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Startseite Hydrothermal Mineral Deposits: Principles and Fundamental Concepts for the Exploration Geologist
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A photograph made from the submersible ALVIN in May 1986 during the first dives to observe black smoker geysers that were initially discovered in the Atlantic Ocean by the NOAA VENTS Program in 1985. The photograph shows vent shrimp (individual length 2 to 5 centimeters) of a new variety swarming over a chimney (heigh about 2 meters) built of polymetallic sulfide minerals through which a black smoker geyser is venting. The scene occurs at a water depth of 3,630 meters (12,000 feet) about 3,300 kilometers (1,800 miles) due east of Miami, Florida at the Mid-Atlantic Ridge, a submerged volcanic mountain range that extends along the center of the Atlantic Ocean. The photograph was made by a research team of scientists from the National Oceanic and Atmospheric Administration (NOAA), the Woods Hole Oceanographic Institution (WHOI), and the Massachusetts Institute of Technology (MIT). This photograph, and its caption, were kindly supplied by Dr. P.. Rona, of the Oceanic and Atmosphere Administration (USA), and published in EOS. 18th of November issue (vol. 67, No. 46). Franco Pirajno Hydrothermal Mineral Deposits Principles and Fundamental Concepts for the Exploration Geologist With 288 Figures Springer-Verlag Berlin Heidelberg New York London Paris Tokyo Hong Kong Barcelona Budapest Professor Dr. FRANCO PIRAJNO Rhodes University Department of Geology Mineral Exploration P.O. Box 94 Grahmstown 6140 South Africa Present address: 2/34 Kennedy Street Maylands 6051 Western Australia e-ISBN-13:978-3-642-75671-9 ISBN-13:978-3-642-75673-3 001: 10.1007/978-3-642-75671-9 Library of Congress Cataloging-in-Publication Data. Pirajno, Franco, 1939- Hydrothermal mineral deposits: principles and fundamental concepts for the exploration geologist / Franco Pirajno. p. cm. Includes bibliographical references and index. ISBN-13:978-3-642-75673-3 I. Hydrothermal deposits. 2. Prospecting. l. Title. TN263.P58 1992 622'.159--dc20 91-34466 This work is subject to copyright. All rights are reserved, whether the whole or part of t; he material is concerned, specifically the rights of translation, reprinting, re-use of illustrations, recitation, broadcasting, reproduction on microfilm or in any other way, and storage in data banks. Duplication of this publication or parts thereof is permitted only under the provisions of the German Copyright Law of September 9, 1965, in its current version, and permission for use must always be obtained from Springer-Verlag. Violations are liable for prosecution under the German Copyright Law. ©! Springer-Verlag Berlin Heidelberg 1992 Softcover reprint of the hardcover I st edition 1992 The use of general descriptive names, registered names, trademarks, etc. in this publication does not imply, even in the absence of a specific statement, that such names are exempt from the relevant protective laws and regulations and therefore free for general use. Media conversion: Elsner & Behrens GmbH, D-6836 Oftersheim 32/3145-543210 - Printed on acid-free paper To my parents Edgardo and Carmen Pirajno Preface This book is intended primarily for exploration geologists and postgraduate students attending specialist courses in mineral exploration. Exploration geologists are engaged not only in the search for new mineral deposits, but also in the extension and re-assessment of existing ones. To succeed in these tasks, the exploration geologist is required to be a "generalist" of the Earth sciences rather than a specialist. The exploration geologist needs to be familiar with most aspects of the geology of ore deposits, and detailed knowledge as well as experience play an all important role in the successful exploration for mineral commodities. In order to achieve this, it is essential that the exploration geologist be up to date with the latest developments in the evolution of concepts and ideas in the Earth sciences. This is no easy task, as thousands of publications appear every year in an ever increasing number of journals, periodicals and books. For this reason it is also difficult, at times, to locate appropriate references on a particular mineral deposit type, although this problem is alleviated by the existence of large bibliographic data bases of geological records, abstracts and papers on computers. During my teaching to explorationists and, indeed, during my years of work as an explorationist, the necessity of having a text dealing with the fundamental aspects of hydrothermal mineral deposits has always been compelling. Metallic mineral deposits can be categorised into three great families, namely: (I) magmatic; (2) sedimentary and residual; (3) hydrothermal. This book deals with the hydrothermal family. Hydrothermal activity, in its broadest sense, is responsible for a very wide range and variety of metallic ore deposits. The intent of this book, however, is not to provide a classification of hydrothermal mineral deposits, for which the reader is referred to Guilbert and Park (1986). The main purpose of this book is to guide the exploration geologist to the most economically and geologically important hydrothermal systems, their possible origin and nature, and the nature of their products, i.e. the ore deposits - and the geotectonic setting in which they occur. As mentioned above there is today a large number of published works dealing with specific aspects of hydrothermal mineral deposits. In general, a great deal of emphasis is placed on laboratory data, such as experimental studies, isotope systematics and geochemistry. As a consequence, field geological observations have suffered a decline. Many of these works are collected into edited volumes, which treat hydrothermal deposits by VIII Preface district or regions, environments of ore deposition, or by geological age. Although by no means exhaustive, I have endeavoured to record in this book the fundamentals of the geology of hydrothermal ore deposits, and the most recent conceptual models which, however imperfect or incomplete, attempt to explain their origin. These models are based on the available evidence and clearly, as more information is published, they will need to be refined, modified or changed altogether. In so doing, I have drawn heavily from the published work of specialists and therefore the theories, concepts and ideas presented are the result of many years of research by Earth scientists from all over the world. Some of the information presented in this book is drawn from unpublished and published works carried out whilst I worked as an exploration geologist in southern Africa, Southeast Asia, Australia, New Zealand and the southwest Pacific islands. Much of the information on southern African deposits was gained during my years as Professor of Exploration Geology at Rhodes University in Grahamstown. During this time, I directed and co-ordinated post-graduate courses in exploration geology, which include field-based studies of ore deposits, their genesis and tectonic settings. This afforded me the unique and pleasurable opportunity of examining and studying a wide range and a great number of mineralised environments, such as only southern Africa can offer. Consequently, two biases - which reflect my experience and provinciality - are discernable in the text. One bias is geographic, as many of the examples discussed are from southern Africa and Oceania. However, I make no apology for this geographical bias, because these land masses are endowed with a wide spectrum of hydrothermal mineral deposits. The other bias is on the selection of topics. These I have chosen largely on the basis of what I believe to be interesting and useful for the explorationist. Finally, the reader must note that I take full responsibility for the contents of this book. If some of the ideas and concepts taken from the literature have been misrepresented, this was unintentional, and in no way reflects a disregard of the original author(s), work. Reference Guilbert J M, Park C F (1986) The geology of ore deposits. Freeman, San Francisco, 985pp Acknowledgements The compilation of this book has benefited from my association with colleagues and post-graduate students in New Zealand, South Africa and Namibia. I lived and worked in New Zealand as an exploration geologist for many years, and I wish to acknowledge the influence that this country and its friendly people have had in shaping my philosophy of life and the benefit derived from my association with its geological community. In particular, lowe gratitude to Drs. I. Speden and R. L. Brathwaite of the Geological Survey, and the late Prof. N. Brothers of the University of Auckland. The "boys" of Gold Mines of New Zealand Ltd.: Greg Walker, Geoff Taylor, Terry Bates, Ian McKenzie, Paul Rutheford, Chris Torrey, Simon Henderson, Phil Bentley, Kajava Davui, are thanked for contributing to a very stimulating professional life. John Kernick (Exploration Manager, Australian Anglo American Ltd.) Dick Lay (Chief Geologist) and John Lawrey (Director) made it all possible. lowe gratitude to the community of the village of Reefton (Westland) for its hospitality, in particular I wish to thank my friends Graeme Jacobs, Maria and Ivan Wilson. The many and detailed field trips undertaken in South Africa and Namibia have provided invaluable experience and knowledge of a treasure house of ore deposits unique in the world. South African mining companies - Anglo American Corp., Gold Fields of South Africa, Genmin, Anglo Vaal and Johannesburg Consolidated Investments - are thanked for their generous support of these field trips, as are their on-site geologists, who proved enthusiastic hosts, and have shared their knowledge and experience with us, frequent visitors from Rhodes University. A special debt of gratitude is owed Gold Fields Namibia and their geologists, Anton Esterhuizen, Volker Petzel, Louis Kruger, Pat Vickers, Ehrart Kostlin and Dudley Corbett, for their continued support of my work in Namibia - a beautiful country and a paradise for any natural scientist. I also wish to acknowledge the contribution of: Prof. R.E. Jacob, Drs. I. Reynolds, J.S. Marsh and R. Scoon; Messrs. C. Mallinson, R. Rossiter, P.J. Ash, V.F.W. Petzel, N. Franey, J. Misiewicz, I. de Klerk, B. Joubert and D. Whitfield of Rhodes University; Drs. A.A. de Gasparis and H. Gewald of Anglo American Corporation; F. Badenhorst and K.H. Hoffmann of the Geological Survey of Namibia. My secretary at Rhodes, x Acknowledgements. Lori Wades on draughted most of the illustrations, while Amanda Butcher read the manuscript and helped with the editing. The last chapter was completed in Perth, Australia, and I wish to express my gratitude to an old friend - Murray Surtees - for having made available the facilities of his office. My wife, Mariateresa, accompanied, supported and encouraged me throughout this long journey. FRANCO PIRAJNO Contents I Introduction ............................................ . Part I Nature and Types of Hydrothermal Solutions and Systems 1 Water and Solutions 1.1 1.2 1.2.1 1.2.2 1.2.3 1.2.4 5 Introduction. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Water: Its Origin and Significance ...................... Planetary Evolution and the Origin of Water ............. Water - Past and Present. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Water in Subduction Zones. . . . . . . . . . . . . . . . . . . . . . . . . . .. Water in the Crust ................................... 1.3 Solutions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .. 1.4 Solubility and Boiling ................................ 1.5 Acid-Base Nomenclature. . . . . . . . . . . . . . . . . . . . . . . . . . . . .. 1.6 Structure of Water - Hydrolysis and Hydration. . . . . . . . . .. 1.7 Redox Potential (Eh) ................................. 1.8 Chemical Potential, Chemical Activity and Fugacity ....... References ................................................ 15 16 17 18 20 21 2 Hydrothermal Solutions ............................... 23 Introduction ........................................ Water of Hydrothermal Solutions ...................... Oxygen and Hydrogen Isotope Systematics of Hydrothermal Fluids. . . . . . . . . . . . . . . . . . . . . . . . . . . . . .. 2.4 Fluid Inclusions ..................................... 2.5 Dissolved Constituents and Metals Partitioning into Hydrothermal Solutions. . . . . . . . . . . . . . . . . . . . . . . . . .. 2.5.1 Partitioning of Metallic Elements into Hydrothermal Solutions. . . . . . . . . . . . . . . . . . . . . . . . . .. 