Lamproite source W Australia W Australia Leucite Hills
Reference McCulloch et al. (1983) Nixon et al. (1984) Vollmer et al. (1984)
143Nd/144Nd 0.5110-0.5114 0.5119-0.5121 0.5118-0.5121
87Sr/86Sr 0.7104-0.7187 No data 0.7053-0.7061
Nature of lamprophyres 8-
213
Calc-alkaline lamprophyres (28)
4 I
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Alkaline lamprophyres (27)
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Ultramafic lamprophyres
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whereas the mantle xenolith (?diamond) contents of LL and UML imply depths of origin greater than other igneous rocks (except kimberlites), their contents of leucite and melilite imply crystallization at among the l o w e s t pressures. Varied macrocryst suites are also characteristic (Table 12). Zoned ovoid Ti-phlogopite megacrysts ('eggs') up to 3 cm long (Le Cheminant & Le Cheminant 1985) appear so far to be unique to lamprophyres. Opinions differ as to the cognate or xenocrystic origin of such phases (see references in Table 12).
FIG. 10. Distribution of initial 87Sr/S6Sr ratios in lamprophyres. The numbers in brackets indicate the total numbers of values (isochron intercepts are taken as single values). The numbers within the histograms identify the data sources and localities as follows. Calc-alkaline lamprophyres: 1 kersantites, Massif Central, France (Cantagrel et al. 1970);2 kersantite and spessartite, Sines, Portugal (Canilho 1971); 3 minettes, Spanish Peaks, Colorado (Jahn et al. 1979); 4 minettes, Navajo, U.S.A. (Powell & Bell 1970; Roden 1981); 5 spessartites & kersantites, Cortlandt, U.S.A. (Ratcliffe et al. 1983); 6 minettes, Lake Huron, Canada (Van Schmus 1971); 7 'lamprophyre', probably lamproite, NW Alps (Dal Piaz et al. 1979). Alkaline lamprophyres: 8 camptonites and monchiquites, Monchique, Portugal (Rock 1976); 9 camptonites, Spanish Peaks, U.S.A. (Jahn et al. 1979); 10 monchiquite, Magnet Cove, U.S.A. (Powell et al. 1966); 11 camptonites and monchiquites, New England, U.S.A. (McHone et al. 1976); 12 'lamprophyre' (camptonite), Girnar, India (Paul et al. 1977); 13 camptonite, Predazzo, Italy (Lucchini & Mezzetti 1969); 14 monchiquite, Pyrenees, Spain (Vitrac-Michard et al. 1977); 15 sannaite, Jombo, Kenya (Rock 1976); 16 monchiquites, Hopi, U.S.A. (Powell & Bell 1970). Ultramafic lamprophyres: 17 alnoite, Oka, Canada (Powell et al. 1966); 18 alnoitic rock, Rangwa, Kenya (Rock 1976); 19 damtjernites, Fen, Norway (Griffin & Taylor 1975); 20 aillikite, Ytteroy, Norway (Priem et al. 1968);21 alnoites, Montana, U.S.A. (Powell & Bell 1970); 22 alnoites, Alno, Sweden (Brueckner & Rex 1980); 23 aillikites and alnoites, Frederikshaab, Greenland (Hansen 1981). Lamproites: 24 SE Spain (Powell & Bell 1970); 25 Leucite Hills, U.S.A. (Powell & Bell 1970; Vollmer et al. 1984); 26 W Kimberley, Australia (Powell & Bell 1970); McCulloch et al. 1983; Jaques et al. 1984).
Globular structures
Felsic 'globular structures' (Phillips 1973) ('amygdales', 'ocelli', 'patches', 'varioles' etc. of other workers) are found in all lamprophyre types (summaries in Popov (1972) and Rock (1977, 1984, 1986)). Although similar structures occur in basalts, they are much less widespread. In Scotland, for example, they distinguish PermoCarboniferous AL from Tertiary and Devonian basalts (Bailey et al. 1924). Globular structures differ from their host rock in their lower colour
2 I4
N.M.S.
Rock
TABLE 12. G e n e r a l i z e d a b u n d a n c e o f x e n o l i t h i c a n d m a c r o c r y s t i c / m e g a c r y s t i c m a t e r i a l in the d i f f e r e n t lamprophyre branches
Calc-alkaline lamprophyres 1
Alkaline lamprophyres 2
Ultramafic lamprophyres 3
Lamproites 4
( X )6 --( x )6
X X
X X
X
x x x
x x x
x x x x
x x x x x x x x x x x
x x x x x x x x x
9 x x x x x x x
? x x x x x x x
--
( x )9
?8
x x
x x x
x x --
x x x --
x x x x
x x x x --
(x) -x x (x )
x x
x ? x
Mantle-type ultramafic xenoliths s
Lherzolites, harzburgites Wehrlites, websterites Phlogopite-hornblende-rich rocks Eclogites/garnet-peridotites Other xenolithic material
Al-rich restites Miscellaneous (meta-)sediments Basalt, gabbro, amphibolite Granitoids Deep crustal granulites Country rocks Megacrysts and xenocrysts 7 Diamond Spinel, ilmenite Orthopyro xene, olivine Garnet ( M g - C r )
Kaersutite 1~ Apatite 1~ Clinopyroxene 1~ Sanidine-anorthoclase l~
-x x --
x x x x
x x x x
Quartz, plagioclase Corundum, kyanite, staurolite, sillimanite, A l - F e garnet
x x x x
x x x
x x , widespread; x , recorded; ( x ), very rare; - - , unknown; ?, unconfirmed. 1For sources see table 5 of Rock (1984); additional information in Kramer & Seifert (1984) and Jaques & Perkin (1984). 2Sources: Carstens 1958; Bayrakov 1964; Yeremenko & Shvakova 1969; Dobretsov & Dobretsova 1969; Azambre 1970; Brooks & Rucklidge 1973; Val'ter & Yeremenko 1974; Brooks & Platt 1975; Chapman 1975; Wallace 1975; Janse 1977; Rock 1977, 1983; Faerseth 1978; Leavy 1980; Larsen 1981; Praegel 1981; Mitchell & Janse 1982; Upton et al. 1983. 3 For sources see table 8 of Rock (1986). 4Sources: Carmichael 1967; Kay et al. 1978; Barton & Van Bergen 1981 ; Nixon et al. 1984; Jaques et al. 1984; Bergman 1987. 6 Navajo province, U.S.A. only (e.g. Ehrenberg 1979). vThose generally attributable to the disaggregation of xenoliths are in italics. 8 Aln6, Isle Bizard diatremes (Rock 1986). 9 Wandagee (W. Australia) only currently known example. 10 Probably includes both cognate megacrysts (phenocrysts) and xenocrysts.
index. M a n y o f t h e m g r a d e f r o m s h a r p l y d e f i n e d s u b s p h e r i c a l ocelli, w i t h d e l i m i t i n g t a n g e n t i a l biotite, to v a g u e p a t c h e s , b a r e l y d i f f e r e n t i a t e d f r o m t h e i r host ( R o c k 1979, 1983). The mineral compositions of such structures are q u i t e d i s t i n c t in t h e f o u r l a m p r o p h y r e b r a n c h e s ( T a b l e 1). G i v e n g o o d e x p e r i m e n t a l e v i d e n c e for s i l i c a t e - c a r b o n a t e liquid i m m i s c i bility ( H a m i l t o n et al. 1979), t h e c a r b o n a t e - r i c h s t r u c t u r e s in U M L p r o b a b l y r e p r e s e n t i m m i s c i b l e c a r b o n a t e d r o p l e t s ( ? c a r b o n a t i t e differentiate).
O p i n i o n s differ m u c h m o r e s h a r p l y o v e r t h e f e l d s p a t h i c s t r u c t u r e s in C A L a n d A L , h o w e v e r , with some workers advocating silicate-silicate liquid i m m i s c i b i l i t y ( P h i l p o t t s 1972, 1976,; E b y 1980) a n d o t h e r s a d v o c a t i n g s e g r e g a t i o n o f latestage liquid into vesicles ( C o o p e r 1979; F o l e y 1984; M a u g e r 1984). A p r o f u s i o n o f g e n e t i c suggestions exists for c a r b o n a t e - f e l d s p a r - c h l o r i t e - q u a r t z s t r u c t u r e s in "CAL a n d A L ( R o c k 1984), b u t i d e n t i f i c a t i o n as a m y g d a l e s or varioles has m u c h to c o m m e n d it ( J a f f e 1952; Z i m m e r l e
Nature of lamprophyres 1977). No assessment of these views can be made at least until microprobe and experimental data are available.
Heteromorphism CAL and AL rock types (Fig. l) are heteromorphic, as indicated by the coexistence of nearly isochemical rock types within single dykes, and by both local and global compositional merging (Rock 1977, 1984). Trivial heteromorphism, in which one rock type is merely a chilled version of another, is also found between holocrystalline and glassy lamprophyres (camptonites-hyalomonchiquites, orendites-wyomingites, jumillites-verites) (Rock 1977; Vollmer et aL 1984). Nevertheless, significant heterogeneities within the different branches also occur (Rock 1984, 1986). Heteromorphism is characteristic of lamprophyres, but it is not universal.
Global distribution Distribution through geological time Because they are partly characterized by texture and 'mobile'-element contents, lamprophyres are unlikely to survive even low-grade metamorphism. The older and/or more foliated they are, the more they may be confused with, say, microdior, ites (as in Johnson & Dalziel 1966). Some Precambrian 'lamprophyres' (Watson 1983 and personal communication) and many 'meta-lamprophyres' (Steiner 1984) are very doubtful. Hence the detailed distribution of lamprophyre magmatism through geological time may never be completely defined. Nevertheless, available age data show that lamprophyres are widely distributed from early Precambrian to Recent times (Table 13). As with kimberlites and carbonatites, certain times appear to have been particularly favourable for their emplacement: for example, the PermoTriassic and Jurassic for AL, and the Caledonian and Hercynian orogenies for CAL. LL alone appear to be restricted to a few age bands (Table 13). Some regional lamprophyre dyke-suites span periods of as much as 100 Ma in several distinct tectono-magmatic events (Baxter & Mitchell 1984).
Volumetric abundance Although individual intrusions are of small volume, numerous regional lamprophyre swarms, with thousands of dykes covering many thou-
215
sands of square kilometres, may be volumetrically equivalent to substantial plutons (Rock 1983). Watson (1984) estimated that lamprophyres and associated felsic minor intrusions make up 15% of Scottish Late Caledonian igneous rocks (i.e. several batholith equivalents), and crustal extensions reach 50% locally (Rock et al. 1985, 1986, 1987; Barnes et al. 1986). Although quantitative comparisons are not yet possible, lamprophyres probably have an aggregate global bulk considerably greater than kimberlites, carbonatites or leucite-bearing rocks and at least equal to nepheline syenites. Within lamprophyres, the order of abundance is almost certainly CAL AL >>UML g LL.
Range of tectonic settings (Table 2) Lamprophyres, like other alkaline rocks, are most abundant in continental rifts, failed arms of triple junctions (Brooks & Platt 1975) and on some oceanic islands (De Almeida 1955, 1961 ; Mitchell-Thom~ 1976). However, they also occur widely in orogenic zones and their peripheries (Himalayas (Viterbo & Zanettin 1959); Alps (Dal Piaz et al.1979); Pyrenees (Barrab~ 1952; Rock 1982); Caledonides (Richey 1939); Hercynides (Knill 1982)), in island arcs (Japan (Yagi et al. 1975); Solomons (Nixon et al. 1980)), in passive to destructive continental margins (W U.S.A. (Snaveley & Wagner 1961)), in anomalous uplifted fragments o f oceanic crust (Gorringe Bank (Cornen 1981, 1982); Bermuda (Aumento & Ade-Hall 1973; Aumento et aL 1974)) and near major transcurrent faults (Alpine Fault, New Zealand (Cooper 1979)). Collectively, therefore, lamprophyres are associated not only with intra-plate magmatism, but also with divergent, convergent and even passive plate-margin magmatism.
Petrogenetic significance of lamprophyres Throughout this section, the words primitive, primary and parental are used in the sense of Rhodes (1981). Previous petrogenetic theories for lamprophyres are legion (Rock 1977, 1984, 1986; Bergman 1987). This section concentrates on (1) unifying petrogenetic characteristics of the whole clan and (2) the increasingly recognized possibility that lamprophyres represent primary-mantle melts--by analogy with widely accepted models for alkaline basic to ultrabasic rocks (Frey et al. 1978). Since 'an essential requirement for any primary basalt candidate is that the bulk compo-
216
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Nature of lamprophyres sition be equivalent to that of a liquid' (Rhodes 1981), we first examine relationships between lamprophyres and their precursor 'magmas'.
Do 'lamprophyremagmas' exist? Owing to their abundant phenocrysts, lamprophyres were assumed by Bowen (1928) never to correspond to fully-liquid magmas. Although lamprophyre phenocrysts need not in fact represent liquidus phases but merely the crystals which grew fastest in such a volatile-rich medium, evidence such as a lack of holohyaline lamprophyres can be used to support Bowen's conclusion (Rock 1984). Lamprophyres might then be attributed to crystal-laden, volatile-charged fluids with only minimal silicate-melt component. These might be generated by gas-phase metasomatism rather than by actual fusion of mantle material, and emplaced by fluidization rather than by normal intrusion (as in Bailey 1984, 1987). However, the lamprophyre combination of [groundmass + phenocrysts + abundant micas, carbonates etc.] could also be taken as a 'frozen' sample representing [melt + suspended crystals + volatile phase]--in fact a 'complete' magma system. Lamprophyres would then be 'closer to intratelluric magma compositions than aphanitic mafic extrusives, the more fashionable yardstick of parental magmas' (Hughes 1982), because they alone retain some representative of the volatile phase. Thus, if it is assumed a priori that primary magmas are fully liquid at the time of generation, natural lamprophyres cannot represent primary magmas. If intratelluric magmas are allowed to be crystal laden and volatile rich, then lamprophyres could uniquely approximate such magmas and many glassy lavas would have to be regarded as fractionated. These alternatives echo the controversy over basalt genesis (i.e. primary origin picrite versus fractionation from primary magma). Two so far largely unaddressed questions are critical here. 1
Are lamprophyre mineralassemblagesprimary ?
Luhr & Carmichael (1981) and Allan & Carmichael (1984 and personal communication) have presented evidence that recent lamprophyres may be erupted as relatively volatile-poor magmas, and that their volatile-rich mineralogy is subsequently acquired by interaction with an aqueous (?deuteric, connate) phase during cooling. The bulk compositions of older lamprophyres may not therefore represent magmas but be determined by syn- or post-magmatic exchanges. 2 A re lamprophyre mineral assemblages in equilibrium ? Commonly coexisting mineral pairs such
217
as analcime with K-feldspar and forsterite or diopside with quartz or alkali feldspars, together with resorbed phenocrysts (especially phlogopite in LL (Mitchell 1985)), may point to disequilibrium. Indeed, many are coming to regard lamprophyre assemblages as 'frozen' mixtures of products and reactants from incompleted reactions (Yoder 1979; Rock 1979, 1986). Unfortunately, while conflicting interpretations of basalt genesis remain unreconciled, and until experimental work is undertaken not only on natural lamprophyres but also on volatile-rich synthetic systems which closely approximate natural lamprophyres, we can only attempt qualitatively to assess whether lamprophyres possess other characteristics consistent with primary status.
Do lamprophyresrepresent primarymagmas? Geochemistry
Among whole-rock parameters widely accepted as indicating primary status are high mg numbers, Ni, Cr and Co. Although acceptable values differ among authors--'primary' mg ranges, for example, from 68%-75% to 70%-83%, and Ni from 240-390 ppm to 200-450 ppm (Rhodes 1981)Fig. 7 shows that many lamprophyres indeed show these 'primary' features. However, many others show lower values and appear to be contaminated or fractionated (Bergman et al. 1985; Laderoute et al. 1985). Absolute contents of MgO around 11%-13% have also been taken (Hart & Davis 1978) to indicate primary status, in which case 10% of analysed CAL, 15% of AL, 20% of UML and 25% of LL would be candidates. However, O'Hara et al. (1975) regarded picritic magmas with 20%-25% MgO as the only primary magmas, in which case only a few lamproites would be candidates. The extreme incompatible-element contents in lamprophyres (Figs 5(b) and 8) present theoretical problems for primary status, for they cannot be generated from mantle materials using known partition coefficients and melting models. Some workers have inferred extreme fractionation of a garnet-bearing assemblage, such as eclogite, to explain steep REE profles (O'Hara et al. 1975). However, similar problems arise with other alkaline basic rocks (Cullers & Graf 1984), and can be reconciled via disequilibrium melting involving a phase rich in REE or a mantle source previously enriched in REE (Venturelli et al. 1984a, b; Hertogen et al. 1985). In view of the inconsistent and sparse Sr-Nd data available (Table 11) and the occurrence of
218
N. M. S. Rock
high Sr ratios in the Australian lamproites (despite their content of diamond and mantle peridotite nodules), the interpretation of unsupported Sr data for other lamprophyres (Fig. 10) must be circumspect. Ratios in AL and UML seem low and consistent enough to rule out both crustal involvement and 'enriched' mantle sources. However, the bias towards higher Sr ratios in CAL has already been interpreted variously in terms of crustal involvement (Powell & Bell 1970) or mantle effects (Roden 1981). Negative spidergram Nb anomalies in CAL and LL (Fig. 5(b)) are characteristic of destructive plate-margin magmatism (Pearce 1982, 1983) and have been interpreted (Thompson et al. 1984) as indicating assimilation of subducted sediment slabs. In that AL and UML have no Nb anomalies, this is at least consistent with the tentative indications of Sr isotopes. However, it is as yet unclear how much weight can be attached to such 'anomalies', given the rather artificial method of data display. Mineralogy
The high Fo content of most lamprophyre olivines (rag number 86-94) (Fig. 2), and the high mg numbers, Cr and Ni of lamprophyre phlogopites have all been taken to support primary status (Jones & Smith 1983; Le Cheminant & Le Cheminant 1985). However, some olivines (especially in CAL and AL) are relatively Fo poor (Fig. 2), and the relative rarity of olivine phenocrysts in CAL could even be taken to deny primary status absolutely. Experimental data
Extensive data are available only for lamproites (Barton & Hamilton 1978, 1979, 1982; Arima & Edgar 1983), where they are consistent with primary origin provided that (1) the mantle source has been previously depleted in garnet and clinopyroxene and enriched in phlogopite and (2) melting occurs under high H20 and F but low CO2 pressures. However, very different melting conditions must apply to UML (probably near kimberlites) (Rock 1986) and to AL (probably near basanites)(Nathan 1968; Rock 1978b). Xenoliths
The abundant mantle-type xenoliths in many lamprophyres (Table 12), and especially the diamond in LL (and ?AL and ?UML), would seem to indicate primary status. However, crystalladen lamprophyres probably have high yield strengths, enhancing ability to carry xenoliths (Sparks et al. 1977; Spera 1980). Xenolith-bearing
lamprophyres might still therefore have fractionated even after xenolith incorporation (cf peridotite-bearing phonolites (Wright 1971)). Larsen ( 1981), for example, argued that one monchiquite had fractionated despite its abundant mantle xenoliths and mg value (71-74) because its olivine and Cr-diopside megacrysts must have crystallized from a magma which was richer still in Mg. Crystallization temperatures and ranges
The consistently late-stage intrusion of lamprophyres and their general lack of contact-metamorphic effects both tend to imply lowtemperature emplacement, which is more consistent with a fractionated condition. However, relatively few dyke-rocks of any composition cause extensive contact metamorphism, and a few lamprophyres have caused pyroxene-hornfels or even sanidinite-facies metamorphism (Rock 1983, 1984; Macdonald et al. 1986). More suggestive still of fractionation or accumulation are the non-cotectic status and long crystallization-temperature ranges of the few natural lamprophyres studied to date (Nathan 1968; Rock 1978b; Esperan~a & Holloway 1985). Tectonic setting
The occurrence of AL and UML on oceanic islands proves that continental-crust involvement is not a prerequisite for their genesis. However, neither CAL nor LL occur in an oceanic setting, and CAL repeatedly show such a close spatial, temporal and almost certainly genetic relationship with granitic plutonism, together with a characteristic abundance of crustal xenoliths (Table 12), that some crustal involvement seems irrefutable in their genesis (Rock 1980, 1984). Although many of the above uncertainties apply to the petrogenesis of other alkaline magmas, they do tend to show the following. 1 Lamprophyres are as diverse genetically as they are compositionally--differing sources and/ or melting conditions can apply even to different examples of one branch within a single igneous complex (Vollmer et al. 1984). 2 Some lamprophyres in each branch approximate primary-mantle melts derived under varying conditions: CAL, Navajo, U.S.A. (Ehrenberg 1979); AL, Wawa, Canada (Mitchell & Janse 1982); UML, Malaita, Solomons (Nixon et al. 1980); LL, W Kimberley, Australia (Jaques et al. 1984). Indeed, LL and UML (along with kimberlites) encompass melts with the deepest origins of all igneous rocks. 3 Other lamprophyres (especially CAL) originate by volatile enrichment, fractionation, mix-
Nature of lamprophyres ing and other multistage processes (Rock 1977, 1984, 1986); Clarke et al. 1983 ; Van Bergen et al. 1983; Macdonald et al. 1986; Nixon et al. 1986).
Differentiation in lamprophyres, and their possible status as parental magmas to hydrous alkaline intrusive rock-suites Whether or not they are primary, lamprophyres repeatedly have the most primitive compositions in a given rock-suite. That they may in turn represent the parental magmas to such suites is supported by several lines of evidence. 1 Availability. In numerous combined major+ minor intrusive alkaline suites (Lovozero and Khibina, U.S.S.R. (Grigor'eva & Savitskij 1979); Borralan, Scotland (Sutherland 1982); Sokli, Finland (Vartiainen et al. 1978)) lamprophyres are the only primitive rocks, the associated major intrusions being exclusively syenitic, ijolitic or carbonatitic. Again, in some entire alkaline provinces (Monteregian and Iberian Provinces (Philpotts 1974; Rock 1982))lamprophyres represent the only primitive liquids intruded throughout the province. 2 M a n t l e xenolith content. In numerous rocksuites (Fen (Griffin 1973); Pyrenees (Azambre 1979)) lamprophyres are the only rocks carrying such xenoliths. 3 Differentation. Single lamprophyre intrusions may carry intermediate to acidic veins, ocelli or schlieren (Mukherjee 1961 ; Philpotts 1972; Rock 1984) or show composite structure (Sathe & Desai 1968 ; Rock 1984, table 8). Entire dyke-suites may show co-magmatic series from primitive lamprophyres to felsic rock types (N~mec 1973; Rock et al. 1986, Barnes et al. 1986).
For these and other reasons, lamprophyres have been equated with parent magmas for the following wide range of hydrous intrusive suites. CAL: 'shoshonitic' hornblende-rich pyroxenite-diorite-norite-granodiorite-granite plutons
219
(Cortlandt, U.S.A. (Ratcliffe et al. 1983)), potassic syenites (Arkansas, U.S.A. (Mullen & Petty 1985); Biella, Italy (Dal Piaz et al. 1979)) and even calc-alkaline granites (Japan (Suzuki & Shiraki 1980); Scottish Caledonides (Macdonald et al. 1986)). AL (camptonites); kaersutite-biotite-rich gabbro-(nepheline) syenite plutons (Monteregian Hills (Philpotts 1974; Bedard et al. 1985)). AL (monchiquites) and UML: ijolite-syenitecarbonatite complexes (Fen (Griffin & Taylor 1975)). One of the most intriguing cases is the Monteregian province (Philpotts 1974), where lamprophyre dykes gradually change from camptonite in the E of the province via monchiquite to aln6ite in the W; the associated plutons change concomitantly from saturated gabbro-syenite via essexite-foyaite to ijolite-carbonatite. The plutonic rocks are cumulates (Eby 1987), and the lamprophyres are among the few candidates for primary magmas. Although Philpotts' view that the three parent lamprophyre types generated the three types of plutonic suites has proved oversimplified (Eby 1985a, b, 1987), there is undoubtedly a genetic connection between representatives of three lamprophyre branches and of all the most globally widespread types of alkaline plutonic suite. In conclusion, lamprophyres are almost always primitive, frequently primary and sometimes parental. They are considerably more abundant and significant than hitherto recognized. This series of reviews will hopefully stimulate new multidisciplinary studies of these fascinating rocks. Stimulus from a more pecuniary direction may well come from the increasing number of deposits of precious stones (Atkinson et al. 1984; Meyer & Mitchell 1985) and metallic minerals (Haughton & Frommerze 1930; Kumar 1968; Burkhalter et al. 1971 ; Rosseykin & Razhmanov 1971 ; Makeyev & Yefimov 1972; Sarkisyan 1973; Shchukin 1974; Daniyelyants & Lyakhov 1975) now reported in (or associated with) lamprophyres !
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--
mechanisms and solubility of titanium in phlogopites from rocks ofprobable mantle origin. Contrib. Mineral. Petrol. 77, 288-95. & - - 1983. A high pressure experimentalstudy
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on a magnesian-rich leucite-lamproite from the West Kimberley area, Australia: petrogenetic implications. Contrib. Mineral. Petrol. 84, 228 34. ATKINSON, W. J., HUGHES, F. E. & SMITH, C. B. 1984. A review of the kimberlitic rocks of western Australia. In: KORNPROBSTJ. (ed.)Kimberlites I: Kimberlites and Related Rocks, pp. 195-224. Elsevier, Amsterdam. AUMENTO, F. & ADE-HALL, J. M. 1973. Deep-drill1972. Petrology of the Bermuda drill core. EOS 54, 485. --, REYNOLDS, P. M. & GUNN, B. M. 1974. The Bermuda seamount: a reactivated section of an older ocean crust. EOS 55, 455. AZAMBRE, M. B. 1970. Les monchiquites et autres roches basiques intrusives accompagnant les sy6nites n6ph61iniques des Corbi+res. C. R. Acad. Sci. Paris, 271, 641-3. BACHINSKI, S. W. & SCOTT, R. B. 1980. Authors' reply. Geochim. cosmochim. Acta 44, 1389-92. & SIMPSON, E. L. 1984. Ti-phlogopites of the Shaw's Cove minette : a comparison with micas of other lamprophyres, potassic rocks, kimberlites, and mantle xenoliths. Am. Mineral. 69, 41-56. BAILEY, D. K. 1984. Kimberlite: 'The mantle sample' formed by ultrametasomatism. In: KORNPROBST, J. (ed.) Kimberlites I: Kimberlites and Related Rocks, pp. 323-34. Elsevier, Amsterdam. 1987. Mantle metasomatism--perspective and prospect. In : FITTON, J. G. & UPTON, B. G. J. (eds) Alkaline Igneous Rocks, Geol. Soc. Spec. Publ. 30, pp. 1-13. BAILEY,E. B., CLOUGH, C. T., WRIGHT, W. B., RICHLY, J. E. & WILSON, G. V. 1924. Tertiary and postTertiary geology of Mull, Loch Aline and Oban. Mem. geol. Surv. Scotland. BAKER, B.H. 1953. The alkaline igneous complex of Jombo Hill. Geol. Surv. Kenya Rep. 24, 32-48. BARNES, R. P., ROCK, N. M. S. & GASKARTH, J. W. 1986. Late Caledonian dyke-swarms of southern Scotland: new data for the Wigtown Peninsula. Geol. J. 21,101-25. BARRABLL. C. 1952. Les roches intrusives fi hornblende brune des Pyr6n6es. 19th Int. Geol. Cong. Fasc., 6, 9-22. BARTON, M. & HAMILTON,D. L. 1978. Water-saturated melting relations to 5 kb of three Leucite Hills lavas. Contrib. Mineral. Petrol. 66, 41-9. -& -1979. The melting relationships of a madupite from the Leucite Hills, Wyoming to 30 kb. Contrib. Mineral. Petrol. 69, 133-42. & 1982. Water-undersaturated melting experiments bearing upon the origin of potassiumrich magmas. Mineral. Mag. 45, 267-78. - & VAN BERGEN, M. J. 1981. Green clinopyroxenes and associated phases in a potassium-rich lava from the Leucite Hills, Wyoming. Contrib. Mineral. Petrol. 77, 101-14. BASTA,E. Z., EL-SHARKAWY,M. A. & MIKHAIL, M. A. 1985. Petrographical and petrochemical studies on some Egyptian lamprophyres. J. Coll. Sci. King Saud Univ. (Riyadh), 16(1), 137-64. BAXTER, A. N. & MITCHELL, J. G. 1984. Camptonitemonchiquite dyke swarms of Northern Scotland:
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NICHOLAS M. S. ROCK, Dept. Geology, University of Western Australia, Nedlands 6009, Western Australia.
