INTRODUCTION TO ORE-FORMING PROCESSES

Introduction to
Ore-Forming
Processes
LAURENCE ROBB

© 2005 by Blackwell Science Ltd
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Robb, L.J.
Introduction to ore-forming processes / Laurence Robb.
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Contents
Prefacevii

INTRODUCTION: MINERAL RESOURCES
Introduction and aims
A classification scheme for ore deposits
What makes a viable mineral deposit?
Some useful definitions and compilations
Natural resources, sustainability, andenvironmental responsibility
Summary and further reading

1
1
2
4
6
11
15

PART 1 IGNEOUS PROCESSES
1 IGNEOUS ORE-FORMING PROCESSES
1.1 Introduction
1.2 Magmas and metallogeny
1.3 Why are some magmas more fertile
than others? The “inheritancefactor”
1.4 Partial melting and crystal
fractionation as ore-forming
processes
1.5 Liquid immiscibility as anore-forming process
1.6 A more detailed considerationof mineralization processes inmafic magmas
1.7 A model for mineralization inlayered mafic intrusions
Summary and further reading

19
19
20

28

37
54

57

2.3

Formation of a magmaticaqueous phase
2.4 The composition andcharacteristics of magmatichydrothermal solutions
2.5 A note on pegmatites and theirsignificance to granite-relatedore-forming processes
2.6 Fluid–melt trace element
partitioning
2.7 Water content and depth of
emplacement of granites –relationships to ore-forming
processes
2.8 Models for the formation ofporphyry-type Cu, Mo, andW deposits
2.9 Fluid flow in and around granite
plutons
2.10 Skarn deposits
2.11 Near-surface magmatichydrothermal processes – the“epithermal” family of
Au–Ag–(Cu) deposits
2.12 The role of hydrothermal fluids
in mineralized mafic rocks
Summary and further reading

71
74

PART 2 HYDROTHERMAL
PROCESSES

75
75

3 HYDROTHERMAL ORE-FORMING PROCESSES
3.1 Introduction
3.2 Other fluids in the Earth’s crustand their origins

79

85

93
96

101

106
108
113

117
122
125

2 MAGMATIC-HYDROTHERMAL ORE-FORMING
PROCESSES

2.1
2.2

Introduction
Some physical and chemical
properties of water

76

129
129
130

vi

CONTENTS

3.3

The movement of hydrothermal
fluids in the Earth’s crust
3.4 Further factors affecting metalsolubility
3.5 Precipitation mechanisms formetals in solution
3.6 More on fluid/rock interaction –an introduction to hydrothermal
alteration
3.7 Metal zoning and paragenetic
sequence
3.8 Modern analogues of ore-forming
processes – the VMS–SEDEX
continuum
3.9 Mineral deposits associatedwith aqueo-carbonic metamorphic
fluids3.10 Ore deposits associated withconnate fluids
3.11 Ore deposits associated withnear surface meteoric fluids(groundwater)
Summary and further reading

138
147
153

166
174

177

189
197

209
214

PART 3 SEDIMENTARY/SURFICIAL PROCESSES
4 SURFICIAL AND SUPERGENE ORE-FORMING
PROCESSES

4.1 Introduction
4.2 Principles of chemical weathering
4.3 Lateritic deposits
4.4 Clay deposits
4.5 Calcrete-hosted deposits

219
219
220
223
233
235

4.6 Supergene enrichment of Cu
and other metals in near surface
deposits
Summary and further reading
5 SEDIMENTARY ORE-FORMING PROCESSES
5.1 Introduction
5.2 Clastic sedimentation and heavy
mineral concentration – placer
deposits
5.3 Chemical sedimentation – bandediron-formations, phosphorites,and evaporites
5.4 Fossil fuels – oil/gas formation andcoalification
Summary and further reading

238
245
246
246

247

266
287
307

PART 4 GLOBAL TECTONICS
AND METALLOGENY
6 ORE DEPOSITS IN A GLOBAL TECTONIC CONTEXT
6.1 Introduction
6.2 Patterns in the distribution ofmineral deposits
6.3 Continental growth rates
6.4 Crustal evolution and metallogenesis
6.5 Metallogeny through time
6.6 Plate tectonics and ore deposits –a summary
Summary and further reading

311
311

References
Index

345
368

312
312
315
319
339
343

Preface

There are many excellent texts, available at both
introductory and advanced levels, that describe
the Earth’s mineral deposits. Several describe thedeposits themselves and others do so in combination with explanations that provide an understanding of how such mineral occurrences form.
Few are dedicated entirely to the multitude ofprocesses that give rise to the ore deposits of theworld. The main purpose of this book is to provide
a better understanding of the processes, as well as
the nature and origin, of mineral occurrences andhow they fit into the Earth system. It is intended
for use at a senior undergraduate level (thirdand fourth year levels), or graduate level (North
America), and assumes a basic knowledge in awide range of core earth science disciplines, aswell as in chemistry and physics. Although meant
to be introductory, it is reasonably comprehensive in its treatment of topics, and it is hoped thatpracticing geologists in the minerals and relatedindustries will also find the book useful as a summary and update of ore-forming processes. To thisend the text is punctuated by a number of boxedcase studies in which actual ore deposits, selected
as classic examples from around the world, arebriefly described to give context and relevance toprocesses being discussed in the main text.
Metallogeny, or the study of the genesis of ore
deposits in relation to the global tectonic paradigm,
is a topic that traditionally has been, and should
remain, a core component of the university earthscience curriculum. It is also the discipline that
underpins the training of professional earth scientists working in the minerals and related industriesof the world. A tendency in the past has been totreat economic geology as a vocational topic andto provide instruction only to those individuals
who wished to specialize in the discipline or tofollow a career in the minerals industries. In more
recent years, changes in earth science curricula
have resulted in a trend, at least in a good many
parts of the world, in which economic geology has
been sidelined. A more holistic, process-orientatedapproach (earth systems science) has led to a wider
appreciation of the Earth as a complex interrelated
system. Another aim of this book, therefore, isto emphasize the range of processes responsible
for the formation of the enormously diverse ore
deposit types found on Earth and to integratethese into a description of Earth evolution andglobal tectonics. In so doing it is hoped thatmetallogenic studies will increasingly be reintegrated into the university earth science curricula.
Teaching the processes involved in the formationof the world’s diminishing resource inventory
is necessary, not only because of its practicalrelevance to the real world, but also because suchprocesses form an integral and informative partof the Earth system.
This book was written mainly while on a
protracted sabbatical in the Department of Earth
Sciences at the University of Oxford. I am very
grateful to John Woodhouse and the departmental
staff who accommodated me and helped to provide
the combination of academic rigor and quietitude
that made writing this book such a pleasure. In
particular Jenny Colls, Earth Science Librarian, wasa tower of support in locating reference material.
The “tea club” at the Banbury Road annexe provided both stimulation and the requisite libations

viii

PREFACEto break the monotony. The staff at Blackwell
managed to combine being really nice people witha truly professional attitude, and Ian Francis,
Delia Sandford, Rosie Hayden, and Cee Pike were
all a pleasure to work with. Dave Coles drafted
all the diagrams and I am extremely grateful for
his forebearance in dealing amiably with a list of
figures that seemingly did not end. Several people
took time to read through the manuscript for me
and in so doing greatly improved both the styleand content. They include John Taylor (copyediting), Judith Kinnaird and Dave Waters (Introduction), Grant Cawthorn (Chapter 1), Philip
Candela (Chapter 2), Franco Pirajno (Chapter 3),
Michael Meyer (Chapter 4), John Parnell andHarold Reading (Chapter 5), and Mark Barley,
Kevin Burke, and John Dewey (Chapter 6). The
deficiencies that remain, though, are entirely myown. A particularly debt of gratitude is owed toDavid Rickard, who undertook the onerous taskof reviewing the entire manuscript; his lucid
comments helped to eliminate a number of flaws
and omissions. Financial support for this projectcame from BHP Billiton in London and the Geological Society of South Africa Trust. My colleagues at Wits were extremely supportive during
my long absences, and I am very grateful to Spike
McCarthy, Paul Dirks, Carl Anhauesser, Johan
Kruger, and Judith Kinnaird for their input inso many ways. Finally, my family, Vicki, Nicole,and Brendan, were subjected to a life-style thatinvolved making personal sacrifices for the
fruition of this project – there is no way of saying
thank you and it is to them that I dedicate thisbook.
Laurence Robb
Johannesburg

Introduction:mineral resources

GENERAL INTRODUCTION AND AIMS OF THE BOOKA SIMPLE CLASSIFICATION SCHEME FOR MINERAL DEPOSITSSOME IMPORTANT DEFINITIONS
metallogeny, syngenetic, epigenetic,
mesothermal, epithermal, supergene,
hypogene, etc.
SOME RELEVANT COMPILATIONS
periodic table of the elements
tables of the main ore and gangue minerals
geological time scale
FACTORS THAT MAKE A VIABLE MINERAL DEPOSIT
enrichment factors required to make ore deposits
how are mineral resources and ore reserves
defined?
NATURAL RESOURCES AND THEIR FUTURE EXPLOITATION
sustainability
environmental responsibility

INTRODUCTION AND AIMS
Given the unprecedented growth of humanpopulation over the past century, as well as therelated increase in demand for and productionof natural resources, it is evident that understanding the nature, origin and distribution of
the world’s mineral deposits remains a strategic
topic. The discipline of “economic geology,” which
covers all aspects pertaining to the description
and understanding of mineral resources, is, therefore, one which traditionally has been, and should
remain, a core component of the university earthscience curriculum. It is also the discipline that

underpins the training of professional earth scientists working in the minerals and related industries of the world. A tendency in the past has been
to treat economic geology as a vocational topic
and to provide instruction only to those individuals who wished to specialize in the disciplineor to follow a career in the minerals industry.
In more recent years, changes in earth sciencecurricula have resulted in a trend, at least in agood many parts of the world, in which economic
geology has been sidelined.
The conceptual development of earth systems
science, also a feature of the latter years of thetwentieth century, has led to dramatic shifts inthe way in which the earth sciences are taught.
A more holistic, process-orientated approach has
led to a much wider appreciation of the Earth
as a complex interrelated system. The understanding of feedback mechanisms has brought
an appreciation that the solid Earth, its oceans
and atmosphere, and the organic life forms thatoccupy niches above, at and below its surface,are intimately connected and can only be understood properly in terms of an interplay of processes. Examples include the links between global
tectonics and climate patterns, and also between
the evolution of unicellular organisms and theformation of certain types of ore deposits. In thiscontext the teaching of many of the traditional
geological disciplines assumes new relevance andthe challenge to successfully teaching earth system science is how best to integrate the wide rangeof topics into a curriculum that provides understanding of the entity. Teaching the processes
involved in the formation of the enormously

2

INTRODUCTION: MINERAL RESOURCES

diverse ore deposit types found on Earth is necessary, not only because of its practical relevance tothe real world, but also because such processes
form an integral and informative part of the
Earth’s evolution.
The purpose of this process-orientated book is toprovide a better understanding of the nature andorigin of mineral occurrences and how they fit into
the Earth system. It is intended for use at a senior
undergraduate level (third and fourth year levels),
or a graduate level, and assumes a basic knowledge
in a wide range of core earth science disciplines,
as well as in chemistry and physics. It is also hopedthat practicing geologists in the minerals andrelated industries will find the book useful as asummary and update of ore-forming processes. To
this end the text is punctuated by a number of
boxed case studies in which actual ore deposits,
selected as classic examples from around the world,
are briefly described to give context and relevanceto processes being discussed in the main text.

A CLASSIFICATION SCHEME FOR ORE DEPOSITS
There are many different ways of categorizing ore
deposits. Most people who have written about
and described ore deposits have either unwittingly or deliberately been involved in theirclassification. This is especially true of textbooks
where the task of providing order and structure toa set of descriptions invariably involves some formof classification. The best classification schemes
are probably those that remain as independent ofgenetic linkages as possible, thereby minimizing
the scope for mistakes and controversy. Nevertheless, genetic classification schemes are ultimately desirable, as there is considerable advantage
to having processes of ore formation reflected in aset of descriptive categories. Guilbert and Park
(1986) discuss the problem of ore deposit classification at some length in Chapters 1 and 9 of
their seminal book on the geology of ore deposits.
They show how classification schemes reflect thedevelopment of theory and techniques, as well as
the level of understanding, in the discipline.
Given the dramatic improvements in the level ofunderstanding in economic geology over the past

few years, the Guilbert and Park (1986) classification scheme, modified after Lindgren’s (1933)
scheme, is both detailed and complex, and befits
the comprehensive coverage of the subject matterprovided by their book. In a more recent, butequally comprehensive, coverage of ore deposits,
Misra (2000) has opted for a categorization based
essentially on genetic type and rock association,
similar to a scheme by Meyer (1981). It is the association between ore deposits and host rock that isparticularly appealing for its simplicity, and thathas been selected as the framework within which
the processes described in this book are placed.
Rocks are classified universally in terms of athreefold subdivision, namely igneous, sedimentary and metamorphic, that reflects the fundamental processes active in the Earth’s crust(Figure 1a). The scheme is universal because rocks
are (generally!) recognizably either igneous or
sedimentary, or, in the case of both precursors,
have been substantially modified to form a metamorphic rock. Likewise, ores are rocks and canoften be relatively easily attributed to an igneousor sedimentary/surficial origin, a feature that represents a good basis for classification. Such a classification also reflects the genetic process involvedin ore formation, since igneous and sedimentarydeposits are often syngenetic and formed at thesame time as the host rock itself. Although many
ores are metamorphosed, and whereas pressure
and temperature increases can substantially modify the original nature of ore deposits, it is evident
that metamorphism does not itself represent a
fundamental process whereby ore deposits are
created. Hydrothermalism, however, is a viableanalogue in ore-forming processes for metamorphism and also involves modification of either
igneous or sedimentary rocks, as well as heat(and mass) transfer and pressure fluctuation. A
very simple classification of ores is, therefore,
achieved on the basis of igneous, sedimentary/
surficial and hydrothermal categories (Figure 1b),and this forms the basis for the structure and layout of this book. This subdivision is very similar
to one used by Einaudi (2000), who stated that allmineral deposits can be classified into three types
based on process, namely magmatic deposits, hydrothermal deposits and surficial deposits formed by

INTRODUCTION: MINERAL RESOURCES(a) Rocks3(b) Ore deposits
IgneousIgneousPart 1

Part 2

Metamorphic

Sedimentary

HydrothermalPart 3

Sedimentarysurficial

(c)

(d)

(e)
Figure 1 Classification of the principal rock types (a) and an analogous, but much simplified, classification of ore
deposit types (b). Photographs show the interplay between ore forming processes. (c) Igneous ore type: the PGEbearing Merensky Reef, Bushveld Complex, South Africa. This unit and the ores within it can be altered andredistributed by hydrothermal solutions. (d) Sedimentary ore type: Au- and U-bearing conglomerate from the
Witwatersrand Basin, South Africa. Quartz veins cutting this unit attest to the action of later hydrothermal fluids inthe sequence. (e) Hydrothermal ore type: quartz-carbonate vein network in an Archean orogenic or lode-gold deposit
from the Abitibi greenstone belt in Canada. The deposit is associated with igneous (lamprophyre) intrusions that
may be implicated in the mineralization process.

4

INTRODUCTION: MINERAL RESOURCES

surface and groundwaters. One drawback of thistype of classification, however, is that ore-forming
processes are complex and episodic. Ore formation
also involves processes that evolve, sometimes
over significant periods of geologic time. For
example, igneous processes become magmatichydrothermal as the intrusion cools and crystallizes, and sediments undergo diagenesis andmetamorphism as they are progressively buried,with accompanying fluid flow and alteration. In
addition, deformation of the Earth’s crust introduces new conduits that also facilitate fluid flow
and promote the potential for mineralization in
virtually any rock type. Ore-forming processes
can, therefore, span more than one of the three
categories, and there is considerable overlap between igneous and hydrothermal and between
sedimentary and hydrothermal, as illustrated
diagrammatically in Figure 1b, and also in theaccompanying photographs of the three major
categories of ore types.
The main part of this book is subdivided
into three sections termed Igneous (Part 1),
Hydrothermal (Part 2), and Sedimentary/surficial(Part 3). Part 1 comprises Chapters 1 and 2, which
deal with igneous and magmatic-hydrothermal
ore-forming processes respectively. Part 2 contains Chapter 3 and covers the large and diverse
range of hydrothermal processes not covered inPart 1. Part 3 comprises Chapter 4 on surficial andsupergene processes, as well as Chapter 5, which
covers sedimentary ore deposits, including a shortsection on the fossil fuels. The final chapter of thebook, Chapter 6, is effectively an addendum tothis threefold subdivision and is an attempt todescribe the distribution of ore deposits, both spatially in the context of global tectonics and temporally in terms of crustal evolution throughEarth history. This chapter is considered relevantin this day and age because the plate tectonic
paradigm, which has so pervasively influenced
geological thought since the early 1970s, provides
another conceptual basis within which to classify
ore deposits. In fact, modern economic geology,
and the scientific exploration of mineral deposits,
is now firmly conceptualized in terms of global
tectonics and crustal evolution. Although there is
still a great deal to be learnt, the links between

Table 1 Average crustal abundances for selected
metals and typical concentration factors that need to beachieved in order to produce a viable ore deposit

Al
Fe
Cu
Ni
Zn
Sn
Au
Pt

Averagecrustalabundance

Typicalexploitable
grade

Approximate
concentration
factor8.2%
5.6%
55 ppm75 ppm70 ppm
2 ppm
4 ppb
5 ppb

30%
50%
1%
1%
5%
0.5%
5 g t −1
5 g t −1

×4
×9
×180
×130
×700
×2500
×1250
×1000

Note: 1 ppm is the same as 1 g t −1.

plate tectonics and ore genesis are now sufficiently well established that studies of ore deposits
are starting to contribute to a better understanding of the Earth system.

WHAT MAKES A VIABLE MINERAL DEPOSIT?
Ore deposits form when a useful commodity is
sufficiently concentrated in an accessible part of
the Earth’s crust so that it can be profitably
extracted. The processes by which this concentration occurs are the topic of this book. As an introduction it is pertinent to consider the range ofconcentration factors that characterize the formation of different ore deposit types. Some of thestrategically important metals, such as Fe, Al,
Mg, Ti, and Mn, are abundantly distributed in the
Earth’s crust (i.e. between about 0.5 and 10%) andonly require a relatively small degree of enrichment in order to make a viable deposit. Table 1
shows that Fe and Al, for example, need to beconcentrated by factors of 9 and 4 respectively,
relative to average crustal abundances, in order toform potentially viable deposits.
By contrast, base metals such as Cu, Zn, and Ni
are much more sparsely distributed and averagecrustal abundances are only in the range 55–75
parts per million (ppm). The economics of mining
dictate that these metals need to be concentrated
by factors in the hundreds in order to form poten-

5

INTRODUCTION: MINERAL RESOURCES

109

Tons produced (1992)

108
Ca

107

5

Sb

10

Ag

104

102

Au

Bi

Hg

Cd

As

W

Fe
Mn

Ni

Sn

Mo

S

Al

Zn

Pb

106

103

Cr

Ti

Mg

Co

Se

PGE

1

10

0

10–6

10–5

10–4

10–3

10–2

10–1

100

101

Crustal abundance (wt%)
Figure 2 Plot of crustal abundances against global production for an number of metal commodities (after Einaudi,
2000). The line through Fe can be regarded as a datum against which the rates of production of the other metals canbe compared in the context of crustal abundances.

tially viable deposits, degrees of enrichment thatare an order of magnitude higher than those applicable to the more abundant metals. The degree ofconcentration required for the precious metals iseven more demanding, where the required enrichment factors are in the thousands. Table 1 showsthat average crustal abundances for Au and Pt arein the range 4–5 parts per billion (ppb) and eventhough ore deposits routinely extract these metals
at grades of around 5 g t−1, the enrichment factors
involved are between 1000 and 1250 times.
Another useful way to distinguish between thegeochemically abundant and scarce metals is toplot average crustal abundances against production estimates. This type of analysis was first carried out by Skinner (1976), who used a plot likethat in Figure 2 to confirm that crustal abundanceis a reasonable measure of the availability of agiven resource. It is by design and of necessitythat we use more of the geochemically abundant
metals than we do the scarce ones. The nature of
our technologies and the materials we use to manufacture mechanical items depend in large measure on the availability of raw materials. As an
example, the technologies (geological and metallurgical) that resulted in the dramatic increasein global aluminum production over the latterpart of the twentieth century allowed iron to be

replaced by aluminum in many products such as
motor vehicles. More importantly, though, Figure
2 allows estimates to be made of the relative rates
of depletion of certain metals relative to others.
These trends are discussed again below.
Mineral resources and ore reserves
In the course of this book reference is made to theterm “ore deposit” with little or no considerationof whether such occurrences might be economically viable. Although such considerations mightseem irrelevant in the present context it is necessary to emphasize that professional institutions
now insist on the correct definition and usage ofterminology pertaining to exploration results,
mineral resources, and ore reserves. Such terminology should be widely used and applied, as it
would help in reducing the irresponsible, and sometimes even fraudulent, usage of terminology, especially with respect to the investor public. Correct
terminology can also assist in the description andidentification of genuine ore deposits from zones
of marginal economic interest or simply anomalous concentrations of a given commodity.
Although the legislation that governs the public reporting of mineral occurrences obviously
varies from one country to the next, there is

6

INTRODUCTION: MINERAL RESOURCES

Exploration
results

Increasinglevel ofknowledge
andconfidence

Mineralresources

Mineralreserves(Reported as in situ
mineralization
estimates)(Reported as
mineable
production
estimates)

terminology is that what is economically extractable in a Third World artisinal operation may notof course be viable in a technically developed FirstWorld economy, and vice versa. The term “oredeposit” has no significance in the professionaldescription of a mineral occurrence and is bestused as a simply descriptive or generic term.

SOME USEFUL DEFINITIONS AND COMPILATIONS

InferredSome general definitions
Indicated

Probable

Measured

ProvedMining, metallurgical, economic,
marketing, legal, environmental,social, and governmental factors(Factors that contribute to the successful
exploitation of a deposit)

Figure 3 Simplified scheme illustrating the conceptual
difference between mineral resources and ore reserves
as applied to mineral occurrences. The scheme forms
the basis for the professional description of ore deposits
as defined by the Australian and South African
Institutes of Mining and Metallurgy.

now reasonable agreement as to the definition ofterms. In general it is agreed that different terms
should apply to mineral occurrences dependingon the level of knowledge and degree of confidence that is associated with the measurementof its quantity. Figure 3 is a matrix that reflects
the terminology associated with an increased
level of geological knowledge and confidence,and modifying factors such as those related tomining techniques, metallurgical extraction, marketing, and environmental reclamation. Exploration results can be translated into a mineral
resource once it is clear that an occurrence of
intrinsic economic interest exists in such formand quantity that there are reasonable prospects
for its eventual exploitation. Such a resource
can only be referred to as an ore reserve if it is apart of an economically extractable measured orindicated mineral resource. One problem with thisThis section is not intended to provide a comprehensive glossary of terms used in this book.
There are, however, several terms that are usedthroughout the text where a definition is either
useful or necessary in order to avoid ambiguity.
The following definitions are consistent withthose provided in the Glossary of Geology (Bates
and Jackson, 1987) and The Encyclopedia of the
Solid Earth Sciences (Kearey, 1993).
• Ore: any naturally occurring material fromwhich a mineral or aggregate of value can beextracted at a profit. In this book the concept
extends to coal (a combustible rock comprisingmore than 50% by weight carbonaceous material)and petroleum (naturally occurring hydrocarbon
in gaseous, liquid, or solid state).
• Syngenetic: refers to ore deposits that form at
the same time as their host rocks. In this bookthis includes deposits that form during the earlystages of sediment diagenesis.
• Epigenetic: refers to ore deposits that form after
their host rocks.
• Hypogene: refers to mineralization caused byascending hydrothermal solutions.
• Supergene: refers to mineralization caused bydescending solutions. Generally refers to theenrichment processes accompanying the weathering and oxidation of sulfide and oxide ores ator near the surface.
• Metallogeny: the study of the genesis of mineral deposits, with emphasis on their relationships in space and time to geological features of
the Earth’s crust.
• Metallotect: any geological, tectonic, lithological, or geochemical feature that has played a role

7

INTRODUCTION: MINERAL RESOURCES

1

MetalsGenerally decreasing electronegativity

H
3

4

Li

Be

11

12

Na

Mg

Metalloids

Ore mineral generally as:Oxide or
SulfidesulfideOxide

2

Non-metals

He

Nativemetal/alloy

Rb- Lithophile Cu- Chalcophile Au- Siderophile Ne- Atmophile

5

6

7

8

9

10

B

C

N

O

F

Ne

13

14

15

16

17

18

AI

Si

P

S

Cl

Ar

19

20

21

22

23

24

25

26

27

28

29

30

31

32

33

34

35

36

K

Ca

Sc

Ti

V

Cr

Mn

Fe

Co

Ni

Cu

Zn

Ga

Ge

As

Se

Br

Kr

37

38

39

40

41

42

43

44

45

46

47

48

49

50

51

52

53

54

Rb

Sr

Y

Zr

Nb

Mo

Tc

Ru

Rh

Pd

Ag

Cd

In

Sn

Sb

Te

I

Xe

55

56

57

72

73

74

75

76

77

78

79

80

81

82

83

84

85

86

Cs

Ba

La

Hf

Ta

W

Re

Os

Ir

Pt

Au

Hg

TI

Pb

Bi

Po

At

Rn

87

88

89

Fr

Ra

Ac

58

59

60

61

62

63

64

65

66

67

68

69

70

71

Ce

Pr

Nd

Pm

Sm

Eu

Gd

Tb

Dy

Ho

Er

Tm

Yb

Lu

90

91

92

Th

Pa

U

Generally decreasing electronegativity
No known use

Figure 4 Periodic table showing the 92 geologically relevant elements classified on the basis of their rock andmineral associations.

in the concentration of one or more elements inthe Earth’s crust.
• Metallogenic Epoch: a unit of geologic time
favorable for the deposition of ores or characterized by a particular assemblage of deposit types.
• Metallogenic Province: a region characterized
by a particular assemblage of mineral deposit types.
• Epithermal: hydrothermal ore deposits formedat shallow depths (less than 1500 meters) andfairly low temperatures (50–200 °C).
• Mesothermal: hydrothermal ore deposits formed
at intermediate depths (1500–4500 meters) andtemperatures (200–400 °C).
• Hypothermal: hydrothermal ore deposits formedat substantial depths (greater than 4500 meters)
and elevated temperatures (400–600 °C).
Periodic table of the elements
The question of the number of elements present
on Earth is a difficult one to answer. Most of theelement compilations relevant to the earth sciences show that there are 92 elements, the majority of which occur in readily detectable amounts
in the Earth’s crust. Figure 4 shows a periodic
table in which these elements are presented inascending atomic number and also categorized
into groupings that are relevant to metallogenesis.
There are in fact as many as 118 elements known
to man, but those with atomic numbers greaterthan 92 (U: uranium) either occur in vanishingly
small amounts as unstable isotopes that are theproducts of various natural radioactive decay
reactions or are synthetically created in nuclear
reactors. The heaviest known element, ununoctium (Uuo, atomic number 118), has been onlytransiently detected in a nuclear reactor and its
actual existence is still conjectural. Some of theheavy, unstable elements are, however, manufactured synthetically and serve a variety of uses.
Plutonium (Pu, atomic number 94), for example,
is manufactured in fast breeder reactors and is

8

INTRODUCTION: MINERAL RESOURCES

used as a nuclear fuel and in weapons manufacture.
Americium (Am, atomic number 95) is also manufactured in reactors and is widely used as theactive agent in smoke detectors.
Of the 92 elements shown in Figure 4, almost
all have some use in our modern technologicallydriven societies. Some of the elements (iron andaluminum) are required in copious quantities as
raw materials for the manufacture of vehicles andin construction, whereas others (the rare earths,
for example) are needed in very much smaller
amounts for use in the alloys and electronics
industries. Only three elements appear at thisstage to have little or no use at all (Figure 4). These
are astatine (At, atomic number 85), francium (Fr,atomic number 87), and protactinium (Pa, atomic
number 91). Francium is radioactive and so shortlived that only some 20–30 g exists in the entire
Earth’s crust at any one time! Astatine, likewise,
is very unstable and exists in vanishingly small
amounts in the crust, or is manufactured synthetically. Radon (Rn, atomic number 86) is an inert
or noble gas that is formed as a radioactive decay
product of radium. It has limited use in medical
applications, but, conversely, if allowed to accumulate can represent a serious health hazard incertain environments.
The useful elements can be broadly subdivided
in a number of different ways. Most of the elements can be classified as metals (Figure 4), with asmaller fraction being non-metals. The elements
B, Si, As, Se, Te, and At have intermediate properties and are referred to as metalloids. Anotherclassification of elements, attributed to the pioneering geochemist Goldschmidt, is based on
their rock associations and forms the basis for
distinguishing between lithophile (associated withsilicates and concentrated in the crust), chalcophile (associated with sulfides), siderophile (occur
as the native metal and concentrated in the core),
and atmophile (occur as gases in the atmosphere)
elements. It is also useful to consider elements interms of their ore mineral associations, with some
preferentially occurring as sulfides and others asoxides (see Figure 4). Some elements have properties that enable them to be classified in more than
one way and iron is a good example, in that it
occurs readily as both an oxide and sulfide.

Common ore and gangue minerals
It is estimated that there are about 3800 known
minerals that have been identified and classified(Battey and Pring, 1997). Only a very small proportion of these make up the bulk of the rocks of the
Earth’s crust, as the common rock forming minerals. Likewise, a relatively small number of minerals make up most of the economically viable ore
deposits of the world. The following compilation
is a breakdown of the more common ore minerals
in terms of chemical classes based essentially on
the anionic part of the mineral formula. Also
included are some of the more common “gangue,”
which are those minerals that form part of the ore
body, but do not contribute to the economicallyextractable part of the deposit. Most of these arealteration assemblages formed during hydrothermal processes. The compilation, including ideal
chemical formulae, is subdivided into six sections, namely native elements, halides, sulfides
and sulfo-salts, oxides and hydroxides, oxy-salts(such as carbonates, phosphates, tungstates, sulfates), and silicates. More detailed descriptions ofboth ore and gangue minerals can be found in avariety of mineralogical texts, such as Deer et al.
(1982), Berry et al. (1983), and Battey and Pring(1997). More information on ore mineral texturesand occurrences can be found in Craig andVaughan (1994) and Ixer (1990).
1 Native elements
Both metals and non-metals exist in nature in thenative form, where essentially only one element
exists in the structure. Copper, silver, gold, andplatinum are all characterized by cubic close
packing of atoms, have high densities, and aremalleable and soft. The carbon atoms in diamond
are linked in tetrahedral groups forming wellcleaved, very hard, translucent crystals. Sulfur occurs as rings of eight atoms and forms bipyramids
or is amorphous.
MetalsGold – Au
Silver – Ag
Platinum – Pt

INTRODUCTION: MINERAL RESOURCES

Palladium – Pd
Copper – Cu
Non-metals
Sulfur – S
Diamond – C
Graphite – C
2 Halides
The halide mineral group comprises compounds
made up by ionic bonding. Minerals such as halite
and sylvite are cubic, have simple chemical formulae, and are highly soluble in water. Halides
sometimes form as ore minerals, such as chlorargyrite and atacamite.
Halite – NaCl
Sylvite – KCl
Chlorargyrite – AgCl
Fluorite – CaF2
Atacamite – Cu2Cl(OH)3

9

Sperrylite – PtAs2
Braggite/cooperite – (Pt,Pd,Ni)S
Moncheite – (Pt,Pd)(Te,Bi)2
Cobaltite – CoAsS
Gersdorffite – NiAsS
Loellingite – FeAs2
Molybdenite – MoS2
Realgar – AsS
Orpiment – As2S3
Stibnite – Sb2S3
Bismuthinite – Bi2S3
Argentite – Ag2S
Calaverite – AuTe2
Pyrite – FeS2
Laurite – RuS2
Sulfo-salts
Tetrahedrite – (Cu,Ag)12Sb4S13
Tennantite – (Cu,Ag)12As4S13
Enargite – Cu3AsS44 Oxides and hydroxides

3 Sulfides and sulfo-salts
This is a large and complex group of minerals inwhich bonding is both ionic and covalent in character. The sulfide group has the general formula
AMXP, where X, the larger atom, is typically S but
can be As, Sb, Te, Bi, or Se, and A is one or moremetals. The sulfo-salts, which are much rarer
than sulfides, have the general formula AMBNXP,
where A is commonly Ag, Cu, or Pb, B is commonly As, Sb, or Bi, and X is S. The sulfide andsulfo-salt minerals are generally opaque, heavy
and have a metallic to sub-metallic lustre.
SulfidesChalcocite – Cu2S
Bornite – Cu5FeS4Galena – PbS
Sphalerite – ZnS
Chalcopyrite – CuFeS2
Pyrrhotite – Fe1–xS
Pentlandite – (Fe,Ni)9S8
Millerite – NiS
Covellite – CuS
Cinnabar – HgS
Skutterudite – (Co,Ni)As3This group of minerals is variable in its properties,
but is characterized by one or more metal in
combination with oxygen or a hydroxyl group.
The oxides and hydroxides typically exhibit ionicbonding. The oxide minerals can be hard, dense,
and refractory in nature (magnetite, cassiterite)
but can also be softer and less dense, forming as
products of hydrothermal alteration and weathering (hematite, anatase, pyrolucite). Hydroxides,
such as goethite and gibbsite, are typically the
products of extreme weathering and alteration.
OxidesCuprite – Cu2O
Hematite – Fe2O3
Ilmenite – FeTiO3Hercynite – FeAl2O4
Gahnite – ZnAl2O4
Magnetite – Fe3O4
Chromite – FeCr2O4
Rutile – TiO2
Anatase – TiO2
Pyrolucite – MnO2
Cassiterite – SnO2
Uraninite – UO2

10

INTRODUCTION: MINERAL RESOURCES

Thorianite – ThO2
Columbite-tantalite – (Fe,Mn)(Nb,Ta)2O6
Hydroxides (or oxyhydroxides)
Goethite – FeO(OH)
Gibbsite – Al(OH)3
Boehmite – AlO(OH)
Manganite – MnO(OH)
5 Oxy-salts
The carbonate group of minerals form whenanionic carbonate groups (CO32−) are linked byintermediate cations such as Ca, Mg, and Fe.
Hydroxyl bearing and hydrated carbonates canalso form, usually as a result of weathering andalteration. The other oxy-salts, such as the tungstates, sulfates, phosphates, and vanadates, areanalogous to the carbonates, but are built aroundan anionic group of the form XO n−
4 .
Carbonates
Calcite – CaCO3
Dolomite – CaMg(CO3)2
Ankerite – CaFe(CO3)2
Siderite – FeCO3
Rhodochrosite – MnCO3
Smithsonite – ZnCO3
Cerussite – PbCO3
Azurite – Cu3(OH)2(CO3)2
Malachite – Cu2(OH)2CO3
Tungstates
Scheelite – CaWO4
Wolframite – (Fe,Mn)WO4
Sulfates
Baryte(s) – BaSO4
Anhydrite – CaSO4
Alunite – KAl3(OH)6(SO4)2
Gypsum – CaSO4.2H2O
Epsomite – MgSO4.7H2O
Phosphates
Xenotime – YPO4
Monazite – (Ce,La,Th)PO4
Apatite – Ca5(PO4)3(F,Cl,OH)

Vanadates
Carnotite – K2(UO2)(VO4)2.3H2O
6 Silicates
The bulk of the Earth’s crust and mantle is madeup of silicate minerals that can be subdivided into
several mineral series based on the structure andcoordination of the tetrahedral SiO 4−4 anionic
group. Silicate minerals are generally hard, refractory and translucent. Most of them cannot be
regarded as ore minerals in that they do not represent the extractable part of an ore body, and thelist provided below shows only some of the silicates more commonly associated with mineraloccurrences as gangue or alteration products.
Some silicate minerals, such as zircon and spodumene, are ore minerals and represent importantsources of metals such as zirconium and lithium,respectively. Others, such as kaolinite, are mined
for their intrinsic properties (i.e. as a clay for theceramics industry).
Tekto (framework)
Quartz – SiO2
Orthoclase – (K,Na)AlSi3O8
Albite – (Na,Ca)AlSi3O8
Scapolite – (Na,Ca)4[(Al,Si)4O8)]3 (Cl, CO3)
Zeolite (analcime) – NaAlSi2O6.H2O
Neso (ortho)
Zircon – Zr(SiO4)
Garnet (almandine) – Fe3Al2(SiO4)3
Garnet (grossular) – Ca3Al2(SiO4)3
Sillimanite – Al2SiO5Topaz – Al2SiO4(F,OH)2
Chloritoid – (Fe,Mg,Mn)2(Al,Fe)Al3O2(SiO4)2(OH)4
Cyclo (ring)
Beryl – Be3Al2Si6O18
Tourmaline –(Na,Ca)(Mg,Fe,Mn,Al)3(Al,Mg,Fe)6Si6O18(BO3)3(OH,F)4
Soro (di)Lawsonite – CaAl2Si2O7(OH)2.H2O
Epidote – Ca2(Al,Fe)3Si3O12(OH)INTRODUCTION: MINERAL RESOURCES

Phyllo (sheet)
Kaolinite – Al4Si4O10(OH)8Montmorillonite – (Na,Ca)0.3(Al,Mg)2
Si4O10(OH)2.nH2O
Illite – KAl2(Si,Al)4O10(H2O)(OH)2
Pyrophyllite – Al2Si4O10(OH)2
Talc – Mg3Si4O10(OH)2
Muscovite – KAl2(AlSi3O10)(OH)2
Biotite – K(Fe,Mg)3(Al,Fe)Si3O10(OH,F)2
Lepidolite – K(Li,Al)3(Si,Al)4O10(OH,F)2
Chlorite – (Fe,Mg,Al)5–6(Si,Al)4O10(OH)8
Ino (chain)
Tremolite-actinolite – Ca2(Fe,Mg)5Si8O22(OH)2
Spodumene – LiAlSi2O6
Wollastonite – CaSiO3
Unknown structure
Chrysocolla – (Cu,Al)2H2Si2O5(OH)4.nH2O
Geological time scale
The development of a geological time scale has
been the subject of a considerable amount ofthought and research over the past few decades
and continues to occupy the minds and activities
of a large number of earth scientists around theworld. The definition of a framework within
which to describe the secular evolution of rocks,
and hence the Earth, has been, and continues
to be, a contentious exercise. The International
Commission on Stratigraphy (ICS is a working
group of the International Union of Geological
Sciences: IUGS) has been given the task of formalizing the geological time scale and this work is
ongoing (www.micropress.org/stratigraphy/). In
this book reference is often made to the timing ofvarious events and processes and the provision ofa time scale to which the reader can refer is, therefore, useful. Figure 5 is a time scale based on the2000 edition of the International Stratigraphic
Chart published and sanctioned by the ICS and
IUGS. In this diagram global chronostratigraphic
terms are presented in terms of eons, eras, periods,
and epochs, and defined by absolute ages in millions of years before present (Ma). Also shown are
the positions on the time scale of many of the ore

11deposits or metallogenic provinces referred to in
the text.

NATURAL RESOURCES, SUSTAINABILITY, AND
ENVIRONMENTAL RESPONSIBILITY

One of the major issues that characterized social and economic development toward the endof the twentieth century revolved around thewidespread acceptance that the Earth’s natural
resources are finite, and that their exploitation
should be carried out in a manner that will not
detrimentally affect future generations. The concept of “sustainable development” in terms of the
exploitation of mineral occurrences implies that
future social and economic practice should
endeavor not to deplete natural resources to thepoint where the needs of the future cannot be
met. This would seem to be an impossible goal
given the unprecedented population growth overthe past century and the fact that many commodities will become depleted within the next 100
years. The challenge for commodity supply overthe next century is a multifaceted one and will
require a better understanding of the earth system, improved incentives to promote more
efficient recycling of existing resources, and themeans to find alternative sources for commodities that are in danger of serious depletion.
There has been a dramatic rise in global population over the past 150 years. The number ofhumans on Earth has risen from 1 billion in 1830
to 6 billion at the end of the twentieth century.
Most predictions suggest that the populations ofmost countries will start to level off over the next
30 years and that global numbers will stabilize ataround 11 billion people by the end of the twentyfirst century. Societies in the next 100 years are,nevertheless, facing a scenario in which thedemand for, and utilization of, natural resources
continues to increase and certain commodities
might well become depleted in this interval.
Production trends for commodities such as oil,bauxite, copper, and gold (Figure 6) confirm thatdemand for resources mirrors population growthand is likely to continue to do so over the next few

QUAT. PERIODNEOGENE

EON

EPOCH

Early

Middle

Late

Early

Middle

Late

Early

Late

Paleocene

EoceneOligocene

MiocenePliocene

Pleistocene

Holocene

251

205

142

65.5

55.0

23.8
33.7

5.32

0.01
1.81

Main Arabian
Gulf reservoirs

MacTung/Troodos
Orapa
Andeanmagmatism(west east
progression)

Los PijiguaosSupergene/exotics
El Salvador
La Escondida/
Chuquicamata
CarlinSkaergaard

Hishikari
El LacoOrange Riverdiamonds
Kasuga

EPOCH
Early

Late
Middle

Early

MiddleLate

Llandovery

Wenlock

LudlowPridoli

Early

MiddleLate

Mississippian

Pennsylvanian

Cisuralian

Guadalupian

Lopingian

543

495

440

417

354

292

Ma
251

..
Rossing

Lachlan
S-and I-typegranites

Red Dog

Viburnum
Trend

Coal measures
Cornubianbatholith

EOARCHEAN

PALEOARCHEANMESOARCHEAN

NEOARCHEAN

PALEOPROTEROZOICMESOPROTEROZOICNEDPROTEROZOIC

3600

3200

2800

2500

1600

1000

Ma
543

4560 Ma(origin of solar system)

PRE CAMBRIAN(Not to scale)

BikitaWitwatersrandKambalda

Hamersley BIF
Great Dyke
Golden Mile

Sudbury Kiruna
Bushveld/
Phalaborwa

Olympic Dam

ArgyleZambianCopperbelt

Figure 5 Geological time scale after the ICS (2000). Also shown are the ages of the various deposits and metallogenic provinces mentioned in
the book.

PHANEROZOICEONPHANEROZOICMaPALEOZOICERACENOZOICMESOZOICPALEOGENECRETACEOUSJURASSICTRIASSIC

EAST
WEST

ERA
PERIODPERMIANCARBONIFEROUS

DEVONIAN
SILURIANCAMBRIAN ORDOVICIANEONPROTEROZOICARCHEANERAINTRODUCTION: MINERAL RESOURCES

(a)

(b)
30

Oil

25
20

Million tons

Billion barrels per year

13

15
10
USA

5

1900

1920

World

1940

1960

1980

120
110
100
90
80
70
60
50
40
30
20
10

2000

Bauxite

1940

Year
(c)

1950

1960 1970
Year

1980

1990

(d)
600

Copper

2000

Gold

1500

400
World reserve baseTons

Million tons

500

300

1000

200
World cumulative
production

100

1950

1960

1970
Year

1980

1990

500

1860 1880 1900 1920 1940 1960 1980 2000
Year

Figure 6 Global production trends for oil (a), bauxite (b), copper (c), and gold (d) over the twentieth century(after compilations in Craig et al., 1996).

decades. World oil production increased precipitously until the late 1970s, but since then avariety of political and economic factors havecontributed to tempering production (Figure 6a),thereby ensuring a longer-term reserve base. A
similar levelling of production is evident for bauxite (Figure 6b) but such a trend is not yet evident
for the precious metals such as gold or platinum.
For some commodities, such as copper (Figure 6c),the world reserve base is also levelling off, a feature that in part also reflects fewer new and large
discoveries. Critical shortages of most naturalcommodities are not likely to present a problem
during the early part of the twenty-first century(Einaudi, 2000), but this situation will deteriorate
unless strategies for sustainability are put into
place immediately.
The depletion of commodities in the Earth’s
crust is particularly serious for those metals thatare already scarce in terms of crustal abundancesand for which high degrees of enrichment are
required in order to make viable ore deposits.
Figure 2 illustrates the point by referring to the
production of iron as a baseline measure against

14

INTRODUCTION: MINERAL RESOURCES

which extraction of other metals can be compared(Skinner, 1976). Those elements which fall above
the Fe production line (notably Au, Ag, Bi, Sb, Sn,
Cu, Pb, and Zn) are being extracted or depleted at
faster rates, relative to their crustal abundances,than Fe. It is these metals that are in most dangerof depletion in the next 50 years or so unlessproduction is ameliorated or the reserve base isreplaced. Conversely, those metals that plot
beneath the Fe production line (such as Ti, Mg,
and Al) are being extracted at slower rates than Fe
and are in less danger of serious depletion during
this century.
One of the ways in which metallogeny can
assist in the creation of a sustainable pattern of
resource utilization is to better understand theprocesses by which ores are concentrated in the
Earth’s crust. The replacement of the global commodity reserve base is obviously dependent on
exploration success and the ability to find newore deposits that can replace those that are being
depleted. It is, of course, increasingly difficult tofind new and large deposits of conventional ores,
since most of the accessible parts of the globe havebeen extensively surveyed and assessed for theirmineral potential. The search for deeper deposits
is an option but this is dependent to a large extent
on the availability of technologies that will
enable mining to take place safely and profitably
at depths in excess of 4000 meters (currently the
deepest level of mining in South African gold
mines). Another option is to extract material frominaccessible parts of the globe, such as the ocean
floor, a proposal that has received serious consideration with respect to metals such as Mn and Cu.
Again, there are technological barriers to suchprocesses at present, but these can be overcome,as demonstrated by the now widespread exploration for, and extraction of, oil and gas from thesea floor. A third option to improve the sustainability of resource exploitation is to extract usefulcommodities from rocks that traditionally havenot been thought of as viable ores. Such a development can only be achieved if the so-called “mineralogical barrier” (Skinner, 1976) is overcome. This
concept can be described in terms of the amountof energy (or cost) required to extract a commodity from its ore. It is, for example, considerably

cheaper to extract Fe from a banded ironformation than it is from olivine or orthopyroxene in an igneous rock, even though both rocktypes might contain significant amounts of themetal. The economics of mining and the widespread availability of banded iron-formation dictate that extraction of Fe from silicate minerals isessentially not feasible. The same is not true ofnickel. Although it is cheaper and easier to extract
Ni from sulfide ore minerals (such as pentlandite)there is now widespread extraction of the metalfrom nickeliferous silicate minerals (garnierite)that form during the lateritic weathering of
ultramafic rocks. Even though Ni is more difficultand expensive to extract from laterite than from
sulfide ores, the high tonnages and grades, as wellas the widespread development and ease of accessof the former, mean that they represent viable
mining propositions despite the extractive difficulties. Ultimately, it may also become desirable toconsider mining iron laterites, but this would only
happen if conventional banded iron-formation
hosted deposits were depleted, or if the economicsof the whole operation favored laterites over ironformations. This is not likely to happen in the
short term, but, if planned for, the scenario does
offer hope for sustainability in the long term. In
short, sustainable production of mineral resources
requires a thorough understanding of ore-forming
processes and the means to apply these to thediscovery of new mineral occurrences. It also
requires the timely development of technologies,
both in the earth sciences and in related fieldsof mining and extractive metallurgy, that will
enable alternative supplies of mineral resources
to be economically exploited in the future.
Mining and environmental responsibility
A global population of 11 billion by the end of the
century presents a major problem in terms of the
supply of most of the world’s mineral resources.
What is even more serious, though, is the enormous strain it will place on the Earth’s fragile environment arising from the justifiable expectationthat future technologically advanced societies
will provide an adequate standard of living, interms of food, water, housing, recreation, andINTRODUCTION: MINERAL RESOURCES

15

material benefits, to all their peoples. In addition
to commodity supply problems, the twenty-first
century will be also be characterized by unprecedented depletion of even more critical resourcesin the form of soil, water, and clean air (Fyfe,
2000). Legislation that is aimed at dealing withissues such as atmospheric pollution and greenhouse gas emissions, factory waste and aciddrainage, the burning and destruction of forests,
the protection of endangered species, overgrazing,and erosion is highly desirable but far from globally applicable because it is perceived as a luxury
that only the developed world can afford.
The study of ore-forming processes is occasionally viewed as an undesirable topic that
ultimately contributes to the exploitation of theworld’s precious natural resources. Nothing could
be further from the truth. An understanding of theprocesses by which metals are concentrated in the
Earth’s crust is essential knowledge for anyone

concerned with the preservation and remediationof the environment. The principles that underpin
the natural concentration of ores in the crust are
the same as those that can be utilized in issues
such as the control of acid mine drainage, andsoil and erosion management. Mining operations
around the world are increasingly having toassume responsibility for reclamation of the landscape once the resource has been depleted. The
industry now encompasses a range of activities
extending from geological exploration and evaluation, through mining and beneficiation, andeventually to environmental reclamation. Thisis the mining cycle and its effective managementin the future will be a multidisciplinary exercise
carried out by highly skilled scientists and engineers. Earth systems science, and in particular thegeological processes that gave rise to the naturalconcentration of the ore in the first place, will becentral to this entire operation.

The discipline of “economic geology” and in particular the field of metallogeny (the study of thegenesis of ore deposits) remains critical to the teaching of earth systems science. A holistic approach
involving the integration of knowledge relevantto the atmosphere, biosphere, and lithosphere is
now regarded as essential to understanding thecomplexities of the earth system. The development of environmentally responsible and sustainable policies with respect to the future supplyof all natural resources will demand a thoroughknowledge of the nature and workings of the earthsystem. Central to this is an understanding of
metallogeny and the nature and origin of the entire
spectrum of mineral resources, including thefossil fuels. The classification of mineral deposits
in terms of process can be simply and effectively
achieved in terms of rock associations, namely
igneous, hydrothermal, and sedimentary. This
breakdown forms the basis for the layout of thisbook.

Blunden, J. (1983) Mineral Resources and Their Management. Harlow: Longman, 302 pp.
Craig, J.R., Vaughan, D.J. and Skinner, B.J. (1996) Resources of the Earth – Origin, Use and Environmental
Impact. Englewood Cliffs, NJ: Prentice Hall, 472 pp.

Ernst, W.G. (2000) Earth Systems – Processes and Issues.
Cambridge: Cambridge University Press, 559 pp.
Kesler, S.E. (1994) Mineral Resources, Economics andthe Environment. London: Macmillan, 400 pp.

Igneous Processes

Igneous ore-forming
processes

METALLOGENY OF OCEANIC AND CONTINENTAL CRUSTFUNDAMENTAL MAGMA TYPES AND THEIR METAL ENDOWMENTTHE RELATIVE FERTILITY OF MAGMAS AND THE “INHERITANCE
FACTOR”“late-veneer” hypothesis
diamonds and kimberlite/lamproite
metal concentrations in metasomatized mantle
S- and I-type granites
PARTIAL MELTING AND CRYSTAL FRACTIONATION AS OREFORMING PROCESSESTRACE ELEMENT DISTRIBUTION DURING PARTIAL MELTING
TRACE ELEMENT DISTRIBUTION DURING FRACTIONAL
CRYSTALLIZATIONMONOMINERALIC CHROMITITE LAYERS
LIQUID IMMISCIBILITY AS AN ORE-FORMING PROCESSSPECIAL EMPHASIS ON MINERALIZATION PROCESSES IN LAYEREDMAFIC INTRUSIONS

sulfide solubility
sulfide–silicate partition coefficients
the R factor
PGE clusters and hiatus models

1.1 INTRODUCTION
Igneous rocks host a large number of different
ore deposit types. Both mafic and felsic rocks
are linked to mineral deposits, examples of whichrange from the chromite ores resulting fromcrystal fractionation of mafic magmas to tin
deposits associated with certain types of granites.
The processes described in this chapter relate
to properties that are intrinsic to the magmaBox 1.1 Diamondiferous kimberlites and lamproites:
the Orapa (Botswana) and the Argyle (Western Australia)
diamond mines
Box 1.2 Partial melting and concentration ofincompatible elements: the Rössing uranium depositBox 1.3 Boundary layer differentiation in granites andincompatible element concentration: the Zaaiplaats tin
deposit, Bushveld Complex
Box 1.4 Crystal fractionation and formation ofmonomineralic chromitite layers: the UG1 chromitite
seam, Bushveld Complex
Box 1.5 Silicate–sulfide immiscibility: the komatiite
hosted Ni–Cu deposits at Kambalda, Western Australia
Box 1.6 New magma injection and magma mixing:
the Merensky Reef, Bushveld Complex
Box 1.7 Magma contamination and sulfide
immiscibility: the Sudbury Ni–Cu deposits

itself and can be linked genetically to its coolingand solidification. Discussion of related processes, whereby an aqueous fluid phase forms or“exsolves” from the magma as it crystallizes,
is placed in Chapter 2. The topics discussed
under the banners of igneous and magmatic–hydrothermal ore-forming processes are intimately
linked and form Part 1 of this book.
A measure of the economic importance of ore
deposits hosted in igneous rocks can be obtained
from a compilation of mineral production data asa function of host rock type. A country like South
Africa, for example, is underlain dominantly by
sedimentary rocks and these undoubtedly hostmany of the valuable mineral resources (especially20PART 1 IGNEOUS PROCESSES

Table 1.1 A comparison of the value of mineral production from igneous and sedimentary rocks in South Africa
Mineralization hosted in

Area (km2)

Value of sales, 1971
(106 US$)% of total area

% of total value

Unit value (US$/km2)

Granites
Mafic layered complexes
Total (igneous)
Sedimentary rocks

163 100
36 400199 500
1 023 900

1973
7288
9261
49 137

13.3
3.0
16.3
83.7

3.4
12.5
15.9
84.1

12 000200 200
46 40047 900Source: after Pretorius (1976).

if the fossil fuels are taken into consideration).
Nevertheless, the value of ores hosted in igneousrocks per unit area of outcrop can be comparablewith that for sedimentary rocks, as indicatedin Table 1.1. Although South Africa is characterized by a rather special endowment of mineral
wealth related to the huge Bushveld Complex,
the importance of igneous-hosted ore deposits isnevertheless apparent.

1.2 MAGMAS AND METALLOGENY
It is well known that different igneous rocks
host ore deposits with different metal associations, and that this must be related somehow tothe environments in which magmas are generated
and the resulting compositional characteristics
they inherit from their various settings. It is
widely recognized, for example, that many of thechalcophile and siderophile elements (such as Ni,
Co, Pt, Pd, and Au) are more likely to be associated with mafic rock types, whereas concentrations of many lithophile elements (such as Li, Sn,
Zr, U, and W) are typically found in association
with felsic or alkaline rock types. This has implications for understanding ore genesis and, consequently, some of the factors related to thesedifferences are discussed below.
1.2.1 Crustal architecture and mineral wealth
Although the greatest concentrations of siderophile and chalcophile elements almost certainly
reside in the mantle and core of the Earth, theseare generally inaccessible due to their very great

depths. In fact, most of the world’s economically
exploitable mineral wealth effectively lies on thesurface or just below the surface of the Earth. The
world’s deepest mine, the Western Deep Levels
gold mine near Johannesburg, South Africa, extends
to just over 4000 m depth and this places an effective limit on ore body exploitation, at least interms of present technologies. Nevertheless, many
mineral commodities are formed much deeper
in the crust than 4 km, with some even being
derived from the mantle. Diamonds, for example,
are hosted in kimberlite magmas that have beenbrought to exploitable depths by a variety of
igneous or tectonic mechanisms. Understanding
ore genesis processes, therefore, requires a knowledge of lithospheric (i.e. crust and upper mantle)
architecture, and also of the origin and nature of
the igneous rocks in this section of the Earth.
The oceanic crust, which covers some twothirds of the Earth surface, is thin (less than 10 km)
and, compared to the continents, has a composition and structure that is relatively simple andconsistent over its entire extent. The upper layer,on average only 0.4 km thick (Kearey and Vine,
1996), comprises a combination of terrigenous andpelagic sediments that are distributed mainly byturbidity currents. They are often highly reduced
and metal charged. This is underlain by a layer,
typically 1–2.5 km thick, that is both extrusive
and intrusive in character and dominantly basaltic
in composition. The basalts are, in turn, underlain
by the main body of oceanic crust that is plutonic
in character and formed by crystallization andfractionation of basaltic magma. This cumulateassemblage comprises mainly gabbro, pyroxenite,and peridotite. Sections of tectonized and meta-

IGNEOUS ORE-FORMING PROCESSES CHAPTER 1

Midocean
VMS
ridge Cu, Co, Zn

Ocean

Mn, Co, Ni
Pelagic sediment

Cr, Ni, Pt

Sheeted dykes
CrPartialmelting

Plagioclase
Pyroxene
Olivine21

Lavas ed) Oceanic lithosphere(pillow

Cumulates

Tectonized dunite and harzburgite

Plagioclaselherzolite
Spinellherzolite

Figure 1.1 Oceanic crustal architecture showing the main types of ore deposits characteristic of this environment.
Only chromite and related deposits (Cr–Ni–Pt) are related to igneous ore-forming processes; VMS (Cu, Co, Zn) andsediment-hosted deposits (Mn, Co, Ni) are discussed in Chapters 3 and 5 respectively.

morphosed oceanic lithosphere can be observed in
ophiolite complexes which represent segments of
the ocean crust (usually back-arc basins) that havebeen thrust or obducted onto continental margins
during continent–ocean collision.
The types of ore deposits that one might expect to find associated with ophiolitic rocks areshown in Figure 1.1. They include the categoryof podiform chromite deposits that are related
to crystal fractionation of mid-ocean ridge basalt(MORB), and also have potential for Ni and Ptgroup element (PGE) mineralization. Accumulations of manganese in nodules on the sea floor,
metal-rich concentrations in pelagic muds, andexhalative volcanogenic massive sulfide (VMS)
Cu–Zn deposits also occur in this tectonic setting, but are not directly related to igneous processes and are discussed elsewhere (Chapters 3
and 5).
The continental crust differs markedly fromits oceanic counterpart. It is typically 35–40 km
thick, but thins to around 20 km under rift zones
and thickens to 80 km or more beneath young
mountain belts. Historically, the continental crust
was thought to comprise an upper zone made up
largely of granite (and its sedimentary derivatives)and a lower, more mafic zone, with the two layers
separated by the Conrad discontinuity (whichmarks a change in seismic velocities, and, therefore crustal density). More recent geophysicaland geological studies clearly indicate that crustal
architecture is more complex and reflects along-lived tectonic and magmatic history, extending back in some cases over 3800 million years(Figure 1.2).
The continents have been progressively constructed throughout geological time by a variety
of magmatic, sedimentary, and orogenic processes
taking place along active plate margins and, to alesser extent, within the continents themselves. In
addition, continental land masses have repeatedly
broken apart and reamalgamated throughout geological history. These episodes, known as Wilson
cycles, have rearranged the configuration of continental fragments several times in the geologicalpast. In the early Proterozoic, for example, it isconceivable that segments of southern Africa andwestern Australia might have been part of thesame continent. The significance of these cycles,
and the pattern of crustal evolution with time, to
global metallogeny is discussed in more detailin Chapter 6.

22PART 1 IGNEOUS PROCESSESCu, Mo, Pb, Zn

Oceanic crus

Sn, W Diamond
Cr, Cu, Ni,
Cu, REE, P PGE, V
Volcanic
arc
Rift

t

Tecto
n

LithosphereI-typeAsthenosphere

Sn, W, Cu, Au U, Th

S-typeContinental
crust

icallyKimberlite

thicke

ned

Co

ntin

ent

al c

rus

t

Figure 1.2 Continental crustal architecture showing the main types of igneous-related ore deposits characteristicof this environment.

The upper crust, which in some continental
sections is defined as extending to the Conrad
discontinuity at some 6 km depth, is made up
of felsic to intermediate compositions (granite todiorite) together with the sedimentary detritus
derived from the weathering and erosion of thismaterial. Archean continental fragments (greaterthan 2500 Myr old) also contain a significant component of greenstone belt material, representing
preserved fragments of ancient oceanic crust. The
lower crust, between the Conrad and Mohorovicic
discontinuities, is variable in composition but
is typically made up of hotter, and usually more
dense, material. This is because temperaturesand pressures in the crust increase with depth ataverage rates of some 25 °C km−1 and 30 MPa km−1
respectively (Kearey and Vine, 1996). The lower
crust is not necessarily compositionally different from the upper crust, but exists at higher
metamorphic grades. It is also likely to be moreanhydrous and residual, in the sense that magma
now present at higher levels was extracted from thelower crust, leaving a residue of modified material.
Some of the lower crust may be more mafic in composition, comprising material such as amphibolite,
gabbro, and anorthosite.
Most of the world’s known ore deposits are,of course, hosted in rocks of the continental crust,and the full range is not shown in Figure 1.2.
Some of the more important igneous rock-related

deposit types are shown and these include diamondiferous kimberlites, anorthosite-hosted Ti
deposits, the Cr–V–Pt–Cu–Ni assemblage of ores
in continental layered mafic suites, and the
Sn–W–F–Nb–REE–P–U family of lithophile ores
related to granites and alkaline intrusions.
1.2.2 Magma types and metal contents
Although their rheological properties are different,
the outer two layers of the Earth, the more rigid
lithosphere and the ductile asthenosphere, arelargely solid. Zones within these layers that are
anomalous in terms of pressure or temperaturedo, however, form and can cause localized melting of the rocks present. The nature of the rockundergoing melting and the extent to which it ismelted are the main factors that control the composition of the magma that is formed. The magmacomposition, in turn, dictates the nature of metal
concentrations that are likely to form in the rocks
that solidify from that magma.
Although it is theoretically possible to form
an almost infinite range of magma compositions(from ultramafic to highly alkaline), for ease of
discussion this section is subdivided into fourparts, each representing what is considered to bea fundamental magma type – these are basalt,
andesite, rhyolite, and alkaline magmas, the latterincluding kimberlite.

IGNEOUS ORE-FORMING PROCESSES CHAPTER 1

BasaltBasalts form in almost every tectonic environment,
but the majority of basaltic magma production
takes place along the mid-ocean ridges, and in
response to hot-spot related plumes, to form oceanic
crust. In addition, basalts are formed together
with a variety of more felsic magmas, along island
arcs and orogenic continental margins. Basaltic
magma may also intrude or extrude continental
crust, either along well defined fractures or rifts(such as continental flood basalt provinces, or theGreat Dyke of Zimbabwe) or in response to intraplate hot-spot activity (which might have been
responsible for the formation of the Bushveld
Complex of South Africa).
Basalt forms by partial melting of mantle
material, much of which can generally be describedas peridotitic in composition. Certain mantle rocks,
such as lherzolite (a peridotite which contains
clinopyroxene and either garnet or spinel), havebeen shown experimentally to produce basaltic
liquids on melting, whereas others, like alpinetype peridotite (comprising mainly olivine andorthopyroxene), are too refractory to yield basaltic
liquids and may indeed represent the residues
left behind after basaltic magma has already been
extracted from the mantle. Likewise, oceanic
crust made up of hydrated (serpentinized) basalt
and drawn down into a subduction zone is alsoa potential source rock for island arc and continental margin type magmatism. Komatiites, which
are ultramafic basalt magmas (with >18% MgO)
mainly restricted to Archean greenstone belts, havea controversial origin but are generally believed
to represent high degrees of partial melting ofmantle during the high heat-flow conditions thatprevailed in the early stages of crust formation
prior to 2500 Ma.
Ore deposits associated with mafic igneous rocks
typically comprise a distinctive (mainly siderophileand chalcophile) metal assemblage of, among
others, Ni, Co, Cr, V, Cu, Pt, and Au. Examinationof Table 1.2 shows that this list corresponds tothose elements that are intrinsically enriched inbasaltic magmas. Figure 1.3 illustrates the relative
abundances of these metals in three fundamental
magma types and the significantly higher con-23

centrations in basalt by comparison with andesite
and rhyolite. The enhanced concentration of these
metals in each case is related to the fact that thesource materials from which the basalt formed
must likewise have been enriched in those constituents. In addition, enhanced abundances also
reflect the chemical affinity that these metals havefor the major elements that characterize a basaltic
magma (Mg and Fe) and dictate its mineral composition (olivine and the pyroxenes).
The chemical affinity that one element has foranother is related to their atomic properties asreflected by their relative positions in the periodic
table (see Figure 4, Introduction). The alkali earth
elements (i.e. K, Na, Rb, Cs, etc.), for example,
are all very similar to one another, but haveproperties that are quite different to the transition
metals (such as Fe, Co, Ni, Pt, Pd). In addition,
minor or trace elements, which occur in such low
abundances in magmas that they cannot forma discrete mineral phase, are present by virtueof their ability either to substitute for anotherchemically similar element in a mineral lattice or
to occupy a defect site in a crystal lattice. Thisbehavior is referred to as diadochy or substitution and explains much, but not all, of the traceelement behavior in rocks. Substitution of a traceelement for a major element in a crystal takes
place if their ionic radii and charges are similar.
Typically radii should be within 15% of one
another and charges should differ by no morethan one unit provided the charge difference canbe compensated by another substitution. Bondstrength and type also effects diadochy and it
preferentially occurs in crystals where ionic bonding dominates.
A good example of diadochic behavior is thesubstitution of Ni2+ for Mg2+ in olivine, or V3+
for Fe3+ in magnetite. Analytical data for the Ni
content of basalts shows an excellent correlation between Ni and MgO contents (Figure 1.4),
confirming the notion that the minor metal substitutes readily for Mg. The higher intrinsic Nicontent of ultramafic basalts and komatiites would
suggest that the latter rocks are perhaps better
suited to hosting viable magmatic nickel deposits,
an observation borne out by the presence ofworld class nickel deposits hosted in the Archean24PART 1 IGNEOUS PROCESSES

Table 1.2 Average abundances of selected elements in the major magma types

Li
Be
F
P
V
Cr
Co
Ni
Cu
Zn
Zr
Mo
Sn
Nb
Sb
Ta
W
Pb
Bi
U
Th
Ag†
Au†
Pt†
S
Ge
As
Cd

BasaltAndesite

Rhyolite

Alkaline magma

Kimberlite

Clarke*

10
0.7
380
3200
266
307
48
134
65
94
87
0.9–2.7
0.9
5
0.1–1.4
0.9
1.2
6.4
0.02
0.1–0.6
0.2
100
3.6
17–30
782
1.1
0.8
0.02

12
1.5
210
2800
148
55
24
18
60
87
205
0.8–1.2
1.5
4–11
0.2
–
1.1
5
0.12
0.8
1.9
80
–
–
423
1.2
1.8
0.02

50
4.1
480
1200
72
4
4.4
6
6
38
136
1
3.6
28
0.1–0.6
2.3
2.4
21
0.12
5
26
37
1.5
3–12
284
1.0–1.3
3.5
0.2–0.5

–
4–24
640
1800
235
–
–
–
–
108
1800
15
–
140
–
10
16
15
–
10
35
–
–
–
598
1.3–2.1
–
0.04

–
–
–
0.6–0.9%
–
–
–
1050
103
–
2200
–
–
240
–
–
–
–
–
–
–
–
–
19
2100
0.5
–
–

20
2.8
625
1050
135
10
25
75
55
70
165
1.5
2
20
0.2
2
1.5
13
0.17
2.7
7.2
70
4
10
260
1.5
1.8
0.2

If no average is available, a range of values is provided. *Clarke is a term that refers to the average crustal abundance.
†
Values as ppb, all other values as ppm.
Source: data from Taylor (1964), Wedepohl (1969), Krauskopf and Bird (1995).

400
BasaltAndesite

Rhyolite

Abundance (ppm)

300

200

100

0

Ni

Co

Cr

V

Cu

Zn

Pt × 10 Au × 100
ppb
ppb

Figure 1.3 Relative abundances of
selected metals in basalt, andesite,
and rhyolite (data from Table 1.2).

IGNEOUS ORE-FORMING PROCESSES CHAPTER 1

2000

Ni (ppm)

1500
1000
500
0
35

30

25

20
15
Wt% MgO

10

50

Figure 1.4 The relationship between Ni and MgOcontents of basalts within which base metalmineralization does not occur (data from Naldrett,1989a).

komatiites of the Kambalda mining district inWestern Australia (see Box 1.5) and elsewhere in
the world.
Andesite
Andesites are rocks that crystallize from magmas
of composition intermediate between basalt andrhyolite (typically with SiO2 contents between
53 and 63 wt%). Their petrogenesis remains
contentious, although it is well known that theytend to occur dominantly in orogenic zones,
either along island arcs or on continental margins
beneath which subduction of oceanic crust is
taking place (Hall, 1996). Discussion about theorigin of andesite revolves around whether it
represents a primary magma composition derived
directly by an appropriate degree of melting of asuitable source rock, or an evolved melt formedby differentiation of a more mafic magma suchas basalt.
Geological observations support the notionthat andesite can be formed both as a primarymagma composition and by in situ fractionation.
The observation that andesitic volcanoes occur
directly above aseismic sections of a Benioffzone (i.e. the subducted slab that produces earthquakes due to movement and fracturing of rock)suggests that melt production (and damping ofseismic waves) has occurred in these areas. This
would support the notion that andesitic magma

25

is produced by direct melting of hydrous oceanic
crust or, more likely, the mantle wedge overlying the subduction zone as it is permeated by
fluids expelled from the subjacent oceanic crust.
Alternatively, andesitic magma can be producedby fractionation of phases such as hornblendeand magnetite from relatively water-rich parent
magmas (Osborn, 1979), or by contamination ofan originally more mafic melt by felsic material
or melt.
Irrespective of the mode of formation of andesite
it is apparent that as a magma type it does notexhibit a primary association with any particular suite of metals or ore deposits. It appears
instead that ore deposits tend to be associatedwith magmas representing the ends of the compositional spectrum, and that intermediate melt
compositions are simply characterized by intermediate trace element abundances. Examinationof Table 1.2 shows that andesites appear to havelittle or no metal specificity and are characterized
by trace element abundances that are intermediate between those of basalt on the one hand andeither granite or alkaline rocks on the other.
Rhyolite
Felsic magmas can also form in a variety of geological environments. They crystallize at depth
to form a spectrum of rock compositions ranging
from Na-rich tonalite to K-rich alkali granite, or
extrude on surface to form dacitic to rhyolitic
volcanic rocks. Very little granite magma forms inoceanic crust or along island arcs that have formed
between two oceanic plates. Where oceanic granite
does occur it is typically the result of differentiation of a more mafic magma type originally formedby mantle melting. Along the mid-Atlantic ridgein Iceland, for example, eruptions of the volcano
Hekla are initiated by a pulse of felsic ash production which is rapidly followed by eruption of
more typical basaltic andesite. This suggests thatthe intervening period between eruptions wascharacterized by differentiation of the magma andthat the accompanying build-up of volatiles may
have been responsible for the subsequent eruption(Baldridge et al., 1973). These observations, among
many others, clearly indicate that granitic melts

26PART 1 IGNEOUS PROCESSES

60
BasaltAndesite

Rhyolite

Abundance (ppm)

50
40
30
20
10
0

Li

Be × 10

F/10

Sn × 10

W × 10can be the products of differentiation of more mafic
magmas in oceanic settings.
Most felsic magmas, however, are derived fromthe partial melting of predominantly crustal material along ocean–continent island arcs and orogenic
continental margins. Although Andean-type subduction zones might facilitate partial melting of
the downgoing slab itself, the much higher proportion of felsic magma formed in this environment compared to oceanic settings points to asignificant role for continental crust as a source.
There is now general agreement that Andeantype subduction-related magmatism receives melt
contributions from both the mantle lithosphereand the continental crust, with the wide-ranging
compositions of so-called “calc–alkaline” igneous
suites being attributed to a combination of bothmagma mingling and fractional crystallization(Best, 2003).
Significant quantities of felsic magma are produced in the latter stages of continent–continentcollision and also in anorogenic continental settings where rifting and crustal thinning has takenplace. Himalayan-type continent collision, forexample, is usually accompanied by crustal thickening associated with intense thrusting, tectonic
duplication and reverse metamorphic gradients.
These processes cause dewatering of crustal
material, which, in turn, promotes partial melting
to form high-level leucogranite magmas derived
from source rocks that often contain significant
proportions of sedimentary material (Le Fort,

U × 10Th

Figure 1.5 Relative abundances of
selected “granitophile” elements
in basalt, andesite, and rhyolite(data from Table 1.2).