2.6 Metal Transport ... . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .. 2.6.1 Complex Ions and Ligands ............................ 2.6.2 Complex Ions in Hydrothermal Solutions . . . . . . . . . . . . . . .. 2.7 Metal Deposition .................................... References ................................................ 23 23 2.1 2.2 2.3 5 5 6 8 12 13 14 25 28 31 33 33 34 37 39 40 Contents XII 3 Hydrothermal Systems ............................... 42 Introduction ........................................ Definition and Types ................................ Magmatic Hydrothermal Systems Related to Shallow and Deep-Seated Plutonism ........................... 3.4 Magmatic-Meteoric Hydrothermal Systems, Related to Volcano-Plutonic and Volcanic Complexes ..... 3.4.1 Magmatic Hydrothermal Systems ...................... 3.4.2 Predominantly Meteoric Hydrothermal (Geothermal) Systems. . . . . . . . . . . . . . . . . . .. . . . . . . . . . . . . . . . . . . . . . . .. 3.4.3 Hot Water-Dominated and Vapour-Dominated Hydrothermal Systems ............................... 3.4.4 Hot Springs, Mud Pools, Geysers, Crater Lakes and Fumaroles ...................................... 3.5 Sub-Sea-Floor Hydrothermal Systems: Spreading Centres and Island Arcs ..................................... 3.5.1 Hydrothermal Systems in Spreading Centres ............. 3.5.2 Hydrothermal Systems in Submarine Volcanic Centres .... 3.6 Rift-Associated Hydrothermal Systems in Sedimentary Basins ................................ 3.6.1 Hydrothermal Systems ............................... 3.6.2 Hydrothermal Systems in Modern Rift Settings ........... 3.7 Hydrothermal Systems of Metamorphic and Crustal Origin 3.7.1 Metamorphism, Metasomatism, Dewatering of Rock Sequences and Fluid Generation ................ 3.7.2 Fluid Pressure, Metamorphic Porosity, Impermeable Barriers and Hydraulic Fracturing .......... 3.7.3 Metamorphic Hydrothermal Systems ................... 3.7.4 Fluid Paths: Faults, Shear Zones and Thrust Faults ....... 3.7.5 Fluids in Subduction Zones ........................... References ................................................ 42 43 3.l 3.2 3.3 44 45 46 50 56 57 59 60 66 69 71 79 83 84 86 88 91 96 97 4 Hydrothermal Alteration .............................. 10 1 4.1 4.2 4.2.l 4.3 4.3.1 4.3.2 4.3.3 4.4 Introduction ........................................ Hydrogen Ion Metasomatism and Base Exchange ......... Chemical Processes Related to Hydrogen Ion Metasomatism Styles and Types of Hydrothermal Alteration ............ Styles of Alteration .................................. Types of Alteration .................................. Other Types of Alteration ............................ Quantification and Monitoring of Hydrothermal Alteration Processes - Data Presentation ......................... Rare Earths Elements in Hydrothermal Alteration Processes Oxygen and Hydrogen Isotope Systematics .............. Metamorphism of Hydrothermally Altered Rocks ........ 4.4.1 4.5 4.6 10 1 102 103 109 110 110 123 132 136 139 142 XIII Contents 4.7 Detection of Hydrothermal Alteration by Spectral Remote Sensing ........................... 150 References ................................................ 152 Part II Crustal Evolution, Global Tectonics, Hydrothermal Mineral Deposits and Mineral Exploration Geotectonic and Metallogenic Analysis of Orogenic Belts 5 Crustal Evolution, Global Tectonics and Mineral Deposits .. 159 5.1 5.2 Introduction. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .. Tectonic Phases in the Earth's Geological Evolution and Related Metallogeny ............................. 5.2.1 The Archean Eon: Phase of Microplate Tectonics ........ 5.2.2 The Proterozoic Eon: Phase ofIntraplate Tectonics ...... 5.2.3 . The Phanerozoic Eon: Phase of Macroplate Tectonics, Sea Floor Spreading and Continental Drift ............. 5.2.4 Conclusions ........................................ References ................................................ 6 162 162 171 179 189 190 Geological Processes and Hydrothermal Mineralisation in Plate Tectonic Settings - Mineral Exploration .......... 197 6.1 Introduction. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .. 6.2 Extensional Plate Tectonics ........................... 6.2.1 Mid-Ocean Spreading Centres ........................ 6.2.2 Intracontinental Rifts ................................ 6.2.3 Passive Continental Margins and Interior Basins ......... 6.3 Compressional Plate Tectonics ........................ 6.3.1 Subduction-Related Settings .......................... 6.3.2 Collision-Related Settings ............................ 6.4 Transform Fault Tectonics ........................... References ................................................ 7 Geotectonic and Metallogenic Analysis of Orogenic Belts 7.1 7.2 7.2.1 7.2.2 7.2.3 7.3 Introduction ....................................... The Pan-African Orogenic Belts of Africa .............. The Arabian-Nubian Shield .......................... The Damara Orogen, Namibia ........................ The Lufillian Fold Belt .............................. Metallogenic Epochs and Geotectonic Environments of Hydrothermal Mineral Deposits of the Orogenic Belts in New Zealand ..................................... Geotectonic Settings of the Tuhuan Orogeny and Related Hydrothermal Mineralisation .............. 7.3.1 159 197 198 199 201 205 206 207 212 215 216 .. 219 219 219 221 226 233 236 236 Contents XIV 7.3.2 Geotectonic Settings of the Rangitata Orogeny and Related Hydrothermal Mineralisation .............. 239 7.3.3 Geotectonic Settings of the Kaikoura Orogeny and Related Hydrothermal Mineralisation .............. 241 References ................................................ 241 Part III Hydrothermal Processes and Activities Related Mineral Deposits 8 Alkali Metasomatism and Related Mineral Deposits 8.1 8.2 8.2.1 8.2.2 8.2.3 8.2.4 8.3 8.3.1 8.4 8.4.1 Introduction ....................................... Alkali Metasomatism in Continental Igneous Systems .... Role of Volatiles in Granitic Magmas .................. Textural Features ................................... Sodic Metasomatism and Albitites ..................... Potassic Metasomatism and Microclinites ............... Alkali Metasomatism in Anorogenic Ring-Type Complexes Fenites. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .. Mineralisation Related to Alkali Metasomatism ......... Mineralisation in Ring Complexes of the Ijolite-Carbonatite Association .................. 8.4.2 Mineralisation in Ring Complexes of the Alkaline Granite Association .................... 8.4.3 Mineralisation Related to Alkali Metasomatism of Pegmatites ....................................... References ................................................ 9 9.1 9.2 9.3 9.4 9.4.1 247 247 247 247 251 253 255 256 259 265 266 271 274 277 Greisen Systems .................................... 280 Introduction ....................................... Greisenisation Processes ............................. Geochemistry ...................................... Greisen-Related Mineral Deposits ..................... Sn and W Geochemistry in the Greisen EnvironmentDeposition of Cassisterite and Wolframite .............. 9.4.2 Sn Deposits Associated with the Acid Phase of the Bushveld Igneous Complex, South Africa ......... 9.4.3 Sn-W Mineralisation at Brandberg West, Damara Orogen, Namibia ............................ 9.4.4 Endo- and Exogreisen Sn Mineralisation at Mount Bischoff, Tasmania ......................... 9.4.5 The Hercynian Sn-W Deposits of Southwest England, Cornwall and Portugal ............................... References ................................................ 280 282 289 290 291 294 304 310 314 322 xv Contents 10 Porphyry Systems and Skarns 10.1 Introduction. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .. 10.2 Tectonic Settings .................................... 10.3 Classification of Porphyry Systems .................... 10.4 Hydrothermal Alteration and Mineralisation ............ 10.4.1 Lowell-Guilbert Model .............................. 10.4.2 Diorite Model ...................................... 10.4.3 Alteration-Mineralisation of Carbonate Wall Rocks (Skarns) 10.5 Mineral Deposits of Porphyry Systems ................. 10.5.1 Panguna and Ok Tedi Porphyry Cu-Au Deposits ........ 10.5.2 Porphyry Cu-Mo Deposits in Chile .................... 10.5.3 Porphyry Mo Deposits of the Colorado Mineral Belt ..... 10.5.4 Porphyry Mo Mineralisation in the Oslo Graben, Norway 10.5.5 Skarn Deposits in the Western USA ................... 10.5.6 Other Types of Skarn Deposits ........................ 10.5.7 Porphyry Sn Deposits in Bolivia ...................... References ................................................ 325 325 326 328 330 333 335 336 344 344 349 355 360 361 364 367 370 11 Fossil and Active Geothermal Systems Epithermal Base and Precious Metal Mineralisation (Including Kuroko-Type Deposits) ...................... 11.1 Introduction ....................................... 11.2 General Characteristics of Epithermal Systems .......... 11.2.1 Main Types of Epithermal Deposits ................... 11.3 Volcanic-Hosted Epithermal Deposit Types ............. 11.3.1 Epithermal Systems of Submerged Volcanic Structures ... 11.3.2 Hydrothermal Alteration ............................. 11.3.3 Mineral and Metal Zoning ........................... 11.4 Transport and Deposition of Precious Metals in Epithermal Systems ............................... 11.4.1 Boiling Depths and Metal Zoning ..................... 1l.5 Active Geothermal Fields ............................ 11.5.1 Geothermal Systems of the Taupo Volcanic Zone, New Zealand ....................................... 11.5.2 Salton Sea, California, USA .......................... 11.6 Volcanic-Hosted Epithermal Mineral Deposits .......... 11.6.1 Hauraki Goldfields, Coromandel Peninsula, New Zealand 11.6.2 Epithermal Au in Lihir Island, Papua New Guinea ....... 11. 7 Sediment-Hosted Epithermal Deposits ................. 11. 7.1 Mineral Belts and Deposit Types of Nevada, USA ....... 11.7.2 Nature of Fluids and Ore Genesis ..................... 11.8 Kuroko-Type Mineral Deposits ....................... 11.8.1 Kuroko Deposits ................................... 11.8.2 Precambrian Volcanogenic Massive Sulphide Deposits ... References ................................................ 375 375 378 380 385 387 390 394 396 400 401 402 405 407 407 415 418 420 426 427 428 440 443 XVI 12 Contents Hydrothermal Processes in Oceanic Crust and Related Mineral Deposits ......................... 450 12.1 12.2 12.2.1 12.2.2 12.2.3 12.3 12.4 12.5 Introduction. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .. Physiography of the Ocean Floor ...................... Mid-Ocean Ridges .................................. Transform Faults and Fracture Zones .................. Seamounts and Volcanic Chains ...................... Birth, Life and Death of an Ocean Basin ............... Oceanic Lithosphere and Ophiolites ................... Heat Flow, Oceanic Crust Metamorphism and the Nature of Related Hydrothermal Solutions ...... 12.5.1 Heat Flow and Oceanic Crust Metamorphism ........... 12.5.2 Nature and Composition of the Hydrothermal Solutions .. Tectonic Settings, Sub-Sea-Floor Hydrothermal Processes, 12.6 Hot Springs and Their Mineral Deposits ............... 12.6.1 Tectonic Settings .................................... 12.6.2 Hydrothermal Processes and Types of Sulphide Deposits . 12.6.3 Sub-Sea-Floor Hydrothermal Mineral Deposits ......... Oceanic Crust-Related Hydrothermal Mineral Deposits ... 12.7 12.7.1 Massive Sulphide Deposits of the Samail Ophiolite, Oman 12.7.2 The Cu Deposits of Cyprus Island ..................... 12.7.3 The Cu Deposits of the Matchless Amphibolite Belt, Namibia ........................................... References ................................................ 13 l3.1 l3.2 13.2.1 13.3 13.3.1 13.3.2 13.4 13.4.1 13.4.2 13.4.3 13.4.4 l3.5 13.5.1 13.5.2 Hydrothermal Mineral Deposits of Continental Rift Environments Introduction ....................................... Continental Rifting .................................. Geophysical Signatures of Continental Rifts ............ Magmatism and Metamorphism Associated with Rifting .. The Nature of Igneous Activity in Rift Systems .......... Metamorphism in Continental Rifts ................... Basin Formation and Volcano-Sedimentary Sequences in Continental Rifts ................................. The Stratigraphic Record of Proterozoic Basins in South Africa ..................................... The Stratigraphic Record of Aulacogens ................ The East African Rift System ......................... The Rio Grande Rift (USA) .......................... Continental Rifting in Space and Time Hydrothermal Mineral Deposits ....................... Early Stages of Continental Rifting .................... Aulacogens and Troughs Intermediate Stages of Continental Rifting .............. 450 451 452 453 454 454 457 460 461 465 468 470 473 476 488 488 491 494 502 507 507 508 509 511 515 516 518 519 522 525 527 530 530 531 XVII Contents 13.5.3 13.6 13.6.1 13.6.2 13.6.3 Advanced Stages of Rifting ........................... Hydrothermal Mineral Deposits in Incipient Rifts ....... The Messina Cu Deposits, South Africa ................ Olympic Dam (Roxby Downs), South Australia ......... Hydrothermal Activity in the Tanganyika Trough, East African Rift System ............................. 13.7 Hydrothermal Mineral Deposits in Aulacogens and Troughs at Intermediate Stages of Rifting ........... 13.7.1 McArthur River and Mt. Isa, Northern Australia ........ 13.7.2 The Sediment-Hosted Exhalative Massive Sulphide Deposits in the Namaqualand Metamorphic Complex, South Africa ....................................... 13.7.3 Stratabound Cu-Ag Deposits of the Irumide Belt in Southern Africa .................................. 13.7.4 The Zambian Copperbelt ............................ 13.7.5 Stratiform and Stratabound Cu Deposits of the Keweenawan Rift .............................. 13.8 Mineral Deposits Related to Advanced Stages of Rifting the Red Sea Deeps .................................. 13.9 Banded Iron Formation (BIF) of Proterozoic Age ....... 13.9.1 The Mineral Deposits of the Transvaal-Griqualand Basins ................... References ................................................ 535 536 536 541 544 545 545 549 555 559 563 565 568 569 572 14 Stratabound Carbonate-Hosted Base Metal Deposits ...... 578 14.1 14.2 14.2.1 14.2.2 14.3 14.4 14.4.1 14.5 Introduction ...................................... . Mississipi Valley-Type Deposits (MVT) ............... . The Viburnum Trend, USA ......................... . Pine Point, Canada ................................. . Alpine-Type Deposits .............................. . Irish-Type Deposits ................................ . Mineral Deposits .................................. . Models of Ore Genesis for the MVT, Alpine and Irish Types .............................. . Karsting .......................................... . Nature and Temperature of Fluids, Source of Metals and Sulphur ....................................... . The Carbonate-Hosted Pb-Zn-Cu-Ag and V Deposits of the Otavi Mountain Land, Namibia ................ . Geology, Structure and Metamorphism ............... . Mineralisation ..................................... . Tsumeb .......................................... . Kombat .......................................... . Berg Aukas ....................................... . 14.5.1 14.5.2 14.6 14.6.1 14.6.2 14.6.3 14.6.4 14.6.5 578 580 580 583 585 586 587 588 590 591 592 594 596 597 602 604 Contents XVIII 14.6.6 Models of Ore Genesis ............................... 605 References ................................................ 609 15 15.1 15.2 15.2.1 15.2.2 15.2.3 15.2.4 15.2.5 15.2.6 15.2.7 15.3 15.4 15.4.1 15.4.2 15.4.3 Crustal Hydrothermal Fluids and Mesothermal Mineral Deposits Introduction ....................................... Metamorphism and Fluid Generation .................. Metamorphic Devolitilisation Reactions ................ Fluid Transport and Migration ....................... Shear Zones ........................................ Metamorphic Vein Systems and Vein Growth ........... Mass Transport and Movement of Metals .............. Au in Hydrothermal Fluids ........................... Oxygen and Hydrogen Isotope Systematics ............. Tectonic Settings .................................... Archean Mesothermal Deposits ....................... The Archean Greenstone Belts ........................ Metallogenesis ...................................... Theories on the Genesis of Archean Mesothermal Au Deposits ................. J5.4.4 Mesothermal Au Deposits of the Barberton and Murchison Greenstone Belts, South Africa .......... 15.4.5 The Golden Mile, Yilgarn Block, Western Australia ...... 15.4.6 The Hemlo Au-Mo Deposit, Superior Province, Canada .. 15.5 Mesothermal Vein Deposits of Phanerozoic Age (Turbidite-Hosted Au) ............................... 15.5.1 The Ballarat Slate Belt, Victoria, Australia .............. 15.5.2 Hydrothermal Lode Systems of Otago-Marlborough and the Southern Alps, New Zealand .................. 15.5.3 The Juneau Gold Belt, Southeast Alaska ............... Mineral Deposits Formed by Multistage Ore Genesis ..... 15.6 15.6.1 Unconformity-Related U Deposits ..................... 15.6.2 Au Mineralisation in the Central Zone of the Damara Orogen, Namibia ...................... 15.6.3 The Possible Role of Metamorphic Fluids in the Origin of the Witwatersrand Goldfields, South Africa .......... References ................................................ 612 612 615 615 617 618 619 622 626 631 633 636 636 641 643 654 661 663 666 667 670 675 676 676 681 683 685 Epilogue .................................................. 693 Appendix ................................................. 697 Subject Index .............................................. 703 Introduction In the study of hydrothermal mineral deposits, the main objective is a basic understanding of ore genesis processes and the eventual building of models, which would best explain empirical observations, and enable deposit types to fit within the local and global geotectonic frameworks. The information required to achieve this must include: (1) origin and composition of the ore solutions; (2) physico-chemical conditions of solution, transport and deposition; (3) temperature, pressure and depth of ore formation; (4) source of the ore constituents; (5) metasomatic processes; (6) local geology and structure; (7) regional geology and tectonic setting; (8) origin and evolution of the magmas that give rise to hydrothermal processes. This information must be obtained through studies of geological characteristics and field relationships, mineralogy, trace and major element geochemistry of ores and wall rocks. The study of fluid inclusions provides data on the nature ofthe solutions, temperature, pressure and depth of ore formation, while stable isotope systematics helps in constraining the source of the fluids and tracing their movement through, and interaction with, the crust. Thus, it can be seen that the study of an ore deposit is a multidisciplinary effort. It requires that information be drawn from "outside" disciplines, such as physics, chemistry, mathematics and computer science, as well as the various branches of the Earth sciences. Regional geological, structural and tectonic studies are of great importance in order to understand the metallogeny of a sector of the Earth's crust. The study of individual deposits and groups of depohl1s contributes to the understanding of the tectonic and metallogenic framework of an area. Conversely, a clear idea of the tectonic setting and geodynamic history of a region, or an orogen, is extremely important in order to predict the types and styles of mineralisation that may occur. The predictive capacity of regional or global geotectonic perspectives is often underestimated or underutilised. As mentioned above one of the ultimate objectives is the construction of an ore deposit model. A mineral deposit model describes the essential characteristics of a group or class of mineral deposits. In modelling ore deposits, therefore, all the available information must be considered. The model mayor may not include a theoretical component, in which ideas and concepts that attempt to explain the origin of the mineralisation and its relationship with the hosting rocks, are proposed. Genetic models tend to evolve from the simplistic, when the data base is limited, to the complex, as this data base grows. Thus, even those deposits whose genetic processes were seen to be well established may be questioned as more and more data become available. A striking example of this is provided by the conglomerate-hosted Au-U ores of the Witwatersrand basin in South Africa (see Chap. 15). Ore deposit models 2 Introduction are inexpensive but effective tools for mineral exploration and they are useful for the preliminary assessment of the mineral resources of a region. Each chapter in this book can be considered on a "stand-alone" basis, so that the reader interested in a specific hydrothermal system can simply select the relevant chapter. There are 15 chapters, which are organised into three parts. Part I is an introduction to the nature and types of hydrothermal systems and wall-rock alterations. Part II deals with crustal evolution and the relationship of hydrothermal mineral deposits to the geological evolution of the Earth and it examines the relationship between tectonic settings and hydrothermal mineral deposits. This is accompanied by brief comments on mineral exploration techniques. Part II ends with a chapter on geotectonic and metallogenic analyses of selected areas in southern Africa and Oceania. In Part III hydrothermal processes are treated more specifically and in greater detail, and are accompanied by discussions on selected examples which, in my opinion, best illustrate these processes. Part III begins with deep-seated magmatic-hydrothermal systems (Chap. 8 on alkali metasomatism; Chap. 9 on greisens). The following chapters reflect the increasing role, with shallowing depth, of meteoric waters in the hydrothermal systems and increasing degrees of water/rock ratios. Thus, Chapter 10 looks at porphyries and skarns, while Chapter 11 examines active and fossil geothermal (epithermal) systems and related volcanic- and sediment-hosted mineral deposits. Kuroko and Precambrian volcanogenic massive sulphide deposits are included in this latter chapter because they represent the submarine equivalent of subaerial geothermal systems. In Chapters 12 and 13 rifting (oceanic and continental) associated hydrothermal activity and mineral deposits are discussed. Oceanic crust environments are examined in Chapter 12 with a discussion on present-day occurrences and their fossil equivalents in the geological record. In Chapter 13 the complex topic of continental rifting and related hydrothermal mineral deposits is examined. It includes examples of present-day mineral deposition in modern rifts. The discussion of deposit types in this chapter encompasses a wide range, from epigenetic to syngenetic, and includes deposits ranging from the Messina breccia pipes, to sediment-hosted massive sulphide deposits, stratiform and stratabound disseminated sulphides and banded iron formations. In most of these deposits the relationship to rifting is based on good evidence, but in a few this is not clear cut and, to add to the problem, there are some deposit types which are of a transitional character. For this reason, it is expedient to consider the rift-related mineralisation in terms of end-member (incipient to advanced) stages of rift evolution. Low-temperature hydrothermal processes related to the dewatering of sedimentary basins, which result in the formation of carbonate-hosted base metal deposits, are examined in Chapter 14. High-temperature metamorphic processes are examined in Chapter 15, which deals with mesothermal vein lodes of Archean and Phanerozoic age, most of which are economically important by virtue of their precious metal content. Ideas and conceptual models on the genesis of Archean lode deposits have flourished in the last decade and for this reason this topic is treated in greater detail. I conclude with a brieflook at hydrothermal deposits which may have originated through a number of genetic processes. A synopsis of hydrothermal processes in the Earth's lithosphere is given in an Epilogue. Part I Nature and Types of Hydrothermal Solutions and Systems Chapter 1 Water and Solutions 1.1 Introduction By way of an introduction to the study of hydrothermal mineral deposits, this chapter reviews what we understand of water, its origin on our planet and its significance in terms of planetary, hydrothermal and ore genesis processes. The nature of solutions, solubility, the phenomena of hydrolysis and hydration, redox potential and fugacity are also briefly considered, before continuing with a discussion of the nature of a large group of mineral deposits, whose origin is attributed to the action of hydrothermal solutions. For a more detailed study of the subject covered during the course of this chapter, the interested reader is referred to the books of Mason (1966), Krauskopf (1979), Masterton et al. (1981), Press and Siever (1982), Holland (1984) and Gill (1989). 