Hawaiian alkaline volcanism David A. Clague SUMMA R Y:Hawaiian volcanoes erupt alkalic lavas during three of four eruptive stages. Stage 1 is characterized by submarine eruption of alkalic basalt and basanite followed by tholeiitic basalt which is exclusively erupted during stage 2. Stage 3 consists ofa post-calderacollapse capping of alkalic basalt and related differentiates. Stage 4 consists of post-erosional eruptions of alkalic basalt, basanite, nephelinite and melilitite. Eruption rates and volumes increase from stage 1 to stage 2 and then decline through stages 3 and 4. Stage 1 and 4 lavas contain deep-seated lherzolite xenoliths (even some garnet-bearing xenoliths in stage 4 lavas) whereas stage 2 lavas contain only very shallow cumulates, veins, dykes and sills; stage 3 lavas contain mainly dunite, wehrlite and various cumulates. The xenolith assemblages reflect the development of magma storage reservoirs which act as hydraulic filters in removing any xenoliths entrained below the magma storage reservoirs. The development of the magma storage system is a direct response to varying magma supply rates during the four eruptive stages. Trace-element, isotopic and rare-gas data all indicate that Hawaiian lavas are generated from heterogeneous sources consisting of at least three distinct compositions. Mixing of the sources or of partial melts derived from these sources is inevitable ; however, the characteristics of the mixing end-members are not well constrained. In a general way, early and late lavas are generated by small degrees of melting of sources enriched with less incompatible elements, whereas the voluminous shield tholeiites are generated by large degrees of melting of sources enriched with more incompatible elements.
Geological setting
Hawaiian eruption sequence
The Hawaiian Islands are the youngest part of a volcanic chain which extends nearly 6000 km across the north Pacific Ocean (Fig. I). This unique geological feature consists of more than 107 volcanoes with a combined volume of slightly greater than 1 million km 3 (Bargar & Jackson 1974). The chain is age progressive (Fig. 2) with volcanoes which are still active to the SE; the volcanoes at the north-western end are about 7580 Ma old (Dalrymple et al. 1981). Hawaiian volcanoes erupt lavas of distinct chemical compositions during different stages in their evolution and growth (Table 1). Three of these stages are characterized by eruption of alkalic to strongly-alkalic lavas. However, the lavas erupted during these alkalic stages probably represent less than 5~/ of the total volume of Hawaiian volcanoes. The remainder is composed of tholeiitic lavas that form the shield-shaped edifices for which the islands are well known. The distribution of the alkalic lavas of the Hawaiian Islands and Ridge and the Emperor Seamounts is reviewed in this paper, and their sequence, timing, volume and compositions, the distribution of the enclosed xenoliths and models for the generation of alkalic lavas in the mantle beneath Hawaii are discussed.
Hawaiian volcanoes evolve through four discrete eruptive stages starting with a pre-shield alkalic stage followed by a tholeiitic shield-building stage, a post-caldera or post-shield alkalic stage and a post-erosional or rejuvenated alkalic stage (Fig. 3). The three latest stages occur widely and are well studied and documented (Stearns 1940; Macdonald & Katsura 1964; Macdonald 1968), but the pre-shield alkalic stage, which includes the early phase of the submarine history of the volcano, has only been described recently from one volcano (Moore et al. 1982). Insight into the pre-shield stage has come from recent studies of Loihi Seamount, a small submarine volcano located about 30 km off the SE coast of Hawaii. The location, small size, seismic activity, fresh glassy lavas and active hydrothermal discharge all indicate that Loihi is an active volcano and the youngest in the linear HawaiianEmperor volcanic chain. Some of the older lavas recovered from Loihi Seamount are alkalic basalt and basanite, whereas most of the youngest lavas are tholeiitic and transitional basalt. This observation led Moore et al. (1982) to conclude,that Loihi Seamount, and perhaps all Hawaiian volcanoes, initially erupted alkalic lavas. In the shield-building stage tholeiitic basalt
Front: FITTON,J. G. & UPTON, B. G. J. (eds), 1987, Alkaline Igneous Rocks,
Geological Society Special Publication No. 30, pp. 227-252.
227
228
D.A. 1620 24 o
Clague 160 o
158 o I
I
500
"~176
1560 I
154 o I
',.',00~------'i
20 o
I
I
I
I
400
DR BEND
Se~Ua~;uU~t 30 30oa -
-~/~ ~Colahan Se. . . . . . ~~'~
o ~
~ w o y ,. Pea;::nd. . . . . .
v ~..~ ~
,~x(j_lq
F..... Frigo,e
Necker l ~ - - ~ 9
I
I
1700
1800
I 170 ~
.~Kouai~50oh Hawaiior _~-r~e~) ~ . , t o n , , . ~ . Deep
I 160 ~
FIo. i. Location map of the Hawaiian-Emperor volcanic chain in the north-central Pacific Ocean. The bathymetry of the principal Hawaiian Islands and Loihi Seamount is shown in the inset (after Chase et al. 1970). Note the striking linearity of the chain and the pronounced bend between the Hawaiian and Emperor sections of the chain. Contour interval 1 km.
flows construct the main volcanic edifice in a relatively short time, perhaps 1 Ma or less (Jackson et al. 1972). Most of the volume of an individual volcano (probably more than 95%) is formed from these voluminous eruptions. The later part of the shield-building stage may include
caldera collapse and the eruption of calderafilling tholeiitic, transitional or alkalic Iavas. In the post-caldera stage a relatively thin cap of alkalic basalt and associated differentiated lavas cover the main shield. These alkalic lavas account for less than 1~ of the total volume of
Hawaiian alkaline volcanism 0
K - A r or 4 ~
9
M o s t reliable K - A r and 4 ~
z~
Fossil a g e s
229
ages ages
0
60 90
0
(3 40 --
<
- - -
HAWAIIAN
RIDGE
~1~
9
-EMPEROR
SEAMOUNTS~
O~ 0
o
o 9
~o I 0
i
2000 DISTANCE
I 4000
F R O M LOIHI S E A M O U N T
6000 (km)
FIG. 2. Plot of the age of the volcanoes in the Hawaiian-Emperor volcanic chain as a function of distance along the chain. For discussion of individual ages plotted see Clague & Dalrymple (1987). Note that in general the volcanoes of the chain increase in age to the N and W away from the Loihi Seamount, Kilauea and Mauna Loa.
TABLE 1.
Hawaiian eruptive products
Eruptive stage
Rock types
Eruption rate
Volume
Post-erosional stage
Alkalic basalt Basanite Nephelinite Nepheline-melilitite
Very low
< 1%
Post-caldera stage
Alkalic basalt Transitional basalt Ankaramite Hawaiite Mugearite Benmoreite Trachyte
Low
--~1%
Shield-building stage
Tholeiitic basalt Tholeiitic picrite Alkalic basalt (?)*
High
95%-98%
Pre-shield stage
Basanite Alkalic basalt Transitional basalt Tholeiitic basalt (?)*
Low
~ 3%
* Wright & Helz (1987) suggest that the shield stage may include rare intercalated alkalic basalt and that the pre-shield stage includes tholeiite. I suspect that tholeiite and alkalic basalt occur intercalated during the transitions from pre-shield stages and from shield stage to post-caldera stages, but that during the main shield-building stage only tholeiitic lavas are erupted. From Clague & Dalrymple 1987.
230
D. A. Clague Post-erosional alkalic . P o s t- c a Id e r a a Ik a li c ~
Ocean
":'('!i::::'::i'ii"i':~'ii'i"i'i':':
ROCK TYPES
~ ~ ' ....~m.",-d:.~,.,-~:~.~.:-..;.:.? ~-.-,:, ~::z:::::;
AIkalic bas alt
ii![i~i i
Nephelinite Melilitite
f.:3:.,:~.:.'...: ?:.:. >.f,
Hawaiite
:i'i:~:':~:i~'::'i"i:~'~-'i:i'i~':~:i~'i~'i"i~~':::::':i'i'i'i
ESTIMATEDVOLUME
<<1%
Alkalic basal! 1%
TMug r aceharite yte Tholeiitic basalt Tholeiilic picrite
97-98%
~ Transitionalbasalt Alkalic basalt Basanite
1-2%
FIG. 3. Cross-sectionthrough a hypothetical Hawaiian volcano showing the age and volume relations of the four eruptive stages. The main rock types and their relative volumes are also shown. the volcano and are erupted soon after cessation of the shield-building stage. The final stage occurs after as much as several million years of volcanic quiescence and erosion. During this post-erosional stage, very small amounts (less than 1%) of SiOz-poor lavas may erupt from satellite vents. An individual volcano may become extinct before this eruptive cycle is complete, but the general sequence is typical of the well-studied Hawaiian volcanoes. The rock types known to occur on each of the volcanoes in the HawaiianEmperor volcanic chain and their inferred eruptive stages are summarized in Table 2. Evidence for the pre-shield alkalic stage exists only at Loihi Seamount. If this stage is present at all Hawaiian volcanoes, it is completely buried by later shieldbuilding tholeiitic lavas. The post-caldera alkalic stage lavas occur relatively late in the eruptive sequence. A number of the volcanoes have erupted predominantly mugearite whereas others have erupted predominantly hawaiite; these are called Kohala type and Haleakala type respectively (Macdonald & Katsura 1962). Farther W along the Hawaiian Ridge and in the Emperor Seamounts, lavas thought to have been erupted during the post-caldera alkalic stage are quite common. Post-erosional alkalic-stage lavas occur on Haleakala, Maui, E Molokai, the Koolau and Waianae volcanoes on Oahu, Kauai, Niihau, Kaula Island and perhaps Kahoolawe. To the W along the Hawaiian Ridge, lavas thought to be post-erosional are scarce but have been recovered from five seamounts and islets. Post-erosional lavas have not been recovered from the Emperor Seamounts.
The assignment of lavas to volcanic stages is difficult for samples recovered by drilling and dredging since the field relations are essentially unknown. We have used geochemical and mineralogical criteria developed from lavas of known volcanic stages in the principal islands (Clague et al. 1980; Frey & Clague 1983) to assign samples recovered from the Hawaiian Ridge and Emperor Seamounts. In general, lavas from the different stages are distinct rock types with the exception of alkalic basalt which can occur in any of the three alkalic stages. At Ojin Seamount and Suiko Seamount drilling has recovered alkalic lavas overlying tholeiitic lavas, as in the Hawaiian Islands. In both cases the overlying lavas are interpreted to be postcaldera-stage alkalic lavas (Kirkpatrick et al. 1980). These observations suggest that the sequence of stages observed in the Hawaiian Islands has occurred repeatedly throughout the nearly 80 Ma history of the Hawaiian-Emperor volcanic chain.
The pre-shield alkalic stage Loihi Seamount is located to the SE of Kilauea and Mauna Loa volcanoes and lies 30km offshore. The seamount rises to within 950 m of the sea surface. Two distinct rift zones extend N and S from a summit caldera 2.8 km wide and 3.7 km long (Malahoff et al. 1982). Moore et al. (1979) describe tholeiitic basalt dredged from Loihi Seamount in 1978 that is fresh and glassy. The young age of the lavas, coupled with seismicity (Klein 1982) and active hydrothermal discharge above the volcano (Malahoff et al. 1982; Horibe et al. 1983; Kim & Craig 1983),
Hawaiian alkaline volcanism TABLE 2.
23I
Rock types and inferred volcanic stages represented along the Hawaiian-Emperor chain Shield-stage 1
2
Kilauea Mauna Loa Hualalai Mauna Kea Kohala
x x x x x
x x x x
Haleakala Kahoolawe W Maui E Molokai Lanai
x x x x x
x x x x
W Molokai Koolau Waianae Kauai Niihau
x x x x x
x x x x x
Kaula Nihoa No. 19 No. 20 No. 21
x x
Necker La Perouse Pinnacles Gardner Pinnacles Brooks Bank St. Rogatein Bank Laysan Island Northampton Bank Pioneer Bank Pearl and Hermes Reef Ladd Bank
Post-erosional alkalic stage
Post-caldera alkalic stage 3
4
5
X
X
X
X
6
7
X
X
9
X
X
X
11
X
X
12
13
14
X
X
X
X
x X
10
X X
X
8
x? X
X
X
X
X
X X
X
X
X
X
X
X
X
X
X X X
X X
X X X
X
x x
x x x x x
x?
x
x?
x X
• •
X
Midway Island Nero Bank No. 57 No. 63 Colahan Seamount
x
Abbott Seamount Daikakuji Yuryaku Kimmei Koko (S)
x (T) •
Koko (N) Ojin Jingu Nintoku Suiko (No. 90)
x x
Suiko Meiji
x x
X
x X
X
X
X
x X*
x (T)
x*
x
X
X
X
X
X
x?
x
From Clague & Dalrymple 1987. 1, Tholeiite; 2, tholeiitic picrite; 3, rhyodacite; 4, alkalic basalt; 5, ankaramite; 6, hawaiite; 7, mugearite; 8, trachyte; 9, phonolite; 10, alkalic basalt; 11, basanite; 12, nephelinite; 13, nepheline-melilitite; 14, tephrite. (T) indicates transitional basalt. * Dredges from seamount no. 63 and Colahan Seamount recovered ankaramite, tephrite, and amphibole-bearing hawaiite that are probably all post-erosional stage lavas.
23 Z
D. A. Clague
indicate that Loihi Seamount is an active volcano and the youngest edifice in the H a w a i i a n Emperor volcanic chain. Moore et al. (1982) describe fresh glassy pillow lavas dredged from Loihi Seamount in 1981 that include basanite, alkalic basalt, basalt transitional between alkalic and tholeiitic basalt, and tholeiitic basalt. They also demonstrated, using thicknesses of palagonite, that the alkalic lavas are generally older than the tholeiitic lavas and suggested that Loihi Seamount, and presumably all Hawaiian volcanoes, have a pre-shield alkalic stage characterized by infrequent small-volume eruptions of alkalic lavas. Representative analyses of the Loihi lavas are given in Table 3. I selected a subset of the Loihi lavas for detailed analysis by a group of investigators for trace elements, radiogenic and stable isotopes and rare gases. Results from many of these studies are presented in a special volume edited by Craig (1983). The results of these detailed studies have broad implications for the generation of Hawaiian alkaline magmas. The broad range of rock types recovered from Loihi Seamount is shown in Fig. 4. Unlike most other Hawaiian volcanoes, the compositional range is continuous from tholeiite to alkalic basalt rather than plotting as discrete tholeiitic and alkalic clusters of analyses.
The lava types recovered have minor-element, trace-element and isotopic characteristics similar to those of sub-aerial Hawaiian lavas (Frey & Clague 1983; Lanphere 1983; Staudigel et al. 1984). These similarities confirm that Loihi Seamount is indeed a Hawaiian volcano. These same data also indicate that the lavas from Loihi Seamount are generated from heterogeneous mantle sources and that the lavas have equilibrated within the garnet stability field (Frey & Clague 1983). Trace-element ratio-ratio plots define linear arrays indicative of mixing but do not identify two end-members; this implies that more than two mixing components are required (Frey & Clague 1983). Radiogenic isotope studies indicate that three or more source compositions are involved in the generation of these lavas (Staudigel et al. 1984). In addition, there is no obvious correlation between radiogenic isotopic ratios of Pb, Nd or Sr and the bulk composition of the lavas or their eruptive ages (Lanphere 1983 ; Staudigel et al. 1984). Helium, argon and xenon isotopic data for glass from the Loihi lavas show that the Loihi samples are strongly enriched in 3He/4He compared with mid-ocean ridge basalt, which is interpreted as indicating a primitive u n d e g a s s e d mantle source for the volatile elements (All6gre et al. 1983 ; Kaneoka et al. 1983 ; Kurz et al. 1983;
TABLE 3. Major-element analyses (weight %) o f Loihi Seamount lavas Tholeiitic and transitional basalts 29-10
23-3
21-2
31-12
4 8 . 4 4 7 . 6 4 6 . 9 46.9 48.9 12.4 12.4 12.5 12.5 14.3 2.02 3.41 3.22 2.78 3.64 10.51 8.63 9.25 9.83 8.96 10.0 7.92 9 . 6 3 9.86 4.65 9.47 10.6 12.6 11.0 11.4 2.19 2.30 2.40 2.35 3.50 0 . 3 4 0 . 4 9 0.55 0.49 0.88 2.37 2.83 2.63 2.34 3.56 0.23 0.26 0.27 0.25 0.46 0.18 0.18 0.19 0.18 0.19 0.34 0.68 0.51 0.52 0.70 0.35 0.40 0 . 1 6 <0.01 0.14 <0.01 <0.01 <0.01 <0.01 <0.01
44.4 10.6 2.24 10.22 16.3 10.1 1.91 0.52 1.68 0.19 0.19 0.57 0.03 0.01
SiO2 A1203 Fe203 FeO MgO CaO Na20 K20 TiO2 P205 MnO H:O + H20CO2
48.7 12.3 2.06 9.57 10.7 10.6 1.96 0.32 2.22 0.20 0.17 0.36 0.01 <0.01
Total
99.2
99.9
Quartz -Hypersthene 22.2 Nepheline --
-17.2 --
1 8 - 8 25-4
99.7 -
99.2
-
8.7 .
-
-
9 9 . 4 99.3 -
6.7 .
24-7
Alkalic basalts
.
-
3.9 .
0.7 10.4
27-4
1 7 - 2 20-4
1 7 - 1 7 15-4
43.1 46.3 45.7 9.49 13.6 13.2 2.68 3.65 3.50 10.19 8.50 8.55 18.5 6.58 7.08 9.55 11.8 12.0 1.93 2.75 3.13 0 . 6 2 0.91 0.99 1.94 3.27 3.37 0.23 0 . 3 8 0.45 0 . 1 9 0 . 1 7 0.16 0 . 7 0 1 . 1 5 0.96 0 . 0 6 0 . 2 0 0.14 <0.01 <0.01 <0.01
42.4 44.7 10.8 14.6 3.99 4.29 9.73 7.75 ll.9 5.52 11.1 10.5 3.13 3.69 1.23 1.61 3.03 3.74 0.47 0.59 0.18 0.17 0.96 1.35 0.15 0.70 <0.01 0.01
99.0 . . 2.4
Basanite
9 9 . 2 99.3 . .
. . 4.7
. . 0.4
99.2
99. l
99.2
4.4
11.8
7.0
From Frey & Clague 1983. Major-element abundances determined by X-ray fluorescence (analysts: J. S. Wahlberg, J. Taggart and J. Baker, U.S. Geological Survey, Denver, CO). Analyses for FeO, H20 +, H20- and CO2 by classical techniques (analysts: H. Neiman and E. Engleman, U.S. Geological Survey, Denver, CO).
Hawaiian alkaline volcanism
233 /
Alkalic b a s a l t J
Thole,re 8
~
Basanite
~]
Transitional
ALKALIC ~
FIELD /
~6
J
%4 2
0
~ 34
I 38
t
I 42
--am
I 46
J
l 50
t
I 54
i
I 58
I
I 62
i 66
S i O 2 ( w t %)
FiG. 4. Alkali-silica diagram for lavas from Loihi Seamount. Fields enclose whole-rock analyses of tholeiitic, transitional and alkalic basalts and basanite. The Loihi lavas are a continuum of compositions from alkalic basalt to tholeiite. Analyses have been published for 12 samples (Frey & Clague 1983); the other 18 are unpublished data of D. A. Clague & J. G. Moore. Rison & Craig 1983). This mantle source has been isolated for much of Earth history and may represent the 'plume' component in Hawaiian lavas. Another lower 3He/4He component resides in olivine xenocrysts and dunite xenoliths and apparently contaminates the 3He/4He of the alkalic lavas. The volcanoes on Hawaii show an inverse correlation of 3He/~He with the volume of the volcano, suggesting that the melts producing smaller edifices are generated from bulk sources having larger percentages of the primitive undegassed mantle component. The Loihi lavas are notably vesicular considering their depth of eruption (1-2.6 kin). Moore & Clague (1982) note that, during vesicle formation, S, CO2 and perhaps up to 0.2 wt.% H20 are lost from the most vesicular alkalic melts. The alkalic lavas have higher FezOJFeO ratios than the tholeiitic lavas, indicating eruption under higher oxygen fugacity. The higher oxygen fugacity could result from loss of hydrogen from the alkalic lavas during vesiculation. Water behaves as an incompatible element in the Loihi lavas and has a positive correlation with K20 such that H20 + --0.9K20. The high gas content of these early alkalic lavas provides a mechanism for initiation of volcanism on the sea-floor and for buoyant transport of magma through the thick relatively cool lithosphere (Moore & Clague 1981). A few of the most alkalic flows from Loihi Seamount contain ultramafic xenoliths. Most of these xenoliths are dunite, but rare lherzolite and olivine cumulates are also present. The xenoliths contain abundant CO2-rich inclusions that were
trapped at depths of 10-18 km. Other decrepitated inclusions indicate that at least some of the xenoliths originated even deeper (Roedder 1983). The mineral chemistry of these dunite xenoliths (Clague et al. 1982b) is similar to that of dunite from the post-erosional Honolulu Volcanics on Oahu which are thought to be recrystallized cumulates related to Hawaiian lavas (Sen 1983). The lherzolite xenoliths occur in several of the older and most primitive alkalic lavas. The lava flows are apparently of small volume since 65 chemically distinct flows were recovered in 18 dredges. Thus the early stages of the growth of Hawaiian volcanoes are characterized by small-volume eruptions of basalt with a large compositional range including alkalic and tholeiitic lavas. Loihi Seamount has already evolved to the point where intrusion of magma occurs frequently enough to sustain a shallow-level magma chamber, believed to underlie the summit caldera, in which many of the younger lavas have differentiated and perhaps mixed.
Tholeiitic shield-building stage The lavas erupted during this stage are tholeiitic basalt and tholeiitic picrite. Occasionally, a pocket oftholeiitic magma is isolated long enough to differentiate to basaltic andesite, andesite (icelandite) and rhyodacite (Sinton 1981). The tholeiitic lavas from each volcano are geochemically distinct (Wright 1971 ; Leeman et al. 1977; Basaltic Volcanism Study Group 1981; Budahn & Schmidt 1984). These sub-alkalic lavas are discussed in detail in these references and they
234
D. A . C l a g u e Likewise, post-caldera alkalic-stage lavas have been recovered from eight of eleven seamounts sampled in the Emperor Seamounts. Analyses of post-caldera alkalic-stage lavas from the principal islands are presented by Macdonald (1968) and Macdonald & Katsura (1964). Several of the volcanoes have been studied in some detail, including Hualalai volcano (Clague et al. 1980), Kohala (Feigenson et al. 1983; Lanphere & Frey 1987), Haleakala (Chen & Frey 1983, 1985) and E Molokai (Beeson 1976; Clague & Beeson 1980; Clague et al. 1983). Additional detailed studies of Mauna Kea (West & Garcia 1982; Wise 1982), and W Molakai (Clague, unpublished data) are in progress. Representative analyses are given in Table 4 and plotted on an alkali-silica diagram in Fig. 5. Samples recovered from the Hawaiian Ridge and Leeward Islands include hawaiite and mugearite cobbles from Laysan Island (Dalrymple et al. 1981) and Midway Atoll (Dalrymple et al. 1977) and miscellaneous samples of alkalic basalt, hawaiite, mugearite and phonolite from other volcanoes (Clague 1974; Garcia et al. in press). Samples from the Emperor Seamounts are alkalic basalt, hawaiite or mugearite, except for a large collection of cobbles from Koko Seamount that includes alkalic basalt, hawaiite, mugearite, benmoreite, sanidine trachyte, anorthoclase trachyte and phonolite (Clague & Greenslate 1972; Clague 1974). These diverse lavas can be related to one another by crystal fractionation of augite, olivine, plagioclase, ilmenite, apatite, titanomagnetite and sanidine.
will not be discussed further here. Throughout this eruptive stage, intrusion of magma occurs frequently enough to sustain a shallow-level magma chamber. It is also probable that a deeper magma storage site develops near the base of the oceanic crust. Post-caldera alkalic stage Lavas from the post-caldera alkalic stage consist mainly of alkalic basalt, hawaiite and mugearite with smaller amounts of ankaramite, trachyte and phonolite. These flows form a thin, usually discontinuous, cap on the tholeiitic shield. Eruption of these alkalic lavas commonly follows collapse of the summit caldera and can occur as part of of the caldera fill or on tholeiitic calderafilling lavas. The eruptions are commonly explosive and form cinder cones that dot the summit regions of the volcanoes. The vents are commonly concentrated in and around the caldera and along the rift zones established during the shieldbuilding stage. The hiatus between the tholeiitic shield stage and the post-caldera alkalic stage is generally of short duration, if it exists at all. The post-caldera alkalic stage has yet to develop on Loihi, Kilauea and Mauna Loa volcanoes. It is present on all the other volcanoes in the Hawaiian Islands except Lanai and Koolau on Oahu (Table 2) although the volumes present on Kauai, Niihau, Kahoolawe and W Molokai are small. In the Leeward Islands and the Hawaiian Ridge, lavas thought to have erupted during the post-caldera alkalic stage are common.
TABLE 4. Major-element compositions (weight %) of representative Hawaiian post-caldera alkalic-stage lavas Volcano Volcanic unit Rock type Sample no.
Mauna Kea
Hualalai Hualalai AOB
Waawaa T
65HU-101 HAW-152
Hamakua AOB
Kohala
Haleakala
Hawi
Kula
Laupahoehoe A
H
C 773 C-673
C-793
AOB
H
M
B
KHll 4 MG9a4MG4b4KHI 4
SiO2 A120 3 Fe20 3 FeO MgO CaO Na_,O K,O TiO, P20 s MnO H20 + H20CO 2
46.62 14.72 2.00 10.71 9.10 10.04 2.82 0.91 2.37 0.27 0.19 0.00 0.04 0.01
62.99 18.04 4.42 -0.32 0.71 7.19 4.87 0.45 0.13 0.29 1.31 -0.11
46.65 14.72 3.63 8.98 7.96 10.14 2.53 0.80 3.22 0.33 0.18 0.63 0.51
44.54 13.52 4.47 7.76 11.92 10.71 1.69 0.59 2.60 0.32 0.18 0.52 0.77
49.79 18.37 6.03 4.82 4.01 6.68 4.58 2.08 2.51 0.56 0.22 0.32 0.24
47.99 15.97
Total
99.88
100.83
100.28
99.59
100.21
100.76
12.31 6.49 10.70 2.80 0.80 3.08 0.42 0.20 . .
51.13 13.61 -13.11 4.45 8.19 3.02 1.10 3.50 0.66 0.26 . . . .
53.18 18.15
99.03
99.31
9.66 2.89 5.40 4.04 2.11 2.14 1.48 0.26 . .