1975). Anorogenic continental magmatism, on theother hand, is usually related to crustal thinning(accompanying plume or hot-spot activity?) andis typified by the production of magmas withbimodal compositions (i.e. basalt plus rhyolite).
A good example is the 2060 Myr old Bushveld
Complex in South Africa, where early mafic magmas intruded to form the world’s largest layered
igneous complex, followed by emplacement of a
voluminous suite of granites.
Ore deposits associated with felsic igneous rocks
often comprise concentrations of the lithophileelements such as Li, Be, F, Sn, W, U, and Th.
Table 1.2 shows that this list corresponds to thoseelements that are intrinsically enriched in rhyolitic magmas and Figure 1.5 illustrates, in bar
graph form, the relative abundances of these elements and, in particular, the higher abundances inrhyolite by comparison with andesite and basalt.
The relative enrichment of certain lithophileelements in rhyolitic magmas is partially related
to their geochemically incompatible nature. An
incompatible element is one whose ionic charge
and radius make it difficult to substitute for anyof the stoichiometric elements in rock-forming
minerals. Thus, incompatible elements tend to
be excluded from the products of crystallizationand concentrated into residual or differentiatedmagmas (such as the granitic magmas that mightform by crystal fractionation of mafic magmas
in oceanic settings). Alternatively, incompatible
elements also tend to be concentrated in crustal

IGNEOUS ORE-FORMING PROCESSES CHAPTER 1

melts derived from low degrees of partial meltingof source rocks that may themselves have beenendowed in the lithophile elements. These concepts are discussed in more detail in section 1.4
below.
A well known and interesting feature of ore
deposits that are genetically associated with
granite intrusions is that the origin and composition of the magma generally controls the natureof the metal assemblage in the deposit (Chappelland White, 1974; Ishihara, 1978, 1981). This control is almost certainly related in part to the metalendowment inherited by the magma from the rocks
that were melted to produce it. Where a felsicmagma is derived from melting of a sedimentaryor supracrustal protolith (termed S-type granites),
associated ore deposits are characterized by concentrations of metals such as Sn, W, U, and Th.
Where it is derived from melting of older igneous
protoliths in the crust (I-type granite) the ore association is typified by metals such as Cu, Mo, Pb,
Zn, and Au. This association is metallogenicallyvery significant and is discussed in more detail in
section 1.3.4 below and again in Chapter 2.
Alkaline magmas and kimberlite
Although most magma compositions can be represented by the basalt–andesite–rhyolite spectrum,some deviate from this trend and are compositionally unusual. For example, magmas that are
depleted in SiO2 but highly enriched in the alkali
elements (Na, K, and Ca) are relatively rare, but
may be economically important as they frequently
contain impressive concentrations of a wide rangeof ore-forming metals (such as Cu, Fe, P, Zr, Nb,
REE, F, U, and Th). In addition, kimberlitic andrelated magma types (such as lamproites) are themain primary source of diamonds.
The most common alkaline mafic magma isnephelinite, which crystallizes to give a range ofrock types (the ijolite suite; Hall, 1996) comprising rather unusual minerals, such as felspathoid, calcic-pyroxene, and carbonate assemblages.
Nephelinite lavas are observed in oceanic settings
such as the Cape Verde and Hawaiian islands, but
are best seen in young (Paleocene to recent) continental volcanic settings such as the East African

27

rift valley, central Europe and southeast Australia.
Old alkaline igneous complexes are rare, oneof the best preserved being the 2050 Myr old
Phalaborwa Complex in South Africa, which ismined for copper and phosphate as well as a hostof minor by-products.
Nephelinite, as well as the associated, but rare,
carbonatite melts (i.e. magmas comprising essentially CaCO3 and lesser Na2CO3), are undoubtedly
primary magma types derived from the mantle byvery low degrees of partial melting under conditions of high Ptotal and PCO2 (Hall, 1996). The relationship between nephelinitic and carbonatitic
magmas is generally attributed to liquid immiscibility, whereby an original alkali-rich silicate
magma rich in a carbonate component exsolves
into two liquid fractions, one a silicate and theother a carbonate (Ferguson and Currie, 1971; Le
Bas, 1987). Low degrees (2%) of partial melting ofa garnet lherzolitic source in the mantle will typically yield olivine nephelinite compositions andthese magmas may be spatially and temporally
associated with basaltic volcanism (Le Bas, 1987).
Nephelinite magma associated with carbonatite,on the other hand, is only considered possible
if the source material also contained a carbonate
phase (such as dolomite) and a soda-amphibole.
This type of mantle source rock is likely to be theresult of extensive metasomatism, a process thatinvolves fluid ingress and enrichment of volatileand other incompatible elements. Melting of a
fertile mantle source rock is probably the main
reason why alkaline magmas are so enriched in
the variety of ore constituents mentioned above.
The extent of metal enrichment relative to average basalt is illustrated in Figure 1.6.
Kimberlitic and related ultramafic magmas
crystallize to form very rare and unusual rocks,
containing among other minerals both mica
and olivine. Kimberlites are rich in potassium(K2O typically 1–3 wt%) and, although derived
from deep in the mantle, are also hydrated andcarbonated. They usually occur in small (<1 km
diameter) pipe-like bodies, or dykes and sills, andcommonly extrude in highly explosive, gas-charged
eruptions. The deep-seated origin of kimberlite isevident from the fact that it commonly transports
garnet lherzolitic and eclogitic xenoliths to the28PART 1 IGNEOUS PROCESSES

700
BasaltAbundance (ppm)

600

Alkaline magma/kimberlite

500
400
300
200
100
0

F

P/10

Cu

Nb

Zr/10

surface, rock types made up of very high pressuremineral assemblages that could only have comefrom the mantle. In addition, a small proportionof kimberlites also contain diamond xenocrysts.
Diamond is the stable carbon polymorph undervery reducing conditions, and at depths in excess
of about 100 km and temperatures greater than
900 °C (Haggerty, 1999).
The origin of kimberlitic magmas is not too different from that of the alkaline rocks describedabove, and high pressure partial melting of agarnet peridotite source rock containing additional
phlogopite or K-amphibole (richterite), as well asa carbonate phase, is viewed as a likely scenario(Hall, 1996). The enrichment of incompatible
constituents (such as K, Rb, H2O, and CO2) inkimberlite, as with alkaline magmas in general,again indicates that metasomatism of the mantle
has played an important role in the provision ofa deep-seated environment capable of producing
highly enriched, or fertile, magmas. These aspects
are discussed in more detail in section 1.3 below.

1.3 WHY ARE SOME MAGMAS MORE FERTILE THAN
OTHERS? THE “INHERITANCE FACTOR”

Geochemical inheritance is clearly an importantfactor in understanding the nature of ore-forming
processes in igneous rocks. Magma may inherit
a surplus of potential ore-forming trace elements
because the source material from which it wasderived was itself enriched in these components.

U × 10Th × 10Figure 1.6 Relative abundances of
selected metals in alkaline magmas(and kimberlite in the case of Cu
and P) relative to average basalt(data from Table 1.2).

Further concentration of incompatible trace elements into residual magma, or of compatible traceelements into crystallizing phases, will take place
during cooling and solidification of the magma,
and these processes are discussed in more detailin section 1.4 below.
A pertinent question that relates to the issue ofgeochemical inheritance is why certain portionsof the Earth’s crust appear to be so much better
endowed in mineral deposits than others. The
spectacular concentrations of, for example, goldand platinum deposits in South Africa, related
to the Witwatersrand Basin and Bushveld Complex respectively, perhaps point to some form of(mantle-related?) enrichment of these metals thatis specific to this region and not to other partsof the globe. Why is it that Sudbury is so rich in
Ni and the Andes are so well endowed with large
Cu deposits, and is there an explanation for thetendency that diamondiferous kimberlites only
occur in ancient, cratonic areas? These questions
are addressed below in relation to some rather
novel ideas suggesting that metals may have been
preferentially concentrated in certain parts of the
mantle and then subsequently transferred into thecrust through a variety of mineralizing agencies.
1.3.1 The “late veneer” hypothesis of siderophile
metal concentration – an extraterrestrial origin
for Au and Pt?
In the very early stages of Earth evolution,
prevailing theory suggests that the originally

IGNEOUS ORE-FORMING PROCESSES CHAPTER 1

to – core/mantle segregation29

to + 500 Myr
intense meteorite flux

Au
Pt
Pt

Pd

Fe, Ni, Au, Pt

Au"Late veneer" ofsiderophile element
enrichment (Au, Pt, Pd, etc.)
Figure 1.7 Schematic representation of the “late veneer” hypothesis for the siderophile (precious) metal enrichmentof the Earth’s lithosphere. During initial segregation of the Earth (to) siderophile metals were comprehensively
partitioned into the Fe–Ni core. Approximately 500 million years later (to + 500 Myr) intense meteorite
bombardment of the Earth added to the siderophile metal budget of the Earth’s lithosphere.

homogeneous, molten planet differentiated into a
metallic core, comprising essentially Fe and FeO
with lesser Ni, and a mantle which had a silicatecomposition. As this differentiation took place
the siderophile metals (i.e. those with a strong
affinity for Fe such as Au and the platinum groupelements or PGE) were comprehensively partitioned into the core. Experiments by Holzheid
et al. (2000) indicate that the average concentrations of elements such as Au, Pt and Pd in the
Earth’s mantle should be at least 10 −4 times lower
than average chondritic abundances. Such concentrations are, in fact, so low that they virtually
preclude the possibility of forming ore deposits
in rocks extracted from the mantle (i.e. the crust).
Yet the actual concentration of these precious
metals in the mantle, although depleted, is only
about 150 times lower than average chondritic
abundances. This depletion might be accommodated by the fact that there have been numerous
ore deposits formed over geological time that haveextracted these metals from the mantle. Anotherexplanation is that the efficiency with whichsiderophile metals are partitioned between metal

core and silicate mantle decreases with increasingpressure (depth) and that this could explain why
the mantle is not as depleted as theory predicts.
Although the latter notion probably applies tonickel, recent experiments suggest that it is notapplicable to the precious metals and that someother explanation must be sought for the higher
than expected concentrations of the latter in the
mantle.
A clue as to why the mantle might be relatively
enriched in siderophile metals lies in the fact that
their abundance ratios (i.e. the abundance of one
element relative to another, such as Au/Pt or Pt/Pd)
are generally similar to chondritic abundance ratios
as determined from analyses of meteorites thathave fallen to Earth. The only way to explain thisis by having a substantial proportion of the precious metals in the mantle derived from meteorites
that impacted the proto-crust during the earlystages of Earth evolution, but after the differentiation of core and mantle (Figure 1.7). This idea,
known as the “late-veneer” hypothesis (Kimura
et al., 1974), suggests that much, if not all, theAu and Pt that is mined from ore deposits on the

30

PART 1 IGNEOUS PROCESSES

Earth’s surface today ultimately had an extraterrestrial origin and that the planet’s own inventoryof these metals is presently locked away in thecore. Since both the meteorite flux itself and thesubsequent distribution of this material throughthe mantle are likely to have been irregular, thishypothesis is also consistent with the heterogeneous distribution of precious metals over the
Earth’s surface.
As a footnote it is intriguing to note that precious metals may not have been the only valuablecommodity introduced to Earth by meteorites.
The enigmatic “carbonado” diamonds found onlyin 1500 Myr old metasediments of Bahia Statein Brazil and the Central African Republic havemineralogical and isotopic characteristics unlike
any diamond of terrestrial origin. Haggerty (1999)
has suggested that carbonado diamonds are derived
from the fall-back of a fragmented carbon-type
asteroid that impacted the Earth’s crust at atime when the relevant parts of Brazil and Africa
formed a single continental entity. Although contentious, the notion of an extraterrestrial origin
for certain constituents of the Earth’s surface (andof life itself?) is one that is likely to continue
attracting attention in the future.
1.3.2 Diamonds and the story they tell
The Earth’s mantle, between about 35 and2900 km depth, is the ultimate source of material
that, over geological time, has contributed, either
directly or by recycling, to formation of thecrust. Considering that the mantle is essentially
inaccessible it has, nevertheless, been the subject of numerous studies by both geologists andgeophysicists and its structure and composition
are now reasonably well known. One reason forstudying the mantle is to understand the originof diamond. This remarkable mineral, together
with the magmas that bring it to the surface,has provided a great deal of information about
the deep Earth, much of which is very relevant tounderstanding those properties of the mantle that
also relate to the source of metals in other igneousore deposits.
Most diamonds are brought to the Earth’s surface by kimberlitic magmas (see section 1.2) or acompositionally similar melt known as lamproite.
Most kimberlites and lamproites are barren, anddiamondiferous magmas only intrude into ancient,stable continental crust that is typically older than2500 Myr, but sometimes as young as 1500 Myr.
The kimberlite magmas that transport diamonds
to the surface, however, are typically much younger
than the rocks they intrude, forming in discrete
episodes in the Mesozoic and Cenozoic eras. Older
intrusive episodes have also been observed in the
Devonian, as well as at around 500 Ma and again
at 1000 Ma (Haggerty, 1999). Diamondiferous
kimberlites must also have been emplaced during
the Archean, since the Witwatersrand conglomerates in South Africa, for example, are known tocontain green detrital diamonds. To further complicate the story, the diamonds themselves tend
to be much older than their kimberlitic host rocks
and range in age from 1500 to 3000Ma, indicatingthat they have resided in the mantle for considerable periods of time prior to their eruption onto
the Earth’s surface. Diamonds did not, therefore,
crystallize from the kimberlite but were introduced to the Earth’s surface as xenocrysts within
the magma (Richardson et al., 1984). Diamond
xenocrysts occur either as isolated single crystals
in the kimberlitic matrix or as minerals within
discrete xenoliths of either peridotite (P-type
diamonds – the more common) or eclogite (E-type
diamonds). The high pressure phase relations thatcharacterize mineral assemblages in these mantle
xenoliths indicate that diamonds are derived fromzones of thickened, sub-cratonic lithosphere, atleast 200 km thick, that extend beneath stable
Archean–Proterozoic shield areas (Figure 1.8). Theselithospheric keels comprise old, depleted, peridotite (i.e. from which mantle melts have alreadybeen extracted) as well as primitive, but younger,
eclogite that has generally not had a melt fraction
extracted and is, therefore, more fertile with respect
to crust-forming elements (Haggerty, 1999).
It is now generally accepted that diamonds weregenerated from deep in the mantle, in the layer
known as the Transition Zone between the lower
and upper mantle at around 400 to 650 km depth.
Because the upper mantle is relatively depleted incarbon (100 ppm compared with 1000–3700 ppm
in the lower mantle; Wood et al., 1996) it isIGNEOUS ORE-FORMING PROCESSES CHAPTER 1

31

LITHOSPHERIC KEEL (ARCHEARATONICN?)
SUB-C
KIMBERLITESLAMPROITE
LT
IC BE
GEN
O
R
O
CRUST
C
0°
Moho
90 C
°
Spinel00
12Harzburgite
LherzoliteC
Spinel°00
Lherzolite
15
Dunite100

Eclogite
E
HIT
P
A
D
GR MON UPPER MANTLE
A
I
D

200

300

CO2
CO2 Eclogite
Metasomatized
H2O
LILE
mantle
H2O
LILE
PLUME
LILE DEPLETED
CO2
H2O
fO2 10–2 to 10–1 FMQ
LILE

H2O: 200 ppm
C: 100 ppm

km

PLUME
400

500

600

TRANSITION ZONELOWER MANTLE

LILE ENRICHEDH2O: 520–1900 ppm
C: 1000–3700 ppm

LILE
CO2

H 2O

MgCO3

fO2>FMQ

H2O

Carbon as
CO2, MgCO3, CCO
Figure 1.8 Schematic diagram illustrating features pertinent to the formation of diamond and the fertilizationof the Earth’s mantle by plume-related magmas and their associated aqueo-carbonic fluids (after Haggerty, 1999).
LILE refers to the large ion lithophile elements; FMQ refers to the fayalite–magnetite–quartz oxygen buffer.

unlikely to be a viable source for the primordialcarbon that makes up diamond. The more fertile
lower mantle is more likely to be the source ofthe carbon, and this is supported by the presenceof very high pressure minerals occurring as tiny
inclusions in many diamonds. However, the upper
mantle is more reduced than the lower mantle,
which, in addition to its high carbon contents, also
contains substantially more water (500–1900 ppm
compared to only 200 ppm in the upper mantle).
The upper mantle is, therefore, more likely to preserve diamond because the mineral’s long-term
stability depends on the existence of a reducing
environment. Carbon in the relatively oxidized,
fluid-rich lower mantle would, despite the higher

pressures, not occur as diamond at all, but as CO2,
CCO, or MgCO3 (Wood et al., 1996). The modelfor diamond formation (Figure 1.8), therefore, suggests that plumes transfer melt and volatiles fromthe lower mantle, and precipitate diamond at
higher levels either in the reduced environment
represented by the Transition Zone or in the keels
extending below thick, cratonic lithosphere.
Thus, the more common P-type diamonds form
when the relatively oxidized carbonic fluids dissolved in ascending plumes interact with reduced
mantle at higher levels and precipitate elemental
carbon. This mass transfer process is referred toas metasomatism and involves the movementof fluids and volatiles from deep in the Earth’s

32PART 1 IGNEOUS PROCESSES

mantle to higher levels. This process is turningout to be very relevant to the concepts of mantle
fertilization and geochemical inheritance. The
more rare E-type diamonds, by contrast, are considered to have crystallized directly from amagma intruded into or ponded below the keels(Haggerty, 1999).
Formation of the kimberlitic magma that transports diamond to the Earth’s surface has also been
attributed to plume activity and the metasomatictransfer of volatile constituents from a fertile lower
mantle into depleted upper mantle. Evidence forthis process comes from the observation that manyof the major episodes of kimberlite intrusion mentioned above correlate with “superchron” events
that are defined as geologically long time periodsof unidirectional polarity in the Earth’s magnetic
field. Superchrons are caused by core–mantle
boundary disruptions which increase the rate ofliquid core convection, causing a damping of thegeomagnetic field intensity but promoting plumeactivity and mantle metasomatism. Intrusion of
diamondiferous kimberlites has also been linked
in time to major geological events, such as continental break-up and flood basaltic magmatism(Haggerty, 1994). England and Houseman (1984)
suggested that enhanced kimberlite intrusion
could be related to periods of low plate velocity
when uninterrupted mantle convection gave rise
to partial melting and volatile production in thelithosphere, and the subsequent formation ofplume activity. Accompanying epeirogenic upliftcreated the fractures that allowed kimberlite
magma to intrude rapidly upwards and, in manycases, to extrude violently onto the Earth’s surface.
This explanation is certainly consistent withthe geodynamic setting of kimberlites, such as
their predominance in Africa during the Mesozoic
still-stand (see Chapter 6). The relative rarity ofkimberlite formation and penetration to the surface

Diamondiferous kimberlites and lamproites: The
Orapa diamond mine, Botswana and the Argyle
diamond mine, Western Australia
Diamond mining around the world was worth some
US$7 billion in 2001, a substantial proportion of which
was derived from exploitation of primary kimberlitic andlamproitic deposits. The biggest single deposit is at Argyle
in Western Australia, which is hosted in lamproite andproduces some 26 million carats per year. Most of the
diamonds produced here, however, are of low value. The
Orapa and Jwaneng deposits of Botswana, by contrast,
produce less than half the number of carats per year, but
their stones are much more valuable. Orapa and Jwaneng
together are the richest diamond deposits in the world.
Kimberlites are by far the most important primary sourceof diamonds (see section 1.3.2) and there are over 5000
occurrences known world wide (Nixon, 1995). By contrast
there are only some 24 known occurrences of lamproite.
Both kimberlites and lamproites are emplaced into the
Earth’s crust as “diatreme–maar” volcanoes which are
the product of highly overpressured, volatile-rich magma.
Kimberlitic or lamproitic magma is injected into the crust,along zones of structural weakness, to within 2–3 km of thesurface. At this point volatiles (H2O and CO2) either exsolvefrom the magma itself (see Chapter 2), or the magma
interacts with groundwater, with the resulting vapor phasecausing violent phreatomagmatic eruption of magma anddisruption of country rock. Figure 1 shows the anatomy
of a diatreme–maar system that has applicability to thenature and geometry of both kimberlites and lamproites.
Kimberlites are richer in CO2 than lamproites and since
CO2 has a lower solubility than water in silicate melts,
kimberlite magmas will usually exsolve a volatile fraction
at lower depths than lamproites (Nixon, 1995). Lamproite
venting is quite often a function of magma interaction withgroundwater, the availability and depth of which dictates
the geometry of the crater. These factors account for the
carrot-like shape of kimberlites compared to a broader,champagne glass shape for many lamproites.
The distribution of diamonds in any one kimberlite
may be highly erratic and there seems to be little or
no relationship between grade and depth (Nixon, 1995).
Some hypabyssal dykes are very rich in diamonds, suchas the Marsfontein mine in South Africa which has
an average grade of 200–300 carats per 100 tons of ore
mined. Diatreme facies kimberlites are often characterizedby multiple injections of magma, some of which are
barren and others economically viable. Kimberlitic pyroclastic sediments in crater facies are also often richer than

IGNEOUS ORE-FORMING PROCESSES CHAPTER 1

Fallback material
Maar
Crater facies

Hypabyssal fecies

Depth (km)Figure 1 Idealized geometry ofa diatreme–maar type volcano,showing the nature of thehypabyssal facies of magma
emplacement along crustal
weaknesses (after Smith et al.,1979). Kimberlites around theworld are seen as hypabyssal,
diatreme, or crater facies,
depending on the extent ofpreservation (or depth of erosion)of the bodies.

Diatreme facies

0
1 Kimberlite
breccia
Diatreme

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

Tuffaceous kimberlite

2

Sill
3

4
Dyke
5

associated diatreme rocks, possibly due to enrichmentof heavy minerals, including diamond, by wind (during
eruption) or water (in the crater lake).
ArgyleThe Argyle diamond mine occurs in a 1200 million year
old lamproite diatreme that has intruded older Proterozoic
sediments in the Kimberley region of Western Australia(Boxer et al., 1989). It forms an elongate 2 km long bodythat varies from 100 to 500 m in width and is the result ofat least two coalesced vents along a fault line. It is minedin the southern section where it has grades in excess of
5 carats per ton. Most of the body is made up of pyroclastic
or tuffaceous material, with marginal breccia and occasional lamproitic dykes. The diatreme was formed whenlamproite magma encountered groundwater in largely
unconsolidated sediments, resulting in multiple phreatomagmatic eruptions and venting of lamproite at the surface.
Orapa
The Orapa diamond mine occurs in a mid-Cretaceous
kimberlite in the north of Botswana, and is well known
for the excellent preservation of its crater facies. The
kimberlite was emplaced in two pulses that merge at
about 200 m depth into a single maar. The diatremes, comprising tuffisitic kimberlite breccia, grade progressively into

Figure 2 Large boulder of country rock in tuffisitic
and lacustrine maar sediments, Orapa Mine,
Botswana.

crater facies made up of both epiclastic and pyroclastic
kimberlite debris. All these phases are diamondiferous.
At Orapa the northern diatreme is believed to have been
emplaced first, followed by residual volatile build-up and
explosive volcanic activity. This was shortly followed by asimilar sequence of events to form the southern diatreme,with the subsequent merging of its crater facies into asingle maar (Field et al., 1997).

33

34

PART 1 IGNEOUS PROCESSES

LadolamCu-Au-deposit
Lihir island
arc
H2O
Ontong-Java plateau
AlkalibasaltCO2

Magma
production

Country rocks
Cu Au Pt Pd
61 0.79 1.75 3.00 b
145 0.3 15.50 21.3 s

Metasomatized mantle
Cu Au Pt Pd
54 1.73 90.2 41.7

H2O H2O
De
CO2
hy
dra
tio
n
sla of su
b
bd u
cte
d

Unalteredharzburgite
Cu Au Pt Pd
9 0.04 2.60 0.09

Figure 1.9 Schematic diagram illustrating the concept of mantle metasomatism and metal enrichment associatedwith subduction, and the subsequent inheritance of an enhanced metal budget by magmas derived from melting ofmetasomatized mantle. Metal abundances of the relevant rock types (in ppm) are from McInnes et al. (1999); the twoanalyses showing metal abundances for the magmatic products of subduction refer to basalt (b) and syenite (s).

is also consistent with this model, since magma
formation requires that a number of coincidentally optimal conditions apply.
1.3.3 Metal concentrations in metasomatizedmantle and their transfer into the crust
Although it has been evident for many years that
there is an association between mantle metasomatism and diamond formation, a link to mantle
fertilization with respect to other constituents,
such as the base and precious metals, has only
recently been suggested. A comparison of Re–Os
isotopic ratios of epithermal Cu–Au ores from thegiant Ladolam mine on Lihir Island, near Papua
New Guinea, with peridotite xenoliths transported
to the ocean surface by volcanic activity from theunderlying mantle wedge, shows that the ore constituents are derived from the mantle. Although
this is not unexpected, closer study of the peridotitic material involved reveals that some of thismaterial has been extensively metasomatized toform a high temperature hydrothermal mineral
assemblage comprising olivine, pyroxene, phlo-

gopite, magnetite, and Fe–Ni sulfides (McInnes
et al., 1999). Metasomatism is considered to be
the result of dehydration of the oceanic slab as it
moves down the subduction zone, yielding fluidsthat migrate upwards into the overlying mantle
wedge. Metasomatized peridotite contains precious and base metal concentrations that are up
to two orders of magnitude enriched relative to
unaltered mantle (Figure 1.9).
Subduction of oceanic crust beneath an islandarc, such as that of which Lihir Island is part, hasresulted in the formation of alkaline basalt whichbuilds up the arc and ultimately forms the hostrocks to the Ladolam Cu–Au deposit. The deposit
itself occurs at a high level in the crust by thecirculation of metal-charged hydrothermal fluids,
a process that is not relevant to the present discussion except for the fact that these fluids havedissolved the ore constituents that they carry fromthe immediate country rocks (i.e. the alkaline
basalts). An interesting feature of the study by
McInnes et al. (1999) is that basaltic and syenitic
country rocks to the deposit (refer to b and s in
Figure 1.9) were analyzed well away from theIGNEOUS ORE-FORMING PROCESSES CHAPTER 1

mineralization itself and also exhibit significant
enrichments in their base and precious metal contents. The inference is, therefore, that although
the ore metals resided originally in the mantle,
they were redistributed, and their concentrations
significantly upgraded, by the high temperature
metasomatic processes observed in the mantle
peridotite. Because it is the metasomatized peridotite that is likely to have been preferentiallymelted above the subduction zone (by virtue of its
hydrous nature and lower solidus temperature), theresulting alkali basalt magma will have inherited
a significant metal endowment, thus enhancing
the chances of creating a substantial ore deposit
during subsequent (hydrothermal) stages of ore
formation. The Ladolam study illustrates threeimportant features: first, that metasomatism andfluid flow are very important processes in redistributing and concentrating incompatible elementsin the mantle; second, that partial melting is themain process by which matter is transferred fromthe mantle to the crust; and, third, that inheritance is critical to the nature and formation of
igneous hosted ore deposits, as well as to whether
subsequent fluid circulation is likely to form
viable hydrothermal deposits.
1.3.4 I- and S-type granite magmas andmetal specificity
As mentioned previously, the different types of
granite, and more specifically the origins of felsicmagma, can be linked to distinct metal associations. Of the many classification schemes thatexist for granitic rocks one of the most relevant,with respect to studies of ore deposits, is the I- andS-type scheme, originally devised for the Lachlan
Fold Belt in southeast Australia (Figure 1.10a)by Chappell and White (1974). In its simplest
form the scheme implies that orogenic granites
can be subdivided on the basis of whether theirparental magmas were derived by partial meltingof predominantly igneous (I-type) or sedimentary(S-type) source rock. In general, I-type granites
tend to be metaluminous and typified by tonalitic(or quartz-dioritic) to granodioritic compositions,whereas S-types are often peraluminous and haveadamellitic (or quartz-monzonitic) to granitic com-35

positions. Also very important from a metallogenic
viewpoint is the fact that I-type granites tend to
be more oxidized (i.e. they have a higher magmaticfO2) than S-type granites, whose magmas wereoriginally fairly reduced because of the presenceof graphite in their source rocks. An approximate
indication of the oxidation state of granitic magmas
can be obtained from their whole rock Fe2O3/FeO
ratio (which effectively records the ferric/ferrous
ratio). Blevin and Chappell (1992) have shownthat an Fe2O3/FeO ratio of about 0.3 provides auseful discriminant between I- (with Fe2O3/FeO> 0.3) and S-type (with Fe2O3/FeO < 0.3) granites,at least for the Australian case (Figure 1.10 b andc). A classification of granites according to oxidation state was, in fact, made relatively early on
by Ishihara (1977), who distinguished between
reduced granite magmas (forming ilmenite-series
granitoids) and more oxidized equivalents (forming
magnetite-series granitoids). The metallogenic significance of this type of granite classification was
also recognized by Ishihara (1981), who indicatedthat Sn–W deposits were preferentially associatedwith reduced ilmenite-series granitoids, whereas
Cu–Mo–Au ores could be linked genetically tooxidized magnetite-series granitoids. Magnetiteseries granitoids are equivalent to most I-types,
whereas ilmenite-series granitoids encompass all
S-types as well as the more reduced I-types.
Although now regarded as somewhat oversimplified, at least with respect to more recent
ideas regarding granite petrogenesis, the S- andI-type classification scheme is nevertheless appealing because it has tectonic implications and canbe used to infer positioning relative to subduction along Andean-type continental margins (seeFigure 1.2). The scheme also has metallogenic
significance because of the empirical observationthat porphyry Cu–Mo mineralization (with associated Pb–Zn–Au–Ag ores) is typically associatedwith I-type granites, whereas Sn–W mineralization(together with concentrations of U and Th) is
more generally hosted by S-type granites. Although
this relationship is broadly applicable it, too, isoversimplified. Some granites, notably those thatare post-tectonic or anorogenic, do not accord withthe scheme, such as the alkaline granites of the
Bushveld Complex, which are polymetallic andPART 1 IGNEOUS PROCESSES

Fractionation

10

Canberra
MainlyMo and
Cu-Au

1

I
S

0.1

lin
e
I-S

Melbourne

I-type granite
S-type granite
0

100
km

Fe2O3/FeO

10
MainlySn-W

Oxidation

Fe2O3/FeO

(a)

(b)

Sn
Mo
Au

(c)

1

I

Oxidation

36

S

0.1

Cu
W

200
0.01

0

100

200

300 400
Rb ppm500

600

Figure 1.10 (a) Simplified map showing the distribution of S- and I-type granites, and associated metallogenic
trends, in the Lachlan Fold Belt of southeastern Australia (after Chappell and White, 1974). (b and c) Plots of
Fe2O3/FeO versus Rb for granites of the Lachlan Fold Belt that are mineralized with respect to Sn–W and Cu–Mo–Au(after Blevin and Chappell, 1992).

contain both Sn–W and base metal mineralization (Robb et al., 2000). A more accurate appraisalof the relationship between magma composition (including oxidation state) and metallogenic
association is given by Barton (1996). Figure 1.11
shows Barton’s scheme, which regards granites
as a continuum of compositional types and their
metal associations in terms of different intrusionrelated ore deposit types.
Adjacent to Andean type subduction zones aclearly defined spatial pattern exists with respect
to the distribution of I- and S-type granite intrusions, as well as associated metallogenic zonation(Sillitoe, 1976; Clark et al., 1990). The leading edge
(i.e. the oceanic side) of the subduction zone tends
to correlate with the production of I-type granite
magmas and is associated with the formation ofporphyry Cu styles of mineralization. By contrast,
the continental side of the subduction zone contains more differentiated granite types that areoften S-type in character and with which Sn–W
styles of mineralization are associated. Other
examples of this type of regional zonation are seenin the Cordilleran granites of the western UnitedStates and in the Lachlan Fold Belt of southeast
Australia, where I- and S-type granites were originally defined (Figure 1.10a). Again, although exceptions and complications do exist, the recognitionand delineation of these patterns is clearly important with respect to understanding the spatial distribution of different types of ore deposits hosted
in granitoid rocks.
The Lachlan Fold Belt provides an excellent
example of the relationships between magma typeand metallogenic association and, as predicted,
granite related mineralization is dominated by
Sn–W in the mainly S-type granites to the west of
the I–S line, whereas largely Mo with lesser Cu–
Au ores are found in the I-type terrane to its east
(Figure 1.10a). Another very important feature of
the metal content of granites is, however, also
apparent in the Lachlan Fold Belt. When the Fe2O3/
FeO ratio (an indication of magmatic oxidation
state) is plotted against Rb content (an indicatorof degree of fractionation) for granites that aremineralized in terms of either Sn–W or Cu–Mo–Au

IGNEOUS ORE-FORMING PROCESSES CHAPTER 1

Mafic

FelsicGraniteSn, W
Zn, F, Mo (Ag, Be)

E
LITHOPHIL

Metaluminous

Cu (Mo)Zn, W (Pb, Ag)
Cu (Au, Mo)
Cu, Zn,(Au, Ag, Fe)

'

HYRY

'PORP

Cu, (Au, Fe)

Mo, F (W)Zn, F, Ag, Be

Au
C

ALI

ALK

F, Li, Be

Monzodiorite

Monzonite

Qtz-monzonite/adamellite

Reduced(low magmatic fO2)

Granodiorite

Oxidized(high magmatic fO2)

Qtz-diorite/tonalite

Peraluminous

DioriteParalkaline/
qtz undersaturated

37

Alkali granite

Figure 1.11 Generalized scheme that links granite compositions and magmatic oxidation state to metal
associations and intrusion-related ore deposit types (modified after Barton, 1996). Ore deposit types referred to are“alkalic,” “porphyry,” and “lithophile” and are discussed in more detail in Chapter 2. Metals shown in bold reflect
the more important associations.

(Figure 1.10b and c) it is clear that metal content is
not simply a function of magma type alone. It isapparent that Cu–Mo–Au related intrusions typically have higher Fe2O3/FeO ratios than those associated with Sn–W mineralization, and are, therefore,
preferentially associated with I-type granites.
What is perhaps more clearly apparent, however,is that Sn and W mineralization is associatedwith intrusions that are more highly fractionatedthan those containing Cu–Mo–Au (Blevin andChappell, 1992). Metal contents are, therefore,
also a function of processes that happen as themagma cools and fractionates and these processes
are discussed in more detail in section 1.4 below.
In fact, mineralization in granites also involveshydrothermal processes that are quite distinct

from either geochemical inheritance or fractionation, and these are examined in Chapter 2.

1.4 PARTIAL MELTING AND CRYSTAL FRACTIONATION AS
ORE-FORMING PROCESSESThe previous section discussed the various magmaforming processes and some of the reasons why
they are variably endowed with respect to theirtrace element and mineral contents. This section
examines how trace elements and minerals behavein the magma, both as it is forming and then subsequently during cooling and solidification.
Trace element abundances can be very good
indicators of petrogenetic processes, during both38

PART 1 IGNEOUS PROCESSES

partial melting and crystal fractionation. Sincemany magmatic ore deposits arise out of concentrations of metals that were originally present
in very small abundances, trace element behavior
during igneous processes is also very useful inunderstanding ore formation.
A trace element is defined as an element thatis present in a rock at concentrations lower than
0.1 wt% (or 1000 ppm), although this limit places
a rather artificial constraint on the definition. In
general trace elements substitute for major elements in the rock forming minerals, but in certain
cases they can and do form the stoichiometriccomponents of accessory mineral phases (Rollinson,1993). Many of the ores associated with igneousrocks are formed from elements (Cu, Ni, Cr, Ti, P,
Sn, W, U, etc.) that originally existed at trace
concentrations in a magma or rock and were subsequently enriched to ore grades by processes discussed in this and later sections.
When rocks undergo partial melting trace elements partition themselves between the melt phaseand solid residue. Those that prefer the solid arereferred to as compatible (i.e. they have an affinity
with elements making up the crystal lattice of an
existing mineral), whereas those whose preferenceis the melt are termed incompatible. Likewise,
during cooling and solidification of magma, compatible elements are preferentially taken up in
the crystals, whereas incompatible elements areenriched in the residual melt. Enrichment of traceelements and potential ore formation can, therefore, be linked to the concentration of incompatible
elements in the early melt phase of a rock undergoing anatexis (see Box 1.2), or the residual magma
during progressive crystallization (see Box 1.3).
Compatible trace elements tend to be “locked up”
in early formed rock-forming minerals and are
typically not concentrated efficiently enough toform viable ore grade material. An exception to thisis provided, for example, during the formation ofchromitite layers, as discussed in section 1.4.3
below.
A brief description of how one can quantify
trace element distribution during igneous processes is presented below, with an indication of
how these processes can be applied to the understanding of selected ore-forming processes.

1.4.1 The conditions of melting
Despite the fact that temperatures in the upper
mantle reach 1500 °C and more, melting (oranatexis) is not as widespread as might be expected
because of the positive correlation that exists
between pressure and the beginning of melting ofa rock (i.e. the solidus temperature). The asthenosphere, which is defined as that zone in the mantle
where rocks are closest to their solidus and where
deformation occurs in a ductile fashion (which
explains the lower strength of the asthenosphererelative to the elastically deformable lithosphere),is the “engine-room” where a considerable amountof magma is formed. Major magma-generating
episodes, however, do not occur randomly andwithout cause, but are catalyzed by processes suchas a decrease of pressure (caused, for example,
by crustal thinning in an extensional regime suchas the mid-ocean ridge) or addition of volatiles
to lower the solidus temperature (such as duringsubduction and metasomatism). An increase inlocal heat supply is generally not important in
the promotion of partial melting, although it ispossible that mantle plumes might be implicatedin this role.
Partial melting, so called because source rocks
very seldom melt to completion, invariably leaves
behind a solid residue. It is a complex processthat is affected by a number of variables, themost important being the nature of the mineral
assemblage making up the protolith, as well as
local pressure, temperature and water/volatile content. Early experimental work on the progressive
fusion of peridotite with increasing temperature(Mysen and Kushiro, 1976) helps to explain theprocess of partial melting in the asthenosphere,
even though the description outlined below is
probably not a particularly good indication of thevery complicated processes actually involved.
In Figure 1.12 melting of peridotite starts justabove 1400 °C with the formation of a small
melt fraction in equilibrium with pyroxenes andolivine. Melting can continue without significant
addition of heat until between 30 and 40% of therock is molten and the clinopyroxene is totallyconsumed. Once the clinopyroxene is gone, further
melting can only continue with addition of heat (theIGNEOUS ORE-FORMING PROCESSES CHAPTER 1

39

1800
(a)
Anhydrous

Temperature (°C)

1700

20 kbar(b)

Ol + Liq

With 1.9 wt% H2O
1600
Ol + Opx + Liq
Ol + Opx +Cpx + Liq
1500

Figure 1.12 The sequential
melting behavior of peridotite in
the mantle at 20 kbar pressure,(a) without water and (b) in thepresence of some 2% water(after Mysen and Kushiro, 1976).

1400
0

10

first inflection point on the curves in Figure 1.12),
which, if present, would then promote the melting of orthopyroxene. Once orthopyroxene has
been consumed, after some 50–60% of the rockhas melted, a further input of heat is required (thesecond inflection point) if olivine is to be wholly
included into the melt product. A small degree ofpartial melting of peridotite will, therefore, yieldmagma with a composition that is dominatedby the melt products of clinopyroxene (an alkali
basalt), with the residue reflecting the bulk composition of orthopyroxene + olivine. In reality
melting processes are more complex and usually
involve the fusion of more than just one mineral ata time. It is also pertinent to note that the presenceof even a small amount of water in the system
will catalyze the extent of melting and also allow
anatexis to occur at lower temperatures.
Anatexis in the crust is better explained by considering the partial melting of a metasedimentary
protolith, to form, for example, an S-type granite.
If the sediment consisted only of quartz then thetemperature would have to exceed 1170 °C (themelting temperature of SiO2 at PH2O = 1 kbar) inorder for it to start melting. If the sediment were

20

30

40

50
Melt (%)

60

70

80

90

100

Partial melt

Qz

Qz

Or
Ab

Ab

Ab
Qz

Or

Or
Ab

Qz

Or
Qz

Or

Qz

No melting
Figure 1.13 The pattern of grain boundary meltingof an arkosic protolith subjected to temperatures of
700–800 °C at PH2O = 1 kbar (after Hall, 1996).

an arkose, however, made up of a recrystallized
assemblage of quartz + orthoclase + albite (as shownin Figure 1.13), then melting would commence atmuch lower temperatures because of the depression of melting points that occurs for binary or
ternary eutectic mixtures. Thus, where quartz +albite are in contact, melting would commence at40

PART 1 IGNEOUS PROCESSES

temperatures around 790 °C and, where quartz +albite + orthoclase meet at a triple junction, melting could start as low as 720 °C. Disaggregationof the arkose protolith would, therefore, takeplace by small increments of partial melt forming
along selected grain boundaries within the rock.
The residue left behind during such a process islikely to be made up of fragmented minerals thatcannot melt on their own at a given temperature.
Melt and residue are also likely to have different
compositions.
The extraction of a partial melt from its residue,
whether it be from an igneous or sedimentary
protolith, is a process which segregates chemicalcomponents and is referred to as fractionation.
Partial melts can be considerably enriched incertain elements, but depleted in others, relative
to the source rock. The mechanisms whereby melt
is removed from its residue may differ, and this isdiscussed in more detail below. More detailed,quantitative descriptions of partial melt modellingare provided in Cox et al. (1979), Rollinson (1993),
and Albarede (1996).
Trace element distribution during
partial melting
In theory there are two limiting extremes by which
partial melting can occur. The first envisages formation of a single melt increment that remains inequilibrium with its solid residue until physicalremoval and emplacement as a magma. This process may be applicable to the formation of highviscosity granitic melts (Rollinson, 1993) and is
known as “batch melting.” It is quantified by thefollowing equation, whose derivation (together
with others discussed below) is provided in Woodand Fraser (1976):
Cliq/Co = 1/[Dres + F(1 − Dres)]enrichment (or depletion) of an incompatible (orcompatible) element in a batch partial melt ispresented in Figure 1.14a, where the ratio Cliq/Co
is plotted as a function of the degree of melt produced (F) for a variety of values of Dres. Marked
enrichments of highly incompatible elements(i.e. those with very small values of Dres) can
occur in small melt fractions, with the maximum
enrichment factor being 1/Dres as F approaches 0.
The second partial melt process is referred to as“fractional melting” and is the process whereby
small increments of melt are instantaneously
removed from their solid residue, aggregating elsewhere to form a magma body. This process may be
more applicable to low viscosity basaltic magmas
where small melt fractions can be removed fromtheir source regions. The distribution of trace elements during fractional melting is quantified interms of the following equation:
Cliq/Co = 1/Do(1 − F)(1/Do − 1)

[1.2]

where: Do is the bulk partition coefficient of theoriginal solid (prior to melting); and the othersymbols are the same as for equation [1.1].
The extent of trace element enrichment anddepletion in a fractional partial melt is shownin Figure 1.14b. For very small degrees of melting
the changes in trace element concentrations relative to the source material are extreme and varyfrom a maximum value of enrichment (1/Do as F
approaches 0) to depletions in the magma as melting progresses. Unlike the batch melt situation,compatible element enrichment can also occur
during fractional melting, but this is only likely
to happen in the unlikely event of more than 70%
partial melting. An example of incompatible element concentration in a granitic partial melt isprovided in Box 1.2.

[1.1]

where: Cliq is the concentration of a trace elementin the liquid (melt); Co is the concentration of
trace element in parental (unmelted) solid; Dresis the bulk partition coefficient of the residual
solid (after the melt is extracted); F is the weightfraction of melt produced.
A diagrammatic illustration of the extent of

1.4.2 Crystallization of magmas
The melt sequence for peridotite described in
Figure 1.12 can be used in reverse to illustrate thecrystallization of an ultramafic magma as it cools(bearing in mind that in reality a magma derived
from complete melting of a peridotitic protolith
at around 1800 °C is unlikely ever to have existed).

IGNEOUS ORE-FORMING PROCESSES CHAPTER 1

Partial melting and concentration of incompatible
elements: the Rössing uranium deposit
This can be tested very simply in terms of the partial(batch) melt equation, given as equation [1.1] in section
1.4.1 of this chapter. If the protolith contained 10 ppm
U and its bulk partition coefficient relative to the melt
fraction was significantly less than 1 (i.e. Dres between
0.01 and 0.1) then a 5% batch melt would yield a magmathat is enriched by factors of between 7 (for Dres = 0.1) and17 (for Dres = 0.01), as shown in Figure 2. This process
would have resulted in uranium concentrations in theleucogranite of between 70 and 170 ppm, which is atbest only one-half of the observed average concentration.
In order to achieve concentrations of around 300 ppm U
by batch melting either the degree of melting would haveto have been very low (i.e. <5%), or the protolith would
itself have to have been significantly enriched, and havecontained substantially more than 10 ppm U.

100

Batch melting
Cliq
CO

=

1
Dres + F (1 – Dres)

17×

Dres = 0.01

10
7×

o
Pr

The Rössing mine in Namibia is one of the largest uranium
deposits in the world, and certainly the largest associatedwith an igneous rock. This deposit is associated withleucogranite (locally termed “alaskite”). The ore occurs
mainly in the form of disseminated uraninite (UO2) distributed unevenly throughout the leucogranite, together
with secondary uranium silicate and oxide minerals suchas beta-uranophane and betafite. Rössing occurs in thecentral zone of the c.500 Ma Damaran orogen andthe leucogranites are thought to be a product of partial
melting of older basement comprising granite and supracrustal (metasedimentary) sequences. Upper amphibolite
to granulite grades of metamorphism apply regionally andmelts are considered to have been derived from depths
not significantly greater than the actual level of emplacement, as indicated by the profusion of leucogranitic dykes
in the region. The Rössing mine is located where severalof the larger leucogranitic dykes have coalesced to form asignificant mass of mineable material. The deposit comprises several hundred million tons of low grade uranium
ore running on average at about 0.031% (or 310 ppm)
U3O8. Although the deposit is low grade, a concentrationof 310 ppm uranium in the leucogranite neverthelessrepresents a significant enrichment factor relative to the
Clarke value (i.e. about 2.7 ppm for U; see Table 1.2). Thisdegree of enrichment is consistent with what might be
expected for a magma formed by a low degree of partial
melting of an already reasonably enriched protolith.

gr
es
siv
e

Cliq
CO

me
lt

ing

Dres = 0.1

5% partial melt

1.0

0.1

0

0.2

0.4

0.6

0.8

1.0

F

Figure 1 Two phases of leucogranite cutting
through high-grade metasediments in the vicinity of
the Rössing uranium mine, Namibia (photograph by
Paul Nex).

Figure 2 Simplified batch melt model showing thedegree of enrichment expected for an incompatible
element such as uranium (with Dres of either 0.1 or0.01) after 5% melting (further details provided in
section 1.4.1).

41

42

PART 1 IGNEOUS PROCESSES

Research by Nex et al. (2001) has shown that partial
melting in the central zone of the Damaran orogeny wasan episodic affair and that only one of several generationsof leucogranite is significantly enriched in uranium. Since
all are likely to have been generated by approximately thesame degrees of partial melting, this would suggest that thesingle enriched episode was derived from a particularly

fertile protolith. It is apparent that deposits representing
enrichment of incompatible trace elements in rocks formedby small degrees of melting, like the Rössing leucogranites,
are not very common in the Earth’s crust. Another example
might be the Johan Beetz uranium deposit in Quebec,
Canada.

(a) Batch melting – liquids(b) Fractional melting

100

100
Cliq
1
=
(1 – F )(1/DO – 1)
DO
CO

Cliq
1
=
CO
Dres + F (1 – Dres)

0.01
10

10

0.0001

Cliq
CO

Cliq
CO

0.1
1

1.0

0.1
0.01

1.0

D=1

2

2

5

5

D = 10
0.1

0

0.2

10
0.4

0.6

0.8

1.0

F

0.1

0

0.2

0.4

0.6

0.8

1.0

F

Figure 1.14 The enrichment/depletion of a trace element in a partial melt relative to its concentration in thesource rock (C liq/Co ) as a function of increasing degrees of melting (F). (a) Batch melting using equation [1.1], and(b) fractional melting using equation [1.2] (after Rollinson, 1993).

The crystallization sequence would have commenced with olivine, followed by orthopyroxene
and clinopyroxene. Together with plagioclase,
this mineral assemblage and the crystallization
sequence is typical for a basaltic magma.
The sequential crystallization of a mafic magmais, of course, another way of fractionating chemicalcomponents, since minerals of one composition
are often physically separated from the compositionally different magma from which they form.
Once minerals are removed from the magma (by aprocess such as crystal settling) there is little or nofurther chemical communication (or equilibration)
between the solid and liquid components of thechamber. This process is referred to as fractional
crystallization or Rayleigh fractionation. The
nature of fractional crystallization, and the geometry of cumulate rocks that form during theseprocesses, are both very relevant to an understanding of ore deposits hosted in igneous rocks. It
is, therefore, necessary to consider these processes
in more detail and, in particular, their relationshipto metal concentration and ore formation. It should
also be emphasized at this stage that emplacement and crystallization of magma is generallyaccompanied by varying degrees of assimilation ofcountry rock by the magma. Consequently, chemical trends reflecting fractionation processes are

IGNEOUS ORE-FORMING PROCESSES CHAPTER 1

accompanied by the effects of magma contamination and this, too, has important implications for
igneous ore-forming processes (see sections 1.4.3
and 1.6).
The form and internal zonation of igneous bodies
The emplacement mechanisms of magma into the
Earth’s crust and the resulting shapes of igneous
intrusions is a complex topic and is reviewed
in Hall (1996), Pitcher (1997), and Best (2003).
Basaltic intrusions are typically quite different
in shape and form from granitic batholiths. In
addition, the mechanisms of crystal fractionation
within a mafic magma chamber are also different to those applicable inside granite intrusions.
These differences are relevant to ore formationin crystallizing magmas, as exemplified by thestratiform nature of chromitite seams in mafic
intrusions such as the Bushveld Complex (seeBox 1.4), compared to tin deposits in granites ofthe same Complex which tend to be concentrated
in disseminated form toward the center of such
intrusions (see Box 1.3).
Mafic intrusions The viscosity of mafic magma
is relatively low, as illustrated by the ease withwhich basaltic magmas flow on eruption. The
densities of mafic minerals are typically greaterthan 3 g cm−3, whereas mafic magma has adensity of around 2.6 g cm−3. The low viscosity ofmafic magma and the high densities of minerals
crystallizing from it imply that minerals such as
olivine and the pyroxenes will typically sink in amagma at velocities of anywhere between 40 and1000 m yr−1 depending on their composition andsize (Hall, 1996). By contrast, less dense minerals,
such as the felspathoids, might float in an alkaline
magma as they have densities less than 2.5 g cm−3.
Plagioclase would float in a basaltic magma atpressures greater than about 5 kbar, but wouldsink in the same magma emplaced at high crustal
levels (Kushiro, 1980).
The sequence of layered rocks that result fromgravitationally induced crystal settling are referred
to as cumulates and their compositions differ
from that of the starting magma. Crystal settlingis, therefore, a form of fractional crystallizationand this process could explain the segregation of43

chemical constituents and their possible concentration into either of the solid or liquid phasesof the chamber. Trace elements that are readily
incorporated (by substitution) into cumulus minerals are referred to as compatible elements. By
contrast, trace elements which are excluded fromthe cumulate assemblage (because they cannot
easily substitute into the crystal lattice sites of the
major rock-forming minerals) are called incompatible elements and they will, naturally, become
progressively enriched in the residual magma as
the cumulate assemblage is formed.
Gravitationally induced settling of minerals
in mafic intrusions would seem to be the mostlogical way of explaining the well layered internal
structure so often observed in these bodies. Earlymodels for explaining the crystallization andlayering in mafic intrusions were derived fromthe classic work of Wager and Brown (1968) on
the Skaergaard intrusion of Greenland. The latteris particularly well suited to petrogenetic study
because it is considered to represent a single pulseof magma that crystallized to completion in aclosed chamber without substantial addition or
removal of magma during solidification. The subhorizontal layers that form in mafic intrusions
arise from the initial accumulation of early formed
minerals (such as olivine and orthopyroxene), followed, in a stratigraphic sense, by accumulations oflater formed minerals such as the pyroxenes andplagioclase. Minor oxide minerals such as chromite
and magnetite are also observed to accumulate
among the major silicate mineral phases. The
Skaergaard intrusion is made up of a chilled marginal phase and a relatively simple internal zone
made up of sub-horizontal rhythmic and cryptic(i.e. where the layering is not visually obvious but
is reflected only in chemical variations) layers,
ranging in composition upwards from gabbro to
ferrodiorite and granophyre (Figure 1.15).
Although it is tempting to attribute the relatively ordered, sub-horizontal layering that is socharacteristic of intrusions such as Skaergaard
to simple gravitationally induced crystal settling,solidification of magma chambers, including
Skaergaard, is a much more complex process(McBirney and Noyes, 1979). The temperature
gradients and associated variations in magma
density that form in cooling magmatic bodies

44

PART 1 IGNEOUS PROCESSESCountryrocks

Skaergaard
intrusion

S

N
MBS

UZ
MZ
HZ

0

1

2

LZ
Platinova"reefs"Au, Pd ore

3km
Figure 1.15 Schematic illustration showing the nature of layering in the Skaergaard intrusion, east Greenland(after Anderson et al., 1998). The relative stratigraphic position of the Au- and Pd-bearing Platinova Reefs is shown inrelation to the Marginal Border Series (MBS) and the Layered Series, which is subdivided into the Hidden Zone (HZ),
Lower Zone (LZ), Middle Zone (MZ), and Upper Zone (UZ).

(a)

(b)
2.61

Layer 2

2.59

2.58

50%
40
30
20

0

Ol
ivin

15

e

10

2.57

5

Layer 1

“Fountain”

Layer 2

“Plume”

HeightOpx(40 + pla
:60 g)

Density (g cm–3)

2.60

Opx

Layer 11400

1350 1300 1250
Temperature (°C)

1200

1150
New
magma

Figure 1.16 (a) Density variations in a fractionating magma similar in starting composition to that of the Bushveld
Complex. (b) Contrasting behavior of a new magma injected into a density stratified magma chamber within which
crystal fractionation has already occurred (after Campbell et al., 1983; Naldrett and von Gruenewaldt, 1989).

have been shown to result in pronounced density
stratification, with the formation of liquid layers
through which elements diffuse in response toboth chemical and temperature gradients (Huppert
and Sparks, 1980; Turner, 1980; Irvine et al., 1983;
McBirney, 1985). Density variations over time ina magma chamber are also a product of crystal

fractionation itself. In Figure 1.16a it is apparent
that during early stages of olivine crystallization
the residual magma density decreases because thechemical components extracted by olivine havehigher densities than that of the starting liquid.
The trend changes, however, with the appearance
on the liquidus of a mineral such as plagioclase,IGNEOUS ORE-FORMING PROCESSES CHAPTER 1

which has a density lower than that of the magma.
In such a situation it is possible, after an intervalof crystallization, for a residual magma to become
more dense than when it started solidifying.
This pattern of density variation has major
implications for the behavior of crystals settlingin a magma chamber, particularly when a new
injection of magma takes place into an alreadyevolved chamber. In Figure 1.16b two situations
are considered. If the new magma has a densitythat is greater than the liquid residue in thechamber (as might be the case if the new magma
was injected fairly early after crystallization hadcommenced), then a rather muted fountain-likefeature will form and mixing between the newand evolved liquids will be limited to a layer
along the base of the chamber. By contrast, if the
new magma is injected late in the crystallization
sequence, there is a possibility that its density
will be less than that of the residual liquid and amore impressive plume-like feature would form
(see Figure 1.16b) and the new magma would rise
to its own density level, or to the very top of thechamber. In the latter scenario turbulent mixing
between the new and residual liquids is likely
to be more complete. Injection of a plume thatreached and interacted with the roof of the magma
chamber has been suggested for the Bushveld
Complex (Schoenberg et al., 1999). This plume has
also been strongly implicated in the formationof both the chromite and PGE mineralization in
the Bushveld (Kinnaird et al. 2002).
At a first glance, layered mafic intrusions such
as Skaergaard may seem to be the products ofrelatively straightforward, gravitationally induced,
crystal settling of minerals as they appear sequentially in the growth sequence. In detail, the nature of
the layers may be quite complex and the normal,
usually rather thick, cumulate assemblages arelikely to be affected by complex processes involving the relative rates of thermal and chemical
diffusion. In addition, convective currents in themagma chamber, as well as turbidity currents inwhich a crystal-laden flow of magma surges downa slope on the margins or floor of the chamber,
add to the complexities of crystallization. Finally,
the composition of the minerals crystallizing at
any one time will be sensitive to perturbations in
the chemistry of the magma and rapid changes

45

caused either by injection of new magma or contamination of existing magma by wall rock. The
processes resulting in crystal fractionation, density
variations and changes in magma composition are
all critical to the formation of ore grade concentrations of chromite, magnetite, platinum groupelements, base metal sulfides and even gold inlayered mafic complexes. As the processes are
better understood, such intrusions are increasingly attracting the attention of explorationists
worldwide and even well studied bodies, such
as Skaergaard, are now known to be mineralized.
Although the presence of sulfide minerals at
Skaergaard was recognized early on (Wager et al.,1957), the recognition of a potentially significant
Au and Pd resource only came much later (Birdet al., 1991; Andersen et al., 1998). The interaction
between magmatic processes and the role thatthey play in mineralization are discussed in moredetail below, and also illustrated in case studies.
Felsic intrusions Granite intrusions do not exhibit
the well defined sub-horizontal layering thattypifies large mafic intrusions. This is largely due
to the fact that felsic magmas are several ordersof magnitude more viscous than mafic ones (i.e.
up to 106 Pa s for rhyolite compared to 102 Pa s
for basalt at the same temperature; Hall, 1996).
In addition, the density contrasts that exist in
crystallizing mafic magma chambers are not as
marked as they are with respect to quartz andfelspars forming in felsic chambers. Crystal settling is inconsequential in all but a few granite
plutons, such as the relatively hot, hydrous alkaline granites whose parental magmas were
emplaced at shallow depths in the crust. It isnevertheless evident that substantial fractionation
does occur in felsic intrusions and there are manyrecorded examples of internal zonation in granite
plutons (Pitcher, 1997). Most plutons record a concentric zonation, with the outer zones preserving
more mafic compositions (i.e. diorite, tonalite,
granodiorite) and rock types becoming progressively more fractionated toward the center. Sincemany granite plutons intrude at shallow crustal
levels and most of their heat is, therefore, lost to
the sidewalls, this type of zonation is logicallyattributed to crystallization that commenced
from the sides and roof of the magma chamber

46

PART 1 IGNEOUS PROCESSES

and progressed inwards. Although crystals are notbeing removed from the magma by the settlingprocess applicable in mafic intrusions, they are
effectively isolated from the residual melt by the
crystallization front which advances in towards
the center of the chamber. This process, referred
to as sidewall boundary layer differentiation, canalso, therefore, be regarded as a form of crystal
fractionation and is characterized by concentration of incompatible elements in the center ofthe intrusion where the final increments of differentiated granite melt accumulate.
Trace element distribution during fractional
crystallization
There are several ways by which crystallization
can be modeled but the most appropriate would
appear to be where crystals are removed from thesite of formation with only limited equilibration
between solid and liquid phases. In such a casetrace element distributions can be described interms of Rayleigh fractionation, in which:Cliq /Co = F(D − 1)

[1.3]

for trace element concentration in the residual
magma (liquid), andCsol /Co = DF(D − 1)

[1.4]

for trace element concentration in the crystallizing assemblage (solid), where: Co is the original
concentration of trace element in the parental
liquid; D is the bulk partition coefficient of thefractionating assemblage; F is the weight fractionof melt remaining.
The extent of trace element enrichment anddepletion during fractional crystallization of amagma is shown in Figure 1.17. The enrichmentof an incompatible trace element in the residual
melt relative to the original concentration in theparental magma is demonstrated in Figure 1.17a,where abundances can be seen to increase exponentially as crystallization proceeds (i.e. as thefraction of melt remaining, F, decreases). Compatible elements will, of course, continuously decreasein the residual melt as they are extracted into
the solid phases. Figure 1.17b shows, however,that the relative concentration of a compatible
element in the crystallizing assemblage will start
off as enriched only for rocks formed in the earlystages of fractionation. Relative concentrations
will decrease as crystallization progresses because
the magma rapidly depletes itself of the compatible elements as they are withdrawn into the
crystallizing assemblage.
Rayleigh fractionation equations apply reasonably well to low viscosity basaltic magmas where
the effects of crystal fractionation are very evident.
As mentioned previously, though, granitic magmas
do not exhibit well defined igneous stratificationand they tend to solidify by inward nucleation ofa sidewall boundary layer. This in situ style ofcrystal fractionation, where the crystallizing frontis spatially distinct from the magma residue, isakin to Rayleigh fractionation, but the equations
that govern trace element distribution have to
be modified, as shown below, to accommodate
the melt fraction that is returned to the magma
chamber as the solidification zone moves progressively into the chamber (Langmuir, 1989):
Cliq /Co = (Mliq/Mo)[f(D − 1)/(D(1 − f) + f)][1.5]

where: Mliq is the mass of liquid remaining in
the magma chamber; Mo is the initial mass of themagma chamber; f is the fraction of liquid in
the solidification zone returned to the magma
chamber; and the other symbols are as in equation [1.4].
The distribution trends of compatible andincompatible trace elements during this style ofprogressive inward solidification are essentiallysimilar to Rayleigh fractionation (Figure 1.17)
although the degrees of concentration, especiallyat low values of f, are not as extreme (Rollinson,1993). This accords with observations that theeffects of crystal fractionation in granites tend to
be less evident than in mafic intrusions.
An example of trace element concentration
during inward nucleation of a sidewall boundary
layer in granite is provided by the Zaaiplaats
Sn deposit in the granitic phase of the Bushveld
Complex in South Africa (see Box 1.3). ThereIGNEOUS ORE-FORMING PROCESSES CHAPTER 1

47(b) Rayleigh fractionation –residue, instantaneous solid(a) Rayleigh fractionation – liquids
100

100
Cliq
= F (D-1)
CO

10

Csol
CO

10

0.01

Cliq
CO

Csol
CO

0.1
1

1.0

= D.F (D-1 )

0.1
1

1.0
2

2
5

5

D = 10

D = 10
0.1

0

0.2

0.4

0.6

0.8

1.0

F

0.1

0

0.2

0.4

0.6

0.8

1.0

F

Figure 1.17 Trace element distribution during fractional crystallization. (a) The enrichment/depletion of a traceelement in the residual magma relative to its concentration in the original melt (Cliq/Co) as a function of the fractionof the remaining magma (F) using equation [1.3]. (b) The enrichment/depletion of a trace element in a crystallized
assemblage (immediately removed from its melt) relative to its concentration in the original melt (Csol/Co) as afunction of the fraction of the remaining magma (F) using equation [1.4] (after Rollinson, 1993).

Boundary layer differentiation in granites
and incompatible element concentration:
the Zaaiplaats tin deposit, Bushveld Complex
The granites of the Bushveld Complex, which overlie thebetter known layered mafic intrusion, occur as large sill-like
intrusions and are also known to be strongly fractionated(McCarthy and Hasty, 1976; Groves and McCarthy, 1978).
The granites of the Bushveld Complex are enriched in Sn
and several deposits occur in the more highly fractionatedparts of the suite. One of the better studied examples is the
Zaaiplaats tin mine, which obtains the bulk of its tonnage
from a zone of low-grade, disseminated cassiterite (SnO2)
mineralization that occurs within the central portion of
the granite (Figure 1).
A cross section through the ore body, from diamond drill
core, defines the low grade zone of disseminated cassiterite
toward the center of the granite sheet. This is consistent

with the suggestion that solidification occurred from themargins inwards and that Sn was fractionated by processes
akin to boundary layer differentiation. This type of crystal
fractionation and its implications for ore-forming processes
can be evaluated in terms of trace element distribution
patterns. If boundary layer differentiation is modeled interms of Rayleigh Fractionation (equations [1.3] and [1.4]in section 1.4.2), then the degree of crystallization required
to concentrate Sn in the residual magma to the levels
observed in the disseminated zone at Zaaiplaats canbe calculated. Assuming that Sn is incompatible in the
crystallizing granitic magma (i.e. D = 0.1) then some 96%
crystallization is required in order to achieve an enrichment
factor of 20 × (Figure 2). Coetzee and Twist (1989) havePART 1 IGNEOUS PROCESSES

S

N

Pegmatite

f

re
ophy
Gran
Zone

of Sn

Chilled marginal
graniteent
nrichm(>35

Surface

n)

ppm S

e

Granite(2–5 ppm Sn)
0

200
100
Meters300

Figure 1 Simplified cross section through the Zaaiplaats tin mine in the northern portion of the Bushveld
Complex, South Africa (after Coetzee and Twist, 1989).

shown that unmineralized granite of the type that hosts tin
mineralization in the Bushveld Complex contains between
8 and 14 ppm Sn, whereas mineralized portions of thesuite, like the disseminated zone at Zaaiplaats, averagearound 270 ppm Sn. The disseminated mineralization
at Zaaiplaats is, therefore, consistent with Sn concentration by crystal fractionation. After an advanced degreeof solidification, it is suggested that the Sn content of
the residual magma was sufficiently enriched to promotecassiterite crystallization. Other factors are also necessary
in order to stabilize cassiterite in granites (Taylor andWall, 1992), and these include fO2 and magma composition (specifically the Na/K ratio). These factors, together
with the degree of enrichment required, account for thefact that cassiterite is seldom seen as an accessory mineralin granites.

Rayleigh fractionation

100

Cliq
CO

= F (D–1)

20×
10

96% fractional crystallization

48

Cliq
CO

1.0

0.1

0

Fra
cti

ona
lc

D = 0.1

0.2

0.4

ryst
alliz
ation

0.6
F

0.8

1.0

Figure 2 (left) Rayleigh fractionation modelshowing the degree of enrichment expected for an
incompatible element (D = 0.1) such as Sn after
96% crystallization (further details provided in
section 1.4.2).

IGNEOUS ORE-FORMING PROCESSES CHAPTER 1

are many examples of trace element concentrations in layered mafic intrusions and subsequent
sections discuss some of the complex processes
involved in accumulation of ore grade Pt, Cu,
and Ni deposits such as those of the Bushveld
Complex, Kambalda, and Sudbury (see Boxes 1.4,
1.5, 1.6, and 1.7). A case where crystal fractionation alone, without the added complication offeatures such as magma replenishment or contamination, has resulted in significant enrichment
of Au, Pd, and S is provided by the Platinova Reefs
of the Skaergaard Complex. Figure 1.15 shows that
this zone of potentially ore grade precious element
enrichment occurs toward the top of the Middle
Zone and formed after a substantial proportion of
the magma chamber had crystallized. The incompatible nature of Au, Pd, and S with respect tothe early formed crystals (dominantly plagioclase,
olivine, and pyroxene with lesser magnetite andapatite) resulted in a progressive enrichment ofthese metals in the residual magma (Anderson
et al., 1998). Once the concentration of sulfur had
reached saturation levels (believed to be between
0.16 and 0.3 wt% in the Skaergaard magma)
the metals precipitated out as a Au–Pd alloy to
be further concentrated as inclusions within
the sulfide minerals that now occur distributed
interstitially among the normal cumulus mineral
assemblage. The concentrations of Au and Pd
that resulted from the progressive in situ crystal
fractionation of the Skaergaard magma amounts
to about 2 ppm for each metal and it has been
estimated that some 90 million tons of potential
ore grade material is available in the Platinova
Reefs (Anderson et al., 1998).
1.4.3 Fractional crystallization and theformation of monomineralic chromitite layers
Crystal settling, convective fluid flow, anddiffusion-related chemical segregation acrossdensity stratified layers are the processes whichgive rise to the characteristically sub-horizontal,
well ordered layering evident in most layered
mafic intrusions. These processes do not, however, account for the occasional developmentof monomineralic layers of chromite or magnetite so spectacularly developed in many layered49

mafic intrusions around the world. Nor do they
explain the formation of the massive pods ofchromitite that occur in the lithospheric portionsof ophiolite complexes. This section provides
some insight into the processes by which layers
and pods of massive chromitite can form in mafic
cumulate rocks.
There are many mafic intrusions that contain layers of near monomineralic chromite or(vanadium-rich) magnetite that are typically 0.5–
1 meter in thickness and extend laterally for tensof kilometers, representing enormous reserves of
Cr and Fe–V ore (see Box 1.4). Notable among
these are the Bushveld Complex in South Africa
and the Great Dyke in Zimbabwe. The formation of such layers, which might comprise up to
90% of a single mineral (chromite or magnetite),would appear to require that normal crystallization of silicate minerals (dominated by olivine,
the pyroxenes, and plagioclase) be “switched off”
and replaced by a brief interlude where only thesingle oxide phase is on the liquidus. A simple, but
elegant, explanation of this process was provided
by Irvine (1977) with respect to the formation ofchromitite seams.
The Irvine model
The Irvine model refers to part of the phase
diagram for a basaltic system in which only theolivine–chromite–silica end-members are portrayed in a ternary plot (Figure 1.18a). The normal
crystallization sequence in a basaltic magma withstarting composition at A (Figure 1.18b) wouldcommence with olivine as the only mineral on
the liquidus, settling of which would result in theformation of a dunitic cumulate rock. Extractionof olivine from the magma composition at A would
result in evolution of the magma composition
away from the olivine end-member composition
and toward the cotectic phase boundary at B. At
B a small amount of chromite (around 1%) would
also start to crystallize together with olivine,
and the magma composition would then evolve
along the cotectic towards C. At C the SiO2 content of the magma has increased to a level where
olivine and chromite can no longer be the stableliquidus assemblage and orthopyroxene starts to50

PART 1 IGNEOUS PROCESSES

(a)

(b)
40

D

Orthopyroxene

Wt
%

SiO

2

QuartzArea coveredby b, c, dOlivine30

C
Chromite

20
B

A

10
Olivine

Chromite

0.5

1.0
Wt% Chromite

1.5

2.0

(c)

SiO

2

Olivine

(d)

40

50

D

40
D

SiO

20
E
A
1.0
Wt% Chromite

H

20
E

10

0.5

G

B

10

Olivine

C

30

%

F

Wt

Wt

2

30

%

SiO

2

Mixing line between
magmas at D and E

GCo
traj ntam
ect inatory ion50

1.5

Olivine

B

A

0.5

1.0
Wt% Chromite

1.5

2.0

Figure 1.18 A portion of the ternary system quartz–olivine–chromite (a) showing the nature of crystallization in a
mafic magma (b). Scenarios in which magma mixing (c) and magma contamination (d) occur as mechanisms forpromoting the transient crystallization of only chromite are also shown (after Irvine, 1977).

crystallize to form a bronzitite cumulate rock. From
this stage magma composition evolves toward D.
Continued fractional crystallization will eventually lead to the appearance of plagioclase together
with orthopyroxene on the liquidus, but felspar
compositions are not reflected on the simplified
phase projection shown here. This crystallization sequence cannot lead to the formation of achromite seam and the latter mineral would only
occur as an accessory phase in the early formed
cumulates.

In order to make an ore deposit something
unusual, or different from the norm, needs to takeplace. One way of disturbing the normal crystallization sequence is to introduce, at point D, a
new magma with a composition at E (i.e. not as
primitive as the original starting liquid), that is
injected into the chamber and allowed to mix withthe evolved liquid at F (Figure 1.18c). Minglingof the two liquids represented by D and E would
result in a mixture whose composition must lie
somewhere along the mixing line DE. Exactly

IGNEOUS ORE-FORMING PROCESSES CHAPTER 1

where along DE this mixture would lie depends
on the relative proportions of D and E that are
mixed together. For most mixtures (at point F,
for example, in Figure 1.18c) the magma composition would lie within the stability field ofchromite and for a brief interval of crystallization(from F to G) only chromite would crystallize

from the mixture. Being relatively dense, chromite
would settle fairly efficiently and a single, nearmonomineralic layer of chromite would form.
In a very large magma chamber, such as the one
from which the Bushveld Complex must havedeveloped, the crystallization of such a layer
could lead to the formation of an ore body with

Crystal fractionation and formation of monomineralicchromitite layers: the UG1 chromitite seam,
Bushveld Complex
The Bushveld Complex is the world’s largest layered mafic
intrusion (Figure 1), covering an area of some 67 000 km2.
It also contains a substantial proportion (more than 75%)of the world’s chromite reserves. Several major chromitite
seams (at least 14 in number) occur within the Critical
Zone of the Bushveld Complex and these are subdivided
into three groups termed the Lower Group (LG1 to LG7),
the Middle Group (MG1 to MG4), and the Upper Group(UG1 to UG3). The LG6 chromitite seam is the mostimportant in terms of production and reserves and can betraced for over 160 km in both the western and eastern
portions of the complex.
In section 1.4.3 of this chapter the mechanisms bywhich monomineralic layers of chromitite could form ina layered mafic intrusion were discussed with specificreference to the Irvine model. The UG1 chromitite layer is

N
50
km

Zaaiplaats
Sn mineBushveld Complex

Granite andrhyolite
Upper zone

Dwars
River
Magnetitelayers3000 m0

Main zone
Rustenburg
Pretoria
Johannesburg

51

Merensky
Reef
Chromitite
layersUG
MG
LG

Critical zone
Lower zone

Figure 1 Simplified geological map and section through the Bushveld Complex. The distribution ofchromitite layers or seams in the Critical Zone is shown in relation to the PGE-rich Merensky Reef and thevanadium-rich magnetite seams at the base of the Upper Zone. Also shown is the distribution of associatedgranite and rhyolite that overlies the layered mafic intrusion.

52

PART 1 IGNEOUS PROCESSES

of particular interest because it is hosted essentially withinanorthositic rocks (rather than olivine and orthopyroxene
cumulates), and is also spectacularly well exposed ina gorge of the Dwars River in the eastern portion of
the Bushveld Complex. The UG1 shows some intriguingfeatures that require special explanation. One of the interesting aspects of the UG1 chromitite layer is the way inwhich the seam bifurcates. In some sections the seam splits
into several thinner layers, whose cumulative thickness issimilar to that of the single seam. Chromitite splits tend tooccur in areas where anorthosite layers thicken into domalfeatures, with the bifurcations opening out towards thecore of the dome. Nex (2002) has explained this by suggesting that normal “sedimentation” of chromite (according to
Irvine’s model) was interrupted by liquefaction of the
footwall crystalline assemblage to form a slurry of plagioclase feldspar and melt that erupted at the magma–cumulate interface. The bottom-up accumulation of the slurry is
responsible for the formation of the domal features in
the anorthosite and also serves to dilute the top-down settling of chromite grains. There is, therefore, a correlation
between liquefaction of the footwall cumulates, doming

potentially vast Cr reserves. Experimental confirmation of these processes has been provided
by Murck and Campbell (1986).
Once the magma composition reaches point G(after extraction of only chromite) on the cotectic,
crystallization will again be dominated by olivine
and the rocks that form in the hanging wall of thechromite seam will again contain only accessory
amounts of chromite. It should also be mentionedthat another way of forcing the magma composition into the chromite field is shown in
Figure 1.18d, where the magma at point E (or anywhere along the cotectic for that matter) becomes
contaminated with siliceous material (perhaps
by assimilation of crustal material forming either
the floor or the roof of the magma chamber). The
contaminated magma would have a composition
that lies somewhere along the mixing line joining E to the SiO2 apex of the ternary diagram.
This composition would also lie transiently inthe chromite field and result in the formationof a monomineralic cumulate layer of chromite(between H and G). This is an indication of how
important contamination of magma can be toigneous ore-forming processes.

Figure 2 Bifurcating chromitite seams exposed atDwars River in the eastern Bushveld Complex.
of the anorthosites, and splitting of chromitite layers. This
type of feature, together with many other intriguing
textures and relationships, indicates that crystallization
processes in magma chambers can be very complex.