1.2 Water: Its Origin and Significance An oxide of hydrogen, water is composed of two hydrogen atoms and one oxygen atom. It is a volatile polar molecule in which the oxygen atom possesses a weak negative charge and the hydrogen atoms a weak positive charge. The polarity of this molecule is important, making it a good ligand for cations and enabling processes of hydration and hydrolysis (Sect. 1.6). Pure water dissociates according to: H20 =H++OHThe OH- ion is known as the hydroxyl group, which commonly enters the structures of a large group of minerals, such as amphiboles and micas. Perhaps surprisingly water is a minor constituent of our planet. Slightly more than 70 % of the surface of the Earth is covered by a thin film ofliquid water. In fact if the planet's dimensions are taken in due account, the oceans constitute less than 0.03% of the total mass of the Earth. Nevertheless, water is not an uncommon constituent in other parts of the solar system, where it may be abundant as water-ice in the outer planets. By contrast the "rocky", or terrestrial planets (Mercury, Venus, Earth, Mars) are relatively dry. Water on Earth occurs as a liquid phase in the oceans, rivers, lakes and underground in the pores of the solid surface materials; as a vapour in the atmosphere and as volcanic emissions; as a solid phase in the polar ice 6 Water and Solutions caps and mountain glaciers; and as a chemical constituent of the hydrous minerals in the crust and mantle. All of this water participates in two major geological cycles: the endogenic cycle, which derives its energy from the interior of the Earth, giving rise to tectonic, igneous and metamorphic processes; and the exogenic cycle at and close to the surface, which derives its energy from the Sun, and promotes the interaction of the hydrosphere, atmosphere and lithosphere (hydrologic and weathering cycles). Juvenile water is that which has never been part of the surface cycle. While its existence may be postulated its detection remains elusive in scientific investigation, because of the obvious contamination with recycled water. One possible way of detecting juvenile water is in those minerals of mantle origin for which a unique isotopic composition must be postulated (Sheppard 1986). Water is a vital constituent of the physical world, being readily found in all three states of matter liquid, solid and gas. Used as a standard for physical constants, it is a powerful solvent, which carries and deposits substances, and in which life abounds. Life, in the sense that we understand it, most probably originated in water, and to remind us of this, all living things are made up of light constituents of which water is by far the major one. 1.2.1 Planetary Evolution and the Origin of Water The Earth was formed approximately 4.6 Ga ago. The planets of the solar system closest to the Sun - Mercury, Venus, Earth-Moon system, Mars and the asteroid belt - are rocky and volatile-poor compared to the outer planets. The planets were formed by the collision and cold accretion of nebular grains and larger bodies known as planetesimals, whose compositions differed at different distances from the parent solar nebula. The solar nebula is the raw material of our system, and is estimated to be comprised of a 98% mixture of H2 and He, less than 2% volatiles (H 20, CH4 , NH 3, etc), and approximately 0.5 % solids containing Si0 2, MgO, FeO, FeS etc. (Lewis and Prinn 1984). Metals, oxides and silicates dominate in the region between Mercury and the asteroids, whereas further beyond water-ice, methane, ammonia and other light materials are more abundant. Nevertheless, comparatively small amounts of these volatiles were present in the accreting rocky debris at the time of the Earth's formation. The newly-formed terrestrial planets were subsequently heated through three main processes: impact energy, gravitational compression and radioactivity. As heat was produced far more rapidly than it could flow out into space, melting began to take place. This initial melting led to the differentiation of the planetary bodies into a series of shells with progressively lighter materials towards their surfaces. On Earth aluminium, silicon, potassium and sodium accumulated as a slag to form the outermost shell (crust) towards the surface, followed by an inner shell (mantle) composed of Fe and Mg silicates with denser and more closely packed crystal structures. The sinking of heavier metallic materials in turn gave rise to the formation of an innermost shell (core), believed to consist mostly of an Fe-Ni alloy. It follows that the Earth's crust has different elemental abundances compared to the whole Earth, as shown in Table 1.1. 7 Water: Its Origin and Significance Table 1.1. Average composition of sea water in part per thousand (ppt), and as percent of total. (After Mason 1966) Crust 0 Si AI Fe Mg Ca K Na Whole Earth 46% 28% 8% 6% 4% 2.4% 2.3% 2.1% other <1% Fe 0 Si Mg Ni S Ca Al 35% 30% 15% 13% 2.4% 1.9% 1.1% 1.1% other <1% A natural consequence of heating and partial melting is volcanic activity that is, the expression of the transfer of heat and mass out onto and towards the surface. It is here that the process of volatile release took place, as it does today through volcanoes and at mid-ocean ridges. Outgassing of water and volatiles, including CO 2 , CO, H 2S, H 2 , N 2, CH 4 , NH 3, HF, HCI, Ar, Ne etc., occurred at a great rate (big burp hypothesis) during this early period of differentiation. Planetary bodies with sufficient gravity, such as the Earth and Venus, and to a lesser extent Mars, were able to retain most of these volatiles, with the exception ofH 2 and He which escaped into space. It is calculated that the juvenile volatiles were outgassed during the first 108 -109 years (Pollack and Yung 1980). On Earth the temperature must have been low enough to allow condensation of water with much CO 2 in solution, and the leaching of other elements from the surface materials. This water collected first of all into pools and then into larger basins to form the oceans. It has been proposed that these larger ba~ns were maretype basins, such as those formed on the Moon, initially formed by the flux of asteroid-size objects (Frey 1977). Most of the Earth's water took up residence in the oceans, and CO 2 in carbonate rocks, while the other volatiles formed a gaseous shell enveloping the entire planet (atmosphere). Studies in the dating of minerals by ionprobe techniques suggest that some form of continental crust must have been in existence since at least 4.1 Ga (Froude et al. 1983). The presence of sedimentary rocks as old as 3.8 Ga indicate that a hydrosphere and atmos·phere must have been well-established since at least that time, suggesting consequences in later developments far more profound than is usually acknowledged by planetary geologists. For example, an unresolved question remains as to whether a granitic crust was formed at one particular early stage or some 500 Ma later, as suggested by the ion-probe dating mentioned above. Equally unresolved is the question of whether the early crust evolved by progressive differentiation of products of partial melting of a solid mantle, or, as seems to be the case for the lunar highlands, by early crystallisation directly from planetary-wide melting. The mass of the Earth' crust relative to the rest of the planet is only about 0.5% and, as pointed out by Taylor and McLennan 8 Water and Solutions (1985), it, like the hydrosphere, may be unique to our planet, since nothing like it has as yet been found on other planets, with the possible exception of Venus. In other words, the retention of a hydro-atmospheric envelope was probably one of the essential ingredients for the development and evolution of a granitic crust. The weathering cycle allowed the accumulation of sedimentary packages, which through metamorphism and partial melting formed crustal melts of granitic composition. Recent hypotheses on the origin and evolution of the primordial continental crust have been proposed by Kroner (1985) and Lowman (1989). Both these authors envisage an important role for mantle plume (hot spot) volcanism as a crustbuilding mechanism. Lowman (1989) proposes that an early crust developed by non-plate tectonic global differentiation followed by basaltic underplating and overplating. This basaltic volcanism would have been initiated by basin-forming meteorite impacts at least 4 Ga ago. Kroner (1985) for the building of the early crust, places his emphasis on hot spot volcanism, differentiation and re-melting. 1.2.2 Water - Past and Present The atmosphere and hydrosphere together make up the gaseous-liquid envelope of the Earth, and, as outlined above, were first formed by the early outgassing of the planet. Condensation of liquid water and its accumulation in low-lying areas, or protobasins, formed the early oceans, which explains the discontinuity of the liquid envelope around the Earth. The hydrosphere comprises all the waters at the surface of the Earth, oceans, lakes, rivers and ice. Present-day oceans cover approximately 366 X 106 km 2, or 70.8%, of the surface of the Earth. The mass of sea water constitutes roughly 98% of the total hydrosphere, and fresh waters the remaining 2 %, of which about 22 % is groundwater. Approximately 77 % of the fresh water is locked up in ice, and less than 1 % is surface meteoric water. This makes water a precious commodity. The history of the hydrosphere and atmosphere is essentially the result of the interplay of volcanic activity, tectonism and the increasing role of biological activity (Holland 1984). Evidence from both terrestrial and marine environments indicates that the salinity of the early oceans had already been established since the formation of the first large bodies of water more than 4 Ga ago. Thus the primeval oceans must have contained significant concentrations of Cl-, Na+, Ca2+ and Mg2+, with Na+ and Cl- being the dominant cation and anion respectively (Holland 1984). That the early oceans were saline is also corroborated by the presence of sedimentary deposits of evaporitic origin in rock sequences of Archean age, such as those of the Warrawoona Group (3.5 Ga) in Australia. In the rocks of this Group, welldeveloped, rosette-shaped minerals replaced by silica are interpreted to represent casts of evaporite minerals such as gypsum (Holland 1984). There is also by now irrefutable evidence that the oxidation state of the atmosphere was gradually acquired through the geological ages, and that during the first 500 Ma it had at least a mildly reducing character. Evidence from paleosols, the Witwatersrand paleoplacers in Soutli Africa, with their detrital pyrite and uraninite grains, and the banded 9 Water: Its Origin and Significance Table 1.2. Average abundance by weight of elements in the crust and the whole Earth (After Press and Siever 1982) Ion ppt percent CINa+ soi+ Mg2+ Ca2+ K+ HC0 3+ BrH 2B0 3 Sr2+ 18.98 10.56 2.65 1.27 0.40 0.38 0.14 0.065 0.026 0.008 55.05 30.61 7.68 3.69 1.16 1.10 0.41 0.19 0.07 0.03 Total 34.48 99.99 Table 1.3. Chief salts of marine evaporites. (After Krauskopf 1979) Chlorides Formula Halite Sylvite Carnallite Bishofite NaCI KCl KMgCI 3 ·6H 20 MgCI 2.6H 20 Sulphates Kainite Anhydrite Gypsum Polyhalite Kieserite KMgCIS04·11/4H 20 CaS04 CaS04·2H20 K2MgCa2(S04)4·2H20 MgS0 4·H 20 iron formations, suggest an atmosphere poorer in oxygen during the Archean than today. The composition of sea water is defined by salinity, which is the total amount of dissolved solids per kilogram of water. Table 1.2 gives an average composition of sea water. The precipitates obtained from the evaporation of sea water (evaporites) can constitute important non-metallic ore bodies. Leaching of ancient evaporites may influence the composition and precipitation of sulphide species, as is the case for the sulphide brines formed in the Red Sea (see Chaps. 