AOB
H
M
C -140s C -1435 C-1415
59.04 18.37 -6.26 1.73 3.39 6.07 2.71 1.39 0.59 0.23
44.72 13.86 5.69 9.82 5.07 10.89 3.07 1.09 3.96 0.61 0.23 0.40 0.33
45.66 17.01 4.96 7.73 4.60 8.09 5.01 1.55 3.41 0.75 0.20 0.37 0.40
51.90 17.10 3.38 6.64 2.26 5.70 6.65 2.72 2.11 0.84 0.17 0.34 0.29
99.78
99.74
99.74
100.10
Hawaiian alkaline volcanism
235
TABLE4--Continued Volcano
W. Maui
E. Molokai
W. Molokai
Niihau
Volcanic unit
Honolua
E. Molokai
W. Molokai
Paniau
H
AOB
Rock type
Sample no.
H
M
B
C-1534 C-1554 C-1514
SiO2 Al:O3 FezO3 FeO MgO CaO Na20 KzO TiO2 P205 MnO H20 + H20CO2
47.74 13.35 4.26 10.35 5.09 9.90 2.78 0.61 3.69 0.59 0.19 0.56 0.62 .
55.26 18,16 4.54 4.04 1.97 3,81 5.58 2.45 1,51 1.06 0.23 0,66 0.77 .
56.08 18.75 6.59 1.02 0.95 2.82 5.57 2.69 1.40 0.69 0.15 1.45 1.39 .
Total
99.73
100.04
99.55
Volcano
Midway Atoll
Rock type
AOB
H
M
H
M
KLPA-286 KLPA-296 KLPA-336
.
45.64 14.04 4.09 10.16 8.11 9.17 2.96 0.98 3.30 0.59 0.20 0.36 0.26 .
48.15 17.20 4.52 7.20 3.78 9.36 3.66 1.12 3.40 0.62 0.19 0.20 0.28 .
51,40 17.67 5.11 4.95 2,97 6,04 5.08 2.00 2,34 1.09 0.21 0.38 0.41
99.97
99.72
99.69
AOB
H
M
B
SiO2 A1203 F%O 3 FeO MgO CaO Na20 K20 TiOz P205 MnO H2 O+ H20CO2
48.60 15.26 8.01 2.84 5.20 7.94 3.54 1.68 2.69 0.39 0.35 2.17 1.26 --
51.04 16.44 8.28 1.87 2.72 5.25 4.46 2.64 2.40 0.85 0.27 1.83 0.93 --
44.4 15.2 -11.9 5.3 8.5 3.2 1.7 4.3 1.0 0.13 ----
45.6 15.7 7.2 6.3 3.5 6.7 4.4 1.8 2.7 2.5 0.17 1.49 0.64 0.01
47.4 16.1 9.5 3.0 2.4 6.8 4.6 2.4 2.5 2.1 0.14 0.91 0.89 0.05
52.7 17.7 3.2 6.3 1.2 2.8 5.4 3.5 0.91 1.30 0.03 1.50 1.30 0.08
Total
99.93
98.98
95.7
98.9
98.9
97.9
M
7ON11-237
46.4 15.5 7.35 6.7 3.81 6,73 4.10 1;31 4.15 0.79 0.t5 1.2 0.62 <0.01
47.46 15.22 4.55 8.25 6.05 9.97 3.08 0.48 3.31 0.40 0.15 0.54 0.55 0.01
46.76 14.81 -11.81 5.27 9.86 3.68 1.29 3.88 0.62 0.15 ---
50.38 17.16 -8.47 2.97 6.13 4.37 2.61 2.11 0.70 0.13 ---
--
--
100.02
98.15
95.03
98,8
D1-2 s D I - l l s
Ojin Nintoku Suiko Seamount Seamount Seamount T
11899 1 1 8 6 9
H
PH-37
KokoSeamount
Sampleno.
Laysan Is.
T
P
43-821~ 43 721~ 43-4 l~ 43-67 l~ 43 86 l~ 43-781~ 43-701~ 59.8 18.0 4.8 2.2 0.5 1.9 5.8 5.6 0.40 0.16 0.16 0.76 0.44 0.10 100.7
H
AOB
AOB
Avg 11
Avg 11
Avg 11
46.3 16.5 6.3 6.2 5.1 8.6 3.6 1.2 2.7 0.50 0.15 1.5 1.3 0.10
45.5 14.3 6.2 7.8 5.9 9.3 3.4 0.87 3.0 0.39 0.19 1.1 1.0 0.05
63.9 17.5 2.8 0.4 0.1 0.4 7.l 5.5 0.02 0.04 0.09 0.44 0.30 0.08
59.2 19.2 2.8 0.3 0.1 0.4 8.6 4.9 0.02 0.05 0.18 2.47 0.27 0.17
47.8 15.9 8.0 4.1 3.8 6.6 4.3 1.7 2.9 1.2 0.10 1.2 2.1 0.06
98.7
98.7
99.8
100.1
100.0
Rock types : AOB, alkalic olivine basalt; A, ankaramite; H, hawaiite; M, mugearite; B, benmoreite; T, trachyte; P, phonolite. Sources: 1, Clague et al. 1980; 2, Basaltic Volcanism Study Project 1981; 3, Macdonald 1968; 4, Feigenson e t al. 1983; 5, Macdonald & Powers 1968 ; 6, Beeson 1976; 7, Clague, unpublished data; 8, Dalrymple e t al. 1977; 9, Dalrymple et al. 1981 ; 10, Clague 1974; 11, Kirkpatrick et al. 1980.
The compositions of many of the lavas erupted during the post-caldera alkalic stage reflect differentiation due to crystal fractionation. The exact site of magma storage and differentiation is less clear, although evidence from Hualalai Volcano in particular suggests that fractionation leading to hawaiite occurs in a magma storage zone at greater than 15 km depth rather than in a shallow sub-caldera magma chamber. The key evidence supporting this conclusion is that fractionated lavas contain recrystallized dunite, and
wehrlite xenoliths contain C O 2 vapour inclusions trapped at a depth of about 15 km or greater (Roedder 1965; Clague et al. 1980; Kirby & Green 1980). The presence of these xenoliths in fractionated magmas implies that the parental magmas fractionated at depths greater than 15 km prior to incorporating the xenoliths during migration to the surface. At shallower levels these fractionated lavas also incorporate cumulate gabbroic xenoliths related to oceanic crust, the earlier tholeiitic shield-building stage and earlier
236
D. A. Clague p
,
,
,
!
,
i
,
i
,
i
,
,
,
Kohala Hualalai Mauna Kea
I0
Alkalic field
,
/
o
+ o z
6 oleiitic field
4
-
Tholeiitic field L~
Haleakala
./
0 (a)
J ~
,
i
i
I
I
L
I
I
i
I
L
(b)
14 ~]
fl
Koko
M i d w a y atoll [5555] Laysan I s l a n d
/ ~
Suiko
3
oN
l/
(
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--
2 /
I0 [Z]
Alkalic field 8
+ o
6
z
4
5'
leiitic field
/
A///~
~
.~ ~"
{r
i ,-~
54
38
42 ' 416
Alkalic field
5~0
i
54
/
~
West M o l o k a i
~
Niihau
5~8
i
6~2
Q_b~q:~
i
Si02 (wt.%)
~
~ ~ T h o l e i
East Molokai
E'~ Waianae
~ /
0
Nintoku
i
66
34
I ~
38
I
42
itic field
J
4'6
i
I
5o
i
5'4
I
I
58
i
~2
L
ee
Si02 (wt.%) (d)
FIG. 5. Alkali-silica diagrams for lavas from the post-caldera alkalic stage: (a) lavas from Hualalai, Mauna Kea and Kohala volcanoes on the island of Hawaii (data from Macdonald & Katsura (1964), Macdonald (1968), Clague et al. (1980), Feigenson et al. (1983) and Lanphere & Frey (1987)); (b) lavas from Haleakala and W Maui (data from Macdonald & K atsura (1964), Macdonald (1968) and Macdonald & Powers (1968)); (c) lavas from E and W Molokai, Waianae and Niihau (data for Waianae from Macdonald (1968), for E Molokai from Beeson (1976) and for W Molokai and Niihau from Clague (unpublished data)); (d) lavas from Midway Atoll (Dalrymple et al. 1977), Laysan Island (Dalrymple et al. 1981), Koko Seamount (Clague 1974) and Ojin, Nintoku and Suiko Seamounts (Kirkpatrick et al. 1980). alkalic magmas of the post-caldera alkalic stage (Clague, unpublished data). The eruptions are generally less voluminous and less frequent than eruptions of the tholeiitic shield-building stage, although individual eruptions may be comparable in volume with the largest tholeiitic-shield-stage eruptions (Macdonald & Abbott 1970). In general the post-caldera alkalic-stage lavas have less radiogenic Sr and more radiogenic Nd isotopic ratios than the underlying tholeiites (Clague 1982; C h e n & Frey 1983; Clague et al. 1983; Lanphere & Frey 1987) but have greater enrichment of incompatible trace elements and higher ratios of light rare-earth elements to heavy rare-earth elements. The enrichment of incom-
patible trace elements leads to post-caldera alkalic-stage lavas with higher Rb/Sr ratios and lower SVSr/86Sr ratios than the underlying tholeiites. These inverse correlations between isotopic ratios and their respective parent-todaughter abundance ratios can be explained by mixing of melts derived from enriched and depleted sources (Chen & Frey 1983, 1985) and will be discussed in a later section. The post-caldera alkalic-stage lavas contain a variety of xenoliths including dunite, wehrlite, various cumulate rocks and fragments of dykes, sills and veins. In general, the percentages of dunite and wehrlite decrease in the more differentiated lavas whereas the percentages of dykes, sills and cumulates all increase. In addition, the
Hawaiian
alkaline volcanism
cumulate xenoliths in the more differentiated lavas are dominantly pyroxene cumulates whereas those in alkalic basalt are dominantly olivine cumulates (Jackson 1968). Rare-gas compositions of dunite are similar to those in midocean ridge basalt and suggest that these dunite xenoliths may be related to the formation of oceanic crustal rocks (Kaneoka & Takaoka 1978, 1980; Kurz et al. 1983). The xenolith assemblages are notable for their lack of lherzolite or any type of garnet-bearing peridotite. Post-erosional alkalic stage Following the post-caldera alkalic-stage Hawaiian volcanoes are dormant for up to 2.5 Ma during which time deep erosional valleys are cut in the tholeiitic shield and overlying post-caldera alkalic cap. On many of the volcanoes this dormant period is ended by eruption of very small volumes of alkalic to strongly-alkalic lavas, generally from vents unrelated to the pre-existing rift system. The youngest volcano to have posterosional alkalic-stage eruptions is Haleakala. Further W, Kahoolawe, W Maui, E Molokai, Koolau and Waianae on Oahu, Kauai, Niihau and Kaula Island have post-erosional lavas, although only Haleakala, Koolau, Kauai and Niihau had more than a few small post-erosional eruptions. Lavas erupted during this stage are dominantly alkalic basalt and basanite, but nephelinite and nepheline melilitite are common among the posterosional Honolulu Volcanics on the Koolau shield on Oahu, and the Koloa Volcanics on Kauai (Winchell 1947; Macdonald 1968; Jackson & Wright 1970; Clague & Frey 1982; Clague et al. 1982; Chen& Frey 1983). Many of these lavas are primitive as shown by their high MgO and Ni contents and the presence of mantle xenoliths. Representative analyses are given in Table 5 and plotted on an alkali-silica diagram in Fig. 6. The erosional period between the post-caldera and post-erosional alkalic stages is not of random duration (Fig. 7) but increases systematically from less than 0.4 Ma at Haleakala to about 2.5 Ma at Niihau (Clague et al. 1982a). Jackson & Wright (1970) proposed that the Honolulu Volcanics might have been generated by uplift as Oahu passed over the Hawaiian Arch. They argued that the Hawaiian Arch, an isostatic response to volcanic loading of the oceanic crust, followed the active volcanic centre by several hundred kilometres and several million years. Clague & Dalrymple (1987) showed that the rate of volcanic migration along the chain is increasing and that the post-erosional volcanic stages follow the formation of the shields not by a constant time interval but by a constant distance
237
of 190_ 30 km. It seems likely that the energy source for post-erosional magma generation is the rapid change from subsidence to uplift as the volcanoes override the Hawaiian Arch. Post-erosional lavas are quite rare from the seamounts to the W of the principal Hawaiian Islands (Table 2), possibly reflecting the lack of large volcanic edifices capable of flexing the lithosphere sufficiently to provide the energy source for magma generation but also reflecting the small chance of recovering these lavas which cover only small areas of the principal islands. The lack of post-erosional lavas from the Emperor Seamounts probably reflects the wide spacing of volcanoes in the Emperor Seamounts and the young thin oceanic crust on which these volcanoes were built. By the time the next younger volcano formed, the previously constructed volcano was already beyond the zone of flexure. The young oceanic crust beneath the Emperor Seamounts results in a smaller distance between the load and the surrounding arch because the lithosphere is less rigid. Post-erosionallavas from the Honolulu Volcanics, the Koloa Volcanics and Kaula Island contain abundant xenoliths of spinel lherzolite and dunite (Jackson 1968; Garcia et al., in press). In addition, several vents in the Honolulu Volcanics, notably Salt Lake Crater, and the tuff at Kaula Island, contain garnet-bearing peridotite (Jackson & Wright 1970; Garcia et all, in press). Xenoliths are rare or absent in post-erosional lavas from Niihau, E Molokai, W Maui and Haleakala. The garnet-bearing peridotite xenoliths originate at depths greater than 50 km and the lherzolite originates at depths of the order of 30-50km (Sen 1983). Clague & Frey (1982) conclude that the varieties of xenoliths in the Honolulu Volcanics lavas are all unrelated to the magmas that brought them to the surface. Frey (1980, 1984) and Frey & Roden (1987) suggest that the spinel lherzolite xenoliths represent depleted oceanic lithosphere modified by processes related to Hawaiian magmatism. The presence of relatively deep-seated xenoliths in some of the post-erosional lavas suggests that (1) the host magmas continuously ascend more rapidly than the xenoliths sink and (2) the host magmas cannot therefore slow or stop in any magma storage reservoirs between the depth at which they incorporate the xenoliths and the surface. In several detailed studies of post-erosional lavas, particularly the Honolulu Volcanics and the Hana Volcanics on Haleakala, it is concluded that the mantle source for these lavas is enriched in incompatible elements and that the enrichment event occurred relatively recently (Clague & Frey
238
Clague
D. A.
TABLE 5. Major-element composition (weight %) o f representative Hawaiian post-erosional alkalic-stage lavas Volcano Volcanic unit
Haleakala Hana
W Maui Lahaina
E Molokai Kalaupapa
Niihau Kiekie
Ladd Bank
Rock type
N
B
B
AOB
B
AOB
AOB
N
Sample no.
C1391
C1291
C-1152
C-1302
MoE23
74KAL-14
C-1682
H1G-23-16
SiO2 A1203 FezO3 FeO MgO CaO Na20 K20 TiO2 P205 MnO H20 § H20CO2
41.55 14.13 3.99 11.36 6.04 11.79 3.84 1.67 4.10 0.63 0.19 0.05 0.23 .
42.53 12.43 2.68 11.23 12.17 11.80 2.35 0.81 2.92 0.55 0.16 0.03 0.17
44.07 13.68 4.01 8.35 11.74 10.22 3.45 1.35 2.16 0.40 0.20 0.20 0.38 .
44.02 12.41 2.14 10.99 13.40 10.95 2.39 0.62 2.05 0.45 0.20 0.17 0.45
43.02 11.82 5.26 9.19 13.76 10.37 2.74 0.74 2.17 0.33 0.16 0.88 0.21
45.8 13.7 3.68 9.11 10.4 10.9 2.29 0.27 1.59 0.20 0.19 0.57 0.61 0.19
44.67 15.17 1.98 10.39 11.65 11.10 2.27 0.27 1.01 0.16 0.18 1.30 0.43 --
39.7 13.2 -13.45 9.20 12.16 2.61 1.13 3.73 0.59 0.15 ----
Total
99.57
100.24
100.65
99.50
100.58
96.0
.
. 99.83
Volcano Volcanic unit Rock type Sample no.
100.21
.
Koolau Honolulu NM
N
65MOLL-25 68BP-25
Kauai Koloa B
AOB
65PAL-15 65KAL-15
NM
N
B
AOB
C-1892
C-186 z
C-1882
C-1932
44.34 10.37 4.14 9.74 11.09 12.67 2.65 0.96 2.38 0.97 0.17 0.16 0.39 .
45.42 12.39 3.22 9.04 11.54 11.72 2.56 0.80 1.98 0.73 0.16 0.48 0.35
100.03
100.39
SiO2 A1203 Fe203 FeO MgO CaO NaEO K:O TiO2 P205 MnO H20 § H20CO2
35.52 11.01 9.82 7.13 11.46 12.17 4.84 1.72 2.78 1.05 0.25 0.58 0.63 0.26
40.15 12.01 5.80 7.20 13.27 12.73 4.13 1.06 1.90 0.98 0.23 0.14 0.14 0.03
42.27 12.11 4.25 9.17 12.96 10.57 3.62 0.87 2.42 0.69 0.21 0.33 0.31 0.01
45.14 13.37 3.85 9.00 11.46 10.39 2.94 0.65 1.97 0.35 0.20 0.19 0.15 0.05
39.91 9.13 5.29 8.16 15.88 11.97 3.16 1.53 2.76 0.81 0.18 0.54 0.63 .
Total
99.33
99.95
99.91
99.86
99.95
41.86 11.04 4.42 9.51 12.00 11.86 3.06 1.20 3.24 0.50 0.17 0.52 0.59 . . 99.97
Rock types: AOB, alkalic olivine basalt; B, basanite; N, nephelinite; NM, nepheline-melilitite. Sources: 1, Macdonald & Powers 1968; 2, Macdonald 1968; 3, Naughton et al. 1980; 4, Clague et al. 1982a; 5, Clague & Frey 1982; 6, Clague 1974. 1982; C h e n & F r e y 1983; R o d e n et al. 1984). N u m e r o u s studies h a v e s h o w n that the posterosional lavas h a v e the least r a d i o g e n i c isotopic ratios o f Sr a n d the m o s t r a d i o g e n i c isotopes o f N d , but h a v e h i g h a b u n d a n c e s o f i n c o m p a t i b l e trace e l e m e n t s ( L a n p h e r e & D a l r y m p l e 1980; Clague & F r e y 1982; C l a g u e et al. 1982a; C h e n & Frey 1983; Stille et al. 1983; U n r u h et al. 1983; F e i g e n s o n 1984; H e g n e r et al. 1984; R o d e n et al. 1984). T h i s is the s a m e inverse r e l a t i o n s h i p b e t w e e n isotopic ratios a n d their r e s p e c t i v e
p a r e n t - t o - d a u g h t e r a b u n d a n c e ratios n o t e d for post-caldera alkalic lavas a n d tholeiitic lavas. T h e s e r e l a t i o n s h i p s require c o m p l e x p e t r o g e n e t i c m i x i n g m o d e l s t h a t will be d i s c u s s e d in a later section.
Experimental constraintsmdepth of origin C l a g u e & Frey (1982) r e v i e w e d the e x p e r i m e n t a l data bearing on the generation of Hawaiian
239
Hawaiian alkaline volcanism
8
"~
Co
+
4
Alkalic field
Alkalic field
/f1/~
/
oa Volconics
//
/f/~'~J
C)
~
Honolulu V o l c a n i c s Hana V o l c a n i c s
E:Z] Lahaina V o l c a n i c s
2
Tholeiitic field E~Kalaupapa Basalt
0
~
54
58
I
42
i
I
46
i
I
50
i
I
54
i
I
58
i
62
/
\ ~ -~/"
/ ~//~/~f ', ~. J/ ~ ' ~ ' ~ ~ / / ] / ~ / ~ / Th:leiitic i
66
54
SiO 2 (wt.%)
,/"
58
I
42
L
EEB .--, '- J v
field
I
46
EEE Kauta Island
A Seamount63
i
50
Seamount 57 Ladd B a n k (Seamount 51) Seamount 2 0
I
54
i
I
i
58
I
62
i
66
Si02(wt.%)
(a)
(b)
FIG. 6. Alkali-silica diagrams for lavas of the post-erosional alkalic stage: (a) data for the Hana Volcanics on Haleakala from Macdonald & Powers (1968), the Lahaina Volcanics on W Maui from Macdonald (1968), the Kalaupapa Basalt on E Molokai from Clague et al. (1982a) and the Honolulu Volcanics on Koolau volcano, Oahu, from Clague & Frey (1982); (b) data for the Koloa Volcanics on Kauai from Macdonald (1968), Feigenson (1984), and D. A. Clague (unpublished data), the Kiekie Basalt on Niihau from Macdonald (1968) and D. A. Clague, G. B. Dalrymple, M. H. Beeson & E. D. Jackson (unpublished data), from Kaula Island from Garcia et al. (in press) and from seamounts 20, 57 and 63 and Ladd Bank from Clague (1974). Note that few of the posterosional lavas are differentiated. Lavas from seamount 63 display a differentiation trend similar to post-calderastage lavas; however, they contain amphibole and clinopyroxene compositionally similar to that in post-erosional stage lavas from the Hawaiian Islands. alkalic basalt to nepheline melilitite. Brey (1978) and Yoder (1979) give more detailed reviews and discussion of the experimental data. There is a general consensus that an increasing C 0 2 / H 2 0 ratio leads to increasingly silica-undersaturated
6
5
3 <
2
---
v
T
T
A
[ i
v
-~ o
= <_ ~
o 0 z~
<
o0
/ ~ ~ / o /~ o ~ .N.~ e,i
<' ~
g
4
-
~
< ~- < ,..' ~o ,,,
< "'
1.1.1~ T / -
T
/ ~
/o.,,~247
-
" /.~'~,,.~-/I T/R~.~ - / I
I I
Y ~
c~ / ~/
P II ',
//
:~]~ [/~ ~ "~ <0.4-,_ .~" I P? ? I ? I 100 200 aoo 400
iI
10
I P
-~ ~
zA
I
500
DISTANCE TO KILAUEA (km)
Fro. 7. Age of tholeiitic (T), post-calderaalkalic (A) and post-erosional (P) stages for the Hawaiian Islands that have both age data and post-erosional lavas plotted as a function of distance to Kilauea. The duration of the quiescent period decreases systematically from 2.5 Ma on Niihau to less than 0.4 Ma on Haleakala. (From Clague et al. (1982a) where the data sources are cited.)
I
600
melts, largely because of the expansion of the primary-phase field of orthopyroxene, an effect also caused by increased pressure. However, no consensus exists on the exact compositions produced from the peridotite-CO2-H20 system as the degree of melting increases (Eggler 1978; Wyllie 1978, 1979). The transition from hypersthene-normative to nepheline-normative melts is poorly constrained, but Wyllie (1979) suggests that the transition to larnite-normative liquids occurs near 2.5-3.5 GPa, depending on the CO2/ H 2 0 ratio. Since all the alkalic lavas have rareearth patterns suggesting equilibration with residual garnet, it is concluded that the m i n i m u m depth of equilibration is about 47 km whereas the more silica-undersaturated post-erosional lavas are probably equilibrated at depths of 75110 km (Clague & Frey 1982; Sen 1983). It is interesting to note that these depth estimates are near the base of the oceanic lithosphere which is estimatedtobe90kmthickforthe90-100Ma old crust beneath Hawaii (Detrick & Crough 1978). There is also a general consensus that the alkalic lavas all originated deeper than the tholeiitic lavas which produce seismic swarms and tremors as deep as 60-70 km (e.g. Klein 1982).
Isotopic constraints on mantle sources
Abundant isotopic data exist for post-caldera and post-erosional alkalic lavas from Hawaii. However, most of the older data are of low precision
24o
D. A. Clague i
i
T
i
[~1 MORB 0.51:52
~
MORB Transitional
Haleakala
E[~] West Maul [Z] East M o l o k a i Mauna K e a Waianae Hualalai
Alkalic
0"5151
~
Tholeiite
0"5150 -g z 0"5129
\ \\\
~\\
O. 5128 0.5127
\\
,\
the\ \
Bulk ear
] (a)
0-7050
\ \~,
\\ \
\\
\ \\\
\
0.5126 0.7020
\\ \\\
0-7040 ~ x ~ 0 . 7 ' 0 5 0 87Sr/86Sr
i
0"71030
0"7( )20
0"7040
0"7050
87Sr/86Sr (b)
eL% \ 0-5129
x
\
E]Z] MORB Hana Volcanics [Z~ Lahaina Volconics [~Z Koloa Volcanics HonoluluVolcanics m KalaupapaBasalt
\\\ \\\\\\\
0-5128 0"5127
\,,
0.5126 0"7020
0"7050
(c)
0.7040 87Sr/86Sr
Bulk
earth "~
xI
0.7050
FIG. 8. 143Nd/l~4Nd versus 87Sr/S6Sr diagrams for (a) the pre-shield stage, (b) the post-caldera alkalic stage and (c) the post-erosional alkalic stage. In a general way the post-erosional lavas have sources with the least radiogenic Sr, the pre-shield and post-caldera lavas have sources with more radiogenic Sr and Hawaiian tholeiites have sources with the most radiogenic Sr. The reverse is true for Nd isotopes. However, considerable overlap occurs, indicating that the source compositions are not homogeneous for all Hawaiian tholeiites, nor for all Hawaiian post-erosional alkalic-stage lavas, nor for all Hawaiian post-caldera-stage lavas. See text for discussion. Data sources are Staudigel et al. (1984) for Loihi Seamount, Chen & Frey (1983, 1985) for the Hana and Kula Volcanics, Stille et al. (1983) for the Waianae Volcanics, Stille et al. (1983) and Roden et al. (1984) for the Honolulu Volcanics, Feigenson (1984) for the Koloa Volcanics, Frey et al. (1984) for Mauna Kea, D. A. Clague, F. A. Frey & A. Kennedy (unpublished data) for E Molokai Volcanics and the Kalaupapa Basalt, and Tatsumoto et al. (1987) for the Lahaina Volcanics and Honolua Basalt of W Maui. and were usually obtained on samples for which no other g e o c h e m i c a l data have been published. More recent studies include integrated isotopic analyses of Sr, N d and occasionally Pb and Hf. Available data for STSr/S%r and 143Nd/144Nd are plotted in Fig. 8 for alkalic lavas from all three eruptive states. In general, the pre-shield and post-caldera alkalic lavas are isotopically similar, but the post-
erosional lavas have less radiogenic Sr and more radiogenic N d isotopic ratios. H a w a i i a n tholeiitic lavas have isotopic values that overlap those of the post-caldera lavas but generally have more radiogenic Sr and less radiogenic N d ratios and occupy the portion of the H a w a i i a n field towards bulk Earth values (Staudigel et al. 1984). Some H a w a i i a n volcanoes show a systematic variation in, for example, Sr isotope ratios as a function of
Hawaiian
2 41
alkaline volcanism
stratigraphic position from more radiogenic in further propose that the Kilauea end-member the tholeiites to progressively less radiogenic in represents the deep-seated mantle plume and that post-caldera and post-erosional lavas (Chen & the remaining two end-members may reside Frey 1983; Clague et al. 1983; Feigenson 1984; within the lithosphere. White & Hofmann (1982) Lanphere & Frey, in press). The arrays of Sr and offer the alternative view that one of the three Nd isotopic data can be explained by the mixing mixing components is recycled crust, whereas the of two components, one similar to bulk Earth and remaining sources are the same depleted and the other more similar to the depleted source of undepleted sources advocated by Chen & Frey mid-ocean ridge basalt. This relatively simple (1983,1985). model is complicated by the fact that the lavas with the least radiogenic Sr and the most radiogenic Nd isotopic ratios also have the most incompatible-element enrichment (Chen & Frey Rare-gas constraints on mantle 1983). These relations require more complex sources mixing of those two sources and small-percentage partial melts derived from at least one of the Isotopic analyses of rare gases, particularly argon sources. C h e n & Frey (1985) present an array of and helium, indicate complex mixing relations between three or more sources which Kaneoka models of this type and geochemical-balance calculations which suggest that small-percentage and Takaoka (1980) infer to be an undegassed partial melts of a depleted source similar to that plume source, a mid-ocean ridge degassed source of mid-ocean ridge basalt mix with enriched and an atmospheric component (Fig. 9). They mantle to produce the sources for tholeiitic and infer that the characteristics of these sources are as follows: the plume source has 3He/4He = 5 x alkalic lavas from Haleakala. 10-5 (36 times atmospheric) and 4~ Pb isotopic data (Tatsumoto 1978; Sun 1980; Stille et al. 1983; Staudigel et al. 1984) indicate 36Ar=400; the ridge source has 3He/aHe = 1 • 10 -5 (seven times atmospheric) and 4~ that even the relatively complex models proposed 36Ar= 16 000; and the atmospheric source has by Chen & Frey (1983, 1985) are still inadequate = 295.5. in that at least three isotopically distinct sources 3He/4He = 1.384 x 10-6 and 4~ Most data from the Hawaiian Islands are for are required to explain the Pb data. Tatsumoto (1978) demonstrated that the volcanoes of the tholeiitic basalt or xenoliths enclosed in alkalic Loa trend (Mauna Loa and Hualalai) have more basalt. However, some data, mostly 3He/4He radiogenic Pb than the volcanoes of the Kea ratios, exist for alkalic lavas from the pre-shield trend (Mauna Kea, Kilauea and Kohala). In alkalic stage (Kaneoka et al. 1983; Kurz et al. addition, the Pb isotopic ratios of tholeiitic and 1983; Rison & Craig 1983), the post-caldera alkalic lavas of the post-caldera alkalic stage are similar for the Kohala and Mauna Kea volcanoes. Haleakala e Hualalai dunite Tatsumoto suggested that this last observation o Kilauea tholeiite • Salt Lake tuff (xenoliths) indicated 'a close relationship between the tholv Hualalai tholeiite 9 Loihi lavas IOO 9 Loihi dunite o Loihi olivine phenocrysts eiitic and alkalic basalt (post-caldera) sources of 10-4 7o 5o each volcano' (Tatsumoto 1978, p. 71). Tatsu4o moto also proposed that one of the mixing end5o T~T~ ,. n . n ~ . ' ~ \ members of the source was the oceanic lithos20 < r'r phere. Lanphere et al. (1980) suggested that the ~o oceanic lithosphere was the source of the lavas of ~z I0 -5 the Hawaiian-Emperor volcanic chain and that the underlying plume provided mainly the heat required to melt the overlying lithosphere. Recent integrated isotopic studies of lavas from Loihi Seamount (Staudigel et al. 1984) and 10-6 I0 3 I0 4 10 5 10 2 Oahu (Stille et al. 1983) suggest that the three 4OAr/36Ar source end-members are (1) a highly depleted source, (2) a slightly depleted source and (3) a plot showing source intermediate between (1) and (2) which FIG. 9.3He/aHe v e r s u s 4~ Hawaiian data from Kaneoka & Takaoka (1978, may be the deeper source. Staudigel et al. (1984) 1980), Kyser & Rison (1982), All~gre et al. (1983) and propose that these isotopic end-members are Kaneoka et al. (1983). The mixing end-members most closely represented by the post-erosional shown are the atmosphere (A), a degassed mid-ocean Honolulu Volcanics, the shield-stage Koolau ridge source (R) and an undegassed plume (P). Mixing lavas and the shield-stage Kilauea lavas. They models shown are from Kaneoka et al. (1983). i