Other mechanisms for the formation ofchromitite layers or pods
Although the Irvine model very neatly explains
many of the characteristics of chromitite layers
it is unlikely to apply to all situations and there
are several other mechanisms that might pertain
to the accumulation of monomineralic layers or
pods. Two of the most likely, since they havebeen confirmed experimentally, include changesin oxygen fugacity (fO2) and total pressure (PTot)of the crystallizing magma.
An increase in fO2 will promote the stabilityof chromite and possibly allow the mineral to
crystallize alone for a period of time (Ulmer, 1969).
Increasing fO2 in the magma could be achievedby a devolatilization reaction such as [1.6] below(after Lipin, 1993).
4FeCO3 ↔ 2Fe2O3 + 4CO + O2

[1.6]

where CO2−
3 in solution breaks down to form
carbon monoxide and free oxygen. However,
because CO2 and not CO is likely to be the dominant carbon species in basalt, it seems unlikely

IGNEOUS ORE-FORMING PROCESSES CHAPTER 1

that oxidation of a magma will be easily achieved
and the process is generally not called upon as an
explanation for chromite accumulation.
By contrast, small increases in PTot of the magma
have now been shown to occur readily in basaltic
chambers. Observations in the Kilauea volcano(Hawaii) have shown that a pressure increase(of up to 0.25 kbar) can occur in the roof of themagma chamber as a result of CO2 exsolution
and expansion of the gas bubbles as they stream
upwards in the magma (Bottinga and Javoy, 1990;
Lipin, 1993). An increase in total pressure in themagma chamber will have the effect of shifting
the phase boundary between olivine and chromite
such that the field of the latter phase expands.
This would have the same effect as that predicted in the Irvine model, in that chromite would
crystallize alone until such time as the ambientpressure is restored (by egress of magma, volatiles
or both). This mechanism has been proposed forthe formation of chromitite layers in the StillwaterComplex (Montana), where it is suggested that
CO2 streaming and associated pressure increaseaccompanied ingress of a new magma pulse into
the chamber (Lipin, 1993). Exsolution of volatiles
and fluids from a magma is discussed in moredetail in Chapter 2.
A note on podiform chromitites
Although not as large as the deposits in layered
mafic intrusions, podiform chromitite ores inophiolite complexes represent important resourcesin many parts of the world. The chromitite ores
are typically found as irregular, stratiform to
discordant, pods within dunitic and harzburgitic
host rocks which themselves are often intensely
deformed. In detail the ore textures, characterized
by nodular and orbicular associations of chromite
and olivine, suggest that the mingling of two
magmas has given rise to crystallization of thechromitite ores. Ballhaus (1998) has suggested
that the sites of chromite mineralization represent zones in the oceanic lithosphere where low
viscosity, olivine-normative melt mingled with amore siliceous, higher viscosity magma. The two
melt fractions remain segregated, at least for thetime it takes to accumulate chromitite ore. For

53

thermodynamic reasons chromite nucleates preferentially in the ultramafic melt globules, withcrystals forming initially along the metastableliquid interface (where mixing takes place at asmall scale and a situation akin to the Irvine model(Figure 1.18) pertains) and then progressively
throughout the globule. Diffusion of chromiumfrom the siliceous magma, where no chromite
nucleation has occurred, across the liquid interface into the ultramafic melt globules, also takes
place, with the result that the latter might ultimately be entirely replaced by chromitite. Because
the siliceous magma acts as a chromium reservoir,
the richest ores are considered to occur where thevolume of the latter is high relative to that of theultramafic globules within which accumulationof chromitite is occurring.
1.4.4 Filter pressing as a process of crystal
fractionation
The separation of crystal phases from residual
melt during the solidification of magma is generally attributed to gravitational segregation where
crystals of either higher or lower density thanthe magma settle or float to form horizontallylayered cumulate rocks. As mentioned previously, this is an oversimplification of the actual
processes, which often involve complex density
and chemical diffusive controls. Another mechanism by which crystal melt segregation canoccur is the process known as filter pressing. The
residual magma within a network of accumulating crystals in a partially solidified chamber canbe pressed out into regions of lower pressure suchas overlying non-crystalline magma or fracturesin the country rock. The process is considered
to apply even in more viscous granitic magma
chambers where evolved, water-saturated melts
are filter pressed into adjacent fractures createdduring hydrofracturing.
Anorthosite hosted Ti–Fe deposits
Large massif-anorthosite intrusions of Mesoproterozoic age, located in the Paleohelikian andGrenvillian orogenies extending from North
America into the Sveconorwegian province, are54PART 1 IGNEOUS PROCESSES

the hosts to very important Ti and Fe deposits ofmagmatic origin (Force, 1991a; Gross et al., 1997).
Well known examples of such deposits include
Sanford Lake in the Adirondacks of New Yorkstate, USA, the Lac Tio deposit in the Allard Lakeregion of Quebec, Canada, and the Tellnes deposit
in southern Norway. The deposits are typicallythought to be related to large differentiated intrusive complexes made up mainly of anorthosite,
gabbro, norite, and monzonite rocks emplacedin the late tectonic to extensional stages of theorogenic cycle. Although mineralogically variable, the more important category economicallyis the andesine anorthosite type (or Adirondacktype), which contains ilmenite–hematite as theprincipal ore minerals. The Ti–Fe oxide ore
accumulations occur as stratiform layers anddisseminations within the intrusive complexes
themselves, or as more massive, higher-grade,cross-cutting or dyke-like bodies.
These deposits are clearly a product of in situ
crystal fractionation. Early extraction of a plagioclase-dominated crystal assemblage results in concentration of Fe and Ti in the residual magmas,
which crystallize to form ferrogabbro or ferrodiorite. Titaniferous magnetite or hemo-ilmenite(depending on the magma composition) also
crystallize with disseminated layers formed by
crystal settling and accumulation on the chamberfloor. The more massive discordant bodies are
considered to be a product of the pressing out of an
Fe–Ti oxide mineral slurry – the slurry concentrates to form an intrusive body often along themargins of the largely consolidated anorthosite
complex, or into fractures and breccia in the hostrocks. The Lac Tio orebody is an irregular, tabular
intrusive mass some 1100 m long and 1000 mwide, whereas the Tellnes ores form part of a
14 km long dyke (Gross et al., 1997).

1.5 LIQUID IMMISCIBILITY AS AN ORE-FORMING PROCESS
Liquid immiscibility is the segregation of two
coexisting liquid fractions from an originallyhomogeneous magma. The two fractions may bemineralogically similar (silicate–silicate immiscibility) or be very different (silicate–oxide, silicate–carbonate or silicate–sulfide immiscibility). The
phenomenon of immiscibility is best observed in
extrusive rocks where rapid quenching prevents
the segregated products from being rehomogenized. Philpotts (1982) noted two compositionallydistinct glasses interstitial to cumulus minerals
in a tholeiitic basalt. Their compositions were
essentially granitic, on the one hand, and an
unusual mafic assemblage, comprising pyroxene,
magnetite–ilmenite, and apatite, on the other.
The August 1963 eruption of the Kilauea volcano
in Hawaii provided evidence of another form ofliquid immiscibility in which a directly observable sulfide melt separated from a cooling, basaltic
magma (Skinner and Peck, 1969). The Duluth
Complex in Minnesota contains several small
occurrences of massive Cu–Ni sulfide mineralization, as well as rare discordant bodies of ilmenite–magnetite–apatite (in oxide to apatite proportionsof about 2:1) referred to as nelsonites. These occurrences are considered to provide evidence thatboth sulfide and Fe–Ti–P immiscible fractions
separated from the Duluth magma and that these
represent a compositional continuum of immiscible products (Ripley et al., 1998). Although its
occurrence is difficult to prove, and despite thefact that it is not a major process during magma
evolution, immiscibility is very important as anore-forming process in mafic magmas and canlead to the formation of large and important
deposits such as the PGE sulfide deposits of theMerensky Reef in the Bushveld Complex, South
Africa (Box 1.6), the Ni–Cu sulfide ores at
Kambalda in Western Australia (Box 1.5), and at
Sudbury in Ontario (Box 1.7). Various types ofimmiscibility are discussed below, with special
emphasis on silicate–sulfide immiscibility.
1.5.1 Silicate–oxide immiscibility
It is well known that unusual, discordant bodies
of magnetite–apatite or ilmenite/rutile–apatite(nelsonite) are preferentially associated withsome alkaline rocks, as well as with anorthosite
complexes. Early experimental work showed that
it is possible to create two immiscible liquids, one
quenching to form a mixture of magnetite andapatite in proportions of about 2:1, and the other aIGNEOUS ORE-FORMING PROCESSES CHAPTER 1(a)fO2 = 10–1 bars

FeO

Immiscibility
gap

KAISi3O8(b)–9

fO2 = 10

KAISi3O8

SiO2
bars

FeO

SiO2

Figure 1.19 The nature of the “immiscibility gap”
(shaded) that exists in felsic magmas under conditionsof variable oxygen fugacity. The tie-lines indicate thecoexisting liquid compositions (after Naslund, 1976).

rock that is dioritic in composition (Philpotts,1967). Further experimental studies showed thatfor a fairly broad range of rock compositions underconditions of high fO2, an immiscible FeO melt
will separate from a magma of felsic composition(Naslund, 1976). A large immiscibiliity field exists,
for example, in the system KAlSi3O8-SiO2-FeO(Figure 1.19) at approximately atmospheric fO2
conditions, but the gap is greatly diminished in amore reducing environment. The existence ofimmiscibility is also enhanced in magmas withhigh concentrations of P, Ti, and Fe, but the fielddiminishes with increasing Ca and Mg (Naslund,1983).

55

The observation that oxidized iron-rich magmas can segregate into two immiscible liquids,
one which is Fe-rich and the other a more normal
silicate composition, has relevance to the formation of Fe- and Ti-rich magmatic segregations innature. In slowly cooled plutonic rocks resorption
reactions are largely responsible for the rehomogenization of segregated liquids, such that immiscibility is not often seen. In extrusive magmas,
however, quenching can occur, or a more dense
oxide liquid could separate from its silicate counterpart to be forcibly injected into a different partof the magma chamber, or into fractured countryrock. Such an immiscible Fe- or Ti-rich fluid
could also form a discrete magma or lava flow, apossibility that provides a theoretical basis forunderstanding the very important magnetite–
hematite–apatite ores in, for example, the Kiruna
district of northern Sweden (Freitsch, 1978).
Another example of what appears to be a flow ofimmiscible magnetite lava that separated from
andesitic lavas has been documented at the El
Laco volcano in northern Chile (Naslund et al.,2002). A word of caution though – in the cases ofboth Kiruna and El Laco there is considerable
debate about the origin of the iron ores, and in thelatter example in particular, a strong case has
been made suggesting a hydrothermal origin for
the Fe mineralization (Rhodes and Oreskes, 1999;
Sillitoe and Burrows, 2002). The question of silicate–oxide immiscibility as a viable ore-forming
process is still very contentious, despite the factthat the process has been proven experimentally.
1.5.2 Silicate–sulfide immiscibility
By contrast with the process of silicate–oxide immiscibility, where there is some controversy overits extent in nature, the existence of silicate–sulfide immiscibility in mafic magmas is widely
accepted as a common feature of magma crystallization. Experimental data in the system SiO2FeO-FeS (MacLean, 1969) confirms that silicate
liquid can coexist with sulfide liquid over a large
volume of the system (Figure 1.20; two-liquidfield). Magma at A in Figure 1.20, crystallizing
Fe-rich olivine (fayalite), would evolve along A–A′and eventually intersect the two-liquid phase

56PART 1 IGNEOUS PROCESSES

SiO2

2-liquids
2-liquids

3-liquids

Cr
Tr

Fe2SiO4(fayalite)AA'
2-liquids

Fay

alit

e

WustiteFeO

boundary where the residual melt would comprise conjugate silicate and sulfide melts.
Sulfur as sulfide is dissolved in magmas bydisplacing oxygen bonded to ferrous iron. Sulfide
solubility is, therefore, a function of FeO activity
in the magma, but is also controlled by oxygen
fugacity (fO2), decreasing as fO2 increases (MacLean,
1969). Sulfide solubility, or the amount of sulfide
dissolved in the magma at saturation, will varyas a magma progressively crystallizes and, at any
point when saturation is reached, small immiscible globules of sulfide melt will form. Sulfide saturation can be achieved as solidification proceeds
and magma temperature falls, or by an increase
in fO2, or by a decrease in the amount of ferrous
iron in the magma (such as might occur during
extraction of an Fe-rich phase; see Figure 1.20). As
will be shown later, other factors, such as additionof externally derived sulfur, or ingress of new
magma, can also promote saturation and the formation of an immiscible sulfide phase.
The immiscible sulfide melt that segregatedfrom basaltic magma on the Kilauea volcano contained approximately 61% Fe, 31% S, 4% Cu, andTroilite

FeS

Figure 1.20 Phase equilibria
established experimentally inthe system SiO2–FeO–FeS (after
MacLean, 1969). The field of two
coexisting silicate and sulfide
liquids is shown by the stipple
accentuated line. Oxidation of
the magma would shift the phaseboundary in the direction shownby the double arrow, and expand
the field of two-liquids. A magma
represented by composition A andcrystallizing fayalite would evolve
along the line A–A′. Silica phases
are represented by tridymite (Tr)and cristobalite (Cr).

4% O. It subsequently solidified to form mineralssuch as pyrrhotite, chalcopyrite, and magnetite(Skinner and Peck, 1969). Segregated sulfide melts
forming in this fashion clearly have enormous
potential to host concentrations of metals withboth chalcophile and siderophile tendencies, suchas base (Cu, Ni, Co) and precious (Au, Pt) metal
ores. There are many large and important ore
deposits associated with the development of an
immiscible sulfide fraction in mafic and ultramafic magmas (see Boxes 1.5, 1.6, and 1.7). Central
to the formation of all these deposits are threefundamental steps:• the appearance of a substantial fraction of
immiscible sulfide melt;• creation of conditions whereby the sulfide
globules can effectively equilibrate with a large
volume of silicate magma;• effective accumulation of the sulfide globules
into a single cohesive layer or spatial entity.
The processes whereby ore deposits are createdduring silicate–sulfide immiscibility are, therefore, complex and multifaceted, and are discussed
in more detail in the following section.

IGNEOUS ORE-FORMING PROCESSES CHAPTER 1

57

0.4

A

0.1
Figure 1.21 Variation in sulfide
solubility as a function ofprogressive crystallization in a
mafic magma such as that fromwhich the Bushveld Complex
formed (after Naldrett and Von
Grünewaldt, 1989).

Sulfidesolubility
curveC(1180°C)D

00

1.6 A MORE DETAILED CONSIDERATION OF
MINERALIZATION PROCESSES IN MAFIC MAGMAS

1.6.1 A closer look at sulfide solubility
The question of whether an immiscible fraction
will develop in a magma or not is related to theamount of sulfur required to achieve sulfide saturation. This concept is referred to simply as sulfide solubility. As mentioned previously, sulfide
solubility decreases with increasing oxygen content in a magma because solution of sulfur appears
to be controlled by the following equilibrium
reaction (after Naldrett, 1989a):
FeO (melt) + 1/2S2 ⇔ FeS (melt) + 1/2O2

Orthopyroxene ACAD
0.2

Olivine

Wt% sulfide

0.3

Plagioclase plus orthopyroxene(+clinopyroxene)

B (1300°C)[1.7]

Sulfide solubility increases as a function ofincreasing temperature and FeO content of themagma, but decreases with increasing pressureand SiO2 content. The data for sulfide solubility(reviewed in Naldrett, 1989a) have been used

10

20

30
40
50
Crystallization (%)

60

70

to construct a generalized solubility curve for afractionating Bushveld type mafic magma andare shown in Figure 1.21. The sulfur content at
sulfide saturation is seen to decrease from amaximum of about 0.4 wt% sulfide at the start ofcrystallization to below 0.1 wt% as olivine andorthopyroxene crystallize, but then levels off after
plagioclase forms.
A magma with an original sulfide content of
0.3 wt% would initially be undersaturated and its
position would plot below the curve in Figure 1.21.
Extraction of olivine and pyroxene as cumulus
minerals would cause the sulfide content toincrease in the residual magma until the saturation limit was attained (i.e. after about 10% crystallization). At this point immiscible globules ofhomogeneous sulfide melt would form and themagma would remain saturated with respect tosulfide. The amount of sulfide remaining in themagma will, however, decrease as dictated by thesolubility curve. If the high density sulfide globules
are extracted from the magma by gravitational

58PART 1 IGNEOUS PROCESSES

settling, then the distribution of trace elements between sulfide and silicate liquids can be modeled
in terms of crystal fractionation mechanisms.
After solidification the cumulate rocks that hadformed during this stage of the crystallization
sequence would contain minor disseminationsof sulfide minerals (such as pyrrhotite, FeS, andchalcopyrite, CuFeS2) among the olivine andorthopyroxene crystals. Alternatively, since thesulfide globules have a high density, they mightalso accumulate as a separate sulfide layer toward
the base of the magma chamber. The decrease insulfide solubility with progressive crystallizationof a mafic magma, as well as the shape of the solubility curve shown in Figure 1.21, have important
implications for the understanding of ore-forming
processes in these rocks and this diagram isreferred to again below.
1.6.2 Sulfide–silicate partition coefficients
The ability of an immiscible sulfide fraction toconcentrate base and precious metals will obviously depend on the extent to which metals partition themselves between sulfide and silicate melts(i.e. the magnitudes of the relevant sulfide-silicate
partition coefficients). All metals that have chalcophile tendencies are likely to partition strongly
into the sulfide phase rather than remain in thesilicate melt, and the data presented in Table 1.3
confirms that both base and precious metals
are markedly compatible with respect to sulfide
minerals. The partition coefficients indicate that
although Cu, Ni and Co partition strongly into
the sulfide phase, scavenging of platinum groupelements (PGE)* by sulfide melt is even more
efficient. Values of partition coefficients up to
100 000 for the PGE (Table 1.3) indicate that thepresence of an immiscible sulfide fraction is potentially a very powerful concentrating mechanism.
In reality, however, even though the sulfide–silicate partition coefficients for the PGE are so

Table 1.3 Estimates of sulfide–silicate partition
coefficients for base and precious metals in ultramafic
and mafic magmas

Komatiite
27% MgO
19% MgO
Basalt

Ni

Cu

Co

Pt

Pd

100
175
275

250–3000
250–2000
250–20004060
80

104–105
104–105
104–105

104–105
104–105
104–105

Source: from Naldrett (1989a), Barnes and Francis (1995),
Tredoux et al. (1995).

high, one seldom finds economically viable concentrations of these elements in layered mafic
complexes. One reason for this is that the original
concentrations of PGE in magmatic reservoirs arevery low. Another reason is that, even in magmas
where immiscible globules of sulfide do form,
the concentration mechanism may be diminishedbecause sulfides are not able to communicate(chemically) with the entire magma reservoir that
contains the metals. In the case of the Skaergaard
intrusion, for example, sulfide disseminations
occur mainly in the upper portions of the layered
body and sulfide immiscibility is believed to haveoccurred relatively late in the crystallization
sequence (Anderson et al., 1998). The sulfide
minerals at Skaergaard are depleted in Ni becausea significant proportion of the latter metal was
extracted by early formed olivine and orthopyroxene and was no longer available to be scavenged
by the sulfide phase. On the other hand, theSkaergaard sulfide zone is cupriferous, and is alsoenriched in Au and Pd, as both these metals wereconcentrated into the residual magma as incompatible elements during crystal fractionation. The
Skaergaard sulfides have only, therefore, equilibrated with the more differentiated parts of themagma chamber.

* Six elements make up the platinum group – they are platinum (Pt), palladium (Pd), and rhodium (Rh), collectively referred
to as the Pd subgroup, and osmium (Os), iridium (Ir), and ruthenium (Ru), referred to as the iridium subgroup. Both the Pd andIr subgroups have strong affinities with Fe–Cu–Ni sulfides, whereas the iridium subgroup is also often associated withchromite and olivine cumulates.

IGNEOUS ORE-FORMING PROCESSES CHAPTER 1

The effects of early removal of trace elements
are demonstrated in Figure 1.17, where it is seenthat partitioning of a highly compatible trace element (e.g. D = 10 in Figure 1.17b) leads to enrichments in the solid cumulates only for the first
20–30% of crystallization (i.e. for values of F from
1.0 to about 0.7) and, thereafter, the solids are
increasingly depleted in that element. This isbecause a high-D trace element will comprehensively enter the early formed solids, leaving theresidual magma strongly depleted. Later solids
will have to equilibrate with a depleted reservoir
and will likewise be depleted, irrespective of themagnitude of the partition coefficient. An immiscible sulfide melt within a silicate magma canalso be regarded as a cumulus mineral phase andthe same situation would, therefore, apply. In
fact, the problem would be particularly acute for
the PGE with respect to an immiscible sulfide
fraction because their partition coefficients are sohigh, and the original abundances in the parental
magma so low (generally only a few ppb), that all
but the very first formed sulfide fraction will
be depleted in these metals. One does, therefore,
have to ask the very pertinent question as to howlarge and economically viable concentrationsof low abundance metals actually form. Or, put
another way, how do the enormous concentrations of PGE-bearing sulfide ores such as those of
the Bushveld Complex form when early depletionof compatible metals apparently limits the extent
to which viable concentrations of metals can beachieved?
1.6.3 The R factor and concentration of low
abundance trace elements
If a globule of sulfide melt equilibrates with an
infinite reservoir of magma then the concentration of a trace metal in that, or any other similar,
globule is given simply by the product of the partition coefficient and the initial concentration of
the metal in the coexisting silicate melt. In realityof course the sulfide globule is only likely to interact with a small, restricted mass of silicate melt,
which is why the problem alluded to at the endof the previous section is very relevant. Clearly,
the greater the proportion of silicate relative to59

sulfide melt, the higher and more persistent theconcentration of compatible trace metals in that
globule is likely to be. Campbell and Naldrett(1979) quantified the problem by proposingthat the concentrations of low-abundance traceelements, such as the PGE, into an immiscible
sulfide phase should be calculated in terms not
only of the original abundance in the magma
and the relevant partition coefficients, but also of
the silicate/sulfide liquid mass ratio, termed the“R” factor. An expression which incorporates
the R factor into the normal distribution law isprovided in equation [1.8] (after Campbell andNaldrett, 1979):
Csul = CoD(R + 1)/(R + D)

[1.8]

where: Csul is the concentration of a trace elementin the sulfide fraction; Co is the original traceelement concentration in the host magma; D isthe sulfide–silicate partition coefficient; and R is
the mass ratio of silicate magma to sulfide melt.
Although the R factor is simply a mass fraction,
it is a parameter that also effectively records theextent to which the immiscible sulfide fraction
interacts with the silicate magma from which itis derived (and which also contains the metals
that need to be concentrated if an ore deposit is tobe formed). If a sulfide globule sinks through alengthy column of magma (or is caught up in aturbulent plume) it is effectively interacting witha large volume of silicate liquid, and this can beequated to a high R factor (Campbell et al., 1983;
Barnes and Maier, 2002). Even though a sulfide
globule may rapidly deplete the surrounding
magma at any one instant of time by virtue of avery high sulfide–silicate partition coefficient, it
may be provided with the opportunity to scavenge
compatible elements by continuously moving
through undepleted magma. A low R factor isanalogous to a situation where a sulfide globule
is static, or is removed early on from the magma,
so that it is not able to efficiently scavenge
compatible elements even though the partition
coefficients may be very high.
The way in which the concentrations of compatible trace metals in sulfide melts vary as afunction of both partition coefficient (D) and R is

PART 1 IGNEOUS PROCESSES

(a)

(b)
105

Ni (R = 106)8

D = 100 000

500

6Pt (R = 10 )Ni (wt%)

400
300

Ni
ppm
350
Xi =

pb

2

101

275

=5p

D = 100

4

Pt

102

5

D = 1000Ni (R = 103)
D=

D = 10 000

X
10 i

103

6

D=

Csul/Co

104

200

Pt (ppm)

60

100

Pt (R = 103)
104

103
R factor

102

101

6

5

4
3
Log R

2

1

Figure 1.22 (a) Plot showing the relationships between partition coefficients (D), the R factor (i.e. the mass ratio of
silicate melt to sulfide melt), and the degree of enrichment (Csul/Co) of compatible trace elements in the sulfide phase(after Barnes and Francis, 1995). (b) Diagram showing the effect of variations in the R factor on the concentration of
Ni and Pt in an immiscible sulfide fraction that is in equilibrium with a basaltic magma (after Naldrett, 1989b).

shown in Figure 1.22a (after Barnes and Francis,
1995). For cases where D is much larger then R,
the enrichment factor (Csul/Co) approximates thevalue of the R factor (i.e. the sloped portions of
the curves in Figure 1.22a). Conversely, when the
R factor is much larger than D, then the enrichment factor is approximately the value of D
(Figure 1.22a – the flat parts of the curves).
Another way of emphasizing the importance of
the R factor is to examine its influence on theconcentrations in sulfide melt of two compatible
trace elements, one with a moderately high partition coefficient (for example, Ni with D = 275) andthe other with a very high partition coefficient (Ptwith D = 100 000). Figure 1.22b shows that where
the R factor is low (say 103) Ni concentrations
in sulfide fractions are high (i.e. typical of mostNi-sulfide ores) because of the combination ofhigh D and high initial Ni content (350 ppm) of
the parental magma. By contrast, the Pt concentrations in the same sulfide fraction will be low
because, despite the very large D, they have not
had the opportunity of scavenging a substantial
mass of the element from a magma that initially
had a very small Pt abundance (5 ppb). By contrast, in situations where the R factor is high (say106) the Ni concentrations of the sulfide fraction
will not be significantly higher than the lower R

factor case, but the Pt contents will have increased
substantially because of more extensive interaction between the immiscible sulfide fraction andthe silicate magma.
Variations in the R factor have real implications for the grades of sulfide phases within individual mafic intrusions. In the Muskox intrusionof northern Canada, for example, disseminated
Cu–Ni sulfides on the margins of the body contain only moderate concentrations of PGE, whereas
the more voluminous interior of the intrusion
contains sulfides with higher PGE grades (Barnes
and Francis, 1995). This is attributed to a lower
mass ratio of silicate to sulfide melt, and therefore
a lower R factor, on the margins of the intrusion.
Considerations such as this might be of considerable relevance to an exploration strategy for thistype of deposit.
1.6.4 Factors that promote sulfide immiscibility
It is conceivable that situations in which sulfide
saturation is attained relatively late in the crystallization history of a magma could be disadvantageous to the formation of, for example, Ni- or
PGE-sulfide ores if the latter metals had alreadybeen extracted from the residual magma. It is also
possible that late sulfide saturation could equate

IGNEOUS ORE-FORMING PROCESSES CHAPTER 1

with a low R factor simply because the volumeof residual silicate magma is low. A scenario
whereby magmas become sulfide saturated relatively early in the crystallization history of an
intrusion is considered to be advantageous for thedevelopment of a wide range of magmatic sulfide
deposits. There are several processes by which
sulfide saturation can be promoted and these are
discussed below.
Addition of externally derived sulfur
Perhaps the simplest way of promoting saturation
in a magma whose composition is sulfur undersaturated is to increase the global amount ofsulfur in the melt by addition from an externalsource. Some of the komatiite-hosted Ni–Cu
deposits, such as Kambalda in Western Australia(Box 1.5) and several examples in Zimbabwe
and Canada, are hosted in lava flows that haveextruded onto sulfide-bearing footwall sediments
such as chert, banded iron-formation, or shale. The
komatiitic host lava commonly cuts down into
its footwall, a feature attributed to a combination of thermal erosion and structural dislocation.
Good geological and geochemical evidence exists
to suggest that the komatiitic magma arrived on
surface undersaturated with respect to sulfur, andthat saturation was achieved by a combination of
crystal fractionation and assimilation of crustallyderived sulfur derived from footwall sediments(Lesher, 1989). This concept is consistent withfactors such as the localization of Ni–Cu ore
within the footwall embayments created by extruded komatiitic lavas and is discussed in moredetail in Box 1.5.
Injection of a new magma and magma mixing
As mentioned in section 1.3.2 above, large magma
chambers are seldom the result of a single injection of melt and most witness one or more replenishments during the crystallization sequence.
The Bushveld Complex and the Stillwater Complex in Montana contain important chromite
deposits as well as PGE-enriched base metal ores
associated with a sulfide-rich layer (the Merensky
reef in the Bushveld and the J-M reef at Stillwater).

61

These two, economically important, layered mafic
intrusions were both characterized by crystallization that was interrupted by periodic injectionsof new, less differentiated magma. The magma
replenishment events broadly coincide, in bothcases, with the development of PGE-enrichedsulfide horizons. In the Bushveld Complex very
detailed Rb–Sr and Re–Os isotopic measurements
indicate that the new magma was more radiogenic (perhaps due to crustal contamination) but
less differentiated than the magma remaining in
the chamber at the time of injection (Kruger andMarsh, 1982; Lee and Butcher, 1990; Schoenberg
et al., 1999). Further discussion on magma replenishment and its role in mineralization in othermafic intrusions is presented in reviews by
Naldrett (1989a, b, 1997, 1999).
The sulfide solubility diagram in Figure 1.21
can be used to demonstrate how the mixing of new
and residual magmas can promote sulfide saturation. Although the proportions of cumulus minerals being extracted from the magma, as well as
the sulfur solubility curve itself, may not be universally applicable, the principles could be modified for application elsewhere. An initial magmacomposition represented by A in Figure 1.21, containing 0.3 wt% sulfide, would be undersaturated
relative to sulfide (i.e. sulfide solubility at thisstage is 0.4%). Extraction of olivine from themagma (A–B in Figure 1.21) would decrease its
FeO content and result in sulfide saturation
with formation of an immiscible sulfide fraction.
With continued crystallization and segregation of
sulfides the magma would remain saturated, following the solubility curve as shown in Figure 1.21.
If after some 20% crystallization (at point C justprior to the appearance of plagioclase as a cumulus
phase) the chamber is replenished with the injection of a new magma (similar to A) and mixing
occurs, then the composition of the mixture would
lie somewhere along the mixing line between A
and C. If the mixture is at AC it would clearly beundersaturated again and any sulfide globules
present would be resorbed back into the magma.
If, however, the chamber is replenished with theinjection of a new magma at point D (after 35%
crystallization, and after the introduction of plagioclase as a cumulus phase), then a mixed magma

62PART 1 IGNEOUS PROCESSES

Silicate–sulfide immiscibility: the komatiite-hosted
Ni–Cu deposits at Kambalda, Western Australia
An important category of magmatic Ni–Cu sulfide deposits
is that related to mafic and ultramafic volcanic rocks(komatiites) in Archean greenstone belts. Deposits of thistype occur in Canada (the Thompson Ni belt of Manitoba)
and Zimbabwe (Trojan and Shangani mines), but the largest
and richest occurrences occur in the Kambalda regionof Western Australia. Although a great deal is now known
about these deposits (Lesher, 1989) this case study focuses
specifically on the source of sulfur in the immiscible
globules of sulfide within the komatiitic lava flows thathost the mineralization.
Komatiites are ultramafic extrusive rocks that were firstdescribed in the Barberton greenstone belt, South Africa, byViljoen and Viljoen (1969). They comprise mainly olivine+ clinopyroxene and typically contain MgO > 18 wt% and
low alkalis. As indicated in section 1.2.2 of this chapter
they are characterized by high Ni contents. Komatiites
are extruded onto the Earth’s surface as high temperature,
low viscosity lava flows characterized by a variety of
volcanic forms of which pillowed lava and bladed spinifex
textures are the most diagnostic. Most komatiites are notmineralized and their Ni is resident mainly in cumulus
olivine, suggesting that on extrusion these lavas were undersaturated with respect to sulfide.
Evidence from the Kambalda region indicates thatvoluminous eruption of hot, low viscosity komatiitic lava
flows formed extensive sheet flows and gradient-controlledlava rivers or channels (showing many of the features
seen in parts of present day Hawaii). Hot lava rivers are
believed to have thermally eroded discrete channels into
the previously consolidated footwall and some of the

Ni–Cu ores at Kambalda exhibit sulfide concentrations at
the base of well defined, linear channelways (Figure 1).
Another feature of this type of ore is that the mineralized
channelways are devoid of interflow sediments eventhough such sediments occur laterally away from the ore
zones at that level. These sediments comprise a varietyof carbonaceous and sulfidic shales as well as sulfidic
chert and banded iron-formation (Bavinton, 1981). The
nature of these sediments has led to suggestions thatthey represented a source of sulfur and this is supported
by S isotope studies showing similarities in the isotope
ratios of ores and interflow sediments (Lesher, 1989). One
implication of this idea is that assimilation of sedimentby thermally eroding komatiitic lava channels enhanced
the sulfide content within the magma and promoted local
sulfide saturation and immiscibility.
Sulfur saturation and immiscibility of a sulfide fraction
early in the magma crystallization history is the fundamental process responsible for mineralization in komatiitic
lava flows. Chalcophile metals, in particular Ni and lesser
Cu, were scavenged from the turbulent, flowing komatiite
magma by the immiscible sulfide globules, which eventually accumulated as massive sulfide ore along the bottomof the channelways (Figure 1). The disseminated ore thatoverlies the massive sulfide ore reflects the static buoyancy contrast that existed between komatiitic crystal mushand massive sulfide. External derivation of sulfur has also
been suggested for the promotion of sulfide saturation in
the very large Noril’sk–Talnakh Cu–Ni–PGE deposits in the
Russian Federation (Grinenko, 1985) and also for the Duluth
Complex in Minnesota (Ripley and Al-Jassar, 1987).

Spinifex textured
komatiite
Lava channel
Disseminated(thermally eroded footwall)
...
ore
.
. . . . . . . . . . . . . . . . . .. . . . . . . . .
. . . . . . . . . . . . . . . . . . .. . . . . . . . .
.
.
.
.. . . . . .. . .
.
.
.
. . . . .. . . . . . .
.

Sulfidicinterflow
sediment

Massive
Ni-Cu ore
Massive basalt

Pillowed komatiite

Figure 1 The characteristics of komatiite-hosted Ni–Ci deposits in the Kambalda region, Western Australia(after Solomon et al., 2000).

IGNEOUS ORE-FORMING PROCESSES CHAPTER 1

New magma injection and magma mixing:
the Merensky Reef, Bushveld Complex
250 km. The origin of the Merensky Reef is a contentious
issue (Naldrett, 1989a) and arguments range from purely
magmatic models to those involving interaction withmagmatic-hydrothermal fluids (Cawthorn, 1999). There is
ample evidence that the Merensky Reef and its contained
PGE mineralization have been involved with a hydrothermal fluid, and this topic is discussed in more detail in
Chapter 2. For the purposes of this discussion, however, itis accepted that the Merensky unit (i.e. the mineralized
reef and the rocks immediately overlying it, called theBastard unit) owes its origin to the turbulence and magma
mixing that accompanied a magmatic replenishment
event (Schoenberg et al., 1999).
Mineralization in the Merensky Reef is evident in
the presence of disseminated base metal sulfides, mainly
chalcopyrite and pyrrhotite–pentlandite, with which minor
PGE sulfides (such as braggite, cooperite, laurite, andUpper Zone

Ferro-gabbronorite

In addition to its huge chromite reserves, the Bushveld
Complex also hosts the world’s largest reserves of platinumgroup elements (PGE). About 80% of the world’s PGE
reserves are located in the complex, from three specifichorizons, namely the Merensky Reef (where exploitationcommenced and which itself contains about 22% of
the world’s total PGE reserves; Misra, 2000), the UG2chromitite layer, and the Plat Reef. A simplified outline
and section of the Bushveld Complex, showing the extent
of the Merensky Reef, is presented in Box 1.4. This casestudy focuses specifically on the Merensky Reef as it wasformed at the time when a major injection of new magma
occurred into the chamber, and it also marks a regional
mineralogical hiatus separating the Critical Zone from theMain Zone (Kruger and Marsh, 1985). It is typically represented by a 1 meter thick, coarse-grained (or pegmatoidal),feldspathic pyroxenite that extends along strike for about

Gabbronorite

Upper

Norite andAnorthosite

Lower

Main Zone

Pyroxenite marker

Merensky and Bastard units
Orthopyroxenite

Lower
Harzburgite

Figure 1 Sr isotope variations(in terms of the initial 87Sr/86Sr
ratio or Ro) with stratigraphic
height in the Bushveld Complex(after F.J. Kruger, personalcommunication).

Lower
Zone

Critical
Zone

Upper

0.704

0.705

0.706

0.707

0.708

Initial 87Sr/86Sr ratio

0.709 0.710

63

64

PART 1 IGNEOUS PROCESSES

Figure 2 The edge of a small pothole showing
anorthosite in contact with pyroxenite. The
Merensky Reef should occur at the contact between
the two rock types.

moncheite) and PGE metal alloys are associated. As ageneral statement, magma replenishment and mixing are
believed to have been responsible for the formation of an
immiscible sulfide fraction, and the concentration of PGE
into these sulfides, at the time of Merensky Reef formation(Naldrett, 1989a). It must be emphasized, however, thatmineralization in the Merensky Reef, as well as in the UG2chromitite and Plat Reef, is a very complex process andthis is discussed in somewhat more detail in section 1.6
and also again in Chapter 2.
Evidence for a new magma pulse at the Merensky unit
is most obvious in terms of mineral–chemical trends. The
most convincing evidence for the input of a new magma,
however, comes from a dramatic shift in the initial 87Sr/
86
Sr ratio (Ro) of rocks below and above the Merensky unit
(Figure 1). Such a change cannot be achieved by crystal
fractionation and must record the input of new magma. In
this case the new magma had a higher initial 87Sr/86Sr ratio(a)
New Merensky
magma

InitialruptureMovement
on ofMigrati l liquid
residua

Footwallcumulate
assemblage

Recrystallized/
metasomatized zone(b)

Continued
down-dip
extension

Movement

Normal Merensky
Reef
Merensky unit magma
Flatterdown-dip
wall

Steep
up-dipwall

Potholereef

Recrystallized/
metasomatized zone

Figure 3 Schematic crosssections illustrating the
progressive development of
potholes in the Merensky Reefby syn-magmatic extension in
the footwall cumulate rocks justprior to and during injection of a
new magma (after Carr et al.,1999).

IGNEOUS ORE-FORMING PROCESSES CHAPTER 1

than the original magma pulse, a trait that is attributed
to crustal contamination of the new magma increment.
Magma mixing, sulfide segregation, and mineralization in
the same interval are also, therefore, likely to be related
to this event. Recognition of replenishment events should
be regarded as a useful criterion for the exploration of
PGE mineralization in layered mafic intrusions.
Another intriguing feature of the Merensky Reef linked
to injection of new magma is the presence of enigmatic
potholes (Figure 2) that are well revealed in undergroundworkings. Potholes are elliptical areas, 100–200 m indiameter, where the Merensky Reef cuts down into its
footwall, often by as much as 20–30 m. They are of considerable inconvenience to mining operations and also
reflect changes in the mineralogy of the PGE, in that metalalloys tend to dominate over PGE minerals (i.e. sulfides,

composition at, for example, AD would be oversaturated with respect to sulfide. This would resultin the segregation of an immiscible sulfide fraction, the mass of which exceeds the expectedcotectic proportion. A period of enhanced sulfide
production could be expected, therefore, to accompany the injection of a new magma at point D.
Magma mixing is, therefore, a process that canlead to the production of an immiscible sulfide
fraction. The process seems to be linked with theformation of some very important PGE-sulfide
bearing horizons such as the J-M Reef in the
Stillwater Complex and the Merensky Reef in the
Bushveld Complex. It is considered again in
the context of a more holistic model in section
1.7 below.
Magma contamination
In the discussion on the formation of monomineralic chromitite seams (section 1.4.3 and Box 1.4) it
was mentioned that one way of forcing a magmacomposition off its cotectic crystallization path
into the stability field of chromite was to contaminate the magma with foreign material rich in
SiO2. The same process will also result in a lowering of the sulfur content required for sulfide
saturation, thereby promoting the formation of an
immiscible sulfide fraction. Figure 1.23 shows a
1200 °C isothermal section of the SiO2-FeO-FeS
system in which the silicate–sulfide immiscibil-

65

tellurides, etc.). Their origin is highly debated (see Buntin
et al., 1985) and has been attributed to either downwarderosional processes (such as convective scouring or
thermal erosion) or upward fluid/melt migration (involving focused vertical migration of filter-pressed interstitial
magma or a hydrothermal fluid). A model that accommodates the Sr isotopic characteristics of the Merensky
unit, however, envisages syn-magmatic extension of the
footwall cumulus assemblage just prior to injection of the
new magma (Carr et al., 1999). The pull-apart rupture
grows in a down-dip direction (perhaps due to subsidence
and slumping) and is then filled by the new magma toform the Merensky unit, with the mineralized reef at its
base (Figure 3). Potholes are another feature that illustrate
just how complex ore-forming processes actually are
when deposits are examined in detail.

1200 °C

Fayalite+ liquid

SiO2

Tridymite+ liquids

Y

Immiscibility field2 liquids(sulfide and silicate)

B
A

FeO

Wustite + liquid

X

FeS

Figure 1.23 The ternary system FeO–SiO2–FeS
at 1200 °C showing how the addition of silica toa homogeneous, sulfide-undersaturated magma will
force its composition below the solvus into the field of
two liquids, one at Y (silicate-rich) and the other at X
(sulfide-rich) (after Naldrett and MacDonald, 1980).

ity field is identified. The composition of a homogeneous magma undersaturated in sulfide at point
A can be forced into the field of two liquids by theaddition of SiO2 (toward B). At B the two liquids inequilibrium will have compositions at Y (silicaterich) and X (sulfide rich). This process could
have applied to the formation of the Ni–Cu ores
at Sudbury (see Box 1.7) where contamination of

66

PART 1 IGNEOUS PROCESSES

Magma contamination and sulfide immiscibility:
the Sudbury Ni–Cu deposits
HuronianSupergroup

N
0

10
km

WhistleFoy

GraniteFraser
GneissLevackMcCreedy W.
Whitewater Group
Ministic
Frood, Stobie
Creighton
Grean
Copper Cliff

Sudbury

Worthington

Figure 2 Typical appearance of the Ni–Cu sulfidebearing “sublayer” at Sudbury showing its
heterogeneous and contaminated nature.

Ni–Cu mineralization associated with the Sudbury Complex in Ontario, Canada represents a fascinating and uniqueore occurrence. Although small (only 1100 km2) in comparison to the Bushveld Complex, the Sudbury layered
mafic intrusion has for a long time been the world’s leading
producer of nickel, and contains over 1.5 billion tons ofore at an average grade of some 1.2 wt% Ni, together with

Sudbury
Complex
Contact sublayer/
Offset dyke
Fault

Figure 1 Simplified geologicalmap of the Sudbury Complex
showing the relationshipbetween the main mass of thelayered mafic intrusion and the
Ni–Cu mineralized sublayer
unit and offset dykes (afterPrevec et al., 2000).

substantial Cu and PGE credits (Misra, 2000). It is evenmore fascinating for the widely held view that it is the
product of a large meteorite impact that struck Earth about
1850 million years ago (Krogh et al., 1984). The meteorite
impacted into a composite crust containing Archean granite
gneisses and the Paleoproterozoic volcano-sedimentary
Huronian Supergroup.
The main mass of the Sudbury Complex is made up of adifferentiated suite of norite, quartz gabbro, and granophyre.
The Ni–Cu ores, however, are found in an enigmatic mafic
unit at the base of the succession termed the sublayer(Figure 1). The interior of the structure is underlain by achaotic sequence of brecciated debris and volcaniclastic
material interpreted as the fall-back from the meteorite
impact. The granitic and sedimentary floor to the structure
also contains dramatic evidence for a violent meteorite
impact in the form of breccia, pseudotachylite (frictional
melt rock formed by very high strain rates), and shattercones. A widely accepted model suggests that Sudbury
magmatism was triggered by the meteorite impact event,or more specifically by partial melting of crustal source
rocks, during the rebound and pressure release that
immediately followed impact (Naldrett, 1989a).
Ni–Cu sulfides (mainly chalcopyrite and pyrrhotite–
pentlandite) are found along the basal contacts of thesublayer as massive ore which grades upwards into more

IGNEOUS ORE-FORMING PROCESSES CHAPTER 1

disseminated mineralization. Significant sulfide mineralization also occurs in the brecciated floor of the structure where
sulfide melt percolated downwards into embayments andfractures. The offset dykes also contain steeply plunging
pods of sulfide mineralization. An enigmatic feature of the
Sudbury ores, however, is their high Ni contents relative to
Cu. High Ni magmatic sulfide ores are generally associatedwith ultramafic rocks (see, for example, Box 1.5 for a description of the Kambalda deposits), with more differentiated rocks being characterized by higher Cu/Ni ratios. An

important, and relevant, feature of the Sudbury ores in thesublayer is their intimate association with exotic ultramafic
inclusions. These inclusions are now known to be thesame age as the main Sudbury Complex itself and display
no mantle isotopic or geochemical signatures (Prevec et al.,2000). Their presence has important implications for thesource of metals in the Sudbury deposits.
An intriguing aspect of the relationship between meteorite impact and Ni–Cu sulfide mineralization at Sudbury
is the notion that the catastrophic impact was responsible(a)Granitegneiss500

Sr (ppm)

400
North
range
300

South
range

200

Mafic-ultramafic
inclusions

100
BasaltFigure 3 (a) Plot of Sr versus
MgO showing the compositional
fields of the mineralized
sublayer from the north andsouth ranges of the Sudbury
structure, and that of themafic–ultramafic inclusions
within the sublayer. The
sublayer is a mixture of melts
derived by near total meltingof granite–gneiss and tholeiitic
basalt of the Huronian
Supergroup. The inclusions
formed by olivine and pyroxene
fractionation of a high-Mg
melt at and differentiatedtoward lower MgO contents(after Prevec et al., 2000).
(b) Schematic and simplified
cross section through a sublayerhosted Ni–Cu sulfide orebody in
the Sudbury Complex showing
some of the features of thePrevec et al. (2000) model forthe origin of the sublayer.

50Olivinecumulate
10

15

20
25
MgO (wt%)

30

35

40

(b)

Granitegneiss

Norite

Su

bl

ay

er

Sudbury
complex(main mass)

Breccia
Massive sulfide
mineralization
Mafic/ultramafic
inclusions

ction
Sublayer inje
Basaltslurryalong footwall(magma, contact
inclusions,sulfide)67

68

PART 1 IGNEOUS PROCESSES

for the extensive and widespread contamination of magma
as it intruded into the dust and debris of the impact scar
(Naldrett et al., 1986). The rocks of the Sudbury Complex
are highly contaminated by crustal material and this isreadily apparent in their high silica and potassium contents relative to normal continental basalts. One of thefeatures of the Sudbury ores is that this type of contamination might have been responsible for the promotion of
sulfide immiscibility and mineralization, as discussed
in section 1.6.4 of this chapter. In detail, however, thepicture is much more complex and the sublayer is now
thought to have an origin that is different to the remainderof the intrusive complex (Prevec et al., 2000; Lightfootet al., 2001). The main mass of intrusive magma is thought
to have been derived by wholesale melting of granite–
gneiss and volcano-sedimentary target rocks immediately
after meteorite impact and elastic rebound. These crustallycontaminated melts crystallized a dominantly plagioclase+ pyroxene assemblage to form the bulk of the Sudbury

the host magma with siliceous country rock isbelieved to have facilitated the formation of an
immiscible sulfide fraction which scavenged Ni
and Cu from the magma to form the ore deposits.
1.6.5 Other magmatic models for mineralizationin layered mafic intrusions
The preceding sections described processes such as
magma replenishment and sulfide immiscibility
that have come to be regarded as providing good“first-order” explanations for mineralization inlayered mafic intrusions. In detail, however, it isevident that additional mechanisms are required
to explain the many differences that exist from one
deposit type to the next, and also the enigmatic
patterns of metal distribution evident from very
detailed study and more rigorous exploration(Cawthorn, 1999; Barnes and Maier, 2002). Suggestions have been made, for example, that evenin rocks as well mineralized as the Critical Zoneof the Bushveld Complex, sulfide saturation might
not have been achieved and the appearance of asulfide melt was not the most important featureof the mineralization process. Instead, it is argued
that, under conditions of low sulfur fugacity (fS2),
PGE might crystallize directly from the silicate
magma in the form of various platinum groupcomplex, from which the offset dykes were also tapped.
The sublayer, however, is considered to have been derived
by melting of a much higher proportion of Huronianbasaltic target material such that its parental magma wasmore mafic than that of the main mass. It crystallized an
assemblage dominated initially by olivine, forming ultramafic cumulate rocks at depth in the structure (Prevec et al.,2000). This particular magma was also sulfur saturated
and segregated significant volumes of immiscible Ni–Cu
sulfide melt, as well as accumulating sulfides from above.
It was then emplaced along the brecciated footwall contact as a basaltic slurry, comprising magma, ultramafic
inclusions, and sulfide melt, at a relatively late stage inthe evolution of the complex. Although convoluted, ameteorite impact model would appear to best explain
the complex geology of the Sudbury Complex and toaccommodate the intricate processes of anatexis, crustal
contamination, and sulfide segregation required to explain
its contained ores.

minerals (PGM, such as braggite, laurite, malanite,moncheite, cooperite) or alloys. Although it isdifficult to conceive how this might happen given
the very low abundances of the PGE in themagma, a novel mechanism involving clusteringof PGE ions has lent credibility to this as an alternative process to sulfide segregation (Tredoux etal., 1995; Ballhaus and Sylvester, 2000).
PGE clusters
Detailed studies of PGE mineralization haveshown that these elements are heterogeneously
distributed in rocks and minerals, even at a
sub-microscopic level. The application of therelatively new field of cluster chemistry has
indicated that the heavy transition metals in amagma will tend to coalesce by metal–metalbonding into clusters of 10 to 100 atoms (Schmid,
1994). Tredoux et al. (1995) have suggested that
this mechanism might also apply to mineralization in natural PGE ores such as the Merensky
Reef of the Bushveld Complex and the J-M Reef at
Stillwater. Theory predicts that heavy transition
metals form more stable clusters than light ones
and that the PGE will therefore tend to cluster
preferentially relative to Cu and Ni. Likewise, theheavy PGE (i.e. higher atomic numbers – Os, Ir,

IGNEOUS ORE-FORMING PROCESSES CHAPTER 1

and Pt) will form more stable clusters than thelight PGE (Ru, Rh, and Pd). Clustering may, therefore, provide a mechanism for fractionating metals in a magma chamber.
Clustering of atoms and molecules in magmas
is a complex process and is controlled by variables
such as temperature and composition. Figure 1.24
schematically illustrates the nature of PGE clustering in magma. Initial coalescence of metals
occurs by metal–metal bonding, but the clusters
are then stabilized by an envelope of ligands (seeChapter 3) such as sulfur or aluminosilicate
species. Although there is no direct experimental
evidence for PGE clusters in magmas (Mathez,1999), circumstantial evidence comes from features such as light/heavy PGE fractionation andmicrobeam detection of minute PGE metal alloys
in natural materials. It has also been noted thatcumulus olivine- and chromite-rich rocks, suchas the harzburgites and chilled margins of the
Bushveld Complex, which show no association
with sulfide segregations, are notably enriched in
PGE relative to less mafic rocks (Davies andTredoux, 1985; Ballhaus and Sylvester, 2000).
Such an enrichment could be attributed to inclusion of PGE clusters into early formed cumulus
phases such as olivine and chromite. Stability of
PGE clusters and their precipitation from themagma as discrete platinum group minerals or
metal alloys is likely to be promoted by a low fS2
environment.
The existence of PGE clusters in mafic magmas
has important implications for understanding themineralization process. If PGE clusters do form
then crystal-chemical considerations and partitioning behavior might be less important thanmechanical concentration mechanisms involving
trapping of clusters as micro-inclusions in any
suitable cumulus phase, be it silicate, oxide, or
sulfide. Figure 1.24 illustrates two options, one
where PGE clusters are incorporated into sulfide
globules and the other where they are trapped as
micro-inclusions in either olivine or chromite(Tredoux et al., 1995; Ballhaus and Sylvester, 2000).
Detailed analyses of PGE in minerals of theMerensky Reef indicate that cluster theory mightwell explain some of the metal distribution patterns not previously explained by conventional69

sulfide partitioning behaviour. Micro-inclusionsof Os–Ir–Pt alloy are found in pyrrhotite and pentlandite, as are other discrete aggregates of Pt withlesser Pd and Au. In olivine and chromite, microinclusions are dominantly the Pt–Pd–Au type
(Ballhaus and Sylvester, 2000). The latter category
is particularly significant as it suggests that the
Bushveld magma contained abundant Pt– (Pd–Au)
clusters early on in its crystallization history andthat concentration of these metals (by trappingthem as inclusions in early cumulus phases) canbe achieved without sulfide saturation having
occurred at all. Admittedly, the concentrationsof PGE in silicate phases have not as yet proved to
be of economic interest, although the process
should remain of interest to platinum explorationists in the future. Chromitite layers can, bycontrast, contain very significant concentrationsof PGE, as in the case of the huge reserves associated with the UG2 chromitite seam in the
Bushveld Complex. This association is discussed
in more detail below.
The role of chromite in PGE concentration
The very significant concentration of PGE in
chromitite layers such as the UG2 in the
Bushveld Complex suggests that mineralization
processes other than sulfide scavenging musthave played a role. The UG2 also contains a different assemblage of the PGE than the Merensky
Reef, for example, and is relatively enriched in
Ru, Os, and Ir, but with similar Pt contents (vonGrünewaldt et al., 1986). At a first glance thiswould seem to suggest that heavy PGE clusters
had been preferentially concentrated as metalalloy inclusions in cumulus chromite grains (similar to the process shown in Figure 2.24), a model
first suggested by Hiemstra (1979). More detailed
work, however, reveals the following characteristics of PGE mineralization in chromitite layers:
1 The chromitite layers do, in fact, containminor sulfide minerals, both interstitially to thechromite grains and as inclusions (of which laurite RuS2 is predominant; Merkle, 1992) withinthem.
2 Pyrrhotite is virtually absent from the chromitites probably because of reaction between sulfides

70PART 1 IGNEOUS PROCESSES

Stabilization(ligand envelope)

Coalescence(metal–metal bonding)

Key:
PGE
Fe
S

(PGE cluster
10–100 atoms)(S-poor pathway)(S-rich pathway)

PGE
inclusions inolivine andchromite

PGE cluster captured
by sulfide melt

Figure 1.24 Schematic diagram illustrating the formation of PGE clusters in a magmatic system and their eventual
inclusion either in an immiscible sulfide fraction (S-rich pathway) or in a silicate or oxide cumulus phase such as
olivine or chromite (S-poor pathway) (after Tredoux et al., 1995; Ballhaus and Sylvester, 2000).

and chromite which partitions additional Fe into
chromite and liberates sulfur (as shown in equation [1.9], after Naldrett and Lehmann, 1988):
4Fe2O3 (chr) + FeS ⇒ 3Fe3O4 (chr) + 0.5S2[1.9]

3 The assemblage Pt–(Pd–Au) occasionally occurs
as discrete metal-alloy inclusions in chromite(Ballhaus and Sylvester, 2000), indicating that theother PGE are associated with sulfide minerals,
either as inclusions within, or interstitially to,chromite.
These characteristics led Barnes and Maier (2002)
to suggest that an association between PGE con-centration and chromitite layers could be related
to both sulfide accumulation and metal clustering. In their model sulfide segregation under highR factor conditions leads to the accumulationof PGE-enriched sulfide cumulates together with
the normal silicate minerals and chromite. The
sulfide phase reacts with chromite consuming Fe
and liberating S, resulting in a Cu–Ni rich residual
sulfide assemblage that crystallizes minor chalcopyrite and pentlandite. The localized loweringof fS2 that results from sulfide–chromite interaction promotes the direct crystallization of PGM
or metal alloys from clusters in the magma sothat PGE concentration in the immiscible sulfide

IGNEOUS ORE-FORMING PROCESSES CHAPTER 1

fraction is dictated by clustering principles rather
than strict partitioning behavior. The unusual
PGE content of chromite-rich cumulate assemblages arises from the fact that Ru and, to a lesser
extent, Rh partition strongly into the oxide phase
whereas other PGE (Pd and Pt) are specifically
excluded from the chromite lattice and remain in
solution. Since Ru and Rh are fractionated into
the chromite, the remainder of the PGE cluster(i.e. Os, Ir, and Pt) destabilizes and is likely to precipitate as PGM or metal alloys. The latter arescavenged by any interstitial sulfides present inand around the chromitite layer.
Application of composite models, as illustrated
above, seems to provide an adequate explanation
for several of the intriguing characteristics of PGE
mineralization in both sulfide-rich and sulfidepoor environments. Detailed observations, in particular intricacies such as the location of PGE
in different mineralogical sites, clearly provide
the sort of constraints that are needed to improve
our understanding of the relevant ore-forming
processes.
Hiatus models
One of the tendencies of traditional models forthe formation of magmatic deposits is to consider
the processes in terms of accumulation of cumulus phases, be they silicate or oxide minerals, or
sulfide melt. An intriguing thought by Cawthorn(1999) has suggested that the opposite, namely thelack of accumulation of cumulus minerals, might
be the essential ingredient in the formation of
PGE and base metal sulfide mineralization styles
in certain types of deposits. The Merensky Reef,
for example, represents a layer in the Bushveld
magma chamber where a major injection of new
magma occurred and where there is a pronounced
mineralogical hiatus in the magmatic “stratigraphy” (see Box 1.6). Magma replenishment could
create a zone of mixing where the liquidus temperature is substantially above that relevant tolocal crystallization, resulting in a transient stage
when no silicate minerals formed. This would
also result in a hiatus in the solidification process,equating to a period, at least in terms of simplified
gravitative settling, with little or no accumula-

71

tion of minerals on the chamber floor. Sulfide accumulation and/or metal clustering could, nevertheless, have continued during this interval so thatenrichment and mineralization was assisted by
the absence of phases that might normally contribute to dilution of the PGE. The view of Ballhaus
and Sylvester (2000) with respect to Merensky Reef
formation is somewhat analogous. They suggestthat PGE concentration occurred by metal clustering in the magma to form an in situ stratiform
PGE anomaly. The anomaly was effectively “frozen
in” at the hiatus created by magma replenishment and local sulfide saturation, as the sulfide
globules scavenged PGE clusters from the magma
and then settled into the underlying cumulate
mush.
Fluid-related infiltration of PGE
There is abundant mineralogical evidence tosuggest that many layered mafic intrusions, aswell as their contained mineralization, have beeneffected by alteration related to hydrothermal
fluids derived either from the evolving magma
itself or from an external source. There is little
doubt that such fluids have played a role in theremobilization of existing magmatic sulfide mineralization. A more pertinent question, however,relates to whether such hydrothermal processes
may not in fact have been dominant in ore formation (Schiffries and Rye, 1990; Boudreau andMeurer, 1999). Discussion of this topic falls outside the confines of this chapter but it is revisitedin Chapter 2.

1.7 A MODEL FOR MINERALIZATION IN LAYEREDMAFIC INTRUSIONSThe Bushveld Complex in South Africa, together
with the Great Dyke of Zimbabwe and theStillwater Complex of Montana, represent theprime examples of chromite and PGE-sulfide ore
formation in layered mafic intrusions. These examples have many features in common and thedetailed work of researchers such as Irvine (1977),
Naldrett (1989a, b), Naldrett and von Grünewaldt(1986), Naldrett and Wilson (1991), and Campbell

72PART 1 IGNEOUS PROCESSES

et al. (1983), as well as many others, facilitates
the compilation of a relatively simple model toexplain the occurrence and formation of these very
important deposits. The nature and origin of important Cu–Ni sulfide deposits such as Sudbury(Ontario), Noril’sk-Talnakh (Siberia), and Duluth(Minnesota) are somewhat different, but theirformation can nevertheless be explained by someof the processes described below. The model
described in this section represents a useful summary, but for ease of explanation is simplified,both conceptually and with respect to the actual
nature of the ore deposits that it represents.
Readers should be cautious in applying this model
too rigorously to the real situation. Additional
detail can be obtained from the original reference
articles, in particular Naldrett (1989b, 1997).
Previous sections of this chapter have emphasized the following features as being importantwith respect to the formation of ores in igneous
intrusions:
1 Crystal fractionation and gravity-induced crystal settling.
2 Density stratification of magma chambers andthe ability of magmas to undergo density changes
as crystallization proceeds.
3 Repeated recharge of chambers by injection of
fresh magma.
4 The ability of fresh magma to find its own density level and to turbulently mix with the residual
magma.
5 The existence of transient periods of crystallization when only a single phase such as chromite ison the liquidus.
6 The formation of immiscible globules of sulfide
liquid once the magma becomes saturated orsupersaturated with respect to sulfide.
7 The very high partition coefficients of siderophilicelements for the immiscible sulfide fraction.
These features are all used to construct themodel in Figure 1.25, which shows a schematizedlayered mafic intrusion containing elements ofmineralization representing well known deposits
such as the PGE-sulfide bearing Merensky, J-M(Stillwater), and LSZ (Lower Sulfide Zone, Great
Dyke) reefs, and the chromitiferous UG and LG(upper and lower groups respectively) seams of
the Bushveld Complex. The UG2 chromitite isparticularly important because it contains vast

reserves of Cr as well as PGE occurring together inthe same, very extensive, horizon.
One empirical observation which has not beenpreviously mentioned, but which turns out to beof considerable importance to understanding boththe position and tenor of these ores, is the “stratigraphic” level in the layered intrusion whereplagioclase first appears as a cumulus phase. This
seemingly innocuous point in the evolution of an
igneous intrusion (clearly marked in Figure 1.25)
divides ore horizons which tend to be less PGEenriched below it from those above it, which tend
to be both more enriched and more substantial(at least in the Bushveld and Stillwater). In Figure 1.16a it is clear that an evolving magma will
become more dense than when it started once
significant plagioclase is extracted. This has important implications for the behavior of newly
injected melt into the magma chamber and affects
the formation of fountains and plumes (Figure
1.16b) as well as the magnitude of the R factor,
which in turn controls metal concentration into
the immiscible sulfide fraction (Figure 1.22). Thus,
with reference to Figure 1.25, any new magma
which is introduced into the chamber early in its
crystallization sequence (i.e. prior to the appearance of plagioclase on the liquidus) will tend tobehave as a fountain. The limited degree of mixingthat does occur (point AC in the “Irvine model,”Figure 1.25) will force the magma composition
into the chromite field and give rise to a chromitite
layer associated with ultramafic cumulates suchas those, for example, of the Great Dyke. The magmacomposition at this stage will be undersaturated
with respect to sulfide (point AC in the “Naldrett
and von Grünewaldt model,” Figure 1.25) and nosulfide ores will, therefore, form. As crystallization
continues, however (A to B in the “Naldrett and
von Grünewaldt model,” Figure 1.25), saturation
will be achieved and immiscible sulfides will
segregate to form disseminations of sulfide ore inan orthopyroxene cumulate rock. The very highpartition coefficients of Cu, Ni, and the PGE into
the immiscible fraction will mean that even withlimited buoyancy and mixing the earliest sulfides
will be enriched in these elements.
As further sulfides exsolve, however, they willrapidly become depleted in low abundance metals
because the residual magma in equilibrium with

IGNEOUS ORE-FORMING PROCESSES CHAPTER 1High RfactorMerensky and J-M
0.4

UG-2
Gabbroic cumulates
Ultramafic cumulates(bronzitites, harzburgites)
LG

QuartzAC

OlivineLow R factor‘Fountain’

DC“Naldrettand vonGrünewaldt
model”B
A

LSZ

Sulfide (wt%)‘Plume’

73

AD

AC
C

D
AD
A

“IRVINE
MODEL”

0020
Crystallization (%)

70

Chromite

Figure 1.25 Generalized model showing the nature of igneous processes that give rise to some of the importantstyles of chromite and PGE–base metal sulfide deposits associated with layered mafic intrusions. LG and UG refer tothe Lower and Upper Group (UG2 specifically) chromite seams of the Bushveld Complex; LSZ is the Lower Sulfide
Zone of PGE mineralization in the Great Dyke; Merensky refers to the Merensky Reef of the Bushveld Complex and
J-M is the J-M reef of the Stillwater Complex, both of which contain PGE–sulfide mineralization. The difference
between ultramafic and gabbroic cumulates in this model is marked by the first appearance of cumulus plagioclasein the latter (modified after Naldrett, 1997).

these sulfides, which itself originally only hadvery low abundances of the PGE, becomes depleted(refer to Figure 1.22 for a quantification of thiseffect). This feature characterizes the Pt–Pd oresof the Lower Sulfide Zone of the Great Dyke, forexample, where despite the existence of Cu–F
sulfides over a 2–3 meter interval, the ores contain viable grades of PGE over only a few centimeters at the base of the mineralized zone.
By contrast, if a new injection of magma takes
place after the first appearance of plagioclase
then it will likely have a lower density than theresidual magma and emplace itself in a turbulent,buoyant fashion at a relatively high level in thechamber. The more efficient mixing that occurs
will again promote chromite precipitation (point
AD in the “Irvine model,” Figure 1.25). Thesechromite seams will be associated with either
gabbroic or ultramafic cumulate rocks, dependingon the relative proportions of liquids A and D thatare mixed together, and it is also possible thatthey will be better developed. Mixing of primitive
magma A with an evolved liquid at D will also, in
this case, result in a mixture that is sulfide oversaturated (point AD in the “Naldrett and vonGrünewaldt model,” Figure 1.25), resulting in theformation of a sulfide immiscible fraction. The
plume effect that is created by the injection of
the lower density primitive magma will result inbuoyancy and a high degree of convective circulation of these sulfides, which translates into a highR factor. PGE are again strongly partitioned into
the sulfide fraction but the grades will be higher
and more evenly distributed because the sulfides
have the opportunity of equilibrating with amuch larger mass of the magma. This situation isapplicable to the formation of the Merensky and
J-M reefs, both of which are formed well above thefirst appearance of plagioclase cumulate rocks. It
also applies to the formation of the UG2 seam
where PGE-sulfides are intimately associatedwith a major chromitite layer.

74PART 1 IGNEOUS PROCESSES

There are many different types of ore deposits
associated with igneous rocks; there is also a wide
range of igneous rock compositions with whichore deposits are linked. Magmas tend to inherit
their metal endowment from the source area fromwhich they are partially melted. Fertile sourceareas, such as metasomatized mantle or sedimentary rock, are usually themselves a product ofsome sort of metal concentration process. Felsicmagmas crystallize to form granites, or their
extrusive equivalents, and are associated withconcentrations of elements such as Sn, W, U, Th,
Li, Be, and Cs, as well as Cu, Mo, Pb, Zn, and Au.
Incompatible elements in felsic magmas are concentrated into the products of very small degrees
of partial melting or into the residual magma at anadvanced stage of crystallization. Such processes
do not, however, very often result in econom-

ically viable ore deposits. Crystal fractionationin mafic magmas, on the other hand, results inimportant concentrations of elements such as Cr,
Ti, Fe, and V, while associated sulfide immiscibility in these rocks results in accumulations ofPGE, Cu, Ni, and Au. Layered mafic intrusions
are very important exploration targets for thissuite of metals worldwide. Primary diamonddeposits represent the very unusual situation
where deep-seated mafic magma vents to the surface as explosive diatreme-maar type volcanoes,
bringing with them older, xenocrystic diamondfrom fertilized mantle. In both mafic and felsicmagmas the latter stages of crystallization areaccompanied by the exsolution of a dominantly
aqueous and carbonic fluid phase that ultimately
plays a very important role in ore formation. It is
this process that is the subject of Chapter 2.

For those readers wishing to delve further into magmatic ore-forming processes, the following references tobooks and journal special issues will help:

Naldrett, A.J. (1999) World-class Ni–Cu–PGE deposits:
key factors in their genesis. Mineralium Deposita, 34,
227–40.
Taylor, R.P. and Strong, D.F. (eds) (1988) Recent advancesin the geology of granite-related mineral deposits.
Canadian Institute of Mining and Metallurgy, special
volume 39, 445 pp.
Whitney, J.A. and Naldrett, A.J. (1989) Ore deposition
associated with magmas. Reviews in Economic
Geology, volume 4. El Paso, TX: Society of Economic
Geologists, 250 pp.

Economic Geology, volume 80 (1985) Special Issue on
the Bushveld Complex.
Kirkham, R.V. et al. (eds) (1997) Mineral Deposit Modeling. Geological Association of Canada, Special Paper 40.
Misra, K.C. (2000) Understanding Mineral Deposits.
Dordrecht: Kluwer Academic Publishers, 845 pp.
Naldrett, A.J. (1989) Magmatic Sulfide Deposits. Oxford
Monographs on Geology and Geophysics. Oxford:
Clarendon Press, 186 pp.

Magmatic-hydrothermal
ore-forming processes

SOME PHYSICAL AND CHEMICAL PROPERTIES OF WATERMAGMATIC-HYDROTHERMAL FLUIDS
water solubility in magmas
first boiling, second boiling
granite-related magmatic-hydrothermal ore
deposits
COMPOSITION AND CHARACTERISTICS OF MAGMATICHYDROTHERMAL FLUIDSPEGMATITES AND THEIR SIGNIFICANCEMETAL TRANSPORT IN MAGMATIC-HYDROTHERMAL FLUIDS
fluid-melt partitioning of trace elements
WATER CONTENT AND DEPTH OF EMPLACEMENT OF GRANITIC

Box 2.1 Magmatic-hydrothermal fluids associated with
granite intrusions. 1 La Escondida porphyry Cu deposit,
Chile
Box 2.2 Fluid flow in and around granite plutons – theCornish granites, SW EnglandBox 2.3 Magmatic-hydrothermal fluids associated with
granite intrusions. 2 The MacTung W skarn deposit,
Yukon, Canada
Box 2.4 Magmatic-hydrothermal fluids in volcanic
environments – the high- and low-sulfidation epithermal
gold deposits of Kyushu, JapanMAGMASFLUID FLOW IN AND AROUND GRANITE INTRUSIONSORIGIN OF PORPHYRY CU, MO, AND W DEPOSITSPOLYMETALLIC SKARN DEPOSITSEPITHERMAL AU–AG–(CU) DEPOSITS
ROLE OF FLUIDS IN MINERALIZED MAFIC ROCKS

2.1 INTRODUCTION
In Chapter 1 emphasis was placed on the concentration of metals during the igneous processes
associated specifically with magma formationand its subsequent cooling and crystallization.
Little mention was made in the previous chapter
of aqueous solutions (or hydrothermal fluids),
even though it is well known that such fluids
have often played a role in either the formation or
modification of various magmatic deposits. Theinteraction between igneous and hydrothermal
processes is an extremely important one for theformation of a wide variety of ore deposit types,
especially in near surface environments wheremagmas and fluids are spatially and geneticallylinked. This chapter introduces the concept of“magmatic-hydrothermal” fluids and concentrates
on those fluids specifically derived from within
the body of magma itself. Particular emphasis isgiven to the fluids derived from granitic magmas
crystallizing in the Earth’s crust.
There is little doubt that the majority of ore
deposits around the world either are a direct product of concentration processes arising from thecirculation of hot, aqueous solutions through the
Earth’s crust, or have been significantly modified
by such fluids. A wide variety of ore-forming processes are associated with hydrothermal fluids
and these can be applied to both igneous andsedimentary environments, and at pressures andtemperatures that range from those applicable
at shallow crustal levels to those deep in the

76

PART 1 IGNEOUS PROCESSES

lithosphere. Many different types of fluids are
involved in hydrothermal ore-forming processes.
The most primitive or “juvenile” of these are
the magmatic-hydrothermal fluids that originate
from magmas as they cool and crystallize at
various levels in the Earth’s crust, and a numberof important ore deposit types are related to theconcentration of metals that arise from circulation of such solutions. This chapter has particular
relevance to ore deposits such as the large familyof porphyry Cu and Mo deposits, as well as the“high-sulfidation” epithermal Au–Ag deposits
that often represent the surface or volcanic manifestations of porphyry deposits. In addition, theformation of greisen-related Sn–W ores, polymetallic skarn mineralization, and pegmatiterelated deposits are also explained in this chapter.
Discussion of processes involved in the formationof hydrothermal ores that are not directly linked
to magmatic activity follows in Part 2 of the book
(Chapter 3).

temporary hydrogen bonds are formed between it
and the oxygen atoms of the neighboring watermolecule (Figure 2.1). Water is, therefore, a liquid
polymer at STP, and this is responsible for its
many anomalous properties.
The physical properties of water that make it so
important for chemical, biological, and geologicalprocesses include:
1 A high heat capacity, which means that it canconduct heat more readily than other liquids.
2 High surface tension implying that it can easily“wet” mineral surfaces.
3 A density maximum at temperatures just above
the freezing point, which means that solid-H2O(ice) will float on liquid-H2O.
4 A high dielectric constant and, hence, an ability
to dissolve more ionic substances, and in greater
quantities, than any other natural liquid.
The last property is particularly relevant to theformation of ore deposits, since water is largely
responsible for the dissolution, transport, and,hence, concentration of a wide range of elements
and compounds, including metals, in the Earth’s
crust.
The chemical properties of water are dominatedby its high dipole moment, a feature related to thefact that the water molecule is not symmetricalbut comprises two hydrogen atoms that are notdirectly opposite one another but offset relative
to the diametric axis of the single, much larger,
oxygen atom (Figure 2.1). The H2O molecule is,therefore, an electric dipole in which the centerof positive charge does not coincide with thecenter of negative charge. Consequently, water

2.2 SOME PHYSICAL AND CHEMICAL PROPERTIESOF WATER

Water (H2O) is a liquid at room or standard temperature and pressure (STP). By contrast, other lightmolecules (CH4, NH3, H2S, etc.) at STP exist as
vapor and form gaseous hydrides. H2O boils at
a much higher temperature than the hydrides
formed by other light elements. The main reason
for this is that the hydrogen of a water molecule
has a very strong affinity for oxygen such that

(a)

(b)
Lone pairs of electrons–ve–veHydrogen O
bond
H
+ve
–ve

Oxygen atomO
H+ve
Hydrogen atoms

Lone pair ofelectrons
H

+ve

H

Figure 2.1 Idealized illustration of
the water molecule and the natureof hydrogen bonding.

MAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2

molecules can align themselves in an electricalfield or orientate themselves around other charged
ionic species. The clustering of water moleculesaround a dissolved ion is referred to as hydration,and this promotes the stability of ions in aqueous
solution. The dielectric constant is a measure of
a fluid’s ability to separate ions or other dipoles
and is a function of the number of dipoles per unit
volume. The high dielectric constant of water
also has the effect of enhancing dissolution andincreasing solubility.
In nature water exists as three different phases,
solid (or ice), vapor (or steam), and liquid(Figure 2.2a). H2O-ice typically exists below
about 0 °C and forms when water molecules arearranged in a hexagonal crystalline structure
similar to many rock forming silicate minerals.
When H2O-ice melts only a small proportion of
the hydrogen bonds are broken and H2O-liquid
retains a large measure of tetrahedral coordination. At temperatures close to the freezing point,
this structure is more tightly knit than ice, which
explains why the liquid form of water is more
dense than its solid, and ice floats. The densityof H2O-liquid varies mainly as a function of temperature, but also pressure, as illustrated in Figure 2.2b. Density defines the difference between
the liquid phase (typically around 1 g cm−3 at room
temperatures) and its coexisting vapor (where
densities are one to two orders of magnitude
lower). Lines of equal density (or volume) in P–T
space are referred to as isochores (Figure 2.2c).
The phase boundary along which liquid and vapor
are in equilibrium (i.e. between the triple point
and the critical point in Figure 2.2a) defines theequilibrium or saturation vapor pressure, also
known simply as the boiling point curve. On the
Earth’s surface, water effectively “boils” whenequilibrium vapor pressure exceeds the prevailing
atmospheric pressure and vapor can bubble offfrom the liquid.
As temperature rises and the density of H2Oliquid decreases, the latter must coexist in equilibrium with a vapor whose partial vapor pressure(and density) is increasing. The boiling point temperature of pure water increases progressively
with pressure until a maximum limit is reached,called the critical point, at 374 °C and 221 bar

77(Figure 2.2a). The critical point is where it is
no longer possible to increase the boiling pointby increase of pressure and is effectively definedas the stage where there is no longer a physicaldistinction (i.e. a contrast in density) between
liquid and vapor. The densities of liquid and vapor
at the critical point have merged to a value ofaround 0.3 g cm−3 (Figure 2.2c). Since the terms“liquid” and “vapor” no longer have any meaningat the critical point, the term “supercritical fluid”(see section 2.3.3 below) is used to describe thehomogeneous single phase that exists at pressuresand temperatures above the critical point.
For an ideal gas which obeys the Gas Law, therelation between pressure (P), temperature (T),and volume (V) or density is expressed by the wellknown equation:
PV = RT[2.1]

where R is the gas constant.
Water only behaves as an ideal gas at very hightemperatures and low pressures and equation[2.1] cannot, therefore, be used to predict aqueous
phase relations under natural conditions in the
Earth’s crust. Consequently, many attempts havebeen made to introduce correction factors andmodifications to the ideal gas equation such that
it will more accurately reflect non-ideal behavior.
The modified forms of the equation, of which
there are many that have been derived by both
empirical and theoretical means, are referred toas equations of state. The simplest, general formof an equation of state, which incorporates twocorrective terms, is an expression known as vander Waals’s equation:
P = [RT/(V − b)] − a(V)

[2.2]

where a is the corrective term that accounts forthe attractive potential between molecules, and b
is a term that accounts for the volume occupiedby the molecules.
A widely utilized equation of state for water athigher pressures and temperatures is the modified
Redlich–Kwong equation, a more detailed discussion of which, together with its applicability
to the chemical and thermodynamic properties

78

PART 1 IGNEOUS PROCESSES

(a)

H 2O

Pressure (bars)l

Criticalpoint

Liquid water(liquid)

221

ca
iti
cr d
r
e i
up fluS

Ice(solid)

Water
vapor

0.06
Triple
point
0.008

(gas)

100
Temperature (°C)

374
900

(b)

1500

600

1000
300

Pressure(atmospheres)

Depth (meters)

2000

cp

500

100

200

300
400
Temperature (°C)

500

600

0.50

0.60

0.70

0.80

0.75

0. 0
30 .40Pressure (bars)

500

0.85

0.90

0.95

600

1.00

(c)

Isochores

400

20
0.
5
0.1

300
cp
200

0.08

Liquid
r
po
0.03
Va

100

0.05

0.01
100

150

200
250
300
Temperature (°C)

350

400

Figure 2.2 (a) Pressure–temperature phase diagram(not to scale) for pure H2O showing
the occurrences of the three phases(solid/ice, liquid water, water
vapor/gas). The Triple Point, where
solid, liquid, and vapor coexist,
is at 0.008 °C and 0.06 bar. The
Critical Point, beyond which there
is no longer a physical distinction
between liquid and vapor, is at
374 °C and 221 bar.The criticaldensity is 0.322 g cm−3.
(b) Schematic visual representationof H2O fluid densities in relation to
the boiling point curve and criticalpoint, cp (after Helgeson, 1964).
(c) Part of the H2O phase diagram
showing actual variations of liquidand vapor equilibrium densities inpressure–temperature space(isochoral densities in g cm−3).

MAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2

of magmatic fluids, can be found in Holloway(1987).

2.3 FORMATION OF A MAGMATIC AQUEOUS PHASE
2.3.1 Magmatic water – where does it come from?
In the very early stages of the Earth’s development
there was probably very little water either in theatmosphere or on the surface of the planet. It
seems likely that the incipient oceans developed
as soon as the early crust had stabilized and rainwater ponded on its surface. Substantial bodies ofwater (oceans or seas) existed by at least 3800 Myr
ago as evident from the preservation of earlyArchean sediments and subaqueously deposited
volcanic rocks. Some of this early water wasderived from the degassing of volcanic magmas asthey extruded onto the early crust, and is referred
to as juvenile water. As plate tectonic processes
progressively dominated Earth processes, water
has been subjected to extensive recycling and it
is likely that much of the fluid introduced bymagmatic activity at the surface in more recent
geological times is no longer juvenile, although it
is still referred to as magmatic.
Subduction of oceanic crust that has beenhighly altered and hydrated by percolating sea
water is a scenario that can be used to explain the
H2O contents of most arc-related andesitic andbasaltic magmas. Arc-related magmas probably,
therefore, contain dissolved water derived frommixing of primitive, mantle-derived fluids and sea
water. A minor component from meteoric fluid(i.e. derived from the hydrosphere) is also possible. It is interesting to note that hydrous basaltic
magmas that pond at the base of the crust, andwhose heat is responsible for initiating anatexisof the lower crust to form granitic melts, may
contribute their water to these melts. Whitney(1989) has suggested that underplating of a felsicmagma chamber with a more dense mafic magma
will result in diffusive transfer of elements andvolatile species across the boundary layer.
Most of the water present in granitic magmas
is, however, derived from the dehydration of minerals in the crust that were themselves melted to79

form the magma. The concept of “fluid-absent” ordehydration melting (i.e. where melting occurs
without the presence of free water in the rock)is regarded as a realistic model for the generationof granites in the Earth’s crust. The process
can best be explained by considering the phaseequilibria of essentially three minerals, namely
muscovite, biotite, and hornblende. Although thedetailed reactions are complex (see Clemens andVielzeuf, 1987; Whitney, 1989; Burnham, 1997),
the approximate conditions for dehydration melting of muscovite, biotite, and hornblende in
relation to an average geothermal gradient are
shown in Figure 2.3. The amount of H2O contained within these three minerals decreases fromaround 8–10% in muscovite, to 3–5% in biotite
and 2–3% in hornblende. Accordingly, the water
activity in melts formed at breakdown of thesehydrous minerals will vary considerably and amagma derived from anatexis of a muscovitebearing precursor is likely to contain more dissolved water than one derived by melting of anamphibolite. Source material comprising dominantly muscovite and progressively buried alongthe 25 °C km−1 geotherm in Figure 2.3 will start tomelt at point A. The water content of that firstformed magma will be 7.4 wt% (Burnham, 1997).
If the source material contained mainly biotite
as the hydrous phase, melting would begin at asignificantly higher temperature and pressure,at point B, and in this case the first-formed melt
would contain only 3.3 wt% H2O. Likewise,
dehydration melting of an amphibolitic sourcerock would only produce a melt at even higher
P–T (point C, Figure 2.3), and this would contain
2.7 wt% H2O. It is apparent, therefore, that granites of variable composition derived from different levels in the crust will initially contain very
different H2O contents.
It is clear that substantial volumes of magmaticwater will be added to the Earth’s crust as granite
magmas progressively build the continents. It isrelevant to point out that melting of a muscoviteor a muscovite + biotite-bearing source rock
(represented in nature by rocks such as metasediments) is likely to yield peraluminous, S-typegranite compositions (see section 1.3.4). The
types of granites with which most Sn–W–U ore

PART 1 IGNEOUS PROCESSES

Amphibolite (2.7% H

2 O)

A

H2 O)
Biotite (3.3%

Muscovite (78(solidus)

12

Granodiorite

Pressure (kb)2

.4% H O)

16

Basalt (solidus)

20

–1

25°

km

C

B

70

60

50

40

30

20
4
10

0

600

800
1000
Temperature (°C)

assemblages are typically associated will, therefore, be relatively “wet,” and contain high initial
H2O contents. S-type granites generally correlatewith the relatively reduced “ilmenite-series”
granites of Ishihara (1977), which inherit their
low fO2 character by melting of graphite-bearing
metasedimentary material. By contrast, meltingof biotite- or biotite + hornblende-bearing source
material (represented by meta-igneous rocks) will
yield metaluminous, I-type granite compositions.
The porphyry Cu–Mo suite of ore deposits areassociated globally with the latter and are typically relatively “dry” compared to S-type granites.
Ishihara’s “magnetite-series” granites are often,
but not always, correlatable with I-type granites
and these are characterized by higher fO2 magmas(Ishihara, 1981).
2.3.2 H2O solubility in silicate magmas
It is evident from the above discussion thatmagmas derived in different tectonic settings arelikely to have had variable initial H2O contents,

Depth (km)80

Figure 2.3 Pressure–temperatureplot showing the approximate
conditions under which
dehydration melting of muscovite-,biotite-, and hornblende-bearing
assemblages would take place inrelation to a 25 °C km−1 geothermal
gradient. Melts formed at A, B, andC respectively are likely to containdifferent initial water contents. The
solidus curves for granite and basalt
are for water-saturated conditions(redrawn after Burnham, 1997).

and that this is partly a function of the amountof water supplied by the source material during
melting. There is, however, a maximum limit to
the amount of H2O that a given magma can dissolve, and this is defined by the solubility of waterin any given magma. The solubility of H2O insilicate magmas is determined mainly by pressure
and to a lesser extent temperature. Figure 2.4ashows the results of experimental determinationsof water content in melts of basaltic, andesitic,
and granitic, or pegmatitic, compositions. The
experiments suggest that water content is strongly
dependent on pressure, with magmas at thebase of the crust (circa 10 kbars) able to dissolve
between 10 and 15% H2O. It would also appear
from Figure 2.4a that for any given pressure,felsic melts are able to dissolve more water thanmafic ones.
When water dissolves in a magma it exists
essentially as hydroxyl (OH) groups, although it
is likely that at higher pressure and water contents discrete molecular water (H2O) also exists(Stolper, 1982). The solubility of water in silicate

MAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2

Depth (km)
0

5

10

15

20

25

(b)
30

35

40
12

20

1100°CH
do 2 O
mi
na
nt

(a)

81

)
°C
00
(11
)
°C
00
de
An
(11
alt
s
Ba

kb

8

8

0
(8 2

fH2O(kb)

H2O (wt%)

ite
an
Gr

10

te
ati
gm
e
p

)
°C
60
–6

10
kb

10
15

6
6

site

5

4

nt

2
20

1

3

5
P (kb)

ina

m

OH

7

9

11

0

do

2 kb

0

0.1

kb

b

4k

H2O saturation

0.2
0.3(XH2O)2

0.4

0.5

Figure 2.4 (a) Experimentally determined solubilities (in wt%) of H2O in silicate melts as function of pressure.
1, basalt (at 1100 °C); 2, andesite (at 1100 °C ); 3, granitic pegmatite (at 660–820 °C) (after Burnham, 1979).
(b) Calculated estimates of the mole fraction of water (XH2O) (squared) in hydrous magmas at 1100 °C as a functionof the water fugacity (fH2O) and pressure (after Stolper, 1982). H2O-saturation is defined by the encompassing curve.

magmas is thought to be governed by the following equilibrium reaction:
H2O(molecular) + Oo ↔ 2OH[2.3]

where Oo refers to the oxygen that bridges orpolymerizes the silicate structure of the magma.
Low viscosity basaltic magmas comprise a smaller
proportion of bridging Oo than more highly
polymerized granitic melts. Basaltic melts may,therefore, accommodate fewer OH groups inOo substituted sites, which could explain theirinability to dissolve as much water as granite. At
high pressures, however, water solubilities are
essentially independent of magma composition.
As a generalization, the mole fraction (squared) ofwater dissolved in any magma is proportional to
the water fugacity (fH 2O or the effective concentration in a non-ideal solution) of the magma andthis is strongly pressure dependent. An indicationof the amount of water dissolved in a hydrous
melt as a function of pressure and water fugacity
is provided in Figure 2.4b. At low pressures, where

total solubilities are low, OH groups are the dominant species of dissolved water in silicate magmas, whereas at higher pressures water solubility
is dominated by molecular H2O (Stolper, 1982).
2.3.3 The Burnham model
The importance of processes whereby zones of
H2O-saturated magma are formed and localized
toward the roof of a granite intrusion, and theirsignificance with respect to granitoid relatedore deposits, was emphasized by the pioneering
work of C. W. Burnham (1967, 1979, 1997). The
concept has stimulated much fruitful researchand continues to receive experimental and theoretical refinement through the work of Whitney(1975, 1989), Candela (1989a, b, 1991, 1992, 1997),
Shinohara (1994), and many others.
When a granitic magma crystallizes the liquidus
assemblage is dominated by anhydrous minerals
and the concentration of dissolved incompatible
constituents, including H2O and other volatile
species, increases by processes akin to Rayleigh82

PART 1 IGNEOUS PROCESSES

0
Solidified
portions

1

S

Countryrock

H2O
saturatedcarapace45

1.0

1000°C

1000°C

Depth (km)

3

6

1.5

Residual
melt

7

2.0

8
2

1

0

1

Distance (km)fractionation (see Chapter 1). At some stage, either
early or late in the crystallization sequence,
granitic magma will become water-saturated,
resulting in the exsolution of an aqueous fluid toform a chemically distinct phase in the silicate
melt. This process is called H2O-saturation, but
it is also often referred to as either “boiling” or“vapor-saturation.” These terms often lead tosemantic confusion and the footnote* should
provide some clarity on the issue. Because theaqueous fluid has a density that is considerably
lower (usually less than or around 1 g cm−3, see
Figure 2.2c) than that of the granitic magma(which is typically around 2.5 g cm−3) it will

2

Pressure (kb)

0.5

2Figure 2.5 Section througha high-level granodioritic
intrusion undergoing progressive
crystallization and showing the
hypothetical position in space ofthe H2O-saturated granite solidus(S), as well as the zone (in gray)where aqueous fluid saturation
occurs in the residual magma(after Burnham, 1979).

tend to rise and concentrate in the roof, or carapace, of the magma chamber. Although some ofthe original OH− in the magma may be utilized
to form hydrous rock-forming minerals (such as
biotite and hornblende), the amount of magmatichydrothermal water formed in this way can bevery substantial.
The concept of the formation of a zone of
H2O-saturated magma in a high-level granite intrusion (2 km depth) which initially contained some
2.7 wt% H2O is schematically illustrated in Figure 2.5. At these shallow depths H2O-saturation isachieved after only about 10% crystallization,
when the water content of the residual magma* Unless the pressure, temperature, and composition of a magmatic aqueous solution are specified, one cannot say whether
it exists as liquid or vapor, or a homogeneous supercritical phase (see Figure 2.2). In such a case it should be referred to by thegeneric term “H2O fluid.” Since the magmatic aqueous phase is so much less dense than the silicate melt from which it
was derived, and because it may contain other low solubility volatile species such as CO2 or SO2, it is often referred to as
the “vapor” or “volatile” phase. In addition, a homogeneous supercritical fluid is one that would effectively fill its container,
and in this sense should be regarded as a gas or vapor, even though its density might be much higher than a gas as wemight envision one at the Earth’s surface. Accordingly, in this book the terms H2O-saturation, boiling, degassing, and vaporsaturation are used interchangeably. The concepts are discussed again in section 2.4.4.

MAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2reaches 3.3 wt%. At low pressures such as thosepertaining in Figure 2.5, the fluid does in fact boil,
since the equilibrium vapor pressure equals thatof the load pressure on the magmatic system andbubbles of gas (i.e. water vapor plus other volatiles
such as CO2) vesiculate. The process whereby
vaporsaturation is achieved by virtue of decreasing pressure (i.e. because of upward emplacementof magma or mechanical failure of the chamber) is
called “first boiling” and is particularly applicableto high level systems. As mentioned earlier, it
is also possible to achieve saturation with respect
to an aqueous fluid by progressive crystallization of dominantly anhydrous minerals underisobaric conditions, and this process is referred to
specifically as “second boiling.” Second boilinggenerally occurs in more deep-seated magmatic
systems and occurs only after a relatively advanced stage of crystallization. As shown later,
the differences between first and second boiling,
and more specifically the timing of H2O fluid saturation relative to the progress of solidification of
the magma, is very important in understanding
how different granite related ore deposits form
(see sections 2.6 and 2.7).
In addition to the strong dependence of boilingon pressure, it is obvious that fluid saturation will
also be a function of the original water content
of the initial melt. Melts that are more enrichedin water and volatiles will achieve saturation earlier (relative to the progression of crystallization)
than those that are poorer in these components.
Experimental studies have confirmed the effects
of pressure and initial water content on fluid
saturation. Figure 2.6 compares the attainment of
H2O-saturation in a typical granite melt at highand low pressures, and in terms of crystallization
sequences and initial H2O content. In a situation
deep in the Earth’s crust (i.e. 8 kbar), where theoriginal granite melt contained 2 wt% H2O,
crystallization would have commenced with
nucleation of plagioclase at temperatures around1100 °C, followed by the appearance of K-feldspar
and quartz on the liquidus at lower temperatures(A–A′–A″, Figure 2.6a). H2O-saturation is achievedat temperatures that are just a few degrees above
the solidus (at A′) and only after over 80% of themelt had crystallized. The solidus is intersected83

at 650 °C, at which point the granite has totallysolidified. In the unlikely event that this granite
was initially H2O saturated at 8 kbar it would havecontained at least 12 wt% H2O and, as illustrated
by the crystallization path B–B′ (Figure 2.6a),would not have started to crystallize plagioclase
until the melt temperature had cooled to around
750 °C. In this case solidification would haveprogressed quite rapidly between 750 and 650 °C
and entirely in the presence of H2O fluid.
At shallower levels in the crust (2 kbar) the situation is quite different (Figure 2.6b). The same
granite magma composition would be saturated
in water if it originally contained only 6–7 wt%
H2O and in this situation (path D–D′) crystallization in the presence of H2O fluid would take placeover a wider temperature interval than at greater
depth. The same crystallization path at 8 kbar
would have existed over much of its temperature
range in the H2O undersaturated field. A granitic
melt at 2 kbar with low initial water content(2 wt%; path C–C′–C″ in Figure 2.7b) would also
crystallize over a significant temperature intervalin the undersaturated field, as with deeper in thecrust, but in this case H2O-saturation would beachieved at a higher temperature (around 700 °C
at C′) and after only some 60–70% crystallization.
These experimental data reinforce the conceptthat an aqueous fluid will exsolve from a granitic
melt as a normal consequence of its crystallization. The Burnham model has great relevance tothe formation of a wide range of ore deposit types.
The porphyry Cu–Mo suite of deposits, epithermal precious metal ores, and polymetallic skarn
type deposits are all examples of deposits whose
origins are related to the processes conceptualized
in this model. More detailed discussion of these
deposit types is presented in later sections.
A note on the mechanical effects of boiling
An aqueous fluid constrained to the roof zone of agranite magma chamber will have limited effecton the concentration of metals unless it has theopportunity to circulate efficiently in and aroundthe intrusive complex from which it is derived.
The appearance of an exsolved H2O fluid withina magma is, however, also accompanied by the

84PART 1 IGNEOUS PROCESSES

(a)
A

B

H2O – saturation

1200

L

1000

T (°C)

L+V
Liq

uid

PI+L

us

P = 8 kb

800
PI+L+V
PI+Af+L

PI+Af+Q+L

PI+Af+L+V

A'

Solidus600

PI+Af+Q+L+V

A"
2

B’

PI+Af+Q+V

4

6
8
10
Wt% H2O

12

14

(b)
C

1200

D

L
L+V

qu

Li

PI+L

P = 2 kb

T (°C)

us

id

1000

PI+L+V

–s
a tu

C'

PI+Af+L+V

O

PI+Af+Q+L

800

rationPI+Af+L

H2

Solidus

PI+Af+Q+L+V

PI+Af+Q+V

600
C"
2

D’
4

6
8
Wt% H2O

10

12

Figure 2.6 Plots of temperature
versus H2O content showing
the crystallization sequencesfor granitic melts cooling andsolidifying at (a) deep crustal levels(8 kbar) and (b) shallower crustal
levels (2 kbar). The bold lines inboth cases refer to the H2Osaturation curve, and also theliquidus and solidus (after Whitney,1989). Pl, plagioclase; Q, quartz;
Af, alkali feldspar; L, melt; V, H2O
fluid. Crystallization paths referto hypothetical situations where
the granite was either initiallyoversaturated in H2O (B–B′ andD–D′) or markedly undersaturatedin H2O (A–A′–A″ and C–C′–C″).

MAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2

Breccia pipe

85

Quartz vein
0.0

0
1

S

1000°CS

6

1.0

1000°C

saturatedcarapace

4
5

1.5

7
S

Figure 2.7 Section througha high-level granodioritic
intrusion showing the nature ofhydrofracturing and breccia pipe
formation that could form aroundthe apical portion of a granite body(after Burnham, 1979).

H2 O
S

Depth (km)3

Pressure (kb)

0.5

2

2.0

8
2

release of mechanical energy, since the volume
per unit mass of silicate melt plus low density
H2O fluid is greater than the equivalent massof H2O saturated magma (Burnham, 1979). At
shallow levels in the crust the volume change
accompanying H2O fluid production may be as
much as 30% (at Ptotal of 1 kbar). This results inoverpressuring of the chamber interior and cancause brittle failure of the surrounding rocks. The
hydrofracturing that results from this type of
failure usually forms fractures with a steep dip,
because expansion of the rock mass takes placein the direction of least principal stress, which
is usually in the horizontal plane. Hydrofractures
tend to emanate from zones of H2O fluid production in the apical portions of the granite body andmay propogate into the country rock and evenreach surface. This concept is schematized in
Figure 2.7.
Experimental work has confirmed that highlevel granite emplacement enhances the possibility of brittle failure, both in the intrusion itself
and in the surrounding country rocks (Dingwell
et al., 1997), thereby providing excellent ground1

0
Distance (km)12

preparation for the efficient circulation of orebearing fluids. The factors that help to promotebrittle failure in high level granite-related oreforming systems include volatile saturation, which
increases magma viscosity because of dehydration,
bubble vesiculation, and rapid cooling.

2.4 THE COMPOSITION AND CHARACTERISTICS OF
MAGMATIC-HYDROTHERMAL SOLUTIONS

2.4.1 Quartz veins – what do they tell us about
fluid compositions?
Quartz veins are the products of precipitation ofsilica from hot aqueous solutions percolating
through fractures in the Earth’s crust. As pressure
and temperature increases, water becomes anincreasingly powerful solvent and can dissolve
significant quantities of most rock-forming minerals. The solubility of quartz in water increases
to about 8 wt% at temperatures of 900 °C andpressures up to 7 kbar (Anderson and Burnham,
1965). When dissolved in water, silica exists in86PART 1 IGNEOUS PROCESSES

(a)

0.5
0.5
0.4

0.3
0.3
0.2
0.2

Solubility (molal)

Solubility (molal)

0.4

0.1
0.1
700
600
6
5
Pre
4
ssu
re (k
b)

500
400
3

2

300

)

°C

e(

ur
rat

pe

Tem

(b)

0.10
0.10
0.08

0.06
0.06
0.04
0.04

Solubility (molal)

Solubility (molal)

0.08

0.02
0.02

0
10

300
9.6

9.2
pH

200
8.4

)

re

100

8.8

Tem

p

tu
era

(°C

Figure 2.8 The variation of quartz
solubility in aqueous solution
as a function of (a) pressure andtemperature and (b) temperatureand pH (after Rimstidt, 1979).

MAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2

the form H4SiO4. This is a typical constituent ofhydrothermal solutions, and explains the common occurence of quartz in veins. The variationof quartz solubility in aqueous solution is, however, quite complex (Rimstidt, 1997) and otherparameters in addition to P and T, such as pH andsalinity, play an important role. The combined
effects of P, T, and pH on quartz solubility inaqueous solutions is shown in Figure 2.8.
2.4.2 Major elements in magmaticaqueous solutions
In addition to silica, water can dissolve significant amounts of other major elements, such as
the alkali metals. In a now classic experiment,
Burnham (1967) reacted granite with pure water
under a variety of conditions and showed thatat high pressures and temperatures (10 kbar and
650 °C) the total solute content of the solution
was about 9 wt% and comprised Si, Na and K in
proportions that are approximately the same as thegranite eutectic (or minimum melt) composition(Figure 2.9). This indicated that material precipitating from an aqueous solution at high P–T could
have the same composition and mineralogy as
granite that crystallized from a silicate melt (i.e.
quartz + plagioclase + K feldspar in approximately
equal proportions). At progressively lower pressures and temperatures, however, the total solutecontent of the solution decreased (to a minimumof about 0.7 wt% at 2 kbar), with the alkali metals(i.e. Na + K) also decreasing relative to silica(Figure 2.9). Close to the surface, therefore, theproducts precipitating from an aqueous hydrothermal solution comprise mainly silica.
In addition to their ability to dissolve cationicspecies (i.e. electron acceptors) such as Na+, K+,
and Si4+, magmatic aqueous solutions can alsotransport significant amounts of Ca2+, Mg2+, andFe2+, as well as a variety of anionic substances, inparticular Cl−. The anions are referred to as ligands (electron donors) and there are several others
which are also commonly found in aqueous solutions, including HS−, HCO3− , and SO42−. These
additional components are important in ore-forming processes and are discussed in more detailbelow. Further discussion of solution chemistry,87

and of the ability of hydrothermal fluids to transport different metals, is presented in Chapter 3.
2.4.3 Other important components of magmaticaqueous solutions
Rocks are mainly made up of some ten major
element oxides that are used to form the moreabundant rock-forming minerals. Any given rocksample also comprises most other known elements, albeit for many of them in vanishingly
small quantities. Their detection, for the mostpart, depends on the barriers imposed by analytical technology. To a certain extent, the same istrue of the solute content of aqueous solutionsin the Earth’s crust, which, although dominatedby a few highly soluble species, contain a wide
range of other constituents, some of which areeasily detectable, with the remainder occurring in
only minute quantities. The trace constituents ofan aqueous solution cannot simply be dismissed,
especially in an ore-forming context, as it is these
ingredients that distinguish an ore-forming fluid
from one that is likely to be barren.
The typical composition of a magmatic aqueous solution in the Earth’s crust is probably best
obtained by direct analysis of waters produced
adjacent to active volcanoes or geothermal springs.
Although it is now known that such fluids are notentirely magmatic in origin (most have significant
contributions from meteoric waters), their composition gives a reasonable indication of the types
and quantities of dissolved components in aqueous solutions. Table 2.1 shows the chemicalcompositions of volcanic and geothermal fluids
from a variety of places around the world. The
data show that these natural fluids are variable incomposition and this reflects the different rocks
through which they have been circulating, andalso, possibly, contamination by other types of
fluid. Typically the solute content of magmatic
fluids is dominated by the alkali and alkali-earth
metal cations and by chlorine as the major ligand,although exceptions do occur (e.g. Rotorua;
Table 2.1).
A few additional comments are relevant. Thereis often a significant amount of carbon dioxide
associated with magmatic fluids and this is88PART 1 IGNEOUS PROCESSES

Qtz

2.2

90

10

2.0
3.0
2.0

80

4.0

2.0

70

20

4.0
4.0
4.0

60

6.0
6.0
8.3

7.6
7.7

2

4

T
8
so otal(w lute
t%
) 6

10

6.0

50

30

Compositional trend withdecreasing pressure

(Te
mp
era
tur
e

90
0–
60
0°C

)

SiO2

5.7
6.0

40

50

8.4

40
9.8

60

7.8

10.0

30

70

Spruce Pine pegmatite
10.0 = pressure (kb)

20

80

90

10

Ab

90

80

70

60

NaAISi3O8

50
Wt%

40

30

20

10

Or
KAISi3O8

Figure 2.9 Normalized compositions of aqueous fluid in equilibrium with a granitic pegmatite at pressures varying
from 2 to 10 kbar (pressure shown next to each point on the diagram) and temperatures between 600 and 900 °C.
The auxiliary diagram on the left shows the approximate variation in total solute content of the aqueous solutions
as a function of pressure (after Burnham, 1967).

discussed again below. The amount of sulfur inmagmatic fluids is generally low, but this may
reflect the fact that at high crustal levels SO2
partitions into the vapor phase on boiling. The
oxidized and reduced forms of sulfur are essentially mutually exclusive and exist either as the
SO 42− complex (with S6+), or as the reduced HS−
complex (with S2−). Relatively oxidized I-typemagmatic fluids tend to comprise sulfur as SO 42−

which fractionates into the aqueous liquid phase.
Porphyry Cu and Mo type deposits, therefore, areassociated with abundant sulfide minerals in theform of pyrite and chalcopyrite. S-type magmas,
which are more reducing because of equilibration
with graphite-bearing metasediments, exsolveaqueous fluids that contain mainly H2S or HS−
and have lower total sulfur contents. The presenceof reduced sulfur species promotes the stability of

MAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2

89

Table 2.1 Typical solute concentrations (in mg kg−1 or ppm) of the main components that comprise natural aqueous
solutions from a variety of geothermal and volcanic fluids
Location

pH

Na+

K+

Ca2+

Mg2+

Fe2+

Cl−

HCO3−

−
SO2−
4

White Island
Mahagnao
Rotorua
Sea water

1.1
5.8
6.8
–

8630
20 340147
10 560

960
4840
20
380

2010
2900
8
400

3200
95
1
1270

6100
<0.1
<0.1
–

7300
46 23513
18 900

<1
20
560
140

6600
138
<5
2710

Sea water is shown for comparison.
Source: after Giggenbach (1997).

sulfide minerals down to lower temperatures,such that the Sn and W oxide minerals (cassiteriteand scheelite) associated with S-type granites
tend to form early in the mineralizing sequence of
these systems (Burnham, 1997).
It is important to point out that Table 2.1
illustrates the major dissolved components ofnatural aqueous solutions. The concentrationsof ore metals, which should be regarded as trace
constituents in hydrothermal fluids, are usuallyconsiderably lower than those shown in the table.
Recent analytical developments have allowed
measurements of magmatic ore-forming fluids
directly from the tiny volumes of fluid trapped inindividual fluid inclusions and this has provided
some important new insights into the compositions of both the aqueous liquid and vapor phases
in granite ore-forming systems.
Magmatic fluid compositions from fluidinclusion analysis
Ore-forming fluids can perhaps best and mostdirectly be studied by examining the fluid inclusions (typically 5–30 µm in diameter) that exist
particularly in quartz, but also in many otherrock- and ore-forming minerals. Studies of porphyry Cu and Mo systems have shown that theearly generation of fluid trapped in primary fluidinclusions is characterized by high temperaturesand high salinities, and also provides evidence ofliquid–vapor phase separation, or boiling (Roedder,1984). The high salinities of these ore-forming
fluids are confirmed by the fact that their associated fluid inclusions often contain daughter

crystals (i.e. tiny minerals that precipitated fromthe fluid on cooling after it was trapped within
the inclusion; see inset microphotograph in
Figure 2.10a). The majority of these daughter
crystals are identified as halite (NaCl) and, to alesser extent, sylvite (KCl), and confirm the factthat magmatic fluids contain significant amountsof K, Na, and Cl. A plot of fluid compositions
derived from microthermometric homogenization experiments on individual fluid inclusions
for a wide range of porphyry Cu deposits is shownin Figure 2.10a. A more detailed description ofthese data and their significance can be found in
Roedder (1984).
A much more accurate and complete assessmentof porphyry Cu-related ore fluid composition,
however, is obtained from the quantitative analysis of individual fluid inclusions by laser-ablationinductively coupled plasma mass-spectrometry(LA-ICP-MS). This remarkable technique allows
recognition and analysis of a wide range of cationicspecies in the tiny volume of inclusion fluid,and down to levels of about 1 ppm. Figure 2.10bshows the results of LA-ICP-MS analysis of two
coexisting fluid inclusions in quartz from the Bajode la Alumbrera porphyry Cu–Au deposit (Ulrich
et al., 2001). The two coexisting fluid inclusions,
one a vapor-rich inclusion and the other a hypersaline liquid-rich inclusion containing several
daughter crystals, indicate that the ore fluid was
boiling as it was trapped. Their analysis reflects theextent to which elements were partitioned between
liquid and vapor, as well as the metal abundancesin each. The plot shows that the majority of elements analyzed partition preferentially (by a factor

90PART 1 IGNEOUS PROCESSES

(a)

H2O

Ice
Hydrohalite

NaCl

Porphyry Cu
fluid inclusions

KCl

KCl

NaCl

(b)

Concentration (vapor-rich fluid inclusion)(ppm)

Th = 720°C
P = 680 bars
100 000

1:1
FeK

10 000Na

Vapor-richfluid inclusion

1000

Cu
Mn
Zn
Pb

100
Ba
10

Rb

Hypersaline,
liquid-rich
fluidinclusion

1000

10 000 100 000

Sr
Cs
Bi

1

Ag

Ce

1

10

100

Concentration (hypersaline liquid-rich fluid inclusion)(ppm)

Figure 2.10 (a) Plot of the system
H2O–NaCl–KCl showing thecompositional range of fluids in
fluid inclusions from a variety ofdifferent porphyry Cu deposits(after Roedder, 1984). Inset
microphotographs show (top) a
hypersaline (containing a vapor
bubble, saturated brine, and adaughter crystal) fluid inclusion,and (bottom) a trail of vapor-richfluid inclusions. (b) Plot of element
concentrations (obtained by laserablation ICP-MS analysis) in avapor-rich fluid inclusion versus
element concentrations in a related
hypersaline liquid-rich fluidinclusion. Schematic illustrationsof the nature of the two fluidinclusions analyzed in order tocreate the plot are shown next to
the respective axes. The data
pertain to the Bajo de la Alumbrera
porphyry Cu–Au deposit inArgentina. The plot is redrawn after
Ulrich et al. (2001).

MAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2

of about 10 to 20 ×) into the liquid phase rather
than the vapor (Figure 2.10b). As expected, theNa and K contents of the ore fluid are high, but
so is the Fe content. The liquid also contains
significant quantities of Mn, Cu, Pb, and Zn insolution. Interestingly, Cu behaves differently tothe other metals in that it is preferentially fractionated into the vapor phase, a feature also previously noted by Lowenstern et al. (1991) andHeinrich et al. (1999).
2.4.4 Carbon dioxide in magmatic fluids
After water, CO2 is the most common gas emanating from volcanic eruptions and its presence isimportant in a variety of ore-forming processes,
albeit in a more obscure way than for H2O. Manydifferent studies have confirmed that the solubility of CO2 in magmas is typically an order ofmagnitude lower than that of H2O. It dissolves asmolecular CO2 in normal felsic or mafic melts,
but in alkaline magmas it also exists as carbonate
ionic complexes in solution (Lowenstern, 2001).
As an indication, rhyolite and basalt melts at
3 kbar will normally contain between 1000 and2000 ppm dissolved CO2 compared to values atleast 4–5 times higher in a leucitite melt. CO2
solubilities, therefore, increase as a function ofpressure and magma alkalinity.
Because CO2 solubilities are much lower thanthose of H2O, it follows that CO2 will exsolve
early such that vapors generated at an early stageof solidification will tend to be more CO2-richthan those forming later. Carbon dioxide will
partition preferentially into the vapor phase andit is probable that many normal magmas, both
felsic and mafic, will be vapor-saturated at significant crustal depths, even to the mid-crust(Lowenstern, 2001). Other low solubility gases arealso likely to be associated with the vapor phasecreated by early CO2 saturation or effervescence.
Volatiles that are generated in systems thathave undergone protracted differentiation, or
reach shallow crustal levels, will tend to be H2Odominated since the relatively insoluble gases like
CO2, as well as N2 and others, will already havebubbled away. It is, therefore, possible for magmas to become volatile-saturated before, and inde-

91

pendently of, H2O fluid saturation. The presenceof CO2 in an evolving aqueous fluid within a crystallizing granite will promote immiscibility
between vapor and saline liquid phases of thesolution. Such processes can be very importantduring ore-forming processes since they promote the precipitation of metals from solution(Lowenstern, 2001). Effervescence of CO2 fromthe fluid will also promote certain types of alteration in the host rocks and increase pH in theremaining fluid, further influencing ore-forming
processes (see Chapter 3 for more discussion ofthese and related topics). Thus, although CO2
does not appear to play a direct role in the transport and concentration of metals in hydrothermal
solutions, it nevertheless has a role to play in the
distribution and precipitation of metals, especially via the volatile or vapor phase.
2.4.5 Other important features of magmatic fluids
An important characteristic of magmatic aqueous
solutions at low pressures is that they tend tosegregate into two phases of differing densities
and composition. This feature of hydrothermal
solutions is often overlooked, and it also further
complicates the description of these fluids. It isevident that the so-called boiling-point curve(Figure 2.2) is actually part of a solvus which separates low density H2O from high density H2O.
As emphasized above, this distinction becomes
meaningless in P–T space above the critical point(or the solvus) where densities of liquid and vapor
have merged and only a single “fluid” phase
exists. The critical point of pure water, at 374 °C,
220 bar, and 0.322 g cm−3, migrates towards higher
temperatures and pressures as dissolved salts(usually in the form of alkali chloride complexes)
are added to water. A representation of the criticalpoints for hydrothermal solutions of differingsalinities and the effect on fluid densities (isochores) is shown in Figure 2.11a. The locus ofcritical points, as a function of increasing salinity,
represents a solvus boundary which governs theexistence or otherwise of immiscible behavior inaqueous solutions.
The phase relationships of magmatic aqueous
solutions (brines) are also, therefore, dependent92

PART 1 IGNEOUS PROCESSES

4

9
0 .7

CP

0.5
8

0 .6

1.11

1200

0 .9 9

0.92

(a)
1400

Pressure (bars)

1000
800
CP

600
400
LiquidCP

200

Pure water
10 wt% NaCl
25 wt% NaCl

Vapor
200

400
500
Temperature (°C)

300

600

700

(b)
Critical point curve
1000
1750
A'
900

A

B'
800

B
2000 bars
C

15

B"

00

13

00

700

10

rs

Solvus curves

ba

0
75
0
60

0

600

50

Temperature (°C)

00

A"

500

400

0

50
NaCl (wt%)

100

on this solvus which, depending on the pressure,temperature, and composition of the fluid, dictates whether a single homogeneous phase exists
above the critical point, or two immiscible phases
below it. Figure 2.11b shows a projection of the
H2O–NaCl system onto the T–X plane in which
the solvus curves for various pressures are shown.
The phase relationships show that, for a granite at

Figure 2.11 (a) Boiling point (liquid– vapor) curves forpure water and for two aqueous solutions containing
10 and 25 wt% NaCl, showing the shift of critical point
to higher temperatures and pressures with increasing
salinity. Selected isochores, with relevant densities(in g cm−3), are also shown for each of the three cases.
The diagram is redrawn after Roedder and Bodnar(1980). (b) A projection of the system H2O–NaClshowing the solvus curves, for a variety of pressures
between 500 and 1750 bars, that define whether ahomogeneous aqueous solution will segregate into two
immiscible phases or not (after Bodnar et al., 1985).

600 bar in which an exsolved H2O fluid hasformed with a composition at point A, the fluid
exists below the relevant solvus. Such a fluid
must segregate into two phases, a small proportion of brine containing 78 wt% NaCl (A″) and amore voluminous low density fluid (A′) containing only 1 wt% NaCl (Burnham, 1997). The same
granite at 1 kbar (point B, Figure 2.11b) would also

MAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2

exsolve two aqueous solutions, but with different
compositions. The brine would in this case contain 53 wt% NaCl (B″), whereas the remaining
fluid would have 2 wt% NaCl (B′). The granite at
2 kbar (point C), however, would exist above thesolvus (which falls off the diagram in Figure 2.11b)
and would exsolve a single, homogeneous fluid
phase containing about 12.5 wt% NaCl. It should
be noted that the presence of CO2 in the fluid will
promote immiscibility to much higher pressuresand, as mentioned above, this is one of the important roles that this common gas has to play in oreforming processes. These features of an H2O fluid
exsolving from a granite magma are extremely
important to the understanding of ore-forming
processes because metals partition themselvesdifferently between the high solute content brine
and the more pure H2O phase. This topic is discussed in more detail in section 2.6 below.
The discussion of magmatic-hydrothermal
fluids and ore-forming processes in the preceding
sections has given rise to some confusing terminology. As a summary, and before we proceedfurther, the following definitions are presented inorder to clarify some of the concepts referred tothus far.
• H2O-saturation. At some stage, either earlyor late in the crystallization sequence, granitic
magma will become water saturated, resulting in
the formation of a chemically distinct, aqueous
phase in the silicate melt.
The magmatic aqueous phase can exist as a liquid,
vapor, or homogeneous supercritical fluid. In thelast case it should also be regarded as gas since it isa substance that would fill its container. For thisreason the magmatic aqueous fluid is often simply referred to as the “vapor phase.”• Vapor-saturation, also referred to as boiling,occurs when the equilibrium vapor pressure of
the magma equals that of the load pressure on
the system and bubbles of gas (i.e. steam orwater vapor, CO2, N2, SO2, H2S, etc.) nucleate in
the magma. Vapor-saturation of low solubility
volatiles such as CO2 may be unrelated to, andcan precede, H2O-saturation.
The process of H2O-saturation can be achieved in
two ways, either by progressive crystallization ofmagma, or by decreasing the pressure of the system:93• First boiling refers to the case where vaporsaturation is achieved by virtue of decreasingpressure (i.e. because of upward emplacementof magma or mechanical failure of the chamber)and is particularly applicable to high level systems. It reflects the fact that solubilities ofvolatile phases in a melt increase as a functionof pressure.
• Second boiling refers to the achievement of
H2O fluid saturation by progressive crystallization of dominantly anhydrous minerals underisobaric conditions. It pertains to more deepseated magmatic systems and occurs only aftera relatively advanced stage of crystallization.
• Immiscibility refers to the tendency for brine
solutions at low pressures, or in the presence ofCO2, to segregate into two phases, one a dense,
more saline brine and the other a lower density,
low salinity aqueous solution.

2.5 A NOTE ON PEGMATITES AND THEIR SIGNIFICANCE TO
GRANITE-RELATED ORE-FORMING PROCESSES
Pegmatites are commonly regarded as rocks
derived from magma that may have crystallized
in the presence of a magmatic aqueous fluid. They
are defined as very coarse-grained rocks, typicallyassociated with granites and comprising the
major granite rock-forming minerals. The large
crystals of quartz, feldspar, and muscovite thatmake up the bulk of most pegmatites are often
extracted for industrial purposes. In addition, they
can comprise a wide variety of minor minerals of
more exotic and semi-precious character, such as
tourmaline, topaz, and beryl. Pegmatites also contain concentrations of the large ion lithophile andhigh field strength elements, such as Sn, W, U,
Th, Li, Be, B, Ta, Nb, Cs, Ce, and Zr. Their origin
is a topic that provides a fruitful area of researchbecause of the clues they provide for ore-forming
processes and igneous petrogenesis in general.
Although there are several classification
schemes for pegmatites, the model by Cerny(1991) is particularly convenient because it separates them into characteristic metal assemblages
that have an implied genetic connotation. Twofamilies of pegmatites are recognized, the Nb–Y–F

94

PART 1 IGNEOUS PROCESSES

suite associated with sub-alkaline to metaluminous (largely I-type) granites and the Li–Cs–Ta
suite, also enriched in boron, and typically associated with peraluminous (dominantly S-type)granites. Some pegmatites clearly form as theresult of small degrees of partial melting and form
minor dykes and segregations in high grade metamorphic terranes. Other pegmatites are spatiallyassociated with the cupola zones of large granite
intrusions and may be genetically linked tothe most highly differentiated, water-saturated
portions of such bodies. Many pegmatites, on theother hand, are almost certainly not the productsof water-saturated granite crystallization. The
origin of pegmatites is one of the more fascinatingproblems in igneous petrology and, even though
the processes are now well understood, theyremain complex and somewhat contentious. Boththe early and later developments in the understanding of these rocks are instructive to oreforming processes, and are discussed below.
2.5.1 Early models of pegmatite genesis
Prior to the classic paper by Jahns and Burnham(1969), pegmatites were viewed as the products ofextreme crystal fractionation of mainly granitic
magmas. Experimental work, however, suggested
that pegmatites were distinguished from “normally textured” granite by crystallization in thepresence of an exsolved H2O fluid phase. Morespecifically, Jahns and Burnham (1969) suggested
that the transition from granite to pegmatite
marked the point at which H2O fluid saturation
occurred in the crystallization sequence and that
pegmatites, therefore, formed in the presence ofan immiscible H2O + volatile fraction. The existence of H2O fluid in the magma was used to
explain the large crystal size of pegmatites, thelatter growing, because of extended crystallization due to either volatile-related depression of
the granite solidus, or more efficient diffusionof major elements into low viscosity H2O fluidsaturated magma. Mineral zonation in pegmatites
was attributed largely to the separation of relatively dense silicate melt from the low density
H2O fluid phase, and the attendant segregationof alkalis such that Na remained in the silicate

melt and K partitioned in the aqueous phase. In
addition, it was envisaged that precipitation ofdissolved constituents from the H2O fluid phase
could explain the pockets of exotic minerals(including gemstone-quality crystals formed in an
open space or vug from fluids enriched in Li, B, Cs,
Be, etc.) so often characteristic of pegmatites. The
Jahns and Burnham model was particularly attractive because it was supported by experimental
evidence demonstrating that an aqueous fluid
at high temperatures and pressures is capable ofdissolving very significant proportions of solute(see Figure 2.9).
2.5.2 More recent ideas on the origin ofpegmatites
The Jahns–Burnham model represented a majoradvance in the understanding of pegmatite genesis,
as well as the crystallization and mineralizationof granites generally. More recent work, however,has shown that there are other factors, in addition
to H2O-saturation, which need to be considered
in order to fully understand the formation ofthese rocks and their associated ore deposits. The
research of D. London and co-workers (1990,
1992, 1996) in particular has shown that it ispossible to generate pegmatites from H2O undersaturated granitic melts by undercooling themagma below its normal liquidus temperature.
They have demonstrated experimentally thatkinetic delays in the initiation of crystallizationin felsic magmas that are rapidly undercooledmeans that melts persist in a metastable condition and result in non-equilibrium crystal growth.
Many of the features of pegmatites, including
mineral zonation, large grain size, variable textures, and highly fractionated chemistry, can be
replicated experimentally in terms of the slow
crystallization response to undercooling of granitic
melts. This work has enabled pegmatites to again
be explained by magmatic crystal fractionation
process (as had been suggested prior to the Jahns–
Burnham model) rather than to silicate melt–H2O
fluid exsolution and alkali element segregation.
The London model has removed a fundamental
difficulty always faced by the Jahns–Burnham
model, namely how to explain the symmetricMAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2

Single phase melt
(supercritical)

750

Temperature (°C)

mineral zonation typical of so many pegmatites
in terms of segregation mechanisms that, being
gravitationally induced, should have been intrinsically asymmetric. It also provided a rather
neat explanation for why pegmatitic textures areuncommon in more mafic rocks. Mafic melts
nucleate crystals much more readily than theirfelsic equivalents, resulting in a tendency toward
equilibrium crystallization conditions and theformation of more uniform textures.
It should be emphasized, however, that more
recent work on pegmatites has not, by any means,
precluded the important role of water in theformation of these rocks. A fluid inclusion studyof the Tanco Li–Cs–Ta pegmatite in Manitoba,
Canada, showed that it crystallized from a mixture of alumina–silicate melt plus H2O (withminor CO2) fluid over a temperature intervalthat extended from about 700 °C down to below
300 °C (Thomas et al., 1988). These observations
were, in fact, interpreted as support for the
Jahns–Burnham model. The processes are further
complicated when one takes into account the
major influence of other volatile species, suchas B, F, and P, on pegmatite formation (London,
1996). These three elements, individually andcollectively, lower the granite solidus temperature to below 500 °C and increase the range oftemperatures over which magmatic crystallization occurs. They also dramatically increase
the solubility of H2O in the melt and promote
the crystallization of quartz. In a study of the
Ehrenfriedersdorf Sn–W–B–F–P-rich pegmatites of
the Erzgebirge in southeast Germany, it has beenshown that fluid inclusions are made up of coexisting silicate-rich H2O-poor and silicate-poor
H2O-rich melts (Thomas et al., 2000). On progressive reheating from 500 to 700 °C at 1 kbar pressure, the latter inclusions become poorer in H2O,
whereas the former undergo a steady increase inwater content (Figure 2.12). At just above 700 °C
the two inclusion populations are effectively
identical, with both containing about 20 wt%
H2O. This indicates that the silicate-rich H2Opoor and silicate-poor H2O-rich melts were in fact
derived from a single, originally homogeneous,melt. Even at low pressures (the Ehrenfriedersdorf
pegmatites crystallized at about 1 kbar) these

95

700

Silicate
melt
(A)
650

Fluid-richmelt (B)

600
Melt A + melt B

550
500
0

10

20
30
H2O (wt%)

40

50

Figure 2.12 Plot of H2O content versus temperature
for rehomogenized (at 1 kbar) silicate-rich H2O-poor(A) and silicate-poor H2O-rich (B) melt inclusions fromthe Ehrenfriedersdorf pegmatite. The curve defines
a solvus with a critical point (above which there is nolonger a distinction between the two melts and they
are miscible) at about 700 °C and 20 wt% H2O content(after Thomas et al., 2000).

B–F–P-rich pegmatitic magmas exhibit complete
miscibility between silicate melt and aqueous
fluid such that water can be regarded as infinitely
soluble in this melt. On cooling below 700 °C,
however, immiscibility occurs to form two coexisting melts. One of these is a silicate melt withlow H2O content, while the other is a silicate
melt with very high H2O content. The latter,
however, has a density, viscosity, and diffusivity
that is more akin to an aqueous solution than toa silicate melt. Although its properties are waterlike, it is by definition different from the exsolved
H2O fluid envisaged to have formed in the Jahns–
Burnham model.
It is apparent, therefore, that a magmatic aqueous phase has to be represented by a fluid that ismore complex than simply exsolved H2O. At highpressures and temperatures water can containsignificant quantities of dissolved silica and alkalis which, on cooling, will yield precipitates withbulk compositions not dissimiliar to a granite.
At lower pressures H2O fluid will exsolve from anormal granitic magma, such that crystallization

96PART 1 IGNEOUS PROCESSES

will occur in the presence of an aqueous phase. At
low pressures and in the presence of significant
concentrations of elements such as B, F, and P,
granite melt solidus temperatures are significantly depressed and H2O solubility increases tosuch an extent that H2O-saturation might notoccur at all. In this case two immiscible melts
form, one of which is H2O-rich and has physicalproperties akin to an aqueous solution with highsolute content. Pegmatites could conceivably form
in all these situations.

2.6 FLUID – MELT TRACE ELEMENT PARTITIONING
The rise of magma to shallow levels of the Earth’s
crust will inevitably result in saturation withrespect to an aqueous fluid or vapor phase. Even
though the diffusivity of H2O is relatively low inmost felsic magmas, the buoyant, low density
H2O-fluid, together with entrained melt andcrystals, will move to the apical portions of themagma chamber (Candela, 1991). In an equilibrium situation, trace components of the magma
must then partition themselves between melt,
crystals, and H2O-fluid. The extent to which tracecomponents, such as metals, distribute themselves between these phases is a process that is
potentially quantifiable in terms of partition
coefficients. It is possible to derive equations
that express the partitioning of trace constituents
between melt and H2O-fluid in exactly the same
way as was discussed for the partitioning of traceelements between melt and crystals in section
1.3.4 of Chapter 1. However, considerations ofmelt–fluid partitioning are complicated by thefact that it is necessary to consider partitioning behavior prior to water saturation, and also
because the sequestration of incompatible constituents by the H2O-fluid depends on solubilities
of both the fluid and the magma from which it isderived. A detailed account of the quantificationof melt–fluid partitioning behavior is provided by
the work of Candela and co-workers, summarizedin Candela and Holland (1984, 1986), Candela(1989a, b, 1992), and Candela and Piccoli (1995).
Before the results of this work are discussed in
more detail, brief mention will first be made ofsome of the classic experimental work that provided initial insights into why hydrothermal
fluids are such efficient scavengers of metals and,
therefore, so important in ore-forming processes.
Again, more details on this topic are provided in
Chapter 3.
2.6.1 The classic experiments of H. D. Holland
and others
In a now classic set of experiments, Holland (1972)
showed that the solubilities of many metals in amagmatic H2O-fluid are strongly dependent on its
Cl− concentration, in addition to other parameterssuch as temperature and pH. When experimentally reacting an aqueous solution with a granitic
melt it was observed that metals such as Zn, Mn,
Fe, and Pb are strongly partitioned into the H2Ofluid and that the ratio of metal concentration
between melt and fluid is exponentially proportional to the Cl− concentration (Figure 2.13a andb). These experiments demonstrated quite clearly
that certain metals will only readily dissolve in an
aqueous solution if the latter contains appreciable
amounts of the chloride anion (Cl−). Furthermore,metals will exhibit different magnitudes of partitioning into H2O-fluid according to their variableabilities to bond, or form complexes, with Cl−.
Figure 2.13b also shows that the total Fe
content of the H2O-fluid varies as a function oftemperature, and is higher for any value of Cl− at
750 °C than it is at 650 °C. This confirms theintuitive notion that solubility normally increases
as a function of temperature. The diagram also
shows that the Al content of the fluid remains
constant (because the total Fe content changes
in the same way as does the total Fe/Al ratio)
and that it is not affected by the Cl− concentrationof the aqueous fluid. This indicates that not all
metals will have their solubilities controlled by
Cl− concentrations in the fluid (although they
may still be affected by the presence of other ligands) and that some metals may not even choose
to partition into the aqueous phase at all.
The dependence of metal solubility on Cl− concentrations presupposes that the ligand itself will
partition effectively into the H2O-fluid phase.
Kilinc and Burnham (1972) and Shinohara et al.

MAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2

97

(a)
80
Zinc

[Zn]f
[Zn]m

60

40
f/m

– 2

KZn α [Cl ]
20

0

1.0

2.0

3.0

4.0

5.0

6.0

Cl– concentration in H2O fluid (mol kg–1)
(b)
Iron

8
7

0.12

0.06

0°

ΣFe = a + b . [Cl– ]3

75

0.08

C

0.10

ΣF
e
ΣF /Al
e
Σ
ΣF
75
Fe
0°C
e/
65
Al
0°
C

0.14

°C

0
65

0.04

Cl− (melt) + OH− (melt) ⇒ HCl (fluid) + O2− (melt)[2.4]
Reaction between granitic melt and a chloridebearing aqueous fluid showed that the Cl− ion partitions strongly into the fluid, but that the valueof the partition coefficient varies in a complex
fashion (Figure 2.14). At low pressures where the

5
4
3
2
1

0.02

(1989) have confirmed experimentally that chloride partitions strongly into the magmatic aqueous
phase, according to the reaction:

6

Molal (ΣFe/Al)v

0.16

Moles ΣFev

Figure 2.13 (a) The ratio of theconcentration of Zn in H2O-fluid[Zn]f to that in a granitic melt[Zn]m as a function of the Cl−
concentration; experiments
carried out at various temperatures
between 770 and 880 °C andpressures between 1.4 and 2.4 kbar
(after Holland, 1972). (b) The total
Fe content (left ordinate) and thetotal Fe to Al ratio (right ordinate)in H2O-fluid that equilibrated witha granitic melt as a function of theCl− concentration; experiments
carried out at temperatures of
650 and 750 °C and a pressure of
2.0 kbar (after Burnham, 1967).

0.2

0.4
0.6
0.8
1.0
Cl– concentration in H2O fluid (moles)

1.2

H2O-fluid phase segregates into two immiscible
fractions (one vapor-rich and the other liquidrich), the partitioning of Cl− into the H2O-fluid
phase is fixed by the saturation of Cl− in the coexisting vapor-rich phase. At higher pressures thepartitioning of Cl− into a homogeneous H2O-fluid
phase (i.e. above the critical point) is controlled by
the concentration (or activity) of Cl− in the melt,
but increases with increasing pressure.
The early experiments of Holland and others
showed how important the composition, and more
specifically the nature and content of the ligand

98PART 1 IGNEOUS PROCESSES

0.6 kb
3.5 kb

0.04
1.2 kb–Concentration Cl in melt (moles)

0.06

0.02

0

2.2 kb

4.2 kb4

2

6.0 kb6

8–

Concentration Cl in aqueous phase (moles)
Figure 2.14 Variations in the concentration of Cl−
in a silicate melt with that in an exsolved H2O-fluidfor a variety of different pressures. Cl− is seen to bepartitioned strongly into the aqueous phase relative
to the melt. At low pressures, where liquid–vapor
immiscibility occurs (0.6 and 1.2 kbar), theconcentration of Cl− is fixed by the saturation of Cl−
in the coexisting vapor-rich phase; at higher pressures,for any given concentration of Cl− in the melt ahomogeneous H2O-fluid phase will contain more
Cl− at higher pressures (after Shinohara et al., 1989).

species, of the aqueous fluid phase is to metalsolubility in aqueous solutions. Metal contents
of aqueous solutions are not, of course, only
dependent on the Cl− concentration and manyother factors affect the ability of hydrothermal
fluids to transport metals. A more complete discussion of this topic is presented in section 3.4
of Chapter 3. In the section below, a more detailed consideration is given to the partitioning of
metals between silicate melt and an exsolved
magmatic H2O-fluid in the light of recent theoretical and experimental work.
2.6.2 Complications in the fluid–melt
partitioning of metals
The partitioning of metals between a silicate melt
and an exsolved H2O-fluid is a complex matter.
The extent to which metals are partitioned
between melt and fluid is variable and is largely

controlled by the Cl− (and other ligand) concentration of the fluid which, in an evolving
system, itself changes continuously. In addition,
however, factors such as temperature, pressure,the amount of water exsolved relative to theamount of water remaining in the silicate melt
(i.e. when H2O-fluid saturation is achieved during
crystallization), and the oxygen fugacity (fO2)of the silicate–fluid system also influence metal
partitioning. This means that a partition coefficient cannot be regarded as a constant because
it pertains only to a very specific, transient, setof conditions. Melt–fluid partition coefficients,
therefore, show wide variations as the magmatohydrothermal system evolves.
Candela (1989) pointed out that the metal content of an exsolved H2O-fluid will depend to alarge extent on when in the crystallization sequence water saturation is achieved. This is because
crystal–melt partitioning can be as efficient ameans of distributing and concentrating trace
metals in a magma as can H2O fluid–melt partitioning. It is particularly important to distinguishbetween “first boiling” (when vapor-saturation iscaused by pressure decrease) and “second boiling”(when crystallization of a dominantly anhydrous
assemblage causes the H2O content of the residual melt to rise to the saturation water content for a given level of emplacement). First
boiling processes are simpler to model mathematically and pertain to ore-forming processes
associated with high level granite intrusions
such as those associated with porphyry Cu andMo deposits. Second boiling requires a more complicated mathematical model since the metals
have to partition among melt, crystal, and H2O
fluid phases.
Fluid–melt partitioning during “first boiling”
Candela (1989) has provided equations thatquantify the distribution of a trace component
between H2O fluid and silicate melt for a firstboiling situation. These particular equations areapplicable to those components whose partitioning does not change as a function of the ligand
concentration of the fluid. They are somewhat
similar in form to the exponential equations

MAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2[2.5]
where: Cifluid is the concentration of a component(i) in the H2O fluid at any instant in the evolution of the aqueous phase; Difluid/melt is the H2O
fluid/silicate melt partition coefficient for theiis the initial concentration ofcomponent (i); CO
the component (i) in the silicate melt (definedwater
is theat the instant of water saturation); CO
initial concentration of water in the silicate melt
(defined at the instant of water saturation); and F
is the ratio of the mass of water in the silicate
melt (at any given instant after water saturation)
to the initial mass of water in the silicate melt.
The concentration of the component (i) in theassociated silicate melt (Cimelt) is simply given by:iCimelt = Cifluid/Dfluid/melt[2.6]

where Cifluid is the value obtained from equation[2.5].
Knowing the magnitude of the partition coeffii), which changes as a function ofcient (Dfluid/meltpressure and, hence, saturation water content, theconcentrations of a component such as chloride
in an evolving aqueous phase, as well as in theassociated silicate melt, can be calculated. Diagrammatic representations of Cl concentration
trends in an evolving silicate melt–H2O-fluid
system are shown in Figure 2.15. The appearanceof a magmatic aqueous fluid will typically result
in depletion of the volatile component in the melt
(Figure 2.15b) and this will be accompanied bya concomitant decrease of the same component
in the H2O-fluid (Figure 2.15a), as the systemevolves toward complete solidification. The
efficiency with which the aqueous phase extracts
components, such as chlorine or other chloridecomplexing metals, from the melt is effectively
measured by the ratio of the total or integrated
amount of the component in the H2O-fluid to theamount initially present in the melt at the time ofwater-saturation (Candela, 1989). The efficiency

CI in vapor (mol kg–1)

water Dfluid/melt−1iiCO(1 − [(1 − F)CO])
Cifluid = Dfluid/melt

(a)
5.0
4.0
3.0
2.0

Chlo

rine

1.0

0.8
0.6
0.4
Proportion H2O remaining in melt

0.2

(b)
4000
Cl in melt (ppm)

applicable to fractional melting and crystallization presented in Chapter 1, and are expressedas:99

Fluorine

3000

Chlo

rine

2000
1.0

0.8
0.6
0.4
Proportion H2O remaining in melt

0.2

Figure 2.15 Calculated concentrations of chlorine in(a) an aqueous fluid exsolved from a granitic magma,and (b) the residual magma itself, both as a functionof the proportion of water remaining in the melt.
The calculations were carried out for a situation wherewater saturation is achieved in a “first boiling” scenario
with a magma rising adiabatically upwards from 5.4 km
depth (where water content at saturation is 5 wt%) to
the surface (after which 1 wt% water remains in theClrock). COis assumed to be 4000 ppm and DClfluid/melt
is calculated to vary from 39 at the point of water
saturation to 1 as the magma reaches the surface(after Candela, 1989b).

factor will be dependent not only on obviousi, but also on parameters
criteria such as Dfluid/meltsuch as F (the ratio of the mass of water in the silicate melt at any given instant after water saturation to the initial mass of water in the silicatemelt). Note that not all volatile components will
act in the same way, as is the case for fluorine,
whose concentration in the melt is predicted toremain relatively constant during first boiling(Figure 2.15b).

100

PART 1 IGNEOUS PROCESSES

Calculation of fluid–melt partitioning is more
complex for components such as Cu or Zn, whose
behavior not only varies as a function of pressure,but is also strongly controlled by the Cl− concentration of the fluid phase. Likewise, the partitioning of metals into the aqueous fluid phase in asituation where “second boiling” has taken placealso involves many additional assumptions anda more complex mathematical derivation. The
detailed discussion of both of these situations isbeyond the scope of this book, and the interested
reader is referred to Candela (1989) for moredetail. It is pertinent to note, however, that any
metal which bonds strongly with Cl− in an aqueous solution (such as Cu) will tend to exhibit partitioning behavior that is similar to that shown in
Figure 2.15. In the case of metals such as copper,
therefore, the maximum concentration into theaqueous fluid occurs immediately after watersaturation has been achieved. By contrast, metals
such as molybdenum, which do not bond strongly
with Cl−, and are not, therefore, controlled byconcentrations of this ligand, partition differently
and may be most efficiently concentrated in
the very last portions of the evolving H2O-fluid
phase. The contrasting behavior of Cu and Mo inthe magmato-hydrothermal environment, suggested by these theoretical calculations, has been
confirmed by direct experimental evidence andthis is discussed in the following section.
2.6.3 Experimental confirmation of Cu, Mo, andW partitioning behavior
In a series of experiments run at 750 °C and 1.4
kbar, Candela and Holland (1984) were able tomeasure the extent to which Cu and Mo partition
between granitic melt and a coexisting aqueous
fluid containing both Cl− and F−. The results ofthese experiments are shown in Figure 2.18 andclearly demonstrate that Cu is a metal whose
partitioning behavior is strongly controlled by theCl− concentration of the fluid phase, whereas Mo
remains unaffected by the latter. The value ofCuvaries from about 1, for low salinity fluids,
Dfluid/melt
up to about 50 at high salinities (Figure 2.16a) andCu= 9.1[Cl−].
is determined by the relationship Dfluid/melt
By contrast, although Mo is also preferentiallyMo
partitioned into the H2O-fluid (Dfluid/melt= 2.5), itspartition coefficient is fixed and remains constantirrespective of the Cl− concentration (Figure 2.16b).
It is interesting to note that the partitioning
behavior of both Cu and Mo is unaffected by thepresence of F− in the aqueous solution. This isbecause fluorine tends to partition more strongly
into the silicate melt (DFfluid/melt = 0.2–0.3) thaninto the aqueous fluid and is, therefore, generally
not present as a complexing agent for metals in
the fluid phase.
The fluid–melt partitioning behavior of W is
not as well constrained as that for Cu and Mo,
although experiments do suggest that its properties are similar to those for Mo, but that its
partition coefficient is even lower (DW
fluid/melt ≈ 1:
Manning and Henderson, 1984; Keppler andWyllie, 1991). However, it is also known that
W behaves as an incompatible element in termsof its crystal–melt partitioning behavior in relatively reduced (i.e. S-type) granitic magmas
solidifying at some depth in the Earth’s crust. By
contrast, Mo behaves in a more compatible
manner in the same magma types such thatfractionation will tend to result in increasing
W/Mo ratios (Candela, 1992). The efficiency ofmetal extraction by an aqueous phase exsolving
late in the crystallization sequence of such agranite will, therefore, favor concentration of W
over Mo, despite the lower Dfluid/melt values. The
crystal–melt partitioning behavior of W also
contrasts with that of copper, which tends to act
as a compatible element in virtually any magmacomposition (Candela and Holland, 1986). A
value of DCu
crystal/melt = 3 was calculated for acrystallizing basaltic magma composition and islikely to be >1 for a granite too, since Cu is able tosubstitute very efficiently into accessory sulfide
phases and, to a lesser extent, into biotite andmagnetite.
It is apparent that both fluid–melt and crystal–melt partition coefficients play an importantrole in the distribution of metals in and aroundcrystallizing granite plutons. These parameters,
when coupled with considerations of granite
type, depth of emplacement, and the timing ofwater-saturation relative to the crystallization
sequence, can be used to explain the nature andMAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2

101

(a)
60

50

40
f/m
= 9.1 [Cl–]
DCu

f/m
DCu

30

20

10

10

20

30

40

50

60

40

50

60

[Cl– ] (mol kg–1)
(b)
10
9
8

f/m
= 2.5
DMo

7
6
Figure 2.16 (a) The relationshipbetween H2O fluid/silicate melt
f/mpartition coefficient for Cu (DCu)
and the Cl− concentration of theaqueous fluid. (b) The relationshipbetween the H2O fluid/silicate
melt partition coefficient for Mo
f/m(DMo) and the Cl− concentration of
the aqueous fluid (after Candela andHolland, 1984).

Df/m
Mo

5
4
3
2
1

origin of many different magmatic-hydrothermalore deposit types. A model that summarizes
some of the features discussed in this chapter ispresented in section 2.8 below as an explanation
for the characteristics of granite-related Cu, Mo,
and W deposits.

10

20

30
–

–1

[Cl ] (mol kg )2.7 WATER CONTENT AND DEPTH OF EMPLACEMENT OFGRANITES – RELATIONSHIPS TO ORE-FORMING PROCESSES

Many of the ore deposits associated with granites,such as porphyry Cu and epithermal Au–Ag ores,
are related to magma emplacement at high levels

102PART 1 IGNEOUS PROCESSESof the crust where H2O-fluid production andhydrofracturing can take place. These types ofdeposits are commonly located in the volcanic or
subvolcanic environment and have formed asmuch from the action of surface-derived (or meteoric) water as they have from the circulation ofmagmatic waters. Other deposit types, such as
porphyry Mo and granitoid hosted Sn–W deposits,
are generally associated with magmas emplaced
at deeper levels in the crust. The depth of
emplacement of a granite magma, together withrelated parameters such as magma composition
and initial water content, plays a very importantrole in determining the nature and origin of ore
deposits associated with felsic igneous rocks, andthese topics are discussed below.
Section 1.3.1 described how granitic magmas
derived from the melting of different source
materials, at different pressures and temperaturesin the crust, would contain variable initialamounts of water. A melt derived by anatexis of arock comprising mainly muscovite, for example,
would contain in the region of 7–8 wt% H2O. By
contrast, dehydration melting of an amphibolitic
source rock could only produce a melt at higher
pressure and temperature (i.e. deeper in thecrust) and this would contain only 2–3 wt% H2O(Figure 2.3). It was suggested on this basis that
S-type granite magmas would initially containmore water than I-type granite magmas. In termsof this model the drier I-type granite magmas
would be derived from the deep crust (possiblywith contributions from the upper mantle), whereas
S-type granites come from material melted in themid- to lower-crust. Several workers have used
these concepts to develop models that link granite emplacement depths with their metallogenic
characteristics (Hyndman, 1981; Strong, 1981).
Figure 2.17a shows the same P–T diagram asthat in Figure 2.3 but inverted to reflect the surface (i.e. low pressures) at the top of the diagram.
Hypothetical zones of melting are shown for eachof three cases where the water required to initiate
melting is supplied by the breakdown of muscovite, biotite, and amphibole. If sufficient melt isallowed to accumulate and then to rise upwards
in the crust along an adiabatic cooling path, it isapparent, at least theoretically, that each of thesemagmas would crystallize at different levels in
the crust. Adiabatic upward movement of magma(i.e. where conductive heat loss to the wall rocks
is ignored) would involve cooling at a rate of about
1.5 °C kbar−1 and in P–T space would approximately follow the steep curves that define mineral phase boundaries. Conceptually, therefore,
magmas could rise upwards in the crust until theyintersect the water-saturated granite solidus, bywhich time they would have become completely
solid and could not intrude any further. In reality
crystallization is likely to have occurred prior tothis level because of heat loss to the wall rocks
and the water-saturated solidus effectively represents the depth above which a magma is unlikely
to be emplaced. These considerations suggestthat S-type granite magmas would be emplacedat mid-crustal depths (4–5 kbar), as dictated byintersection of the muscovite breakdown curvewith the water-saturated granite solidus. By contrast, intersection of the biotite or amphibole
breakdown curves with the water-saturated granite solidus indicates that an I-type magma could
move to much shallower crustal levels (1 kbar orless) before completely solidifying (Figure 2.17a).
Figure 2.17b is a schematic crustal profile
showing the relationships between depth of
emplacement and the metallogenic character ofvarious granite-related deposit types. I-type magmas generated deep in the lithosphere usuallyform adjacent to subduction zones and commonlyreceive a contribution from mantle-derived mafic
melts. Forming at high temperatures (1000 °C
or more) and being relatively dry (H2O contents< 3–4 wt%) they will rise to shallow levels of the

Figure 2.17 (Opposite) Strong’s model showing therelationship between level of granite emplacement andmetallogenic character. (a) Pressure–temperature plot
showing the approximate conditions where fluid-absent
melting of amphibolite- (I), biotite- (II), and muscovitebearing (III) protoliths would occur, and the expectedlevels in the Earth’s crust to which melt fractions
would rise adiabatically, as a function of the watersaturated solidus. (b) Schematic diagram illustrating
the emplacement style and metallogenic character of
granites formed under each of the conditions portrayedin part (a) (after Strong, 1988).

(a)

0

0
D
C

4

A

III

6

Granitemelt(8.4% H2O)

II

Water-saturatedgranite s
olidus

7
8
9

400

500

600

I
Dioritemelt(2.7% H2O)

700

8
12
16
20
24

Granodiorite
melt(3.3% H2O)

lidus
Muscovite so

5

AmphibolitBe s s ed4

Volatiled ep rPressure (kb)3

Biotite solidus

solid
u

s

e solidus

2

Depth (km)

1

800

28
32
36

900

Temperature (°C)(b)

GRANOPHILE DEPOSITS(Sn, W, U)

PORPHYRY/EPITHERMAL DEPOSITS(Cu, Mo, Au)(Ground water)

0

DFirst boiling

1
A
2

Connate
water

4

C

8
Veins

4

Metamorphic
water

Pa

rti
al

5

B

Se

m

el

ts

12
Pegmatites

16(Barren)

20

co

nd

bo

ilin

g

6

III

24

7
8
9

28

Muscovite (–biotite) granitoid

32

II

Biotite granitoid
Amphibole (–biotite) granitoid

36

I

Depth (km)Pressure (kb)

3

104

PART 1 IGNEOUS PROCESSES

crust and may even extrude to form substantial
volcanic structures. Such magmas will typicallyexsolve a magmatic vapor phase by first boiling,
an event that will also promote hydrofracturing,
brecciation, and the widespread circulation ofhydrothermal solutions in and around the sites ofmagmatic activity. These are the environments inwhich porphyry Cu, as well as epithermal Au–Ag
deposit types, occur.
By contrast, S-type magmas are generated in themid- to lower-crust by partial melting of a sourcerock that comprises a substantial proportion ofmetasedimentary material. These melts will format relatively low temperatures (around 700 °C)
and will initially comprise significant H2O dissolved in the magma. Such a magma type will
crystallize in the mid-crust, not too far from itssite of generation, and will typically be barren. If
substantial crystal fractionation occurs, however,incompatible trace elements will become concentrated in residual melts, and in those rocks representing the crystallized products of differentiatedmagma (see Box 1.3, Chapter 1). H2O-fluid saturation will also eventually occur in the residual
magma by second boiling, to form pegmatites andrelated deposits. The concentration of volatiles,

Magmatic-hydrothermal fluids associated with
granite intrusions: 1 La Escondida porphyry copper
deposit, Chile
Chile is the world’s leading copper producer, with an
output in 2000 of about 4.6 million tons of fine copper,
or 36% of global production (Camus and Dilles, 2001).
Northern Chile in particular contains several giant porphyry copper deposits, including world-renowned minessuch as Chuquicamata, El Teniente, and El Salvador. The
porphyry deposits of this huge metallogenic province
formed during five discrete magmatic episodes extending
from the Cretaceous to the Pliocene. The most prolific ofthese is the late Eocene–Oligocene period of magmatism,
during which time at least 10 major porphyry copper
deposits were formed, including the two biggest copper
mines in the world, La Escondida and Chuquicamata.
La Escondida was discovered as recently as 1981 and in1999 produced over 800 000 tons of fine copper from areserve base of more than 2 billion tons of ore at 1.15 wt%
copper (Padilla Garza et al., 2001).
La Escondida, together with other giant porphyry systems
such as Potrerillos, El Salvador, Chuquicamata, El Abra, andCollahuasi to its north and south, lies on the Domeyko
fault system (Figure 1). This faulting was dominantly transcurrent during Eocene–Oligocene magmatism. Locallydeveloped areas of transtension and dilation are considered to have created an environment conducive tohigh level ascent of magma and accompanying magmatichydrothermal mineralization (Richards et al., 2001). The
Escondida district actually comprises six separate deposits(of which La Escondida is the largest) contained withina system of left-lateral strike-slip faults. Mineralization
is associated with quartz monzonite (adamellite) andgranodiorite intrusions within which alteration evolved
from early potassic to sericite–chlorite and quartz–sericite.
These assemblages have been overprinted by a younger

South
AmericaN

Domeyko
Fault

Collahuasi
El Abra(39–37 Ma)
Chuquicamata(34–31 Ma)

22° S

Antofagasta
La Escondida(38–34 Ma)

26° S

El Salvador(41–39 Ma)
Potrerillos(37 Ma)
0

200
km

70° W

Figure 1 Simplified map showing the position, in
northern Chile, of the Domeyko fault system and thesix giant porphyry copper deposits located on it.

(a)Lithology
Pit outline
1997

Rhyolitic
dome
f

f

f

f

f
Escondida
stock

(b)

Andesitef
Alteration

Qz-Ser

Chl-Ser (Bt)

Advanced argillic
Epid-Chl
Chl-Ser (Bt)

Kfel-Chl-Ser(c)Figure 2 North–south sections
through La Escondida showing
the lithologies, alteration
zonation, and mineralization
characteristics (after Padilla
Garza et al., 2001).
advanced argillic alteration assemblage, and both hypogene sulfide and supergene sulphide + oxide ores occur(Padilla Garza et al., 2001).
Cross sections through La Escondida (Figure 2) show thatthe system is related to a multiphase, 38 Myr old, intrusion,
the Escondida stock, that cuts through Paleocene andesites.
This was capped, 35 million years ago, by a rhyolite dome.
Three phases of sequential alteration are recorded in thedeposit. The earliest stage involves potassic alteration(K-feldspar + biotite) with associated silicification and propylitic (chlorite–sericite) alteration, and is linked to magnetite, chalcopyrite, and bornite. This was followed byphyllic alteration (sericite–chlorite–quartz) with whichchalcopyrite, pyrite, and molybdenite are associated. The
final hydrothermal alteration stage involves acid-sulfate oradvanced argillic alteration (quartz–pyrophyllite–alunite).
Precipitation of the hypogene sulfide assemblages, mainly
during the first two stages of alteration, is attributed to adrop in temperature as the magmatic system cooled andwas uplifted, as well as to progressive dilution of magmatic
fluids by meteoric solutions (Padilla Garza et al., 2001).
An important aspect of the ore-forming system at La
Escondida is the development of supergene mineralization(see Chapter 4 and Box 4.2) in areas where the threestages of hydrothermal alteration overlap (i.e. where maximum sulfides were concentrated). Supergene processes
were active between 18 and 15 Ma and were responsiblefor 65% of the total copper resources of the mine (Figure 2).

Mineralization

Leached capping

Chalcocite blanket (supergene)Primary Cu-Fe sulfides (hypogene)

The upper portions of the supergene zone comprise aleached cap characterized by low Cu and Mo contents
and limonite, hematite, and goethite. This is underlain by
a supergene-enriched sulfide blanket containing mainly
chalcocite, with lesser covellite and digenite. The best
grades (up to 3.5 wt% Cu) occur in the thickest parts of thesupergene zone, and also coincide with zones of intense
vein stockwork and overlapping hydrothermal alteration.
Copper “oxide” minerals such as chrysocolla, brochantite,
and atacamite also occur, mainly in late fractures.