3 and 13). Table 1.3 lists some of the more common and important salts of marine evaporites. Sea water also contains gases in solution (0 2 , N2 , CO 2 , Ar, H 2 S) as well as a host of other elements induding Li, C, AI, Si, P, Ti, V, Mn, Fe, Co, Ni, Cu, Zn, As, Se, Rb, Water and Solutions 10 ATMOSPHERIC CIRCULATION I nput from submarine volcanism and hydrothermal plumes From subduction, metamorphi sm and partial melting Fig. 1.1. Endogenic and meteoric cycles of H 20 and CO 2 (After Press and Siever 1982) Mo, I, Ba. Elements such as Au, Ag, Hg, Pb and U are present in even lesser quantities than chloride, carbonate and sulphide species. H 2S may be locally abundant, especially in stagnant organic-rich bottom waters (e.g. Black Sea) and in the areas of mid-ocean ridges. Meteoric, or terrestrial, waters though minor components of the hydrosphere, are important in that they are responsible for the weathering, erosion and transport of material from the land masses. Figure 1.1 shows the hydrologic cycles from ocean to atmosphere, to land and back to ocean. It is estimated that meteoric water carries some 27.4 X 10 14 g of dissolved matter per year. The average salinity of river water is about 100 ppm. Table 1.4 gives the average composition of dissolved solids in river water. The amount of dissolved solids carried by the rivers annually is but a small fraction of the total mass of solids contained in the sea (49.5 X 1021 g). The material added to the oceans by the rivers is augmented by that from submarine volcanism along the active mid-ocean ridges, which injects the most common anions in the sea (CI-, sOi-, B, Br). To determine the budget of dissolved matter, it is important to know the average length of time that each element remains in the sea (residence time). This is calculated by dividing the total amount of anyone element in solution by the amount introduced each year by terrestrial waters. In terms of this residence time, elements with the highest abundanc~ (Na, CI, Mg, Ca, K) are found to have the longest residence time, whereas those with lowest abundance (e.g. Fe, AI, Zn, Cu, Mn, Pb, Au, Ag) have the shortest residence time. Figure 1.2 shows the relationship between oceanic residence time of elements anq their concentrations, as calculated by Whitfield (1982), who has presented a view of the oceanic dynamic system using 11 Water: Its Origin and Significance Table 1.4. Average composition of dissolved solids in river water expressed in parts per million (ppm), and as percent of total. (After Mason 1966) Ion HC0 3Ca2+ Si0 2 sOi- Cl- Na+ Mg2+ K+ N0 3Fe Total ppm wt% 58.5 15 13.1 11.2 7.8 6.3 48.6 12.5 11.0 9.3 6.5 5.3 3.4 2.0 0.8 0.6 4.1 2.3 I 0.7 120 100.0 TIME IN 102 104 10 3 YEARS 10 5 10 6 107 Ca Mg S K Qj 1: ..... VI OJ 0 .§. 60- 4 - -- - - -N - z 0 ~ l _ - - Si F P Ba I c:: !Z lOB Na Zr Cd U c:: UJ ::t 3: 10-12 Pb La Ce Li Sr MINOR ELEMENTS Rb TRACE ELEMENTS SbGa ~ W Th Hg Sn MAJOR ELEMENTS B - - Br -AI - - - Zn - As N· l1l1 -- - C Ti V I U Mn U Cr Sc Cs ~10-8-Fe ~ Cl Au Eu i:3 Vl ltns lOs 01s Is lOs 100s lOOOs OCEANIC RESIDENCE TIME. Earth day (4' 6 Ga =24hrs.l Fig. 1.2. Residence time of elements in oceanic water in relation to concentration (After Whitfiela" 1982) the Earth-day method, in which the 4.6 Ga of the Earth's history is compressed into 24 h. At the present rate, rivers transport enough material every "hour" to fill the ocean basins, while every 18 "seconds" the amount of dissolved material transported by rivers would double the amount of salt in the oceans. The longest residence time of the major elements is only about 1 "minute", most of the trace elements remain in 12 Water and Solutions the sea less than I "second". Sediments in the sea are less than 1 "hour" old, as the deposited material is continuously subjected to tectonic uplifting and subduction. Despite this dynamic flux, Whitfield points out that geological evidence suggests sea water composition has remained constant for at least 11 "hours", or 2000 Ma. One of the most remarkable characteristics of the oceans is that the concentration of those elements which may be harmful to living organisms (e.g. As, Se, Hg, Pb etc.) does not increase, as might be expected from the annual input from the rivers. Whitfield (1982) explains the chemical budget of the oceans in terms of a steadystate model; in other words, material is added and removed at a constant rate. He also suggests that while the composition of sea water is largely controlled by geological processes, the biotic masses may also influence modern-day seas in their control of the elements' distribution, as indicated by the correlation between the biological importance of an element and its concentration (Fig. 1.2). Some scientists believe that the composition of sea water has not remained constant throughout the geological ages. Spencer and Hardie (1990), for example, consider that the composition of sea water is controlled by the mixing of two major inputs, namely Na-HCO r S04 from rivers and Ca-Cl from hydrothermal discharges at mid-ocean ridges. Therefore the concentration of these ionic species in the sea could have varied with time as a function of the intensity of hydrothermal discharges from mid-ocean ridges and climatic changes. These authors argue that a smaller development of mid-ocean ridges would result in lesser fluxes of Ca and CI and higher solubilities of Na+, Mg2+, S02-4 and HCO- 3 from river input and viceversa in the case of larger mid-ocean ridges. This would also affect the composition of marine evaporites, in which aragonite rather than calcite and Mg-rich salts would predominate for small ocean ridges fluxes, whilst calcite and K-rich salts would predominate in the other case. 1.2.3 Water in Subduction Zones It has been calculated that, given the present rate, the entire oceanic crustal mass would be subducted within about 1000 million years, which implies there must be a strong return flow (Fyfe 1978). Endogenic processes are responsible for this continuous return flow, which has very important implications as hydrothermal activity occurs in those regions in which this return flow is most efficient. Largescale release of fluids (0 2, H 20, CO 2, S) occurs during subduction as volatile and water-bearing materials are dragged down into the mantle. At depth along the subducting slab, hydrated minerals such as amphibole and mica lose their bound water and other volatiles. During subduction of oceanic lithosphere, re-equilibration to the higher temperature and pressure of amphibolite facies in a region of 20 X 10 X 1 km releases a mass of fluid of some 20 km 3 • A proportion of this fluid moves upward through hydrofracturing, or via thrust planes (Fig. 1.3). Fyfe (1978) points out further that in addition to the volatile phases, other elements such as K, Rb, Na, U are also recycled into the mantle. In the mantle wedge above the subducting slab, mantle melts form aided by the release of volatiles. Magmas generated in subduction environments can contain in excess of 0.5% water, and 13 Water: Its Origin and Significance Volcanic gases and hydrothermal plumes Cl, S02' B, Br, He,etc. HZO, COZ' S, NZ' HZ Upward moving volatiles: Dehydra t i on of oceanic crust Fig. 1.3. Idealised scheme of circulation of H 2 0 and other volatiles in the lithosphere-hydrosphereatmosphere system. MOR Mid ocean ridge active volcanism at convergent plate boundaries is a powerful contributor of volcanic gases into the atmosphere, mainly H 2 0, CO 2 and S. At mid-ocean ridges the uprise of mafic melts probably results in the addition of juvenile gases - including water - into the ocean, although a large amount of sea water is consumed during the penetration of the water in the crust and subsequent hydration reactions. It is estimated that the total addition of juvenile water produced from volcanism and hot springs is about 0.8% of the total quantity of water circulating in the hydrosphere and atmosphere, the remaining 99.2% being recycled water (Mason 1966). 1.2.4 Water in the Crust Meteoric waters in surface materials constitute groundwater, and their infiltration is controlled by porosity and permeability. Below the groundwater table, rocks are almost completely saturated with water, which deeper down heats to become a brine as it dissolves more and more solids. It is likely that water penetrates to depths of at least 15 km, since geophysical data support the presence of conducting layers at depths of 15-20 km. These layers are interpreted to be an effect offree water in the deep crust, present either as intergranular films, or as water contained in major crustal discontinuities (e.g. underthrust sheets), or both (Taylor and McLennan 1985). The abundance of water at great depths in the crust was confirmed by a Russian drilling programme in shield areas. In the Kola peninsula, for example, zones of hydraulic fracturing and mineralising fluids have been observed at depths 14 Water and Solutions of between 4 and 9 km (see Chap. 2). In the absence of magmatic activity, the origin of this water is thought to be dehydration in the high temperature environment of amphibolite to granulite facies metamorphism in the lower crust. In oceanic areas water probably penetrates to depths of about 8 km. 1.3 Solutions Ore-forming fluids may occur as molecular solutions, or colloidal solutions. Hydrothermal fluids are generally assumed to be liquid solutions in which water is the solvent. A true, or molecular solution is defined as a homogeneous system containing a dissolved substance (solute) distributed uniformly in a dissolving substance (solvent). Solutions can exist in all three states of matter: gas, solid or liquid. All gas mixtures are solutions (air, for example); whereas solid solutions are characterised by the substitution or interposing of foreign atoms, or ions, in the lattice sites, or between the lattices of mineral-forming elements. Solids, or gases become solutes when dissolved in liquids. If both components are liquid, it is usual to call the more abundant component the solvent. A solution is said to be saturated if it is in equilibrium with the undissolved solute. Therefore, an undersaturated solution is characterised by a lower concentration of solute with respect to the saturated solution; a supersaturated solution, on the other hand, contains more than the equilibrium concentration of solute. This latter type is usually unstable. There are two types of solute in water: non-electrolytes and electrolytes. In nonelectrolytic water solutions the solute dissolves as molecules and therefore they do not conduct electric current. Examples of non-electrolytic solutions are methyl alcohol, or sugar in water. Electrolytic solutions conduct electric current through the dissolved substances which are present as electrically charged ions. NaCI is a typical example which, when dissolved in water, dissociates into its constituent ions: NaCI(s) Na+ (aq) = + Cl- (aq). Other examples are: Na2S04(S) HCI (aq) = = 2Na+(aq) H+ (aq) + S042- (aq); + Cl- (aq). The concentration of solutes in aqueous solutions is expressed as moles of solutes per litre of solution, defined as molarity or molar concentration, and indicated as "M". A mole is the number of a given element's molecules, or atoms, contained in its unit molecular or atomic mass. For example, one mole of O2 contains 6.022 X 1023 O2 molecules (Avogadro's number) in exactly 32 g (atomic mass 16 X 2), and one mole of H 20 contains 6.022 X 1023 H 20 molecules in 18.02 g of H 20 (2.016 + 16.00 atomic masses of H2 and 0 respectively). A one-molar solution (1 M) therefore contains one mole (1 mol) of solute per litre of solution. A colloidal solution is made up of a dispersed phase, solid, liquid or gas, and diffused in a medium which can be either solid, liquid or gas. A review of the general Solubility and Boiling 15 properties of colloidal systems can be found in Krauskopf (1979, p.120-139), from whom the following is summarised. In ore-forming colloidal systems, solid phases are dispersed in liquid (sol), or gaseous media. The size of the particles is in the range of 10-3 to 10-6 mm (1000 to 111m). Colloidal particles have a very high surface area to volume ratio and behave as electrically charged particles. Their charge is due to adsorption of ions, making colloidal particles repel each other, so that settling is difficult, or impossible. The addition of an electrolyte to the system will neutralise the colloidal particle to enable flocculation (precipitation). When a colloid containing a certain adsorbed ion on its particles, is added to an electrolyte solution consisting of different ions, some of these electrolytic ions will be adsorbed, displacing the original ions, which will then pass into the solution. This is the phenomenon of ion exchange, and an important mechanism in explaining the distribution of cations between solutions and wall rocks. The high concentration of metals in manganese ores can be attributed to adsorption of cations on manganese dioxide sols. Ions that are also commonly adsorbed are H+ and OH-. In general, sulphides, silica and manganese dioxide sols are negatively charged, whereas oxide and hydroxide sols are positively charged. Temperature is an important parameter for colloids in hydrothermal solutions. High temperatures favour flocculation because increased kinetic movements allow the particles to come into contact with one another, and the larger particles thus formed will flocculate (although this is not always true, because gold sols are stable at temperatures of up to lSO°C). On a geological time scale colloids are unstable and tend to crystallise, so that the materials examined by the geologist may not show their pristine colloidal state. Amorphous or cryptocrystalline substances with colour banding may have had a colloidal origin. Certain textures of metallic ores in hydrothermal veins and some finely crystalline ores in open cavities are also thought to form from colloidal solutions. Silica in hot springs may also be of colloidal origin. At the Sleeper bonanza Au deposit in Nevada (USA) evidence of colloidal solutions, responsible for the transport of Au and silica, is provided by alternating bands of barren colloform quartz and mineralised opaline quartz (Saunders 1990). 