i
L
I
i
242
D. A.
alkalic stage (Kaneoka and Takaoka 1980; Kurz et al. 1983) and the post-erosional stage (Kyser & Rison 1982; Rison and Craig 1983). The alkalic lavas from Loihi Seamount have 3He/4Heatm = 21-24, whereas post-erosional lavas from the Hana Volcanics on Haleakala and the Honolulu Volcanics on Oahu have (3He/4He)atm=8-14. Post-caldera alkalic lavas from the Kula Volcanics on Haleakala have (3He/4He)atm=27-32, whereas a single post-caldera alkalic lava from Mauna Kea has (3He/4He)atm = 7.9. It is difficult to reconcile these variations in 3He/4He, and similarly large variations in 4~ with simple mixing models. Hawaiian lavas with the highest 3He/4He isotopic ratios, and therefore the lavas believed to be derived from the most primitive undegassed sources, are post-caldera alkalic lavas from Haleakala and tholeiitic lavas from Loihi Seamount. Similarly, Hawaiian samples with 3He/4He ratios most similar to that of mid-ocean ridge basalt, and therefore believed to be derived from the most degassed sources, include post-caldera alkalic lavas from Mauna Kea, post-erosional alkalic lavas from Haleakala and Oahu, dunite xenoliths from post-caldera alkalic lavas on Hualalai (Kaneoka & Takaoka 1980; Kyser & Rison 1982; Kurz et al. 1983; Rison & Craig 1983) and a variety of xenoliths from post-erosional alkalic lavas from Oahu (Kaneoka & Takaoka 1980; Kyser & Rison 1982). Interpretation of rare-gas data is complicated by a number of factors (see discussion in Kaneoka (1981)) but the most striking observation is that helium isotopic ratios equilibrate rapidly between melt and xenocrysts (or xenoliths) owing to high rates of diffusion at magmatic temperatures, whereas the heavier-rare-gas isotopic ratios commonly display disequilibrium between coexisting quenched melt (glass) and enclosed xenocrysts or xenoliths (Kaneoka et al. 1983). This observation suggests that 3He/4He ratios determined for melts probably represent the cumulative history of the melt including its source characteristics and the characteristics of the mantle and crust through which it migrates. The rare-gas isotopic ratios observed in the lavas must therefore reflect the ratios and abundances of gases of both the source and the intervening oceanic lithosphere as well as the percentage partial melting of the source and the ascent velocity of the melt. Models that explain the rare-gas characteristics of Hawaiian lavas will probably be similar to the complex partial-melting and mixing presented by Chen & Frey (1983, 1985) to explain trace-element and Sr- and Nd-isotope characteristics of lavas from Haleakala. These models will be discussed in the following section. The available data, particularly for the Loihi Seamount, suggest that those
Clague samples with the most radiogenic Sr and the least radiogenic Nd isotopic ratios also have the highest 3He/4He ratios indicating the least degassed source. These data suggest that the hypothesized plume source both is the least degassed and has the most radiogenic Sr and the least radiogenic Nd of the multiple sources involved in the generation of Hawaiian magmas.
T r a c e - e l e m e n t c o n s t r a i n t s on m a g m a sources Voluminous trace-element data exist for Hawaiian lavas and numerous additional traceelement and isotopic studies are currently under way. As with the isotopic data, much of the earlier data were obtained for poorly located samples or for samples that had not been analysed for major elements. Few of these early studies attempted to evaluate coherent suites of samples from single volcanoes and fewer still evaluated known stratigraphic sequences of samples from single volcanoes. In the past 5 years integrated studies of stratigraphically controlled samples have shown that much of the variation between volcanoes in Hawaii also occurs within the lavas from single volcanoes. Detailed trace-element studies have been published for post-caldera alkalic lavas from E Molokai (Clague & Beeson 1980; Clague et al. 1983), Hualalai (Clague et al. 1980), Haleakala (Chen & Frey 1983, 1985), Kohala (Feigenson et al. 1983; Lanphere & Frey 1987) and Mauna Kea (Frey et al. 1984). All Hawaiian alkalic lavas analysed so far have rare-earth patterns that are enriched in the light rare earths, generally to a much greater degree, than Hawaiian tholeiites. Figure 10 shows chondrite-normalized rare-earth patterns for alkalic lavas from the pre-shield, post-caldera and posterosional alkalic stages. The rare-earth data present the classic problem of the generation of strongly enriched rare-earth patterns in alkalic magmas from worldwide locations (Kay & Gast 1972). In order to generate these patterns from chondritic source compositions, extremely small degrees of partial melting are required. More recent models propose an enriched source, thereby allowing the magmas to be generated by higher percentages of partial melting (see Clague & Frey (1982) for a discussion). Other trace-element data show that the lavas from the different eruptive stages cannot be generated from similar sources. In particular, the post-erosional lavas are derived from sources with lower Zr/P205 and K/Ba ratios and variable KzO/P20 5 ratios (see Figs 11 and 12). The sources
Hawaiian
alkaline
for the post-erosional lavas are apparently more depleted than those for the post-caldera or preshield lavas. A detailed study of the pre-shield alkalic (and tholeiitic) lavas from Loihi Seamont (Frey & Clague 1983) indicates that these lavas are geochemically similar to the post-caldera alkalic lavas erupted along the Hawaiian-Emperor volcanic chain. However, most of the alkalic samples from Loihi are relatively unfractionated whereas those from the post-caldera stage have commonly evolved to hawaiite, mugearite and, more rarely, trachyte. Trace-element ratios, such as K/Ba and Zr/P, isotopic ratios and rare-earth abundance patterns are similar for the pre-shield and postcaldera alkalic lavas (Frey & Clague 1983). On the basis of K/Ba ratios, the mantle residuum for the E Molokai alkalic basalts retains phlogopite whereas that for transitional lavas does not (Clague & Beeson 1980). Clague et al. (1980) proposed that the source of Hualalai alkalic basalts is garnet bearing and relatively enriched in the light rare-earth elements. Feigenson et al. (1983) proposed that the source for alkalic basalt from Kohala has fiat to slightly-convex-upward chondrite-normalized rare-earth patterns and argued that a metasomatic pre-enrichment of the source was not required, Their model is significantly different from that of Chen & Frey (1983, 1985) which explains the trace-element and isotopic characteristics of the shield tholeiites, the post-caldera alkalic lavas and the posterosional alkalic lavas from Haleakala. Clague & Frey (1982) presented a detailed trace-element analysis of the post-erosional Honolulu Volcanics on Oahu and concluded that lavas ranging from nepheline-melilititeto alkalic basalt were generated by 2~-11X partial melting of a homogeneous garnet (less than 10%) lherzolite source that was carbon bearing. The source had recently been enriched and had a chondritenormalized La/Yb ratio of 4.4. During the melting process phlogopite, amphibole and a Ti-rich phase (oxide?) remain in the residuum, but apatite is completely melted. This model emphasizes the bulk composition of the source rocks and the physical conditions of the melting process. Other studies of post-erosional lavas are less comprehensive but include evaluation of the Kalaupapa Basalt on E Molokai (Clague et al. 1982a), the Koloa Volcanics on Kauai (Feigenson 1984) and the Hana Volcanics on Haleakala (Chen & Frey 1983, 1985). The Feigenson (1984) model is similar to that of Chen & Frey (1983), except that the primitive-mantle-mixing endmember is replaced by a combination of ancient subducted crust and primitive mantle.
volcanism
243
The Chen& Frey (1983, 1985) model proposes mixing of partial melts derived from an enriched mantle composition and very small percentage partial melts derived from a depleted mid-ocean ridge basalt source (Fig. 13). The model is specifically designed to explain the inverse correlations of radiogenic isotopic ratios and their respective parent-to-daughter abundance ratios. In other words, at Haleakala the lavas with the most radiogenic 878r/86Sr ratios (the tholeiites) have the lowest Rb/Sr ratios whereas the lavas with the least radiogenic 87Sr/86Sr ratios (the post-erosional alkalic lavas) have the highest Rb/ Sr ratios. This same pattern has now been documented for E Molokai (Clague et al. 1983) Kohala (Lanphere & Frey 1987), Hualalai (Clague 1982), Koolau (Roden et al. 1984) and Kauai (Feigenson 1984). The pattern does not appear to be universal for Hawaiian volcanoes as tholeiitic and post-caldera alkalic lavas from both Mauna Kea (Frey et al. 1984) and Waianae (Stille et al. 1983) have identical 87Sr/S6Sr and 143Nd/a44Nd isotopic ratios. Chen & Frey (1983) propose a geodynamic model in which partial melts of the deeper enriched-mantle source extract very small percentage partial melts of their wall-rocks, which are assumed to be a depleted mid-ocean ridge basalt source, during their ascent. An alternative model (Fig. 14) is that small-percentage partial melts of the enriched source are mixed with depleted mid-ocean ridge source (Clague et al. 1983) (this model was also discussed and rejected by C h e n & Frey (1985)). This alternative is the equivalent of a metasomatic model wherein the small-percentage partial melts of the enriched source are added to the depleted lower lithosphere which is then partially melted to create the various Hawaiian lavas. Roden et al. (1984) combine the Chen & Frey (1983) model to generate the source compositions of the Clague & Frey (1982) model for generation of the Honolulu Volcanics post-erosional lavas. The isotopic homogeneity of the Honolulu Volcanics lavas requires that the source of the nepheline-melilitite to alkalic basalt must have been created by mixing prior to the partial melting that generated the lavas. This is a rigid constraint since the Chen & Frey (1983) model forms each magma batch by partial melting of enriched mantle followed by mixing with smallpercentage melts derived from the wall-rocks. It seems improbable that the mixing of variablepercentage partial melts of the enriched mantle and variable-percentage partial melts of the wallrock invariably occur in proportions which will yield isotopically homogeneous lavas. The compositions of the Honolulu Volcanics lavas thus
D. A. Clague
244
,~176 l 20
8C 6O
!Mauna
40
~ rr
20
~o
2
20 Haleakala
I0
7 15
,a&
4~ Sm~u&'g
40 -/m'gLu
East Molokai!
IO 7
30
(a)
20
70 50
La Ce "
Cbl) IOC
IOC
70 50
50
La Ce (b2)
_~0
T~n Yb Lu
Nd
s" ~u & -~
Tm Yb Lu
z45
Hawaiian alkaline volcanism 5OO
i
i
i
i
K6k6 seamount
300 5OO
Phonol ire
300
Sanidine 300 200 Anorthoclase trachyte
200
20
lugearite ant ;enmorei te
120 i O0 -k'------....~ Hawa iite
"10
30
5 10 5 L
La Ce (b3)
I
Nd
L
Sm Eu Gd Tb
Ho
I
~L~
tm Yb Lu
I
La Ce (c)
1
Nd
I
I
Sm Eu Gd'(b
14o
TmYb Lu
FIG. 10. Chondrite-normalized rare-earth diagrams. (a) Lavas from Loihi Seamount (Frey & Clague 1983; D. A. Clague, unpublished data). The Loihi lavas show increasing relative light-rare-earth enrichment from the tholeiites through the transitional lavas to the alkalic lavas. (b) Lavas from the post-caldera alkalic stage from Mauna Kea (Frey & Roden 1987), Kohala (Feigenson et al. 1983; Lanphere & Frey 1987), Hualalai (Clague e t al. 1980; Basaltic Volcanism Study Project 1981), the K ula Formation of Haleakala (Chen & Frey 1983), Niihau (D. A. Clague, G. B. Dalrymple, M. H. Beeson & E. D. Jackson, unpublished data), Midway Island (F. A. Frey & D. A. Clague, unpublished data), Koko Seamount (D. A. Clague, unpublished data) and Ojin, Nintoku and Suiko Seamount (Bence et al. 1980). Rock types from Koko Seamount lavas are shown separately to emphasize the effects of crystal fractionation including the progressive development of a Eu anomaly due to sanidine fractionation and the curvature in the heavy-rare-earth elements that reflects apatite fractionation. In general the post-caldera lavas from along the chain have patterns of similar shape whose absolute enrichments are due to varying degrees of crystal fractionation. (c) kavas from the post-erosional alkalic stage including the Hana Volcanics (Chen & Frey 1983), the Kalaupapa Basalt (Clague et al. 1982a), the Honolulu Volcanics (Clague & Frey 1982), the Koloa Volcanics (Feigenson 1984; D. A. Clague, unpublished data) and the Kiekie Basalt (D. A. Clague, G. B. Dalrymple, M. H. Beeson & E. D. Jackson, unpublished data). The post-erosional lavas generally have greater relative light-rareearth enrichment and lower absolute heavy-rare-earth abundances than those of the post-caldera alkalic lavas.
246
D. A. Clague
500
individual volcanoes. These models can then be expanded to explain the differences between volcanoes in the Hawaiian Islands and along the Hawaiian-Emperor volcanic chain. It is also clear that the origin of alkalic and stronglyalkalic lavas in Hawaii cannot be determined independently from the origin of the more voluminous tholeiitic lavas.
East M o l o ~
400
:7
////
300 Loti
~ill
l 1
Volume relationsmmagma storage at depthmxenolith assemblages
2OO
100
. Hualal
04
0.8
1.2
Pe 9 ( wt%l
FIG. 11. Zr-P205 plot showing that the pre-shield Loihi lavas (horizontally shaded field) have similar Zr/ P2Os ratios to those of the post-caldera lavas (stippled fields) of the Hualalai, Niihau and E Molokai Volcanics, the Kula Volcanics of Haleakala and the Paniau Basalt of Niihau, whereas the post-erosional lavas (diagonally shaded fields) of the Koloa Volcanics, Kiekie Basalt, Honolulu Volcanics and the Kalaupapa Basalt have distinctly lower Zr/P205 ratios. The Kauai field shown is for the tholeiites of the shield stage and shows the similarity of the sources for the shield stage and post-caldera alkalic lavas. (From Frey & Clague (1983) modified after a figure in Clague et al. (1980.))
lend support to the metasomatic model proposed by Clague et al. (1983). These various interpretations illustrate the problems still to be resolved in unravelling the mantle source characteristics of basaltic magmas in general and Hawaiian basalts in particular. It seems possible that the three mixing end-members include two enriched sources, one the deep plume and the other a source more like that beneath Iceland, and a depleted mid-ocean ridge source. The depleted mid-ocean ridge and the enriched ridge sources probably reside in a heterogeneous lithosphere formed at the E Pacific Rise about 90 Ma ago. The main difference between the two enriched sources is the 90 Ma of radioactive decay that the ridge source has undergone since formation. Successful mixing models will have to integrate Sr, Nd and Pb isotopic data with major- and trace-element data for stratigraphically controlled samples from
The volumes of post-caldera and post-erosional lavas on individual volcanoes are inversely correlated. Thus, moderate to large volumes of post-caldera alkalic lavas occur on the Mauna Kea, Hualalai, Kohala, Haleakala, W Maui, E Molokai and Waianae volcanoes, whereas moderate to large volumes of post-erosional alkalic lavas occur on Haleakala, Koolau, Kauai and Niihau only. Thus Haleakala is the only volcano to have relatively large volumes of both postcaldera and post-erosional lavas. Elsewhere, if post-erosional lavas are abundant, post-caldera stage lavas are either absent or present in exceedingly small volumes. This relationship suggests that the source regions for these alkalic stages are related to one another. If the mantle source produces abundant post-caldera alkalic magmas, there is little remaining to contribute to post-erosional alkalic magmas. However, if little or no post-caldera alkalic magma is produced from the mantle source, abundant post-erosional alkalic magmas may be produced. In a general way the volumes of lavas produced during each stage reflect the degree of partial melting of the mantle sources. Thus Hawaiian volcanism commences with rather small-volume small-percentage partial melts that erupt infrequently. These are the early pre-shield alkalicstage lavas. This is followed by large-volume large-percentage partial melts that erupt frequently. These are the tholeiitic-shield-stage lavas. Post-caldera alkalic-stage lavas are once again small volumes of small-percentage partial melts that erupt relatively infrequently. The strongly-alkalic post-erosional lavas represent very small volumes of very small-percentage partial melts that erupt very infrequently. These relationships are shown diagrammatically in Fig. 15. The magma storage system within and beneath the growing volcano also evolves as a function of eruption rate so that during the earliest infrequent alkalic eruptions no magma storage system can be maintained. As the eruption rate increases, a
Hawaiian alkaline volcanism 1000
t
Honolulu~ K o l o a ~ / ~
700 F
~
2000
Kula
~
Ea~tokai
300 [KiekieN~.~~ - ~ ~ |Kaulapapa .,,~x\'~ ~Y.[
247 K()KO (A)
Iooo 800
Hualalai (A)
~-//L AYSAN (A)
400 300
i'/~-
MOLOKAI (A)
80 60
~
Hualalai(T)
0.1
'
HUALALAI (A)
'--SUIKO(.6,)
20Oi00
O lO IS
6 J IN (A)
~ - - ~
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Ol
I 1 I 01610181I I 04 I(3. 20 K (wf.%)
I I I I 40 60
(b)
{a)
013 '01.5 '017' 'II.0 K [wt%)
'
310 '510
FIG. 12. (a) K-Ba plot showing that the pre-shield Loihi lavas (horizontally shaded fields) have similar K/Ba ratios to those of the lavas of the post-caldera alkalic stage (stippled fields) and the tholeiitic shield stage of Hualalai, whereas the post-erosional lavas (diagonally ruled fields) have distinctly lower K/Ba ratios indicating derivation from more depleted sources. (From Frey & Clague (1983) modified after a figure in Clague et al. (1980).) (b) K-Ba plot showing fields for tholeiitic lavas from Kauai (Napali Member of the Waimea Canyon Basalt (D. A. Clague, unpublished data)) and the Suiko and Ojin Seamounts (Kirkpatrick et al. 1980). Also shown are the fields for post-caldera alkalic-stage lavas from Suiko, Nintoku and Ojin Seamounts (Kirkpatrick et al. 1980), Laysan Island (Dalrymple et al. 198 !) and Koko Seamount (Clague 1974). These data demonstrate that the K/Ba ratios of post-caldera alkalic-stage lavas and their source rocks have been constant during the last 64 Ma of activity along the Hawaiian-Emperor volcanic chain.
0"58 0'54 0"50 L~ 0'46
-
~bs 0"O5 ['~,-.: ~O I i'x~ %%
~1 ~
Upper E. Molokai . . . . . I 0"1L ! ' ~ ' ;4 Lines of constont ~ I~'~" Ir~'x~'bc ~'~.% mixing pr~176176 _~ I i '!"~x.i ~~x %' o "-"'c-["~lHnna, ~ oL , I ', ""4.~"h ~,..Honolulu I~.b< ~ rn . . . . . .<.% ~ 0"3~'--.~ ~ ~'Loih'
J
o-4
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% cb~
o'
l,
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Mixing lines
L nes of constant mixing proportions
0.58
Mixing lines
O.52r
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/
o.2
~/ /i ~U/// / ~
~
,/~ i
I // //~
/'~ .,-"/
f.~ "~--~
I
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~ "5 "~ :.,=
I
' Honomonu o=onu
o ce . . . .
i ....
0.70.30 0.7040 87Sr/86Sr
J , I
0-7050
FIG. 13. Plot of La/Ce versus 87Sr/S6Sr for various lavas from the Hawaiian Islands. The mixing lines are from Chen & Frey (1983) and show models in which small-percentage partial melts of a mid-ocean ridge source are mixed with a primitive bulk-Earth composition source. The broken lines are lines of constant mixing proportions. In general, Hawaiian lavas can be generated from sources composed of 0.5%-3% of 0.1%- 1% partial melts of a depleted source mixed with 99.5%-97% bulk-Earth source. See text for discussion.
I
M
0.702
i
i
0"705 0.704 8ZSr/86Sr
i
0.70{
FIG. 14. Plot of La/Ce versus 87Sr/86Sr for various lavas from the Hawaiian Islands. This model (Clague et al. 1983) has the opposite mixing relations to those shown in Fig. 13. In this model small-percentage partial melts of the bulk-Earth source are mixed with the depleted mid-ocean ridge source. This model has a positive correlation between the percentage partial melting of the bulk-Earth source and the amount of melt mixed with the depleted source, whereas the model shown in Fig. 13 has an inverse correlation between these parameters.
248 108
D. A. Clague i0 0
Volcanic stages ~_lf-2~-3"LErosion I 4
Stage I
Rock .ty ~__ss
gI
Tholeiitic pictrite
I Pre-shield alkolic
iO-~j 107
3o
5 Post-caldera
alkalic 4 Post-erosional alkalic
20~
g
io-21
o~ Transitional basalt
i,h
106 _
io-31 105 104
I0
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io-SI
o
~;
3 4 Time (Ma)
'
5
0
Alkalir basalt Basanite Nephelinite
_ Spinel 30 Iherzolite X ---X E 40 --
~- 5o 123
FIG. 15. Plot showing the eruption rates, eruption frequency and percentage partial melting during the evolutionary stages of a hypothetical Hawaiian volcano. During stage 1, the pre-shield alkalic stage, eruption rates are low and eruptions are infrequent, perhaps 1 in 100 years. The lavas erupted are smallpercentage melts (less than 15%) of relatively depleted sources and consist of basanite to transitional basalt. This stage is shown as lasting for 250 000 years but may be considerably longer. The earliest lavas are unsampled but may be nephelinite, melilitite or perhaps even more alkaline compositions. Stage 2, the tholeiitic shield stage, is characterized by high eruption rates and frequent eruptions of largepercentage partial melts (15% 40%) generated from relatively enriched sources. The eruption frequency and rate shown are based on estimates for Kilauea (Dzurisin et al. 1984). The maximum percentage of partial melting is taken from Wright & Helz (1987). Stage 3, the post-caldera alkalic stage, is characterized by eruptions every 100 to 1000 years at moderate to low eruption rates. The lavas are generated by 10%15% partial melting of relatively depleted sources and consist of transitional to alkalic basalt and related differentiates. Stage 4 follows a period of erosion and volcanic quiescence, this erosional period is between 0.4 and 2.5 Ma long. The post-erosional alkalic stage is also of variable duration but can be as long as 2.0 Ma as shown here. Eruptions are very infrequent, perhaps 1 every 10 000 to 100 000 years, and have very low eruptive rates. The lavas are generated by less than 10~ partial melting of sources nearly as depleted as the source of mid-ocean ridge basalt. The erupted lavas include alkalic basalt, basanite, nephelinite and melilitite.
shallow m a g m a storage system develops w h i c h is c h a r a c t e r i z e d by an overlying caldera. Loihi S e a m o u n t has already evolved to this stage. D u r i n g the period with the greatest e r u p t i o n rates, not only is m a g m a continuously stored in this shallow m a g m a reservoir but a n o t h e r d e e p e r staging area a p p a r e n t l y develops. D u r i n g the
70
S )inel Iherzolite
S )inel Iherzolite
Spinel Iherzolite
I -
i~
Garnet Iherzoite
Melilitite
60
Stage 4
X . . . .
20
"~ Tholeiite
o "..IT_ P, 3
~
Stage 3
O~n~crust
2 Tholeiitic shield "C
Stage 2
5 SL
40
I
Garnet Iherzoite
Garnet Iherzoite
-
_
Garnet Iherzoite I
\Magma /I \,~accumulationl
j
I
J
/,I
2 \\
Magma 80 - -// \\ accumulation
i Magma I / i Xaccumulation! ' ' /~1 ~ X--X
90 I00"
Magma (/i~ accumulation/.//!~
FIG. 16. Cross-sections showing the structure beneath a typical Hawaiian volcano during the four eruptive stages. The stability fields for spinel lherzolite and garnet lherzolite are from Sen (1983). The lines denoted by x - x show the maximum depth from which xenoliths are brought to the surface. Zones of magma accumulation are shown. The ocean crust is shown flexing beneath the accumulating load of the volcano and thickening due to dyke injection. Stage 1 lavas erupt onto ocean crust and come from perhaps 80 km deep and carry rare xenoliths of spinel lherzolite to the surface. No shallow storage zones exist. Stage 2 lavas come from 65-70 km deep (the depth of the deepest earthquakes (Klein 1982)) but pass through two magma storage zones, a deep zone (D) at the base of the thickened down-warped crust and a shallow zone (S) within the volcano and underlying a summit caldera. The magma storage zones filter out all xenoliths from deeper levels so that the erupted lavas contain only xenoliths derived from above 3 + km. Stage 3 lavas come from 75-85 km but pass through a deep magma storage zone at the base of the crust. The shallow storage chamber apparently solidifies as the eruption rate declines at the end of stage 2 or the beginning of stage 3 (see Fig. 14). Xenoliths from above the deep storage zone are carried to the surface and include dunite, wehrlite and a variety of gabbroic rocks. Stage 4 lavas come from depths of around 100 km and move directly to the surface. Xenoliths from within the garnet stability field are occasionally brought to the surface; these probably originate at depths of about 80 km. The lavas from stages 2 and 3, and the late parts of stage 1, erupt from distinct rift zones whereas those from stage 4, and perhaps the early part of stage 1, erupt randomly.