Figure 3 View of the Escondida open pit showing
the nature of altered host rocks (photo courtesy of
Rio Tinto)

106PART 1 IGNEOUS PROCESSES

as well as elements such as boron and phosphorus,
will, together with exsolved water in the residual
magma, depress the temperature of the watersaturated granite solidus to lower temperatures(500–600 °C) such that the magma could continue its upward migration to higher levels in thecrust (Figure 2.17b). Fluid pressures would probably not be sufficient to fracture rock at these
depths, but other structural controls would nevertheless promote fluid circulation and it is theseenvironments in which Sn greisen deposits, porphyry Mo deposits, polymetallic skarn ores, andmesothermal veins might form.

2.8 MODELS FOR THE FORMATION OF PORPHYRY-TYPE
CU, MO, AND W DEPOSITS
Considerable progress has been made in understanding the processes that occur in a crystallizing granite intrusion in order to form geneticallyassociated base metal deposits (Candela andHolland, 1984, 1986; Candela, 1989, 1991, 1994;
Candela and Piccoli, 1995). As a generalization,
porphyry Cu–Mo deposits are associated with arcrelated “calc–alkaline” or I-type magmas generated adjacent to Andean type subduction zones.
Sn–W deposits are more often associated with Stype granites that are derived by partial meltingof continental crust that would normally include
a significant proportion of metasedimentarymaterial. In certain settings the latter deposits arealso regarded as “porphyry-type” even though
they may differ significantly in their geologicalcharacteristics from the Cu–Mo types. Each ofthese deposit types is considered in turn below.
2.8.1 The origin of porphyry Cu–(Mo) andporphyry Mo–(Cu) type deposits
The family of deposit types known as porphyry
Cu–Mo ores can be subdivided into two groups,
one in which Cu is the dominant exploitable
metal (together with minor Mo and occasionallyAu), and the other where Mo is the dominant
exploitable metal (with minor Cu and sometimes
W). The two groups are referred to below as theCu–(Mo) and the Mo–(Cu) porphyry types, respect-

ively. Porphyry deposits are the world’s mostimportant source of both Cu and Mo and, especially in the circum-Pacific region, there are many
world-class deposits that exploit these metals. A
description of one of the world’s great porphyry
Cu–(Mo) deposits, La Escondida in Chile, is provided in Box 2.1.
As illustrated in Figure 2.18, both types areassociated with the generation of oxidized I-typegranite magmas associated with melting processes
adjacent to subducted oceanic crust. Porphyry
Cu–(Mo) can be explained in terms of a body ofmagma with a relatively low initial H2O content(inherited from the fluid absent melting of anamphibolitic protolith) rising to high levels in
the crust before significant crystallization takes
place. It is considered likely that some melt
fractions from this high level magma chamber
will be tapped off and extrude on the surface.
These fractions will crystallize to form volcanicand subvolcanic (porphyry) suites of rocks whose
compositions will not be highly differentiated (i.e.
granodioritic or rhyodacitic) because of the low
degree of fractionation that has taken place prior
to extrusion. Because the magma is emplacedat low load pressures the saturation water content
will be relatively low and probably not significantly different from the initial water content.
Vapor-saturation will, therefore, occur early inthe crystallization sequence, essentially due to“first boiling.” Even though Cu is a compatible
element in a crystallizing granitic melt (sequestration of Cu into accessory sulfide phases andbiotite result in DCu
crystal/melt > 1), the lack of crystallization means that very little of the metal
will have been removed from the melt by thetime water-saturation occurs. The vapor phase, bycontrast, is characterized by high Cl− concentrations and it will, therefore, efficiently scavenge
Cu from the silicate melt.
In this setting, however, the situation withrespect to Mo is different. Mo is an incompatible element in a crystallizing granitic meltMo< 1) but, because of the low degree of(Dcrystal/melt
crystallization, its concentration in the residual
magma will not have increased significantly
prior to water saturation. On saturation, Mo will
partition into the H2O-fluid phase but, because its

MAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2

Su

1 Cu-(Mo)

Cu

Mo

rface

Mo
Dcrystal/melt= 0.02
Mo
Dfluid/melt = 2.5

Cu
Dcrystal/melt=2
Cu
Dfluid/melt= 9.1[Cl– ]Exsolved
H2O

Cu

Cu
Mo
OH–

Mo

Melt

OH–
Mo

107

Crystals

2 Mo-(Cu)

Cu
Cu

Cu

(2% H2O)

I-typemagmas(oxidized)

Cu Mo
OH–

w
Dcrystal/melt= 0.05
Mo

Mo
Melt

OH–

3

w
Dfluid/melt=1

W-(Mo)

Mo

Su

bd

uc

Cu

Mo
te

d

oc

ea

(3% H2O)
ni

c

cr

us

t

Mo
W

W

Mo

W

Mo

W

(8% H2O)
S-typemagmas(reduced)

Figure 2.18 Schematic model for the origin and formation of porphyry-type Cu, Mo, and W deposits. Detailed
descriptions are provided in the text (modified after the models of Candela and Holland, 1984; 1986; Strong, 1988;
Candela, 1992).

partition coefficient is relatively small and unafMo
fected by the Cl− concentration (Dfluid/melt= 2.5) it
will never attain very significant concentrationsin the fluid phase. In this setting, therefore, a highlevel granodioritic I-type magma will exsolve anaqueous fluid phase that is highly enriched in Cu,
but only moderately so in Mo, and form a typical
porphyry Cu–(Mo) deposit.
A different scenario can be constructed for theformation of Mo–(Cu) porphyry deposits, where itis envisaged that the parental I-type magma might
originally have contained a slightly higher initialwater content (perhaps due to the fluid absent
melting of a biotite-bearing protolith) than the
Cu-dominant situation. As shown in Figure 2.18,
this magma would not normally rise to the same

shallow crustal levels as its drier equivalent.
The saturation water content of the magma isalso significantly higher in this situation and,
consequently, a greater degree of crystallizationneeds to take place before water-saturation isachieved (in this case vapor-saturation mightoccur by second boiling). As the magma crystallizes so will Cu be extracted from the melt
and end up distributed somewhat evenly throughout the rocks representing the marginal zonesof the pluton. This Cu will not, of course, beavailable for partitioning into the aqueous fluid
once saturation is achieved. Mo, on the other
hand, is incompatible and its concentration willcontinue to increase in the residual melt. Whensaturation does occur, Mo will be concentrated

108PART 1 IGNEOUS PROCESSES

even further into the H2O-fluid phase becauseof its favorable partition coefficient. Despite thefact that Cu will be strongly partitioned into
the highly saline H2O-fluid, its abundance in themelt is now significantly depleted due to earlier
crystal–melt sequestration and the vapor phasewill not be significantly concentrated in Cu.
The extraction of a melt fraction from a moredeep-seated magma chamber will give rise to apluton that is less likely to reach the surface andwhose composition is more highly differentiateddue to the greater degree of crystal fractionation
prior to extraction. Porphyry-style mineralization associated with the circulation of a magmatic hydrothermal fluid in and around such an
intrusion may well have Mo concentrations inexcess of Cu, giving rise to Mo–(Cu) porphyrytype mineralization.
2.8.2 The origin of porphyry W-type deposits
The above model can also be used to explain someof the characteristics of W-dominated, porphyrytype ore deposits that are associated with many
magmatic arc settings. Candela (1992) has emphasized the relationship that exists between W-richporphyry-type deposits and the more reduced,
ilmenite-bearing, S-type granites that crystallize at
relatively deep levels of the Earth’s crust and whose
setting is illustrated under (3) in Figure 2.18.
Under oxidizing conditions, such as those applicable to the I-type, magnetite-bearing granites
discussed above, the efficiency of extracting W
from a magmatic-hydrothermal system into anore body is relatively low and this metal is not a
normal component of porphyry Cu–Mo deposits.
Under reducing conditions, however, W behaves
as an incompatible element in terms of its crystal–melt partitioning behavior and its concentration will increase during crystal fractionation.
When a magma forms from the partial melting ofa metasedimentary precursor, the melt formed
will be relatively hydrous and will tend to crystallize relatively deep in the crust. Such a meltis also likely to be peraluminous and relatively
reduced since it will have equilibrated with metasedimentary material that might have containedorganic carbon. The saturation water content underthese conditions will be high and a significant
degree of crystallization will occur before saturation is attained (by second boiling). Because W is
incompatible under these conditions its concentration will rise in the residual melt, whereas
Mo concentrations will tend to decrease because
it behaves in a more compatible way under reducing conditions. When the H2O-fluid phase does
exsolve it will interact with a highly differentiated melt that is significantly enriched in W. The
fluid phase will scavenge some W but the relatively low partition coefficient (DW
fluid/melt ≈ 1) will
ensure that most of the concentration is achievedprior to water saturation. The deposits that formin this setting will be associated with deep-level,
highly differentiated S-type granites with metal
concentrations dominated by W and only minor
Mo.

2.9 FLUID FLOW IN AND AROUND GRANITE PLUTONS
Discussion has focused in this chapter thus far on
the mechanisms by which a magmatic aqueous
fluid phase is formed and its role in the transportof metals derived essentially from the melt. Of
equal importance with respect to the formationof magmatic-hydrothermal ore deposits are considerations of the flow patterns in and aroundthe intrusion that forms the fluid source, and also
the length of time that magmatic-hydrothermal
fluids can remain active subsequent to magma
emplacement. Fossilized fluid flow pathways are
revealed by recognizing the effects of alterationand mineralization that a hydrothermal solution
imposes on the rocks through which it circulates(see Chapter 3). The prediction of where such
fluids have circulated, however, is best achievedby modeling the thermal effects of an evolving
fluid phase in terms of both conductive andconvective heat loss in and around a theoretical
intrusion. There have been many attempts tomodel fluid flow around granite plutons and thesehave proven useful in predicting where optimal
fluid flux, and hence mineralization, should belocated in relation to the intrusion.
Cathles (1981) used a thermal modeling approachto demonstrate that intrusions cool very rapidly,

MAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2

109

(b)

(a)

T + 20 000 years
107
1

100°C

10

2a
Intrusive geometry

3

600°C

T25% (years)

0.10 md
0.25 mdetiv
ec ingv
n ol
Co co
1.0 md1052.75 km

500°C

Co
n
cooductiv
ling e

3a

300°C

4

0.0 md62=
ψSurface T = 0°C

1042.75 km3.75 km(c)

T + 100 000 years

ψ = 0.5
100°C

200°C
1

10
a (km)

Figure 2.19 Models illustrating the thermal and fluid flow characteristics in and around a cooling igneous intrusive
body. (a) Plot of the time taken for the temperature of an intrusion to fall to 25% of its initial temperature (T25%) asa function of the intrusion dimensions (a in km and defined in the inset diagram). Two situations, one where heatis lost by conduction alone (steep slopes) and another where a combination of conduction and convective heat loss
occurs (shallow slopes), are shown. Permeabilities (in millidarcy, md) are shown for the two cases (after Cathles,
1981). (b) Thermal model showing fluid flow and temperatures in and around an intrusion some 20 000 years after
emplacement. For clarity, streamlines (quantified in terms of the stream function ψ) are shown on the left andisotherms on the right hand side, although the model is symmetric about the intrusion. The intrusion was initiallyat 750 °C with a permeability of 0.07 md and the country rocks at 0 °C with a permeability of 0.13 md. (c) Similar
model for the situation 100 000 years after emplacement. Parts b and c are after Norton and Cathles (1979).

at least in terms of a geological time scale. In
Figure 2.19a model curves are plotted which
relate the time for an intrusion to cool to 25% of
its initial temperature, and the size and geometryof the intrusion. The models can be erected forcases where heat is lost either by conduction alone,or by a combination of conduction and convective
fluid flow. For the case where a small (1–2 kmwide) intrusion cools by conduction of heat to thewall-rocks, solidification will be complete andmagmatic-hydrothermal fluid circulation effectively over in about 105–106 years. If cooling is ac-

companied by convective heat loss caused by thecirculation of a hydrothermal fluid through a permeable network of fractures (such as would be thecase if hydrofracturing had occurred; see section
2.3.3 and Figure 2.7) then such a pluton could be
reduced to 25% of its initial temperature within
as little as 104 years. Thus, in a highly permeable
system where fluid flow is maximized by fluid
exsolution and hydrofracturing, heat loss will bevery rapid and a single intrusion will not be able
to sustain a geologically long-lived period of hydrothermal circulation. This feature could work

110PART 1 IGNEOUS PROCESSES

against the development of a viable ore body, if it
is considered that only magmatic fluids are responsible for ore formation. It should, however, be notedthat mineralization in and around magmatic intrusions can also be the result of externally derived
fluids and that these can circulate in response toother factors besides thermal gradients.
Norton and Cathles (1979) have been able tomodel the evolution of hydrothermal fluid flow
with time in and around a granite intrusion.
Figure 2.19b and c illustrate the isotherms andfluid flow streamlines that are likely to exist
around a granite intrusion at two discrete periods(i.e. 2 × 104 and 105 years after magma emplacement and fluid saturation) for situations where
heat is lost by both conduction and convection.
Soon after intrusion of even a small granite body
(Figure 2.19b) a substantial thermal anomaly iscreated for up to 2 km around the intrusion and an
active system of circulatory fluid flow is set up
from both the sides and top of the body. Aqueous
solutions establish a convective cell where flow
is upwards from the intrusion, and circulatesback downwards at some distance away. These
solutions could incorporate waters from thesurrounding country rocks as well as the magmatic fluid derived from the cooling pluton
itself. An indication of the fluid flux around the
intrusion is obtained in the diagrams from the gradients of the stream function which defines thestreamlines (i.e. the closer together the streamlines, the greater the fluid flux). The demise of thehydrothermal fluid cell is demonstrated in Figure

2.19c, where it is clear that the thermal anomaly
providing the energy for fluid circulation, as wellas the fluid convection itself, has diminished toinsignificant proportions within 105 years ofmagma emplacement and fluid saturation. Thisagain reinforces the view that ore-forming eventsare likely to be short-lived in and around small
granite intrusions representing just a single pulseof magma. Likewise, the ore bodies that form
under these sorts of conditions are also likely tobe small and perhaps sub-economic. Larger,
multi-episodic intrusions are required in order tocreate a system where a substantial magmatichydrothermal ore body will form.
A good example of where the nature of fluid
flow in and around a granite intrusion contributes
greatly to understanding the distribution of magmatic-hydrothermal mineralization is provided
by the polymetallic (Sn–W–Cu–Pb–Zn) depositsof the Cornubian batholith in Cornwall andDevon, southwest England. In this classic mining
district mineralized veins occur along the margins of individual granite plutons, but also extend
out into the surrounding metasedimentary country rocks. Mineralization is also characterized bya pronounced regional zonation in the distribution of metals. These patterns can be explained interms of the nature of fluid flow in and aroundindividual granite plutons, as well as the shape of
the intrusions and the extent to which they havebeen exhumed. A description of mineralizationand metal zoning in the Cornubian batholith isprovided in Box 2.2.

Fluid flow in and around granite intrusions:polymetallic mineralization associated with thegranites of Cornwall, southwest England
The Cornubian granite batholith, emplaced at around 280–
290 Ma during the Carboniferous to Permian Variscan
Orogen, is associated with one of the classic polymetallic
mining districts of the world. The batholith outcrops asa series of five major plutons, in addition to a host ofminor granite bodies and related dykes, in the counties ofCornwall and Devon, southwest England (Figure 1). The
batholith is believed to have been emplaced in response
to subduction caused by convergence of Laurasia andAfrica (the assembly of Pangea), with similar granites andstyles of mineralization also evident elsewhere along
the orogenic belt, notably in Spain, Portugal, France, andthe Czech Republic in Europe, and New Brunswick–
Newfoundland in North America. In southwest Englandmining can be traced back to the Bronze Age, but thepeak of activities occurred in the nineteenth century.
Historically, the district has been a predominantly Sn(2.5 million tons of cumulative metal production) andMAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2

111

DartmoorN
Bodmin
0

10

20

30

km

Devon

St AgnesCligga

Cornwall

St Austell

Carnmenellis

Land'sEnd

MineralizedveinsExposedgraniteTin
CopperLead-zinc

Figure 1 Distribution of the major exposed plutons in Cornwall and Devon making up part of the Cornubian
granite batholith, and the distribution of Sn, Cu, and Pb–Zn mineralization in and around these plutons(after Atkinson and Baker, 1986).
Cu (2 million tons of cumulative metal production) producer, although other metals such as Fe, As, Pb, Zn, W,
U, and Ag have also been extracted (Alderton, 1993).
Presently the district is still a major producer of kaolinite(China clay) for the ceramics industry (see section 4.4.1
of Chapter 4).
The Cornish granite, as with many similarly mineralized
bodies elsewhere along the Variscan belt, is peraluminous in composition, enriched in Rb, Li, F, B, and Be, andcan be classified as an S-type granite (Willis-Richards andJackson, 1989). Mineralization in the granite is related
to the early exsolution of a voluminous aqueous phasefrom the magma which is responsible for widespread
alteration of pluton margins and much of the mineralization (Jackson et al., 1989). Alteration is typically zoned
and characterized by the sequence tourmaline – potassic
alteration (K-feldspar and biotite) – sericite – chlorite as
one progresses away from mineralization. Magmatichydrothermal fluids are also responsible for the formationof mineralized quartz veins that are both endo- andexogranitic. Particularly characteristic of the Cornubianmineralization is the presence of sheeted, greisen-borderedvein sets which generally represent the sites of preferential

Sn–W mineralization. These veins consist essentially ofquartz and tourmaline with variable cassiterite and wolframite, and a greisen alteration halo comprising quartz,
muscovite, Li-mica, and topaz. These vein systems often
extend for several kilometers out into the surrounding
country rock and record a well defined metallogenic zoning or paragenetic sequence. The ore mineral zonation
progresses from an early oxide-dominated assemblage(cassiterite and wolframite together with tourmaline) toa later sulfide assemblage comprising chalcopyrite, sphalerite, and galena, together with fluorite and chlorite. The
zonation pattern is evident both vertically within a single
vein system and laterally away from a granite body as shownin Figures 1 and 3. Although metal zonation is complex(Willis-Richards and Jackson, 1989), the progression from
Sn–W to Cu to Pb–Zn reflects an ore fluid that evolved,
in terms of both cooling and chemical changes due to
fluid–fluid and fluid–rock interactions.
Computerized numerical simulations of the fluid flowin and around the Cornubian granites have provided useful insights into the distribution of mineralization, as wellas the controls on metal zonation that is such of featureof the district. One of the characteristics of the Cornubianbatholith is that mineralization is preferentially distributed
along one margin of the exposed granite pluton, such as
the northern margin of the Carnmenellis pluton and thesouthern margins of the St Austell and Bodmin plutons(Figure 1). Two of the factors that influence the pattern ofmagmatic-hydrothermal fluid flow in and around granite
plutons are the shape of the body and its depth of
emplacement (Sams and Thomas-Betts, 1988). Maximum
fluid flow tends to be concentrated on the more gently
dipping margins of a granite body, which accounts for theasymmetry in distribution of mineralization. In addition,
progressive unroofing of a granite by erosion will cause
the locus of convective fluid flow to migrate outwards into
the surrounding sediments. A model for the southeastern
margin of the St Agnes–Cligga Head granite (Sams andThomas-Betts, 1988), much of which is under the sea,
illustrates how progressive unroofing of the body causes
the locus of vertical fluid flow to move progressively awayfrom the pluton margins. Thermal and chemical evolutionof the mineralizing solutions over the time interval represented by unroofing could account for the successive
increments of metal deposition and the regional zonationof metals observed.

Figure 2 (left) Sheeted Sn–W-bearing, greisenbordered veins exposed at Cligga Head, Cornwall.
(b)

(a)

Fluid flow
Zones
(x 10–8 kg m–2 s–1) Sn Cu Pb-Zn
Coastline
40

GranitePorphyry dyke
Mineralized vein(showing dip direction)

20

N
A

CliggaHead AB–20 Original surface
St Agnes
Head(c)

Lead-zinc zone

Exhumation

Copper zone

1
2
3
4
5
6

Tin zoneLine of section
B

0

1

2
km

3

Depth (km)Present surface

Granite

A

B
10

20
Distance (km)30

Figure 3 The regional pattern of metal zonation around the southeast margin of the St Agnes–Cligga Head
granite (a). Numerical simulation of fluid flow in and around the granite as a function of progressive unroofingof the granite with time (b and c; both diagrams after Sams and Thomas-Betts, 1988).

MAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2

113

Skarn deposit types

8

Figure 2.20 Plot of SiO2 versus[FeO + Fe2O3 + CaO + Na2O]/K2O
showing the relationship between
the composition of igneous
intrusions and the dominant
metal in various skarn deposits(after Meinert, 1992).

[FeO + Fe2O3 + CaO + Na2O]/K2O

Fe
6

Au
4

Cu
Zn
2

2.10 SKARN DEPOSITS
The word “skarn” is an old Swedish term thatoriginally referred to the very hard rocks composed dominantly of calc–silicate minerals (i.e.
Ca-rich garnet, pyroxene, amphibole, and epidote)that identify the rather unusual alteration assemblages associated with magnetite and chalcopyrite deposits in that country. It is now widelyused to refer to the metasomatic replacement ofcarbonate rocks (limestone and dolomite) by calc–
silicate mineral assemblages during either contact or regional metamorphic processes. Mineraldeposits associated with skarn assemblages arereferred to as skarn deposits, and are typically the
product of contact metamorphism and metasomatism associated with intrusion of granite into
carbonate rocks. A wide variety of deposit types
and metal associations are grouped into the category of skarn deposits, and these include W, Sn,
Mo, Cu, Fe, Pb–Zn, and Au ores. The different
metals found in skarn deposits are a product ofthe differing compositions, oxidation state, andmetallogenic affinities of the igneous intrusion,as described in Chapter 1 (Einaudi et al., 1981;
Misra, 2000). A simple diagram relating granitoid

W
Mo
60

65

70

Sn

75

SiO2 (wt%)

composition to skarn deposit type is shown in
Figure 2.20. As a general rule Fe and Au skarn
deposits tend to be associated with intrusions ofmore mafic to intermediate compositions. Cu, Pb,
Zn, and W are linked to calc–alkaline, magnetitebearing, oxidized (I-type) granitic intrusions, andMo and Sn with more differentiated granites thatmight be reduced (S-type) and ilmenite-bearing. It
should be noted that there are exceptions to thisgeneral trend.
Skarn deposits can be classified into calcic or
magnesian types, depending largely on whether
the host rock is a limestone or dolomite. They arealso described as either endo- or exo-skarns,
depending on whether the metasomatic assemblage is internal or external to the intruding pluton. Most of the large, economically viable skarn
deposits are associated with calcic exoskarns.
Tungsten skarns produce the bulk of the world’s
W production and are typically associated with
intrusion of calc–alkaline intrusions, emplaced
relatively deep in the crust. Examples include theKing Island mine of Tasmania and the MacTungdeposit in the Yukon territory of Canada (seeBox 2.3). Copper skarns, by contrast, are often
associated with high level porphyry-style intrusions

114

PART 1 IGNEOUS PROCESSES

Magmatic-hydrothermal fluids associated with granite
intrusions: 2 The MacTung tungsten skarn deposit,
Yukon, Canada
The northern Cordillera of Canada, extending from British
Columbia into the Yukon Territory, is a highly prospective
zone that contains several major porphyry, SEDEX, andVMS base metal mines, as well as gold deposits. In addition, several large W–Cu–(Zn–Mo) skarn deposits occur
in the region, including the world’s largest known depositin the Macmillan Pass area. The deposit, known as
MacTung, has yet to be mined but contains ore reservesof some 63 million tons at 0.95 wt% WO3, with minor
copper (Misra, 2000).
The MacTung deposit occurs in Cambrian–Ordovician
clastic and carbonate sediments deposited on a continental margin in a platformal shelf setting (Atkinson and
Baker, 1986). The sequence is characterized by alternating siltstones, carbonaceous shales, and limestones. It is
intruded by the 90 Myr old Cirque Lake stock, a biotiteand garnet-bearing quartz monzonite of probable S-typeaffinity. This granite is implicated in the mineralization
process, although the actual source of the magmatichydrothermal fluids associated with alteration and metal
deposition is believed to be a hidden intrusion at depth tothe south of the deposit (Figure 1). Two distinct zones ofskarn mineralization are evident at MacTung. The lower
zone is hosted within a folded limestone breccia, whereas
the upper zone straddles three separate lithological units,
namely a lower limestone breccia, an intermediate pelitic
unit, and an upper unit comprising alternating shale andlimestone. Prograde hydrothermal alteration varies as afunction of the host sediment composition and reactivity
with fluids. Non-reactive and less porous shales are characterized mainly by quartz veining with narrow bleachedhalos and low concentrations of scheelite. Highly reactive
and porous limestones are the main hosts to skarn mineralization, and are pervasively altered. A distinct zonation isevident and alteration progresses from marginal limestone
cut by occasional garnet–pyroxene veins, to a pervasively
altered intermediate zone comprising extensive limestone replacement by garnet–pyroxene skarn, and finallya core of pyroxene and pyroxene–pyrrhotite skarn. Ore
grades improve progressively from the margins to thecore, where WO3 and Cu concentrations are better than
1.5 and 0.2 wt% respectively (Atkinson and Baker, 1986).

eon

Mt Allen

er

z
re

o

p
Up

ne

f
r

we

e
or

zo

Lo

e/
on

est
Limale
sh

a

cci

ale
Sh

re
eb

n

o
est
Lim
e
on
est
m
i
L

ed
raliz

e

Min

Hidden granite
stock (?)

e

elop

env

f

Fluid path throughpervious, reactive
limestone (pervasiveskarn)
Fluid flow throughimpervious, non-reactive
lithologies (quartz veins)

Figure 1 Generalized geology of the MacTung deposit and a cross section of the ore zone showing
mineralized lithologies and the possible location of a hidden intrusion suggested to be the source of theore-bearing magmatic-hydrothermal fluids (after Atkinson and Baker, 1986).

MAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2

115

Retrograde alteration is minimal at MacTung and is characterized by late quartz and calcite veins with coarse grained
scheelite, and minor amphibole formation representing
hydration of the pyroxene skarn.
The pattern of alteration zonation, which is discordant
to the Cirque Lake stock contact, as well as the distribution of quartz veins, suggest that fluids migrated up-dip
and toward the stock. For these reasons the source ofmagmatic-hydrothermal fluids involved in skarn formation is believed to be a hidden intrusion at depth to thesouth of the mineralization, as suggested in Figure 1. It is
not known whether such an intrusion is merely part of
the Cirque Lake stock or an unrelated event.
Figure 2 The Cirque Lake stock intruding
Cambrian–Ordovician clastic and carbonate
sediments in the vicinity of the MacTung skarn
deposit.

and many porphyry copper systems that intrude
carbonate host rocks have copper skarns associated with them. A classic example is the Bingham
district of Utah, USA, which contains not only ahuge porphyry Cu deposit, but also the world’s
largest Cu skarn deposit. In addition, the nonporphyry ores at Bingham contain economically
viable Pb–Zn–Ag ores in limestones that are distal
to the copper mineralization (Einaudi, 1982). It isinteresting to note that the world’s largest goldmine, at Grasberg in the Ertsberg district of West
Papua, also exhibits a porphyry–skarn association
and actually produces gold as a by-product of
the copper mining process (Meinert, 2000). Goldskarns associated with porphyry Cu mineralization are associated with emplacement of highlevel, oxidized, magnetite-bearing granitoids.
Other gold-specific skarn deposits, where Au occurs in association with a Bi–Te–As metal assemblage, are linked to more reducing, ilmenite bearing
granitoid intrusions (Meinert, 2000). Fe skarns,
which occasionally form large, economically
viable magnetite deposits, such as at Sverdlovsk
and Sarbai in Russia, are associated with more
mafic gabbroic to granodioritic intrusions, andare typified by endoskarn alteration and sodium
metasomatism (Einaudi et al., 1981). Tin skarns
are generally associated with highly differentiatedS-type (or ilmenite-bearing) granitoids, a good

example of which is the Renison Bell mine in
Tasmania.
Even though there are so many different metal
associations in skarn deposits, the processes bywhich they form are similar, namely granitoid
emplacement and magmatic-hydrothermal activity, albeit at different levels in the crust. An
association with granite intrusion cannot always
be demonstrated, but is usually inferred. Skarn
deposits typically form as a result of three sequential processes (Einaudi et al., 1981; Meinert,
1992). These are isochemical contact metamorphism during early stages of pluton emplacement
and crystallization, followed by open systemmetasomatism and alteration during magmatic
fluid saturation, and, finally, draw-down and mixing with meteoric fluids (for a definition see
Chapter 3) during cooling of the pluton.
Prograde – isochemical contact metamorphism(Figure 2.21a)
As the granite pluton intrudes the country rocksediments are subjected to contact metamorphism and the formation of a variety of hornfelsic
textures. The mineral assemblages that form at
this stage reflect the composition of the lithotypes within which they form. Contact metamorphism is largely a thermal effect although fluids

116PART 1 IGNEOUS PROCESSES

(a)

(b)
Carbonate
Exoskarn (metasomatic)

Descending
meteoric waters

Stage 2

Stage 3ticagm
Ma idsflu

Isochemicalalteration
Stage 1

Crystallizing
graniteintrusion

PROGRADE:
Stage 1 Isochemical (contact
metamorphism)
Stage 2 Metasomatic (magmatic
fluid exsolution)

RETROGRADE:
Stage 3 Meteoric fluid influx,
metal precipitation

Figure 2.21 The evolution of intrusion-related skarn deposits showing the three sequential stages of formation.
(a) Prograde stages, and (b) retrograde stage (modified after Corbett and Leach, 1998).

are likely to circulate during this process andare largely a product of prograde metamorphic
reactions and, therefore, comprise mainly H2O
and CO2 (see Chapter 3 for a description of metamorphic fluids). In dolomitic units metamorphic
mineral zonation approximates the sequence
garnet–clinopyroxene–tremolite–talc/phlogopite,reflecting increased distance and progressively
more hydrous assemblages away from the intrusion. In limestone units the mineral zonation isgarnet–vesuvianite + wollastinite–marble. Thereis no mineralization associated with this stage,
although the process of dehydration close to thepluton margins may be important for increasing
porosity of the source rocks and facilitating fluid
flow during later episodes of mineralization.
Prograde – metasomatism and replacement(Figure 2.21a)
The second stage in the formation of skarn
deposits involves H2O-fluid and vapor-saturationof the intruding magma (as a function of either
first or second boiling or both) and the egress ofthe fluid phase into the surrounding contact
metamorphic halo. At deeper crustal levels fluid

MAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2

flow is likely to be concentrated along discrete
structural or bedding parallel conduits, whereas at
higher levels fluid flow will be more pervasively
distributed, perhaps due to hydrofracturing, ina broad halo in and around the granite cupola.
Metasomatic mineral assemblages will be similar
to those formed during contact metamorphism
but alteration will be more pervasive and coarsergrained, and will replace earlier formed assemblages. Si, Al, and Fe, as well as other components,
will be introduced into the calcareous sediments
by the aqueous magmatic fluid, while Ca, Mg, andCO2 are locally derived and also introduced into
the metasomatic system. Sulfide mineralization
does not form at this stage, although magnetiteand scheelite (in W skarns) do precipitate in thewaning stages of prograde metasomatism.
Retrograde – meteoric fluid influx and main
metal precipitation (Figure 2.21b)
All magmatic-hydrothermal systems undergo
progressive cooling and decay of the high temperature magmatic fluid system. As fluids become
progressively dominated by shallow meteoric
waters, a series of complex retrograde reactions
takes place, as well as the precipitation of themain stages of base and precious metal, sulfiderelated mineralization (Einaudi et al., 1981). The
retrograde alteration assemblages are superimposed onto earlier metamorphic and metasomatic
minerals and, typically, this process is recognizedby paragenetically late formation of epidote,biotite, chlorite, plagioclase, calcite, quartz (allafter various garnet types), tremolite–actinolite,
and talc (after pyroxenes) and serpentine (after
olivine). Sulfide ore minerals, as well as magnetite and hematite, occur as disseminations, orveins, that cut across prograde assemblages.
Assemblages such as pyrite–chalcopyrite–magnetite characterize proximal settings, whereas
bornite and sphalerite–galena are typically moredistal in occurrence. The paragenetically late precipitation of most skarn-related ores suggests thatmetal precipitation is related to decreasing temperature of the ore fluids (and a resulting drop insolubilities), fluid mixing, or neutralization of theore fluid by reaction with carbonate lithologies.

117

Mixing of the magmatic ore fluid with the latemeteoric component, and related redox reactionsin the fluid, may be additional controls on the ore
formation process. A more detailed description of
precipitation mechanisms in hydrothermal solutions is provided in Chapter 3.

2.11 NEAR-SURFACE MAGMATIC-HYDROTHERMAL
PROCESSES – THE “EPITHERMAL” FAMILY OF

AU–AG–(CU) DEPOSITS
Exploration for gold in the circum-Pacific region,
especially since the 1970s, has led to the discovery of a large number of world-class gold deposits
associated with either active or geologically recent volcanic environments. These deposits are
now regarded as an important and very prospective category of gold deposit type, termed epithermal deposits. The term “epithermal” is derived
from Lindgren’s (1933) classification of ore deposits and refers to those that formed at shallow
crustal levels (i.e. the epizone). Many studies ofthis ore-forming environment, and in particularcomparisons with active, modern analogues suchas the Taupo volcanic zone on the north island of
New Zealand, have shown that epithermal deposits
typically form at temperatures between 160 and
270 °C and pressures equivalent to depths ofbetween 50 and 1000 m (Cooke and Simmons,
2000; Hedenquist et al., 2000).
There are two contrasting styles of mineralization that are now recognized in epithermal deposits, and these are referred to as high-sulfidationand low-sulfidation types. These terms refer specifically to the oxidation state of sulfur in the ore
fluid, the chemistry and pH of which also relates
to the nature of alteration associated with each type.
Unfortunately, the geological literature contains a confusing plethora of synonyms for these
two types. Table 2.2 attempts to summarize theseterms and also illustrates the main characteristics of each. Note that the two terms do not
relate to the abundance of sulfur, as this is highlyvariable in each deposit type, but in some publications “high sulfur” and “low sulfur” have beenequated with high-sulfidation and low-sulfidation
respectively.

118

PART 1 IGNEOUS PROCESSES

Table 2.2 Characteristics of high- and low-sulfidation epithermal deposits
High-sulfidation
Oxidized sulfur species (SO2, SO42−, HSO4− )in ore fluid/vapor
Also referred to as
Gold–alunite, acid–sulfate, alunite–kaolinite
FluidsAcidic pH, probably saline initially, dominantly magmatic

Low-sulfidation
Reduced sulfur species (HS−, H2S) in ore fluid/vapor

Adularia–sericite, hotspring-related
Near-neutral pH, low salinity, gas-rich (CO2, H2S),
dominantly meteoric

Alteration assemblage
Advanced argillic (zonation: quartz–alunite–kaolinite–
illite–montmorillonite–chlorite)

Adularia–sericite (zonation: quartz/chalcedony–calcite–
adularia–sericite–chlorite)

Metal associations
Au–Cu (lesser Ag, Bi, Te)

Au–Ag (lesser As, Sb, Se, Hg)

High- and low-sulfidation epithermal deposits
can be viewed as end-members of processes related to fluid evolution and circulation in andaround volcanoes. High-sulfidation deposits occur
in proximal settings and are commonly foundwithin or close to the volcanic vent itself. The
fluids involved with mineralization are derived
directly from the magma as a product of vaporand fluid-saturation and are usually boiling in theore-forming environment. The fluids are very
acidic (pH of 1–3) and oxidized, carrying the oxidized S4+ or S6+ species as SO2, SO 42−, or HSO −4 insolution. As this fluid boils and SO2 and CO2
are partitioned into the vapor phase, the remaining liquid carries a surplus of H+ which makes
it very acidic (pH = 1; Hedenquist et al., 2000).
This acidic fluid is also capable of leaching mostof the major elements from the host volcanic
or volcano-sedimentary rocks through which it
circulates, resulting in vuggy textures and an
advanced argillic style of alteration (see Table 2.2
and also the more detailed account of hydrothermal alteration in Chapter 3). By contrast, lowsulfidation deposits are associated with fluidsthat are similar to those involved with hotspringsand other geothermal manifestations in areas ofenhanced heat flow. These fluids have equilibrated with their host rocks and generally comprise a dominantly meteoric component, although
it is likely that this will have been mixed with an

evolved magmatic fluid if active volcanism islocated nearby. Consequently, low-sulfidation
deposits may form within the volcanic edifice,especially during the waning stages of magmatic
activity when draw-down of meteoric fluids isperhaps more likely. More typically they may
form at locations that are somewhat removed fromthe focus of volcanism. The fluid involved is nearneutral and has low salinities, but as with highsulfidation environments, it is also likely to haveboiled in and around the zone of ore formation.
The characteristics of high- and low-sulfidation
deposits are shown schematically in Figure 2.22.
This diagram suggests that a spatial and genetic
link exists between the two types, but it should beemphasized that this may not be the case. Many
gold districts contain either low-sulfidation epithermal deposits (such as the major deposits in
Nevada, USA, including Round Mountain, Comstock Lode, Midas, and Sleeper) or high-sulfidation
deposits (such as many of the Andean deposits, including Yanacocha, Pierina, and El Indio–Tambo).
There is, however, an increasing number of cases
where spatial and genetic links are evident. One
area where both high- and low-sulfidation epithermal deposits occur in fairly close proximity to one
another is in Kyushu, the southernmost island ofJapan (Box 2.4). It is evident in Figure 2.22 that
there should also be a genetic link between porphyry Cu–(Mo–Au) deposits formed in the sub-

MAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2

119

Magmatic-hydrothermal system
Geothermal systemFne luid
ut
ra
liz
a

°

°

30

300

200

g
ilin

0°

°

tio

High sulfidation(Au, Cu)–acidic fluids–advanced argillicalteration

600

Low
sulfid(Au, A ation
g)
–Neu
t
–Adu ral fluids
laaltera ria, sericitetion

CO2, SO2, HCI
Crater lake

CO2,
H2S

Acid springs
solfataras

Bo

Hotsprings, mud pools
fumaroles

Volcano

n

ric
teo
Me

600°

r
te
wa

Porphyry Cu, (Mo, Au)

Saline magmatic fluid
Liquid flow
Vapor flow

Figure 2.22 The geological setting and characteristics of high-sulfidation and low-sulfidation epithermal deposits.
A genetic link between high-sulfidation epithermal Au–Cu and sub-volcanic porphyry type Cu–Au deposits is alsosuggested (after Hedenquist et al., 2000).

volcanic environment and high-sulfidation Au–Cu
deposits at surface. This link can be demonstrated
in the case of the Lepanto epithermal Au–Cu andthe Far Southeast porphyry Cu–Au deposits on
Luzon in the Phillipines. These two deposits are
adjacent to one another in late Pliocene rocks, withthe mineralization in both occurring over a brief
300 000 year time period at around 1.3 Ma (Arribas
et al., 1995). It is also pertin-ent to note that lowsulfidation Au–Ag mineraliza-tion is located within a few kilometers of the Lepanto–Far Southeast
system, indicating that spatial and genetic links
between high- and low-sulfidation epithermal
and porphyry styles of min-eralization are likely.
The fact that their coexistence is only rarely documented is, at least partially, therefore, a functionof erosion and lack of preservation.

2.11.1 Gold precipitation mechanisms inepithermal deposits
The detailed mechanisms of gold precipitationin high- and low-sulfidation deposits are complex
and vary depending on the geological setting andnature of fluids involved (Cooke and Simmons,
2000). The following section briefly considers
gold precipitation mechanisms as applicable toepithermal deposits, although the principles discussed here also apply to other gold deposit types.
The reader is encouraged to read the relevant
sections on metal solubility and precipitation
mechanisms in Chapter 3 (sections 3.4 and 3.5)
before continuing with this section.
Anionic species, or ligands, of chloride (Cl−) andsulfide (specifically HS−) are regarded as being

120

PART 1 IGNEOUS PROCESSES

Magmatic-hydrothermal processes in volcanic
environments: the Kasuga and Hishikari epithermal
Au–Ag deposits, Kyushu, Japan
Kyushu, the southernmost major island of the Japanesearc, is the country’s principal gold producing region. Late
Cenozoic volcanic activity in this region has given riseto several epithermal Au–Ag deposits, all related to either
extinct or waning magmatic-hydrothermal systems arising from this volcanism. Both styles of epithermal gold
mineralization are preserved on Kyushu, namely highsulfidation (or acid-sulfate) mineralization represented by
the Nansatsu type ores of which Kasuga is an example,
and low-sulfidation (or adularia–sericite) mineralization
represented by the very rich Hishikari deposit.
KasugaThe Nansatsu district in the southern portion of Kyushu

has been the site of calc–alkaline volcanism for the past
10 million years. Several deposits occur in the district,including Akeshi, Iwato, and Kasuga. Kasuga is a small
deposit that produces about 120 000 tons of ore annuallyat an average Au grade of about 3 g t−1. The ore body is
associated with a residual high-silica zone in andesites of
the Nansatsu Group, which overlies basement metasediments of the Cretaceous Shimanto Supergroup (Hedenquist
et al., 1994). Alteration in the open pit is characterized bya central quartz-rich zone in which the volcanic host rockhas been almost entirely leached of all elements except
Si (Figure 1a). The quartz body is surrounded by a zone ofadvanced argillic alteration (comprising alunite, dickite,
and kaolinite), which in turn grades out into a propylitic
zone (chlorite + illite). Ore minerals consist mainly of pyrite(a)Kasuga open pit

m

Se-S

Pr
Gravels

0100m(b) 400

Alteration:
Quartz bodyAl
AluniteSe-Sm Sericite-smectite
Dt-Se-Py Dickite-sericite-pyrite
Se-Ch
Sericite-chlorite
Pr
Propylitic

Al
Dt-Se
Py
Se-Ch

Quaternary
pyroclastic flows

Dt-Se-Py
Se-Ch

AlteredNansatsuGroup
volcanics

Pr

Shimanto Supergroup(basement)

Shishimano
dacites/mudstones etc.

Recent pyroclastic
flows(meters)

200
0
–200

Hishikari andesites

Yamada Zone

Honko-Sajin Zone

(0.98–1.62 Ma)(0.60–1.15 Ma)(0.66–1.10 Ma)
Shimanto Supergroup(65–70 Ma)

0

m

500

Figure 1 (a) Cross section through the Kasuga deposit showing the nature and geometry of the alteration
zoning (after Hedenquist et al., 1994). (b) Cross section through the Hishikari deposits illustrating theconcentration of mineralization at one particular elevation (after Izawa et al., 2001).

MAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2

121

the zonation in the alteration halo and the late oxidationof pre-existing assemblages (Hedenquist et al., 1994).
Hishikari

Figure 2 Gold and silver rich Hosen vein
photographed underground at the Hishikari mine,
Kyushu.
and enargite with later native sulfur, covellite, and goethite.
The Au–Ag (electrum) is largely contained within thequartz body in the center of the alteration halo, althoughconsiderable remobilization into late oxidized phases
also occurs. Mineralization is linked to the exsolution ofa metal charged fluid phase from the host andesite, withsubsequent segregation into vapor and liquid phases. Highly
acidic vapors are implicated in the leaching of volcanic
rocks to form the porous quartz body. The metal charged
fluid phase subsequently percolated through the porousrock and precipitated Cu, Ag, and Au. Progressive mixingof this fluid with meteoric waters was probably the main
precipitation mechanism, and was also responsible for

essential in order to solubilize gold in aqueous
solutions (Seward and Barnes, 1997). In reduced,
near-neutral pH, aqueous solutions Au is likely
to be transported as the Au(HS) −2 complex and thisis also likely to be the preferred medium for the
movement of gold in low-sulfidation environments.
By contrast, at higher temperatures (>300 °C) andfor solutions that are both more acidic and saline,
gold is preferentially transported as the Au(Cl) −2
complex, and this mode of transport probably
applies to high-sulfidation environments. Since
the mechanism whereby gold is transported could
be fundamentally different in the two epithermal
deposit types, it follows that the chemical andphysical controls that precipitate the gold fromthe hydrothermal fluids are also likely to differ.

The Hishikari deposit, discovered in 1981, is a very rich
low-sulfidation epithermal deposit containing about 260
tons of mineable gold (Izawa et al., 2001). It is locatedsome 20 km to the northwest of the recently active
Kirishima volcano in south-central Kyushu. The deposit
occurs as a series of vertical, en-echelon veins largely
within the Cretaceous Shimanto Supergroup basement, butalso extending short distances into overlying Pleistocene
andesites. The veins consist mainly of quartz and adularia,
with lesser smectite clay. Vein formation and mineralization is believed to be linked to volcanism that started
about 1.6 Ma and terminated at about 0.7 Ma. Vein related
adularia, however, has been dated at between 1.15 and0.60 Ma, suggesting that the geothermal system andmineralization commenced about 0.5 Myr after volcanism(Izawa et al., 2001). The spectacularly rich Hosen No. 1
vein (which averages over 3000 g t−1 Au and 2000 g t−1 Ag!)
records evidence of three phases of vein deposition. Eachconsists mainly of quartz, adularia, and electrum. Minor
pyrite and chalcopyrite also occur but only during the earlyphases of deposition. The high grade ores at Hishikari arerestricted to a specific elevation, extending over about
200 m of vertical extent (Figure 1b), and veins either terminate or become sub-economic above and below thisrange. This, together with fluid inclusion evidence, suggeststhat boiling of a dominantly meteoric fluid was responsible
for the precipitation of gold and silver from solution.

For low-sulfidation deposits gold precipitationis relatively straightforward and is linked to one,or both, of the two processes that characterize
fluid evolution in this environment, namely boiling and fluid mixing. Boiling of an ore fluid inthis case will result in loss of H2S to the vapor
phase, which causes destabilization of the Au(HS) −2
complex and precipitation of Au, as describedin equation [2.7] (after Cooke and Simmons,
2000):
Au(HS) 2− + H+ + 0.5H2 ↔ Au + 2H2S

[2.7]

Mixing of an oxidized meteoric water with thesame ore fluid would also have the effect of precipitating gold, as shown in equation [2.8]:

122PART 1 IGNEOUS PROCESSES

Au(HS) −2 + 8H2O ↔ Au + 2SO42− + 3H+ + 7.5H2[2.8]
Evidence for boiling-induced Au precipitation isprovided from modern geothermal systems that
exploit steam to drive electricity-generating turbines, such as at Broadlands in New Zealand(Cooke and Simmons, 2000). It is well known
in these power stations that the siliceous scalethat plates the inside of pipes and accompanies
the flashing of water to steam is often enrichedin both Au and Ag. In actual low-sulfidation
deposits the narrow vertical interval over whichvein-hosted mineralization occurs (such as at
Hishikari; Box 2.4) is another indication that boiling acts as a fundamental control on ore precipitation. Although it is clear that fluid mixing also
plays a role, in deposits such as Creede in Colorado,the evidence for a widespread role for this mechanism is less clear (Cooke and Simmons, 2000).
For high-sulfidation deposits the Au deposition
mechanisms are less well understood and more
complex, since it is feasible to transport gold, notonly as a Au(Cl) −2 complex, as suggested above, butalso as a bisulfide complex (Au(HS)) in fluids witha high sulfur activity. It is also considered possiblethat Au is transported together with Cu in thevapor phase (see section 2.4.2 above). If gold is transported as a bisulfide complex in high-sulfidation
environments then the precipitation mechanisms
are also likely to be related to boiling and fluid
mixing, as for low-sulfidation systems. If, on theother hand, gold is transported as a chloride complex, then boiling and extraction of oxidized sulfur
species (SO2 or SO 2−4 ) into the vapor phase will
have little effect on its stability. Fluid mixing,
between a hot, acidic and saline Au(Cl) −2 bearingore fluid and a cooler, neutral meteoric solution,
could, however, be an important precipitation
mechanism, since it will have the effect of increasing pH and decreasing salinity (by dilution)of the ore fluid. Studies at Lepanto, for example,
indicate that fluid mixing did occur with variations in the ratio of magmatic to meteoric fluids
ranging from 9:1 to 1:1 over the ore deposit(Hedenquist et al., 1998).
Arribas et al. (1995) have provided a modelfor the formation of high-sulfidation epithermal

systems that reflects the complexities related topossible variations of gold speciation in this environment. The model envisages two stages of ore
formation. The first involves degassing of hot,acid rich magmatic vapors that are responsible forintense leaching of the host rock to form the porous
vuggy quartz zone in the fumarolic conduit andthe advanced argillic alteration halo around it(Figure 2.23a). The vapor phases could also mix
with meteoric waters to form an acid sulfate fluidthat has a low Au solubility, but is also implicatedin the alteration process. The secondary porosityand permeability created during this alteration
process is considered to be a necessary preparatory
stage for the influx of later metal-bearing fluids.
Subsequent ore deposition can occur in one of two
ways. A hot, acidic, and saline ore fluid, carrying
gold as a Au(Cl) −2 complex, could be derived directlyfrom the subjacent magma and move directly up
into the alteration zone. Ore precipitation would
occur as a result of mixing and dilution of thisfluid by cooler meteoric waters (Figure 2.23b).
Alternatively, it is suggested that Cu and Au are
initially removed from the magma in the vapor
phase and that these metal charged gases mix
with heated ground waters circulating around the
intrusion to form a low salinity fluid in whichgold is transported as Au(HS). This fluid would thenprecipitate metals by boiling in the near surface
environment, or mixing with meteoric waters, orboth (Figure 2.23b).

2.12 THE ROLE OF HYDROTHERMAL FLUIDS IN
MINERALIZED MAFIC ROCKS

Much of this chapter has concentrated on therelationships between felsic magmas and thefluids that exsolve from them. As mentioned
at the beginning of this chapter, however, it isalso feasible for a mafic melt to exsolve a magmatic fluid phase, although the resultant massfraction of exsolved water may be lower than in
felsic melts. In mafic rocks the role of a magmatic
fluid in ore-forming processes is commonly overlooked, even though it is now apparent thatsuch fluids may be important in the concentration of metals such as Cu, Ni, and the platinum

MAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2

Alteration

Ore deposition

Fumaroles

(a)

123

(b1)

(b2)

0
Cool meteoric
water
300

°C

400

km

°C

1
Magmatic vapor(SO2, HCI)

2

Magmaticbrine

Ore deposition
by mixing withshallow
meteoric
water

500

°C
Alteration
Vuggy quartz
AluniteKaolinite
Sericite
Potassic

Acid magmaticbrine
transports
gold asAuCl2–

Ores
Alterationenvelope
Mixing produces
reduced, acid,low salinity, Mixing
water carrying
Au as AuHSund

Hot growater

Gas phasemetaltransport(as AuS, CuS)

Figure 2.23 Two stage model for the formation of high-sulfidation epithermal deposits (after Arribas et al., 1995).
(a) Initial stage where a dominantly magmatic vapor phase is responsible for leaching of the country rock anddevelopment of an advanced argillic alteration halo around the main fumarolic conduit. (b1) Ore deposition stage, inthis case where gold is transported as a chloride complex; and (b2) ore deposition stage where gold is transported as abisulfide complex.

group elements (PGE) in layered mafic intrusions(Mathez, 1989a, b).
The solubility of H2O in a melt is higher thanthat of CO2 and it follows, therefore, that thelatter will exsolve before the aqueous species. The
first fluids to form from a crystallizing mafic melt
are rich in CO2 (as well as CO in more reduced
environments), but they will evolve to more H2Orich compositions as crystallization progresses. In
addition the fluid will also contain appreciablechlorine (as HCl or FeCl2) and sulfur (as HS− or
SO 2−
4 depending on fO2). The change in fluid composition from CO2 to H2O-dominated occurs
quite rapidly in fluids derived from mafic melts
and is marked by the precipitation of C (graphite)from the fluid. These fluids are, therefore, unlikely,
at least initially, to be able to dissolve much
metal, but this would change as the proportion ofwater in the solutions increases.
Considerations of metal concentration in mafic
rocks, such as those discussed at some lengthin Chapter 1, assume that distribution patterns of
Cu, Ni, and PGE reflect crystal–melt or melt–
melt partitioning during progressive crystallizationor silicate–sulfide immiscibility. These assumptions presuppose that the ore deposits observed

have not re-equilibrated with a vapor phase andthat the metal distribution patterns are genuinely
magmatic and not overprinted by hydrothermal
processes. This is not always the case and mineralogical observations in serpentinites fromeastern Australia, for example, show that certainof the PGE (in particular Pd, Pt, and Rh) tend to bereadily redistributed by hydrothermal solutions,whereas others (specifically Os, Ir, and Ru) remain
unaffected by such processes (Yang and Seccombe,
1994). The rare cases where nuggets, overgrowths
and dendrites of PGE are found in placer deposits, and in laterite enrichments, also indicatethat under certain conditions these metals arelabile and can be put into solution even at low
temperatures (see Chapter 4).
In the case of the PGE it is apparent, with thepossible exception of palladium (Pd), that these
metals tend to exhibit low solubilities in magmatic-hydrothermal fluids at high temperaturesand under fairly reduced (low fO2) conditions. In
the case of Pd it appears that this metal is fairly
soluble in saline solutions at elevated temperatures (Sassani and Shock, 1990). In general, however, the presence of ligands such as Cl− enhances
PGE solubilities only slightly and significant

124PART 1 IGNEOUS PROCESSES

dissolution of these metals is only feasible underhighly oxidized, acidic conditions (Wood et al.,1992). Hydroxide (OH−) complexation is also
regarded as unlikely to contribute toward PGE
solubility in hydrothermal solutions, although it
may be important in surficial waters and PGE–
OH− complexes may be implicated in placer andlaterite deposits. Likewise, bisulfide (HS−) complexes result in very low (ppb) quantities of Pt
and Pd being transported in most geologicallypertinent solutions. These observations tend tosupport the conventional view that metal concentrations in mafic rocks have not been markedly
modified by hydrothermal processes. However,under conditions of very low fS2 (where the metal
and not the metal sulfide is stable) it appears that
PGE solubilities in a highly saline brine might behigh (102–103 ppb), mainly because PGE complexation with Cl− is not detrimentally affected by
the presence of sulfide ligands. Under such conditions, a fluid could play an important role in theredistribution of these metals.
2.12.1 The effects of a magmatic fluid on PGE
mineralization in the Bushveld Complex
In the Bushveld Complex the grades of platinumgroup element (PGE) mineralization and therelative proportions of these elements remain
uniform over tens of kilometers of strike, in both
the Merensky Reef and the UG2 chromitite seam(Cawthorn et al., 2002). By contrast the mineralogy of the PGE varies markedly. The normal stratiform reef is generally dominated by PGE sulfide
minerals such as cooperite, braggite, and laurite.
Atypical situations, represented by mineralized
potholes (see Box 1.6) and unusual discordantPt-rich dunite pipes, are characterized by PGE–Fe
alloys and tellurides. The general consensus isthat the sulfide-dominated mineralogy of normal
reefs is the product of magmatic processes, as discussed in Chapter 1, but that the sulfur-deficient
PGE mineralogy in potholes and pipes reflects
areas where a volatile-rich fluid phase has interacted with the magmatic ores, resulting in localized low fS2 conditions.
The Bushveld Complex contains several unusual rock types such as iron-rich ultramafic pipes,

as well as iron-rich pegmatites (with concentrations of Pb, As, Sb, and Bi) and plagioclase–
amphibole–phlogopite veins, that have been usedas evidence for vapor- or fluid-saturation in thelate stages of crystallization of the Bushveld magmas. Schiffries (1982) regarded the Fe- and Pt-rich
dunite pipes as metasomatic in origin and implicated an aqueous magmatic brine that reacted withthe country rocks at around 600 °C and 3.5 kbar.
More recent work has indicated that these bodies are probably discrete magmatic intrusions,but that they did act as channelways for later,lower temperature hydrothermal fluids (Cawthorn
et al., 2000). The sulfide-poor, Pt-dominated mineralogy of these bodies is considered to represent
re-equilibration of an original magmatic sulfide
assemblage by later hydrothermal fluids.
One of the still contentious issues concerning
the Bushveld Complex, however, is the extent towhich the major mineralized horizons, such as
the Merensky Reef and the UG2, owe their PGE
and base metal mineralization to hydrothermalrather than magmatic processes. Detailed mineralogical studies have shown that in certain environments magmatic fluids must have played a role
in redistribution of sulfur and recrystallization of
the PGE. Where the UG2 chromitite horizon iscut by a dunite pipe, for example, it contains thesame distinctive Pt–Fe alloy and Pt–arsenide(sperrylite) phases as the pipe itself. By contrast,
the UG2 horizon well removed from such bodies
comprises the normal, magmatic PGE–sulfide minerals such as braggite and cooperite (Peyerl, 1982).
In addition, the Merensky Reef itself often
exhibits the same sort of variations in PGE
mineralogy as does the UG2, but in this case theoccurrence of sulfide-deficient PGE minerals is
usually related to the formation of “potholes”
and not necessarily to cross-cutting dunite replacement pipes (Figure 2.24). The Merensky Reef
potholes were described in Chapter 1 (see Box 1.6)
and are attributed to syn-magmatic faulting of
the footwall cumulates just prior to injection of a
new magma pulse, with accompanying metasomatism of rocks within the pothole structure byhydrothermal fluid. The occurrence of desulfidized PGE mineral assemblages in the potholes is
consistent both with this process and with theMAGMATIC-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 2

Normal
Pt-Fe alloys + basemetal sulfides
Pt-Pd sulfides
Laurite Ru(Os,Ir)S2
Tellurides andCu–Au alloy

Normal reef

Figure 2.24 Schematic illustration
showing the nature of “potholes”
in the Merensky Reef of the
Bushveld Complex and thedistinctive mineral assemblagesof these features relative tonormal reef (after Kinloch, 1982;
Mathez, 1989b).

0.1

125

Contact Pothole
0

92.6

89.0

11.5

0

0.1
10.8

87.0
1.5

0.4
7.0

Pothole depression

Contact
reef
Pothole reefZone of fluid–rockinteraction

fact that potholes were probably also the sites of
fluid-saturation and enhanced circulation. The
presence of saline, but relatively oxidized, fluids
is considered to have been responsible for modifying the primary magmatic ore mineralogy, destabilizing the dominant base metal and PGE sulfide
phases, with the resultant loss of sulfur and reaction between PGE and Fe to form alloys or telluride and arsenide minerals (Kinloch, 1982). It is
important to note that, during the hydrothermal
overprint, PGE and base metals do not appear to
be redistributed and the metal budget, therefore,remains constant (Cawthorn et al., 2002). Thereseems little evidence for a solely hydrothermal
origin for the base and precious metal ores of theMerensky Reef, or any other major mineralizedpart of the Bushveld Complex, although it is clearthat magmatic fluids have reacted extensively
with the rocks of this layered mafic intrusion.
In summary, it would appear that in contrast
to granitoid-related systems, magmatic fluids inmafic melts have played a relatively small rolein the mineralization processes that accompany
the crystallization of these rock types. Vaporsaturation does occur but its influence as a mineralizing agent would generally appear to be limited
to re-equilibration of existing magmatic PGE andbase metal sulfides, in rocks where metasomatism is clearly demonstratable, such as potholes,
Fe-rich dunite pipes, pegmatites, and veins.

All magmas contain the constituents that, on
crystallization, combine to exsolve discrete fluidand vapor phases. Most magmas will exsolve substantial quantities (up to several wt%) of water, as

well as an order of magnitude or so less of carbon
dioxide, and these are the two dominant magmatic-hydrothermal fluids. Water in particular
has the ability to dissolve significant quantities of126PART 1 IGNEOUS PROCESSES

anionic substances, in particular Cl−, which inturn promotes the dissolution of other alkali andtransition metal cations. The magmatic aqueous
phase can exist as a liquid, vapor, or homogeneous
supercritical fluid. The process of H2O-saturation
can be achieved in two ways, either by decreasing
the pressure of the system (called first boiling) orby progressive crystallization of magma (second
boiling).
H2O-saturation is particularly relevant to oreforming processes during the emplacement andcrystallization of granitic magmas at moderate toshallow crustal levels. This environment gives
rise to the formation of a wide variety of important ore deposit types including porphyry Cu andMo deposits, polymetallic skarn ores, graniterelated Sn–W deposits, and the family of volcanicrelated epithermal Au–Ag–(Cu) deposit types.
Many metals will partition strongly into the liquid or vapor that forms on H2O-saturation and, insuch cases, mineralization accompanies the alteration of host rocks, both within and external to theintrusion. The formation of either Cu-dominantor Mo-dominant porphyry deposits reflects a

subtle interplay between the depth of intrusion ofa granitic body (itself a function of the original
water content), the timing of H2O-saturation relative to the progress of crystallization, and the
behavior of metals during melt–fluid partitioning.
Egress of fluids and vapor from the magma andtheir subsequent circulation are dependent on thepermeability of the surrounds, and may be modified by boiling-related hydrofracturing. Polymetallic skarn deposits reflect the interaction between
the exsolved magmatic fluids from different
types of granite and calcareous sediments. H2Osaturation and boiling in volcanic environments,
producing significant volumes of volatile-richvapor, is conducive to the formation of epithermal
deposits. High- and low-sulfidation epithermal
deposits reflect end-members in a continuum ofmagmatic-hydrothermal processes that progressively incorporate more non-magmatic waters as
the volcanic system wanes, or as one moves awayfrom the volcanic center. Many ore deposit types
are the product of fluids that are unrelated toa magmatic source and these are the subject ofChapter 3.

For those readers wishing to read more about magmatichydrothermal ore-forming processes, the following
references to books and journal special issues will be
useful.

istry of Epithermal Systems. El Paso, TX: Society of
Economic Geologists, 298 pp.
Reviews in Economic Geology, Volume 4: Whitney,
J.A. and Naldrett, A.J. (eds) (1989) Ore Deposition
Associated with Magmas. El Paso, TX: Society of
Economic Geologists, 250 pp.
Reviews in Economic Geology, Volume 13: Hagemann,
S.G. and Brown, P.E. (eds) (2000) Gold in 2000. El
Paso, TX: Society of Economic Geologists, 559 pp.
Seltmann, R., Lehmann, B., Lowenstern, J.B. andCandela, P.A. (eds) (1997) High-level silicic magmatism and related hydrothermal systems. Special issue,
Journal of Petrology, 38(12), 1617–807.
Taylor, R.P. and Strong, D.F. (eds) (1988) Recent Advancesin the Geology of Granite-related Mineral Deposits.
Canadian Institute of Mining and Metallurgy, Special
Volume 39, 445 pp.
Thompson, J.F.H. (ed.) (1995) Magmas, Fluids and Ore
Deposits. Mineralogical Association of Canada, Short
Course Handbook, 23, 525 pp.

Corbett, G.J. and Leach, T.M. (1998) Southwest Pacific
Rim Gold–Copper Systems: Structure, Alteration andMineralization. Special Publication 6. El Paso, TX:
Society of Economic Geologists, 237 pp.
Geochemistry of Hydrothermal Ore Deposits, ed. H.L.
Barnes. 1st edn, 1967. Holt, Rinehart and Winston
Inc., pp. 34–76. 2nd edn 1979. John Wiley and Sons,
pp. 71–136. 3rd edn, 1997. John Wiley and Sons,
pp. 63–124 and 737–96.
Misra, K.C. (2000) Understanding Mineral Deposits.
Dordrecht: Kluwer Academic Publishers, 845 pp.
Pirajno, F. (1992) Hydrothermal Mineral Deposits. New
York: Springer-Verlag, 709 pp.
Reviews in Economic Geology, Volume 2: Berger, B.R.
and Bethke, B.M. (eds) (1985) Geology and Geochem-

Hydrothermal
Processes

Hydrothermal ore-forming
processes

ORIGIN OF FLUIDS IN THE EARTH’S CRUSTDEFORMATION, PRESSURE GRADIENTS, AND HYDROTHERMALFLUID FLOWMETAL SOLUBILITIES IN AQUEOUS SOLUTIONS
the nature of metal–ligand complexes andPearson’s Principle
FLUID –ROCK INTERACTIONS AND ALTERATIONPRECIPITATION MECHANISMS
physico-chemical processes
adsorption
biologically mediated processes
METAL ZONING AND PARAGENETIC SEQUENCES
MODERN ANALOGUES OF HYDROTHERMAL ORE-FORMING
PROCESSES

the VMS–SEDEX continuum
ORE DEPOSITS ASSOCIATED WITH AQUEO-CARBONIC
HYDROTHERMAL FLUIDS

Box 3.1 Fluid mixing and metal precipitation:
the Olympic Dam iron oxide–copper–gold deposit,
South Australia
Box 3.2 Alteration and metal precipitation: the Golden
Mile, Kalgoorlie, Western Australia – an Archean
orogenic gold depositBox 3.3 Exhalative venting and “black smokers” on thesea floor: the Cyprus-type VMS deposits
Box 3.4 Sedimentary exhalative processes: the Red Dog
deposit, Alaska, USA
Box 3.5 Circulation of orogeny driven aqueo-carbonic
fluids: Twin Creeks – a Carlin type gold deposit,
Nevada, USA
Box 3.6 Circulation of sediment-hosted connate fluids.
1 The Central African Copperbelt
Box 3.7 Circulation of sediment-hosted connate fluids.
2 The Viburnum Trend, Missouri, USA

orogenic, Carlin-type, and quartz pebble
conglomerate hosted gold deposits
ORE DEPOSITS ASSOCIATED WITH CONNATE FLUIDS
stratiform sediment hosted copper (SSC) andMississippi valley type (MVT) lead–zinc deposits
ORE DEPOSITS ASSOCIATED WITH METEORIC FLUIDS
sandstone-hosted uranium deposits

3.1 INTRODUCTION
This chapter extends the concept of hydrothermal
mineralization to deposits related to fluids derived
from sources other than magmatic solutions.
Such fluids include those formed from metamorphic dehydration reactions, from the expulsion ofpore fluids during compaction of sediment, andfrom meteoric waters. It also considers sea water

as a hydrothermal fluid with specific reference tothe formation of base metal deposits on the ocean
floor. Unlike the previous chapter, which was
mainly concerned with granite-related ore deposits,
the present chapter discusses a much broader
range of ore-forming processes and environments.
Hydrothermal ore-forming processes are ubiquitous and there is scarcely an ore deposit anywhere on Earth that has not been formed directlyfrom hot aqueous solutions flowing through thecrust, or modified to varying degrees by such
fluids. This view is supported by the example
in section 2.12 of Chapter 2, where a compellingcase for hydrothermal precipitation of PGE mineralization in the potholes of the Merensky Reef,

130

PART 2 HYDROTHERMAL PROCESSES

Bushveld Complex, traditionally regarded as
igneous in origin, is presented. Likewise, the Au–U
ores of the Witwatersrand Basin can no longer
be regarded simply as paleoplacer deposits andhydrothermal processes have clearly played asignificant role in their formation (see section 3.9
below). Similarly, oil and gas deposits (see Chapter 5) have migrated to their present locations in
the presence of hot water, during processes akinto those discussed below for hydrothermal solutions. Many of the giant ore deposits of the world
owe their origins to the flow of hydrothermal
fluids in the Earth’s crust and the ability of aqueous solutions to effectively scavenge, transport,
and concentrate a wide range of economicallyimportant components.
Over the past few decades in particular, a great
deal of research has been directed toward better understanding the complexity of hydrothermal processes. Concepts such as the source ofhydrothermal solutions, their passage throughthe Earth’s crust, and the precipitation mechanisms involved in the formation of ore bodies are
now relatively well understood. The three editions of Barnes’s Geochemistry of Hydrothermal
Ore Deposits (1967, 1979a, 1997) provide an
account of the progress of this research over several decades. There are, of course, some features ofhydrothermal ores about which we understandvery little, and these include the ages and duration of ore-forming processes, the detailed recognition of ancient fluid pathways, and the depthsof fluid flow in the crust, as well as the relationship between global tectonics and metallogeny(Skinner, 1997). The last topic is a particularly
important one (see Chapter 6) that applies not
only to hydrothermal ores, but to the entire rangeof mineral deposit types. Finally, the role ofmicroorganisms in the formation of ore deposits
is a topic that is beginning to attract attention andmay turn out to be much more important thanpreviously thought.
It will be evident that the separation of“magmatic-hydrothermal” and “hydrothermal”
processes into two sections (Chapters 2 and 3
respectively) is not so much a conceptual necessity as a requirement of the organization andstructure of the book. The two chapters should beseen as covering a spectrum of processes ranging from magmatic fluid flow to shallow levelmeteoric infiltrations. In order to emphasize thecontinuum, this chapter is terminated with asummary diagram that attempts to relate thesources of hydrothermal solutions to ore deposit
types, and applies to both Chapters 2 and 3.

3.2 OTHER FLUIDS IN THE EARTH’S CRUST AND
THEIR ORIGINS

Figure 3.1a shows that, in addition to magmatic
fluids, there are four other major water types on ornear the Earth’s surface. Although they may all
have had similar origins, each of these fluid reservoirs is different in terms of its composition andtemperature and will, therefore, play different
roles in the formation of ore deposits. The majorwater types are defined as sea water, meteoric
water, connate water and metamorphic water,
listed typically in order of increasing depth (andtemperature) in the crust. A fifth fluid reservoir,
where waters are derived from a mixing of two or
more other water types, is also described below,
specifically because mixed fluids can be very
important in certain ore-forming environments.
In their present settings each of these fluids
can be identified because the environment fromwhich the fluid came is known. In the geologicalpast, however, where only indirect manifestations of fluid flow are apparent, it is much moredifficult to determine the origin of a particularwater type. Fortunately, the hydrogen and oxygenisotope characteristics of water are reasonably
diagnostic of its source and can be used to deduce
the origins of an ancient reservoir. Figure 3.1b is
a plot of δD versus δ18O that shows trends andfields which serve to fingerprint the major fluid
reservoirs on or near the Earth’s surface. The
origins of the various water types and their stable
isotopic characteristics are briefly described below.
3.2.1 Sea water
The oceans collectively represent the largest fluid
reservoir on the Earth’s surface (Figure 3.2), covering some 70% of the surface and containing about

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

131

(a)
Sedimentary
basin

Ocean

Meteoric

ConnateSea water
Metamorphic
Magmatic
Chapter 2

Metamorphosed rocks

Igneousintrusion
(b)

+1

0)

SMOW

lf
Gu

La
titu
ric High dinalswa
ter L hifts
lin ow
e(
δD
=

–40

ast

Co

in

Metamorphic
water(300–600°C)

as

b
an

ig

ch

Mi

teo

B

–80

Me

δD (% )

Figure 3.1 (a) The major types ofliquid water that exist at or near the
Earth’s surface; magmatic water
was specifically discussed in
Chapter 2, while the other fluids
are discussed in this chapter.
(b) Plot of hydrogen (δD permil)and oxygen (δ18O permil) isotopic
ratios for various water types(after Taylor, 1997). Standard Mean
Ocean Water (SMOW) is defined
to be zero for both δD and δ18O.
Some ore-forming environments
are clearly related to mixed fluid
reservoirs; in the cases shown(see A and B) mixing of meteoricand connate fluids has taken place.

8×

δ 18
O

0

ta

A

Primarymagmatic water

sin

ba

r
be

Al

A Steamboat Springs, Nevada–120

B Salton Sea
Connate fluids–20

98% of its free water. As mentioned in Chapter 2,
the Earth’s surface has been covered by substantial volumes of water since very early in Earth history. Sea water is relatively well mixed at a global
scale, and is weakly saline because of reaction
with both continental and oceanic erosion products over time. The principal dissolved constituents in sea water are the cations Na+, K+,
Ca2+, and Mg2+ and anions Cl−, HCO −3 , and SO42− ,which typically occur at a total concentration(or salinity) of around 35 g of solids per kg of sea
water (3.5 wt%).

–100
δ18O (% )1020

Sea water is extensively circulated through theoceanic crust and is responsible for widespread
alteration and metal redistribution in this portionof the Earth’s crust. The drawdown of sea water
into major faults associated with the mid-ocean
ridges and its subsequent emergence from exhalative vents or “black smokers” is a major oceanographic discovery that has revolutionized theunderstanding of volcanogenic massive sulfide(VMS) deposits (see section 3.8 below). VMS
deposits occur in many different parts of theworld, and in rocks of all ages, confirming the132

PART 2 HYDROTHERMAL PROCESSES

Precipitation
0.107
Ice (43.4)

Precipitation
0.398

Lakes/rivers(0.13)

Groundwater(15.3)

Mid-oceanridge (blacksmoker)Ocean(1400)
ff

f

f

Figure 3.2 Simplified diagram illustrating the water budget on or close to the surface of the Earth (after Berner andBerner, 1996). Precipitation fluxes are in 106 km3 yr−1 and reservoir volumes in 106 km3.

Table 3.1 Representative concentrations (in ppm) of major ionic species in natural meteoric waters
Species (ppm)

Sea water

Black smoker sea water(at 100–300 °C)

Rain water

Meteoric springwater (at 6 °C)

Meteoric geothermal
water (at 300 °C)

Na+
Cl −
SO42−
Mg2+
Ca2+
K+
Si4+
Metals (Fe, Cu, Zn, Mn, etc.)