1.4 Solubility and Boiling Boiling and solubility are instrumental in the separation of components from the solutions, and therefore of great importance in the deposition of ores from hydrothermal fluids. The solubility of ionic substances is the result of interaction between polar H 20 molecules and the dissolving ions. Solubility depends on: 1) the force of attraction between H 20 molecules and the ions of the solid, which tends to bring the solid into solution; and 2) the force of attraction between oppositely charged ions which will tend to prevent the solid from going into solution. Generalized solubility rules are as follows (Masterton et al. 1981): 16 Water and Solutions N0 3ClS02C02OHS2- All nitrates are soluble. All chlorides are soluble, except for AgCl, Hg 2C1 2. All sulphates are soluble except CaS04, SrS04, BaS04, PbS0 4, Ag 2S04. All carbonates are insoluble except those of Group elements (II, Na, K, etc.). All hydroxides are insoluble except Group I elements, Sr(OH)2, Ba(OH)2. All sulphides are insoluble, except Groups I and II elements. The above solubility rules illustrate that certain substances are insoluble, particularly the base metal sulphides. This is reiterated by Krauskopf (1979), in what he calls the "dilemma" of hydrothermal solutions. He writes that "most of the familiar ore metals - lead, zinc, copper, silver, molybdenum and mercury - occur directly as sulphides, or as complex minerals containing arsenic or antimony in addition to sulphur. Geological occurrences indicate that these compounds were deposited from solutions in water at moderate temperatures; yet these same compounds in the laboratory are among the most insoluble with which a chemist has to deal. What sort of solutions can be imagined to carry these very insoluble substances in nature?" (Krauskopf 1979, p.393). This question is considered further when we take a look at metal transport and the effects of ligands in Chapter 2. Both temperature and pressure have major effects on solubility. Usually the dissolution of a solid is an endothermic process, brought about by the necessary absorption of heat to break down the crystal lattice. Hence an increase in temperature (~H>O) increases the solubility and the reaction given below moves to the right: solid + solvent = solution; ~H>O. Dissolving a gas into a liquid evolves heat (exothermic) gas + liquid = solution; (~H <0), so that: ~H <0. This means that gases become less soluble as temperature increases. Pressure can also have major effects on gas-liquid systems: for a given temperature an increase in pressure increases the solubility of a gas. A pressure rise increases the concentration of molecules in the gas phase, which in turn is counteracted by more molecules entering the solution. The equilibrium of the second equation above also moves to the right. Conversely a drop in pressure will cause the gas to exsolve from the solution. Boiling of a common liquid occurs at a temperature at which its vapour pressure equals the pressure above it. Boiling of a hydrothermal solution takes place for the same reason, with the immediate result that dissolved gases and other volatile compounds, such as CO 2 and H 2S, are removed from the solution. The process of boiling of hydrothermal solutions is important because it results in the precipitation of ore elements (e.g. Au, As, Sb , Ag etc.). 1.5 Acid-Base Nomenclature A substance is said to be an acid if it produces hydrogen ions (H+) in water solution, and a base if it produces hydroxide ions (OH-). This terminology is not to be Structure of Water - Hydrolysis and Hydration 17 confused with the geological usage of acid, basic and alkaline. Silica, being the principal oxide constituent of all rock systems, serves as a basis to determine the acidity of a rock. Thus a high concentration of Si02 (> 66 wt. %) defines an acid rock. An intermediate rock contains between 52 and 66 wt. % Si0 2 , whereas values ofless than 52 wt. % define basic rocks. The basicity of a rock can also be determined by the abundance of metal oxides, such as MgO and FeO; the alkalinity of rocks by the oxides ofNa and K. The above geological usage originated many years ago when Si0 2 was considered an acid oxide, and MgO,FeO etc. as basic oxides. This terminology. is well-entrenched in the current geological language and is considered useful for practical purposes. Water, whether pure or in solution, dissociates into a hydrogen cation (H+) and a hydroxyl anion (OH-). Tlfe dissociation constant of water (Kw) at 25°C is given as: Kw = (H+) X (OH-) = 1.0 X 10- 14 . The product (H+) X (OH-) is constant for all water solutions, which at 25°C is always 1.0 X 10- 14 ; thus the concentration ofH+ or OH- can be calculated and the acidity or alkalinity of a solution specified (Krauskopf 1979). A neutral solution has equal concentrations ofH+ and OH-, which is 10-7 MH+. A strong acid has 1 MH+, and a strong base 10- 14 MH+.1t is more convenient to express acidity in terms of pH rather than H+, and this is calculated using a negative logarithmic exponent of the hydrogen concentration according to the formula: pH = -IOglO [H+] = 10glO 1/[H+]. Thus a strong acid has a pH of 0, a neutral solution 7, and a strong base a pH of 14. 1.6 Structure of Water - Hydrolysis and Hydration The structure of the water molecule is of special importance to its chemical behaviour, as discussed in Franks (1982), Neilson and Enderby (1986) and Brimhall and Crerar (1987). It has been determined that the structure of water-ice is analogous to that of tridymite, and that the oxygens in ice are tetrahedrally coordinated (the tetrahedron being the building block of all silicate minerals in the Earth's crust). Therefore water appears to have a pseudo-crystalline arrangement similar to the quartz structure, which possibly serves to explain its higher density relative to ice (Paton 1978). The oxygen ion in a water molecule is much larger than the hydrogen ion, with the result that the molecule can be described as a sphere, which - though neutral- has a positive charge on the side of the two hydrogens and a negative charge on the side of the oxygen. Consequently, the isolated molecule has a polar character and behaves in solution like a small magnet. This polar character is the key to the hydration and hydrolysis of silicate minerals, and thus important in weathering and hydrothermal alteration processes. In the process of hydration water molecules are attracted to and - by virtue of their polar charges - become orientated around other attracting ions, forming hydration shells (Brimhall and Crerar 1987). Dissolution occurs when successive 18 Water and Solutions layers of water molecules completely surround the ion. Polar water molecules can enter crystal lattices and orientate themselves against charged mineral surfaces. Where these molecules come into mutual contact, a lubricated surface is produced, for example clays which become slippery when wet. Water of crystallisation is part of the mineral lattice, as in gypsum (CaS04.2H20), or zeolitic minerals. On heating, this water is given off without breaking the lattice, and can be restored to reform the mineral. Hydrolysis is the effect of the dissociation of water molecules into H+ and OHions. The process of hydrolysis is responsible for the breakdown of silicate minerals and involves the addition of H+ and OH- to bonding sites in the mineral lattice. Hydrolysis is defined as the reaction between water and the ion of a weak acid or a weak base (Krauskopf 1979, p.37). An example of hydrolysis is the reaction of fayalite with water at neutral pH: 2FeSi0 4 + 4H 20 = 2Fe 2+ + 40H- + 2H 2Si04. Hydrolysis reactions tend to be accelerated under conditions of low pH, as for example in the vicinity of oxidising sulphide ore bodies. The presence of H+ ions in acid waters promotes the attack on silicate minerals resulting in the liberation of cations. These cations may remain in the vicinity and become fixed as stable secondary mineral assemblages, while others go into solution and are transported elsewhere. The mobility of these cations under different physico-chemical conditions has important implications for exploration geochemistry and for the evaluation of gossans (Reynolds 1982). Thus hydrolysis is dependent on the concentration of H+ ions, and any process which affects their concentration will also affect the speed and intensity of the hydrolysis process. The breaking of molecular bonds by water in a silicate melt, and consequent lowering of viscosity and consolidation temperature is a similar process. Hydrolysis of silicate minerals is very important for hydrothermal alteration, because the hydrogen ions penetrate the silicate lattices where they compete with cations (K, Ca, Na etc) to attach themselves to oxygen ions. The larger concentration of the charge in the hydrogen ions predominates, resulting in the displacement of the cations which are transferred from the silicate into the solution, while H+ enters the silicate structure, producing in it drastic changes which convert the old silicate into a new mineral, such as sericite or illie. In silicate minerals the presence of the various ions tends to consume more hydrogen than hydroxyl ions, so that if the reaction is to proceed more H+ ions have to be supplied. Volcanic gases (HCl, H 2S) can provide these H+ ions and are major acidifying agents during processes of hydrothermal alteration. 1.7 Redox Potential (Eh) An oxieation-reduction reaction is one in which transfer of electrons takes place from one element to another. The process by which certain elements lose electrons while others gain them can be quantitatively measured. This can be done by 19 Redox Potential (Eh) assigning a potential difference to half-reactions, choosing as a standard the halfreaction of the hydrogen couple: H2 = 2H++2e-. To this half-reaction potential is assigned the arbitrary value of 0 (E= 0.00 V). If the sign is positive, the half-reaction has a greater tendency to release electrons than does the reaction shown above, and the converse is true if the sign is negative. Thus, Cu with an oxidation potential of +0.16 displaces Ag with a potential of +0.22, whereas Zn (-0.76) will displace Cu (+0.16). This means that Zn will release electrons to both Cu and Ag, but Cu only to Ag. Tables of standard oxidation potentials can be found in most chemistry textbooks. The term redox potential is used as a synonym for oxidation potential and is given the symbol Eh. The redox potential is an important parameter for characterising aqueous solutions and the oxidising and reducing conditions of a geological environment. Conventionally, it is assumed that a high Eh indicates an oxidising system, whereas a low Eh indicates a reducing system. Oxidation involves the loss of electrons, and reduction the gain of electrons. If a copper wire is immersed in a solution of silver nitrate (AgN0 3), the copper will dissolve and metallic silver will precipitate. The copper wire releases electrons (e-) which combine C18 with silver cations in the solution to precipitate neutral metallic silver. This reaction involves oxidation-reduction (redox) and is shown thus: Cu(s) = Cu+2+2e-; 2Ag+(aq)+ 2e- = 2Ag(s). ® ® Eh 0·8 -0.2 -0·1 0 0·1 0·2 03 0-4 Eh Cu+ solubility Cu2+solubility 0·4 Environment in contact 'With atmosphere '-...,. .......... ........ 2 3 4 5 Euxinic marine Ground water Ocean water Rain water Pb gossan ....... 0·0 ~2 H2 ........ Environment isolated '- from atmosphere -0·4 ........ 0 2 4 6 8 10 ........ 