Hawaiian alkal&e volcanism
249
TABLE 6. Distribution of Hawaiian xenofiths* Eruptive stage
Lava types
Dykes, sills, veins
Cumulates
Metamorphic rocks
Post-erosional alkalic
Alkalic basalt Basanite Nephelinite Nepheline-melilitite Alkalic basalt
< 1%
1%, olivine cumulates dominant
99%, dunite ,~ wehrlite,~ lherzolite > harzburgite, garnet peridotite locally
Post-caldera alkalic
Ankaramite
Hawaiite Mugearite Trachyte
3%, veins dominant 35%, olivine 62%, dunite >>wehrlite cumulates > pyroxene cumulates 14%, dykes and sills 53%, pyroxene 33%, dunite>>wehrlite dominant cumulates > olivine cumulates 14%, dykes and sills 57%, pyroxene 29~, dunite >>wehrlite dominant cumulates > olivine cumulates
Shield-building
Tholeiite
75%, dykes and sills 25%, olivine dominant cumulates dominant
None
Pre-shield alkalic (Loihi)
Alkalic basalt Basanite
None
99%, dunite >>lherzolite
1%, olivine cumulates
* Data for all but the pre-shield alkalic stage from Jackson (1968). The percentages indicate the relative percentages of dykes, sills and veins, cumulates and metamorphic rocks. The symbols ~, >, >> indicate the relative abundance of each lithology. post-caldera alkalic stage, as eruption rates wane, the shallow magma reservoir can no longer be maintained and it crystallizes. Magmas continue to slow their ascent and to differentiate within the deeper staging area, probably located near the base of the oceanic crust at a depth of about 20 km. By the time of the post-erosional alkalic stage, the magmas ascend rapidly from the mantle without slowing or stopping in any magma storage zones. The development of magma storage or staging areas is schematically shown in Fig. 16. The main evidence for the preceding model is the assemblages of xenoliths that occur in lavas of the different eruptive stages (Table 6). Although it would be useful if these xenolith populations reflected the mantle through which the lavas ascend, it seems more likely that they reflect the development of magma storage reservoirs which, once formed, act as hydraulic filters that remove xenoliths carried up from greater depths (Clague & Dalrymple, 1987). Thus the presence of lherzolite in lavas of only the preshield and post-erosional alkalic stages reflects the lack of shallow reservoirs during these times of low eruption rates. The tholeiitic lavas apparently pass through a double filter and generally contain only fragments of veins, dykes and shallow olivine cumulates. The post-caldera alkalic-stage lavas erupt after the shallow reservoir
has crystallized, but the deeper staging area still filters out any spinel lherzolite or garnet-bearing deep-origin xenoliths. The lavas, however, do pick up a shallower xenolith assemblage ofdunite, wehrlite and gabbro that includes rocks that are part of the oceanic crust and cumulates from earlier eruptive stages of the volcano.
Conclusions Hawaiian volcanoes are constructed of lavas erupted during four sequential stages. Stage 1 is characterized by submarine eruption of alkalic basalt and basanite followed by eruption of tholeiitic lavas. Stage 2 is characterized by eruption of enormous volumes of tholeiitic basalt which form the shield volcanoes. Stage 3 follows collapse of the summit caldera and consists of a cap of alkalic basalt and related differentiates which commonly erupt from vents aligned along the rift zones developed during stage 2. Stage 4 follows a period of volcanic quiescence of as much as 2.5 Ma and is characterized by infrequent small-volume eruptions of nephelinemelilitite, nephelinite, basanite and alkalic basalt from vents unassociated with pre-existing structures. Early stage 1 and stage 4 lavas are relatively primitive and contain deep-origin xenoliths of
z5o
D. A. Clague
lherzolite. Stage 2 and 3 lavas are generally m o d e r a t e l y differentiated indicating substantial residence time in sub-caldera and perhaps d e e p e r m a g m a storage sites. E r u p t i o n rates and xenolith populations in the lavas of the four stages enable a model to be constructed in w h i c h stage 1 lavas are small volumes of small-percentage partial melts that erupt infrequently; no m a g m a storage system is m a i n t a i n e d and deep-seated xenoliths are brought to the surface. Stage 2 lavas are large volumes of large-percentage partial melts that erupt frequently; two m a g m a storage zones are m a i n t a i n e d b e n e a t h the caldera and b e n e a t h the base of the oceanic lithosphere. These storage zones filter out all deep-origin and crustal xenoliths. Stage 3 lavas are again small volumes of small-percentage partial melts that erupt infrequently; the shallow storage zone is no longer m a i n t a i n e d but the deeper zone still filters out deep-origin xenoliths. Crustal xenoliths are c o r n -
m o n in these lavas. Stage 4 lavas are very small volumes of very-small-percentage partial melts that erupt infrequently following a long volcanic quiescence. As for early stage 1 lavas, no m a g m a storage zones are m a i n t a i n e d and deep-origin xenoliths are brought to the surface. The m a g m a sources are not h o m o g e n e o u s through these eruptive stages. In general the early stage 1 and stage 4 lavas are derived from the most depleted sources and the stage 2 lavas from the most enriched source. Rare-gas contents of H a w a i i a n lavas indicate that at least one of the source compositions involved in g e n e r a t i o n of the lavas is a primitive, p r e s u m a b l y deep, undegassed c o m p o n e n t . Various mixing models can be constructed to explain the source compositions, but the characteristics of the e n d - m e m bers are not well constrained. At least three c o m p o n e n t s are required by radiogenic isotopic data.
References ALLEGRE, C. J., STAUDACHER,T., SARDA, P. & KURZ, M. 1983. Constraints on evolution of Earth's mantle from rare gas systematics. Nature, Lond. 303, 762-6. BARGER, K. E. & JACKSON, E. n. 1974. Calculated volumes of individual shield volcanoes along the Hawaiian-Emperor chain. U.S. Geol. Surv. J. Res. 2, 545-50. BASALTIC VOLCANISM STUDY GROUP. 1981. Basaltic Volcanism on the Terrestrial Planets, Pergamon Press, Oxford. BEESON, M. H. 1976. Petrology, mineralogy and geochemistry of the East Molokai Volcanic Series, Hawaii. U.S. Geol. Surv, Prof. Pap. 961, 53 pp. BENCE, A. E., TAYLOR, S. R. & FISK, M. 1980. Majorand trace-element geochemistry of basalts from Ojin, Nintoku, and Suiko Seamounts of the Emperor Seamount Chain: DSDP-180D Leg 55. In JACKSON, E. D., KOIZUMI,I., et al. (eds.) Initial Reports of the Deep Sea Drilling Project, 55, pp. 599-606. U.S. Government Printing Office, Washington DC. BREY, G. 1978. Origin of olivine melilitites--chemical and experimental constraints. J. Volcanol. geotherm. Res. 3, 61-88. BUDAHN, J. R. & SCHMITT, R. A. 1984. Petrogenic modelling of Hawaiian tholeiitic basalts. A geochemical approach. Geochim. cosmochim. Aeta, 49, 67 87. CHASE,T. E., MENARD,H. W. & MAMMERICKX,J. 1970. Bathymetry of the north Pacific. Scripps Institution of Oceanography Charts 2, 7, 8, Scripps Institution of Oceanography, La Jolla, CA. CHEN, C.-Y. & FREY, F. A. 1983. Origin of Hawaiian tholeiites and alkalic basalt. Nature, Lond. 302, 785-9. - - & - - 1985. Trace element and isotope geochemistry oflavas from Haleakala Volcano, East Maui:
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regions for Hawaiian basalts-rare earth element evidence for mantle heterogeneity. In: DICK, H. J. B. (ed.) Magma Genesis, Bull. Oregon Dep. Geol. Miner. Ind. 96, pp. 169-83. MACDONALD, G. A. 1968. Composition and origin of Hawaiian lavas. In: COATS, R. E., HAY, R. L. & ANDERSON, C. A. (eds) Studies in Volcanology, Geol. Soc. Am. Mere. 116, pp. 477-522. -& ABBOTT, A. T. 1970. Volcanoes in the Sea, University of Hawaii Press, Honolulu, HI. & KATSURA,T. 1962. Relationship of petrographic suites in Hawaii. In : MACDONALD,G. A. & KUNO, H. (eds) The Crust of the Pacific Basin, Am. Geophys. Union Monogr. 6, pp. 187-95. & -1964. Chemical composition of Hawaiian lavas. J. Petrol. 5, 82-133. & POWERS, H. A. 1968. A further contribution to the petrology of Haleakala volcano, Hawaii. Bull. Geol. Soc. Am. 79, 877-88. MALAHOFF, A., MCMURTY, G. M., WILTSHIRE,J. C. & YEn, H.-W. 1982. Geology and geochemistry of hydrothermal deposits from active submarine volcano Loihi, Hawaii. Nature, Lond. 289, 234-9. MOORE, J. G. & CLAGUE, D. A. 1981. Loihi seamount lavas: volatile contents. Eos, 62, 1083. -& CLAGUE, O. A. 1982. Volatile contents of Loihi lavas, Hawaii. Generation of Major Basalt Types, IA VCEI-IAGC Scientific Assembly, Reykjavik, Iceland, p. 100. & NORMARK, W. R. 1982. Diverse basalt types from Loihi seamount, Hawaii. Geology 10, 88-92. --, NORMARK, W. R. & LIPMAN, P. W. 1979. Loihi seamount--a young submarine Hawaiian volcano. Hawaii Symp. on Intraplate Volcanism and Submarine Volcanism, Abstracts, p. 127. Hilo. Hawaii. NAUGHTON, J. J., MACDONALD, G. m. & GREENBERG, V. A. 1980. Some additional potassium-argon ages of Hawaiian rocks: the Maui Volcanic Complex of Molokai, Maui, Lanai and Kahoolawe. J. Volcanol. geotherm. Res. 7, 339-55. RISON, W. & CRAIG, H. 1983. Helium isotopes and mantle volatiles in Loihi Seamount and Hawaiian island basalts and xenoliths. Earth planet. Sci. Lett. 66, 407-26. RODEN, M. F., FREY, F. A. & CLAGUE, O. A. 1984. Geochemistry of tholeiitic and alkalic lavas from the Koolau Range, Oahu, Hawaii: implications for Hawaiian volcanism. Earth planet. Sci. Lett. 69, 141-58. ROEDDER, E. 1965. Liquid CO2 inclusions in olivinebearing nodules and phenocrysts for basalts. Am. Mineral. 50, 1746-82. 1983. Geobarometry of ultramafic xenoliths from Loihi seamount, Hawaii on the basis of CO2 inclusions in olivine. Earth planet. Sci. Lett. 66, 369-79. SEN, G. 1983. A petrologic model for the constitution of the upper mantle and crust of the Koolau Shield, -
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DAVID A. CLAGUE, U.S. Geological Survey, 345 Middlefield Road, Menlo Park, CA 94025, U.S.A.
Geochemistry of ocean island basalts from the South Atlantic: Ascension, Bouvet, St. Helena, Gough and Tristan da Cunha Barry L. Weaver, David A. Wood, John Tarney and Jean Louis Joron S U M M A R Y: Basaltic and hawaiitic lavas from the S Atlantic Ocean islands of Ascension, Bouvet, St. Helena, Gough and Tristan da Cunha have been analysed for major elements and a wide range of trace elements. There is marked chemical (particularly trace-element) diversity between these islands which parallels observed Pb, Sr and Nd isotopic variations, and implies considerable large-scale heterogeneity in the source regions for ocean island volcanism. Trace-element and isotopic variation within individual islands suggests significant small-scale source heterogeneity. Abundance ratios between highly-incompatibletrace elements (Rb, Ba, Th, U, K, Ta, Nb and La) appear not to be fractionated during partial melting (except at low degrees of melting in the production of Tristan da Cunha lavas) and can be used to infer source characteristics. Lavas from the islands of Ascension, Bouvet and St. Helena have comparable highlyincompatible trace-element ratios (e.g. La/Nb, Th/Ta, La/Th, Th/U, Ba/La and Ba/Nb), with the exception of the apparent depletion of Rb and K relative to other highlyincompatible elements in St. Helena lavas. Tristan da Cunha and Gough lavas have similar La/Th ratios but much higher La/Nb, Th/Ta, Th/U, Ba/La and Ba/Nb ratios than do lavas from the other islands. These differences reflect depletion in Nb and Ta and enrichment in Ba relative to other highly-incompatibleelements in Gough lavas. The anomalous behaviour (enrichment) of Nb and Ta compared with other highlyincompatible trace elements, and specific consideration of relations between Ba, La and Nb, suggest that ancient subducted ocean crust, rather than primitive or depleted mantle, is the main component of the mantle source for ocean island basalts. Gough, Tristan da Cunha and Walvis Ridge lavas seem to be derived from such a source which has been contaminated by a component with high Ba/Nb, La/Nb and Ba/La ratios. Correlation between Ba/Nb and La/Nb in Gough basalts and hawaiites implies that these lavas contain variable proportions of this component. Pelagic sediment has the high Ba/Nb, La/Nb and Ba/La ratios required of this contaminant, and a few per cent of such material in the mantle source for Gough, Tristan da Cunha and Walvis Ridge lavas would account for their trace-element chemistry.
Introduction Ocean island basalts (OIBs) are generally regarded as being derived from chemically-anomalous mantle sources, and are associated with hot-spot activity (Morgan 1971, 1972). The strong enrichment in incompatible trace elements evident in these volcanics (e.g. Wood et al. 1981) represents a relatively recent event in the source region, as isotopic constraints demand that the magmas are derived from mantle sources with a time-integrated depletion in, for instance, Rb relative to Sr and Nd relative to Sm (O'Nions et al. 1977). The relationship of this source to the continental crust and the mantle source for 'normal' (depleted) mid-ocean ridge basalts (NMORB) is of considerable interest, and has been extensively studied using isotope systematics (O'Nions et al. 1977; Sun 1980; White & Hofmann 1982; Zindler et al. 1982; White 1985). Four main hypotheses have been advanced to explain the origin of chemically-anomalous OIB. These hypotheses suggest that the origins of these volcanics are as follows:
1 From primordial lower mantle which has undergone little or no chemical fractionation since shortly after the formation of the Earth (Schilling 1973; Sun & Hanson 1975; Dupr6 & All6gre 1980; All6gre 1982). 2 From subducted ocean crust which has had a considerable residence time at the boundary between the upper and lower mantle (Hofmann & White 1982; Ringwood 1982; Vollmer 1983; White 1985). 3 From material which was originally part of the subcontinental mantle, but which is now resident in the deep asthenosphere (McKenzie & O'Nions 1983; Cohen et al. 1984). 4 From a mantle source which has been contaminated by recycled continental crustal material (Hawkesworth et al. 1979; Cohen & O'Nions 1982; White 1985). However, relatively few comprehensive traceelement data are available for ocean island lavas which might provide additional important constraints on these possibilities. Here we report major and detailed trace-element data for OIBs
From: FITTON,J. G. & UPTON,B. G. J. (eds), 1987, Alkaline Igneous Rocks,
Geological Society Special Publication No. 30, pp. 253-267.
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B . L . W e a v e r e t al.
254
from five islands in the S Atlantic Ocean: Ascension, Bouvet, St. Helena, Gough and Tristan da Cunha (Fig. 1). These ocean islands are well characterized isotopically (O'Nions & Pankhurst 1974; O'Nions et al. 1977; Sun 1980; Cohen & O'Nions 1982; White & Hofmann 1982), and display considerable inter-island (and often intra-island) isotopic heterogeneity (Sun 1980; Cohen & O'Nions 1982). They are individually representative of islands which are both typical and extreme in terms of the isotopic variation observed in ocean island volcanics. Gough and Tristan da Cunha lavas have Nd and Sr isotope compositions which fall on the mantle array and lie close to supposed bulk Earth 143Nd/144Nd and 87Sr/S%r ratios (O'Nions et al. 1977), although He isotopes are suggestive of a non-primitive source (Kurz et al. 1982). Compared with MORB and other OIBs from the Atlantic Ocean, Gough and Tristan da Cunha lavas display anomalous 2~176 2~ 2~ and 2~176 ratios, having non-radiogenic 2~176 ratios and somewhat radiogenic :~176 ratios. However, St. Helena lavas plot somewhat to the left of the Nd-Sr mantle array with 143Nd/144Nd ratios greater than and 87Sr/86Sr ratios less than bulk Earth values. Additionally, St. Helena lavas have unusually radiogenic Pb isotope compositions (Sun 1980). Ascension and Bouvet lavas might be termed more isotopically typical of ocean island volcanics; they plot on (or in the case of Ascension possibly slightly to the left of) the Nd-Sr mantle
C *o
F.
A
'"
~
FIG. 1. Map of the main ocean islands in the Atlantic Ocean to the S of latitude 40~ : M, Madeira; C, Canary Islands; CV, Cape Verde Islands; F, Fernando de Noronha; A, Ascension; SH, St. Helena; T, Trindade; TC, Tristan da Cunha; G, Gough; DT, Discovery Tablemount; B, Bouvet.
array, although Bouvet lavas have less radiogenic Nd and more radiogenic Sr than do Ascension lavas. Pb isotope ratios for Ascension and Bouvet lava,s are comparable with those of the majority of Atlantic Ocean islands (e.g. the Canary Islands, the Azores, the Cape Verde Islands and Fernando de Noronha (Sun 1980)).
Sampling and analytical techniques Samples from Ascension (Daly 1925; Harris 1983), Bouvet (Verwoerd et al. 1976; Imsland et al. 1977; LeRoex & Erlank 1982), St. Helena (Daly 1927; Baker 1969) and Tristan da Cunha (Baker et al. 1964) were obtained from the collections of the British Museum (Natural History), London. The samples from Gough Island were obtained from the collections of LeMaitre (1962) at the Department of Earth Sciences, University of Cambridge. Relevant geological details for each island, some petrographic data and a limited amount of geochemical data are available in the above cited publications. In all instances samples were chosen to cover the full compositional range of lavas erupted on each island, although only data for basaltic and hawaiitic rocks are reported here. Data for the major elements and the trace elements Cr, V, Ni, Zn, Rb, Sr, Ba, Zr, Nb, Y, Th, La, Ce and Nd were obtained by X-ray fluorescence (XRF) analysis at the University of Leicester using a Philips PWl400 automatic spectrometer. Major-element determinations were made using an Rh X-ray tube on glass fusion beads utilizing a lithium tetraborate-lithium metaborate flux and with calibrations produced against international standards. Major-element data have been corrected to take account of weight loss on ignition of the sample at 900~ Trace-element concentrations were determined on pressed powder pellets using calibrations produced against international standards and standard spikes. Mass absorption corrections were applied using the intensity of the Rh Kct Compton scatter peak for the elements Ni, Zn, Rb, Sr, Y, Zr, Nb and Th and by using the intensity of the W L0r Rayleigh scatter peak combined with Fe K~ counts to cross the Fe absorption edge for the elements Cr, V, Ba, La, Ce and Nd. A more detailed description of the analytical procedures is given by Weaver et al. (1985). Data for the the trace elements Cr, Ni, Sc, Co, Cs, Rb, Ba, Zr, La, Ce, Eu, Tb, Hf, Ta, Th and U were obtained by epithermal neutron activation analysis (ENAA) at C.N.R.S., Universit6 Pierre et Marie Curie, Paris. Full details of the
Geochemistry of ocean island basalts analytical procedure, the precision and the accuracy are given by Chayla et al. (1973) and Jaffrezic et al. (1980). Where there is overlap in the tabulated data (Table 1) between the two analytical techniques we have reported the XRF data for the elements Cr, Ni, Rb, Ba, Zr and Ce and the NAA data for Th and La. A standard Fe203/FeO ratio of 0.22 was used in the calculation of CIPW normative compositions.
Geochemistry Major elements Representative analyses of lavas from the islands of Ascension, Bouvet, St. Helena, Gough and Tristan da Cunha are presented in Table 1. In this paper we have attempted to select only those lavas which by chemical criteria appear to have undergone limited low-pressure crystal fractionation. Our main criteria in selecting samples have been based on uniformity of trace-element ratios rather than major-element composition. Essentially, only those lavas which have a basaltic or hawaiitic major-element chemistry are considered. In Fig. 2 normative compositional data for basalts and hawaiites from the islands are plotted in terms of the components ne-di-ol-hy-qz. Considerable variation in normative composition exists between islands, and often within individual islands. Bouvet basalts and hawaiites are uniformly ol-di-hy normative and have relatively high di/ol and di/hy ratios (Fig. 2). St. Helena basalts and hawaiites (with the exception of one sample) are ne normative, while Ascension basalts and hawaiites are somewhat variable in composition, ranging from ne normative to moderately Di
/
Ol
k
Hy
FIG. 2. Plot of normative components ne-ol-di-hy-qz in basalts and hawaiites from Bouvet (I~), Ascension (-k), Gough ( 9 St. Helena (D) and Tristan da Cunha (0). All normative compositions were calculated using an Fe203/FeO ratio of 0.22.
255
hy normative (Fig. 2). There is an interesting contrast between the chemistry of Gough and Tristan da Cunha basalts and hawaiites. The Tristan da Cunha samples range from moderately to mildly ne normative, with some samples straddling the di-ol join with approximately constant ol/di ratios. Gough samples, however, are extremely variable in normative mineralogy, ranging from hawaiites with a negligible amount of normative ne (essentially lying on the di-ol join) through basalts and hawaiites with quite variable proportions of ol, di and hy to a hawaiite with a small amount of normative qz (Fig. 2). Clearly, few (if any) of the samples considered in this study are representative of primary mantle melts; in general Mg numbers (100 Mg/(Mg+ FEZ+)) and Ni and Cr contents (e.g. Table 1) are low, and those lavas with Mg numbers which might be appropriate for primary mantle melts are generally olivine- and/or clinopyroxenephyric. Indeed, in these islands the major-element compositions are largely controlled by removal or accumulation of the assemblage olivine+ clinopyroxene + plagioclase + Fe-Ti oxide in basaltic and hawaiitic lavas (Zielinski & Frey 1970; LeRoex & Erlank 1982; Harris 1983; LeRoex 1985).
Trace elements In this section we are primarily concerned with assessing the compositional variation in the mantle sources for OIBs. The most useful elements for this purpose are those trace elements which have very low bulk crystal-liquid partition coefficients D during both partial melting and fractional crystallization processes. In particular, ratios between such elements should be representative of the ratios in the mantle source regions. Cs, Rb, Ba, Th, U, K, Ta, Nb and La are highly-incompatible trace elements for which D is very low (less than or equal to 0.01), and the behaviour of these elements will be emphasized. In evolved compositions (mugearites and trachytes) from each island ratios between some of these trace elements become influenced by lowpressure crystal-liquid processes. In the less evolved basaltic and hawaiitic compositions considered here there is no evidence for highlycompatible trace-element fractionation attributable to low-pressure crystallization. However, the degree to which, in particular, alteration and partial melting are capable of affecting abundance ratios needs to be assessed.
Alteration In general the abundances of the large-ion lithophile (LIL) elements Cs, Rb, Ba, Th, U, K
256
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Geochemistry o f ocean &land basalts and Sr are susceptible to variation owing to lowtemperature alteration processes (Hart et al. 1974; Clague & Frey 1982). Element mobility accompanying sub-aerial alteration and hydration is evident in a number of the ocean island lavas studied, although the effects are not always consistent. Cs abundances are particularly unreliable, and can be either enriched or depleted relative to immobile trace elements (Nb, Ta, La) in both altered and otherwise fresh lavas. U is moderately mobile upon alteration, usually with a loss of U leading to an increase in Th/U ratio in altered lavas (Fig. 3). K and Rb appear to be somewhat less mobile than Cs and U, the mobility
257
of K and Rb generally resulting in an increase in the K / R b ratio. There is also some evidence that Ba and Th may display more limited mobility. These effects clearly contribute to the scatter evident in the LIL element plots of Fig. 3, and it is necessary to use care in the interpretation of the behaviour of some of these elements.
Partial melting For trace elements to be good indicators of mantle source composition they must not be strongly fractionated during partial melting (e.g. by low degrees of melting or owing to the effect of K2o
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FiG. 3. Plots of Th against other highly-incompatibletrace elements for the islands of Ascension ( , ) , Bouvet (1), St. Helena (F1), Gough (O) and Tristan da Cunha (0).