10 000
20 000
2700
1300
410
400
0.5–10depleted

6000–14 000
15 000–25 000
0
0
36
26
20
enriched

1.0
1.1
1.5
0.2
0.4
0.5
1.2
depleted

23
3.1
11
2.4
5.1
1.0
17
depleted

187
21
103
0
0.5
27
780
enriched

Source: after data in Krauskopf and Bird (1995), Scott (1997).

importance of sea water as a hydrothermal fluid
source. Table 3.1 illustrates the concentrations ofmajor ionic species in sea water and compares thisto fluids exhaled from black smoker vents, where
significant concentrations of metals have takenplace due to interaction between hot sea waterand oceanic crust. Sea water is also compared torain water and meteoric derived groundwaters,
which again have higher dissolved solute contents at higher temperatures.
3.2.2 Meteoric water
Meteoric water has its immediate origin within
the hydrological cycle and has, therefore, been incontact with the atmosphere. In a geological context it refers to groundwater that has infiltrated
into the upper crust, through either rainfall or
seepage from standing or flowing surface water.
In this sense sea water infiltrating into the crustshould also be regarded as a source of meteoric
groundwater. Groundwater is the second largest
reservoir of liquid water (Figure 3.2) and generally
exists close to the surface in the interstitial porespaces of rocks and soil. It does not refer to thewater that forms part of the crystal structure ofhydrous minerals, nor does it refer to the micrometer-sized fluid inclusions that occur withinmany of the rock-forming minerals of both thecrust and mantle. Meteoric fluid can neverthelessHYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

penetrate along fractures to deep levels in thecrust and is, therefore, involved in widespread
circulation throughout the crustal regime. It is
responsible for the formation of many different
types of hydrothermal ore deposits, but especiallythose characterized by relatively low temperaturetransport and precipitation, such as sandstonehosted and surficial uranium ores (see section
3.11 and Chapter 4).
The oxygen and hydrogen isotope compositionsof meteoric waters, both on the surface and as
groundwater, vary systematically over the entire
globe as a function of latitude and elevation. The
linear relationship between δD and δ18O values(the meteoric water line in Figure 3.1b) exists
because deuterium and hydrogen, as well as 18O
and 16O, are systematically fractionated between
liquid water and water vapor during the evaporation–condensation processes of the hydrologicalcycle, and this fractionation increases as a function of temperature (Craig, 1961). Thus, most
meteoric waters have δD and δ18O values thatvary along a straight line (δD = 8 × δ18O + 10;
Figure 3.1b) and which help to broadly identify
the climatic conditions or latitude from where
the water came.
3.2.3 Connate water
Water that is included within the interstitial porespaces of sediment as it is deposited is referred toas connate or formational water. Originally thiswater is either meteoric or sea water, but it undergoes substantial modification as the sedimentis buried, compacted, and lithified. The various
stages of diagenesis that result in the transformation from uncompacted particles of sediment tolithified sedimentary rock produce aqueous solutions that evolve with time and depth. Such fluids
invariably move through the sequence and areoften involved in the formation of ore deposits.
The progressive burial of sediment to depths ofaround 300 meters results in a rapid reductionof porosity and the initial production of a substantial volume of water. Shales are initially very
porous and the early stages of burial can result in
the production of up to 3500 litres of water foreach cubic meter of sediment deposited (Hanor,

133

1979). Over 75% of the interstitial pore fluid islikely to be expelled from a shale by the time itis buried to 300 meters depth and its porosity
will, accordingly, decrease rapidly at this earlystage. By contrast, uncompacted sandstones are
initially less porous than shales and will release
only 700 litres of water for each cubic meterof sediment deposited (Hanor, 1979). Calculated
average rates of water release for shales and sands
are shown as a function of depth of burial in
Figure 3.3a.
An additional source of formation water is the“structural” or “bound” water that occurs either
as loosely bound H2O or OH− molecules withinclay mineral particles in argillaceous sediments.
Such water is able to leave the host mineral once
temperatures of 50–100 °C have been reached.
This is shown as the episodic peak-like increases
in the rate of release of water from the sedimentin the depth profiles of Figure 3.3a. The natureand volume of bound water release will obviously
vary from one situation to the next and depend on
factors such as the local geothermal gradient andthe type and proportions of clay host minerals.
The two main stages of connate fluid production(i.e. pore fluid and bound water) are also schematically illustrated in Figure 3.4, where they are compared to the production of metamorphic fluidsthat occurs at somewhat higher temperatures andgreater burial depths.
The temperature of connate fluids increases
with depth in the sedimentary sequence, with theexact rate of increase being a function of the local
geothermal gradient. The latter varies typicallybetween 15 and 40 °C km−1. Fluid pressures will
also increase with depth, although the nature ofpressure variation with progressive burial is likely
to be complex and depend on the amount of pore
fluid present in any given sedimentary unit andthe interplay between hydrostatic and lithostaticpressure gradients. A more detailed descriptionof the relationships between hydrostatic andlithostatic pressure is presented in section 3.3.2
below, where its significance with respect to the
movement of hydrothermal fluids in sedimentary
sequences is discussed. Figure 3.3b illustrates theway fluid pressure varies with depth in the Gulf
Coast sediments of the southern USA and points

134

(a)

PART 2 HYDROTHERMAL PROCESSES

(b)

0

1

0

1
ti
sta
ho
Lit

Release of interstitialwater on compaction

Depth (km)

ure

3

nt

ress

tic p

3

ie
ad
gr

rosta

Depth (km)2
Episodic release of“bound” water from clays

e
ur
ss
re
cp

Hyd

2

ient

grad

4

Gulf
Basin

4

Shales
Sands

5
0.01

(d)

0

2

2

Gulf
Basin Alberta

1000

Depth (km)

1

3

500
Pressure (bars)

0

0

1

Depth (km)

(c)

5

0.1
1
10
Rate of water release (l m–1 of burial)

3
Michigan

Alberta44Michigan

ShalesGulf Basin
Sands

5

1.1
1.2
Density (g cm–3)

1.3

5

100

200
300
Salinity (% )

400

500

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

135

to significant deviations above the hydrostatic
gradient (i.e “overpressuring”) caused by low permeability horizons within the sequence which
impede the expulsion of fluids so that pore watersupports the weight of the overlying strata andhigh fluid pressures result (Hanor, 1979). Zones ofoverpressured fluids dictate the nature of fluid
flow and mass/heat transfer, which has important
implications for the migration of oil brines (seeChapter 5) and the formation of a variety of
sediment-hosted hydrothermal ore deposits.
Connate fluids also undergo increases in density and salinity with depth, as illustrated in Figure 3.3c and d. The density increases are related toincreases in pressure and salinity, although there
is a limit to such a trend since temperature is also
augmented and this has an inverse effect on density. Deep basinal waters are commonly saline tothe extent that they are generally unpotable. The
increase in salinity is sometimes related to interaction of connate waters with evaporitic horizons
which contain easily dissolvable minerals such as
halite, sylvite, gypsum, and anhydrite. Salinity
increases are also, however, apparent in sedimentary sequences which do not contain evaporites
and in such cases the process of “membrane-” or“salt-filtration” probably applies. In argillaceous
layers, where clay particles are strongly compacted, the interaction of the net negative charge(caused by atomic substitutions within the crystal lattice) around each particle creates what is
termed a “Gouy layer” (Berner, 1971). The latter
acts as a semipermeable membrane or filter
which serves to reject anions in solutions passing
through the shale layer. Since anions are excluded
from the filtrate then so too must be cations, inorder to maintain electroneutrality. The result isthat any solution passing through the compacted

clay will emerge with a lower salinity. Upwardly
migrating brines will, therefore, become less saline, whereas the dissolved salt content of deeper
residual waters will increase. Although doubts
have been expressed about the efficacy of forcing
solutions upwards in a sedimentary sequence
across semipermeable shale layers, the processof salt filtration is generally regarded as a feasible
explanation for the widespread existence of subsurface connate brines. The topic is of considerable importance to ore genesis, as the availabilityof ligands in hydrothermal fluids, especially Cl−,is crucial to increasing the solubility and transport efficiency of metals in solution (see Chapter
2 and section 3.4 below).
The salinity characteristics of connate waters
will vary from one sedimentary basin to another,as seen in Figure 3.3d. Likewise, the δD andδ18O characteristics of such fluids will also varydepending on the climatic characteristics and latitude of the host basin. Since the original trappedpore fluids in most sedimentary sequences will beeither meteoric or sea water, their hydrogen andoxygen isotope compositions will initially reflect
either the appropriate point along the meteoric
water line or SMOW (Figure 3.1b). As the fluids
evolve, through both temperature increases andinteraction with the host rocks, their stable isotope characteristics will also change and define
trends that typically deviate from a position on
the meteoric water line along paths of shallower
slope. Fluids will tend to reflect higher δ18O
values as their temperatures and salinities increase in sedimentary sequences (Taylor, 1997).
In Figure 3.1b, for example, the low latitude Gulf
Coast formational waters define a trend that isquite distinct from the fluids in the high latitude
Alberta Basin of Canada.

Figure 3.3 (Opposite) The formation andcharacteristics of connate waters. (a) Depth profile
illustrating the rate of water release due to initialcompaction and porosity reduction followed bydehydration of clay minerals such as montmorillonite.
(b, c, and d) Depth profiles showing typical pressures,
fluid densities, and fluid salinities, respectively, in acompacting sediment pile. Different profiles refer tothe Michigan, Alberta, and Gulf basins in the USA(after Hanor, 1979).

3.2.4 Metamorphic water
As rocks are progressively buried and temperatures
exceed about 200 °C, the process of diagenesis
evolves to one of metamorphism. Metamorphism
is a multifaceted process but essentially involves
the transformation of one mineral or mineral
assemblage to another whose stability is more in
equilibrium with the prevailing conditions at

136PART 2 HYDROTHERMAL PROCESSES

7

5
4

garnetbiotite}6

Biotite + quartz

8

Kaolinite + quartz

3“Bound” water produced
during breakdown of clays

2}

Proportion of water released (wt%)

9Muscovite + chlorite

Pore fluid produced
during compaction

10

Metamorphic fluid

pyrophyllite

Connate fluid

1
0
0

100

200

300

400

T (°C)

higher pressure and temperature. Of importancein this discussion is the transformation ofhydrous silicate and carbonate minerals to newly
formed assemblages that contain lesser volatiles
than before. Dehydration and decarbonation
reactions during prograde metamorphism are,therefore, very important processes that produce
substantial volumes of metamorphic water inthe mid- and lower-crust. An example from the
Witwatersrand Basin in South Africa (Figure 3.4)
shows that at around 300 °C a metamorphic
reaction involving the breakdown of kaolinite topyrophyllite will result in the production of metamorphic water. The latter arises from the factthat kaolinite contains more water than does
pyrophyllite so that the excess water liberated by
the phase transformation has to be expelled into
the sedimentary sequence. This reaction is followed at around 400 °C by the transformation of
muscovite and chlorite to biotite, when more
fluid will be produced. Both H2O and CO2 (thelatter specifically when the breakdown reactions
involve carbonate minerals), as well as volatiles
such as CH4 and sulfur species, are produced asmetamorphic fluids during prograde reactions of
this type.

500

600

Figure 3.4 The relationshipbetween connate fluid production
during diagenesis and metamorphic
fluid production due to systematic
dehydration reactions involving
the breakdown of one mineral
assemblage to another containing
less water. The conditions
pertaining here were calculated
specifically for the WitwatersrandBasin (after Stevens et al., 1997).

Fluids produced in low to medium grade metamorphic terranes are dominated compositionallyby H2O, CO2, and CH4 in approximately thatorder of abundance. High grade rocks tend to be
dominated by dense CO2 with lesser amounts ofassociated water and methane. Most metamorphic fluids have low salinities and low concentrations of reduced sulfur. Fluids of this nature areglobally implicated in the formation of orogenic
gold deposits, which represent a very important
category of gold deposit types (Phillips et al.,
1994; see Box 3.2). Although it has not been
mentioned before, it is clear that CO2 may be
an important component of hydrothermal solutions, not only in those originating through metamorphic devolatilization but also in magmaticwaters.
Carbon dioxide is, after water, the most abundant component in hydrothermal solutions and its
phase equilibria relative to those of water are
briefly considered below. It is a larger moleculethan water and is non-polar (see Chapter 2), which
accounts for its lower melting and critical points.
It is of little importance as a solvent other thanfor non-anionic species such as the hydrocarbons,especially CH4.

137

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

A part of the phase diagram for pure CO2 isshown in Figure 3.5a and when compared to asimilar diagram for H2O (Figure 2.2c) is seen tohave remarkably similar densities and phase transition characteristics to water, except that it has

lower melting (−56.6 °C) and critical (31.1 °C)
points. At high temperatures H2O and CO2 arecompletely miscible and form a single fluid phase
where the one compound is dissolved in the other.
At lower temperatures H2O and CO2 become

CO2

(a)
400

00

95

0.

1.

05
1.

1.1
0

300

1.1
5

Pressure (bars)

Isochore

85

0.

200
Triple
point

100

5
0.7 0.65
0.55

Criticalpoint

Liquid

0.45

0.15

Vapor

0
–60

–40

–20

0
20
Temperature (°C)

40

0.10 0.02

60

H2O–CO2(b)
600

H2O–CO2 with 6%
NaCl at 0.5 kb
500

H2O–CO2 with 6%
NaCl at 1.5 kb

H2O–CO2 with
2.6% NaCl

Figure 3.5 (a) Part of the phase
diagram for CO2 showing thevariations in fluid density (i.e.
isochores in g cm−3) as a functionof pressure and temperature. The
diagram is similar to that for H2O(see Figure 2.2c) except for the
lower melting and critical points.
(b) Various solvus curves for thesystem H2O–CO2; the solvus (i.e.
the curve which defines where in
P–T–X space H2O and CO2 unmix)
moves to higher temperatureswith decreasing pressure andincreasing salinity of the mixture(compilations after Brown, 1998).

Temperature (°C)

400
A

300
One phase (miscible)
Two phase (immiscible)

200

Pure H2O–CO2
at 2 kb

100

0

0

20

40

60

Mole % CO2

80

100

138

PART 2 HYDROTHERMAL PROCESSES

meteoric and connate waters, but which may
overlap to a certain extent with magmatic waters(Figure 3.1b). The reasons for this are complex
and related to the variable nature of metamorphic
protoliths and the ubiquitous effects of fluid–rock interaction during regional metamorphism(Taylor, 1997).
3.2.5 Waters of mixed origin

Figure 3.5 (cont’d) (c) Microphotograph showing
the typical appearance of a H2O–CO2-bearing fluidinclusion. Prior to entrapment this fluid was ahomogeneous mixture of mutually soluble H2O and
CO2. On cooling it subsequently exsolved to form
immiscible fractions of H2O (outer fluid wetting thewalls of the inclusion) and CO2 (inner globule of fluid).
A vapor bubble occurs within the inner CO2 globule.

immiscible (like oil and water) and exist as two
separate phases (usually as H2O liquid and CO2
supercritical fluid). This relationship is shown in
Figure 3.5b, where the solvus (i.e. the curve which
defines where unmixing of the single phase mixture occurs) for pure H2O–CO2 at 2 kbar identifies
a plateau at around 250–300 °C, above which thetwo compounds occur as a single miscible phase.
As the temperature falls the single phase fluid
progressively unmixes until at room temperature
H2O and CO2 are essentially immiscible. Thisphenomenon is clearly illustrated in fluid inclusions that have trapped single phase H2O–CO2
mixtures above the relevant solvus. When viewed
in a laboratory at room temperature, these fluidinclusions are typically characterized by a socalled “double bubble.” The larger “bubble” isactually a globule of immiscible CO2 fluid withinliquid H2O, whereas the smaller inner “bubble”
comprises a mixture of H2O and CO2 vapor(Figure 3.5c). The widespread occurrence of H2Oand CO2-bearing fluid inclusions in the rockrecord attests to the ubiquity of aqueo-carbonic
fluids especially in the deep crust.
Metamorphic waters are characterized by relatively high δ18O values (+10 to +30 permil) which
are typically quite distinct compared to both

Although the previous discussion has emphasized
the nature and characteristics of different individual fluid reservoirs in the Earth’s crust, it isunreasonable to expect that in nature such waters
would necessarily retain their integrity withtime. In the upper crust especially, it is expectedthat fluids will undergo mixing and effectively
become hybrids with more than just a single
source or origin. It is apparent that the mixing ofore-bearing hydrothermal fluids may be an important consideration in the precipitation of metals
from such solutions. Magmatic and meteoric
waters, for example, often undergo mixing at the
volcanic to subvolcanic crustal levels associatedwith porphyry copper and epithermal Au–Ag oreforming environments. Mixing of connate andmeteoric waters is also a process known to characterize sedimentary basins forming in rapidly
filled continental rift environments. The major
effect of fluid mixing is to promote the precipitation of metals from ore-forming solutions and, consequently, this topic is discussed in more detail in
section 3.5.1 below.

3.3 THE MOVEMENT OF HYDROTHERMAL FLUIDS IN
THE EARTH’S CRUSTIn order to be effective as a mineralizing agent,hydrothermal fluids need to circulate throughthe Earth’s crust. The main reason for this is thatthey need to interact with large volumes of rockin order to dissolve and transport the metals
required to form hydrothermal ore deposits. The
flow of an ore fluid should preferably be focusedso that the dissolved constituents can be concentrated into an accessible portion of the Earth’s
crust that has dimensions consistent with thoseHYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

of a potential ore body. The study of hydrothermal
fluid movement has received a great deal of impetus from the oil and gas industry, where it isparticularly important to understand the pathways of hydrocarbon-bearing brines in sedimentary
sequences in order to explore for these types ofdeposits (see Chapter 5, section 5.4). Exploration
for metallic hydrothermal ores has also benefited
over the past few years from a better understanding of fluid flow in the Earth’s crust, especiallywith respect to Archean lode gold ores andMississippi Valley type Pb–Zn deposits (see Boxes
3.2 and 3.7 respectively).
In Chapter 2 the nature and duration ofhydrothermal fluid flow associated with magmatic intrusions was briefly discussed. The present section looks at hydrothermal fluid flow in
the Earth’s crust in more general terms and also
at a variety of scales. It emphasizes the role that
deformation and the resulting structural features
play, in terms of both movement and focusingof fluid flow in the crust. It is these features thatare of particular importance to the formation andlocation of metallic hydrothermal ores.
3.3.1 Factors affecting fluid flow at a crustal scale
The question of how large volumes of fluid canmove around at deep levels in the Earth’s crust,
where rocks are highly compacted and have low
intrinsic permeabilities, is one that has long taxed
researchers in a variety of disciplines. Movementof fluid is typically a response to either a thermal
or a pressure gradient in the Earth’s crust. The latter, in particular, is often related to deformation
which causes both stress and strain to vary in
the rocks being deformed. Deformation is widely
regarded as having played a major role in controlling fluid flow throughout the crust (Sibson et al.,
1975; Oliver, 1996) and its effects are discussed
in more detail in section 3.3.3 below. Since
crustal deformation is invariably linked to global
tectonic processes it is clear why the latter are so
intimately related to the formation of hydrothermal ores (see Chapter 6).
Another question that is pertinent to understanding the nature of fluid movement in the
Earth’s crust is that of pervasive as opposed to139channelized flow. The two scenarios can be
described in terms of “intrinsic permeability,”
where fluids infiltrate pervasively along grainboundaries and microcracks in a rock, and“hydraulic permeability,” where fluid flows along
major cracks that are sufficiently interconnected
to allow a distinct fluid channelway to develop(Oliver, 1996). Pervasive fluid flow can take placein porous rocks at relatively shallow levels in thecrust, a situation that allows convective fluid
movement to occur. Figure 3.6 illustrates a variety of tectonic scenarios at a crustal scale and thenature and distribution of major fluid flow vectors associated with each. At relatively shallow
crustal levels pervasive fluid flow, along a porous
aquifer, for example, often occurs in response to a
gravity-driven hydraulic head formed by uplifted
topography (Figure 3.6a). This is believed to be thedominant type of groundwater flow in continental regions and flow rates of between 1 and10 m yr−1 could develop depending on the permeability of the aquifer (Garven and Raffensperger,
1997). A similar style of long-range fluid flow
is now known to be associated with periods oforogenic compression which effectively squeeze
the fluid out of one package of rock into channelways (thrusts?) or high permeability aquifers inanother (Figure 3.6b). Orogeny-driven fluids havebeen implicated in the formation of the major
Mississippi Valley type Pb–Zn ore province of thesoutheast USA (see Box 3.7).
In oceanic crust fluids are believed to flow in
response to thermal gradients formed becauseof the high heat flow that characterizes the midocean ridges (Figure 3.6c). Sea water is forced
down into the crust along faults and, althoughessentially channelized, flows largely in response
to convection, reappearing close to the zones ofmaximum heat flow. Thermally driven convective fluid flow can occur, as mentioned in Chapter 2, in close proximity to high level magmatic
intrusions. It can also occur in deep intracratonic
rift basins where high heat flow or fluid density
gradients result in convective flow, but only inaquifers where permeabilities are sufficiently high(Figure 3.6d).
Large-scale flow of fluid at a crustal scale canalso occur in association with the dilatancy (i.e.

140PART 2 HYDROTHERMAL PROCESSES

Maximum flow rate: 1–10 m yr–1(a) Gravity driven
Hydraulichead
Upliftedregion

2
200

0
km

Maximum flow rate: 0.1–1 m yr–1(b) Orogeny driven

Thrust terraneCompression

2
0

20
km

(c) Thermally driven
Ocean floor

Mid-oceanridge

Sea water f
2–3°C

Exhalative
vents
300°C

f

2

f

f
0

100
km

f

350°C isotherm(d) Thermally drivenfMaximum flow rate: 0.1–1 m yr–1

Intracratonic basin or rift

2

200

0
km

Maximum flow rate: >10 m yr–1(e) Dilatancy/fault driven

5

Seismic pumping

Normal fault

Earthquake
focus

5

0
km

Figure 3.6 Various tectonic
scenarios illustrating themechanisms by which major fluid
movements in the Earth’s crusttake place. (a) Gravity-driven fluid
flow in response to the creation of ahydrostatic head in an area of uplift.
(b) Orogeny-driven fluid flow in
response to rock compaction in afold and thrust belt. (c) Thermally
driven fluid flow in the ocean
crust in response to high heat flowat mid-ocean ridges. (d) Thermally
driven fluid flow in a permeable
unit within a basin or rift.
(e) Dilatancy- or fault-driven
fluid flow along major, seismically
active structural features byseismic pumping and fault valve
mechanisms (modified after
Garven and Raffensperger, 1997).

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

141

SS

1200

Depth (m)

SH

SS

25

SH Hydrostatic

a

KP

3600

Pa

“Overpressure” dueto compaction
disequilibrium

–1

m

SS
4800

0

change in volume) of a rock mass that accompanies faulting and seismic rupture (Figure 3.6e).
Orogeny-driven fluids may be partially associatedwith this process, but more typically dilatancy
accompanies faulting and shearing and this particular process is considered in more detail in
section 3.3.3 below.
3.3.2 A note on hydrostatic versus lithostaticpressure gradients
As mentioned previously, one of the main factors
which controls fluid migration pathways is theexistence of a pressure differential in the hostrock. In a completely dry rock the pressure at any
depth is effectively provided by the weight of themass of rock above that point and is known as
lithostatic pressure. In Figure 3.7 the lithostaticpressure gradient is demonstrated by the line ofslope 25 kPa m−1, and reflects the average density

–1

m

10 K

Figure 3.7 Plot of depth versus
fluid pressure to illustrate thedifference between hydrostaticand lithostatic pressure gradients.
The dashed line shows theexcesses in hydrostatic pressure (i.e.
overpressure) that can accumulate
in low permeability shale horizons(shown as SH relative to sandstone(SS) layers in the column on the left
hand side) because of the inabilityof these rocks to efficiently drain
off pore waters on compaction(after Hunt, 1979).

Lithostatic

2400

300

600

900

1200

Fluid pressure (kg cm–2)(i.e. about 2.5 g cm−3) of the rock sequence overthat interval. The distribution of stress (i.e. force
per unit area) in a rock at any point along a depth
profile at lithostatic gradients will vary withorientation, and will be maximized in a verticalsense and minimized horizontally. By contrast,
the pressure on a rock located at the bottom of theocean will be given by the hydrostatic pressure
gradient (Figure 3.7), which has a slope of about
10 kPa m−1, reflecting the lower density of water
relative to rock (i.e. about 1.0 g cm−3). Stresses
under hydrostatic pressure are distributed equally in all directions, which accounts for the
relative “weightlessness” of objects immersed
in water.
Uncompacted sediment on the ocean floor containing abundant pore water that is still in directcontact with the overlying ocean water column
will also be pressurized along the hydrostatic gradient. As the sediment is buried and compaction

142

PART 2 HYDROTHERMAL PROCESSES

proceeds, water will be driven off. If, however,the pore water remains interconnected then theapplied load on that rock will still be carried by
the water and pressures will be made up of partlyhydrostatic and partly lithostatic components andhave values intermediate to the two gradients.
When drainage of pore fluid from the sediment isgood then a condition known as compaction equilibrium will accompany the progressive burial
and lithification of that rock. If, however, theremoval of water is impeded by low permeability
then compaction will likewise be retarded (andporosity maintained at a higher value) and pressures will increase to values above hydrostatic, acondition known as “overpressuring.” Fine andcoarse sediment will expel pore waters at different rates during burial, and will therefore compact along different pressure gradients. Fluidpressures will usually be significantly higher inless permeable rock units such as shales relative
to well drained rocks such as sandstones. Figure 3.7 illustrates this effect with respect to asequence of alternating sandstone and shale andplots the overpressures that occur in compacting
shale horizons relative to sandstones which maintain compaction equilibrium and hydrostaticpressures. Eventually, however, with increasing
depth in normal crustal profiles pore fluid pressures increase from hydrostatic to lithostatic (sothat fluid pressure equals rock pressure) and thisoccurs at between 5 and 10 km depth dependingon the nature of the rocks. The nature of hydrostatic and lithostatic pressure gradients in the
Earth’s crust, and the deviations from these endmember scenarios, are critical to understanding
the nature of fluid flow in all rock types.
3.3.3 Deformation and hydrothermal fluid flow
At progressively deeper levels in the crust, rockporosity is reduced and so is the fluid volume contained within that rock. It is also more difficult
for fluids to move pervasively through a rockand under these conditions fluid migration can
only take place along channelways represented
by structural discontinuities that form during
deformation. Most of the fluid movement that
occurs at deeper crustal levels, and is relevant to

the formation of mineral deposits, is locatedwithin faults that may, as a result, become loci formineralization. Fault displacements are generallyaccomplished by increments of rapid movementthat trigger earthquake events, usually in theupper 15 km (the seismogenic regime) of the crust(Sibson, 1994). The relationships between earthquake events, fault propogation, and fluid flow
have become a major and important area of study
that has relevance not only to ore genesis but tohydrology and earthquake prediction. It is, forexample, well known that seismic events trigger awide range of hydrological events that can varyfrom the sudden drying up of wells to significant
increases in the flow of rivers adjacent to faults(Muir Wood, 1994). The work of Sibson and coworkers (Sibson, 1986, 1987; Sibson et al., 1975,
1988) in particular has emphasized the importance of fault-driven fluid flow (Figure 3.6e) to theformation of hydrothermal ore deposits and this isdiscussed in more detail below.
Sibson et al. (1975) first introduced the term
seismic pumping to describe a theoretical modelin which it was envisaged that the cyclicity ofstress variations in and around a rupturing, seismically active fault system would affect local fluidpressures and promote the flow of fluids along
the fault (Figure 3.8a). Prior to a seismic event
friction between the opposing faces of a fault prevents rupture such that shear stress increases andthe adjacent rock undergoes dilation (Figure 3.8a
and b). The dilatant strain that develops in therock around the fault causes cracks to form, into
which fluids flow. Consequently, fluid pressures
are predicted to fall in the build up to fault failure. At the instant the fault ruptures and anearthquake occurrs, however, shear stress drops
significantly (∆τ in Figure 3.8b) but fluid pressure
increases, resulting in the upward expulsion of
fluids along the fault. As frictional forces arereinstated and shear stresses increase again, theentire cycle could repeat itself, causing episodic
fluid flow along fault-related channelways in thecrust. The seismic pumping model provided an
explanation and a mechanism for moving fluids
through the crust related to cyclical variations inboth the stress and strain state of rocks around
faults.

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3(a)Earthquake,seismic failure(b)
Build-up of
tectonic
shear
stress

σ2

Fault
plane

σ1
σ3

Stressorientation
Surface

Flu

Onset ofdilatancy

Fluid fills Collapse
dilatantof dilatantcracks
zone

∆τ

Shear
stress (τ)

ids

143

FluidsFluidpressure

Fault
plane

Dilationzone

Fluid flow
arounddilatant
zone

Inflow

OutflowTime
Earthquake
focus

Quartzfilled
veins

Limit of shear dislocation

Figure 3.8 Early model explaining the episodic flow of fluid along a seismically active fault zone. (a) Block diagram
illustrating the geometry of a fault generated seismic pumping system. (b) Explanation of the changes in tectonic
shear stress and fluid pressure with time before and after a seismic event, and the accompanying fluid flow direction
around the dilatant zone. Maximum outflow of fluid from the zone of dilatancy occurs as the seismic failure event
occurs and fractures close (after Sibson et al., 1975).

In reality the factors controlling channelized
movement of hydrothermal solutions through the
Earth’s crust are more complex than the mechanisms outlined in the Sibson et al. (1975) model.
The model is also somewhat unrealistic since
there seems to be little or no evidence to substantiate the sort of crack propagation in rocks
adjacent to faults that was envisaged in the model(Sibson, 1994). It is nevertheless still apparent
that stress/strain cycling in rocks in and around
faults is a major control of crustal fluid movement. The relevance of such processes to theformation of mesothermal lode-gold deposits has,for example, been demonstrated by Sibson et al.
(1988). These types of deposits (see Box 3.2)
are typically located within high angle (i.e. subvertical) reverse faults (Figure 3.9a). Such faults
are an enigma since compressive stresses in the
Earth’s crust generally form thrust faults that are
shallowly inclined (i.e. 25–30° to a horizontal
maximum principal compression). Friction theory suggests that reactivation of high angle faults
in a horizontal compressive stress regime canonly occur when the local fluid pressure reaches
or exceeds the lithostatic load. The subsequentmodel proposed by Sibson et al. (1988) envisages

144

PART 2 HYDROTHERMAL PROCESSES(a)(b) (Plan view)

Suction pump
FluidsEn echelon
strike slip
faultsFault valvesinVe

Dilational
fault
jog

ittl

Br

Mesothermal
lode-gold
depositσ3(c)Mid-crust

Plith

Epithermal
gold deposit

Fault
ruptureePrincipal
compressive
stress(σ1)

Activemetamorphism

High angle
reverse fault –
orogenic
lode gold

Fault
valve

PfluidDilatant
fractures

Fault
rupture

Phyd

e

ctil

Du

FluidsStrike-slip fault
jog – epithermal
gold

Suctionpump

Time
Figure 3.9 Models explaining the nature of hydrothermal fluid flow in (a) high angle reverse faults in a horizontal
compressive stress regime considered to be applicable to the formation of mesothermal lode-gold deposits such as
the Mother Lode in California and in many Archean greenstone belts; and (b) en echelon strike-slip fault arrays
associated with extensional stresses and considered applicable to some shallow level epithermal gold deposits.
(c) Plot of fluid pressure with time showing the cyclic fluctuations envisaged for the fault valve model (applicableto high angle reverse fault-related mesothermal lode-gold deposits) and the suction pump model (applicable tostrike-slip related fault jogs and shallow level epithermal gold mineralization). The diagrams are after Sibson (1987)
and Sibson et al. (1988).

that sites of mesothermal lode-gold mineralization are located in structures where fault rupture
has been caused by the attainment of lithostatic
fluid pressures. Figure 3.9c illustrates how fluidpressure in a seismically active fault builds up tolevels approaching the lithostatic load because
the upper crust represents an impervious cap to
the fluid reservoir beneath. Fault rupture occurs
as fluid pressures reach the lithostatic equivalentand the seismic event is accompanied by dilatancy and the development of a fracture permeability upwards and along the fault (Sibson et al.,1988). As the fault fails and shear stresses aresubstantially reduced, fluids are also discharged
into the open space created by the fault rupture
itself. The dramatic decrease in fluid pressure(back toward hydrostatic values) that results fromthe creation of open space could promote theprecipitation of dissolved constituents within thehydrothermal solution, causing the fault to reseal
itself. Once this has happened fluid pressures arelikely to increase again. This crack–seal mechanism is now known as the “fault valve model” andappears to explain many of the characteristics ofmesothermal lode-gold deposits (Figure 3.9a and c).
A slightly different mechanism of fluid movement is envisaged with respect to strike slip faulting associated with extensional stresses. Dilational
fault jogs are commonly found between the end of
one rupture plane and the beginning of another in
extensional en echelon fault arrays (Figure 3.9b),
and such sites are particularly favorable for epithermal gold mineralization at high crustal levels(Sibson, 1987). Fault jogs are typically character-HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

3.3.4 Other factors affecting fluid flow andmineral precipitation
Previous discussion has centered on fluid movement in large-scale structures such as faults thatare tens to hundreds of kilometers in length. Thissection considers smaller-scale fluid migration as
evident at mesoscale (as individual quartz veins)or microscale (as fluid inclusion traces in individual minerals) levels. At these scales effective flowof fluid through a rock is determined by hydraulicconductivity, or the extent to which fractures
along which fluids might flow are interconnected(Odling, 1997; Cox et al., 2001). In a rock containing very few fractures or cracks, and where poros-

Below percolation threshold

Above percolation threshold

Percolationthreshold

Permeability

ized by the development of extensional fractures
which are preserved as a network of quartz veins.
Fault jog related quartz vein arrays may also be
mineralized, an example of which is the Martha
Hill (or Waihi) gold deposit of the Coromandel
Peninsula in New Zealand (Brathwaite andFaure, 2002). The development of extensional
fractures in the fault jog results from the transfer
of fault slip from one en echelon fault segment to
the other during seismic activity along the fault
trace. The opening up of open space at highcrustal levels and the discharge of fluid into
these spaces results in a rapid fluid pressure drop
(Figure 3.9c) which could, at least within 2–3 km
or so of the surface, also be accompanied by boiling of the fluid. The mechanical energy released
by boiling would result in further hydraulic fracturing and brecciation, with enhanced fluid circulation and mineral precipitation. This particular
process is referred to as the “suction pumpmodel” (Sibson, 1987) and appears to have applicability to strike slip fault systems and epithermal fluid flow at shallow crustal depths (Figure
3.9b and c).
There is no doubt that deformation-controlled
fluid movement along major fault systems in the
Earth’s crust is of major importance to the formation of hydrothermal ore deposits over a considerable range of crustal depths. This is confirmed by
the fact that many hydrothermal ores are locatedwithin structural discontinuities and also that
exploration companies frequently target structural features in their quest for mineral deposits.

145

Fracture density
Figure 3.10 The small scale considerations of fluid
flow in a rock depend on the creation of permeability
and a percolation threshold, which is in turn a functionof connectivity and fracture density. The percolationthreshold is reached when one cluster of fractures
becomes large enough to span the sample region(after Odling, 1997).

ity is low, permeability may be effectively zeroand little or no fluid will flow through the rock
(Figure 3.10). As the fracture density increases,
so too does the probability of two cracks interconnecting. Eventually a point is reached, termed
the percolation threshold, where permeability
suddenly increases dramatically and fluid flow
across a finite volume of rock becomes possible(Figure 3.10; Odling, 1997). The attainment of apercolation threshold, at any scale in a rock mass,
is clearly necessary if effective fluid circulation is
to take place.
How do we know that a fluid had passed
through a rock?
The evidence that a fluid had once flowed througha rock mass is provided either by alteration (seesection 3.6 below) or by the presence of veins that

146

PART 2 HYDROTHERMAL PROCESSES

Dissolution of quartz
Precipitation of calcite

ath

ing p

Heat

Decreasing
calcitesolubility

Increasing
quartzsolubility

Increasing
temperature

er

c cc
c
c

Aquif
q

q
qq

h

g pat

n
Cooli

Pore fluid
flux

Precipitation of quartz
Dissolution of calcite

Figure 3.11 Schematic diagram showing a convecting pore fluid circulating in a porous sandstone and thecontrasting behavior of mineral precipitation and dissolution. In this model it is assumed that quartz solubility
decreases, but calcite solubility increases, as fluid temperatures decrease (after Byørlykke, 1994).

are typically filled with quartz or a carbonate
phase (i.e. calcite, dolomite, siderite, etc.), together
with other less common gangue minerals. Veinsare formed by an assemblage of minerals thatprecipitate from hot aqueous solutions as theypassed through a fracture, effectively fossilizing
and preserving the fluid conduit in the rockrecord. Fracture fill will develop in one of two
ways (Bjørlykke, 1994):
1 By diffusion of solids from the surrounds andprecipitation within a fracture or open space (i.e. avug). Since a mineral such as quartz has a lower
solubility in aqueous solution at lower pressuresand temperatures (see section 2.4.1 in Chapter 2),
it follows that it is likely to precipitate in a fracture because the open space will exist underhydrostatic pressure gradients compared to thefracture walls which carry higher pressure lithostatic loads.
2 By precipitation of minerals from fluids flowing
through a fracture. In this case material can bebrought in from a distant source and will precipitate in the open space as a function of many different processes (see section 3.5.1 below), such astemperature decrease, rock alteration, solubility
decrease, and boiling.
The factors controlling precipitation of minerals
are varied and complex and this topic is discussed
in more detail with respect to ore-forming constituents in sections 3.4 and 3.5 below. Minerals
also precipitate from fluids flowing pervasively
through a porous rock mass and this is the mech-anism responsible for diagenesis, whereby sedimentary particles are cemented together andlithified (i.e. turned into rock). Assuming equilibrium between pore fluid and host rock, thevolume of minerals precipitated from a fluid
circulating through a rock mass can be calculated
from the relation (after Bjørlykke, 1994):
Vm = Ftsin β(dT/dZ)α T/ρ

[3.1]

where Vm is the volume of mineral precipitated;
F is the fluid flux; t is time (seconds); β is the angle
between direction of flow and isotherms in therock; dT/dZ is the geothermal gradient; α T is afunction which reflects solubility in terms oftemperature, and ρ is the density of the mineralbeing precipitated.
A schematic illustration of mineral precipitation, and conversely of mineral dissolution, isshown in Figure 3.11, where pore fluid is considered to be convecting freely through a porous sanstone. Because the solubility of quartz is reduced
down a temperature gradient it will tend to beprecipitated along a cooling path, but dissolved
along a heating path. The volumes of quartz precipitated in such a situation can be calculated
from equation [3.1] above. Conversely, a mineralsuch as calcite exhibits retrograde solubility (itssolubility decreases as a function of increasing
temperature) and will, therefore, behave in the
opposite sense to quartz, tending to dissolve at
the sites of quartz precipitation and vice versa

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3(Figure 3.11). Different types of mineral precipitates in a sandstone aquifer can, therefore, provide
information about the nature and hydrology of
fluid flow. In this type of environment calculations will show that the precipitation of evensmall volumes of quartz or calcite requires huge
fluid fluxes, so that mass transfer becomes effective only when sediments have high permeabilities, or fluid fluxes are very focused (Bjørlykke,
1994). Although cooling is important, precipitation of minerals from circulating fluids is not only
temperature-dependent and many other factors
control the formation of ore constituents fromhydrothermal solutions (see section 3.5 below).

3.4 FURTHER FACTORS AFFECTING METAL SOLUBILITY
Solubility is defined as the “upper limit to theamount of dissolved metal that a hydrothermal
fluid can transport” (Wood and Samson, 1998).
In Chapter 2 the concept of metal solubility inhydrothermal aqueous solutions was introduced
and mention made of Holland’s (1972) classic
experiments in which the solubility of metals
such as Pb, Zn, Mn, and Fe was found to vary as
an exponential function of the chloride ion (Cl−)
concentration of the fluid. This work emphasized
the fact that efficient transport of metals byhydrothermal fluids in the Earth’s crust can beachieved only if aqueous solutions contain otherdissolved ingredients with which the metals canbond, thereby promoting metal dissolution. For
example, the solubility of a mineral such as
galena (PbS) in pure water, even at high temperatures, is extremely small. Wood and Samson(1998) calculated that the amount of Pb that could
be dissolved in an aqueous solution (with 10 −3
molal H2S and in equilibrium with galena at
200 °C) would be a trivial 47.4 ppb. If 5.5 wt%
NaCl were added to this solution, however, thesolubility of Pb at the same temperature would
increase significantly to 1038 ppm.
In addition to Cl−, a variety of different ligands(electron donors or electronegative ions with lonepairs of valence electrons) will affect the solubility of metals in solution. The ligand combineswith a metallic ion by the formation of a coor-

147

dinate bond, which is a type of covalent bond inwhich the shared electron pair is provided by oneof the participating molecules (the ligand in thiscase) rather than the normal situation where theshared electrons are supplied from each of the twoparticipating components. A ligand that promotes
the formation of complexes by coordination bonding will typically increase the solubility of metals
in aqueous ore-forming solutions.
Considerations of how metals go into solution
in hydrothermal fluids can be made with respect
to the concept of Lewis acids and bases. In the latter concept the conventional definition of acids
and bases (i.e. substances capable, respectively, ofcontributing H+ ions to, or taking up H+ ions from,a solution) is expanded to situations where H+
ions are not present, and this has particular relevance to geological situations (Gill, 1996). A
Lewis acid is defined as an atom or molecule that
can accept a lone pair of valence electrons,
whereas a Lewis base can donate an electron pair
to a bond. “Hard” Lewis acids are strongly electropositive metals (with high charges and/or
small atomic radii) such as the alkali and alkaline
earth metals (e.g. Na+, K+, Mg2+, Ca2+) that form
ionic bonds with strongly electronegative elements like oxygen (O2−). “Soft” Lewis acids havean abundance of easily accessible electrons intheir outer shells and prefer to form covalentbonds with soft bases. Soft acids are typified by
the chalcophile metals (i.e. those that have anaffinity with sulfur, such as Cu, Pb, Zn, Ag, Bi,Cd) which tend to form covalent bonds with ligands of low electronegativity, such as S2−. The
principle stating that in a competitive situation
hard metals (acids or electron acceptors) will tend
to complex with hard ligands (bases or electron
donors), and soft metals with soft ligands, iscommonly referred to as Pearson’s Principle, andunderpins the solubility behavior of metals inhydrothermal solutions. A classification of metals and ligands in terms of the hard–soft breakdown and applicable to ore-forming processes ispresented in Table 3.2. The “borderline” category
is added to accommodate the fact that somemetals (such as Fe and Pb) can complex readily
with both hard and soft ligands, forming a range ofminerals such as sulfides (pyrite FeS2, galena PbS)

148

PART 2 HYDROTHERMAL PROCESSES

Table 3.2 Classification of some metals and ligands in terms of Lewis acid/base principles
HARD METALSBORDERLINE

SOFT METALS

Li+ Na+ K+ Rb+ Cs+
Be2+ Sr2+ Ba2+ Fe3+
Ce4+ Sn4+ Mo4+ W4+ V4+ Mn4+
As5+ Sb5+ U6+

Divalent transition metals(Zn2+ Pb2+ Fe2+ etc.)

Au+ Ag+ Cu+
Hg2+ Cd2+ Sn2+ Pt2+ Pd2+
Au3+ Tl3+⇓⇓⇓

HARD LIGANDSNH3OH− F− NO3− HCO3− CH3COO−2−
CO 2−
3 SO 4
PO43−−

BORDERLINE

SOFT LIGANDS−− −HS I CN− H2S S2O 2−
3

Cl Br

The major ligands in natural hydrothermal solutions are shown in bold. Caution is required in the application of the hard–softmodel since the structure of water changes at higher temperatures and metal–ligand interaction will, likewise, change(Seward and Barnes, 1997).
Source: after Pearson (1963).

and carbonates (siderite FeCO3, cerrusite PbCO3).
Likewise Cl− can be an effective complexing agent
for both hard and soft metals, and therefore promotes the solubility of a wide range of metals inhydrothermal aqueous solutions.
The classification of aqueous ionic species into
hard and soft Lewis acids/bases assists considerably in understanding the nature of metal–ligand
complexation, and, therefore, the controls on solubility, in hydrothermal solutions (Seward andBarnes, 1997; Wood and Samson, 1998). In addition, however, it is well known that temperatureplays a very important role in determining thedegree to which metals enter solution. In mostcases stabilities of metal–ligand complexes suchas PbCl+ and ZnCl+ will increase by several ordersof magnitude as temperature increases fromambient values to 300 °C (Seward and Barnes,
1997). This is illustrated, together with data forother metal–chloride complexes, in Figure 3.12.
By contrast, an increase in pressure will have theopposite effect to temperature and the stabilitiesof metal–ligand complexes will tend to decrease
because of bond dissociation and the formation offree metal ions. In general, however, the effect onsolubilities of pressure increases are minimal andmore than offset by the significant changes associated with temperature increases (Seward andBarnes, 1997).

3.4.1 The important metal–ligand complexes inhydrothermal solutions
Wood and Samson (1998) have provided an
excellent summary of solubility data for a wide
variety of metal–ligand complexes. These data aredescribed below in terms of hard, borderline, andsoft metals. The summary given below provides
an indication of the most likely metal–ligand
complex that will exist in a natural fluid as afunction of variables such as oxidation state, pH,temperature, and fluid composition. However, itcannot and should not be used to try to simply
predict the nature of metal speciation in any
situation where the many parameters that control
solubility are imperfectly known. The concept of
hard–soft Lewis acids and bases is an idealized
scenario predicated on the assumption that there
is competition between metals and a variety ofligands. In natural situations metals may, for example, complex with the most suitable available
base and the formation of actual metal–ligand
complexes and their solubilities will be quite different to those predicted on theoretical grounds.
Hard metals
Tungsten (W) Tungsten tends to occur in natureas the hexavalent aqueous cation W6+, althoughHYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

149

8

Zn
Cl +

7
6
5
log β1

+

l3
UC

4

+

Cl
Pb

°
Cl
Ag

3
Figure 3.12 Plot of the effective
stability of a metal complex(expressed in terms of β1, whichis the equilibrium formation
constant) versus temperature(after Seward and Barnes, 1997).
The stabilities of metal–chloride
complexes shown increase by
several orders of magnitude as
temperature increases, as will theirsolubilities in aqueous solution.

+

l

C
Mn

2

+

l
FeC

1
0
–1

the pentavalent form W5+ can also occur undermore reducing conditions. These ions are hard
Lewis acids and are likely to complex with hardbases such as O2−, OH−, F−, and CO 2−
3 , although
the chloride ion is relatively unimportant as acomplexing agent under most geological conditions. The majority of dissolved tungsten species
in hydrothermal solutions occur as tungstates and−have the forms WO 2−
4 , HWO 4 , and H2WO4. The
stabilities of these tungstate species are obviously
dependent on hydrogen ion concentration (i.e. onpH), such that H2WO4 is only stable under acidic
conditions, whereas WO 42− is stable at alkaline
pH. Significant quantities of tungsten can be carried in solution by these tungstate complexes andno other species are required to form most scheelite and wolframite deposits.
Molybdenum (Mo) Molybdenum tends to be
more easily reduced in nature than tungsten andtypically occurs in a variety of aqueous valence
states as Mo6+, Mo5+, Mo4+, and Mo3+. It is also a

100

200
Temperature (°C)

300

400

hard Lewis acid and like tungsten will complex
with hard bases, in particular, O2− and OH−.
Chloride complexing takes place only under very
acidic conditions and typically it is the oxyhydroxide complexes (such as MoO +2 and Mo(OH) +3)
that are the most likely to occur in natural aqueous solutions.
Uranium (U) Uranium is characterized by twovalence states in nature, as the quadrivalenturanous ion, U4+, and the hexavalent uranyl ion,
U6+. The uranous ion is generally characterized
by very low solubilities in aqueous solutions andmost fluid-related transport takes place as U6+. Asa hard acid, U6+ will complex with hard bases
such as O2−, OH−, F−, HCO 3− , and CO 32−.
Borderline metals
Arsenic (As) The trivalent cation As3+ dominates the valence state of arsenic in nature andin solution the neutral H3AsO3 complex is known

150PART 2 HYDROTHERMAL PROCESSES

to be stable over wide ranges of temperature(above 200 °C), Eh, and pH. Other less protonatedarsenate complexes are also stable at higher pH
values (as for tungsten) and generally chloride
complexes do not play a role in fluid transport.
Under conditions of high sulfide concentration
and at low temperatures, As–sulfide complexes(such as AsS2(SH)2−) may become more important,although their exact stoichiometric proportions
are unclear.
Antimony (Sb) The aqueous geochemistry of
antimony is similar to arsenic, with Sb3+ predominating and Sb(OH)3 occurring as the main stablecomplex in low-sulfide aqueous solutions. In thepresence of high concentrations of reduced sulfide
species in the fluid a number of thioantimonite
complexes (such as HSb2S −4 ) are also likely to bestabilized and contribute to increasing antimony
solubility.
Iron (Fe) The majority of iron transported inaqueous solution through the Earth’s crust isin the ferrous or divalent form (Fe2+), while ferriciron (Fe3+), which forms under relatively oxidizing
conditions, is much less soluble. With its intermediate Lewis acid properties, iron does notexhibit a preference for either hard or soft bases
and a variety of metal–ligand complexes are
implicated in the dissolution of ferrous iron inaqueous solutions, including Cl−, OH−, and HCO 3−.
Most experimental studies indicate that FeCl+and FeCl2 are the main complexes involved in thehydrothermal transport of ferrous iron, especiallyat high temperatures and salinities. The freehydrated ion itself, Fe2+, is also implicated in fluid
transport where hydrolysis has taken place, as are
Fe–bicarbonate complexes under more alkalineconditions. Fe–bisulfide complexes are generally
not involved in the transport of iron in hydrothermal solutions although they may be important inocean floor exhalative environments (i.e. blacksmokers).
Manganese (Mn) Manganese is also a borderline
Lewis acid and, like iron, Mn2+ complexation inhydrothermal solutions is likely to be dominatedby the chloride ligand (i.e. MnCl+ and MnCl2),
with hydroxide and bicarbonate complexes also
contributing to Mn solubility.
Tin (Sn) Tin is a metal that exhibits both hardacid quadrivalent (as Sn4+) and borderline divalent(as Sn2+) traits such that it can be solubilized bycomplexation with a number of different ligands.
Under oxidizing conditions it has been found that
the Sn4+–hydroxychloride complex, Sn(OH)2Cl2 ,is the dominant species, but that its solubility islow. Under more reducing conditions both Sn4+and Sn2+ can complex with the simple chloride
ion, forming very soluble complexes of the form
SnCl nX−n (i.e. where X is the valence state, either 2
or 4, and n is the ligation number for the chloride
complex). The divalent Sn2+–chloride complexes,
formed under more reducing conditions, exhibit
much higher solubilities than the quadrivalent
Sn4+–hydroxychloride complex that exists in amore oxidized state. Sn–hydroxide complexes(Sn(OH)4 and Sn(OH)2) are very stable under alkaline, lower temperature conditions but their solubilities are again much lower than those exhibitedby the dominant Sn–chloride complexes formed at
higher temperatures and lower pH. Contrary to
expectation, tin does not complex significantlywith fluorine despite its association with highly
fractionated fluorite/topaz-bearing granites.
Zinc (Zn) The dominant divalent zinc cation,
Zn2+, is a borderline Lewis acid that complexes
with a variety of ligands, including Cl−, HS−, OH−,
HCO −3 , and CO 32− . At relatively low temperaturesand high pH, and in fluids with high bisulfide
concentrations but low salinities, a variety of
Zn2+–bisulfide complexes are stable, including
Zn(HS)2, Zn(HS) −3 , and Zn(HS) 2−
4 . In contrast, at
higher temperatures and under more acidic conditions, a host of chloride complexes predominate,including Zn(Cl)+, Zn(Cl)2, Zn(Cl) −3 , and Zn(Cl) 2−
4 .
Lead (Pb) Lead is the softest of the borderline
Lewis acids and forms stronger bonds with Cl−
and HS− ligands than with OH− or HCO −3 . The
valence state and complexing behavior of lead isvery similar to that of zinc. Pb2+–bisulfide com-

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

plexes are more stable than chloride complexes in
low temperature, neutral to alkaline, low salinity
solutions. Different Pb2+–bisulfide complexes form
as a function of pH, and these include Pb(HS)2under acidic conditions and Pb(HS) −3 at higher pH(i.e. >6). At higher temperatures and lower pH,
however, a variety of Pb2+–chloride complexes
form and these include Pb(Cl)+, Pb(Cl)2 , Pb(Cl) −3 ,and Pb(Cl) 42−. Pb2+–carbonate or Pb2+–bicarbonate
complexes can form in CO2-bearing fluids but
their solubilities are typically low and they are
unlikely, therefore, to have relevance in oreforming processes. There is, however, compelling
evidence that lead, and to a lesser extent zinc,
are transported by organo-metallic complexes (seesection 3.4.2 below), especially in relatively low
temperature environments such as those related
to the formation of Mississippi Valley type Pb–Zn
deposits (see Box 3.7), where hydrothermal fluids
are known to contain hydrocarbons. The acetate
ligand (CH3COO−) is one organic molecule that has
been implicated in base metal transport, although
there are others, such as thiols or sulfur-bearingorganic compounds, that may also be important.
Soft metals
Copper (Cu) In nature, copper occurs in bothCu+ and Cu2+ valence states, although the monovalent state is predominant in most hydrothermal
fluids. It is a relatively soft acid and forms
stable Cu+–chloride and Cu+–bisulfide complexes,
as well as Cu+–hydroxide complexes if otherligands are not available. Movement of copperin hydrothermal solutions under a wide rangeof conditions is most likely to take place by
Cu+–chloride complexation in the form CuCl −2 .
Although copper–bisulfide complexes will formin high-sulfur environments, they are not likely
to play a significant role in most hydrothermal
ore-forming environments.
Gold (Au) Gold is the softest metal ion considered in this summary and prefers complexation
with soft ligands such as the bisulfide ion, rather
than chloride. Gold typically occurs in a monovalent state (Au+) in aqueous solutions and will form151

the complex AuHS at low pH and Au(HS) −2 under
weakly acidic to basic conditions. The bisulfide
complexes predominate in most hydrothermal
fluids and are stable over a wide temperature
range. Under very oxidizing, saline, and acidic conditions, it is possible to form both Au+– andAu3+–chloride complexes in the form of Au(Cl) −2
and Au(Cl) −4 respectively. The former complex
might prevail in certain high temperature environments such as porphyry coppers, whereas thelatter is only likely to be important at low temperatures in near surface settings. Au does form very
soluble complexes with cyanide (Au(CN)−2), whichis the reason why cyanidation is used so effectively at low temperatures to recover gold fromore. The cyanide molecule is, however, unstable
at higher temperatures and the complex has no
relevance to hydrothermal gold transport.
Silver (Ag) Silver is also a soft metal ion, but not
as soft as gold. It is, therefore, more likely to form
complexes with a borderline ligand such as chloride than gold is, although solution chemistry willagain be dominated by Ag+–bisulfide bonding.
The predominant complexes are AgHS and Ag(HS)−2
under acidic and neutral to alkaline conditions
respectively. Several Ag+–chloride complexes are
known to be stable over a wide range of conditions pertaining to hydrothermal fluid flow andthese include AgCl, Ag(Cl) −2 , and Ag(Cl) 2−
3 .
Mercury (Hg) Mercury, occurring predominantlyas Hg2+ in aqueous solutions, is a soft metal ion
and will complex preferentially with bisulfide
ligands. Most mercury transport under neutral toalkaline and moderately reducing conditions islikely to take place by the formation of complexes
such as Hg(HS)2, HgS(HS)−, and HgS 2−
2 . Harder ligands probably do not play much of a role in mosthydrothermal fluids. Of interest and related to its
low boiling point, however, is the fact that mercury can be transported as a vapor, or as elemental
Hg in either aqueous solutions or hydrocarbonbearing fluids.
In summary, it is apparent that the borderline
chloride ion is the most important ligand that

152

PART 2 HYDROTHERMAL PROCESSES

promotes the dissolution and transport of metals
in hydrothermal fluids, followed by bisulfide. A
variety of other ligands such as OH−, HCO −3 , CO 32− ,−−F−, SO 2−
4 , NO 3 , NH3, and CH3COO are of lesser
importance because they are either stable underabnormal conditions or exist at very low concentrations in natural fluids. In addition to its ability
to complex with a wide range of both hard andsoft metal cations, Cl− is also the major anionicspecies in most natural fluids. This ensures thatmany natural fluids are very capable of transporting large quantities of metal in solution and it isfor this reason that the circulation of hydrothermal solutions in the Earth’s crust is such animportant ore-forming process. Even thoughmetal–ligand complexes may be readily stabilizedin natural aqueous solutions, it is neverthelessapparent from both direct analysis and theoretical
calculation that the actual concentrations of metals dissolved in most hydrothermal fluids are very
low. This means that the formation of a viable
ore deposit still requires large volumes of fluid
passing through a highly focused point in the
Earth’s crust (i.e. high fluid/rock ratios), as well as
efficient precipitation mechanisms to take metals
out of solution and concentrate them in the hostrock. The compilation shown in Figure 3.13
shows the minimum concentrations of precious
and base metals (i.e. the vertical dashed lines)
estimated to be necessary for the formation of adeposit, and also the ranges shown for several
examples where such data are available. Solutions
that deposit Au and Ag typically contain much
lower concentrations, in the range 1 ppb to 1 ppm,
than those associated with Cu, Pb, and Zn ores. In
the latter case massive sulfide deposits are formed
from fluids that typically carry 1–100 ppm ofmetal, whereas the fluids associated with MVT
and porphyry deposits are relatively enrichedwith concentrations in the range 100–1000 ppm.
3.4.2 A brief note on metal–organic complexes
A number of different hydrothermal ore types
that are hosted in intracratonic and rift-related
sedimentary sequences are known to have formed
from fluids containing significant amounts of dis-solved organic compounds. It is also well known
that crude oil and bitumen invariably contains
minor amounts (<1%) of inorganic material that
includes dissolved metals. It is not surprising,
therefore, that considerable attention has recently
been given to the roles that metal–organic complexes might have played in hydrothermal ore
forming processes, especially in Mississippi
Valley-type (MVT) Pb–Zn and stratiform, red-bed
or shale-hosted base metal ores (see Parnell, 1994;
Gize, 1999).
Organic compounds in connate waters tend tooccur either as neutral hydrocarbon molecules(alkanes, aromatics, etc.) or as charged anionic species such as the carboxylic acids (acetic acid, oxalic
acid, etc.). Most neutral hydrocarbons (with the
exception of methane, CH4) have only limited
aqueous solubility and this decreases exponentially as the hydrocarbon molecule gets bigger andmore complex. Organic acid anions, on the other
hand, may have very high solubilities in water andthe dominant aqueous anion, acetate (CH3COO−),
can exist at concentrations up to 2000 mg l−1(Hanor, 1994). Organometallic complexes (i.e.
those in which a metal is covalently bonded to thecarbon of an adjacent hydrocarbon molecule) only
rarely occur in nature, whereas metal–organic
complexes, in which a metal cation is attached tothe electron-donor atom (such as oxygen, sulfur,
or nitrogen) of an organic ligand, are widely stablein natural aqueous solutions, as well as in ore
fluids (Giordano and Kharaka, 1994).
In ore fluids metal–organic complexes will only
exist if they are sufficiently stable to be able tocompete with the normal inorganic complexes
discussed previously. Modeling has shown thatPb and Zn can be transported as metal–organic
complexes involving acetate or oxalate (C2O 2−
4 )2−
ligands (e.g. Zn(CH3COO−)2 and Pb(C2O 2−4 ) 2 ) in
reduced fluids with greater than 10−5 molal H2S orHS−. In more oxidized fluids, such as those more
applicable to red-bed related stratiform base metal
deposits, Pb and Zn would be transported predominantly as metal–chloride complexes, but Pb– andZn–acetate complexes could still account for up
to 5% of metal transport (Giordano and Kharaka,1994).

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

Skarn

153

740–7700

100–4600

Porphyry

Base metal ores

80–4000

10900
9600
200
131

Mississippi Valley

1330
540

Vein

100–1000
20–2200

9.2
4.7
3.5

Massive Sulfide

7.0
6.5
5.0
1.2

Precious metal ores

Figure 3.13 Base and precious
metal concentrations in oreforming hydrothermal solutions,in each of five major categories ofdeposit type. Base metals include
Cu, Pb, and Zn (filled dots), andprecious metals comprise Au, Ag,
and Hg (open circles). The numbers
next to individual points or ranges
refer to average concentrations or
ranges in a variety of different
situations, and are all after Seward
and Barnes (1997).

40–55
38–62
35
20–29

0.1–1.0 (Ag)0.008 (Ag)0.0015Vein0.001– 0.01

0.001

–3

–2

3.5 PRECIPITATION MECHANISMS FOR METALSIN SOLUTION

The previous section has shown that it is possible
to stabilize many different metal–ligand complexes
under wide ranging conditions in natural hydrothermal solutions, thereby promoting the kind ofsolubility levels required to make effective oreforming fluids. Once a metal is in solution, however, it then needs to be extracted from that fluid

–1

0
1
2
Log concentration (ppm)

3

4

5

and concentrated in a portion of the Earth’s crustthat is sufficiently restricted and accessible to makean economically viable ore body. It is obvious thata wide range of precipitation mechanisms are likely
to be effective since any mechanism that will
destabilize a metal–ligand complex and, therefore,
reduce the metal solubility, will cause it to be deposited in the host rock through which the hydrothermal solution is passing. The following sections
examine the physical and chemical controls on

154PART 2 HYDROTHERMAL PROCESSES

metal precipitation, as well as adsorption andbiologically mediated processes. The latter is particularly important as it is now widely acceptedthat the geochemical pathways by which both
metals and non-metals are distributed in the
Earth’s crust are affected to varying degrees by theaction of microorganisms. The concept of biomineralization is becoming increasingly important tothe understanding of ore-forming processes andthis topic is likely to be the subject of much more
research in the future.
3.5.1 Physico-chemical factors affecting
metal precipitation
The basic principles that control ore deposition
have been summarized by Barnes (1979b) and
Seward and Barnes (1997). At shallow crustal
levels ore deposition will take place by open spacefilling, whereas deeper down where porosity is
restricted, replacement of existing minerals tends
to occur. Decrease in temperature is the factor
that, intuitively, is regarded as the most obvious
way of promoting the precipitation of metals fromhydrothermal fluids. At depth, however, temperature gradients across the structures within which
fluids are moving tend to be minimal and metal
precipitation will be neither efficient nor wellconstrained to a particular trap zone. Depositionof metals in such a case is achieved more effectively by changing the properties or compositionof the hydrothermal fluid. If ore solution occurs
by metal–chloride complexing then precipitation
could occur very efficiently by increasing the pHof the ore fluid (see Figure 3.14). An example isprovided by the reaction of an acidic ore solution
with a carbonate host rock and where precipitation
is promoted by digestion and replacement of thehost. Two coupled reactions illustrate the process
in terms of calcite dissolution by an originallyacidic, Zn2+–chloride-bearing ore solution:
CaCO3(SOLID) + 2H+ ⇔ Ca2+ + H2CO3(AQUEOUS)[3.2]
and subsequent sphalerite precipitation caused byincrease of pH in the fluid as hydrogen ions are
consumed:

ZnCl2n − n + H2S(AQUEOUS) ⇔ ZnS(SOLID) + 2H+ + nCl[3.3]
It is apparent from reactions 3.2 and 3.3 that precipitation of metal–chloride complexes could also
take place by increasing the H2S concentrationof the fluid (perhaps by fluid mixing or bacterial
sulfate reduction) or by decreasing the Cl− concentration (again perhaps by mixing of the ore
fluid with dilute groundwater). As another option
to calcite dissolution, an increase in fluid pH
could also be brought about by boiling of the fluid,which results in acidic volatiles being partitioned
preferentially into the vapor phase. There are,therefore, a number of ways of reducing the solubility of metal–chloride complexes and promoting
metal precipitation.
The factors that promote precipitation of metal–
sulfide complexes may be somewhat different
to those for chloride complexes (Barnes, 1979b).
Although fluid mixing (or dilution) and temperature decreases will promote metal precipitation, it
is oxidation of the ore fluid (see Figure 3.14) thatis particularly effective in decreasing the solubilities of metal–sulfide complexes. Oxidation (orloss of electrons) causes a decrease in the pH andalso the total sulfide concentration, thereby promoting metal precipitation, as shown in reaction 3.4:
Zn(HS)−3 + 4O2(AQUEOUS) ⇔ ZnS(SOLID) + 3H+ + 2SO 2−
4[3.4]
Precipitation of ores from metal–sulfide complexes are controlled by oxidation, decreases inpH (or acidification), and decreases in sulfide
concentration in the fluid, all tendencies thatare opposite to those controlling metal–chloride
disassociation. A more detailed discussion of theactual geological processes that affect ore deposition controls is presented below, together withdifferent scenarios where they might apply. An
understanding of the geological factors relating tometal deposition is obviously very important to
the exploration geologist who needs to have a firm
grasp of the concepts that control where and whyhydrothermal ore bodies form.

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

Temperature
Since the stabilities of many metal–ligand complexes increase as a function of temperature(Figure 3.12) it is clear that cooling of a fluid will
generally have the effect of promoting ore deposition. Temperature decrease is particularly effective
for destabilizing metal–chloride complexes because their solubilities are much more sensitive
to temperature changes than are those of equivalent sulfide complexes. As mentioned previously,
however, at deep crustal levels where fluid androck are in equilibrium and exist at the same temperature, thermal gradients are minimal and cooling is a slow and ineffectual process. Under thesecircumstances the concentration of metals andthe formation of hydrothermal ores are generally
not mediated by temperature decreases.
In the near surface environment, however,rapid decreases in the temperature of an ore
fluid are much more likely to happen and insuch environments cooling is likely to play animportant role in ore formation. The ocean floor
undoubtedly represents the prime example ofwhere a dramatic reduction in the temperatureof ore-forming fluids plays the dominant rolein controlling metal deposition. Volcanogenic
massive sulfide deposits typically form at sites
where hot (up to 350 °C) brines with very highmetal contents are vented onto the ocean floor as
black smokers (see section 3.8.1 and Box 3.3).
Precipitation of base metals (mainly Cu and Zntransported as chloride complexes; Scott, 1997)
and, in certain cases, precious metals (Au transported as a gold–bisulfide complex; Scott, 1997)
is virtually instantaneous as the ore fluids mix
with an essentially infinite volume of very cold
(2–4 °C) sea water. In this environment it is thevery rapid cooling through virtually 350 °C thatcauses a highly efficient precipitation of metal in
the immediate vicinity of the exhalative vent.
Black smokers are probably the most dramaticand efficient example of hydrothermal ore deposition anywhere on Earth.
The volcanic settings which give rise to thedeposition of epithermal Au–Ag deposits (seeChapter 2, section 2.11 and Box 2.4) are also characterized by steep geothermal gradients and high155

conductive and convective heat loss by fluids tothe surrounding rocks. Cooling may, therefore, playa role in controlling metal solubilities althoughin such environments it is more likely to be thedecrease in pressure and associated phase separation (i.e. boiling and effervescence; see below) thatplays the dominant role in ore deposition.
Pressure
Pressure variations do not dramatically effect thesolubilities of metal–ligand complexes although
it is clear that pressure increases will lead toa volume reduction which, in turn, promotes
the dissociation of complexes to ionic species(Seward, 1981). In effect a decrease in pressure
tends to favor an increase in solubility and, therefore, works in the opposite sense to temperature.
Pressure reduction is not a process that, on its
own, is typically associated with ore formation. A
major exception occurs, however, when pressure
reduction is accompanied by an event such as
boiling or effervescence. These specific processes
are very important in ore deposition and are generally promoted by decreases in pressure. In thissense pressure only has an indirect influence on
ore formation.
Phase separation (boiling and effervescence)
Although one might intuitively expect that boiling is promoted by an increase in temperature,
there is no doubt that, in the upper levels of the
Earth’s crust, the transformation from liquid tovapor usually occurs because of a decrease in fluidpressure (see trend A–A′ in Figure 3.5a). At deeper
levels in the crust a decrease in fluid pressure
could also be accompanied by effervescence,or the transition from a single phase H2O–CO2
mixture to one where H2O and CO2 unmix (seeFigure 3.5c). Both these processes are potentiallyvery important as mechanisms of precipitating
metals from ore-forming solutions because they
dramatically change the prevailing conditions
under which metal–ligand complexes are stable.
In section 3.3.3 the relationships between
fluid flow and deformation were discussed andthe fault-valve and suction-pump models were156

PART 2 HYDROTHERMAL PROCESSES

described with reference to the formation ofmesothermal lode-gold and epithermal Au–Ag
deposits respectively. The metamorphic fluids
generally attributed to formation of mesothermal
lode-gold deposits are characterized by single
phase H2O–CO2 mixtures and it is possible thatthe episodic decreases in fluid pressure accompanying crack–seal processes in a fault or shear(Figure 3.9a and c) could promote precipitationof ore components. This is because there is atendency for the H2O–CO2 solvus to move tohigher temperatures with decreasing pressure(Brown, 1998). A typical mesothermal lode-gold
fluid comprising H2O + CO2 at 350 °C and with
2.6% NaCl, for example, will initially exist as asingle homogeneous liquid just above the relevantsolvus (shown as point A; Figure 3.5b). A pressure
decrease will have the effect of moving the solvus
plateau to higher temperatures, so that the fluid
enters the immiscible, or two phase, field segregating CO2 and H2O. CO2 effervescence willundoubtedly affect the prevailing fluid properties
and likely promote precipitation of both gangue
and ore components, as suggested in the faultvalve model.
Similarly, in the near surface environment
where en echelon strike slip faulting results in
fault jog formation, the resulting rupture-related
dilation will instantaneously reduce pressure(Figure 3.9c) to such an extent that a fluid could
move from a liquid to the vapor state. Boiling will,of course, only occur if the pressure reduction
traverses the boiling point curve for the relevant hydrothermal fluid composition (refer tophase diagrams in Figure 2.2). The boiling awayof an aqueous vapor from a hydrothermal fluid
will result in the residual enrichment of solute,as well as partitioning of other volatile species
into the gas phase, with probable increase of pH
in the remaining fluid. Such processes will dramatically modify prevailing fluid properties andthe pH increase will, as mentioned previously,
be particularly effective in precipitating metal–
chloride complexes from solution. When accompanied by temperature decreases, prolonged
boiling will likely bring about solubility reduction of most dissolved species. As suggested by
Sibson (1987) with respect to epithermal ore-

forming environments, boiling is likely to be avery efficient mechanism for ore deposition atshallow crustal levels.
Fluid mixing/dilution
The mixing of two fluids is widely regarded
as another important mechanism for reducing
solubility in ore-forming solutions and promoting
metal precipitation. This is particularly the casewhen a relatively hot, metal-charged ore fluid
mingles with a cooler, more dilute solution.
Mixing of the two fluids would result in coolingof the hotter with modification of the prevailing ore fluid properties and destabilizationof existing metal–ligand complexes. There areseveral examples that point to the importanceof fluid mixing during ore deposition. In a classic
stable isotope study, Sheppard et al. (1971) demonstrated that precipitation of chalcopyrite at
the interface between the potassic and phyllicalteration zones associated with porphyry copper
deposits (see Box 2.1, Chapter 2) coincides withthe zone where fluids of dominantly magmaticorigin mingled with those derived from a meteoric source. A similar situation is known toexist with respect to precipitation of Au and Ag
ores in epithermal deposits (Box 2.4, Chapter 2).
Hedenquist and Aoki (1991) showed that magmatic fluids venting directly from a volcano without any mixing or dilution below the surface areless conducive to the formation of economically viable deposits since ground preparation (oralteration) is minimized and the resultant ore
accumulations at the surface have a low preservation potential. By contrast, systems characterized
by mixing of magmatic waters with a meteoric
water carapace 1–2 km below the surface willmore likely be associated with the formationof economically viable mineralization. This isbecause mixing of a hot, saline, metal-charged
magmatic fluid with a cooler, more dilute meteoric water promotes acid leaching of the hostrocks, increases their permeability, and forces thefluids to condense and precipitate their dissolved
metal solute. Hedenquist and Aoki (1991) suggested that interaction of magmatic fluids with ameteoric fluid blanket near the surface of activeHYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

157

Fluid mixing and metal precipitation: the Olympic
Dam iron oxide–copper–gold deposit, South Australia
The Olympic Dam deposit, some 500 km north of Adelaide
in South Australia, was discovered in 1975 beneath 350 mof sedimentary cover. It represents one of the most spectacular discoveries of modern times, containing around 2 billion tons of Cu, U, Ag, and Au ore (Solomon and Groves,1994). The ore body is hosted in a breccia complex within
the apical portions of the 1588 Myr old (Creaser andCooper, 1993) Roxby Downs granite, the latter being a K-,
U-, and Th-enriched A-type or anorogenic granite. Mineralization is generally attributed to mixing of two entirely
different fluids, one of possible magmatic derivation, andthe other an oxidized meteoric fluid (Reeve et al., 1990;
Oreskes and Einaudi, 1992; Haynes et al., 1995).
The breccias that host the ore comprise a variety of fragments, the majority of which are either granitic or hematitic,
together with volcanic fragments, sedimentary rocks, andmassive sulfide ore. The breccias are extremely iron richand themselves comprise a huge iron resource made up ofhematite and minor magnetite. Polymetallic mineralization,in the form of chalcocite, bornite, chalcopyrite, pitchblende, argentite, gold, and a variety of rare earth element(REE) minerals, is most closely linked to the hematitic
breccia. All breccias are cut by mafic and felsic dykes thatare approximately the same age as the Roxby Downs
granite, indicating that magmatism, brecciation, and mineralization were broadly coeval (Johnson and Cross, 1991).
Ore genesis at Olympic Dam (Solomon and Groves,1994) was initiated by the high level intrusion of fertile
A-type granite, with accompanying fluid exsolution andboiling, which gave rise to explosive volcanic activity andan early phase of brecciation. The fluid involved here isinferred to be of magmatic origin and moderately oxidizing. It was also clearly Fe-rich and precipitated an early
generation of magnetite as well as REE-bearing minerals.
Cooling of the magmatic system subsequently led to
draw-down of groundwater (perhaps derived from a saline

South Australia
Olympic Dam
Roxby Downs

Adelaide
Volcaniclastic
rocks

WhenanShaft

Granite breccia

Hematitequartzbreccia

Hematite
breccia

TN MN

1

0

Roxby Downs
Granitekm

Figure 1 Simplified, subsurface geological map of the Olympic Dam ore deposit (after Haynes et al., 1995).

(a)Geological setting
Playa
lake

Olympic Dam
breccia complex

Dykes

BasaltRoxby Downs granite(b)

Fluid mixing

Oxidizing
meteoric
water
Hotter, magmatic
fluid

Ore precipitation
Ore leachingplaya lake above the deposit; Haynes et al., 1995) that inturn led to phreatic explosive activity, further brecciation, and hematite precipitation. The spatial association
between hematite breccia and polymetallic mineralization
indicates that a low temperature, highly oxidizing, meteoric
fluid was implicated in the dissolution and transport ofsubstantial amounts of Cu, Fe, and U, possibly from thegranite, or overlying volcanic rocks, or both. This fluid
encountered and mixed with the less oxidized magmatic
solutions (containing reduced sulfur species), destabilizing metal–ligand complexes and causing precipitation of

Figure 2 Geological model forthe formation of the Olympic
Dam deposit (after Haynes et al.,1995). The geological setting
and the distribution of a lower,
hotter, moderately oxidizing
fluid of possible magmatic
derivation, and an upper, saline,
highly oxidizing meteoric fluid,
is shown in (a). Drawdown of
meteoric fluid together withuprising plumes of hotter waterpromote mixing of the fluids
and metal precipitation, as
shown in (b).

metal sulfides together with iron oxides. Figure 2 illustrates
a model showing the existence of two fundamentallydifferent fluid types at different crustal levels, with ore
precipitation occurring at the interface between the fluids.
It is envisaged that fluid circulation and mixing continued
episodically for a considerable period of time, perhaps
in part stimulated by the high radioelement (K, U, Th)
content of the granite and the resultant heat generated
by radioactive decay. The longevity of the ore-forming
process is consistent with the complexity and enormous
size of the ore deposit.

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

volcanic regions may be an important prerequisite to the formation of viable epithermal Au–Ag
mineralization.
In another example demonstrating the importance of mixed fluids in ore formation, it has beensuggested that sediment-hosted ore deposits associated with continental rifts environments are theproduct of the mixing of meteoric and connate
fluids. Using the active mineralizing systems of
the Salton Sea geothermal system (see section 3.7
below) as an analogue, McKibben et al. (1988)
have suggested that the mixing of upwellinghigh-salinity brines of essentially connate origin
with descending meteoric waters of lower temperature and salinity may be a viable mechanismof promoting metal precipitation from the former.
It has been suggested that this type of model
might apply to the formation of several sedimenthosted ore deposits such as the stratiform Cu–Co
ores of the central African Copperbelt, and thesedimentary exhalative (SEDEX) Zn–Pb ores ofMount Isa and Broken Hill in eastern Australia(McKibben et al., 1988; McKibben and Hardie,
1997).
Mixing of two fluids with distinct temperature,
pH, and redox characteristics is a process thatis likely to have widespread applicability to ore
genesis in a variety of crustal settings. Anotherimportant is the Olympic Dam described inBox 3.1.