12 14 pH Fig.1.4A, B. Examples ofEh-pH diagrams. A Solubility ofCu as a function ofEh values (After Rose et al. 1979). B Eh and pH limits of some natural environments (After Krauskopf 1979; Rose et al. 1979) 20 Water and Solutions The net reaction is: Cu(s) + 2Ag+(aq)= Cu+ 2(aq) + 2Ag(s). The first equation above is a half-reaction in which electrons are lost, and is therefore oxidation; the second equation is the half-reaction in which electrons are gained, and is reduction. The net result, expressed in the third and last equation, is therefore the oxidation-reduction reaction, in which Cu is the reducing agent and is oxidised, and Ag is the oxidising agent which is reduced. The same can be said for oxidation numbers, where a fictitious charge is given to an ion. Thus, for the net reaction above, Cu is oxidised because its oxidation number is increased from 0 to +2; reduction, on the other hand, results from a decrease in oxidation number, and so Ag is reduced with a decrease in its oxidation number from +2 to O. Eh and pH are two important variables which control the mobility of elements in a geological system, and consequently Eh-pH diagrams are useful to show conditions under which processes of oxidation and reduction can occur in nature. Examples of Eh-pH diagrams are given in Fig. 1.4. 1.8 Chemical Potential, Chemical Activity and Fugacity The terms chemical potential, chemical activity and fugacity are frequently used in discussions of hydrothermal alteration involving aspects of physical chemistry, chemical equilibria and thermodynamics. The chemical potential indicates the capacity of one compound to react with another, and represents a particular energy level for the given compound in a phase. The chemical potential is analogous to gravitational potential energy, in which the stable state is at the lowest potential. The chemical potential can be expressed by the formula: Ila = (dG/dXa)p,T,Xb,X c , etc. In the equation above, the chemical potential Il of a substance a equals the change of an extensive state function G (internal energy, volume, mass) resulting from the addition of small amounts of a, while P,T and all other substances b, c etc remain constant (Best 1982). In a hydrous silicate melt at equilibrium the chemical potential Il of water equals the potential of water vapour, which in turn equals the potential of water in the hydrated minerals (Best 1982). The activity of a chemical species is denoted by "a" and is approximately proportional to concentration and is expressed as moles per litre. In solid solutions, or mixtures of liquids, chemical activity is related to the mole fraction of the constituent. In gases, chemical activity can be approximated by partial pressure (Rose et al. 1979), and is expressed by: a=m, where m is the concentration and a is the activity coefficient (Hemley and Jones 1964). References 21 Fugacity is a measure of the escaping tendency of a vapour or gas and its departure from ideal behaviour. For example the tendency of liquid water to vaporise or escape into the vapour phase. In a mixture of gases the fugacity f of a gas is proportional to its mole fraction X, or concentration in the mixture, and to the total pressure, according to the relation: Fugacity is not restricted to gases, but is also accepted as a measure of flow of matter from high to low potentials of solids and liquids. For a more detailed and comprehensive treatment of this topic the reader should consult Carmichael et al. (1974, p.292). References Best M G (1982) Igneous and metamorphic petrology. Freeman, San Francisco, 630 pp Brimhall G H, Crerar D A (1987) Ore fluids: magmatic to supergene. In: Carmichael I S E, Eugster H P (eds) Thermodynamic modelling of geological materials: minerals, fluids and melts. Reviews in mineralogy, vol 17. Min Soc Am, pp 235-321 Carmichael I S E, Turner F J, Verhoogen J (1974) Igneous petrology. McGraw-Hill Book Co 739 pp Franks F (ed) (1982) Water. A comprehensive treatise. Plenum, New York, 484 pp Frey H (1977) Origin of the Earth's ocean basins. Icarus, 32:235-250 Froude D 0, Ireland T R, Kinny P D, Williams I S, Compston, W, Williams, I R, Myers, J S (1983) Ion microprobe identification of 4,100-4,200 Myr-old terrestrial zircons. Nature (London) 304: 616-618 Fyfe W S (1978) The evolution of the Earth's crust: modern plate tectonics to ancient spot tectonics? Chern Geol 23:89-114 Gill R (1989) Chemical fundamentals of geology. Unwin Hyman, London, 291 pp Hemley J, Jones W R (1964) Chemical aspects of hydrothermal alteration with emphasis on hydrogen metasomatism. Econ Geol 59:538-569 Holland H D (1984) The chemical evolution of the atmosphere and oceans. Univ Press, Princeton, 582 pp KrauskopfK B (1979) Introduction to geochemistry, 2nd edn. McGraw-Hill Kogakushu New York, 617 pp Kroner A (1985) Evolution of the Archean continental crust. Annu Rev Earth Planet Sci 13:49-74 Lewis J S, Prinn R G (1984) Planets and their atmospheres. Origin and evolution Academic Press, New York London, 470 pp Lowman P D (1989) Comparative planetology and the origin of continental crust. Precambrian Res 44:171-195 Mason B (1966) Principles of geochemistry, 3d edn. John Wiley & Sons, New York, 329 pp Masterton W L, Slowinski E J, Stanitski C L (1981) Chemical principles, 5th edn. Holt-Saunders international editions, Saunders, Philadelphia, 641 pp Neilson G W, Enderby J E (1986) Water and aqueous solutions. Hilger, Bristol, 349pp Paton T R (1978) The formation of soil material. Allen & Unwin, 143 pp Pollack J B, Yung Y L (1980) Origin and evolution of planetary atmospheres. Annu Rev Earth Planet Sci 8:425-487 Press F, Siever R (1982) Earth, 3d edn. Freeman, San Francisco, 613 pp Reynolds I M (1982) Weathering of ore minerals and textural evaluation of gossans. Unpublished course notes. Rhodes Univ Press Grahamstown, S Afr, 183 pp Rose A W, Hawkes H E, Webb S (1979) Geochemistry in mineral exploration, 2nd edn. Academic Press, New York London, 657 pp 22 Water and Solutions Saunders] A (1990) Colloidal transport of gold and silica in epithermal precious metal systems: evidence from the Sleeper deposit, Nevada. Geology 18:757-760 Sheppard S M F (1986) Characterisation and isotopic variations in natural waters. Reviews in mineralogy, vol 16. Min Soc Am, pp 165-183 Spencer R J, Hardie L A (1990) Control of sea water composition by mixing of river waters and midocean ridge hydrothermal brines. Geochem Soc Spec Publ 2: 409-420 Taylor S R, McLennan S M (1985) The continental crust: its composition and evolution. Blackwell Oxford, 312 pp Whitfield M (1982) The salt sea-accident or design? New Sci 94:14-17 Chapter 2 Hydrothermal Solutions 2.1 Introduction A hydrothermal fluid is defined as a hot (ca. 50 - >500°C), aqueous solution (hydro = water; thermal = hot), containing Na, K, Ca, Cl as major components, as well as many other elements (e.g. Mg, B, S, Sr, CO 2 , H 2S, NH 4 , Cu, Pb, Zn, Sn, Mo, Ag, Au etc.) as minor constituents (Skinner 1979). The terms fluid and solution are here used interchangeably, although fluid in the strict sense refers to a phase at a supercritical temperature in which a liquid can no longer exist. The pressure required to cause condensation at a given critical temperature is called the critical pressure. There are certain known conditions in nature where a hydrothermal solution is in fact a fluid. We begin this chapter by having a look at those kinds of waters that may constitute a hydrothermal solution. A brief introduction is given on oxygen and hydrogen isotope systematics and their role in understanding the nature and interactions of hydrothermal fluids and wall rocks. As the only direct observation and sample material of hydrothermal solutions come from fluid inclusions and hot springs, an outline of the characteristics and methodology involved in the study of fluid inclusion is also given. It is emphasised that fluid inclusions provide important insights into the physical and chemical environments of ore deposition. This leads us to a discussion of dissolved constituents, where a series of tables provides some understanding of the composition of natural solutions. The transport of apparently insoluble metals had long baffled chemists until the role ofligands in the dissolution and transport of ore elements - and hence its importance in the study of hydrothermal mineral deposits - was discovered. Various mechanisms of deposition of metals from solutions conclude this chapter. 2.2 Water of Hydrothermal Solutions The water of hydrothermal solutions can be derived from the following sources: meteoric, sea water, connate, metamorphic,juvenile or magmatic. Most hydrothermal solutions are of mixed origin in which one or more of the above sources can predominate. Meteoric waters include rain water, lake and river waters, and groundwaters. These waters can penetrate deep into the crust and may become heated and 24 Hydrothermal Solutions mineralised, thereby acquiring the properties of hydrothermal solutions. Stable isotope systematics indicate that in volcanic regions the waters of hot springs and pools are largely, if not exclusively, of meteoric origin. Oceanic crust, in and around mid-ocean ridges, allows penetration of sea water to several kilometres below the sea floor. As a result, the sea water is heated, transformed into a fluid highly enriched with metals and driven by convection, and is subsequently discharged at the sea floor as a submarine mineral-depositing hot spring (Bonatti 1975). Water trapped during the deposition of sediments and produced during diagenetic reactions is known as connate water, or formation water. It is calculated that some 20% by volume of unmetamorphosed sediments in the Earth's crust consist of pore water (Hanor 1979). It is also widely recognised that hydrothermal fluids can develop during burial diagenesis, and attain high salinities and temperatures. This water is essentially unbound water, that is, not bound in the lattice of rock-forming minerals. Removal of interlayer water from clays, gypsum and organic matter is an important aspect of diagenetic processes. Expulsion of fluids during burial and diagenesis takes place as a result of reduction of porosity, and the volume of water released can be considerable. An average shale can yield 3.5 X 10 3 litres of water for every I m 3 of solid material deposited. Temperatures during diagenesis may range from a few degrees below O°C up to 250-300°C, but water expulsion during diagenetic changes occurs at temperatures between 90 and 120°C. (Hanor 1979). During the evolution of a sedimentary basin, expulsion of fluids takes place and they migrate upward and towards the margins of the basin, and, where sulphur is added from any source, precipitation of sulphides may occur at favourable sites. Metamorphic waters are derived from the dehydration of hydroxyl-bearing minerals (bound water) through rising pressure and temperature (metamorphic dewatering). Smirnov (1976) envisaged the presence of a hydrothermal zone ahead of the metamorphic front during prograde regional metamorphism. Henley et al. (1976) suggested that metamorphism could be a source of energy leading to hydrothermal vein mineralisation. The presence of volatile-rich fluids liberated during metamorphism is commonly accepted, and they can be considered as dilute brines generally containing H 20, CO 2 and CH 4 . Direct observation of waters of probable metamorphic origin comes from the Kola peninsula, where the world's deepest well was drilled to a depth of 12 000 m. Between 4500 and 9000 m the well intersected a zone of "disaggregated metamorphic rock", within which abundant hot and highly mineralised water was found. At such depths, this water is likely to be of metamorphic origin released from bound water in rock-forming minerals. Also present in the well were abundant gases such as He, H 2 , N 2, CH 4 , CO 2 and various hydrocarbons, most of which are also believed to have been liberated by deep metamorphic processes (Koslovsky 1984). It is interesting to note that the dehydration of rocks at those depths is accompanied by microfracturing or hydraulic disaggregation. Temperatures of around 180°C were measured at 10 000 m, whereas sulphur isotopic determinations indicate an origin of the sulphur from the mantle. A surprising finding from these results is that the well intersected the Conrad Discontinuity (where density and velocity of seismic waves rise from 2.5 to Oxygen and Hydrogen Isotope Systematics of Hydrothermal Fluids 25 2.75 g/cm 3 and 5 to 6 km S-1 respectively) which is taken to mark the transition from upper (granitic) to lower crust. This discontinuity, however, is found to be the lower boundary of the disaggregation and hot metamorphic water zones, where the rock returns to its normal density for that depth and water is no longer present. It was further established that in the zone of disaggregation, rock fragments are cemented by sulphides of Cu, Zn, Ni, Fe and Co. These discoveries indicate that gas-waterrock interactions are active at great depths, and that conditions exist for the formation of hydrothermal ore deposits entirely powered by metamorphism (see Chaps. 