258
B. L. Weaver et al.
residual mineralogy). This is clearly a problem with alkali basalts, as the degree of melting may be less than 5~o-10~ (Clague & Frey 1982). At these low degrees of melting, phases other than olivine, orthopyroxene, clinopyroxene and garnet (such as amphibole and/or phlogopite), which are capable of fractionating Ba, K, Rb and Th from other highly-incompatible trace elements (Philpotts & Schnetzler 1970), may be residual. From the systematic behaviour of highlyincompatible trace elements outlined below we believe that for Ascension, Bouvet, St. Helena and Gough there is little chemical fractionation attributable to partial melting. However, anomalous values for some trace-element ratios in Tristan da Cunha lavas suggest that these ratios may not be representative of the source. The geographical proximity of Gough and Tristan da Cunha implies that these islands are related to the activity of the same hot spot, and indeed Nd, Sr, Pb and He isotopes indicate a chemically similar source for their lavas. However, the strongly ne-normative nature of Tristan da Cunha lavas relative to Gough lavas (Fig. 2) may be the result of a smaller degree of partial melting. Additionally, significant differences in rare-earthelement (REE) fractionation between the two islands (as expressed by Cen/Yn ratios which have values of 8.0 and 10.8 for Gough and Tristan da Cunha respectively (Table 2)) could be interpreted in terms of a lower degree of partial melting in the production of the Tristan da Cunha lavas. Similar 143Nd/144Nd ratios (Cohen & O'Nions 1982; White & Hofmann 1982) indicate that the source for Tristan da Cunha lavas has had a similar time-integrated Sm/Nd ratio to that of the source for Gough lavas. Relationships amongst the trace elements Rb, Ba, K and Th are somewhat ambiguous regarding the possible role of amphibole and/or phlogopite
as residualphases in the mantle source for Tristan da Cunha lavas. Partition-coefficient data suggest that DR > ORb ~ DBa for amphibole, whereas DK ,~ DRb > DBa for phlogopite (Philpotts & Schnetzler 1970). T h / N b and Th/Ta ratios (Fig. 3; Table 2) are not fractionated signficantly during different degrees of melting in the production of Gough and Tristan da Cunha lavas. Variations in Ba/ Nb, K / N b and R b / N b ratios (Table 2) demonstrate that Ba and K are equally partitioned into the residue in the production of Tristan da Cunha lavas, but that Rb behaves in a more incompatible fashion. However, for residual amphibole and/or phlogopite Ba should behave more incompatibly than K. Clearly it is not possible to constrain the residual mantle mineralogy tightly during the partial-melting event which generated the Tristan da Cunha lavas, except that phases retaining K, Ba and (to a lesser extent) Rb appear to be required. Source characteristics Abundances of La, Ta, Rb, K, Ba and U are plotted against Th (used as a general index of fractionation) in Fig. 3. In addition, selected element ratios (with associated l a standard deviations) for each island are tabulated in Table 2. Mean La/Th ratios are constant for the islands with the exception of Tristan da Cunha, which has a lower La/Th ratio owing to slightly more compatible behaviour of La during the low degree of partial melting. A rather larger variation in Th/Ta ratios between the islands is evident, with Gough and Tristan da Cunha having higher Th/ Ta ratios (1.47 and 1.50 respectively) than those of lavas from Ascension, Bouvet and St. Helena (range 1.00-1.09). Plots of Rb, K, Ba and U against Th display somewhat more scatter than those of Ta and La, largely owing to alteration
TABLE 2. Means and l a standard deviations for selected element ratios in basaltic and hawaiitic lavas from Gough, Tristan da Cunha, St. Helena, Ascension and Bouvet Islands Ce,/Y, Zr/Nb La/Th Ba/Th Rb/Th Th/U Th/Ta La/Nb Th/Nb Ba/Nb K/Nb Rb/Nb Ba/La Gough
8.0 0.9
6.8 0.8
9.1 0.6
154 20
9.5 0.9
4.86 0.28
1.47 0.15
0 . 9 7 0.105 16.1 0.13 0.009 2.8
432 47
0.99 0.15
16.6 1.6
Tristanda Cunha
10.8 1.2
4.2 0.4
7.8 0.8
103 12
8.2 1.0
4.50 0.08
1.50 0.12
0.86 0.108 11.4 0 . 0 6 0.007 0.8
307 16
0.88 0.09
13.2 0.7
St. Helena
7.3 0.9
4.5 0.2
8.9 0.5
77 8
4.9 0.5
3.78 0.10
1.09 0.07
0 . 6 9 0.078 0 . 0 3 0.005
5.9 0.2
179 9
0.38 0.04
8.7 0.7
Ascension
4.6 0.5
5.3 0.4
8.7 1.0
92 12
6.9 0.8
3.54 0.17
1.05 0.09
0.65 0.04
0.073 0.006
6.8 0.5
230 17
0.46 0.05
10.3 1.2
Bouvet
3.9 0.1
6.9 0.3
9.3 0.4
83 4
7.1 0.4
3.68 0.19
1.00 0.03
0 . 6 8 0.073 0.04 0.003
6.0 0.3
272 13
0.52 0.02
8.9 0.4
Geochemistry of ocean island basalts effects. However, it is apparent that St. Helena lavas have lower Rb/Th ratios (average 4.9) than those of Ascension and Bouvet lavas (average ratios 6.9 and 7.1), which are in turn lower than the Rb/Th ratios in Gough lavas (average 9.5). A very similar pattern is evident for K/Th ratios, with the exception that Bouvet lavas have a higher K/Th ratio than do Ascension lavas. Both Gough and Tristan da Cunha lavas (when altered samples are excluded) have considerably higher Th/U ratios (average 4.86 and 4.50 respectively) than those of Ascension, Bouvet and St. Helena lavas (range of average ratios, 3.54-3.78). There is high degree of variation in Ba/Th ratios, often within each island, but Ascension, Bouvet and St. Helena lavas have rather consistent ratios of 77-92, whilst the Ba/Th ratio of 154 in Gough lavas is high. It is evident from the foregoing that, in general, the trace-element characteristics of the lavas from Gough and Tristan da Cunha (where fractionation during melting affects some element ratios) are very different from the characteristics of the lavas from Ascension, Bouvet and St. Helena. Within the latter three islands the St. Helena lavas appear to be somewhat depleted in K and Rb relative to other highly-incompatible trace elements. It is also apparent that within each island there is often a degree of heterogeneity in element ratios (expressed by the standard deviation) which cannot be simply attributed to analytical uncertainty or alteration effects. This observation is in accord with the well-documented isotopic heterogeneity within lavas from a single island (Sun 1980; Cohen & O'Nions 1982). Figure 4 is a plot of Th against Zr, a less incompatible element that is likely to be fractionated from highly-incompatible trace elements (e.g. Th) at low to moderate degrees of partial
400
200 /
/
-
121
Th
J
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J
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110
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FIG. 4. Plot of Th against Zr for basalts and hawaiites from the islands of Ascension ( , ) , Bouvet (*), St. Helena (U]), Gough (C)) and Tristan da Cunha ($).
259
melting. The lavas of Bouvet and Ascension define good linear trends on this diagram which intersect the Zr axis near the origin, indicating slightly more compatible behaviour of Zr than Th during melting. A rather wide variation in Zr/Th ratios in Gough lavas results in an imprecise correlation between Zr and Th, with an extrapolated trend which intersects the Zr axis at a rather higher Zr value implying a significantly more compatible behaviour of Zr than Th during partial melting. Similarly, good correlation of Zr and Th abundances in lavas from St. Helena and Tristan da Cunha have high intercepts on the Zr axis (particularly in the case of Tristan da Cunha). Although difficult to quantify, owing to uncertainties in the degree of relative light REE enrichment in mantle source, high Ce,/Y, ratios for St. Helena, Gough and Tristan da Cunha lavas (7.3, 8.0 and 10.8 respectively) suggest a lower degree of partial melting in the generation of these lavas than for the Ascension and Bouvet lavas (with mean Cen/Y, ratios of 4.6 and 3.9 respectively). It is also clear that for the fomer three islands Zr does not behave as a highly-incompatible element during partial melting. Thus, depending upon the degree of melting, use of ratios employing Zr (e.g. Zr/Nb ratios) will not necessarily reflect source characteristics for OIBs. This becomes apparent from a plot of Zr against Nb (Fig. 5) where Tristan da Cunha lavas display a strong covariation of Zr and Nb which has a considerable intercept on the Zr axis. The distinctive trace-element characteristics of Gough and Tristan da Cunha as opposed to Ascension, Bouvet and St. Helena basalts and hawaiites is well displayed on plots of La, K20 and (particularly) Ba against Nb (Fig. 5). Average La/Nb ratios in Ascension, Bouvet and St. Helena lavas are rather uniform at 0.65-0.69 (Fig. 5; Table 2), but significantly higher at 0.97 in Gough lavas (Fig. 5; Table 2). There is some variation of K/Nb ratios between Ascension, Bouvet and St. Helena lavas with, as already noted, St. Helena basalts and hawaiites (K/Nb = 179) being somewhat depleted in K (Fig. 5) compared with Ascension and Bouvet basalts and hawaiites (K/ Nb values of 230 and 272). The K/Nb ratio of Gough lavas is high at 432 (Fig. 5; Table 2). Extremely large differences in Ba/Nb ratios are evident between lavas from the different islands (Fig. 5); Ascension, Bouvet and St. Helena lavas have consistent Ba/Nb ratios (Fig. 5) of 5.9-6.8, whereas the average Ba/Nb ratio in Gough lavas (Fig. 5) is very high at 16.1 but with considerable dispersion of values about the average ratio (Table 2). Representative plots of incompatible element abundances normalized to an estimated primor-
260
B . L . W e a v e r et al. La
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FIG. 5. Plots of Nb against other incompatible trace elements for basalts and hawaiites from the islands of Ascension (~-), Bouvet (#), St. Helena (R), Gough (O) and Tristan da Cunha (0). dial mantle composition (Wood et al. 1981) for basaltic and hawaiitic lavas from Ascension, Bouvet, St. Helena, Gough and Tristan da Cunha appear in Fig. 6. The order in which the elements are plotted corresponds to decreasing D from right to left in the diagram for equilibration of basaltic melt with four-phase lherzolite (Wood 1979). These diagrams illustrate some interesting features of the trace-element chemistry in addition to those described above. The general patterns are of increasing normalized abundance from right to left in the diagrams, although amongst the highly-incompatible elements Cs-K there is a general decrease in normalized abundance with increasing incompatiblity (Fig. 6). The anomalous behaviour of Sr relative to Ce and Nd is due to plagioclase fractionation. Hf abundances define consistent negative anomalies relative to Zr and P, and Zr/Hfratios are constant between all five islands (range of average ratios from 43 to 45). These values are considerably higher than the primordial Zr/Hf ratio of 31
(Wood 1979), but similar to Zr/Hf ratios observed in other alkali basalt suites (Frey et al. 1978; Clague & Frey 1982). Maximum normalized abundances occur for Nb and Ta (Fig. 6), and Nb/Ta ratios are very uniform (within analytical error) both within and between islands (range of average Nb/Ta ratios from 13.7 to 14.1). Gough lavas are notable for the similar normalized abundances of Nb, Ta, La, K and Ba (Fig. 6). From Nb to Cs there is a general decrease in normalized highly-incompatible element abundances, with the exception of the islands of Tristan da Cunha and (particularly) Gough where Ba is enriched (producing a positive Ba spike) relative to other highly-incompatible LIL elements (e.g. Cs, Rb, Th, U and K). The two most striking features of Fig. 6 are the large positive anomaly present for Nb and Ta in Ascension, Bouvet and St. Helena lavas, and the disappearance of this anomaly in Gough (and to an extent Tristan da Cunha) lavas with the simultaneous development of a positive Ba anomaly. The
Geochemistry of ocean island basalts 100
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F]6.6. Plots of incompatible element abundances normalized to primordial mantle (Wood et al. 1981) for selected basalts and hawaiites from the islands of Ascension, Bouvet (stippled field), St. Helena, Gough and Tristan da Cunha. magnitude of the N b - T a anomaly (as expressed by La/Nb and Th/Ta ratios in Table 2) is constant in Ascension, Bouvet and St. Helena lavas, and typical of OIBs from other islands in the Atlantic Ocean (B. L. Weaver et al., unpublished data) as well as of continental alkali basalts (e.g. Thompson et al. 1984).
Discussion Amongst the highly-incompatible trace elements it is evident that the behaviour of, in particular, Ba, Nb and Ta may provide strong constraints on potential source reservoirs for OIBs. For instance, partial melting (especially at low degrees of melting) of primordial (or primitive) mantle should produce melts with highly-incompatible trace-element (Cs-La) abundances which, when plotted on normalized diagrams such as Fig. 6, would be either unfractionated or, more likely, slightly enriched in more highly-incompatible elements (e.g. Cs) over less highly-incompatible elements (e.g. La). It would not be possible to produce significant N b - T a or Ba anomalies (Fig. 6) from such a primordial source. Furthermore, the depletion of Cs and Rb relative to K and La evident from Fig. 6 argues for derivation of OIBs from a source which is itself depleted in more highly-incompatible trace elements. It
would therefore seem that chemically primitive mantle (such as the lower mantle (All6gre 1982)) is not a suitable source for the production of OIBs. In models for the chemical evolution of the crust-mantle system the continental crust and the mantle source for N-MORB are considered to be complementary reservoirs (e.g. Jacobsen & Wasserburg 1979; O'Nions et al. 1979; All6gre et al. 1983; All6gre & Rousseau 1984). To a first approximation those incompatible trace elements which are depleted in N-MORB are enriched in the continental crust. However, this is clearly not the case when consideration is given to the chemical budgets for Nb and Ta. Subductionzone-related magmas, and indeed the bulk continental crust itself, are strongly depleted in Nb and Ta relative to other highly-incompatible trace elements (e.g. Saunders et al. 1980; Thompson et al. 1984). For example, continental margin basalts and andesites typically have La/Nb ratios in the range 2.2-4.7 (data from the compilations of Ewart (1982)), while estimates of the bulk composition of the continental crust (Taylor & McLennan 1981 ; Weaver & Tarney 1984) suggest that the crust itself is similarly depleted, with La/ Nb ratios of 1.7-2.2. Island-arc volcanics have equally high La/Nb ratios (e.g. Tarney et al. 1981). Estimates of the composition of typical NMORB (Sun 1980; Tarney et al. 1981 ; Pearce
262
B . L . W e a v e r e t al.
1983) yield a La/Nb ratio of close to 1.0 which, considering the very similar D values for La and Nb during MORB genesis (Briqueu et al. 1984), represents the La/Nb ratio in the MORB reservoir. A La/Nb ratio of 0.83-0.94 is likely for the chondritic Earth or primordial mantle (Sun 1980; Taylor & McLennan 1981; Thompson et al. 1984). Thus the production of continental crust with La/Nb > 2 from primordial mantle with La/ Nb ~ 0.9 requires a volumetrically large complementary reservoir with a low (considerably less than 1) La/Nb ratio; the N-MORB source does not represent an appropriate reservoir. However, this reservoir would appear to be the source for ocean island (and continental) alkali volcanics with low La/Nb ratios (about 0.7). It has been suggested that oceanic and continental alkali basalts could be the product of partial melting of subducted ocean crust which has been resident in the deep mantle for a considerable period of time (Hofmann & White 1982; Ringwood 1982). Such a model is consistent with Pb isotope data for OIBs which imply derivation from source regions that have maintained high U/Pb ratios (typical of altered oceanic crust) for periods of time of the order of (1.52.0) x 103 Ma (Tatsumoto 1978; Sun 1980; Chase 1981). With regard to the behaviour of Nb and Ta, a key question is: what chemical modification occurs in oceanic crust during subduction ? Traceelement and isotopic considerations of the chemistry of subduction-related magmas invariably require a component derived from the subducted ocean crust (Kay 1984; O'Nions 1984; Wilson & Davidson 1984). This component either could be a fluid produced by dehydration of the slab which carries LIL elements (and La, Ce?) into the overlying mantle wedge, or may involve a degree of partial melting of the slab, again under hydrous conditions (Kay 1984; Wilson & Davidson 1984; Wyllie 1984). The effect of these processes on the trace-element chemistry of the subducted ocean crust would be that LIL elements would be strongly depleted in subducting N-MORB-type ocean crust. However, it is also likely that the conditions pertaining during dehydration (and melting ?) of the ocean crust would stabilize minor phases (especially sphene (Hellman & Green 1979)) in the ocean crust which would retain high-field-strength elements such as Nb and Ta (Saunders et al. 1980; Briqueu et al. 1984; Thompson et al. 1984). Thus Nb and Ta would become strongly decoupled in their geochemical behaviour from other highly-incompatible trace elements, and the subducted ocean crust would develop low La/Nb and Ba/Nb ratios whilst being severely depleted in LIL elements. Unfortunately, lack of knowledge of phase relations
and partition coefficients for the probable mineralogy of the ocean crust in the deep mantle (Ringwood 1982) precludes modelling of subsequent partial melting of this material. It is perhaps probable that moderate degrees of melting (10~-15~) would produce undersaturated liquids with strong light REE enrichment (Hofmann & White 1982), positive N b - T a anomalies and increasing relative depletion in the more highly-incompatible trace elements reflecting the composition of the old ocean crust source. Such melts would mix with the mantle to form an enriched hot-spot component, the composition of which is strongly reflected in the chemistry of lavas from Ascension, Bouvet and St. Helena. This component might further interact and mix with the depleted mantle source of N-MORB (producing transitional or T-type MORB), or could invade and vein (metasomatize) depleted upper-mantle peridotite (Menzies 1983; Hawkesworth et al. 1984) which might itself be subsequently melted to produce alkali basalts. These general features are demonstrated in a plot of Ba/Nb against La/Nb in Fig. 7. As Nb is used as a denominator in the ratios plotted on both axes, mixing trends will be straight lines on
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] 1.2
La/Nb
FIG. 7. Plot of Ba/Nb against La/Nb ratios for basalts and hawaiites from the islands of Ascension (~), Bouvet (,), St. Helena ([-]), Gough (O) and Tristan da Cunha (0). Also plotted are average ratios for basalts from DSDP Holes 525A, 527, 528 and 530A from the Walvis Ridge (data from Humphris and Thompson (1983)), and fields showing the average Ba/Nb and La/ Nb ratios in N-MORB and primordial mantle.
Geochemistry of ocean island basalts
this diagram. Since DBa~
263
ditions of partial melting where DNb,~DLa. Rather, the chemistry is best explained by mixing between a 'typical' hot-spot component (represented by Site 530A basalt or Ascension, Bouvet and St. Helena lavas) and a component with high Ba/Nb and La/Nb ratios. It is clear from the Ba/ Nb-La/Nb relations in Fig. 7 that an N-MORB component or source is not a suitable endmember in any mixing model, as pointed out by Richardson et al. (1982) for Walvis Ridge lavas and LeRoex (1985) for Gough lavas. The wide range in, and good correlation between, Ba/Nb and La/Nb ratios in Gough lavas (Fig. 7) suggests variable proportions of the high Ba/Nb and La/ Nb ratio components in these lavas. This trend is not simply the result of dilution of Nb causing the Ba/Nb and La/Nb ratios to increase; with decreasing magnitude of the N b - T a anomaly (e.g. increasing La/Nb ratio) Ba abundances increase relative to other highly-incompatible trace elements such that there are reasonable correlations between La/Nb, Ba/Th and Ba/La ratios. Thus the end-member with the high Ba/ Nb and La/Nb ratios must be a component which itself has a negative Nb anomaly but a positive Ba anomaly. Such chemical characteristics are typical of pelagic sediments, which are strongly enriched in Ba relative to Th, Rb and Nb and in La relative to Nb (Hole et al. 1984; Kay 1984; Thompson et al. 1984). Terrigenous sediments are also a possible contaminant, but typically have rather low La/Nb, Ba/Nb and Th/Nb ratios which makes them unsuitable end-members in mixing models. Hence the trace-element data are qualitatively suggestive (as are the He isotope data (Kurz et al. 1982)) of a (pelagic) sedimentary component in the Gough, Tristan da Cunha and Walvis Ridge lavas. We can make some attempt at quantifying the mixing process using Ba/Nb and La/Nb ratios, and approach the question of low-pressure mixing of the common parental ocean island magma type with pelagic sediment versus high-pressure mixing of components in the source region of some OIBs. A parental (or primary) OIB would have Ba/Nb and La/Nb ratios of 6.5 and 0.68 (typical of Ascension, Bouvet and St. Helena lavas (Table 2)), and might have a Nb content (corrected for the effects of low-pressure crystal fractionation) of 25 ppm. For an estimate of the composition of pelagic sediment we use the Pacific authigenic weighted mean sediment (PAWMS) of Hole et al. (1984), which has Ba/ Nb and La/Nb ratios of 1070 and 20.6 respectively and an Nb concentration of 1.25 ppm. A mixture of 20~ PAWMS + 80~ parental OIB is required to generate high Ba/Nb and La/Nb ratios (19.6 and 0.93 respectively) which approximate to the
264
B . L . W e a v e r e t al.
values for these ratios in Gough lavas (Table 2). Varying the proportions of nannofossil ooze, diatom ooze, ferruginous clay and pelagic clay in the sediment (PAWMS comprises 95~ nannofossil ooze plus 5~ ferruginous clay) has little effect on element ratios in the sediment or on the amount of pelagic sediment required as a contaminant. Addition of such large proportions of pelagic sediment would have a drastic effect on the major-element composition of the mixture for which there is little evidence. Also, why should such low-pressure mixing (presumably associated with a sub-volcano magma chamber) be apparently restricted geographically to the islands of Gough and Tristan da Cunha in the South Atlantic and to islands in the Indian Ocean (Dupr6 & All6gre 1983)? However, if mixing occurs in the source regions of some OIB the endmembers would be pelagic sediment and subducted ocean crust (the main mantle source component for OIB). The subducted ocean crust would have a Nb content of approximately 3 ppm and Ba/Nb and La/Nb ratios similar to those of the hot-spot material. In consideration of pelagic sediment older than Mesozoic, nannofossil and diatomaceous oozes would not be a component. For a ratio of ferruginous clay (Hole et al. 1984) to pelagic clay (Li 1982) of 2:8, the ancient pelagic sediment might have a Nb content of 12.3 ppm and La/Nb and Ba/Nb ratios of 5.7 and 341 respectively. In this case a mixture of 1~ pelagic sediment plus 99~ OIB source would have a La/Nb ratio of 0.92 and a Ba/Nb ratio of 22, similar to the values for these ratios in Gough lavas. Such a situation is compatible with the injection of a small proportion of sediment into the deep mantle along with subducted ocean crust. Additional constraints on mixing processes in Gough, Tristan da Cunha and Walvis Ridge lavas come from Pb, Sr and Nd isotope data (Sun 1980; Cohen & O'Nions 1982; Richardson et al. 1982). Isotopic trends within Walvis Ridge basalts suggest a limiting composition for the enriched end-member with radiogenic Sr (8~Sr/ 86Sr=0.7058) and non-radiogenic Nd (143Nd/ 144Nd = 0.5122) and Pb (e.g. 2~176 = 17.1) (Richardson et al. 1982). Amongst the Walvis Ridge and Gough and Tristan da Cunha lavas, Site 525A basalts have isotopic characteristics suggesting the greatest proportion of this enriched end-member, which from the Ba/Nb-La/Nb trend in Fig. 7 is equated with the high Ba/Nb and La/Nb ratio component. The isotopic composition of the depleted end-member is not well constrained by the Pb, Sr and Nd data (Richardson et al. 1982). However, Ba/Nb-La/Nb relations (Fig. 7) imply a hot-spot (rather than
MORB) end-member represented by Site 530A basalts, for which isotope data are lacking. The trend defined by Walvis Ridge and Tristan da Cunha lavas on a 2~176176176 diagram (Richardson et al. 1982) suggests that this depleted end-member could have 2~176 and 2~176 ratios similar to those of Ascension and Bouvet lavas. If such a hot-spot component has a 2~176 ratio of approximately 19.5, then about 40~ of the Pb in Gough and Tristan da Cunha (and about 70~ of the Pb in Site 525A) lavas would be derived from the component with high Ba/Nb and La/Nb ratios. A 2~176 ratio for the hot-spot component in excess of those of Ascension and Bouvet lavas would be required, however, as the 2~176 2~176 trend for Gough, Tristan da Cunha and Walvis Ridge lavas (Richardson et al. 1982) is parallel to that for MORB and other OIBs from the Atlantic. Thus the hot-spot component may not have isotopic compositions directly comparable with those of other islands in the S Atlantic. The isotopic composition of the limiting enriched end-member is not apparently compatible with that of typical pelagic sediment. Although non-radiogenic Nd and radiogenic Sr are consistent with the characteristics of pelagic sediment, non-radiogenic Pb is not. Modern Atlantic pelagic sediments generally have rather radiogenic Pb (2~176 ~ 19, 2~176 ~ 15.7 and 2~176 ~ 39.1 (Sun 1980)). However, if the model for derivation of OIB dominantly from old subducted ocean crust with the possibility of contamination occurring in the source is accepted, the sedimentary component of interest will be that subducted with the ocean crust at some considerable time (of the order of (1.52.0) x 103 Ma) in the past. The initial isotopic composition of this sediment is obviously uncertain, but it is likely to have had, relative to upper mantle at that time, radiogenic Sr and Pb and non-radiogenic Nd. The pelagic sediment would have had a Sm/Nd ratio less than the chondritic value (Sm/Nd ~ 0.20), a low Rb/Sr ratio (Rb/Sr 0.05) and very low U/Pb and Th/Pb ratios (# ~ 2 owing to high Pb contents in pelagic sediments of up to 50 ppm (Cohens & O'Nions 1982)). These arguments assume that the sediment maintains its chemical integrity through the subduction process. This is unlikely, but the chemical effects associated with dehydration and/or partial melting of sediment during subduction are very poorly constrained. Subsequent to subduction, growth of radiogenic Nd, Sr and Pb would be inhibited (strongly so in the case of Pb) in this material, and could result in the present isotopic composition predicted for the end-member with high Ba/ Nb and La/Nb ratios observed in the Gough,
Geochemistry of ocean island basalts Tristan da C u n h a and Walvis Ridge lavas. M u c h of the foregoing must r e m a i n speculative until the composition of this c o m p o n e n t can be more rigorously characterized. The variation of tracee l e m e n t ratios (Fig. 7) in G o u g h lavas suggests that a c o m b i n e d detailed trace-element and isotopic study of this island m i g h t be of particular i m p o r t a n c e in a p p r o a c h i n g this problem. In modelling the petrogenesis of Walvis Ridge and G o u g h lavas, R i c h a r d s o n et al. (1982) and L e R o e x (1985) ascribed the unusual isotopic and trace-element characteristics of these lavas to a ' d o u b l e - e n r i c h m e n t ' event in the m a n t l e source
265
region. W e would equate the first trace-element ' e n r i c h m e n t ' event to a melt c o m p o n e n t derived d o m i n a n t l y from old subducted o c e a n crust in the source region for OIB. This event produces OIB with the r a t h e r u n i f o r m trace-element composition of Ascension, Bouvet a n d St. H e l e n a lavas. The second ' e n r i c h m e n t ' event observable in the geochemistry of Gough, Tristan da C u n h a and Walvis Ridge lavas can be e q u a t e d with c o n t a m i n a t i o n of the OIB source region by a small a m o u n t (1%-2%) of ancient pelagic sedim e n t w h i c h was subducted into the deep m a n t l e along with the o c e a n crust.
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HUMPHRIS, S. E. & THOMPSON,G. 1983. Geochemistry of rare earth elements in basalts from the Walvis Ridge: implications for its origin and evolution. Earth planet. Sci. Lett. 66, 223-42. IMSLAND, P., LARSEN,J. G., PRESTVIK, T. & SIGMOND, E. M. 1977. The geology and petrology of Bouvetoya, South Atlantic Ocean. Lithos, 10, 213-34. JACOBSEN, S. B. & WASSERBURG,G. J. 1979. The mean age of mantle and crustal reservoirs. J. geophys. Res. 84, 7411-27. JAFFREZlC, H., JORON, J. L. TREUIL, M. & WOOD, D. A. 1980. A study of the precision attained by neutron activation analysis using international standard rocks GS-N and BCR-I as examples. J. radioanal. Chem. 55, 417-25. KAY, R. W. 1984. Elemental abundances relevant to identification of magma sources. Phil. Trans. R. Soc. Lond., Ser. A, 310, 535-47. KURZ, M. D., JENKINS, W. J. & HART, S. R. 1982. Helium isotopic systematics of oceanic islands and mantle heterogeneity. Nature, Lond. 297, 43-7. LEMAITRE, R. W. 1962. Petrology of volcanic rocks, Gough Island, South Atlantic. Geol. Soc. Am. Bull. 73, 1309-40. LEROEX, A. P. 1985. Geochemistry, mineralogy and magmatic evolution of the basaltic and trachytic lavas from Gough Island, South Atlantic. J. Petrol. 26, 149-86. & ERLANK, A. J. 1982. Quantitative evaluation of fractional crystallisation in Bouvet Island lavas. J. Volcanol. geotherm. Res. 13, 309 38. --, DICK, H. J. B., ERLANK, A. J., REID, A. M., FREY, F. A. & HART, S. R. 1983. Geochemistry, mineralogy and petrogenesis of lavas erupted along the Southwest Indian Ridge between the Bouvet triple junction and 11 degrees East. J. Petrol. 24, 267-318. LI, Y.-H. 1982. A brief discussion of the mean oceanic residence time of elements. Geochim. cosmochim. Acta, 46, 2671-75. MCKENZIE, D. & O'NIONS, R. K. 1983. Mantle reservoirs and ocean island basalts. Nature, Lond. 301,229-31. MENZIES, M. A. 1983. Mantle ultramafic xenoliths in alkaline magmas: evidence for mantle heterogeneity modified by magmatic activity. In . HAWKESWORTH, C. J. & NORRY, M. J. (eds) Continental Basalts and Mantle Xenoliths, pp. 92-110. Shiva, Orpington. MORGAN, J. 1971. Convection plumes in the lower mantle. Nature, Lond. 230, 42-3. -1972. Plate motions and deep mantle convection. Geol. Soc. Am. Mem. 132, 7-22. O'NIoNS, R. K. 1984. Isotopic abundances relevant to the identification of magma sources. Phil. Trans. R. Soc. Lond., Set. A, 310, 591-603. & PANKHURST, R. J. 1974. Petrogenetic significance of isotope and trace element variations in volcanic rocks from the Mid-Atlantic. J. Petrol. 15, 603-34. , EVENSEN, N. M. & HAMILTON, P. J. 1979. Geochemical modelling of mantle differentiation and crustal growth. J. geophys. Res. 84, 6091101. -
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BARRY L. WEAVER, School of Geology and Geophysics, University of Oklahoma, Norman, OK 73019, U.S.A. DAVID A. WOOD, Amoco Europe and West Africa Inc., Amoco House, 1 Stephen Street, Tottenham Court Road, London W 1P 2AU, UK JOHN TARNEY, Department of Geology, University of Leicester, Leicester LE1 7RH, U.K. JEAN LOUISJORON, Group des Sciences de la Terre, Laboratoire Pierre Sue, C.N.R.S., Centre d'Etudes Nucl6aires de Saclay, BP 2, 91190 Gif-sur-Yvette, France.