159

tion 3.5.1 above. Quantification of the effects ofchanges in pH and redox state on, for example,
gold solubility is demonstrated in Figure 3.14,
where calculated gold solubility contours for thedominant Au(HS) −2 complex, as well as for theless important AuCl −2 complex, in a fluid at 200 °C
are shown. A reduced, neutral fluid like that at
A, subject to an oxidation event akin to the one
in reaction [3.4] above, would have its Au(HS) −2
solubility dramatically reduced, as shown by theoxidation trend in Figure 3.14. This would almost
certainly result in significant precipitation ofAu from the fluid. Similiarly, a decrease in fluid
pH from around 7 to 3 would also result in a
reduction of gold solubility and deposition of themetal. As described earlier, the precipitation ofores from metal–chloride complexes may occur
under different conditions altogether. In Figure
3.14 it is clear that the precipitation of Au from anoxidized, acidic fluid like that at B, in which the
AuCl −2 complex is stable, will best be achieved byincreasing pH (perhaps by consumption of H+ during alteration) or by reduction (perhaps caused bysulfate reduction to sulfide).
The relationship between alteration and hydrothermal ore-forming processes is a very important
one and for this reason the topic is covered in
more detail in section 3.6 below. In this section
the link between fluid/rock reactions and theprecipitation of metals from ore fluids is also
discussed.

Fluid/rock reactions (pH and Eh controls)
A widespread manifestation of hydrothermal
mineralization processes is the development ofalteration mineral assemblages in and around thefluid conduit. Alteration is caused by the reaction
between fluids and their wall rocks and is a complex process that has been extensively studied,both in ore genesis research and for the understanding of metamorphic and metasomatic (masstransfer) processes in the crust. The interactionthat occurs between a fluid and its wall rockpromotes metal precipitation because it is yet
another process that changes the prevailing fluid
properties, especially in terms of acidity (pH) andredox state. Reactions describing the factors controlling metal precipitation were presented in sec-

3.5.2 Adsorption
Although metal precipitation is generally considered to have been instigated by changes tothe prevailing fluid properties and accompanying
reduction in equilibrium solubility, it is evident
that ores can also form by adsorption of metal
onto an existing mineral surface. Metal deposition by adsorption can occur from fluids whose
concentrations are below their saturation levels
and the process may, therefore, be important incertain ore-forming environments. Adsorption
is defined as the adherence of an ion in solution to
the surface of a solid (or mineral) with which it isin contact. The process occurs because a mineralsurface will inevitably contain charge inbalances

160

PART 2 HYDROTHERMAL PROCESSES–30
100 ppb

1 ppb

B Neutralization2

Hematite

Au

Reduction

–36

Cl –

–33

Oxidation–39

Log fO2

Acidification

A

–42
1 ppb

10 ppb 100 ppb–45
Pyrite–48

Magnetite

HS) 2

Au(

Pyrrhotite

2

–

–51

4

6
pH

8

created by the fact that metal cations will notalways be fully coordinated with anions such as
O2− or S−. Sites of high charge density (either positive or negative) on a mineral surface represent thelocations where adsorption of oppositely charged
ions is likely to occur. Zones of high charge
density are represented by features such as latticedefect sites, fracture planes, and trace element
substitution sites. In general, a mineral surface incontact with an acidic solution (i.e. a high activityof H+) will contain an abundance of positive
charge and is likely to adsorb anionic complexes.
Conversely, a mineral surface in contact with analkaline solution (with a high activity of OH−)

10

12

Figure 3.14 Log fO2–pH diagram
showing the stability of iron oxideand sulfide minerals in relation
to gold solubility contours for
Au(HS) 2− and AuCl 2− complexes
in a fluid at 200 °C (also at saturated
water vapor pressure, with ∑S = 0.01
and aCl − = 1.0). Oxidation of fluid A
across the pyrite–hematite phaseboundary will result in a four orderof magnitude decrease in thesolubility of the Au(HS) −2 complex
and likely precipitation of gold.
The same result would occur if thefluid became more acidic and pH
decreased by some 5 units. The
converse would apply with respect
to AuCl −2 in fluid B. In this caseAu precipitation would occur in
response to an increase in pH(neutralization) or reduction of thefluid (perhaps caused by an increasein the activity of H2S). Phase
diagram and contours from Wood
(1998) and S. Wood, personalcommunication.

will show a tendency for a surplus negative charge
and will adsorb cations.
Experiments have shown that adsorption ofmany different metals onto the surface of a rangeof silicate, oxide, and sulfide minerals can be an
efficient way of forming an ore deposit. A varietyof parameters affect adsorptivity and the mostimportant of these are the surface area and surface
properties of the adsorbent, as well as the pH of
the solution (Rose and Bianchi-Mosquera, 1993).
Temperature tends to have an inverse effecton adsorption and for this reason it is typically a
low temperature phenomenon. Fine grained clay
particles, as well as other materials with a high

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

(a)

Ox

d

80

ed
idiz

uc
e

60

CuCl+ onto
goethite
25°C
1M NaCled

40

R% adsorption

100

20
6

5

7
pH

9

8

10

(b)
Goethite% adsorption

100
Cu

80

Pb

Zn

60
Co

Ni

7
pH

8

40
20
5

6

9

10

(c)
100

% adsorption

80

Pb
(py)

Hg
(po)

Zn
(po)

60

Cd
(sl)

40

Pyrrhotite (po)Sphalerite (sl)Pyrite (py)

20
2

4

6

8

10

12

14

161

surface area to volume aspect such as diatomaceous earth, are very efficient adsorbers and arewidely used in filtration and the removal of toxic
matter from industrial effluent. With respect toore deposition, however, it is the oxide and sulfide minerals that appear to play an importantrole in the concentration of metals from hydrothermal fluids. Since most metals form cationicspecies in nature, it follows that adsorption will
be promoted by an increase in solution pH. This
trend is demonstrated in Figure 3.15a, where theexperimentally derived adsorption efficiency of
Cu onto goethite (FeO(OH)) is plotted for oxidizing and reducing conditions. Under oxidizing
conditions, such as those applicable to the formation of red-bed hosted stratiform copper deposits(see Box 3.6), Cu will be adsorbed very efficiently
from a CuCl+-bearing solution at pH greaterthan 6 (Rose and Bianchi-Mosquera, 1993). Under
more reducing conditions Cu is only efficiently
adsorbed under more alkaline conditions. Similar behavior to Cu is exhibited by other metals
such as Pb, Zn, Co, and Ni (Figure 3.15b),
although it is clear that the onset of efficientadsorption takes place at different pH values.
Under the specific oxidizing conditions of theexperiment, metals adsorb in the order Cu, Pb,
Zn, Co, Ni with increasing pH. In a hydrothermal
solution that is evolving along an acid neutralization trend caused, for example, by wall-rockalteration (see section 3.6 below) this sequence
could account for any zonation evident in an ore
deposit.
Metal adsorption efficiency is clearly very sensitive to pH, which is, in turn, mediated by the

pH
(d)

60

AuHS°

Py

awite

% Au adsorbed

80

in
Mack

100

rite

40
Py

20
4

rrh
otit
e
6
pH

8

10

Figure 3.15 (left) Plots of percentage adsorption versus
pH for (a) copper as CuCl+ onto goethite under oxidizing
and reduced conditions and (b) a variety of transition
metals onto goethite. Solutions in both cases are at
25 °C and 1 M NaCl, and data are from Rose andBianchi-Mosquera (1993). (c) Plots of percentage
adsorption versus pH for different metals with respect
to various mineral sulfide substrates (after Jean andBancroft, 1986). (d) Percentage adsorption for gold as
AuHS onto different sulfide minerals as a functionof pH. Solutions are at 25 °C and 0.1 M NaCl
(after Widler and Seward, 2002).

162

PART 2 HYDROTHERMAL PROCESSES

1 Physical adsorpton
Van der
Waals
forcesL
Au

2 Chemical adsorpton
L

Covalent
bonding

ElectronAu
transfercerfa
Mineral su

3 Growth
Au clustere–(Zone of highcharge density)
Figure 3.16 Diagram showing the steps involved in adsorbing metal ions onto a mineral surface (after Jean andBancroft, 1985; Knipe et al., 1992). This case considers a gold–ligand (Au–L) complex that adheres initially byphysical adsorption to a negatively charged sulfide mineral surface by weak Van der Waals forces and subsequently
by chemical adsorption and covalent-like bonding to form a stronger attachment. Adsorbed gold ions can diffuse
along the surface to form clusters such as that shown in the adjacent SEM image (inset; after Jean and Bancroft,1985). Further growth of the site of metal accumulation can result in the formation of discrete gold particles thatadhere onto sulfide grains such as pyrite or arsenopyrite.

presence of surface OH−-related charge distribution. Experimental data suggest the existence ofan adsorption edge where an increase from low tovirtually complete adsorption occurs over a narrow range of pH increase. Adsorption efficiency is
not, on the other hand, much affected by the typeof mineral surface. In addition to goethite, which
is a good adsorbent under oxidizing conditions(Figure 3.15a and b), sulfide minerals also attract
metals to their surfaces in more reducing environments. Sulfide minerals are typically characterized by an intrinsic negative surface charge
usually attributed to the presence of S2− (Shuey,1975). Jean and Bancroft (1986) demonstrated thatcomprehensive adsorption of metals like Hg, Pb,
Zn, and Cd can take place under alkaline conditions onto phases such as pyrite, pyrrhotite, andsphalerite (Figure 3.15c). Mercury is, in fact, readily adsorbed onto sulfide mineral surfaces overalmost the entire pH range.
Interestingly, the role of adsorption of gold (asAuHS) from solution by sulfide minerals exhibits
the converse pattern of behavior to the basemetals. Widler and Seward (2002) showed that themost efficient adsorption of Au onto phases suchas pyrite and pyrrhotite occurs under acidic conditions, with a marked decrease in adsorbence atneutral to alkaline pH ranges (Figure 3.15d). The

reason for this particular trend is that the AuHS
complex is itself only stable at low pH, whereas
under neutral to alkaline conditions the negatively charged Au(HS) −2 complex is stable. Since a
negatively charged complex will be repulsed by anegatively charged mineral surface, adsorption
will drop dramatically as pH increases.
The actual process by which metal adsorption
occurs is very complex. Jean and Bancroft (1985)
and Knipe et al. (1992) have suggested a two-stage
process that is schematically illustrated in Figure
3.16. It is envisaged that adsorption is initiated byphysical adsorption where the metal–ligand complex is loosely held at the mineral surface by weakvan der Waals forces. There is no charge transfer
at this initial stage. The second stage of the process is chemical and is promoted by reduction,or electron gain, and formation of covalent-likebonds between the metal ion and mineral substrate. Adsorbed metal atoms may diffuse acrossthe mineral surface to form clusters at sites ofhigh charge density (Figure 3.16). Further growthof the cluster could occur either by electrochemical precipitation or by conventional precipitation
around the earlier formed metal nuclei.
Many sulfide minerals are good metal adsorbents
and this fact clearly has an important bearing on
the formation of different ore deposit types. Adsorp-

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

tion as a metal-depositing process in ore-forming
environments may have been widely underestimated in terms of its importance. Adsorption of
toxic metals onto a variety of different substrates
is also likely to become increasingly importantin combating pollution and for this reason also
warrants much closer study in the future.
As a final point, it should be noted that the
charge imbalances on mineral surfaces give rise to
the fact that many natural materials act as semiconductors. Sulfide mineral surfaces, for example,
exist as both n-type (negative or electron saturated) and p-type (positive or electron depleted)semiconductors (Shuey, 1975) and this influences
metal precipitation onto these surfaces in muchthe same way as described above.
3.5.3 Biologically mediated processes ofmetal precipitation
It is well known that the evolution of all higher
life forms since the end of the Precambrian ischaracterized by the ability of cells to interact
chemically with inorganic metals such as silicon,
calcium, and phosphorus to form shells, teeth,and bones. The interaction of microorganisms
with a variety of other metals is also important
for the formation of authigenic mineral phases
during diagenetic and hydrothermal processes.
The concept of biomineralization is increasingly
being recognized as a significant process in ore
formation, especially with respect to deposit types
such as banded iron-formations, VMS-SEDEX ores,
phosphorites, and supergene enrichments of Cu
and Au. The following section briefly considers
the basic principles of biomineralization and its
significance to ore-forming processes.
The fundamental subdivision of life into
prokaryotes and eukaryotes came about with an
improvement in the understanding of cell microbiology in the mid-twentieth century (Patterson,1999). Prokaryotes are the more primitive microorganism and the ancestral life form on Earth,developing at some stage between 4650 Ma andabout 3500 Ma when they first appear in the fossil
record. They are single celled microorganismsthat can be subdivided into two kingdoms,

163

namely Bacteria (or Eubacteria) and Archaea (orArchaebacteria). Archaea exist in extreme environments, such as volcanic hot-springs and ocean
floor hydrothermal vents (thermophiles) or hypersaline evaporitic waters (halophiles), and utilize
strange metabolic pathways, involving the energy
derived from hydrogen- and sulfur-based chemical reactions, to nourish themselves. Bacteria
include a huge variety of microorganisms thatoccupy virtually every ecological niche on Earth.
Bacteria perform many different functions: someare fermentative (Escherichia coli) whereas others
either contain the enzymes for oxygen metabolism (i.e. they are aerobic) or lack such enzymes(anaerobic).
Eukaryotes evolved from prokaryotes and contain much more complex cellular structures,
including a discrete nucleus, chromosomes, andcharacteristic cycles of cell division (mitosis andmeiosis) and sexual reproduction. They probably
appeared during Paleoproterozoic times although
they could also have formed somewhat earlier(Patterson, 1999). They are almost all aerobic andmetabolize oxygen in discrete packages within
the cell called mitochondria, which are semiautonomous chemical entities, containing DNA
that is different from that of its own cell nucleus
and more akin to prokaryotic DNA. This has led
to the suggestion that more advanced cellular
structures evolved by symbiotic inclusion of
aerobic bacteria (i.e. mitochondria) into primitive
eukaryotic entities. All higher life forms (theAnimalia, Plantae, Fungi, and Protista kingdoms)
consist essentially of eukaryotic cells.
Biomineralization
A capacity to chemically interact with metals is
an inherent characteristic of most microorganisms. Mediation by microorganisms in the formation of a variety of authigenic mineral phases isreferred to as “biomineralization” and the process
has important implications for ore formation. The
major animal phyla incorporated biomineralization as a “lifestyle” soon after their formation
during the Cambrian explosion. It is this processthat allowed the more advanced life forms to164PART 2 HYDROTHERMAL PROCESSES

develop an exoskeleton and teeth and, eventually,
internal bones (the vertebrates). It is, however,apparent that biomineralization occurred muchbefore the Cambrian Period and it has been suggested that the formation of magnetite in bacteria(the so-called magnetotactic bacteria) might represent an ancestral template for unraveling thegenetics of this process in higher life forms (Kirschvink and Hagadorn, 2000). It is the bacterial
kingdom that is implicated in many of the processes linked to biologically mediated precipitation of metals during ore formation.
Biomineralization occurs in two ways
(Lowenstam, 1981; Konhauser, 1998, 2003). The
first is termed “biologically induced biomineralization” and refers to situations where minerals
form as a consequence of the activity of microorganisms on their immediate environment.
The second process is “biologically controlled
biomineralization,” in which the microorganism
directly controls the uptake of inorganic material,
either within the cell (where elements diffuse
through the cell wall) or at the interface between
the cell and its surrounds. These two processes
are briefly described below with respect to iron,which is particularly relevant to many oreforming environments.
Biologically induced biomineralization is controlled by the same equilibrium principles thatgovern metal precipitation in inorganic chemical systems. The metabolism of microorganisms
effects the properties (i.e. pH, Eh) of aqueous solutions in their immediate environment and canpromote the precipitation of authigenic minerals
as a function of their solubility products. More
importantly, ionized cell surfaces can induce
mineral nucleation in much the same way as
discussed in the previous section on adsorption.
A wide range of different minerals can be biologically induced because the phases that form
reflect the composition of the aqueous solutionin which the microbe exists. Ferric hydroxide (orferrihydrite Fe(OH)3) is one of the most common
biominerals, forming in a variety of different environments, and is often upgraded during diagenesis
to more stable phases such as goethite and hematite. Precipitation of iron occurs in progressive

stages, from Fe-staining of the cell walls, to nucleation of small (<100 nm) ferric hydroxide grains
and eventual replacement of bacterial colonies by
the inorganic authigenic phase (Konhauser, 2003).
There are a number of bacteria that passively
induce ferric iron precipitation onto their cell
walls to form ochreous accumulations, a common
one being Leptothrix ochracea, as well as the
Crenothrix, Clonothrix, and Metallogenium genera.
Marine hydrothermal vent sites are characterized
by abundant iron-depositing microbial populations. It has also been suggested that Precambrianoceans were characterized by bacteria capable ofbinding both ferric iron and silica and that thisbiologically mediated process contributed to theformation of banded iron-formations (see Chapter 5). In the present-day surface environment, thebacterium Thiobacillus ferrooxidans is capable ofoxidizing ferrous iron, creating low pH solutions
that have environmental implications during theformation of acid mine drainage from sulfide
mineral-bearing ore deposits and mines. The
fundamental problem associated with acid minedrainage is the production of sulfate and toxic
labile metals in very acid solutions, as expressedin terms of reaction [3.5]:
FeS2 + 8H2O ↔ Fe2+ + 2SO 42− + 16H+

[3.5]

This problem can, however, also be remediated bybacterial activity, as discussed in the section on
sulfate-reducing bacteria (SRB) below.
Biologically controlled biomineralization ischaracterized by the creation of an area internal
to the cell into which ions specific to the requirements of the microorganism are diffused, thereby
creating localized conditions where concentrations are increased until saturation is reached.
Minerals are thus able to form even though theaqueous milieu surrounding the microbe would
not be able to precipitate the same phase. This
process is responsible for the formation, forexample, of linear arrays of magnetite crystals
within magnetotactic bacteria, which enables
them to navigate to preferred redox environments. The initial uptake of metal in this case is
as ferric iron, with subsequent reduction to the

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

(a)

165

(b)
Figure 3.17 Microphotographs of (a) bacterial cell with amorphous ferric hydroxide replacing the cell walls, and(b) magnetotactic spirillum bacteria showing a chain of magnetite crystals developed along the longitudinal axisof the microorganism (photographs courtesy of Kurt Konhauser and Dennis Bazylinski).

ferrous state within the cell, but prior to transport
into the magnetite bearing membrane. Fossil
magnetotactic bacteria have been recognizedfrom the 2000 Ma Gunflint Iron Formation, andthis represents the oldest evidence for controlled
biomineralization.
In terms of ore formation, the role that many
bacteria play is to catalyze the oxidation ofsoluble metals to insoluble oxide phases. One particular group of anaerobic bacteria, the so-called“sulfate-reducing bacteria” (SRB), contribute toore formation in a different way. SRB, such as
Desulphovibrio desulphuricans, oxidize organic
molecules by using sulfate as an electron acceptor, a process that generates dissolved sulfide (HS−or H2S). On diagenesis the sulfide reacts withother metals or minerals to form sulfide phasesthat are either directly implicated in mineralization, or form the precursors to later ore-forming
processes. These low temperature processes canresult in the formation of many different sulfide
minerals, including pyrite (via mackinawite andgreigite) as well as nickel, lead, and zinc sulfides(Konhauser, 2003). In a modern setting SRB can be
useful in the microbial treatment of pollution andacid mine drainage. With reference to reaction[3.5] above, SRB would catalyze the reduction of

the sulfate to form sulfide, which in turn reacts
with and stabilizes labile toxic metals. A bicarbonate ion (HCO3−) is also commonly produced
during these reactions, which has the effect ofneutralizing the previously acid waters. This typeof biological remediation of aqueous pollutants
forms the basis for many operations that have totreat both acidic and metal toxic waters.
Biomineralization is a process that has considerable applicability to many other ore-forming
environments in addition to those involving iron.
Manganese is another redox-sensitive elementthat is biogenically precipitated in much the same
way as iron. Manganous oxides such as birnessite and vernadite are the main phases formed in
the presence of bacteria. Bacterial concentrationsof Mn and Fe on the ocean floor have been implicated in the formation of manganese nodules(see Chapter 5). The formation of phosphate-richsediments has also been linked, at least in part,to biogenic mediation, with apatite overgrowths
having been observed replacing the remains of
cyanobacteria. Large-scale phosphate deposits(phosphorites; see Chapter 5) forming along theeastern seaboard of Australia have been attributed
to the progressive assimilation of sea water phosphorus by benthic bacteria (Konhauser, 2003). The

166

PART 2 HYDROTHERMAL PROCESSES

amorphous silica that forms the characteristic
siliceous sinters above epithermal gold deposits is
also attributed to the presence of a diverse rangeof bacterial genera, the filaments of which act as
nucleation sites for silicification. The siderophile
metals (Au and Pt) are also amenable to biogenically mediated concentration, particularly withrespect to the formation of nuggets in placer
and lateritic environments (Watterson, 1991; see
Chapter 4). One of the most spectacular demonstrations of the role of bacteria in the concentration of metals is derived from sulfur and leadisotopic determinations of sphalerite and galenaores from the world-class Navan deposit in Ireland.
Although also characterized by replacement andepigenetic textural characteristics, the observationof pyritized worm tubes from a related deposit at
Silvermines (Boyce et al., 1999) has shown that
the Irish-type Zn–Pb–Ba ores are SEDEX types
and related, at least in part, to syngenetic hydrothermal venting on the sea floor (see section 3.8.2).
Isotopic studies indicate that some 90% of thesulfides at Navan were derived by the bacterial
reduction of Carboniferous sea water sulfate(Fallick et al., 2001). It is argued that without anample supply of reduced sulfur the Navan deposit
might not have formed to the size that it is,emphasizing the importance of biologically mediated ore-forming processes.
Bacterial biomineralization is an extremely
important process that has wide implications forthe origin and evolution of life itself, and also
for the formation of a wide variety of authigenic
minerals in diagenetic and hydrothermal environments. In terms of ore-forming processes its role
has probably been underestimated and this is
likely to be a fruitful area of research in the future.

3.6 MORE ON FLUID/ROCK INTERACTION – AN
INTRODUCTION TO HYDROTHERMAL ALTERATION

The passage of hydrothermal solutions throughthe Earth’s crust is invariably accompanied byalteration. Alteration of a rock by the fluid passing through it is marked by the development ofa mineral assemblage that is different from theoriginal one and reflects the rock composition, as

well as the properties and amount of fluid thathas traversed the system. Zones of alteration markthe pathways of hydrothermal fluids through thecrust and may represent useful guides for the
exploration of many ore deposit types. The natureof hydrothermal alteration also provides an indication of the fluid properties associated with ore
formation and there is commonly a close relationship between the chemical reactions involved inalteration and those responsible for metal deposition. It is, therefore, appropriate that a summaryof the characteristics of hydrothermal alterationaccompanies any description of ore-forming processes. The subject is, however, a complex one andthe following section is necessarily brief. Moredetailed accounts documenting the evolution ofideas regarding hydrothermal alteration are provided in the relevant chapters of the three editionsof H.L. Barnes’s Geochemistry of Hydrothermal
Ore Deposits (namely Meyer and Hemley, 1967;
Rose and Burt, 1979; Reed, 1997).
Traditionally, alteration has been regarded as aprocess that, in its simplest form, involves boththe hydrothermal fluid itself (essentially waterwith its dissociated components H+ and OH−) andthe dissolved constituents of the aqueous solution. The low to moderate degrees of pervasive,
isochemical alteration that characterize almost
any rock evolving in the presence of a fluid canlargely be explained by hydrolysis or H+ ion metasomatism (see also Chapter 4, section 4.2.2). This
process can be depicted in terms of simple chemical equilibria, or reactions, although it must be
remembered that these are likely to be simplifications of actual alteration processes. For example,
one of the most common forms of alteration
recognized in nature is the reaction of K-feldspar
with water to form muscovite or sericite. This
reaction can be expressed as follows:/2KAlSi3O8 + H+ ⇔(k-fel)1/2KAl3Si3O10(OH)2 + 3SiO2 + K+
(musc)
(qtz)
3

[3.6]

The hydrolysis of K-feldspar to muscovite interms of this reaction requires nothing more thanthe presence of H+ ions in an aqueous solution.

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

The reaction is isochemical and no new ingredients need to be added to the system. If the reaction
proceeds to the right, H+ ions are consumed andthe fluid will become more basic, and will continue to do so until the K-feldspar is used up. The
reaction will also produce quartz as part of thealteration assemblage, as well as K+ ions dissolvedin the aqueous solution.
If the products of reaction [3.6] are permitted tofurther react with H+ ions in solution such thatthe system undergoes an increase in fluid/rockratio, then muscovite would react to form kaolinite, as follows:
KAl3Si3O10(OH)2 + H+ + 3/2H2O ⇔(musc)
3/2Al2Si2O5(OH)4 + K+
(kaol)

167

exchange because an aqueous H+ ion is consumed
and replaced by cations from the reactants. The
accumulation of aqueous base cations in the
evolving fluid will have an effect on the nature ofdownstream alteration, as they may themselves
react with wall-rock. This type of alteration process is known as cation metasomatism. A wide
variety of alteration products can form by cation
metasomatism, examples of which are:
2NaAlSi3O8 + 4(Mg,Fe)2+ + 2(Fe,Al)3+ + 10H2O ⇒(plag)(Mg,Fe)4(Fe,Al)2Si2O10(OH)8 + 4SiO2 + 2Na+ + 12H+
(chl)
(qtz)
[3.9]
and

[3.7]

This reaction illustrates the way in which increasing fluid/rock ratios will change the alteration
mineralogy by continued reaction with the rockafter one set of mineral buffers have broken down(i.e. by consumption of K-feldspar in the caseof reaction [3.6]). Alteration does not, however,necessarily result in the formation of only one
mineral reactant, but generally forms an assemblage of phases. Plagioclase, for example, occurring in the same rock that originally contained
the K-feldspar, might also react with H+ ions insolution to form pyrophyllite, as follows:
2NaAlSi3O8 + 2H+ ⇔ Al2Si4O10(OH)2 + 2SiO2 + Na+(plag)(pyroph)(qtz) [3.8]Thus, reaction of a simple mineral assemblage
with nothing but water could result in a multicomponent alteration assemblage, even if thesystem remains isochemical.
If consideration is now given to an open chemical system where the dissolved constituents ofan aqueous solution are also involved in alteration, then the processes and their products will
be still more complex. The isochemical reactions
above ([3.6], [3.7], and [3.8]) describing hydrogenion metasomatism pertain to situations wherereactants and products have similar bulk compositions. All three reactions also reflect acid–base

3CaMg(CO3)2 + 4SiO2 + 6H+ + 4H2O ⇒(dol)
Mg3Si4O10(OH)2 + 6H2CO3 + 3Ca2+(talc)[3.10]In reaction [3.9] the presence of Mg, Fe, andAl cations in a fluid which equilibrates withplagioclase will result in the chloritization ofthat mineral, with accompanying precipitation ofquartz and significant H+ production. This typeof reaction will result in the fluid becoming more
acidic during progressive alteration, which is contrary to the case for hydrolysis reactions. Reaction[3.10] demonstrates the sort of process whichtakes place when a cation-bearing aqueous solution interacts with a carbonate mineral, such as
dolomite, to form a silicate mineral, such as talc.
These processes are relevant to the formation ofskarn ore deposits where fertile granites intrude
carbonate country rocks, and are referred to as
silication (see section 3.6.1 below).
More recent considerations of alteration processes have suggested that they should be regarded
not only in terms of chemical equilibria, but also
as dynamic systems that evolve with time. Given
the time scale of most geological events, theattainment of equilibrium between rock andfluid is generally feasible, at least on a local scale.
In dynamic systems, however, it is likely thatkinetic considerations predominate over those ofsteady state equilibrium. The mineral assemblages

168PART 2 HYDROTHERMAL PROCESSES

that form as a result of hydrothermal alterationvary as a function of five factors (Reed, 1997),
namely:• temperature;• pressure;• host rock composition;• fluid composition; and• the ratio of fluid to rock in the alteration process (termed the fluid/rock ratio).
In nature these five factors interact in a complex
way to create alteration assemblages that reflect
all of them. The stabilities of mineral assemblages are dictated largely by temperature andpressure, whereas the host rock composition
controls, to a certain extent, which ingredientsare available to make up a particular alterationassemblage. Fluid composition and temperature
have a strong influence on the nature and extent
to which the host rock is dissolved by the aqueous
solution, as well as the nature and concentrationof new constituents introduced into the systemby the fluid. The fluid/rock ratio is considered tobe the “master variable” (Reed, 1997) in alteration
processes because it dictates the extent to which
reactions will move to completion.
Reed (1982, 1997) has shown that, becausehydrothermal fluids migrate along pathways that
represent evolving chemical gradients, alteration
should be viewed in a dynamic sense as a functionof changing fluid/rock ratio. This concept isschematically illustrated in Figure 3.18a, showing
a fracture along which a hydrothermal fluid isflowing and adjacent to which alteration is taking
place by diffusion-related reactions between fluidand country rock. At any instant in time, alteration pinches out downstream along the fracture
because of the progressively diminishing abilityof the fluid to react with the country rock as its
enthalpy and its dissolved ingredients are effectively “used up” (e.g. an originally acidic fluid
may become progressively neutralized as its H+
ions are consumed by acid–base exchange; seebelow). The sequence of alteration minerals outwards or downwards of the fracture or vein reflects
the sequence of reactions that relate to decreasing fluid/rock ratios (Figure 3.18a). The upstream
portion of the vein has effectively “seen” more

fluid (and is, therefore, characterized by a higher
fluid/rock ratio) than the distal sections of thesystem, and this is reflected in a variable alteration assemblage.
The evolution of alteration assemblages in a
dynamic hydrothermal system can be modeled byreacting a fluid of known initial composition, pH,
and temperature with a rock, using constraints
imposed by established thermodynamic, chemical, and equilibrium data (Reed, 1997). The fluid
composition and pH is then changed as the constituents of chemical reactions are consumed (i.e.
as a mineral buffer system breaks down). Figure
3.18b shows how the alteration assemblage
changes in a rock of andesitic composition subject
to reaction with a hydrothermal solution. The
scenario is modeled in terms of a system withhigh initial fluid/rock ratios and an acidic (lowpH) fluid, evolving to lower fluid/rock ratios andalkaline (high pH) conditions. The plot showsthat an acidic fluid circulating with high fluid/rock ratios through the andesite will be characterized by an alteration assemblage comprisingquartz + pyrophyllite + alunite, also referred toas an advanced argillic alteration assemblage(Figure 3.18b). As the alteration systematics
evolve, the same rock, now infiltrated by a morealkaline fluid circulating at a lower fluid/rockratio, would be altered, initially to a sericite +chlorite assemblage and then to a chlorite +epidote + muscovite + albite assemblage. The
latter assemblage is termed propylitic alteration
and also matches the typical greenschist facies
metamorphic assemblage. The expected distribution of ore minerals in the evolving system isshown in Figure 3.18c.
Viewing alteration assemblages in terms of dynamic fluid/rock interaction offers many advantages to the understanding of alteration processes,
as well as to explaining the relationship between
fluid rock interactions and ore deposition. The
implication of time into the models of Figure 3.18
provides a good explanation for zonation andparagenetic sequence, two hallmark characteristics of hydrothermal ore deposits, discussed in
more detail below. An evolving alteration systemthat moves from high to low fluid/rock ratios

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

169

Diffusion halo

Decreasing f/r
Fluid flow

Fracture(b)

Relative proportion

Advancedargillic

serchl0.8
0.6

Andesitic host
300°C
pH = 1

act

ca

ab
ep
par

mus

chl

0.4
0.2

propylitic
micr

pyro

al

py

anh
qz

(c)
mt

Log moles kg –1

–1

hem

po

py

–2
bar

–3

sl
cp

–4
gn

(d)

pH
–2
Log molality

Figure 3.18 Models illustrating
the dynamic nature of fluid/rockinteraction and hydrothermal
alteration as a function of changing
fluid/rock ratio. (a) Conceptualmodel showing one side of adiffusion-related alteration haloaround a fracture through which
fluid is passing from left to right.
As fluid reacts with its host rock itchanges composition and, throughtime, alteration assemblages
evolve both along and outwardof the fracture system (i.e. withdecreasing fluid/rock ratios). (b) The
occurrence and relative proportionsof the principal alteration minerals
forming when a relatively oxidized
aqueous solution (of magmatic
derivation at 300 °C and withinitial pH = 1) reacts with a rock ofandesitic composition, as a functionof changing fluid/rock ratio (ser,sericite; chl, chlorite; ca, calcite;act, actinolite; micr, microcline; ab,albite; ep, epidote; par, paragonite;mus, muscovite; pyro, pyrophyllite;
al, alunite; py, pyrite; anh,anhydrite; qz, quartz). (c) The
precipitation sequence of some of
the major ore minerals as a functionof changing fluid/rock ratio for thesame setting as in (b) (po, pyrrhotite;
mt, magnetite; hem, hematite;bar, barite; sl, sphalerite; cp,
chalcopyrite; gn, galena). (d) The
evolution of fluid pH, as well as
concentrations of certain key
ligands in the fluid, as a function ofchanging fluid/rock ratio (also forthe same setting as b). All figures
after Reed (1997).

Decreasing f/r

(a)

–4

HCO3–

2–
SO 4

pH

–

O3

–6

HC

–8

HS –

1.0

SO 2–
4

0.0
Log f/r

–1.0

1
2
HS– 3
4
5 pH
6
7
8
9
–2.0

170

PART 2 HYDROTHERMAL PROCESSES(i.e. from left to right in Figure 3.18) reflects achange from cation to hydrogen ion metasomatism,and from fluid-dominated to rock-dominated environments. The evolving system also dictates
the compositional changes that occur in the fluid
itself, such as the stepwise increase in fluid pH as
H+ ions are consumed by wall-rock reaction andstabilization of successive mineral assemblages(Figure 3.18d). The redox state of the fluid alsochanges as the alteration system evolves. Figure
3.18d shows that with decreasing fluid/rock ratio
the concentration of SO42− in the fluid decreases as
the HS− content increases. This is because SO 42−
is reduced by reaction with ferrous wall-rocks,
resulting in the precipitation of hematite in thealteration zone and fractionation of sulfide into
the evolving aqueous solution (Reed, 1997). Sulfate reduction is an important source of sulfide
in many hydrothermal ore-forming systems andpromotes base metal sulfide precipitation. It isthe changes in pH and redox state of the fluid thatare so important in promoting the precipitation of
metals from ore fluids.
There are many examples of situations wherereaction between a hydrothermal fluid and its
wall rocks is the main cause of ore precipitation.
This process can be recognized, for example, insamples that reveal textures illustrating the
replacement of primary minerals in the host rockby ore minerals. Such a texture is distinct fromsituations where ore minerals are precipitatedinto open spaces such as fractures or vugs, whereother mechanisms of ore deposition, such as
fluid mixing or boiling, might have prevailed.
Skarn deposits (see section 2.10, Chapter 2) are aparticularly good example of where alteration ofcarbonate rocks by acidic, metal-bearing solutions yields Ca–Mg silicate minerals accompanied by significant polymetallic mineralization.
Carbonate rocks do, of course, have the ability toneutralize acidic solutions and the pH increases
that arise from fluid/rock reactions in thisenvironment promote efficient precipitation of
metal sulfides (Reed, 1997). Likewise, greenstone
belt hosted mesothermal lode-Au deposits are
typically accompanied by extensive zones ofquartz–albite–carbonate–muscovite–pyrite alteration in and around the shear zone or fault along

which fluids were focused (see section 3.9.1 below
and Box 3.2). A common Au precipitation mechanism in these environments is the reaction of a
H2O–CO2, sulfide-rich fluid with ferrous hostrocks (commonly basalt or banded iron-formation) to form pyrite (Phillips, 1986). This causes
destabilization of the dominant AuHS −2 complex
and precipitation of native Au together with thesulfide minerals.
The intimate association between alterationand hydrothermal ore-forming processes is wellestablished. The recognition of alteration styles
and parageneses is now a standard tool in thearray of techniques utilized by exploration geologists and has important bearing on the nature of
the ore deposit itself.
3.6.1 Types of alteration and their ore
associations
This section outlines some of the characteristics
and ore associations of the main alterationstyles associated with a variety of hydrothermal
environments.
Potassic alteration
Potassic (or K-silicate) alteration is characterized
by the formation of new K-feldspar and/or biotite,usually together with minor sericite, chlorite,and quartz. Accessory amounts of magnetite/
hematite and anhydrite may occur associatedwith the potassic alteration assemblage. It typically represents the highest temperature form ofalteration (500–600 °C) associated with porphyry
Cu-type deposits, forming in the core of thesystem and usually within the granite intrusion
itself. A degree of K+ (cation) metasomatism in
addition to hydrolysis is considered to have takenplace during the formation of potassic alterationassemblages because studies of bulk compositions on a volume basis suggest that potassium
has been added to the system. A variation ofpotassic alteration involving substantial additionof Na and Ca (called sodic and calcic alterationand characterized by abundant albite, epidote andactinolite) is documented in some porphyry systems such as Yerington, Nevada, USA.

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

171

Alteration and metal precipitation: the Golden Mile,
Kalgoorlie, Western Australia – an Archean orogenic
gold deposit
The “Golden Mile” is a zone of exceptional gold mineralization located immediately to the east of the twin cities of
Kalgoorlie–Boulder in Western Australia. Gold was discovered here in 1893 and well over 1200 tons of gold has
been won from the area since then. Until 1975 mining
was carried out in several underground operations, asituation which led to a steady decline in production due
to increasing depths and decreasing grades. This situation was reversed in the late 1980s when ownership was
amalgamated and plans initiated to open a huge, low-grade
open-pit operation known as the Fimiston or Super Pit.
This is now the largest gold mining operation in Australia,
producing well over 800 000 oz of gold per year.
The Golden Mile is located in the Norseman–Wiluna(a)greenstone belt of the Yilgarn Craton. Auriferous mineralization is structurally controlled and hosted mainly insteeply dipping shear zones or lodes that are centered on
the Golden Mile Fault (Figure 1a). Hundreds of subsidiarylodes are concentrated in and around this system and are
hosted by a 2675 Myr old ferruginous tholeiitic intrusion(sill) called the Golden Mile dolerite. Although the initialhigh-grade, largely underground, workings were concentrated on the individual shear zones themselves, the systemis accompanied by an extensive alteration halo of low
grade mineralization which now forms the bulk of the ore
in the Super Pit. It is the alteration of doleritic host rockby mineralizing, aqueo-carbonic hydrothermal solutions
that is believed to be the main cause of the pervasive gold(b)
Quartz-pyritegold shear zone

Kalgoorlien Pyke F.
lde
Go

Eas

t Lo

de

Eastern
Lode
Systemf"SuperPit"

Lake View Shaft

er

uld

Bo

N

F.

Carbonate zoneaTr

Chlorite zone

lg

fa
F.
ar

Metasediments
Golden Miledolerite
Metavolcanics
f Golden Mile
fault
1
km

50
m

100

e F.

0

0

AdeladfWesternf
Lode
SystemFigure 1 (a) General geological outline of the Golden Mile, east of Kalgoorlie, Western Australia, and theapproximate position of the “Super Pit,” currently Australia’s largest gold mining operation. (b) The natureof hydrothermal alteration in the Golden Mile dolerite. The East Lode was originally mined as a high grade
underground deposit through the Lake View Shaft. The entire alteration zone now forms the bulk of the ore
for the high tonnage–low grade operation in the Super Pit. Both diagrams are after Phillips (1986).

172

PART 2 HYDROTHERMAL PROCESSES

mineralization in this deposit (Phillips, 1986) and this isdiscussed in more detail below.
A schematic indication of the regional geology of theGolden Mile, and the approximate position of the Super
Pit, is shown in Figure 1(a). The nature of the alteration haloaround quartz–pyrite–gold mineralized shear zones (suchas the East Lode at the Lake View mine, which has now
been subsumed by the Super Pit) is shown in Figure 1(b).
The Golden Mile dolerite was regionally metamorphosed
to a dominantly actinolitic assemblage prior to and duringfolding and ductile deformation of the Norseman–Wiluna
greenstone belt. Metamorphism was superseded by introduction of a reduced, low-salinity aqueo-carbonic fluidthat percolated along hundreds of shear zones cutting preferentially through the relatively competent Golden Miledolerite, progressively altering the host rock in and aroundthem. Alteration is markedly zoned around individual
shear zones, as well as on a broader scale, and is characterized by a paragenetic sequence that commences withearly (metamorphic) actinolite and then progresses successively to chlorite-, carbonate-(siderite), and eventuallypyrite-dominated assemblages (Phillips, 1986). Chlorite(with lesser carbonate, albite, and quartz) alteration forms
a pervasive halo extending for several hundred meters
around zones of shearing, and overprints the regional
metamorphic assemblage. Carbonate assemblages reflecta more intense manifestation of alteration (i.e. higher
fluid/rock ratio) and occur closer to individual shear zones(i.e. in halos extending up to 5 m away from the shear) or
in more ferruginous parts of the differentiated dolerite sill.
The shear zone itself is demarcated by a pyrite–quartz–
muscovite–carbonate assemblage with which high-grade

gold mineralization occurs. Phillips (1986) has argued
that the alteration paragenesis is a product of a single metamorphic fluid that introduced CO2, K, S, and Au into thesystem, and that alteration and mineralization were synchronous. Precipitation of gold is considered to have been
facilitated by the interaction of a reduced auriferous fluidwith Fe-rich host rocks, such as in the following reactions:
FeCO3 + 2H2S + 1/2O2 ⇒ FeS2 + CO2 + 2H2O
(i.e. siderite replaced by pyrite)andFeCO3 + Au(HS)2− ⇒ FeS2 + CO2 + H2O + Au
(gold bisulfide complex destabilized by pyrite formation)
This model emphasizes the role that host rock alteration
plays on metal precipitation, as described in section 3.5.1of this chapter. Although the process is undoubtedly very
important in the formation of many hydrothermal ore
deposits, it seldom takes place in isolation. At Golden Mile,for example, it is now recognized that the late-stage, highgrade gold–tellurium–vanadium ores that characterizevery rich vein systems such as the Oroya shoot were theproduct of the mixing of a (possibly magmatic) oxidized,
SO2-bearing fluid with the more reduced metamorphic
fluid that was responsible for the pervasive alteration
described above (J. Walshe, www.agcrc.csiro.au/projects).
Fluid mixing processes also appear to be broadly coincidental with the transition from a ductile to a brittle
deformational regime, a changeover that is commonly
associated with significant mineralization in orogenic
gold systems.

This alteration style is very common in a variety
of hydrothermal ore deposits and typically forms
over a wide temperature range by hydrolysis offeldspars to form sericite (fine-grained whitemica), with minor associated quartz, chlorite,
and pyrite (see reaction [3.6] and Figure 3.18).
Phyllic alteration is associated with porphyry
Cu deposits, but also with mesothermal precious
metal ores and volcanogenic massive sulfide
deposits in felsic rocks.

indistinguishable from the assemblages which
form during regional greenschist metamorphism.
It comprises mainly chlorite and epidote, together
with lesser quantities of clinozoisite, calcite,
zoisite, and albite. It is a mild form of alteration
representing low to intermediate temperatures(200–350 °C) and low fluid/rock ratios (seeFigure 3.18). This style of alteration tends to beisochemical and forms in response to H+ metasomatism. It characterizes the margins of porphyry
Cu deposits as well as epithermal precious metal
ores.

Propylitic alteration

Argillic alteration

Propylitic alteration is probably the most widespread form of alteration in that it is essentiallyThis alteration style is commonly subdivided into
intermediate and advanced categories depending

Phyllic (or sericitic) alterationHYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

on the intensity of host mineral breakdown.
Intermediate argillic alteration affects mainly
plagioclase feldspars and is characterized by
the formation of clay minerals kaolinite and thesmectite group (mainly montmorillonite). It typically forms below about 250 °C by H+ metasomatism and occurs on the fringes of porphyrysystems. Advanced argillic alteration represents
an extreme form of base leaching where rocks
have been stripped of alkali elements by very
acidic fluids active in high fluid/rock ratio environments (see Figure 3.18). It is characterized bykaolinite, pyrophyllite, or dickite (depending on
the temperature) and alunite together with lesser
quartz, topaz, and tourmaline. It is commonly
associated with near surface, epithermal precious
metal deposits where alteration is associated withboiling fluids and condensation of volatile-richvapors to form extremely acidic solutions.
Silication
Silication is the conversion of a carbonate mineralor rock into a silicate mineral or rock (see reaction[3.10] above) and necessarily involves introduction of additional components into the system(cation metasomatism). It is the main process
which accompanies the prograde stage in theformation of polymetallic skarn deposits whichdevelop when a fertile, acidic, magmatic fluid
infiltrates a carbonate host rock. Carbonate rocks
are a particularly efficient host for metal deposition from hydrothermal solutions because of
their ability to neutralize acidic fluids and their“reactivity,” which enhances permeability andfluid flow. Skarn reactions and the formationof “calc–silicate” mineral assemblages are very
complex and develop over an extended temperature range. Fluid/rock ratios also develop to very
high values in skarn systems and these factors
result in the precipitation of a diverse, polymetallic suite of ores.
Silicification
Silicification should not be confused with silication and refers specifically to the formation of
new quartz or amorphous silica minerals in a rock173

during alteration. Minor silicification develops in
the alteration halos associated with many different ore deposit types and is usually a by-productof isochemical hydrolysis reactions where Si islocally derived (see reactions [3.6] and [3.8]above). The majority of fractures through whichhydrothermal fluids have passed are at least
partially filled with quartz to form veins. The Si
in these settings is usually derived by leaching of
the country rocks through which the fluids are
circulating. Intense silicification, however, forms
as a result of cation metasomatism, where substantial addition of Si4+ in solution is added to thesystem. This type of alteration is characteristic of
the sinter zones in high level epithermal precious
metal ore deposits.
Carbonatization
Carbonatization refers to the formation of carbonate minerals (calcite, dolomite, magnesite,siderite, etc.) during alteration of a rock and is
promoted by fluids characterized by high partialpressures of carbon dioxide (PCO2) and neutral
to alkaline pH. Archean greenstone belt related
lode-gold deposits represent an ore deposit typewhere carbonate alteration is virtually ubiquitousand is accompanied by an assemblage comprisingquartz, muscovite, biotite, albite, and chlorite.
It forms when reaction occurs between a lowsalinity, CO2-rich fluid and its host rock. The
carbonate mineral that forms is a function of
the composition of the host rock and could be
dolomite in association with amphibolite, sideritein a banded iron-formation, or calcite in a granitic
host.
Greisenization
The formation of a greisen is specific to the cupolazones of highly differentiated (S-type) granites
that contain Sn and W mineralization, as well as
significant concentrations of other incompatible
elements such as F, Li, and B. Greisens represent
an alteration assemblage comprising mainly
quartz, muscovite, and topaz, with lesser tourmaline and fluorite, usually forming adjacent toquartz–cassiterite–wolframite veins.

174

PART 2 HYDROTHERMAL PROCESSES

Hematitization
Alteration that is associated with oxidizing fluids
often results in the formation of minerals with ahigh Fe3+/Fe2+ ratio and, in particular, hematitewith associated K-feldspar, sericite, chlorite, andepidote. In the magmatic-hydrothermal environment, occurrences such as the granitoid-hosted,
Olympic Dam-type Cu–Au–Fe–U deposit inSouth Australia are characterized by this style ofalteration. It also typifies other ore bodies, including the stratiform, sediment-hosted Cu–Co oresof the Central African Copperbelt. This style ofalteration appears to be related to redox processes
where highly saline, oxidizing fluids come into
contact with a more reduced host rock environment, or mix with more reduced fluids.

3.7 METAL ZONING AND PARAGENETIC SEQUENCE
A characteristic feature of many hydrothermalore deposits is the occurrence of a regular patternof distribution, or zoning, of metals and minerals
in space. Zoning can be observed at many different scales, ranging from regional patterns of metal
distribution (at a scale of hundreds of kilometers),
through a district scale, to individual ore-bodyrelated variations and even down to the level of asingle vein or hand specimen (Smirnov, 1977;
Guilbert and Park, 1986; Pirajno, 1992; Misra,
2000). A good example of regional metal zonationis the Andes metallogenic province, where the
distributional pattern of ore deposits is related
to continental scale subduction along the western margin of South America. At a district scale,
however, the distribution of metals is more
tightly constrained and better defined, and commonly characterized by a consistent pattern thatis reproduced in other ore-forming environments. District scale zoning is well exemplified
by the Sn–W–Cu–Pb–Zn–Ag–Sb–U sequence ofores associated with the Cornubian batholith insouthwest England (see Box 2.2, Chapter 2), a pattern that is reproduced in other granite-relatedore-forming systems. Zonation at this and smallerscales is clearly related to the systematic evolution of a hydrothermal fluid (i.e. alteration) andthe sequential precipitation of its contained metal
budget by the types of processes discussed in section 3.6 above. Although the processes controlling metal precipitation are complex and varied,
the relatively consistent pattern of metal zonation in so many different ore-forming environments means that the process should be amenable
to rational and unified explanation.
Although there is still some ambiguity in its
use, the term paragenesis is now widely applied tothe association of minerals and metals that characterizes a particular ore type and, therefore, has acommon origin (Guilbert and Park, 1986; Kutina
et al., 1965). Because both space and time need tobe considered in the understanding of zonation,
the phrase paragenetic sequence is now used todescribe the distribution in time of a set of genetically related minerals or metals. The recognitionand significance of parageneses and paragenetic
sequences in hydrothermal ore deposits camefrom early observations of ore-forming environments by some of the founders of modern economic geology, notably Waldemar Lindgren (1933).
On the basis of these empirical observations
Emmons (1936) devised the concept of a “reconstructed vein.” This was an idealized composite
vein that extended from deep in the crust up to
the surface and contained a paragenetic sequenceof ore components that is typical of the zoning
patterns observed in many different hydrothermal ore deposit types. The “Emmons reconstructed vein,” revised by Guilbert and Park
(1986), is summarized in Table 3.3.
Observations suggest that metals such as Mo,
W, and Sn will precipitate early (or at deep levels)from a high temperature hydrothermal solution.
Such solutions are sometimes referred to as
hypothermal. They are followed, in the ideal precipitation sequence, by Cu, and then Zn, Pb, Mn,
and Ag, as the fluid infiltrates upwards in thecrust and cools to form mesothermal solutions.
The precious and volatile metals such as Au, Sb,
and Hg are typically observed to represent the latest stages of the sequence, forming from stillcooler epithermal solutions circulating near thesurface. Extensive studies of hydrothermal ore
deposits leave little doubt that the Emmonssequence shown in Table 3.3 has general applicab-

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

175

Table 3.3 The “Emmons Reconstructed Vein” concept showing empirical observations that relate to typical
patterns of zonation and paragenetic sequences in many hydrothermal ore deposits
Vein

Metal

Ore mineralogy (in bold), and related alteration assemblages

Surface (epithermal)

BarrenMercury

Chalcedony, quartz, barite, fluorite, and carbonate minerals
Cinnabar: chalcedony, quartz, barite, fluorite, and carbonate
minerals
Stibnite: quartz
Gold, electrum, acanthite: quartz, chalcedony, adularia,
alunite, carbonate minerals, silicification, some potassic,
phyllic, and propylitic alteration
Quartz and carbonate minerals
Acanthite, rhodochrosite: quartz and carbonate minerals,some phyllic, argillic, propylitic alteration
Galena: quartz with minor carbonate minerals
Sphalerite: quartz with occasional carbonate minerals,
advanced argillic alteration
Chalcopyrite, tennantite–tetrahedrite: quartz, phyllic,propylitic, argillic alteration
Chalcopyrite: quartz, phyllic alteration
Molybdenite, huebnerite, scheelite, cassiterite: quartz,
potassic alteration
Potassic alteration, anhydrite, carbonate minerals

Antimony
Gold–silver

BarrenSilver–manganeseIntermediate (mesothermal)

Lead
Zinc
Copper–arsenic–antimony
CopperMolybdenum–tungsten–tin

Deep (hypothermal)

BarrenSource: after Guilbert and Park (1986).

ility. Many exceptions, however, do exist andthere is little doubt that the scheme can be much
improved and extended in the light of modern
observations and theory.
The wide ranging applicability of the Emmonssequence suggests that the secular precipitationof metals might simply be related to the decline
in solubility that generally accompanies temperature decrease. Sections 3.5 and 3.6 above, however, have shown that metal solubilities andprecipitation are controlled by a variety of complexly interacting variables that are related toboth the fluid itself and its surrounds. Barnes(1975) first showed that if metals are considered
to have been transported as metal–sulfide complexes then the relative stabilities of these complexes, when corrected for differing concentrationof metals in solution, closely match the Emmons
precipitation sequence. On this basis he identified
the following paragenetic sequence:
Fe – Ni – Sn – Cu – Zn – Pb – Ag – Au – Sb – Hg

By contrast, if metal transport occurs by metal–
chloride complexation then the above sequence
no longer applies and a subset of these metals is
more likely to be precipitated according to thesequence:
Cu – Ag – Pb – Zn
This suggests that paragenetic sequences arelikely to be controlled by more than just a single
variable. Susak and Crerar (1981) demonstrated
that paragenetic sequences need to be considered
in terms of classes of metals that are grouped interms of mineral stoichiometry and the valence
state of the aqueous metal species. They went on
to show that the precipitation sequence of metals
from within one particular class will be controlled essentially by the Gibbs free energiesof the associated minerals. The precipitation of
metals from different classes would, however, becontrolled by the environment within which themetal-bearing solution occurs, and this could be a

176

PART 2 HYDROTHERMAL PROCESSES

100
Zoning classes
MS (q = 2)90

MS2 (q = 2)

WS

2

70

–∆G°f (kcal mol –1)

Precipitation sequence for metals of the same zoning class

M2S (q = 1)
80

60

MnS (alabandite)50

ZnS (spha
lerite)FeS
2 (pyrite)
CdS (gree

nockite)FeS (pyrrhotite)
SnS (herzenberg
ite)

40

30

Cu2S (chalcocite)

PbS (galena)
20

Ag2S (argentite)

HgS (cinn
a

bar)

10
AuS2

0
50

100

150

200

250

T (°C)

more complex, multivariable process. Thus, asimple and reproduceable zonation pattern orparagenetic sequence is likely to apply only tohydrothermal ore-forming environments thatconsist of metals from the same zoning class. Any
exceptions and reversals to the typical Emmonssequence are likely to be explained in terms ofeither disequilibrium, or precipitation of metals
from different classes.
The predicted precipitation sequence for metals
in specifically defined zonation classes is illustrated in Figure 3.19. The precipitation sequence
for metals in solution from the same class can beread on this diagram in terms of decreasing Gibbs

300

350

Figure 3.19 Plot of Gibbs freeenergy versus temperature fora variety of sulfide minerals
categorized according to theirstoichiometry (where M is metal
and S is sulfide) and the valence
state of the dissolved metal species(q). The precipitation sequence formetals in solution from the samezoning class is predicted in termsof decreasing values of free energy(after Susak and Crerar, 1981).

free energy values for a specific temperature.
Thus, the precipitation sequence for metals in the
MS (q = 2) zoning class is Mn–Zn–Cd–Pb–Hg for
any temperature, but this sequence would notnecessarily prevail under conditions where a different stoichiometry and valence state prevails.
Replacement processes
In the previous section the precipitation of minerals and metals was discussed in the context of
aqueous solutions passing through open spaces ina rock, especially vugs and veins at shallow crustal
levels. Different conditions apply to metals that

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

are precipitated during the replacement of minerals deeper in the Earth’s crust. Replacement
occurs when the original minerals in a rock are
dissolved and near simultaneous precipitation of
secondary minerals occurs (Seward and Barnes,
1997). In contrast to the more permeable nearsurface environment where ore deposition takes
place by open space filling, metal precipitation
deeper in the crust occurs as fluids percolate
along poorly interconnected microfractures andpore spaces (see section 3.3.4). Thus, the processof replacement will only persist if porosity is
maintained and if the accompanying fluid–rockreactions are characterized by a reduction in themolar volume of the mineral being replaced. An
increase in molar volume would effectively block
the through-flow of fluid as microfractures andpores are filled by secondary mineral precipitates
whose volumes are greater than the material
being replaced.
Comparison of molar volumes shows that calcite, for example, can be replaced by most calc–
silicate and sulfide minerals, a feature that iscommonly observed in skarn deposits. By contrast,
oxide minerals such as magnetite and ilmenite
are less likely to be completely sulfidized since
the relevant reactions will involve an increase inmolar volume. At best, only partial replacementof oxide by sulfide would occur since the volume increase, together with the accompanying
decrease in microporosity that accompanies thereaction, would limit fluid flow. The sequence of
precipitation of secondary or ore minerals during
replacement processes can best be predicted interms of decreasing molar volumes of the respective solids involved, as discussed by Seward
and Barnes (1997). The mineral sequence shown
below, with relevant molar volumes, is usefulfor the understanding of replacement processes
and the prediction of replacement sequences.
Replacement will only be effective when minerals of smaller volume replace those with larger
volume. In the sequence presented below minerals
will only comprehensively replace those occurring to their left. In addition, mineral replacement
will presumably occur more readily the greater
the volume deficit between original and secondary mineral.

177

microcline ← scheelite ← anhydrite ← calcite
[54]
[47]
[46]
[37]
← kaolinite ← galena ← molybdenite ← siderite
[33]
[31]
[30]
[29]
← cinnabar ← muscovite ← arsenopyrite
[28]
[28]
[26]
← bornite ← pyrite ← sphalerite ← chalcopyrite
[25]
[24]
[24]
[22]
← pyrrhotite ← ilmenite ← hematite
[18]
[16]
[15]
← magnetite ← chalcocite
[15]
[14]
(This sequence is from Seward and Barnes (1979);[molar volumes] in cm3 mol−1.)

3.8 MODERN ANALOGUES OF ORE-FORMING PROCESSES– THE VMS–SEDEX CONTINUUM
The term “volcanogenic massive sulfide,” orVMS, refers to a large family of mainly Cu–Zn(occasionally with minor Pb and Au) deposit types
that formed during episodes of major orogenesis
throughout Earth history (see Chapter 6). They
are also referred to as VHMS deposits, where
the acronym stands for “volcanic hosted massive
sulfide.” VMS deposits occur in a variety of
tectonic settings, but are typically related toprecipitation of metals from hydrothermal solutions circulating in volcanically active submarine
environments. Sedimentary exhalative (SEDEX)
deposits, by contrast, are dominated by a Zn–Pb(with lesser Cu, but commonly Ba and Ag) metal
association and are also related to hydrothermal
fluids venting onto the sea floor, but without
an obvious or direct link to volcanism. Manyof the large SEDEX deposits of the world areProterozoic in age, although several examples,such as Red Dog (see Box 3.4) also formed in
Phanerozoic times. Although there is generally
no spatial or temporal link between SEDEX andVMS deposits, it is widely held that they represent a continuum and are conceptually linked
by the fact that they formed by the same basic
processes (Gilmour, 1976; Plimer, 1978; Guilbert
and Park, 1986; Kirkham and Roscoe, 1993; Misra,
2000). These processes are active and can be178

PART 2 HYDROTHERMAL PROCESSES

studied in modern day environments, as discussed below.
The notion of a continuum between VMS andSEDEX deposit types is, however, contentious
and readers should be wary of over-interpreting
the genetic link between the two deposit types.
The principal VMS and SEDEX deposits apparently formed at different times of Earth history
and in different tectonic settings. In addition,
studies have shown that some of the stratiform
base metal deposits, previously thought to besyngenetic and related to exhalative venting ofhydrothermal solutions onto the sea floor, aremore likely to be the product of replacement
processes and, therefore, epigenetic in origin. A
reinterpretation of this type has, for example,
recently been made with respect to the Mount Isa
deposits in Australia (Perkins, 1997). With thiscautionary proviso in mind, it is neverthelessconvenient to discuss VMS and SEDEX deposits
as a conceptual continuum, at least insofar as theprincipal ore-forming processes are concerned.
Both deposit types can be discussed in terms ofmodern analogues, which is advantageous in thatthe ore-forming processes can be studied directly.
VMS deposits are discussed in the light of thespectacular discoveries on the ocean floors of“black smokers,” whereas SEDEX deposits are
considered in terms of the rift-related hydrothermal activity in the Red Sea and also around theSalton Sea in California.
3.8.1 “Black smokers” – a modern analogue for
VMS deposit formation
Even though it had been known since the 1960s
that warm brines vented onto the ocean floor, it
was the discovery in 1979 of “black smokers” at
21° N on the East Pacific Rise (Francheteau et al.,1979) that provided spectacular insights into thenature of submarine hydrothermal activity and its
potential impact on the understanding of ore
deposition in such environments. Black smokers
are described as hot (up to 400 °C), metal charged,
reduced, and slightly acidic hydrothermal fluidsthat vent onto the sea floor, usually in zones ofextension and active volcanism along mid-ocean
ridges (Figure 3.20). The fluids originate essen-tially from cold (2 °C), alkaline, oxidizing, andmetal-deficient sea water. They circulate throughthe basaltic ocean crust and, in so doing, scavenge
metals to form the hydrothermal fluids now
observed at more than 100 black smoker sites in
the Pacific, Atlantic, and Indian Oceans as well as
the Mediterranean Sea (Scott, 1997). Black smoker
fluids usually vent through tube-like structures,
called chimneys (Figure 3.20), that are built out ofa mixture of anhydrite, barite, and sulfide minerals such as pyrite, pyrrhotite, chalcopyrite, andsphalerite, as well as gangue opaline silica.
Metal-charged fluids venting on the ocean floor
point to the hydrothermal exhalative processes
by which most, if not all, VMS and SEDEX canbe explained. Black smokers venting along midocean ridges represent a direct analogy for one
specific type of the VMS family of deposits,
namely the ophiolite-hosted Cu–Zn deposits suchas those of the Troodos Massif, Cyprus (Box 3.3). An
ophiolite represents a section of a mid-ocean ridge
spreading center that is preserved by obduction
onto continental material, and is characterized
by a sheeted dyke complex, pillowed basalt, andpelagic sediment. A lens of massive sulfide ore,comprising mainly pyrite, chalcopyrite, andsphalerite is located at the ocean floor interface
and this is underlain by a pipe-like zone of disseminated sulfides and intense chloritic alteration representing the conduit along which fluids
passed on their way to the ocean floor (Figure
3.21a). The geometry of this ore deposit type isfairly typical of most other VMS deposits (Figure
3.21b) even though they may form in other tectonic settings such as island arcs, back-arc basins,and fore-arc troughs. Other well known Cyprustype VMS deposits include Lokken in Norway
and Outokumpu in Finland. Famous examples of
VMS deposits formed in other tectonic settings
include the greenstone belt hosted Kidd Creek
and Mattabi deposits in the Archean Superior
Province of Canada, Besshi and the numerous
Kuroko-type deposits of the Japanese island arc,
Rio Tinto and Neves Corvo in the Iberian Pyrite
Belt of Spain and Portugal, the Buchans andBathurst deposits of Newfoundland, Jerome in
Arizona, and the Mount Lyell and Hellyer deposits in Tasmania.

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

179

r

ke

o
sm

Anhydrite withsome sulfide

kac
Bl Chimneys

Collapsedchimney
Anhydrite

Mainly sulfides

Impermeable
crust

Sea floor basalt
Hydrothermal

fluid

Sea water
Cold (2°C)
Alkaline (pH ~7–8)
OxidizingSO4-rich
Metal-deficient

Extension

Hydrothermal
fluid
Hot (up to 400°C)
Acidic (pH ∼4–6)
Reducing
H2S-rich
Metal-rich(Fe, Mn, Zn, Cu)

Extension

Magma

Figure 3.20 Conceptual diagram illustrating the fluid characteristics and circulation pattern in mid-ocean ridge
environments that give rise to the formation of “black smokers” on the sea floor. Inset is a cross section of an
exhalative vent site showing the construction of anhydrite–sulfide chimneys on top of a mound of massive sulfide
mineralization (after Lydon 1988; Scott, 1997).

There is now general consensus that submarine
venting represents a very good analogue for theprocesses that applied during the formation ofmost of the actual VMS deposits preserved orfossilized in the Earth’s crust. Excellent reviewsof the characteristics of actual VMS deposits
and the nature of their ore-forming processes areprovided by Franklin et al. (1981), Hekinian andFouquet (1985), Lydon (1988), Large (1992), andScott (1997). Studies of many actual deposits
confirm that the fluids involved are derived dominantly from sea water, although fluid inclusionand stable isotope studies suggest that in certain
cases a minor magmatic fluid component may
have been incorporated into the circulating fluid
system. The main source of metals is believed

180PART 2 HYDROTHERMAL PROCESSES

(a)
Pelag

ic sedimen

ts

Oxidized
zone
Massivesulfide ore
zone

Mafic
pillow
lavas

Stockworkore zone(disseminatedsulfidesand chloritic
alteration)

Sheeteddyke
complex(b)
Sulfide-bearing
epiclastic

Pyritic
siltstone

Massive sulfide

Stringer sulfide
0

m

20

Quartz-hematite +–magnetite exhalite

Beddedcarbonate-sulfide
lens

to have been the volcanic rocks through which
the sea water was percolating, and evidence insupport of this comes from mass balance considerations and the fact that metal assemblages
in different VMS deposits are consistent with the
expected metal contents and ratios in the associated primary igneous rock. For example, Cyprustype VMS deposits, which reflect leaching of adominantly mafic volcanic source rock, aretypified by a Cu + Zn metal association (Franklin,1993) and this is explained by the fact that mafic
volcanics are characterized by much higher Cu
and Zn contents than their felsic equivalents.
However, the reverse is true of Pb (see Table 1.2,
Chapter 1) and this helps to explain why the

Figure 3.21 (a) Section through atypical ophiolite-hosted, Cyprustype VMS deposit. Footwall rocks
may consist of a sheeted dyke
complex and the associatedvolcanics are often pillowed andhave a tholeiitic composition.
After Hutchinson and Searle(1971). (b) Section showing thecharacteristics of VMS deposits
other than the ophiolite-hosted
type. Associated volcanics may
be intermediate or felsic incomposition and a closer laterallink to chemical and epiclastic
sediments is often apparent. The
lens of massive sulfide ore formedon the ocean floor, underlain by astockwork zone of disseminatedsulfides and intensely alteredvolcanic rock, is typical of VMS
deposits in general (after Large,1992).

Kuroko deposits of Japan, for example, described
as a Zn–Pb–Cu variant of the VMS family (Pirajno,
1992; Franklin, 1993; Misra, 2000) and associatedwith arc-related bimodal (i.e. felsic and mafic)
volcanism, have a different metal association.
By contrast, the source of sulfur is essentiallyfrom the sulfate component of the sea water itself
with reduction of sulfate to sulfide occurring
during fluid–rock interaction prior to venting(Ohmoto et al., 1983). This is demonstrated in
the compilation by Large (1992), which shows
similar trends in sulfur isotope variations of
VMS sulfide ores and sea water derived sulfate
minerals over time (at least for Phanerozoic
deposits).

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

181

Exhalative venting and “black smokers” on thesea floor: the VMS deposits of the Troodos
ophiolite, Cyprus
The 91 Myr old Troodos ophiolite complex in Cyprus
represents an obducted and exposed fragment of oceanic
crust that has been extremely well studied. It is also
famous as a site of copper mining since the Bronze Ageand contains some 90 small to intermediate sized Cu–Zn

VMS deposits hosted in pillowed basalts, mainly along
the northern edge of the complex (Figure 1). The ophiolite
is well exposed and comprises a cross section of theocean crust, from lower mafic and ultramafic cumulates(gabbro and harzburgite that contain podiform chromitite

Mediterranean Sea

CyprusMavrovouni
Nicosia
Skouriotissa
Alestos/Memi
Larnaca

Troodos complex
Mt Olympus
PaphosLimassolFigure 1 Generalized geologyof the Troodos ophiolitecomplex in Cyprus and the
distribution of VMS styleCu–Zn mineralization.

Famagusta

Basaltic pillow lavas
Sheeted dyke complex
Gabbroic/harzburgitic
cumulates
VMS deposits (Cu-Zn)

0 10 20 30km

fHigh-grade
ore
f

f f

Sediments and"umber"Upper pillow lava

Upper pillow
lava

Upper pillow
lava
Figure 2 The nature andgeometry of VMS mineralization
in typical Cyprus deposits.
Mineralization can beconcentrated at the interface
between successive basaltic
lava flows, as shown, or it canoccur between the upper pillow
lavas and the umber sediments(after Constantinou andGovett, 1973).

Lower
pillow
lava
Low-grade
ore
0

100
meters182

PART 2 HYDROTHERMAL PROCESSES

deposits; see Chapter 1), overlain by a sheeted dyke
complex and pillowed basalts. The largest deposits are
Mavrovouni (15 million tons) and Skouriotissa (6 million
tons), although several others have been mined in the
recent past (Herzig and Hannington, 1995).
Mineralization in the Cyprus deposits comprises massive
pyrite and chalcopyrite, with lesser sphalerite, hosted inpillow lavas and associated with zones of intense silicification and chlorite alteration (Figure 2; Constantinou andGovett, 1973). The lavas are overlain by shales and cherts
that are Fe- and Mn-rich (referred to as umber), andmay also contain concentrations of Au. The hydrothermal
fluid responsible for mineralization in the Troodos VMS
ores was modified sea water that discharged throughblack smoker vents located along ridge axis parallel
faults at temperatures of 300–350 °C (Spooner, 1980).
The characteristics of the Cyprus deposits have been compared directly with the TAG site, an active exhalative ventsystem on the mid-Atlantic ridge (Eddy et al., 1998). It isalso evident, however, that VMS mineralization can beassociated with magmatic activity that is removed fromthat occurring along the ridge axis. The intrusion of discrete gabbroic plutons, and the formation of seamounts
and sea floor calderas, for example, are features that pro-

VMS deposits are characterized by well developed metal zonation patterns defined by a typicalsequence from Fe to Fe–Cu to Cu–Pb–Zn toPb–Zn–Ba in an upward and lateral sense. At afirst approximation this sequence reflects thevariable solubilities of these metals at progressively lower temperatures and could be seen as
a paragenetic sequence. In detail, however, thedevelopment of this zonation is more complex
and reflects the evolution of fluids and the growthmechanism of the massive sulfide mound withtime, as described in Figure 3.22a. In this model itis envisaged that the temperature of the ore fluid
increases with time as the deposit grows. Low
temperatures would not be able to dissolve muchin the way of base metals, although such fluids
could transport sulfate complexes and precipitate
anhydrite or barite on mixing with sea water(the so-called “white smokers” seen associatedwith low temperature vents). At temperaturesapproaching 250 °C, however, the solubilities of
Pb and Zn as chloride complexes would be high,reaching 100 ppm under the conditions applicable
to Figure 3.22b. Copper would be poorly soluble,

Figure 3 Pillow lavas exposed adjacent to a small
VMS-type Cu–Zn deposit in the Troodos ophiolite,
Cyprus (photograph by Carl Anhaeusser).
mote fluid circulation and VMS styles of mineralizationin areas that can be well away from the ridge axis. Thisobservation has obvious exploration significance, and theAlestos and Memi deposits in Cyprus are two examples ofclassic VMS type mineralization associated with off-axis
mineralization (Eddy et al., 1998).

but some Au could be transported as a bisulfide
complex. Discharge of such fluids would resultin precipitation of barite/anhydrite together
with sphalerite and galena, as well as minor gold
(Stage 1, Figure 3.22a). As fluids evolve to higher
temperatures (250–300 °C) they are capable ofcontaining significant Cu as a chloride complex
and chalcopyrite will be precipitated in the footwall stockwork zone and at the base of the massive sulfide mound (Stage 2, Figure 3.22a). In thelatter area it has been observed that chalcopyrite
replaces sphalerite. Since the solvent capacity ofa 250 °C fluid with respect to both Zn and Pb
is high, it is feasible for the earlier formed sulfide minerals to be dissolved and reprecipitatedfurther up in the mound, or distal to it. Thiszone-refining type of process accounts for thedevelopment of high grade massive sulfide cappings seen in certain VMS deposits. At still higher
temperatures (300–350 °C) both Cu and Au are insolution as chloride complexes, ultimately giving
rise to further chalcopyrite precipitation, together
with Au and pyrite. In this model the sequence
Fe–Cu–Pb–Zn–Ba is in fact the reverse of theHYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3(a)Stage 1
Massive
sphalerite,
galena183

Anhydrite or
bariteZn, Pbstockwork
150–250°C

Stage 2Massivechalcopyrite

Fe-Cu-Pb-Zn-Ba
zoning

250–300°C

Stage 3Cu stockwork(with Au)
Enriched
Pb, Zn zone

MassivepyriteAlterationzone,
disseminatedsulfides300–350°C

100

1000

10

100
Au

10

1

S

Zn

1
0.1

S
Pb

0.1

eS 2

CuF

0.01

0.01

0.001
Zn-Pb-Au zone

150

paragenetic sequence predicted simply in termsof solubility variations with declining temperature. This scheme also predicts that VMS deposits
could grow by input of metals from below and notnecessarily by sulfide particles sedimenting out

200

Cu zoneCu-Au zone

250
300
Temperature (°C)

350

Au solubilty (ppb)

Figure 3.22 (a) Secular model
explaining the evolution of fluids,
growth of massive sulfide mounds,
zoning, and paragenetic sequence
for VMS deposits. (b) Plot of
temperature versus Cu, Pb, Zn,
and Au solubilities appropriate to
fluid conditions during VMS ore
deposition (pH = 4; 1 M NaCl;
aH2S = 0.001; SO4/H2S = 0.01).
Both diagrams after Large (1992).