3 and 15). Some geoscientists consider juvenile water as that which is assumed to be derived from the mantle. Magmatic waters are those that separate from melts upon cooling, generating a magmatic-hydrothermal system, which is perhaps the most powerful ore-depositing agent. In addition to water, other volatile constituents that may be present in magmas include H 2S, CO 2, S02, HCl, B, F, H 2 . Water contents of magmas range from as little as 0.2 % to as much as 6.5 % by weight. The presence and abundance of these volatiles in a magma is generally related to its composition and the source region from which it originated. It is possible, for example, that a B-rich granite may have derived from melts that have interacted with tourmaline-rich metasediments. Further, the relative abundance of one volatile over another (e.g. B/F) may be very important in determining the type of ore deposit that may be formed on exsolution of the volatile fractions from the late stages of consolidation. The evolution of a magmatic-aqueous phase during the solidification of a hydrous magma is discussed in some detail in Chapter 3. 2.3 Oxygen and Hydrogen Isotope Systematics of Hydrothermal Fluids Analyses of stable (non-radiogenic) isotopes provide useful data for the understanding of the origin, nature and evolution of hydrothermal fluids and their interaction with wall rocks. Since oxygen is the most abundant element, and hydrogen its companion in the water molecule, the ratios of 18 0/ 16 0 and D/H where D is deuterium elH), the heavy isotope of hydrogen, constitute powerful indicators or tracers of the nature and source of hydrothermal fluids. The theoretical, experimental aspects and reviews of studies of stable isotope systematics can be found in a number of excellent publications. The interested reader is advised to consult Valley et al.(1986), Taylor (1979), Faure (1986) and Bowen (1988). The isotopic values are reported as deviations from an arbitrary standard, which in the case of hydrogen and oxygen is taken to be oceanic water (SMOW, standard mean ocean water) with 8 18 0=0 and 8D=0. This deviation, indicated by 8 is given by the formula: 8= (Rsample/Rstandard - 1) X 1000 where R is the ratio of the isotopes being considered (e.g. D/H and 18 0/ 16 0). Thus, positive or negative values represent enrichment or depletion with respect to the standard. 26 Hydrothermal Solutions Sea water, meteoric water and juvenile water are considered to be "reference waters", in that they have defined isotopic characteristics at their source. Consequently, geothermal, connate, metamorphic and magmatic waters are re-cycled from one or more of these reference waters (Ohmoto 1986). In the study of hydrogen and oxygen isotope systematics, it is important to take into account the "meteoric water line" (MWL). Isotopic variation of meteoric waters is dependent on latitude and elevation, with 8D and 8 18 0 values being lower towards higher elevations and higher latitudes. This is due essentially to less evaporation and therefore less vapourfluid fractionation, which tends to favour the partitioning ofthe lighter isotopes into the vapour phase, leaving the fluid enriched in the heavy isotopes. Thus, water condensing from a vapour is enriched in 18 0 and D relative to the vapour, so that air masses become progressively depleted in the heavy isotopes relative to the initial condensate (Taylor 1979). The linear relationship between 8D and 818 0 is defined by: Studies of the isotope characteristics of natural waters by Taylor (1979) and Sheppard (1986); Fig. 2.1) reveal that present-day ocean waters are relatively uniform with 8D of +5 to -7 per mill and 8 18 0 of +0.5 to 1.0 per mill, and with mean values close to 0 per mill for both D and 18 0 (SMOW standard). Isotopic values of ancient sea water, as determined through the analysis of minerals that equilibrated with fossil sea water hydrothermal systems, show a progressive increase o "(3)- -, OCEAN \-IATERS (SHOW) - _ ...... METAMORPHIC \-lATER ( 300 - 6000 () ,..o -50 cf!. / Cl / 'tQ / SEDIHENTS -100 ./ ./ ./ ./ -10 0 +10 +20 +30 60%0 Fig. 2.1. Fields of isotopic compositions of natural waters (primary magmatic, metamorphic, ocean waters). Long dashed lines are fields of formation waters in sedimentary basins: 1 Alberta basin; 2 California basin; 3 Gulf Coast basin. Dashed box partly superimposed on the box of primary magmatic water, represents "juvenile water" (After Ohmoto 1986). The diagram is taken and modified from Sheppard (1986) Oxygen and Hydrogen Isotope Systematics of Hydrothermal Fluids 27 of 018 0 of the sea water since Archean times (-8 to -12 per mill). From 2.5 Ga ago it is calculated that the 018 0 values were between 0 and -3 per mill, and oD of 0 to -25 per mill. Although in no instance can it be stated that truly juvenile water (or for that matter magmatic water) has ever been identified, the nearest D and 18 0 values are obtained from unaltered igneous and mantle rocks and minerals, using mineralwater isotopic fractionation factors at high temperatures (>700°C). These give oD values of approximately -50 to -90 per mill and 018 0 of +5.5 to + 10 per mill (Sheppard 1986). The isotopic composition of geothermal waters indicates in all cases a strong meteoric component. The oD values of geothermal waters are more or less the same as those of the local meteoric waters. Modern geothermal waters show an enrichment in 018 0 80 shift) due to isotopic exchange with wall rocks. The amount of the 018 0 shift increases with temperature and salinity of the fluids, and this probably means that the fluids acquired their heat, and dissolve solids, by interaction with deeply buried and hot rocks (Ohmoto 1986). The 018 0 enrichment is accompanied by a depletion of 18 0 in the wall rocks. We return to this important aspect of isotope systematics in Chapter 4 with a discussion of the isotopic effects during hydrothermal alteration processes. Connate and formation waters show a wide range of 018 0 and oD. The D values of connate waters show a decrease with latitude, as with meteoric waters. The variations in oD and 018 0 of pore fluids and connate waters are largely dependent on whether the original fluid was sea water or meteoric water. Generally, in pore fluids of oceanic rocks, there appears to be a decrease in 0 18 0 with increasing depth and salinity. Measurements taken from samples of ocean floor suggest that alteration of basalt with sea water is responsible for this isotopic characterisation of pore fluids (Ohmoto 1986). The oD and 018 0 values obtained from sedimentary basin fluids by contrast show a wide range of values (oD = +20 to - 150 per mill; 018 0 = + 10 to -20 per mill ). Also, in basin brines there seems to be an overall increase in the isotopic values with salinity and temperature, so that brines with the lowest temperature approach the isotopic composition of present-day local meteoric waters. Metamorphic waters, like juvenile water, are still somewhat unconstrained istopically. They are derived from dehydration of mineral phases during prograde regional metamorphic events, as defined in Section 2.2. Nevertheless, here too the isotopic composition is clearly dependent on the orginal rock type and its history of interaction with fluids. Ohmoto (1986) points out that if volcanic rocks are first altered by sea water, at temperatures less than 200°C, and then undergo metamorphism, their final 018 0 values could be as high as 25 per mill. On the other hand, if volcanic rocks are altered by meteoric waters, then subjected to high-grade metamorphism, the oD and 018 0 values could be much lower. Metamorphic waters have a range of oD values from 0 to about -70 per mill, and a range of 018 0 from + 3 to about 25 per mill. Taylor (1979) attributes the wide range of 018 0 to the original values in the metamorphosed rocks. Magmatic waters may have a range of oD and 018 0 values, depending on the source region of the magma in question (1- or S-type for example), as well as the possible isotopic exchanges with wall rocks during the cooling stages of the melt. Isotopic values, determined on the basis of experimental studies of fractionation factors between melts and water, give a range for magmatic waters ofoD = -30to -75 per mill and, 018 0 = +7 to +13 per mi1l (Ohmoto 1986). e 28 Hydrothermal Solutions 2,4 Fluid Inclusions Fluid inclusions are droplets of fluid trapped in crystals at the time of their growth, or introduced along microcracks and cleavages after crystallisation of the host mineral. They represent samples of hydrothermal fluids, and range in size from a single water molecule up to several millimetres, with an average of about 0.01 mm (Roedder 1979). The number of inclusions in anyone crystal can be very high, with a maximum of 109 inclusions having been estimated in 1 cm 3• Much interest has been directed to the subject of fluid inclusions over the years, which is covered in a wide number of publications. The present discussion is based largely on the works of: Roedder (1979), from whom a concise review is gained, Guilbert and Park (1985), who provide a general introduction; and the text books by Shepherd et al (1985), Hollister and Crawford (1981) and finally Roedder (1984). These all provide essential reading for the exploration geologist who wishes to acquire a more indepth understanding of the topic. Fluid inclusions have many practical uses for the student of ore deposits, providing information on the temperature, pressure, density and composition of the fluids that originated the mineralisation. In terms of their origin, three types of fluid inclusions are recognised, namely: primary, secondary and pseudo-secondary (Fig. 2.2). Primary inclusions are those that became trapped during the growth of the host mineral, and are therefore associated with crystallisation features - such as growth zones; or they may occur isolated due to imperfections of the crystal during its growth (Fig. 2.2). Secondary inclusions are those that form after the growth of the host mineral is completed. They cut across growth zones and even crystal boundaries, and may represent infilling of microcracks by late fluids that might well be unrelated to the ore-forming event. Pseudo secondary inclusions form during the two stages outlined above, and are characterised by their alignment with microcracks that end against a growth zone (Fig.2.2). These definitions are somewhat s / o~ " . • ps-. v ;.~ %0 •• o~, ., //1 - /0" .,s /J s / ., (If'", ,0" • '.' 0 ~ ,'. 0 ' 0 O ',0, , 0 0', P~ 0 o, 0 , o '.0 C"I: v <9 (,., 0' 6" ps : ,e' e 0 0 01 01 01 1 81 ' 01 ':.1 0," 00, ,. 01 (I @' I ~ \' to pi GROWTH ZONING ,p PS S PTI Fig. 2.2. Types of fluid inclusions, and their distribution in a quartz crystal. P Primary; S secondary; PS pseudosecondary (After Shepherd et al. 1985) 29 Fluid Inclusions TYPE I: LIQUID (L) 0 ~0...!l0P.9~e _ _ _ _ _ _ 0 0 ___ _ TYPE II: LIQUID >50%(L) ~ (Q) +VAPOUR <50% (V)\,Q/'~ two- phase v 1--- - - - - - - - - - - TYPE III: VAPOUR >50%(V) e--L~ + LIQUID <50% (L) v~ v two-phase L 1-------------TYPE IV: VAPOUR ONLY (V) Q 0 monophase 0 i--- ----------- TYPE V: LIQUID :!: VAPOUR~s ~s + SOLIDS (5) ~50%~ ~L multiphase ~ v _ _ _ _ _ _ _ _ _ _ 2.... __ _ Fig. 2.3. Classification of fluid inclusions observed at room temperature (After Shepherd et al. 1985) TYPE VI: IMMISCIBLE L1 ~L/~L1 LIQUIDS L1+ L2 ~(('\v :!: VAPOUR v ~ L1 L2 simplistic, and it is often difficult to distinguish between, say, secondary and pseudosecondary, or primary and pseudo secondary inclusions. Tables of empirical criteria for the identification of the genetic types of fluid inclusions are given in Roedder (1979, 1984). In terms of morphology and contents, several important types of fluid inclusions have been described. The classification scheme reported by Shepherd et al. (1985) is summarised below, and diagramatically shown in Fig.2.3. 1. Monophase inclusions: entirely filled with liquid (L). 2. Two-phase inclusions: filled with a liquid phase and a small vapour bubble (L+V). 3. Two-phase inclusions: in which the vapour phase is dominant and occupies more than 50% of the volume (V+L). 4. Monophase vapour inclusion (V): entirely filled with a low density vapour phase (generally mixtures of H 20, CH 4 , CO 2). 5. Multiphase inclusions containing solids (S+L+j-V): contain solid crystalline phases known as daughter minerals. These are commonly hal