M a g m a and fluid evolution in the lavas and associated granite xenoliths of Ascension Island Chris Harris and Simon M.F. Sheppard S U M M A R Y : The mildly alkaline lavas of Ascension Island evolved in a H20undersaturated environment. H20 comprises less than 0.4 wt.% of glassy equivalents of the most evolved lavas (comendites) for which there is no evidence that it has been lost by degassing. In contrast, peralkaline granite xenoliths which are chemically similar to the evolved lavas contain silicate glass and fluid inclusions which demonstrate that the original granite body crystallized from a water-saturated magma. Microprobe analysis of the glass inclusions gives low totals suggesting that the H20 content was high (5-8 wt.%). Amphibole 6D values for the granite xenoliths are significantly different from the whole-rock 6D values of the comenditic obsidians while whole-rock 8180 values are similar. It is suggested that the high water content in the original granite magma is due to sea-water addition and that this explains these differences. Sea-water interaction with high-level magma chambers on Ascension may have been the cause of the extensive pyroclastic volcanism which occurred on the island.
Introduction Ascension Island is one of the closest exposed volcanoes to the mid-Atlantic Ridge in the S Atlantic Ocean. It is situated on oceanic crust of age 7 Ma (van Andel et al. 1978), but the oldest dated rocks on the island are 1.5 Ma in age (Harris et al. 1982). There appears to be no connection between the island's location and any special tectonic features and there is no physical evidence that the island is part of a hot-spot track (van Andel et al. 1973). The geology of Ascension has been described in detail by Daly (1925) and Atkins et al. (1964), while the petrology and isotope geochemistry (Sr, Pb, O and H) w a s discussed by Harris (1983), Harris et al. (1982a, b, 1983) and Sheppard & Harris (1985). About half of the island's exposed surface is covered by lava, and the remainder is covered by a variety of pyroclastic deposits. Associated predominantly with the pyroclastics is a suite of plutonic xenoliths including granites. The chemical and fluid evolution preserved in both lavas and plutonic xenoliths are compared in this brief review. Particular emphasis is placed on the most evolved compositions.
Composition and evolution of the lavas and plutonic xenoliths The lavas range from mildly alkaline basalts to peralkaline rhyolites (comendites) with up to 1.5% normative sodium metasilicate (Harris 1983, and unpublished data). Major- and traceelement variations suggest that the comendites
were derived from the alkali basalts by extensive crystal fractionation (about 80% (Harris 1983)). Taken as a whole, the minerals of the plutonic xenoliths show a similar compositional range to the phenocrysts in the lavas (Harris 1983). The less evolved plutonic rocks have textural features which suggest that they are cumulates. The more evolved plutonic rocks (monzonites, quartz syenites and peralkaline granites) are not obviously cumulates and are closely similar in whole-rock chemistry to the trachytes and comendites. They would appear to represent slowly cooled equivalents of magmas similar to the evolved lavas which did not reach the surface. The most evolved plutonic xenoliths are granites which appear to have been derived from a single plutonic body. They are remarkable in that they contain the rare minerals dalyite (K2ZrSi6Ols) and vlasovite (Na2ZrSi4Oll), and are among the few granites described to contain eudialyte (van Tassel 1952; Fleet & Cann 1967; Harris et al. 1982a). Dalyite and vlasovite are indicative of the peralkalinity of these granite xenoliths.
Fluid evolution in the lavas and granite xenoliths There is no evidence, such as amphibole phenocrysts, to suggest that the basalt to comendite lavas were water rich. The water content of the comenditic obsidians from Ascension is less than 0.4 wt.% (Sheppard & Harris, 1985). These obsidians are wholly glassy, non-vesicular and have relatively high C1 and F contents (Table 1) which are consistent with them not having lost signifi-
From: FITTON,J. G. & UPTON,B. G. J. (eds), 1987, Alkaline Igneous Rocks,
Geological Society Special Publication No. 30, pp. 269-272.
269
C. Harris & S. M. F. Sheppard
z7o
TABLE 1. Selected major- and trace-element analyses
SiOz TiO2 A1203 FeO* MnO MgO CaO Na20 K20 LOI(wt.%) H20 + C1 (ppm) F Nb
1
2
74.39 0.17 10.68 4.38 0.03 0.03 0.20 5.52 4.60
73.36 0.27 12.26 3.42 0.14 0.18 0.26 5.47 4.62
------
0.76 0.36 2500 2100 195
3
4
73.20 75.20 0.30 0.16 12.40 11.04 3.46 3.92 0 . 1 3 0.13 0 . 1 6 0.00 0 . 2 0 0.00 5 . 5 2 4.91 4.62 4.64
5
74.46 0.25 10.98 4.32 0.12 0.13 0.00 5.26 4.48
0.62 0.12 0.56 0.31 0.18 2800 167 212 2100 1700 9OO 199 233 241
FeO*, total iron as FeO; LOI, loss on ignition. Nb by X-ray fluorescence (Oxford); C1 and F by wet chemical methods (CRPG, Nancy). 1, silicate melt inclusion in granite H30(1) recalculated to 100% (original total, 91.22 wt. %); 2, OB6 (obsidian from Middleton's Ridge, Ascension); 3, OB7 (obsidian from Middleton's Ridge, Ascension); 4, H30(1) (granite from Five Mile Post, Ascension); 5, H30(2) (granite from Five Mile Post, Ascension). Analysis 1 by microprobe, University of Nancy (Harris, 1986). Analyses 2-5 by inductively-coupled plasma emission spectrometry, CRPG, Nancy (Harris, 1986, and unpublished data); major-element oxides recalculated to 100%. H20 + by manometric measurement of H2 gas extracted during hydrogen isotope analysis. cant H 2 0 by degassing. Ascension lavas as a whole are therefore typical of oversaturated peralkaline series, as they are very water undersaturated (Bailey & Cooper 1978). Fluid inclusions in Ascension granite xenoliths were first described by Roedder & Coombs (1967) and fluid inclusions in the granites and other plutonic xenoliths from Ascension have been further studied by Harris (1986). The most important findings are as follows. 1 The granites contain inclusions of silicate glass and highly saline aqueous fluid, which in some cases occur within the same cavity. This shows that magma and saline fluid coexisted during crystallization (Roedder & Coombs 1967). 2 Low-temperature (below 300~ fluid circulation was minimal. 3 Vapour-rich inclusions contain CO2 and H20 and represent vapour which coexisted with the saline solutions (Roedder & Coombs 1967). 4 The silicate glass inclusions have small contraction bubbles which homogenize between 715 and 790~ (mean of seven runs, 763~ Micro-
probe analyses of the glass have totals which are consistently between 90 and 92 wt.%. Allowing for possible alkali loss under the probe beam it is suggested that H20 is high (probably between 5 and 8 wt.% (Harris, 1986)). A value of 5 wt.% H 2 0 would correspond to a trapping pressure of about 1.5 kb (Dingwell et al. 1984). 5 There is no strong correlation between the salinity of the high-density saline inclusions (liquid, halite cube and vapour bubble) and the temperature of homogenization of the inclusion. However, those inclusions with the lowest salinities (30-50 wt.% NaC1) are not homogenized at 570~ (the upper temperature limit of the stage). The fluid inclusion suite was interpreted by Roedder & Coombs (1967) to be the result of heterogeneous trapping of saline liquid and aqueous vapour during fluid immiscibility. It is suggested by Harris (1986) that fluid immiscibility and loss of the vapour produced increased the salinity in the remaining saline fluid. Thus the lowest-salinity inclusions are interpreted as fluids which were parental to the higher-salinity fluid inclusions.
Discussion The evidence summarized in the previous section suggests that the evolved lavas were always dry but the original granite body crystallized from a water-saturated magma. There is no textural evidence that the high H 2 0 and its eventual saturation occurred only at the end stages of crystallization of the granite. If this were the case the glass inclusions should show a range of H20 contents. If the original melt from which the granite crystallized contained about 5 wt.% HzO and a similar CI content to the comendite obsidians, the salinity of the fluid produced by exsolution of all the dissolved H 2 0 would be low (less than 10 wt.% NaC1) even if all the dissolved C1 were partitioned into the fluid. The high salinity of fluid inclusions suggests that considerable loss of H 2 0 from the system as vapour must have occurred. Two mechanisms by which the high water content in the original granite magma could have been generated are considered below. 1 High degrees of crystallization of comendite magma led to eventual water saturation. This would require greater than 90% fractionation of comendite magma with initial H 2 0 = 0 . 4 w t . ~ and is not consistent with the small differences in major- and trace-element chemistry between the comendite obsidians and the granite xenoliths (Table 1).
Magma andfluid evolution 2 Water was added from some outside source; in this case sea-water or meteoric water are the most obvious sources. The latter would not seem likely owing to the low rainfall and high evaporation rates on Ascension.
271
constituents and there is no evidence that fO2 increased during crystallization of the granite. This should have been the case if significant H 2 loss occurred. Direct addition of about 5 wt.% H20 to comendite magma would decrease its 61so slightly and this may be why the granites show a slightly greater range in 61sO than the comendite obsidians (Sheppard & Harris, 1985).
Stable isotope constraints on fluid evolution These data are discussed by Sheppard & Harris (1985). The granite whole-rock 81 SO values range from 6.4%o to 6.8%0 and are comparable with those of the obsidians (6.7%o_+0.1; n=3). The range in the lava suite as a whole is from +6.0%o to +6.8%o. Separated amphiboles from the granites have 6D values between - 53%o and -58%o which is rather too high for primary magmatic amphiboles and some 20%o higher than the comendite obsidian whole-rock 6D (-76%o to -83%o). Accepted primary magmatic 6D values for amphiboles and basalts are between -70%o and -85%o (e.g. Boettcher & O'Neil 1980). Factors affecting the final 6D value of these amphiboles are (i) the original 6D of the magma, (ii) the hydrogen isotope fractionation factor between amphibole and fluid, (iii) the amount of fluid lost during degassing and (iv) the speciation of hydrogen in the fluid (e.g. H2, H20, CH4 etc.). Taylor et al. (1983) showed that during degassing of H20 from silicic magmas, the residual melt is depleted in deuterium. This depletion is most pronounced when the fraction of H20 remaining is small (el Rayleigh fractionation). The fraction remaining in the Ascension granites is small (for granite H30(1) H20 + = 0.18 wt.% (Table 1)) and, given the combined effect of an amphibole-fluid fractionation factor (1000 In 0d,mph..2o) of about -20%o (Graham et al. 1984) and open-system degassing, the original fluid 6D was estimated by Sheppard & Harris (1985) to have been close to zero. This is consistent with the fluid's having been derived directly from sea-water. Water may have been derived from dehydration of the basalts, hydrothermally altered by seawater, which presumably make up part of the volcanic cone, but this is not considered to be the principal source of the water (Sheppard & Harris, 1985). Differential loss of H 2 relative to H20 could explain the difference between the 6D of amphibole in the Ascension granite and the accepted primary magmatic values. However, under the prevailing fO2 conditions of the crystallizing magma, on the HM side of the FMQ buffer (Harris 1983), both species are trace
Conclusions and speculations Fluid inclusion data show that granite xenoliths from Ascension Island, unlike the co-genetic comendite lavas, crystallized from a watersaturated magma. Hydrogen isotope data suggest that sea-water was the probable source of the fluid rather than the concentration of primary magmatic water after high degrees of crystallization. Contamination of the granite magmas by sea-water could explain the anomalously high STSr/86Sr ratios reported by Harris et al. (1982a, 1983) for these rocks. Mechanisms by which sea-water addition took place are not well constrained by the data. Any magma chamber in which fractionation took place was probably sited within the volcanic cone and water may have been derived from hydrated wall-rock basalt. An alternative mechanism is direct addition via a caldera. In both these cases (especially the latter) the tendency would have been to produce explosive pyroclastic eruptions because, unlike dry acid magmas, the d P / d T of wet acid peralkaline magmas is negative (Bailey & Cooper 1978). The relatively high proportion of pyroclastics to lava on Ascension (47% by area (Atkins et al. 1964)) may be the result of seawater addition. It is suggested that two types of eruption occurred on Ascension. One was dry and dominated by lava, and the other was pyroclastic dominated and triggered by sea-water interaction. The granite xenoliths of Ascension are the slowly cooled equivalents of the latter. Field and laboratory study of the pyroclastic deposits of Ascension, which have so far been neglected, are needed to substantiate this model. ACKNOWLEDGMENTS: C. Harris is grateful to the
Natural Environment Research Council and the Royal Society for financial support. Many colleagues in Oxford and Nancy provided technical help and advice. We are particularly grateful to Brian Atkins, David Bell and Pierre Coget for their help and encouragement. Tony Ewart, J. R. Cann, E. Roedder, Stuart Smith and Andy Duncan provided valuable comments on earlier versions of the paper.
272
C. Harris & S. M. F. Sheppard
References ATKINS, F. B., BAKER, P. E., BELL, J. D. & SMITH, D. G. W. 1964. Oxford expedition to Ascension Island, 1964. Nature, Lond. 204, 721-4. BAILEY, D. K. & COOPER, J. P. 1978. Comparison of the crystallization of pantelleritic obsidian under hydrous and anhydrous conditions. Progress in Experimental Petrology, NERC Progress Report Dll, pp. 220-33. BOETTCHER, A. L. & O'NEIL J. R. 1980. Stable isotope, chemical and petrographic studies of high-pressure amphiboles and micas: evidence for mantle metasomatism in the mantle source region of alkali basalts. Am. J. Sci. 280A, 594-621. DALY, R. A. 1925. The geology of Ascension Island. Am. Acad. Arts Sci. Proc. 60, 1-80. DINGWELL, D. B., HARRIS, D. M. & SCARFE, C. M. 1984. The solubility of H20 in melts in the system SiO2-A1203-Na20-K20 at 1 to 2 kbars. J. Geol. 92, 387 95. FLEET, S. G. & CANN, J. R. 1967. Vlasovite: a second occurrence and a trictinic to monoclinic inversion. Mineral. Mag. 36, 233-41. GRAHAM, C. M., HARMON, R. S. & SHEPPARD, S. M. F. 1984. Experimental hydrogen isotope studies: hydrogen isotope exchange between amphibole and water. Am. Mineral. 69, 128-38. HARRIS, C. 1983. The petrology of lavas and associated plutonic inclusions of Ascension Island. J. Petrol. 24, 424-70. 91986. A quantitative study ofmagmatic inclusions in the plutonic ejecta of Ascension Island. J. Petrol., 27, 251-76.
---,
BELL, J. D. & ATKINS, F. B. 1982a. Isotopic composition of lead and strontium in lavas and coarse-grained blocks of Ascension Island. Earth planet Sci. Lett. 60, 79-85. -- & --. 1983. Isotopic composition of lead and strontium in lavas and coarse-grained blocks of Ascension Island--an addendum. Earth planet Sci. Lett. 63, 139-41. , CRESSEY, G., BELL, J. D., ATKINS, F. B. & BESWETHERICK, S. 1982b. An occurrence of rareearth rich eudialyte from Ascension Island. Mineral. Mag. 46, 421-5. ROEDDER, E. & COOMBS, D. S. 1967. Immiscibility in granitic melts indicated by fluid inclusions in ejected granite blocks of Ascension Island. J. Petrol. 8, 417-5 I. SHEPPARD, S. M. F. & HARRIS, C. 1985. Hydrogen and oxygen isotope composition of Ascension Island lavas and granites : variation with crystal fractionation and interaction with sea-water. Contrib. Mineral. Petrol., 91, 74-8 I. TAYLOR, B. E., EICHELBERGER, J. C. & WESTRICH, H. R. 1983. Hydrogen isotopic evidence of rhyolite magma degassing during shallow intrusion and eruption. Nature, Lond. 306, 541-5. VAN ANDEL, T.-J., REA, D. K., VON HERZEN, R. P. & HOSKINS, H. 1973. Ascension Fracture Zone, Ascension Island and the mid-Atlantic Ridge. Geol. Soc. Am. Bull. 84, 1527-46. VAN TASSEL, R. 1952. Daylite: a new potassium zirconium silicate from Ascension Island 9Mineral. Mag. 29, 850-7.
CHRIS HARRIS*, Department of Earth Sciences, Parks Road, Oxford OX1 3PR, U.K., and Centre de Recherches Petrographiques et Geochimiques, BP20, 54501 Vandoeuvre-lesNancy C~dex, France. SIMON M. F. SHEPPARD, Centre de Recherches Petrographiques et Geochimiques, BP20, 54501 Vandoeuvres-les-Nancy C6dex, France. * Present address: Department of Geochemistry, University of Cape Town, Rondebosch 7700, South Africa 9
The Cameroon line, West Africa: a comparison between oceanic and continental alkaline volcanism J. G. Fitton S U M M A RY: The Cameroon line is the only known intra-plate alkaline volcanic province which straddles a continental margin. It consists of a chain of Tertiary to Recent volcanoes stretching from the Atlantic island of Pagalu to the interior of the African continent. It therefore provides a unique area in which to study the differences, if any, between the suboceanic and sub-continental mantle sources for alkali basalt. Although the Cameroon line does not have a graben structure, its origin is closely linked to that of the nearby Benue rift. Geochemical and isotopic data show no significant differences between basaltic rocks in the continental and oceanic sectors. However, the more evolved rocks in the two sectors are quite distinct. The continental magmas evolve towards peralkaline rhyolite, whereas those in the oceanic sector evolve towards phonolite. Progressive crustal contamination of the continental magmas accompanied by crystal fractionation is required to explain this distinction. The striking geochemical similarity between basaltic rocks in the two sectors implies that their parental magmas had very similar mantle sources. This source must lie in the asthenosphere. The old lithosphere mantle beneath Africa will be chemically and isotopically different from the young Atlantic Ocean lithosphere mantle and its involvement would be readily detected in the geochemistry of the basalts. Recent models of intra-plate alkaline magma genesis have stressed the importance of metasomatic enrichment of the mantle in large-ion lithophile elements (LILE) as a precursor to magmatism. Evidence for mantle metasomatism is provided by lithosphere-derived mantle xenoliths. Since the lithosphere is the only place where large domains of enriched mantle are likely to be preserved for long periods, it follows that the Cameroon line magmas could not have had a metasomatically enriched lithosphere source. The Cameroon line alkali basalts are chemically and isotopically similar to most other intra-plate (e.g. ocean island and continental rift) basalts which may also, therefore, have an asthenosphere rather than a metasomatized lithosphere source. The asthenosphere is also the source for mid-ocean ridge basalts (MORB) and must have a bulk composition depleted in LILE. Isotopic differences between MORB and intra-plate basalts require this source to be heterogeneous. The Cameroon line and most other intra-plate alkali basalts can be generated by small-degree melting (less than 1%) of a LILE-depleted MORB mantle source containing LILE-rich streaks.
Introduction The Cameroon line (Fig. 1) is a prominent volcanic lineament, 1600 km long, which comprises both oceanic and continental volcanoes. It includes four large central volcanoes in Cameroon and extensive lava plateaux in Cameroon and northern Nigeria, and extends offshore into the islands and seamounts of the Gulf of Guinea. The volcanic rocks are dominantly alkaline, ranging from transitional basalt to nephelinite and alkali rhyolite to phonolite. Chemical analyses of representative volcanic rocks from the Cameroon line are given in Tables 1-3. General reviews of the geology and physical geography of various parts of the region have been written by G+ze (1943, 1953), Hedberg (1968), MitchellThorn6 (1970), Gouhier et al. (1974) and Dunlop (1983). The occurrence of contemporaneous oceanic and continental alkaline volcanism within a
single province makes the Cameroon line a unique area in which to compare oceanic and continental intra-plate volcanism. The results of a project designed to assess the differences between the sub-oceanic and sub-continental mantle sources for alkaline magmas, the melting processes required to produce them and the effects of crustal contamination upon their subsequent evolution are summarized in this paper.
Age of the Cameroon line The earliest igneous activity on the Cameroon line is represented by a number of small (up to about 10 km) intrusive ring complexes composed of granite and syenite with less abundant gabbro and occasional remnants oftrachyte and rhyolite. These extend along virtually the whole of the continental sector (Fig. 2) and give ages ranging
From: FIT'rON, J. G. & UPTON, B. G. J. (eds), 1987, Alkaline Igneous Rocks, Geological Society Special Publication No. 30, pp. 273-291.
273
J. G. Fitton
274
0
Kilometres
~'~i~l~Biu Plat eau !::;;.,;,~. Mandara
500
;::i;;~
M,s ,0
'
Ngaoundere
--ipJ
........ }enue {rough Bambouto ~1~'Manengouba Mt Cameroon
.....o e , / i
I
i%/o"
FIG. 1. Geological map of part of W Africa showing the outcrop of the Cameroon line volcanic rocks (black) and the Cretaceous sedimentary rocks of the Benue trough (stippled). The boundary between continental and oceanic crust (bold broken line) is taken from Emery & Uchupi (1984), chart XI. The Ngaound6r6 fault system (between Bambouto and the Ngaound6r6 Plateau) and oceanic transform faults are from Gazel (1956) and Sibuet & Mascle (1978) respectively.
from 66 to 30 Ma (Cantagrel et al. 1978; Lasserre 1978 ; Jacquemin et al. 1982). Seismic profiles across the oceanic sector (Grunau et al. 1975) show disturbance of sediments of inferred Upper Cretaceous age. Gorini & Bryan (1976) have suggested that uplift of the ocean floor along the Cameroon line has formed an effective barrier between the Niger Delta and the Douala Basin since the Cretaceous. It seems likely, therefore, that the earliest volcanism in the oceanic and continental sectors was essentially contemporaneous. The oldest dated extrusive rocks are trachyte lavas and rhyolitic ignimbrite flows associated with a small granophyre and microsyenite ring complex at Kirawa on the Nigeria-Cameroon border (Fig. 2). The trachytes and rhyolites give Rb-Sr whole-rock ages of 51.2 Ma and 45.5 Ma respectively (Okeke 1980; Dunlop 1983). The younger extrusive centres (Fig. 1) give K - A r ages ranging from 35 Ma (Mandara Mountains) to Recent with the oldest dated oceanic rocks (Principe) at 31 Ma (Dunlop & Fitton 1979; Dunlop 1983; Fitton & Dunlop 1985, and references cited therein), The available K - A r data on the extrusive centres are summarized in Fig. 3. Signs of geologically recent volcanic activity can be found in most of the volcanic --
I
. L - - dl
Continental sector
I
IB,u i
,
Mandara nmlii
H i
I Ngaoundere
* '
1
m 9
~
i
i
Kirawa (46) Golda Zuelva (66) ~0~
Jos Plateau
~ {jL,/~
/
i Bambouto
J . i l Manengouba
' 9
....
r_~
* oko
1 Mt. Cameroon
~o,,(3~.)- ~,0go, (43)
'*1 Etinde
A: i
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Oceanic sector
(Namboe ,.~ ~(63) . ] ~ ~ , ~_ r_- Mayo Dar,e (63) ~ |,q~'~v -
, . . ~ ( - b ~ B a n a (30) ~ Nlonako(44) ~ , ~ K Oup6 (46) If~ ~ ?
}
ld0o
Bioko -X- 9
I I
Kiiometres i l ~
.._
Principe
m
I
500
9
115~
FIG. 2. The distribution of early Tertiary intrusive complexes (&) with their ages (in Ma) given in parentheses. Data from Cantagrel et al. (1978), Lasserre (1978) and Jacquemin et al. (1982). The Benue trough and the outcrop of the late Tertiary to Quaternary extrusive rocks of the Cameroon line are shown in outline.
S&o Tome
*__
in
n
Pagalu ")(" a n n e
0
--
10
i,
20 Ma
30
40
FIG. 3. Summary of available K-Ar age data for the Cameroon line volcanic rocks. Asterisks indicate the presence of morphologically recent cinder cones. (After Fitton & Dunlop 1985.)
The Cameroon line
275
TABLE 1. Chemical analyses o f representative Cameroon line basic volcanic rocks Continental sector N31
C147 C174
C89
Cl12
Oceanic sector C65
C5
FP44
P15
P31
P18
ST72 ST96 ST19AN15
Major elements (wt.~) SiO2 A1203 Fe203(t) MgO CaO Na20 K20 TiO2 MnO P2Os LOI
49.18 43.60 45.47 46.62 43.85 13.69 14.50 15.35 16.86 13.65 10.94 12.66 11.55 12.50 12.52 9.78 7.42 5.47 5.56 9.21 7.94 9.78 9.34 8.50 10.55 3.42 3.32 5.08 3.39 3.47 1.57 1.54 2.86 1.37 1.47 2.17 3 . 9 1 2.58 3.18 3.10 0.15 0.16 0.26 0.19 0.19 0.47 0.74 0.91 0.86 1 . 0 3 0.14 1.79 0.93 1.10 0.19
50.14 47.30 14.21 17.20 12.40 10.32 7.94 4.25 7.60 8.87 3.48 5.26 1.35 2.09 2.24 2.93 0.18 0.20 0.40 0.92 0.08 - 0 . 1 2
Total
99.45 99.42 99.79100.13 99.21 100.01 99.21
44.40 46.77 47.71 37.01 43.80 10.80 12.45 14.07 10.88 14.06 13.33 12.66 11.90 14.28 12.46 10.80 9.22 8.15 12.24 8 . 3 1 12.58 9.88 8.00 13.19 10.62 2.48 2.55 3.05 3.10 3 . 6 1 1.19 0.48 1.67 1 . 5 5 1 . 4 1 3.51 2.38 2.94 4.15 3.17 0.19 0.16 0.16 0.21 0.22 0.63 0.41 0.67 1.42 1 . 0 3 0.00 3 . 1 1 1.60 1.62 0.54
49.25 44.25 41.67 17.43 15.30 10.62 9.50 12.95 13.76 4.01 5.63 15.52 7.08 9.50 9.70 3.74 4.76 2.36 2.72 2.15 0.85 2.51 3 . 8 1 3.12 0.19 0.21 0.19 0.99 1.02 0.82 2.22 0.04 0.62
99.90100.07 99.92 99.95 99.23 99.64 99.61 99.23
Trace elements (ppm) Ni Cr V Sc Cu Zn Sr Rb Zr Nb Ba Th La Ce Nd Y
282 376 295 24 44 96 632 35 239 55 527 4 37 77 33 27
88 129 503 19 43 90 1110 34 250 62 485 0 45 87 44 25
54 116 185 15 35 122 1266 74 476 114 1000 12 109 180 65 34
29 148 46 459 182 232 19 23 22 43 98 96 1 0 0 5 1242 22 32 254 190 46 56 517 874 2 1 46 43 94 89 48 46 28 26
197 275 193 23 57 122 429 31 299 47 458 4 41 64 33 29
18 25 177 10 30 103 1202 51 461 118 650 4 97 199 81 38
8.07 29.70 19.36 . -12.95 12.51 7.12 . 5.05 4.29 0.95
12.52 23.30 17.37 . 11.82 17.14 -5.77 . 4.22 5.64 2.21
215 521 294 30 72 107 823 33 336 65 562 7 58 121 59 32
236 351 215 24 60 113 444 13 131 25 174 2 16 40 22 23
167 231 231 21 42 112 625 38 327 49 439 5 43 88 41 29
193 338 307 26 56 108 1297 40 337 118 987 10 103 203 85 35
171 381 237 18 48 114 1130 47 366 92 718 17 79 157 66 33
7 5 124 6 9 128 1056 62 527 105 854 12 89 180 76 37
43 33 234 16 41 133 1113 60 472 99 728 11 83 163 72 36
569 588 268 25 43 118 779 37 253 57 516 6 58 90 43 30
10.13 -26.43 -20.31 11.53 7.40 -14.64 12.77 24.95 8.66 -9.42 19.33 4.63 4.92 5.94 5.73 8.12 1.62 3.46
8.52 15.05 18.39 . 8.76 22.92 -12.58 . 5.14 6.15 2.49
16.59 32.66 23.51 . . -4.94 4.40 6.59 . . 3.96 4.92 2.43
12.86 14.33 14.21 . 14.35 21.67 -7.51 . 5.30 7.33 2.44
5.12 10.74 16.27
CIP W norms (Fez03/(FeO + Fe203) = 0.3) or
ab an
lc ne
di
hy ol cs mt
il ap
9.41 29.38 17.62 . -15.38 2.45 15.97 . 4.48 4.18 1.12
9.42 18.14 20.77 . 5.91 19.50 -11.49 . 5.28 7.67 1.82
17.25 8.23 10.38 29.18 10.81 27.26 . . . 18.13 ~ 24.68 8 . 1 1 -4.37 6.82 9.49 . . . 4.75 5.14 4.99 6.15 2.19 2.07
8.82 14.09 17.66 . 8.56 23.07 -14.19 . 5.14 6.00 2.48
7.08 2.96 10.67 22.47 14.99 21.95 . . 5.70 -35.02 21.11 -14.83 12.85 5.65 . . 5.43 5.32 6.74 4.71 1.51 1.00
5.25 21.87 -27.00 5.68 6.07 1.99
All the analyses in Tables 1, 2 and 3 were carried out in Edinburgh using X-ray fluorescence techniques (see Fitton & Dunlop 1985). Samples were dried at 110~ before analysis. Fe203(t ) is total Fe expressed as Fe203 . LOI, loss on ignition at 1100~ N31, ol-phyric transitional basalt (5 km N N W of Tila Lake, Biu Plateau, Nigeria); C147, aphyric alkali basalt (hillside to S of Kila, Mandara Mountains, Cameroon)" C174, basanite (waterfall on River Vina, 13 km S of Ngaound6r6, Cameroon); C89, plag(-ol-augite)-phyric transitional basalt (4 km SW of Kumbo, E of Mount Oku, Cameroon); C112, ol-augite-phyric basanite (2 km WSW of Lekwu6 Lelo6, near summit of Bambouto, Cameroon); C65, plag-(-ol-augite)-phyric transitional basalt (S wall of Lac de la Femme, Eboga caldera (Manengouba), Cameroon); C5, plag-augite(-ol)-phyric tephrite (summit of Mount Cameroon) (see Fitton et al. (1983) for more analyses of Mount Cameroon lavas); FP44, ankaramite (coast W of Pico Sta Isobel, Bioko (see Piper & Richardson 1972)); P15, olivine tholeiite from blocks in hyaloclastite breccia (River Cambungo. SE side of Principe); P31, ol-augite-phyric transitional basalt (N-S dyke, Praia de Sta. Rita, N coast of Principe); P 18, olivine nephelinite (1 km SE of S. Antonio, Principe) (total includes 0.30~ F); ST72, ol-hbld-augite-plag-phyric basanite (Maria Luisa, N W side of S~o Tom6)" ST96, hbld-phyric hawaiite (NW-SE dyke, Praia Capitango, S coast of S~o Tom6); ST19, aphyric basanite (recent lava flow from Morro Corregado, N coast of S~o Tom6); A N 15, picritic basalt (N side of Pagalu) (see Piper & Richardson 1972).
276
J. G. Fitton
TABLE 2. Chemical analyses o f representative Cameroon line intermediate and evolved volcanic rocks Continental sector C144 C155
C85
Oceanic sector
C87 C104 Cl16
C63
Pll
P12
P16
ST10 ST57 ST73 ST84STl10
Major elements (wt.~) SiO2 A1203 Fe203(t) MgO CaO Na20 K20 TiO2 MnO P205 LOI
74.08 55.75 72.70 71.13 68.56 58.19 53.99 11.40 22.05 10.74 12.56 13.08 19.50 17.03 3.49 2.29 5.69 3 . 3 1 5.74 4.39 8.95 0.17 0.09 0.04 0.04 0.05 0.80 3.30 0.09 1.08 0.12 0.02 0.44 2.88 5.39 5.26 10.93 4.72 5.68 5.87 7.77 5.29 4.05 5.54 4.57 5.26 4.81 4.26 3.02 0.12 0.35 0.37 0.68 0.35 0.91 1.98 0.10 0.19 0.09 0.03 0.13 0.21 0.17 0.01 0.08 0.00 0.09 0.01 0.20 0.65 0.46 1.09 0.22 0.66 0.60 0.84 0.22
54.90 53.98 56.28 57.39 66.69 48.54 61.66 53.64 22.32 19.54 21.47 21.70 18.05 17.42 18.63 18.83 3.13 4.80 2.38 2.20 2.16 9.93 3 . 4 1 6.41 0.30 1.42 0.07 0 . 1 1 0.08 3.51 0.16 1.76 2.00 5.11 2.00 1.24 0.58 8.16 0.94 4.96 9.53 5.39 8.62 9.43 6.29 5.38 7.93 7.28 6.02 4.41 4.86 6.02 4.37 2.45 5.46 3.65 0.48 1.08 0.16 0.24 0.30 2.65 0.26 1.63 0.12 0.13 0.18 0.15 0.11 0.22 0.24 0.16 0.09 0.24 0.02 0.04 0.07 0.93 0.03 0.44 0.85 4.08 3.75 1.19 1.15 0.34 0.99 0.55
Total
99.24 99.43 99.26 99.46 99.64 99.93 99.98
99.75 100.17 99.80 99.72 99.86 99.53 99.70 99.31
Trace elements (ppm) Ni Cr V Sc Cu Zn Sr Rb Zr Nb Ba Th La Ce Nd Y
1 3 7 0 0 307 4 402 1938 576 9 81 41 153 29 78
2 1 1 3 5 4 5 3 14 0 0 0 0 0 2 0 0 0 0 0 131 141 261 300 466 4 15 6 212 172 207 133 1 2 9 1 1 4 0 5 2 0 7 1 1372 167 255 379 251 406 4 80 31 31 38 39 30 60 20 325 249 105 18 659 398 27 19 267 188 22 26 165 112
2 3 2 0 0 75 1549 101 558 152 2501 6 116 183 73 49
24 22 91 9 14 106 856 87 643 107 701 13 76 148 58 32
3 5 45 0 0 79 959 133 431 79 1087 13 65 93 23 15
11 14 56 2 4 87 1115 180 719 119 758 26 68 111 36 20
. . 36.08 21.47 0.34 32.65 . . 6.28 0.78 . -1.28 0.92 0.21
. . 27.19 35.23 16.79 6.73 . . 6.60 . . 2.69 2.02 2.13 0.60
2 4 5 0 0 185 11 314 1745 136 11 76 98 110 17 19
2 4 15 0 0 114 110 223 965 82 154 34 76 95 18 13
2 5 2 0 0 79 231 124 515 141 1361 20 148 254 78 38
29.97 36.19 37.73 29.02 5.74 -20.78 27.26 1.58 . . 3.98 5.26 . . . . 0.43 0.02 1.00 0.11 0.31 0.46 0.04 0.10
11.58 2.12 26.20 54.00 2.45 --. -. 2.03 -0.88 0.58 0.17
19 23 173 9 21 120 1281 64 599 133 870 18 105 197 79 38
1 4 0 0 0 174 19 200 1414 264 70 39 199 338 88 43
6 5 79 4 7 117 1128 114 751 116 845 17 79 148 52 27
s
C IP W norms ( Fe 20 3/ ( FeO + Fe 20 3) = 0.3)
a
C or
ab on
ne ac ns
di WO
hy ol mt
il ap
29.92 . 24.29 36.62 . -2.85 1.24 0.35 . 4.46 -. 0.23 0.03
-29.11 . . 33.33 27.40 25.76 30.18 . . . 31.61 -1 . 8 8 4.64 1.86 1.17 4.31 0.51 . . . -6.26 0.38 -. . . 0.68 0.71 0.20 0.01
23.18 .
16.18 . . . . . . 31.57 28.79 25.47 18.02 35.83 41.10 45.74 43.21 . 5.83 13.90 --11.27 1.05 2.70 4.69 . . . 2.29 0.91 . . . -1.92 6.13 7.18 . . . 3.04 5.72 . . . --1.56 7.65 . 1.79 3.64 1 . 3 1 0.68 1.74 3.80 0.21 0.01 0.48 1.55 .
. .
------14.72 32.76 21.97 26.48 50.12 34.35 16.40 -8.03 10.66 8.86 15.33 -1.47 -. . 15.08 3.98 11.57 . . ---5.24 1.59 1.91 4.07 0.66 2.63 5.10 0.51 3.14 2.25 0.07 1.07
LOI, loss on ignition. C144, rhyolite (Aiguille Mchirgui, 2 km N of Mogode, Mandara Mountains, Cameroon); C155, phonolite (plug near Beka, 20 km N of Ngaound6r6, Cameroon); C85, rhyolite (flow) (Sabga Pass, 16 km E N E of Bamenda; SW of Mount Oku, Cameroon); C87, rhyolitic welded tuff (6 km SW of Jakiri; SE of Mount Oku, Cameroon); C104, trachyte (Pinyin, N slopes of Bambouto, Cameroon); C 116, trachyphonolite (Bambouto summit, Cameroon); C63, benmoreite (NW rim of Eboga caldera (Manengouba), Cameroon); P11, phonolite (large block in River Fria, SE side of Principe); P12, tristanite ( a s Dais Irmfios (plug), SE side of Principe); P16, trachyphonolite (large block in River Cambungo, SE side of Principe); ST10, phonolite (C~.o Grande (plug), S side of S~o Tom~); ST57, quartz trachyte (NE peak of Ilh6u das Cabras, off N E coast of S~,o Tom~); ST73, hbld-plagphyric mugearite (Maria Luisa, N W side of Sao Tom~); ST84, trachyte (S side of Mizambu (plug), E side of S~o Tomb); ST 110, phonolitic tephrite (River Manuel Jorge, 0.5 km N W of Sta. Luisa, E side of S~o Tomb).
The Cameroon line TABLE 3. Chemical analyses o f representative nephelinite lavas f r o m Etinde, Cameroon C150
C24
Major elements (wt.%) SiO2 39.60 40.04 A1203 12.75 14.56 Fe203(t) 13.91 12.34 MgO 8.03 5.28 CaO 14.98 14.00 Na20 2.68 3.47 K20 1.66 2 . 6 1 TiO2 4.49 4.06 MnO 0.20 0.30 P205 0.67 1 . 1 8 SrO 0.12 0.28 BaO 0.09 0.12 SO3 0.00 0.24 C1 0.04 0.10 LOI 1.00 1 . 3 1
Total
C152 C131 C154 39.99 42.65 46.25 17.54 17.39 19.46 9.69 9.94 6.64 3.74 2.70 1.16 10.87 8.79 6.15 6.94 6.41 7.80 3.87 4.48 6.22 2.95 2.54 1.06 0.28 0.38 0.36 0.89 0.68 0.16 0.28 0.33 0.78 0.11 0.14 0.38 1.94 1.41 0.27 0.37 0.49 0.47 0.79 1.95 3.06
100.21 99.87 100.17 100.18 100.11
Tracee&ments(ppm) Ni 46 12 10 6 2 Cr 126 8 5 6 3 V 456 448 342 416 241 Sc 45 15 4 2 0 Cu 113 64 84 38 5 Zn 93 127 134 194 209 Sr 1054 2401 2395 2825 6553 Rb 89 87 114 349 201 Zr 399 656 626 888 914 Nb 96 223 281 320 295 Ba 788 1067 1006 1254 3424 Th 8 25 21 15 4 La 97 238 225 170 131 Ce 195 489 425 272 168 Nd 84 180 142 86 41 Y 28 47 44 52 37 CIP W norms (Fee03/(FeO + Fe 20 3) = 0.3) or ---18.90 23.20 an 18.36 18.12 13.68 12.61 2.60 lc 7.84 12.40 18.20 6.51 11.83 ne 12.36 15.02 23.79 23.03 34.29 hl 0.07 0.17 0.62 0.83 0.81 th -0.44 3.49 2.57 0.50 di 35.37 36.17 25.23 23.75 16.90 wo . . . . 4.59 ol 7.13 1 . 3 8 1.64 1.05 - cs 2.85 0 . 4 1 1.55 - -mt 5.72 5 . 1 1 3.97 4.12 2.79 il 8.69 7 . 9 1 5.69 4.96 2.10 ap 1.62 2.87 2.14 1 . 6 6 0.39
LOI, loss on ignition; C 150, olivine melanephelinite; C24, nephelinite; C152 and C131, haiiyne nephelinite; C154, leucocratic nosean leucite nephelinite. All the samples were collected from loose blocks in stream beds : C24 from Bonenza, and the others from Batoke.
277
centres (Fig. 3). M o u n t C a m e r o o n is still active and last erupted in 1982 (D&uelle et al. 1983; Fitton et al. 1983). It is clear that volcanic activity on both the oceanic and continental sectors of the C a m e r o o n line has been more or less continuous since the end of the Cretaceous.
Geology and petrology Continental sector Intrusive complexes
Relatively little is k n o w n about the m i n e r a l o g y and geochemistry of these complexes. T h a t w h i c h is k n o w n suggests that they are broadly similar to the ring complexes of the Jos Plateau in N i g e r i a (Bowden et al. 1987). Some of the granite bodies in the M a y o Darl~ area contain traces of cassiterite ( C h a p u t et al. 1954; Gazel et al. 1963). Only the M b o u t o u complex (Fig. 2) has been described in detail. Parsons et al. (1986) have published an account of the m i n e r a l o g y and crystallization history of this complex a n d stress the very mildly alkaline nature of its parental m a g m a . J a c q u e m i n et al. (1982) have carried out a g e o c h e m i c a l and isotopic study of the G o l d a Z u e l v a and M b o u t o u complexes and conclude that the m a g m a s , although initially m a n t l e derived, have interacted with crustal rocks as they crystallized. T h e intrusive complexes were not included in the present study because they are only exposed in the continental sector. F u r t h e r m o r e , the effects of crystal a c c u m u l a t i o n and late-stage alteration processes limit the usefulness of coarse-grained rocks in g e o c h e m i c a l studies. T h e following discussion will therefore be confined entirely to the younger extrusive centres shown on Fig. 1. Biu Plateau
The small t o w n of Biu in N E N i g e r i a is built on a plateau c o m p o s e d of basaltic lava flows. These reach a m a x i m u m thickness of 250 m (J. W. du Preez, cited by Carter et al. 1963) and range in composition from basanite to h y - n o r m a t i v e transitional basalt. The most recent volcanism in the area p r o d u c e d a large n u m b e r of c i n d e r cones aligned in a N N W - S S E direction. These usually have well-defined craters w h i c h have often been b r e a c h e d by small lava flows. K - A r ages on the plateau basalts range from 5.3 M a for the basal lavas to about 0.8 M a for a flow high in the succession ( G r a n t et al. 1972; Fitton & D u n l o p 1985). The lava flows and cinder cones often contain a b u n d a n t peridotite xenoliths and clinopyroxene and anorthoclase megacrysts. A large crater to
278
J. G. F i t t o n
the west of the village of Miringa (14 km N N W of Biu) is particularly noteworthy for its abundant and extensive inclusion suite. These include peridotite xenoliths up to 0.5 m in diameter and large megacrysts of clinopyroxene, garnet, anorthoclase, ilmenite and amphibole (Wright 1970; Frisch & Wright 1971). Mandara Mountains
The Mandara Mountains lie along the Cameroon-Nigeria border and are composed of granite and gneiss of Pan-African age. In the area around the small settlements of Roumsiki and Mogode the basement rocks have been intruded by numerous plugs of peralkaline trachyte and rhyolite which form spectacular spines. Occasional remnants of alkali basalt lava flows are also found in this area. K - A r dating (Fitton & Dunlop 1985) gives ages of 30 and 33 Ma for two samples of basalt and 35 Ma for one of the trachyte plugs. Ngaoundbr~ Plateau
The town of Ngaound6rb in northern Cameroon lies on the western edge of an extensive basalt plateau. The earliest lavas were erupted into two broad valleys eroded into the Precambrian basement by the headwaters of the Rivers Bini and Vina. Continued volcanism built a plateau of alkali basalt across the area and culminated in the building of a central volcano, Nganha. This volcano is composed of basanite capped by trachyte and phonolite flows which have yielded K - A r ages of 10-7 Ma (Gouhier et al, 1974). Trachyte flows also occur as small outliers over the whole plateau as do trachyte and phonolite plugs. The most recent volcanism is represented by cinder cones aligned in a W N W - E S E direction. These have sometimes produced small lava flows. All the basic lavas in the area are ne normative and range from alkali basalt to basanite. Some of the flows contain peridotite xenoliths and clinopyroxene, anorthoclase and rare zircon megacrysts.
although basic lava flows also occur higher in the succession. K - A r ages obtained from Oku lavas range from 23 to 17 Ma (Gouhier et al. 1974; Fitton & Dunlop 1985). The basic lavas range in composition from slightly ne normative to slightly hy normative. The volcanic rocks as a whole form a strongly bimodal suite. Numerous cinder cones (some yielding peridotite xenoliths) and explosion craters on and around Mount Oku provide evidence for recent volcanic activity in the area although very few of the vents produced lava flows. The craters reach 2 km in diameter and many form crater lakes. Two of these (Lake Monoun to the south and Lake Nyos to the N W of Oku) recently released large volumes of carbon dioxide with catastrophic results (SEAN 1985, 1986). The cause has not yet been established. Bambouto
Unlike Oku, Bambouto still retains some of its original volcanic form despite extensive erosion. The highest peaks (up to 2740 m) lie around a calder rim 10 km in diameter. Tchoua (1972) identified two episodes of caldera collapse. As with Oku, the lava succession comprises a strongly bimodal basalt-trachyte suite. Basalt lavas dominate the basal part of the succession but are also found on the upper slopes of the volcano. The Bambouto lavas are generally more undersaturated than the Oku suite. All the basic lava samples analysed are slightly ne normative and the salic rocks range from quartz trachyte to trachyphonolite. Rhyolites have not been found by the present author although Tchoua (1973) has reported extensive ignimbrite sheets on and around Bambouto. K - A r dating of lavas from Bambouto gives ages ranging from 23 to 14 Ma (Gouhier et al. 1974; Fitton & Dunlop 1985). Gouhier et al. also report an age of 5.8 Ma for a basalt sample collected about 20 km N of the caldera rim. There is no evidence for recent volcanic activity on Bambouto.
Oku
Manengouba
Mount Oku (3011 m) is a large deeply dissected volcanic massif composed largely of rhyolite and quartz trachyte flows, domes and plugs. Rhyolitic ignimbrite sheets are common in the area and extend as far south as the town of Bamenda (midway between Oku and Bambouto) which sits at the base of a prominent escarpment of welded tuff. The lower part of the volcanic succession is dominated by flows of basalt and hawaiite,
Manengouba is a well-preserved central volcano whose summit region is occupied by two concentric calderas. A poorly defined outer caldera (Elengoum; 6 km in diameter) encloses the younger perfectly preserved Eboga caldera (3 km in diameter). The highest point on the mountain (2411 m) lies on the Elengoum rim. This outer rim was breached on its eastern side during the building of the Eboga volcano and allowed large
T h e C a m e r o o n line
volumes of lava to flow over the north-eastern flanks of the older volcano. The oldest Manengouba lava flow so far dated gives a K - A r age of 1.55 Ma (Gouhier et al. 1974). A recent detailed K - A r and field study of the volcano (C. A. Hirst, unpublished data) has confirmed this date and established a chronology for the evolution of the volcano. The collapse of the Elengoum caldera probably occurred between 0.8 and 0.6 Ma ago. The Eboga lavas give ages ranging from 0.56 to 0.26 Ma, and the collapse of the Eboga caldera is thought to have occurred at about 0.25 Ma. An isolated trachyte spine intruding the eastern rim of the Elengoum caldera gave a K - A r age of 0.26 Ma. More recent volcanism has produced a S W - N E line of crater lakes and cinder cones across the floor of the Eboga caldera. The lavas range in composition from ne- and by-normative basalts to trachyte, quartz trachyte and rare rhyolite. Intermediate lavas are common. In striking contrast with the lavas of Oku and Bambouto, the Manengouba lavas form a complete compositional continuum. Mount Cameroon
At 4095 m, Mount Cameroon is by far the highest mountain in West Africa and one of Africa's largest volcanoes. The base of the lava pile is below sea level. Mount Cameroon is the only currently active centre of the Cameroon line and has had five recorded eruptions this century (Fitton et al. 1983, and references cited therein). It is a composite volcano built of alkali basalt to basanite lava flows interbedded with small amounts of pyroclastic material. Small cinder cones are abundant in several areas on the flanks of the volcano and on the surrounding lowlands. These cones are often aligned along fissures running SW-NE, parallel to the long axis of Mount Cameroon. The Plain of Tombel, between Mount Cameroon and Manengouba, is likewise dotted with numerous cinder cones and explosion craters, some of which are occupied by lakes. The largest of these, Lake Barombi Mbo near the town of Kumbo, is surrounded by a tuff ring containing abundant peridotite xenoliths. The Mount Cameroon lavas are all ne normative and range from almost aphyric to strongly porphyritic (olivine _+plagioclase). Lavas more evolved than hawaiite have not been reported. The few K - A r dates available (Hedberg 1968) suggest that all the exposed lavas were erupted in the last million years although there is field evidence that the volcanic edifice was built upon a basement of older lava flows. These are best exposed along the coast to the SE of Limbe (formerly Victoria) but have not yet been dated.
279
Etinde
Mount Cameroon is not a single volcano but has a large subsidiary peak, Etinde, on its SW flank. Etinde rises from the rain forest a few kilometres from the coast and is densely forested up to its summit (1713 m). Exposures of lava are rare and deeply weathered but the numerous streams contain blocks of fresh volcanic rock unlike anything found on the main volcanic massif. These provide evidence that Etinde is composed entirely of nephelinite lavas. The lavas were first described by Esch (1901). Attempts to date the Etinde lavas have not yet produced consistent results. A sample of nosean leucite nephelinite gave a K - A r age of about 0.1 Ma (Fitton & Dunlop 1985). R. M. MacIntyre (unpublished data) has obtained K - A r ages of 1.1,0.6 and 6.3 Ma from, respectively, a melanephelinite whole-rock sample and nepheline and hafiyne separated from two different nephelinite samples. It is possible that the apparent age of the hatiyne separate is anomalous and is due to excess radiogenic argon. The lavas appear to form a consanguineous series, and it would therefore be unwise to postulate an eruptive history of over 6 Ma on the basis of this one determination. It seems likely that the lavas are less than a million years old and possibly contemporaneous with some of the Mount Cameroon lavas. More work is needed to establish the relative ages of the two volcanic centres. The lavas form a compositional continuum from olivine melanephelinite through nephelinite and haiiyne-nephelinite to a leucocratic nosean leucite-nephelinite (Table 3). Feldspar is absent throughout the whole series and melilite is observed in a few samples. The observed chemical variation among the lavas can be explained by crystal fractionation dominated by aluminous clinopyroxene. This gave rise to residual liquids close in composition to the nepheline-leucitediopside cotectic in the system nepheline-sanidine-diopside (Fig. 4). These highly evolved compositions contain very high concentrations of some incompatible elements. The nosean leucite-nephelinite lavas, for example, contain around 7000 ppm Sr which is present as an unusual strontian melilite (Fitton & Hughes 1981). Oceanic sector Bioko (formerly Fernando P6o)
Despite being only 50 km from the mainland, Bioko is, geologically, the least well known of the Gulf of Guinea islands. From the time the island gained independence from Spain in 1968 until the coup of 1979, the brutal regime of President
J. G. Fitton
280
was formed in the Brunhes epoch and is thus younger than 0.7 Ma. Numerous cinder cones occur on all three volcanic centres although there have been no well-documented reports of eruptions. What may have been a small eruption on the east side of the island was reported by natives around the beginning of this century (Hedberg 1968).
DIOPSIDE
Principe
N E PHE L INE
Weight
-
SANIDINE
FIG. 4. Normative compositions of three leucocratic nosean leucite--nephelinite samples from Etinde projected into the system nepheline-sanidine-diopside (Platt & Edgar 1972). The norms were calculated using Fe_,O3/(FeO + Fe:O0 values of 0.3, 0.5 and 0.7. Chemical analysis ofC154 gives a value of 0.659 for this ratio (Fitton & Hughes 1981). Macias Nguema ensured complete inaccessibility. The current social and economic state of the island has been described by Winchester (1984) who, 4 years after the execution of the President, claimed that Bioko was 'the nastiest place on earth'. Consequently our knowledge of the geology and petrology of the island is based on preindependence field work and rock collections. The most recent geological expedition to the island was undertaken as part of a palaeomagnetic study of the Cameroon line (Piper & Richardson 1972). The cores they collected were used in the present study. The earlier literature on the geology of Bioko and the other Gulf of Guinea islands has been reviewed by MitchellThom~ (1970). Bioko is a youthful volcanic island with three eruptive centres. Pico Santa Isobel, the highest point (2972 m), forms the northern part of the island and the other two volcanoes, San Carlos and Pico Biao, form the southern part. The lavas range from basanite to hy-normative basalt. Unlike the other Gulf of Guinea islands, more evolved rocks have not been reported from Bioko. A single K-Ar age of 1.1 Ma was obtained by Hedberg (1968) from a basanite sample collected near the town of Santa Isabel on the northern coast. Nearly all the palaeomagnetic sites investigated by Piper & Richardson (1972) are normally magnetized which led these workers to conclude that virtually the whole volcanic edifice
Principe is a small and deeply eroded island surrounded by a broad submarine shelf. The prevailing SE trade winds in the Gulf of Guinea veer SW as they approach the mainland. Therefore the south-western parts of the islands receive the highest rainfall and consequently suffer the most erosion. The structure, volcanic stratigraphy and petrology of Principe have been described by Fitton & Hughes (1977), and a K-Ar chronology has been established by Dunlop & Fitton (1979) with an additional determination by Hedberg (1968). The oldest exposed rocks are hyaloclastite breccias which crop out in river valleys in the southeastern part of the island. These represent a submarine phase in the evolution of the island and contain blocks of fresh olivine tholeiite dated at 31 Ma. The oldest sub-aerial lavas (the Older Lava Series) are exposed mostly in the north and range in composition from transitional basalt to hawaiite. They give K-Ar ages ranging from 24 to 19 Ma (Hedberg 1968; Dunlop & Fitton 1979) and are intruded by numerous dykes and small intrusions compositionally similar to the lavas. The Older Lava Series is unconformably overlain by nephelinite and basanite lava flows which have yielded K-Ar ages of 5.6 and 3.5 Ma. This Younger Lava Series is overlain by phonolite lavas in the northern part of the island. The high ground to the south (up to 948 m) is composed of a chaotic assemblage of lava flows and plugs ranging in composition from tristanite to trachyte and phonolite. Four K-Ar determinations on samples collected from these plugs give ages ranging from 7 to 5.3 Ma. There are no signs of recent volcanic activity on Principe. S~o Tom~ S~.o Tom6 is a roughly conical volcanic island rising from about 3000 m below sea level to a height of 2024 m. The ground rises gently in the drier northern and eastern parts of the island but is much more rugged in the S and W. Phonolite and trachyte spines rising vertically out of the rain forest form a spectacular feature of the southern part of the island where rainfall and erosion rates are highest.
The Cameroon line The oldest rocks on the island are conglomerates, sandstones and shales of the Ubabudo Formation (Hedberg 1968). These crop out in stream sections in a small area in the eastern part of the island and are accompanied by oil seepages. They are almost non-fossiliferous and cannot be 15-o o O o00 9
Ca 9 I0~-
5
3.5 P205
o o
3"0
. 9