Cu, Pb and Zn solubilty (ppm)(b)

Pyrite zone

400

from the sea water, or aggregation of collapsed
chimneys. Finally, it should be emphasized that
biologically mediated precipitation mechanisms,
such as the involvement of sulfate-reducing
bacteria in exhalative vent processes, are likely

184PART 2 HYDROTHERMAL PROCESSES

to be important contributing factors during theformation of both VMS and SEDEX deposits.
Large-scale precipitation of sulfides by bacterial
reduction of sea water sulfate has, for example,
been suggested for the Irish Zn–Pb–Ba deposits(see section 3.5.3 above; Fallick et al., 2001).
Not all VMS deposits are characterized by aclassic lens shaped sulfide mound and a welldefined footwall alteration zone. Some have flator tabular morphologies and appear to be distalfrom a vent, showing little or no evidence of footwall alteration/mineralization. Others exhibit
indications of massive sulfide lenses stacked one
on top of the other or significant metal deposition
below the sea water–rock interface (Large, 1992).
The morphology of atypical VMS deposits maylargely be a function of the physical properties ofthe ore fluids themselves as they infiltrate up
through the footwall volcanics and then vent onto
the sea floor. Sato (1972) has suggested that submarine fluids would be either more or less dense
than sea water as a function of their temperature,salinity, and degree of mixing with cold sea water.
Figure 3.23a (Type I) shows a situation wherean ore solution is more dense than sea water,
giving rise to concentrated metal precipitationthat remains proximal to the site of venting. Thismechanism could explain the formation of highgrade tabular VMS deposits. If the ore solution
and sea water have similar densities venting will
occur but metals will be precipitated close to thesite or in nearby topographic depressions (Type II,
Figure 3.23a). A different product would result
if the black smoker fluids were much less dense
than sea water (Type III, Figure 3.23a), giving riseto buoyant metal-rich plumes that disperse metals into more distal marine sediments. Although
this is not likely to be a particularly efficient process it remains a possible mechanism for explaining distal, low grade deposits associated with ahigher proportion of ocean floor sediment.
Another factor that may be important inconsiderations of exactly where in the system
metals accumulate relates to whether or not
the hydrothermal solutions undergo boiling. The
phase change from liquid to vapor is strongly
pressure-dependent and even though many blacksmoker fluids have been observed venting attemperatures as high as 400 °C, they appear on
the sea floor as fluids because of the high confining pressures that exist deep in the oceans. Thisis confirmed in Figure 3.23b, where the boilingpoint curve for sea water is plotted relative todeep (represented by 21° N East Pacific Rise) andshallow (represented by the Juan de Fuca ridge)
ocean floor settings. Any 350 °C hydrothermal
solution rising adiabatically through the crust
would intersect the deep ocean floor as a liquid
since it would not yet have intersected the boilingpoint curve. Metal precipitation on the sea
floor from the black smoker would, in this case,occur by rapid quenching of the ore solution as it
mixed with a virtually limitless volume of cold
sea water. By contrast, the same fluid moving up
toward a shallower sea floor is likely to intersect
the boiling point curve before it vents to surface(Figure 3.23b). Boiling below the ocean floor is
likely to promote brecciation of the footwalland local metal precipitation, possibly resulting
in extensive stockwork mineralization. Classic
VMS deposition is unlikely to occur and ore
bodies may appear stacked one on top of the other.
Although there are many variants in the familyof VMS ores the processes applicable to thesubmarine exhalative model appear capable ofexplaining many of the features of this very
important class of deposits.
3.8.2 The Salton Sea and Red Sea geothermal
systems – modern analogues for SEDEX
mineralization processes
SEDEX deposits contain more than half of theworld’s known resources of Pb and Zn andare generally represented by bigger and richer
deposits than the VMS category, although there
may be fewer of them. They are typically formed
within intracratonic rift basins and are hosted bymarine clastic or chemical sediments with little
or no direct association with volcanic rocks. Mostof the world’s giant SEDEX deposits are Paleoto Mesoproterozoic in age and well known
examples include HYC (McArthur River), Mount
Isa and Broken Hill in eastern Australia, Sullivanin British Columbia, Aggeneys and Gamsberg
in South Africa, and Rajpura-Dariba in India. A

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

185

(a)
0

Temperature (°C)

I

3m

200

2.5 m
1m

300

II

Type IIIII

Density ofseawaterat 10°C
5m

2m
0.85

Type III

0.90
0.95
1.00
Density (g cm–3)

1.05

(b)
Shallow water
Pb–Zn–Au–Ag

Cu–Pb–Zn

S ea watboiling er (3.5curve wt%
Liq
uid

0

Cu

Va
p

or

Na

100

1000

)
Cl

Sea floor
Boiling and metal deposition
below sea floor

200

2000

“Black smoker” and metal

Sea floor deposition on sea floor
21°N

3000

300
Deep water

400“Critical
point”
sea water
Adiabatic upriseof 350°C hydrothermal
fluid

500
0

significant group of deposits is Paleozoic in age
and these include Meggen and Rammelsberg inGermany, Navan in Ireland, Red Dog in Alaska,
and Howard’s Pass in eastern Canada.

100

Depth (meters)

Juan de Fuca Ridge
Pressure (bars)

Figure 3.23 (a) Plot of temperature
versus density for a variety ofpossible hydrothermal solutions(where 1 M, 3 M, etc. refer to the
NaCl molality or the salinity ofthe solution) relative to sea waterat 10 °C. Different solutions may beeither less or more dense than sea
water. At one extreme they would
form buoyant plumes capable of
distributing metals some distanceaway from the site of venting (TypeIII), whereas at the other they would
precipitate metals proximally tothe vent (Type I). After Sato (1972).
(b) Pressure (or depth) versus
temperature plot showing the sea
water (with 3.5 wt% NaCl) boilingcurve relative to deep (21° N East
Pacific Rise) and shallow (Juan de
Fuca ridge) sea floor. Because of
the strong pressure dependency
on boiling, a 350 °C hydrothermal
fluid rising adiabatically in thecrust would intersect a deep ocean
floor as a liquid and would vent as a
black smoker precipitating metals
into the sea. The same fluid relative
to a shallower ocean would boil
before intersecting the floor suchthat only a vapor phase would reachthe surface. Metal precipitation insuch a case will likely occur
beneath the ocean floor. After
Delaney and Cosens (1982).

Type I

100

200
300
Temperature (°C)

400

4000

5000
500

Modern analogues for the genesis of SEDEX
deposits have been described from the Salton
Sea and Gulf of California (Figure 3.24) andalso the Red Sea. The study of these hydrothermal

186

PART 2 HYDROTHERMAL PROCESSES

systems reveals both syn-sedimentary exhalativeand replacement processes and can, to a certain
extent, be used to discount the often controversial
views that prevail regarding the origin of sedimenthosted stratiform deposits.

Pacific
plate
30

SAN
ANDREAS
FAULT

(a)

North
American
plate

°

Gulf
of
California

Salton Sea geothermal system

USA
MexicoN

EAST
PACIFIC
RISE

0

km

500

(b)
Primary transform fault
Active transform extension
Inactive fault zone
Spreading centerult
San Andreas fa

SALTON
SEA

Salton Sea
geothermal system

U
Me SA
xic
o

Cerro Prieto
VolcanoNA
US icoxMeradoColoiverR

Basement
outcrops
020
km

The East Pacific Rise intersects the North
American plate at the head of the Gulf ofCalifornia (Figure 3.24a). At this point the prevailing stresses along this mid-ocean ridge are transferred onto the continent to form the essentiallydextral San Andreas fault system. The fault cuts
through a dry intermontane region called Imperial
Valley, which is characterized by active, riftrelated sedimentation. Early settlers discovered ahuge salt pan in this valley, named the Salton
Sink, that was mined for a few years in the late
1800s. In 1905, however, a canal, designed tobring irrigation water from the Colorado River tosettler farmers in Imperial Valley, was breachedby spring floods and the entire volume of thislarge river was channeled into the Salton Sink, atthat time close to 100 meters below sea level. It
was two years before the breach was plugged andthe Colorado River reverted to its natural course.
By that time the salt pan had become the SaltonSea, a huge inland lake that covered an area ofseveral hundred square kilometers. The Salton
Sea soon became brackish as the lake waters
dissolved the salt on its floor. Evaporation also
increased the salinity of Salton Sea waters to
the Na–Cl–SO4 dominated composition that isobserved today (McKibben and Hardie, 1997).
Over the past century water from the Salton
Sea has been percolating downwards along faultrelated pathways and through earlier formed
lacustrine sediments to form an extensive hydrothermal system from which geothermal power is
now derived. In addition, hot metalliferous brines,
also derived directly from Salton Sea waters andheated by a combination of the high prevailing
Figure 3.24 (left) The Salton Sea Geothermal System(SSGS) and its location (inset) relative to the Gulf ofCalifornia transform fault region and the dextral San
Andreas fault system (after McKibben and Hardie, 1997).

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

geothermal gradient and active volcanism in
the region, vent from numerous sites, and it isthese that are of particular interest to the studyof ore-forming processes (White, 1963; Skinneret al., 1969). The combination of well understood
fluid compositional and circulation pathways,
set in an active seismic and volcanic, intracontinental rift environment, has made the Salton
Sea geothermal system a classic example of anore-forming environment whose evolution can beobserved, and from which many inferences about
ancient sediment-hosted ore-forming processes
can be made. In addition to its obvious applicability to the understanding of SEDEX ore-forming
processes, the processes observed in this environment might also be relevant to the formationof stratiform red-bed hosted Cu deposits such as
the Kupferschiefer ores of central Europe and theCentral African Copperbelt (McKibben et al.,
1988; McKibben and Hardie, 1997).
Although the Salton Sea geothermal systemdoes not exist in a submarine environment, thegeometry of the system and the processes involved
are considered to be useful for the understandingof sediment-hosted hydrothermal base metal
deposits in which fluids circulated during andsoon after deposition. A variety of conditions needto prevail in order to make SEDEX ore-forming
processes viable. These include an abundant fluid
source, a rich source of metals and complexing
agents, active rifting, and deformation induced
fluid flow. Detailed study of geothermal wells
around the Salton Sea has shown that at depthsof 1000–3000 meters a hot (350 °C), dense (up to
26 wt% dissolved solids), Na–Ca–K–Cl brine exists(McKibben et al., 1988). This fluid also contains
significant concentrations of metals, in particular
Fe, Mn, Pb, Zn, and Cu, derived from local lacustrine sediments. Precipitation of these metals has
been induced by mixing of the hot metal charged
brine with cooler, dilute surface waters about1000 meters below the surface, to form sedimenthosted base metal sulfide veins and associatedchloritic alteration. This is clearly not an exhalative system and the ores that form will appear
epigenetic in character. Ores that form in thisenvironment could also have a stratiform geometry, with metals being deposited by replacement

187

of pre-existing phases such as diagenetic cements
or easily dissolvable detrital minerals. In the
Salton Sea setting it is also conceivable, especiallyin the event of very energetic fluid circulation,that the metal-charged brine could intersect thelake floor and vent ore-bearing solutions in muchthe same way that black smokers are exhaled along
the mid-ocean ridges. It seems feasible, therefore, for SEDEX ore bodies to appear syngenetic
and exhalative in sediments that were depositedat the time of fluid venting, but also distinctly
epigenetic relative to the footwall sedimentary
sequence. A model illustrating the main featuresof SEDEX deposits and incorporating the conceptthat syngenetic exhalative and epigenetic replacement type ores could be coeval and form part ofthe same system is presented in Figure 3.25.
The Red Sea and the VMS–SEDEX continuum
The extension of the East Pacific Rise (mid-ocean
ridge) onto continental North America to form
the San Andreas fault system (Figure 3.24) provides support for the view that a conceptual continuum exists between VMS and SEDEX stylesof mineralization. Present day ocean floor blacksmokers at 21° N on the East Pacific Rise andthe active, metal-charged geothermal systems of
the Salton Sea perhaps represent extremes in therange of VMS and SEDEX ore-forming systems.
The Guaymas Basin in the Gulf of California
is also known to contain terrigenous sedimenthosted styles of VMS mineralization, also known
as Besshi-type ores, from the type location in
Japan. These intermediate styles of mineralization also provide support for the view that VMS
and SEDEX styles of base/precious metal mineralization can be linked in terms of both tectonic setting and hydrothermal systems. Perhaps the best
example of the conceptual link between VMS andSEDEX deposit types, however, is provided by
the Red Sea rift, separating Africa and Arabia. The
Red Sea is an ocean basin in the early stages of
its development, containing active geothermal
venting and deposition of base metal deposits
in clayey muds accumulating in depressions on
the sea floor (Bischoff, 1969). Mineralization isparticularly well developed at a site known as the188PART 2 HYDROTHERMAL PROCESSES

Fluid
ingress
f

f

Fa

rp

s
ult

ca

Exhalative ore

Distal sediments,
barite, chert
Vent
complex

Replacement ore
Fault breccia andfeeder pipe

Permeable unitsfSyn-sedimentary
growth fault
fFigure 3.25 Diagram illustrating
the setting for the formation of
SEDEX-type Pb–Zn ores and a
scenario which incorporates
both exhalative and replacement
concepts for the formation ofthese ores (after compilations
by Goodfellow et al., 1993;
Misra, 2000).

Sedimentary exhalative (SEDEX) processes:
the Red Dog Zn–Pb–Ag deposit, Alaska
The Red Dog mine in northwest Alaska is presently theworld’s largest producer of zinc concentrate, contributing
about 6% of annual global supply. It is a very rich deposit
and in 2002 had reserves of close to 100 million tonsof ore grading 18% zinc, 5% lead, and 85 g ton−1 silver(www.teckcominco.com). It was discovered in 1968 by
an astute bush pilot who noticed a color anomaly in therocks caused by oxidation of the sulfide ores.

The Red Dog ores are hosted in Carboniferous blackshales and limestone of the Kuna Formation in the DeLong mountains of the Brooks Range, northern Alaska.
The host succession has been deformed into a sequenceof thrust allochthons during a Jurassic orogeny. Figure 1
shows an interpretive section through a Devonian–
Carboniferous sequence of sediments in the Brooks Range
prior to thrusting. The sequence of sediments, including

Sea level
Red Dogsequence?

Oceanic side

Devonian–Carboniferous boundary
Chert

Sandstone

Limestone

Conglomerate

Continental side

Black shale

Figure 1 Interpretive, pre-thrusting, cross section through Devonian–Carboniferous sediments in the Brooks
Range, illustrating the depositional setting of host rocks at the Red Dog mine. Syn-sedimentary growth faults
and abrupt lateral facies variations are also a feature of the regional setting (after Moore et al., 1986).

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

Turbidites

Red Dog
f

f

f

Ka
ya
No
k
at
sh
ak
al
sa
e
nd
st
on
e

Site ofexhalative
venting

189

f

Figure 2 Simplified block
diagram showing theenvironment of ore formationin the Carboniferous for the
Red Dog deposit (after Moore
et al., 1986).

Kivalina
unit

limestones and turbiditic assemblages, is consistent withdeposition in a restricted ocean basin, with no indicationof volcanism in the immediate environs. Abrupt lateralfacies changes in the succession point to the presence of
syn-sedimentary growth faults that would have facilitatedfluid circulation through the basin.
Mineralization is stratabound and, despite the thrusting,is relatively flat-lying (Moore et al., 1986). It is marked
by laterally extensive zones of silicification (chert andchalcedony) within which poorly bedded sulfide minerals(sphalerite, pyrite–marcasite, and galena) occur. Barite
is closely associated with the ore and tends to cap themineralized zones. An interesting feature of the ore zones
is the occurrence of sinuous cylindrical structuresinterpreted to be worm tubes, similar to biotic structures
observed in present day exhalative vent environments.
This evidence supports the notion that silica and sulfide
precipitation took place at the rock–water interface on theocean floor and was, therefore, coeval with the deposi-

Atlantis II Deep where muds are estimated to contain around 50 million tons of rich Zn–Cu–Pb–
Ag–Au mineralization. Although the geologicaland tectonic setting of the Red Sea mineralization
is analogous to a mid-ocean ridge VMS setting,the absence of black smokers and the presence ofboth clastic and carbonate sediment hosts to themineralization are features perhaps more akin to
SEDEX ores (Pirajno, 1992).

tion of the sediments. There is also, however, evidence forsulfide replacement of sediments at Red Dog, indicatingthat the dual processes of exhalation and replacement
were taking place (see section 3.8.2 of this chapter).
Figure 2 presents a simplified model of the ore-forming
environment at Red Dog. Fluids circulating through
Carboniferous organic-rich marine sediments scavenged
metals from these host rocks and were vented onto theocean floor at sights created by syn-sedimentary growthfaults. Mineralization at Red Dog is massive, brecciated,and poorly bedded, suggesting that sulfides were precipitated fairly close to the vent sites (Moore et al., 1986).
Variable sulfur isotope signatures in sulfide minerals suggest that precipitation was a product of mixing a buoyant
hydrothermal plume with sea water that was becoming
progressively more oxic over time. It is suggested thatmineralization ceased when the sea water became sufficiently oxidizing to favor barite precipitation over sulfides(Moore et al., 1986).

3.9 MINERAL DEPOSITS ASSOCIATED WITH AQUEOCARBONIC METAMORPHIC FLUIDS

A significant proportion of the world’s gold
deposits are genetically linked to the formation ofmetamorphic fluids that are typically characterized by low salinity, near neutral pH, and mixed
H2O–CO2 compositions. These gold deposits, ofwhich there are many different types that formed

190PART 2 HYDROTHERMAL PROCESSES

MesoscaleOre deposition

First orderfault zone

Megascale
Paleosurface

e e
nit n
ra terragSecond/n ethird orderea tonCarbonation fault zonehscchloritization
Ar reensulfidationgCurrent surface

Metamorphic
devolatilization

Oxidizedmagma?

Brittle-ductile transition

1 Phaseseparation

H2O–NaCl–CO2+CH4
fluid inclusions

2 Fluid-rockreaction(e.g. sulfidation)

Fault
Internal granitoids

Metamorphic
devolatilization

Granulite
facies

Proximalalteration
zone
Distalalteration
zone

Figure 3.26 Schematic illustrations showing the principal features of Archean orogenic gold deposits – many ofwhich also apply to Proterozoic and Phanerozoic examples. The main features at a megascale are that the deposits
are associated with a convergent plate margin where metamorphism results in the production of fluids that arefocused along major structural discontinuities. At a mesoscale major shear zones within or along the margins ofgreenstone belts represent district-scale hosts to mineralization that are accompanied by broad zones of intense
alteration. Individual deposits occur along second- and third-order structures. Ore deposition is a function of fluidrock reaction and/or H2O–CO2 phase separation (after Hagemann and Cassidy, 2000).

over the entire span of Earth history, are associated with regionally metamorphosed terranes
resulting from compression at convergent platemargins, and have been termed orogenic gold
deposits (Groves et al., 1998). They are also widelyreferred to, especially in the older literature, as“mesothermal” or “lode-gold” deposits, as they
commonly precipitated at around 300 °C, forming
quartz-vein or fracture dominated ore systems.
Although orogenic gold deposits are commonly
associated with Archean granite–greenstone terranes (see Chapter 6), important examples of orogenic gold ores are also found hosted in a varietyof Proterozoic and Phanerozoic settings. The dominant and characteristic genetic features that link
all orogenic gold deposits are a synchroneitywith major accretionary or collisional orogenic
episodes and the production of metamorphic –and in some cases magmatic – fluids that precipitate metals at various crustal levels along

deep-seated shear and fracture zones (Figure 3.26;
mega- and mesoscale models). It is widely recognized that these deposits were formed over arange of P–T conditions, occurring in granulite togreenschist facies host environments and ductile
through to brittle structural regimes (Colvine,
1989; Groves, 1993).
In the sections that follow, a variety of orogenic
gold deposits from throughout geological time
are described. In addition, Carlin-type deposits, aswell as the hydrothermal component of mineralization in quartz pebble conglomerate hosted goldores such as the Witwatersrand and Tarkwaian
deposits, are described, as they share several common features with orogenic gold deposits.
3.9.1 Orogenic gold deposits
There are many different types of orogenic gold
deposits and their classification into a single

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

category is bound to be controversial. From a
descriptive and organizational point of view, however, the classification is appropriate and recognition of the different subtypes that undoubtedly do
exist can perhaps best be made in terms of geological time and crustal evolution.
Archean
The Archean was a major gold metallogenic
epoch and many important gold deposits occur
in the granite–greenstone terranes of the world,
examples of which include the Superior Province of Canada, the Yilgarn Craton of Western
Australia, and the Zimbabwe Craton. A feature ofArchean orogenic gold deposits is that they can behosted in a variety of greenstone belt associatedlithologies, including mafic volcanics, metasediments, and banded iron-formations, as well as infelsic plutonic rocks. Despite the range of hostrock lithologies, they are invariably found associated with zones of high strain that are manifest
as brittle, brittle–ductile or ductile deformation,
depending on the crustal depth at which fluids
were circulating. Reviews of the nature and characteristics of Archean deposits can be found in
McCuaig and Kerrich (1998) and Hagemann andCassidy (2000). Figure 3.26 illustrates the mainfeatures of Archean orogenic gold deposits. At amegascale the energy required to form huge zones
of fluid flow, alteration, and mineralization, suchas those preserved in the Abitibi (Superior) and
Norseman–Wiluna (Yilgarn) greenstone belts, isultimately derived from plate subduction andcollision. Crustal thickening, deformation, metamorphism, and synorogenic magmatism all play
important roles in the ultimate origin of fluids
and their focused flow upwards into the crust. A
major part of the fluids implicated in mineralization appears to have been derived from thedehydration that accompanies regional prograde
metamorphism. These fluids are fairly reduced(fO2 around the quartz–fayalite–magnetite or
QFM buffer) and gold is preferentially transported
as the Au(HS) −2 complex (Hagemann and Cassidy,
2000; Ridley and Diamond, 2000). Derivation ofmineralizing fluids from oxidized granitic magma
is also considered likely, although such fluids

191

are probably restricted to environments where aspatial link between mineralization and granite
intrusion is evident, or an association between
Au and metals such as Mo, W, or Cu occurs(Burrows and Spooner, 1987). Gold mineralizationis broadly synchronous with the regional peak ofmetamorphism, although this peak may be in thelower crust rather than at the site of depositionin the upper crust. Fluids tend to be focusedalong major structural discontinuities that havestrike lengths of tens to hundreds of kilometersand extend to considerable depths in the crust.
Examples of well mineralized structural systems
include the Larder Lake–Cadillac fault zone in the
Abitibi belt and the Boulder–Lefroy fault zone in
the Yilgarn craton. Individual deposits and mines
are typically found along second- and third-order
structures within these major discontinuities.
Movement of fluid along these structures is
explained by the fault-valve model, as discussed
in section 3.3.3 above. The sites of mineralization
are marked, at a regional scale, by zones of intensecarbonation and chloritization that can extend
for hundreds of meters around the conduit (seeBox 3.2). Alteration plays a major role in metal
precipitation, and this is particularly the casefor intense wall-rock alteration and sulfidization.
H2O–CO2 phase separation is also considered to
be an important process in gold deposition andmay explain the siting of rich accumulationsof gold in quartz veins (i.e. the bonanza lodes).
The latter mechanism may also explain the common association of gold mineralization with thebrittle–ductile transition, as this zone coincides
broadly with the relevant P–T conditions alongparts of the H2O–CO2 solvus (see Figure 3.5c).
Mixing of metamorphic with meteoric fluids
may explain gold precipitation at high levels in
the crust.
Proterozoic
The Proterozoic Eon is not as important as either
its Archean or Phanerozoic counterparts in termsof gold mineralization (see Chapter 6). There are,nevertheless, several important examples of orogenic gold deposits from this period of time andthese include Ashanti–Obuasi and several other192

PART 2 HYDROTHERMAL PROCESSES

deposits associated with the Birimian orogenyof West Africa, Telfer in Western Australia,
Homestake in South Dakota, USA, Omai on the
Guayanian craton of South America, and the
Sabie–Pilgrim’s Rest goldfield in South Africa.
All these examples are associated with periodsof major orogenesis and largely involve metamorphic fluids very similar in character to those
applicable to the formation of Archean orogenic
gold deposits. Gold deposition is generally lateorogenic and is hosted in major, high-angle thrustfaults (Partington and Williams, 2000). Granite
intrusions, and possibly a magmatic fluid component, are associated with many of the Proterozoic
deposits and, in cases such as Telfer (Rowins
et al., 1997) and Sabie–Pilgrim’s Rest (Boer et al.,1995), a polymetallic association of Cu, Co, andBi, in addition to Au, is evident. In general,
however, the characteristics of Proterozoic orogenic gold deposits are very similar to those of theArchean.
Phanerozoic
Orogenic gold deposits of the Phanerozoic Eon
were formed during two mineralization episodes,
one in the Silurian–Devonian at around 450–
350 Ma and the other in the Cretaceous–
Paleogene at around 150–50 Ma (Bierlein andCrowe, 2000). These gold deposits are again associated with convergent plate margins and hosted
in compressional to transpressional shear zones
that typically cut through thick marine shale
sequences that have been metamorphosed togreenschist facies grades. They are often referred
to in the older literature as “slate-belt hosted” gold
deposits. There are many important examplesof Phanerozoic orogenic gold deposits and theseinclude, as Paleozoic examples, Bendigo, Stawell,and Ballarat in the Lachlan orogen of southeastern
Australia, Haile in the Carolina slate belt, USA,
Muruntau in Uzbekistan, the Urals of central
Russia, and the Meguma terrane of Nova Scotia,
Canada. Younger, Mesozoic examples include theJuneau gold belt of southern Alaska and otherportions of the Cordilleran orogen of Canada
and the western USA, including the Mother Lodeof California (Bierlein and Crowe, 2000).

3.9.2 Carlin-type gold deposits
Although several of the features of Carlin-typeores mirror those of orogenic gold deposits, they
are generally classified separately because of onevery notable difference. While linked to orogenic
activity and hosted along major, deep-crustal
structures, they are not related to compression
but formed after the onset of extensional forcesthat followed earlier, subduction-related processes (Hofstra and Cline, 2000). Carlin-type ores
are located in Nevada, USA, and formed over ashort period of time in the Paleogene Period,
between 42 and 30 Ma. In this sense they are
regarded as an essentially unique deposit type,although it is likely that similar ores will belocated elsewhere, such as in China. In Nevada
they are hosted in Paleozoic carbonate rocks thathave been subjected to severe crustal shortening
and west to east directed thrusting (Figure 3.27aand Box 3.5). The carbonate sequences are structurally overlain by less permeable siliciclastic
rocks and have acted as aquifers along whichmetamorphically derived fluids have flowed.
During mid-Miocene times host rocks were subjected to uplift and extension, which thinned
the crust and gave rise to the Basin and Range
faulting and associated fluid circulation along
the newly formed dilatant structures.
Fluids associated with Carlin deposits are very
similar to the metamorphic fluids implicated inorogenic gold deposits, comprising low salinity,
moderately acidic, reduced H2O–CO2 solutions
in which gold is transported as either Au(HS) −2 orAu(HS). Fluid temperatures are cooler than those
associated with orogenic ores and fall in the range
150–250 °C, which has led to the suggestion thatthey may, however, be meteoric rather than metamorphic in origin. Gold deposition occurs where
normal faults intersect a less permeable cap rock,usually at a shale/limestone contact and in the
crests of fault-propagated anticlines (Figure 3.27b).
The precipitation mechanism is believed to berelated to neutralization of the ore fluid during
carbonate dissolution, although argillic alteration,
silicification, and sulfidization also played a rolein the ore-forming process. In Carlin-type deposits
the gold occurs as micrometer-sized particles or

(a)

120°

118°

116°

Oregon

112°

Idaho

Ma
g

ma
ti

G
et
ch
e

ll

ca

rc

42°

114°

Ba

ttl

40°

e

Ca

rli

n

Easternlimitofmagmatism

Mt

Nevada
Eu
r

ek

lt

a

thru
s

100
ts

er

Be

ro

rO

ie
ev

t

M

ic

n

ge

t

nda
lgo
Go hrust
t

S

b
Ro

Utah

California

Arizona0(b)nressio

omp
rly c

hon

Twin Creeks

Ea

ocht
da all

n

Golco

thon

lloch

ain a
ount

Carbonate sequence

rts M

Robe

Proterozoic–Cambriansequence

Basin andRange
faults

Oceanic crustArchean crust

Later extension

Figure 3.27 (a) Map showing the distribution of Carlin-type gold deposits in Nevada and the geological framework.
(b) Cross section illustrating the fundamental controls behind the location of Carlin-type gold deposits (after Hofstra
and Cline, 2000).

194PART 2 HYDROTHERMAL PROCESSES

Circulation of orogeny-driven aqueo-carbonic fluids:
Twin Creeks – a Carlin-type gold deposit, Nevada
Twin Creeks mine, formed in 1993 after the merger of
the Chimney Creek and Rabbit Creek mines, is one of thelargest Carlin-type deposits in Nevada, with reserves in 1999
of some 90 million tons of ore at an average gold grade of
2.5 g ton−1. Most of the gold at the mine is extracted fromthe huge Mega Pit which produces some 280 000 tons ofore per day. Although gold at the Mega Pit is hosted essentially in limestones of the Ordovician Comus Formation,mineralization elsewhere on the mine (i.e. the Vista Pit)
also occurs in overlying sequences that form a package ofthrusted allochthons (see Figure 3.27a and b) emplaced in
Carboniferous times. Gold mineralization occurs mainlyin the arsenic-rich rims to pyrite grains formed during hydrothermal alteration (i.e. decalcification, dolomitization,0

silicification, and sulfidization) of the Comus Formation
limestones (Thoreson et al., 2000).
In the Mega Pit the mineralized limestones are cappedby the Roberts Mountain thrust which folded the underlying limestones and also possibly acted as a cap that
limited further egress of hydrothermal fluids and concentrated mineralization within the Comus Formation.
Figure 1 shows a cross section through the Mega Pit, illustrating the relationships between the early thrusting andfolding of Comus limestones, and the later extensional
faults along which mineralizing fluids are thought tohave been focused. In this section line gold mineralization is concentrated largely within the core of a prominent
eastward verging recumbent antiform.

100
metersUpper Sill

Mega Pit
outlinePaleocene cover

Roberts
Mountain thrustf
Upper Sill

f

f

ult
Fa

Gold mineralization(in fold core)

DZ

Wes
t

Side

Fau

lt

Main Sill

t

ll

ul

Upp
e

r Si

y
st

f

k
ac
Bl

Fa

Comus
Formation(folded during
thrust event)NafFigure 1 Cross section looking north of the Mega Pit at the Twin Creeks mine (after Thoreson et al., 2000).

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

Mineralization at Twin Creeks, as elsewhere in the Carlin
district of Nevada, is considered to have been a resultof fluid circulation between 42 and 30 Ma (Eocene–
Oligocene) during the onset of extensional (Basin andRange) tectonism and mantle plume activity (Hofstra andCline, 2000). Gold deposition, therefore, took place longafter deposition of the host rocks (in the Ordovician) andtheir subsequent deformation (thrusting and folding) in
Carboniferous times. Although subduction-related magmatism took place in the region between 43 and 34 Ma,
no intrusives are directly implicated in the formationof any Carlin-type deposit and most workers refer to ametamorphic fluid as the mineralizing agent. This fluid
was a moderately acidic, reduced, H2O–CO2 solution ataround 150–250 °C in which gold was transported asbisulfide complexes (Hofstra and Cline, 2000). Gold precipitation is believed to have accompanied cooling andneutralization of the fluid as it interacted with reactive

in solid solution in arsenic rich pyrite, marcasite,or arsenopyrite.
3.9.3 Quartz pebble conglomerate hosted
gold deposits
Quartz pebble conglomerates of Archean andPaleoproterozoic ages represent a very importantsource of gold, although production over the past
century has been dominated by the WitwatersrandBasin in South Africa. The seven major goldfields
associated with the Witwatersrand sequences haveproduced a staggering 48 000 tons of gold metal –some 35% of all the gold produced in the historyof mankind – since 1886 (Phillips and Law, 2000).
Other conglomerate-hosted gold deposits such as
those of the Tarkwa (Ghana) and Jacobina (Brazil)
basins are relatively unimportant by comparison.
Irrespective of their production statistics, quartz
pebble conglomerate hosted gold deposits arecharacterized by considerable controversy regarding the origin of their gold mineralization. Arguments revolve essentially around whether gold
was introduced as detrital particles during sedimentation (a placer process; see Chapter 5) or
precipitated from hydrothermal solutions circulating through the sediments at some stage after
sediment deposition. There is little doubt, however, that in all three cases (i.e. the Witwatersrand,

195

country rocks such as the Comus Formation limestones.
Circulation of mineralizing fluids was focused essentiallyalong major extensional faults, which is the main reason
why major deposits are located along well defined structural trend lines. Individual deposits tend to occur beneatha less permeable, often siliciclastic, cap rock (such as theRoberts Mountain allochthon at Twin Creeks) causing thefluids ascending along the faults to flow laterally into
more reactive sequences such as the Comus Formation.
Carlin-type deposits are unique in the broader familyof orogenic gold deposits and are perhaps intermediatein character between more typical compression-related
orogenic gold ores and low-sulfidation epithermal gold
deposits (see Chapter 2). There is, however, no reason
why similar styles of epigenetic, sediment-hosted deposits,
containing “invisible” or micrometer-sized gold in sulfides,
should not be located in other parts of the world besides
Nevada.

Tarkwa, and Jacobina) the host sediments underwent burial and regional metamorphism andthat the associated circulation of metamorphic
fluids resulted in a clearly epigenetic componentof mineralization. Fluid inclusion studies indicate that certain of these fluids have a mixed
H2O–CO2 composition and have some similarities to those involved with orogenic gold deposits.
One view of the origin of gold mineralization in
the Witwatersrand sediments is that it occurs
along major unconformities, represented by conglomerate deposition, which represent zones ofintense fluid flow and alteration (Figure 3.28a).
Alteration is the product of fluid–rock reactions
involving progressive neutralization of an acidfluid (see section 3.5 above), with the most intense
alteration being characterized by the assemblage pyrophyllite–chloritoid–muscovite–chlorite(Barnicoat et al., 1997; Philips and Law, 2000). In
the Witwatersrand Basin, however, gold mineralization is characterized by a long-lived paragenetic sequence (Robb et al., 1997; Frimmel and
Minter, 2002) and is likely to be related to morethan just one-pass fluid–rock interaction. The
existence of a prominent alteration halo aroundzones of pervasive fluid flow is neverthelessone of the reasons why the Witwatersrand gold
deposits are so large. At the scale of a single
deposit similar effects can be seen, as exemplified

196

PART 2 HYDROTHERMAL PROCESSES(a)

Central Rand GroupS

Witwatersrand basin

Elsburg
sequence
Kimberley
sequenceN

Gold-bearing
conglomerate
packages

Elsburg

Bird sequenceMainsequence

West Rand
Group
Kimberley

Alteration assemblages (+Qtz)
Pyrophyllite, chloritoid, muscovite, chlorite
Muscovite, chlorite
500 m
Albite, epidote, calcite, muscovite, chlorite, paragonite
Stratigraphic boundaries
Principal zones of fluid flow

Bird

20 km

Main

(b)
W
E
Gold-bearing conglomerate

e
rat

me

glo

on

nc

rre

Ba
2

0
metresShear zones
Cr mica, tourmaline, rutile, muscovite alteration
Zone of sulfidation (pyrite)

Shear zones

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

by the João Belo reef in the Jacobina Basin of Brazil(Figure 3.28b). Here gold mineralization coincides
with a section of the conglomerate horizon thatis cut by a number of bedding-parallel shear
zones along which fluids have flowed resulting infuchsite (Cr–mica)–tourmaline–rutile–muscovite–pyrite alteration (Milési et al., 2002). Outside
this alteration zone the conglomerate is reported
to be barren.
The nature of the ore-forming fluid associatedwith the Witwatersrand, Jacobina, and Tarkwa
auriferous conglomerates has been described as
mixed H2O–CO2, fairly reducing, near-neutral,
and with low salinity (Philips and Law, 2000).
Barnicoat et al. (1997) preferred a more oxidized,
acidic fluid for the Witwatersrand Basin whereas
Klemd et al. (1993) suggested that the ore fluids
associated with the Tarkwaian conglomerates
were dominantly high density carbonic with significant contained methane and nitrogen. Detailed
fluid inclusion studies (Drennan et al., 1999;
Frimmel et al., 1999) have indicated that in the
Witwatersrand Basin different fluids are implicatedin at least two distinct episodes of mineralization,and these include mixed H2O–CO2 compositions
as well as aqueous, high-salinity fluids. The originof these fluids is also contentious and options
range from devolatilization of greenschist toamphibolite facies mafic rocks beneath the sedimentary basins, to derivation by prograde metamorphic reactions within the basin itself (Stevens
et al., 1997 and Figure 3.4). The hydrothermal
ore-forming processes of the WitwatersrandBasin in particular are long-lived and complex andthis is the reason why there has been, and stillis, so much debate on the genesis of these ores. It
is, for example, now reasonably well established

Figure 3.28 (Opposite) (a) Simplified north–south
section showing the pattern of alteration at a regional
scale associated with gold-bearing conglomerate
packages in the Central Rand Group of the
Witwatersrand Basin (after Barnicoat et al., 1997);(b) Cross section through the João Belo reef in the
Jacobina Basin showing the relationship between
shear zone related fluid conduits, alteration,and gold mineralization at a deposit scale(after Milési et al., 2002).

197

that the basin was struck by a large meteorite at
2025 Ma and that the catastrophic exhumationof the gold-bearing sedimentary host rocks, aswell as the likely circulation of related meteoric
fluids, must also have played an important role in
the ore formation process (Gibson and Reimold,1999).

3.10 ORE DEPOSITS ASSOCIATED WITH CONNATE FLUIDS
A number of sediment-hosted base metal deposits
are genetically linked to the circulation of connate fluids during diagenesis. In these deposit
types metal transport and deposition is generally
restricted to the sedimentary sequence throughwhich the connate fluids circulate. Ores includedin this category are the important stratiform,
sediment-hosted copper deposits (abbreviated to
SSC, but also called red-bed copper deposits)
and the family of Pb–Zn ores, usually associatedwith carbonate sediments (although some aresandstone-hosted), known as Mississippi Valley
type (or MVT) deposits. Although the two deposit
types are different in several respects, they bothowe their origins to circulating basinal brines.
Sverjensky (1989) has suggested that the majordifferences between SSC and MVT deposits relate
mainly to the type of sediment through which thefluids have traveled prior to ore deposition and also
to the contrasting properties of Cu, Pb, and Zn
in hydrothermal solutions. It is suggested that abasinal brine channeled into an oxidized (red-bed)
aquifer would scavenge all three base metals fromfertile detritus in the basin, but that the resulting
ore fluid would be saturated only with respect to
Cu and undersaturated relative to Pb and Zn. The
tendency in such an environment would be for
fluids to preferentially deposit copper sulfide minerals while the higher solubility metals remain in
solution. Accordingly, SSC deposits tend to be
characterized by a Cu > (Pb + Zn) metal association (Figure 3.29). By contrast, the same original
fluid channeled into a relatively reduced carbonate aquifer would evolve to Zn and Pb saturated
conditions in the presence of dissolved sulfur
species and ore deposition would yield sphalerite> galena >> chalcopyrite, as is typical of most198

PART 2 HYDROTHERMAL PROCESSES

Examples)ifer sy

ate aquCarbonBasinalbrine
sourceeducedstem (rQuartz sandstone aquifer system

Zn-rich ores
Zn > Pb >> Cu

Pb-rich ores
Pb > Zn >> Cu

Evaporite + oxidizedred bedT = 100–150°C
pH = 4–6
fO2 = magnetite/
hematite
Zn = 5 ppm
Pb = 1 ppm
Cu = 0.1 ppm

aquifersystem

Cu-rich ores
Cu > Pb + Zn

Upper Mississippi Valley
Pine Point

LaisvallSoutheast Missouri

Kupferscheifer
Central African Copperbelt
White Pine

Figure 3.29 Diagrammatic representation showing the relationship between a single basinal brine andthe conditions under which it might form SSC, as well as carbonate- and sandstone-hosted, MVT deposits(after Sverjensky, 1989; Metcalfe et al., 1994).

MVT deposits. In the situation where the originalbasinal brine passed through a quartzose sandstone aquifer, the fluid would not evolve significantly from its original metal carrying capacity,with saturated levels of Cu and Pb but undersaturated relative to Zn. Maintenance of the oxidation
state of the fluid at levels appropriate to hematite–magnetite coexistence would nevertheless limit
the mobility of copper and resulting ores would
be galena-rich, consistent with the fact thatsandstone-hosted MVT deposits tend to have Pb> Zn >> Cu metal ratios. It is interesting to notethat in the Central African Copperbelt, for example, Cu(–Co) mineralization is typically hosted
in clastic sediments such as arkose and shale,
whereas Zn–Pb mineralization is found associated with carbonate rocks.
Some of the processes relevant to the formationof SSC and MVT deposits are discussed in moredetail below.
3.10.1 Stratiform sediment-hosted copper(SSC) deposits
SSC deposits worldwide rank second only to porphyry copper deposits in terms of copper production and they represent the most important global

source of cobalt, as well as containing resourcesof many other metals such as Pb, Zn, Ag, U, Au,
PGE, and Re. There are several deposits of varying
ages around the world, but these are overshadowedby the huge resources of only two regions, the Neoproterozoic Central African Copperbelt (Box 3.6)
and the Permo-Triassic Kupferschiefer (or “copper
shale”) of central Europe. Other important deposits include White Pine in Michigan, Udokan
and Dzhezhkazgan in Kazakhstan, Corocoro in
Bolivia, and the Dongchuan district of China(Misra, 2000). The formation of major SSC deposits seems to coincide with periods of supercontinent amalgamation such as Rodinia in the
Neoproterozoic and Pangea in the Permo-Triassic(see Chapter 6).
SSC deposits are typically hosted within an
intracontinental, rift-related sedimentary sequence.
The early part of the sequence was either deposited originally as an oxidized (red-bed), aeolian
to evaporitic assemblage, or was rapidly oxidized
during burial and diagenesis. This sequence isoverlain by a shallow marine transgression that
deposited a more reduced assemblage of shales,carbonates, and evaporites. Basin-derived fluid
circulation was promoted by the high heat flow
conditions accompanying rapid rifting and sub-

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

sidence, and the very permeable environment
created by porous clastic sediments and active
growth faults. The connate fluids circulating at
this stage of basin evolution were saline, relatively oxidized, and pH neutral. Metals, in particular copper, are believed to have been leachedfrom detrital minerals such as magnetite, biotite,
hornblende, and pyroxene which themselves may
have been derived from erosion of a fertile basement assemblage (a Paleoproterozoic magmatic
arc in the case, for example, of the Central African
Copperbelt; Rainault et al., 1999; and Box 3.6).
The aqueous transport of copper under these
conditions was as a cuprous–chloride complex,
probably CuCl 32− (Rose, 1976). Metal deposition
occurred at a redox interface where the oxidized,
metal-charged connate fluids intersected overlying or laterally equivalent reduced sediments
or fluids. The ore deposits are generally zoned,
usually at a district scale, and typical zonation ischaracterized by the sequence barren/hematite –native copper – chalcocite – bornite – chalcopyrite– Pb/Zn/Co sulfides – pyrite.
SSC deposits can be subdivided into two subtypes (Kirkham, 1989). The first, less important,subtype is hosted in continental red-beds thatwere probably originally oxidized, with ores precipitated around locally developed zones of reduction (for example, Dzhezhkazgan in Kazakhstan
and Corocoro in Bolivia). The second subtype,
represented by the much larger Kupferschiefer
and Copperbelt deposits, is characterized bymetal deposition in more reduced shallow marine
sequences that were partially oxidized subsequent to deposition, and where oxidized (red-bed)
sediments occur stratigraphically beneath the orezone. In both cases, however, ore fluids are considered to have interacted with the dominantly
clastic, oxidized sedimentary sequence and it is
this feature that determines the nature and properties of the hydrothermal solutions. Such fluids
are likely to have been characterized by low temperatures (<150 °C and possibly even <100 °C), a
neutral range of pH (5–9) constrained by carbonate or silicate equilibria (K-feldspar/illite), oxidized
conditions (hematite and other ferric-oxyhydroxides stable), and moderate salinities (typically up
to 20 wt% NaCl equivalent), reflecting interac-

199

tion of the fluids with evaporitic rocks. Rose
(1976) showed that under these conditions the
CuCl 32− complex would have been particularly
stable (Figure 3.30), although other aqueous species, such as CuSO4 or Cu(OH) 2+
2 could also haveexisted. Solubility of copper as a chloride complex under such conditions has been estimatedto be as high as 35–100 ppm (Haynes, 1986a). In
Figure 3.30 it is evident that the stability field of
the CuCl 32− complex largely coincides with that ofhematite stability, and that the aqueous sulfur−
species would be SO 2−
4 rather than H2S or HS .
Rose (1976) points out that the conditions described above are particularly well suited to copper
mobilization. By comparison, copper is relatively
insoluble in normal meteoric waters and it is also
unlikely to be easily leached from more reduced
sediments where copper sulfides are stable. An
extensive red-bed sedimentary environment where
fertile detrital minerals readily break down to forma source of copper metal, and where stable CuCl 32−
complexes result in high copper solubilities, is,therefore, ideal for the formation of Cu > (Pb + Zn)
SSC deposits.
The formation of SSC deposits has, in the past,
been a controversial topic and models supporting
syngenetic as opposed to diagenetic origins havebeen much debated (Garlick, 1982; Fleischer,
1984; Jowett et al., 1987; Annels, 1989; Sweeney
et al., 1991). The syngenetic view, in which sulfides were believed to have precipitated directly
out of anoxic sea water in processes akin perhaps
to those giving rise to Mn deposition in thepresent day Black Sea (see Chapter 5), have now
largely been superseded by models that view ore
genesis in terms of either early or late diagenesis.
A diagenetic origin is substantiated by many ofthe prominent features of SSC deposits, especiallythe broadly transgressive nature of ore zones relative to lithologies, the clear relationship of metal
deposition to a redox front, and replacement textures in the sulfide mineral paragenetic sequence.
There are, nevertheless, several processes that canbe called upon to explain ore deposition and metal
zonation in SSC deposits.
It is evident in the Kupferschiefer deposits,
for example, that metal deposition is spatiallyassociated with a reddish oxidized zone (termed

200PART 2 HYDROTHERMAL PROCESSES

0.8

O

2

H

Cu4(OH)6Cl2

2O

Cu2+

CuO

Eh (volts)

0.4

CuCl32–
He

Cu

tite

ma

2S

0.0

SO 2
4

Cu2O

–

H

2S

H

Cu

2O

H

2

HS –
–0.4

3

5

7
pH

9

the “Rote Fäule”) that transgresses different
lithologies and comprises hematite-replacing
early diagenetic sulfides. The oxidized zone represents a chemical front in the footwall of the orezone and against which a zoned sequence of basemetal sulfides (chalcocite–bornite–chalcopyrite–
galena–sphalerite–pyrite) is arranged. Much of
the metal concentration is nevertheless concentrated within the organic-rich shale (i.e. theKupferschiefer itself) since it is this unit that
contained concentrations of early diagenetic
framboidal iron sulfides (such as marcasite or
pyrite) formed by bacterial reduction of sulfate tosulfide. Sawlowicz (1992) has documented evid-

11

Figure 3.30 Eh–pH plot for thesystem Cu–O–H–S–Cl at 25 °C(with ∑S = 10 −4 m and Cl− = 0.5 m).
Also shown are the stability fields
for hematite, various sulfur species,
and the cuprous-chloride complex
for conditions compatible withthe formation of SSC deposits(after Rose, 1976).

ence of copper sulfide replacement of framboidsin the Kupferschiefer, and the applicability ofearly chalcocite replacement of biogenic sulfide
material to SSC deposits in general has been discussed by Haynes (1986b). These features suggest
low temperature processes and point to metal
deposition during the early stages of diagenesis of
the sedimentary succession. A schematic representation of the zonal distribution of metals
against the Rote Fäule in the Kupferscheifer ores
is shown in Figure 3.31a, together with a model by
Metcalfe et al. (1994) that demonstrates the effectof changing the redox state of the ore fluid onsequential metal precipitation (Figure 3.31b). In

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

201

(a)

Fe2+

Fe3+

Cu

Zn

Pb

Rote Fäule
Ore

(b)
100

Figure 3.31 (a) Section showing
metal zonation in typicalKupferschiefer ores in relation to
the transgressive, oxidized “RoteFäule” zone (after Jowett et al.,1987). (b) Model showing theeffects of mixing a typical SSC-typeore fluid with a reduced fluid on
metal precipitation. The sequenceof deposition is the same as thezonal pattern observed in theKupferschiefer (after Metcalfe
et al., 1994).

% total metal in solution

80

Zn

60

Pb

40
Cu
20

0

0

the latter metal precipitation is achieved by mixing the oxidized ore-bearing fluid with a reduced
fluid, rather than reacting the fluid with a reduced
sediment. The effect is nevertheless the same andmetals are precipitated in the sequence Cu–Pb–
Zn as a function of increased mixing of the two
fluids. This is exactly the sequence observed in
the Kupferschiefer ores and confirms that thistype of zoning could be a result of sequential
metal precipitation from mixed connate fluids.
There are other ways of explaining the patternsof metal zoning and paragenetic sequence in SSCtype deposits. Alternative proposals prefer a late

10

20

30

40

50

% reduced water in mixture

diagenetic origin for metal precipitation in SSC
deposits where sulfate reduction to sulfide occurs
at higher temperatures and is not bacteriallyinduced. This could be achieved by fluid mixing(as suggested in Figure 3.31b) or simply by interaction between oxidized fluid and reduced countryrock. Ripley et al. (1985) have proposed a scheme
in which the pattern of zoning in SSC type ores is
explained in terms of either fluid mixing or fluidrock reaction. In a plot of Fe2+ versus Cu+ activities (Figure 3.32) it is evident that a fluid with ahigh initial Cu content will become progressively
more dilute as copper sulfides precipitate. This

202

PART 2 HYDROTHERMAL PROCESSES

T = 100°C
Log f O2 = –54

Chalcocite
Log [α(Cu+)/α (H+)]

–6

e

nit

r
Bo

Chalcopyrite–8
Hematitesaturation
Pyrite

–10

0

2

4
6
Log [α(Fe2+)/α (H+)]

8

Figure 3.32 Plot of Cu+/H+ versus Fe+/H+ activity
ratios to explain the zoning and paragenetic sequencethat is often observed in SSC deposits when a Cubearing fluid reacts with a pyritic shale. In thissituation the original copper content of the fluid ishigh and then decreases as a function of Cu-sulfide
precipitation to form the paragenetic sequencechalcocite–bornite–chalcopyrite–pyrite (after Ripley
et al., 1985).

would produce a paragenetic sequence and zonation pattern that reflects a decreasing Cu content
in the sulfide ores, such as chalcocite – bornite –chalcopyrite – pyrite.
Another feature of SSC deposits is that they
contain appreciable amounts of other metals,
such as Ag, Pb, and Zn (Kupferschiefer) and Co(Copperbelt), all of which form part of a broadly
zonal arrangement of ores. Such patterns may bedifficult to explain in terms of solubility contrasts
when metal concentrations are well below theirsaturation levels, and in such cases a mechanism
such as adsorption may be relevant (see section
3.5 above). Figure 3.15b, for example, shows thevariations in adsorption efficiency as a function of
pH. It is feasible that variable adsorption efficiencies could control the uptake of different metals

from red-beds into connate fluids as well as theirsubsequent deposition into nearby reducing sediments. This could explain regional zonation patterns in SSC ores, although the details of how thismight relate ultimately to paragenetic sequences
are likely to be complex (Rose and BianchiMosquera, 1993).
3.10.2 Mississippi Valley type (MVT)
Pb–Zn deposits
MVT deposits, like SEDEX and SSC ores, owe
their origin to fluid circulation and metal transport/
deposition within sedimentary basins. Unlike thesyngenetic to diagenetic time frame of SEDEXand SSC deposit formation, however, MVT ores
are distinctly epigenetic and metals can be deposited tens of millions of years after sedimentdeposition. The deposits form from relatively low
temperature fluids (<150 °C) and are broadly
stratabound, mainly carbonate-hosted, and dominated by sphalerite and galena with associatedfluorite and barite. Most MVT deposits containsignificantly more sphalerite than galena, with
the notable exception of the Viburnum Trend insoutheast Missouri, where the reverse applies.
The sandstone-hosted MVT deposits (such as
Laisvall and some deposits in southeast Missouri)
represent a conceptually intermediate category
between SSC and carbonate-hosted MVT ores
(Figure 3.29) and are characterized by Pb > Zn >>Cu. MVT deposits contain a significant proportion of the world’s Zn and Pb reserves and there
are many other deposits around the world in addition to the numerous mines in the MVT “typearea” of the southeastern USA. In the latter region
the main metal-producing districts include theViburnum Trend of southeast Missouri (which ismainly a Pb-producing region), the Tri-State (i.e.
Oklahoma, Kansas, Missouri) region, the Illinois–
Wisconsin region and the Mascot–Jefferson City
area of east Tennessee. Other well known MVT
districts or deposits elsewhere in the world
include the Pine Point and Polaris deposits ofnorthern Canada, the Silesian district of Poland,
Mechernich in Austria, the Lennard Shelf district,
Sorby Hills and Coxco in Australia, and Pering inSouth Africa (Misra, 2000).

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

203

Circulation of sediment-hosted connate fluids:
1 The Central African Copperbelt
The stratiform, sediment-hosted Cu–Co deposits hosted
within the Neoproterozoic Katangan sequence of Zambia
and the Democratic Republic of the Congo (DRC) form
the Central African Copperbelt, undoubtedly one of thegreat metallogenic provinces of the world. Several worldclass deposits (such as Nchanga, Konkola, and Tenke
Fungurume), together with dozens of smaller mines,

combine to make this region one of the foremost copperproducing areas of the world, as well as its most important
Co supplier. Although figures for the entire Central African
Copperbelt are unreliable, it has been estimated that theregion still contains resources in excess of 150 million
tons of Cu metal and 8 million tons of Co metal (Misra,
2000).
Group

Less than 570 Ma

Upper
Kundelungu

Glacial diamictite
Lower
Kundelungu
Glacial diamictite(circa 750 Ma)

Katangansequence
MwashyaLess than
880 MaStratiform
Cu–Co
mineralization
Roan

Nchanga
granite(880 Ma)
Muva
metasediments
GneissPaleoproterozoic
basement(2050–1870 Ma)

Lufubumetavolcanics

Figure 1 Simplified stratigraphic profile through the rocks of the Central African Copperbelt showing therelationship between the Paleoproterozoic granite–gneiss and volcano-sedimentary basement and theoverlying Katangan sequence. Stratiform Cu–Co mineralization is largely hosted in a variety of sedimentary
lithotypes in the lower part of the Roan Group, but does also occur higher up in the succession and up into the
Mwashya Group. The age constraints used in this compilation are after Rainaud et al. (2002a), Master et al.
(2002), and Key et al. (2001).

204

PART 2 HYDROTHERMAL PROCESSESThe Katangan sediments were deposited in an intracratonic rift formed on a Paleoproterozoic basement
comprising 2050–1870 Myr old granite–gneiss, andmeta-volcanosedimentary sequences (Figure 1; Unrug,1988). The basement is thought to represent a deformed but
fertile magmatic arc that contained the ultimate sourceof Cu (in porphyry type intrusions; Rainaud et al., 1999),
and possibly Co too. Initiation of rifting and sedimentdeposition followed soon after intrusion of the Nchanga
granite at 880 Ma. Initially sedimentation was dominatedby siliciclastic deposition, but this was followed by playa
lake and restricted marine incursions to form evaporites,
shales, and carbonate rocks. A major glacial diamictite (the“grand conglomerat”) occurs at the top of the MwashyaGroup, the age of which (circa 750 Ma) makes it a likely
correlative of the mid-Neoproterozoic “Snowball Earth”
glaciation (Robb et al., 2002). The uppermost parts of the
Katangan sequence (i.e. the Upper Kundelungu Group andits correlatives) now appear to be substantially younger(i.e. less than 565 Ma) than the underlying portions andare regarded as a separate basin that formed as a foreland
to the Pan-African orogeny (Wendorff, 2001).
Mineralization in the Central African Copperbelt is amuch debated and contentious issue. Ideas on ore genesis
in the 1960s and 1970s were dominated by syngenetic
models, where metals were accumulated essentially at thesame time as the sediments were being deposited (Fleisher
et al., 1976). More recent work has shown that mineralization is more likely to be related to later fluids circulating
through the sediments, either during diagenesis and compaction of the sequence (Sweeney et al., 1991) or significantly later during deformation and metamorphism (Unrug,1988). Although much work is still required, it seems likelythat metals were sequentially precipitated by processes
related to reduction–oxidation reactions between sedimentary host rocks and a fertile connate fluid. A highly
oxidized, saline, and sulfate-rich connate fluid, derived byinteraction with evaporites in the Roan Group, is regarded
as the agent that scavenged Cu and Co from the sediments themselves. The Cu may have been introduced into
the basin in Cu-rich magnetite or mafic mineral detritus
eroded from the fertile Paleoproterozoic magmatic arc
in the hinterland (S. Master, personal communication).

The majority of MVT deposits worldwide
formed in the Phanerozoic Eon, but more
specifically, in Devonian to Permian times; some
deposits, such as Pine Point and the Silesian
district, also formed during the Cretaceous–
Paleogene period (Leach et al., 2001). The former,
far more important, metallogenic epoch coincides

Figure 2 Folded metasediments hosting stratiform
Cu–Co sulfide ore at the Chambishi Mine, Zambia.
The bedding parallel ore is cut by an axial planarcleavage (photograph by Lynnette Greyling).
Co, on the other hand, may have been derived from
penecontemporaneous sills that intrude the upper RoanGroup (Annels, 1984). Cu and Co would be highly
soluble as chloride complexes in a connate fluid of thistype (see section 3.10.1 of this chapter) and it is likely thatvery efficient leaching of these metals during diagenesis
and oxidative alteration of the host rocks would haveoccurred. Precipitation of these metals would occur as thefluid reacted, either with more reduced strata (containing
organic carbon or diagenetic framboidal pyrite), or witha second, more reduced, fluid. Metal zonation, which isa feature of mineralization throughout the Copperbelt,
could, at least in part, be a product of the variable reduction potentials of metals such as Cu, Co, Pb, and Zn. An
early or late diagenetic model for the origin and timing ofstratiform mineralization in the Central African Copperbelt
is consistent with the fact that it pre-dates the major
period of compressive deformation that affected the entire
region during the Lufilian orogeny (590–510 Ma; Rainaud
et al., 2002b). Figure 2 shows the folded nature of thehost rock strata at the Chambishi Mine in Zambia; thestratiform Cu–sulfide ore here is folded and cut by an axial
planar cleavage.

with the assembly of Pangea when major compressional orogenies were active over many partsof the Earth (see Chapter 6). The later period is more
specifically related to the Alpine and Laramide
orogenies and, likewise, associated with compressional tectonic regimes. The timing of MVT
deposits in relation to associated orogenic activity

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

Topographically elevatedorogen205

MVT deposit

Foreland basin

Tectonic
compresson
0

100
km

Probable trace ofpre-mineralization fault

Figure 3.33 Diagram illustrating
the concept of hydrologicalcontinuity between acompressional orogenic belt anda foreland sedimentary basinthrough which orogenically andtopographically driven fluids
flow, and within which MVT
Zn–Pb deposits form. Inset shows
characteristics of bedded anddiscordant mineralization at adeposit scale. Mineralization may
be bedded and associated withdolostone/limestone sequences, orin discordant carbonate breccia
zones (after Garven et al., 1993).

provides an important clue as to the causes of
fluid flow and hydrothermal activity involved intheir formation. It is generally accepted that topographically driven fluid pathways were critical tothe development of large MVT ore districts, andthat the carbonate host rocks maintained a hydrological connection to orogenic belts active during
the period of ore deposition (Rickard et al., 1979;
Oliver, 1986; Duane and De Wit, 1988). Other
features believed to conceptually link most MVT
deposits include a low-latitudinal setting, wherehigh rainfall ensured an adequate fluid reservoir,
and the presence, somewhere in the fluid flow
system, of a high-evaporation sabkha environment that resulted in the high fluid salinities
necessary for viable levels of metal solubility andtransport (Leach et al., 2001).
Figure 3.33 illustrates the main characteristicsf

Dolostoneorlimestonef
Mineralized
breccia zones

Bedded-type
orebody0

Discordant-type
orebody40

metersof MVT deposits. Regional fluid flow is stimulated
by compressional orogeny that results in thrustfaulting and uplift and this, in turn, creates a
topographic head and fluid flow down a hydrological gradient. Fluid flow in these tectonic settings
occurs over distances of hundreds of kilometersand such fluids are implicated in the migrationof hydrocarbons as well as metals (Rickard, 1976;
Oliver, 1986). At the site of metal deposition MVT
ores are sometimes bedded and focused along
conformable dolostone/limestone interfaces, but
are more commonly associated with discordant,
dissolution-related zones of brecciation. The hydrothermal fluids that are linked to MVT deposit
formation are typically low-temperature (100–150 °C), high-salinity (>15 wt% NaCl equivalent)brines, with appreciable SO 2−
4 , CO2 and CH4
and associated organic compounds and oil-like206

PART 2 HYDROTHERMAL PROCESSES

H2S + Zn2+ ⇔ ZnS + 2H+

droplets (Gize and Hoering, 1980; Roedder, 1984).
These compositional characteristics are not unlike
oil-field brines and reflect protracted fluid residence times in the sedimentary basin.
In many MVT deposits ores occur as cement
between carbonate breccia fragments suggestingthat brecciation occurred either before (perhaps asa karst related feature) or during metal deposition(Anderson, 1983). Hydrothermal dissolution ofcarbonate requires acidic fluids and could occur
by a reaction such as:[3.12]The production of additional Ca2+, Mg2+, and CO2
in the fluid as a result of carbonate dissolution
provides the ingredients for later precipitation of
calcite and dolomite, a feature that is observed insome MVT deposits where secondary carbonate
gangue minerals cement previously precipitatedore sulfides (Misra, 2000).
The rather unusual nature of the ore fluid
associated with MVT ore formation raises someinteresting questions concerning metal transportand deposition (Sverjensky, 1986). Zn and Pb are
borderline Lewis acids (see section 3.4.1 above)and can complex with both chloride and bisulfide
ligands. At high temperatures and under acidic
conditions the metal–chloride complex tends tobe more stable, whereas at lower temperatures

CaMg(CO3)2 + 4H+ ⇔ Ca2+ + Mg2+ + 2H2O + 2CO2
[3.11]
The fluid itself might originally have been acidic,
or hydrogen ions were produced by precipitationof metal sulfides according to the reaction:–48

T = 100°C
NaCl = 3 m
ΣS = 10–2 m

Calcite stable

Log f O2

–52

–54

ΣPb = 10–5 m
HematitePb relatively insoluble–50

Pb relatively

solu

bl

e

Py

rite

Magnetite
SO42–

Redu

ced S

H2S
–

HS

= 10 –5m–56
Sufficient reduced
S to form sulfide
ore
–58

3

4

5

6

7

8

9

10

11

pH
Figure 3.34 fO2–pH plot showing key mineral and aqueous species stabilities under the conditions shown.
Solubility contours of ∑Pb = 10 −5 m and reduced sulfur = 10 −5 m are also shown. Higher solubilities lie to the left ofand above the Pb contour, and below the sulfur contour (after Anderson, 1975).

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

and neutral to alkaline pH the metal–bisulfide
complexes dominate. There seems little doubt,
given the prevailing conditions, that metal–
chloride complexes predominate during MVT ore
formation (Sverjensky, 1986). Anderson (1975)
pointed out that viable Zn and Pb solubilities are,however, only achieved when fluids are either
relatively oxidized (i.e. the sulfur in the fluid−
is transported as SO 2−
4 and not as H2S or HS ), orat low pH. Figure 3.34 shows that if a fluid was
sufficiently oxidized for SO 42− to have been stablethen high Pb solubilities could have occurred overa range of pH. That the sites of MVT ore deposition are characterized by sulfide minerals, however, implies that metal precipitation must havebeen accompanied by the reduction of sulfate
to sulfide. In some MVT deposits, such as thePine Point district, such a process is feasible andore deposition has been attributed to the mixingof a sulfur-deficient ore fluid with another carrying significant reduced sulfur (i.e. H2S or HS−) atthe site of ore deposition (Beales and Jackson,1966). Alternatively, the ore fluid might itself
originally contain significant SO 2−
4 , with reduction and metal precipitation occurring, again either

207

by mixing with a reduced fluid, or simply byreaction of the fluid with organic matter. In
many MVT deposits, however, textures and geochemistry are incompatible with a fluid mixing
model and ores appear to have precipitated over a
protracted period of time. If fluid mixing and/or
sulfate reduction did not occur then the ore fluid
itself would have needed sufficient reduced sulfur
to form sulfide minerals (i.e. in excess of 10−5 m;
see Figure 3.34). In a reduced fluid metal–chloride
complexes will only be stable under acidic conditions (low pH) and it has been suggested that
this could not apply to MVT fluids that haveequilibrated with carbonate host rocks. It appears,
however, that the CO2 contents of fluids associated with some MVT deposits are sufficiently
high to form low pH solutions (Sverjensky, 1981;
Jones and Kesler, 1992). Such acidic fluids will
promote the stability of metal–chloride complexes resulting in high metal solubilities. Suchfluids also promote ground preparation by creating dissolution breccias at the site of ore deposition. This in turn results in metal precipitation as
the ore fluid is neutralized (equation [3.11] above)by reaction with the carbonate host rock.

Circulation of sediment-hosted connate fluids:
2 The Viburnum Trend, Missouri
The Viburnum Trend, also known as the “New Lead Belt,”
in southeast Missouri is the world’s largest lead-producing
metallogenic province. In the mid-1980s production plus
total reserves from this remarkable, 65 km long, belt ofmineralization amounted to some 540 million tons of ore
at an average grade of 6% Pb and 1% Zn (Misra, 2000).
The Viburnum Trend occurs to the west of the St Francois
Mountains, a Precambrian basement outlier that is unconformably overlain by a Cambro–Ordovician sequenceof siliciclastic and carbonate sediments. The host rocks
to Pb–Zn ore are mainly Cambrian dolostones of theBonneterre Formation, although mineralization took place
more than 200 million years later during Carboniferous
times. The Pb–Zn ores of the Viburnum Trend are regarded
as a classic example of Mississippi Valley type (MVT)
deposits.

The galena–sphalerite dominated mineralization of theViburnum Trend is hosted within a sequence of carbonate
rocks that formed during a shallow marine incursion overa continental shelf (Larsen, 1977). A stromatolitic reefdeveloped around the northern and western margins of apositive feature (the St Francois Mountains) on this shelf,
resulting in the formation of biostomal dolomites thatare underlain by the porous Lamotte sandstone and overlain by the relatively impervious Davis shale (Figure 2).
Accurate age dating of the actual mineralization in the
Viburnum Trend, and indeed for much of the related
Pb–Zn MVT styles of mineralization throughout the southeastern USA, indicates that it formed during a relatively
brief interval of time in the late Carboniferous (i.e. between
about 330 and 300 Ma; Leach et al., 2001). This interval
coincides with a period of Earth history characterized

208

PART 2 HYDROTHERMAL PROCESSES

Missouri

Indian Creek Mine

PotosiViburnum
Trend
Old
Lead
Belt
No. 27
Viburnum Mines
No. 28

No. 29Magmont Mine
Buick Mine
Brushy Creek MineSt Francois
Mountains

Fletcher Mine

Ozark Lead Mine

LegendOutcrops of Precambrianbasement
Cambrian and Ordovician
sedimentary rocks
Intensivemineralization
Weakmineralization
Major faults
0

10
km

Figure 1 Geological setting of the Viburnum Trend Pb–Zn mineralization and the major mines in this belt
(after Kisvarsanyi, 1977).

209

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3

Fletcher mine
No. 30 shaft

W
Davis shaleE

Pb-Zn sulfide ore

Clastic carbonate

Stromatolitic

reefBonneterre
Formation

Precambrian basement
Lamotte sandstone

0300
metersFigure 2 West–east section through the Fletcher Mine, Viburnum Trend, showing the close association
between mineralization, basement topographic highs, and stratal pinch-outs (after Paarlberg and Evans, 1977).
by major compressional orogeny, during which time theamalgamation of the supercontinent Pangea was taking
place (see Chapter 6). In the southeastern USA this eventis recorded by the Ouachita collisional orogen and it issuggested that it was this event that stimulated fluid flow
throughout the region. Orogeny- or topography-driven
fluid flow (see sections 3.3.1 and 3.10.2 of this chapter)focused hydrothermal solutions into connected paleoaquifers such as the Lamotte sandstone, with subsequent
metal precipitation occurring in response to features suchas fluid interaction and dissolution of carbonate rocks. Aninteresting feature of many of the Viburnum Trend deposits

3.11 ORE DEPOSITS ASSOCIATED WITH NEAR SURFACEMETEORIC FLUIDS (GROUNDWATER)
Few hydrothermal ore deposits are linked to thecirculation of ambient meteoric fluids, or groundwaters, in the near surface environment sincemetals are generally not transported in thesetypes of fluids. A notable exception is the familyof sediment-hosted uranium deposit types associated with low temperature meteoric fluid flow
(Nash et al., 1981). One category of these deposit
types is the calcrete-hosted, uranium–vanadiumores that form in arid regions. These deposits are
related to high rates of evaporation in surficial
environments and, consequently, are discussed
in Chapter 4. The other category, the sandstonehosted uranium ores, is related to groundwaterflow and redox precipitation mechanisms, and isdiscussed below.

is the close association between basement topography,
overlying sedimentary strata pinch-outs, and mineralization. Figure 2 shows a west–east section through the
Fletcher Mine where mineralization is constrained to azone in the Bonneterre dolostones that is directly above abasement topographic high. Stratal pinch-outs against thisbuttress suggest that fluids, perhaps originally flowing
through underlying successions such as the Lamottesandstone (where some mineralization does occur) or the
lower stromatolitic reefs, might have been forced upwards
into the clastic carbonates, and that this was an importantfactor influencing metal precipitation.

3.11.1 A brief note on the aqueous transport anddeposition of uranium
In nature uranium occurs in two valence states,as the uranous (U4+) ion and the uranyl (U6+) ion.
In the magmatic environment uranium occurs
essentially as U4+ and in this form is a highly
incompatible trace element that occurs in only a
few accessory minerals (zircon, monazite, apatite,sphene, etc.) and is concentrated into residual
melts. The uranous ion is, however, readily oxidized to U6+ in meteoric waters. In most naturalaqueous solutions uranous complexes are insoluble, but U6+ forms stable complexes with fluoride(under acidic conditions at pH < 4), phosphate(under near neutral conditions at 5 < pH < 7.5) andcarbonate (under alkaline conditions at pH > 8)
ligands. The oxidized form of uranium is, therefore, easily transported over a wide range of210PART 2 HYDROTHERMAL PROCESSES

1.0

T = 25°C
PCO2 = 10–2 atm

0.8

0.6

Eh (volts)

0.4

2

UO2(CO3)22–

UO2

UO2CO3°

O
2+

H

2O

3.11.2 Sandstone-hosted uranium deposits
UO2(CO3)34–

4+

U
0.2

0

H

Uraninite
UO2

2O

H

2

–0.2
–0.4

2

4

6
pH

controls, uranium precipitation may also be promoted by fluid mixing, changes in fluid pH, andadsorption. A more complete overview of uranium ore-forming processes is presented in Nashet al. (1981).

8

10

12

Figure 3.35 Eh–pH diagram showing relevant aqueous
uranium species for the conditions specified. For most
meteoric waters in the near neutral pH range, thedominant aqueous species are likely to be U6+-oxide or-carbonate complexes. These will be precipitated fromsolution by a reduction of Eh to form U4+-oxide, oruraninite (after Langmuir, 1978).

pH conditions, whereas the reduced form of
the metal is generally insoluble. Figure 3.35, an
Eh–pH diagram for the system U–O2–CO2–H2O at
25 °C, shows that for most meteoric waters in
the near neutral pH range, the dominant aqueous
species are likely to be U6+–oxide or –carbonate
complexes. Uranium can be precipitated by lowering Eh of the fluid to form a uranous oxide, inthe form of either uraninite or pitchblende.
Consequently, many low temperature uranium
deposit types are a product of oxidation–reduction
processes where the metal is transported as U6+
and precipitated by reduction. In the presence ofhigh concentrations of other ligands, uranium canbe transported as a variety of different complexes,
and at higher temperatures (>100 °C) and underacidic conditions a molecule such as UO2Cl+ may
be stable (Kojima et al., 1994). In addition to redox

Most of the world’s productive sandstone-hosted
uranium deposits occur in the USA. These deposits occur mainly in three regions, the Colorado
Plateau region (around the mutual corners ofUtah, Colorado, New Mexico, and Arizona states),south Texas, and the Wyoming–South Dakota
region. Smaller, generally non-viable, deposits are
known from elsewhere, such as the Permian
Beaufort Group sandstones of the Karoo sequence
in South Africa and the Lodève deposits in France.
The deposits are generally associated with Paleozoic to Mesozoic fluvio-lacustrine sandstones andarkoses and are stratabound, although different
geometrical variants exist. Most of the deposits in
the Colorado Plateau region are tabular and associated with either organic material or vanadium(Northrop and Goldhaber, 1990). Deposits in
Wyoming and south Texas, by contrast, are sinuous in plan and have a crescent shape in crosssection. Although still stratabound these ores are
discordant to bedding and are typically formed at
the interface between oxidized and reduced portions of the same sandstone aquifer (Reynolds andGoldhaber, 1983; Goldhaber et al., 1983). It is thelatter type that is commonly referred to as rollfront uranium deposits. The genesis of the two
subtypes is quite different and they are discussed
separately.
Colorado Plateau (tabular)
uranium–vanadium type
Important features of tabular sandstone-hosted
deposits are that they can occur stacked one
on top of the other, and that the ore zones arebounded above and below by sandstones containing a dolomitic cement. The generally acceptedmodel for the formation of these deposit types isbased on work done in the Henry Basin of Utah,
where stacked U–V ore bodies are considered
typical of the Colorado Plateau region as a wholeHYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3(Northrop and Goldhaber, 1990). The ore bodies ofthis area occur predominantly within the Jurassic,
sandstone-dominated Salt Wash Member of theMorrison Formation, although on a regional scale
they are slightly discordant with respect to bedding (Figure 3.36a). The ores are believed to haveformed at the interface between two discrete, lowtemperature meteoric fluids. One of these fluids,
a relatively stagnant but saline basinal brine containing Na+, Mg2+, and Ca2+ cations and Cl− andSO 2−
4 anions, underlies a low salinity, meteoric
fluid that flows readily along aquifer horizons
bringing with it highly soluble metal species suchas UO2(CO3) 22− and VO+. Mixing of these two
fluids results in precipitation of metals in subhorizontal tabular zones reflecting the fluid interface(Figure 3.36b). The main uranium ore mineralin these deposits is a uranous silicate, coffinite(U(SiO4)1−x(OH)4x), which is associated eitherwith syn-sedimentary organic (plant) debris, ortogether with a vanadium-rich chlorite in theintergranular sandstone pore spaces.
Fluid mixing and metal precipitation occurred
within a few hundred meters of the surface suchthat fluid temperatures were unlikely to haveexceeded 30–40 °C. The stagnant basinal brine isconsidered to have been locally derived, but interacted with evaporitic sequences, which explains
its high Mg/Ca ratios. The overlying meteoric
fluid was essentially groundwater that scavenged
U6+ and V4+, as well as other oxide-soluble metals
such as Cu, Co, As, Se, and Mo, from overlying,uranium-enriched tuffaceous rocks, or frombreakdown of detrital Fe–Ti oxide minerals in
the aquifer sediments through which it flowed.
Mixing of the two fluids and associated diagenetic
reactions ultimately gave rise to ore formation,
the details of which are presented in Northropand Goldhaber (1990). One of the products of
fluid mixing in this environment was a dolomite cement that precipitated from pH-related(Ca,Mg)CO3 oversaturation within the basinalbrine. The pH changes in the mixing zone appear
to have been a product of the formation of
vanadium-rich clays (smectite–chlorite) during
compaction and sediment diagenesis. Vanadium
is incorporated into the dioctahedral interlayerhydroxide sites (as V(OH)3) of the clay minerals,
which reduces the [OH−] activity of the fluid and211

decreases its pH. The same chemical controls arealso implicated in the formation of coffinite ore,although the actual processes are more complex
and involved two stages. The uranyl–dicarbonate
complex most likely relevant to this meteoric
fluid is stable over only a limited pH range(around 7–8 in the conditions applicable to Figure3.35). A decrease in fluid pH caused by fluid mixing and V-clay formation causes a lowering of
UO2(CO3) 22− solubility which promotes the adsorption of UO +2 molecules onto the surface ofquartz grains. Once the fluid is reduced (probably
by bacterially induced SO4 reduction and production of H2S) to stabilize U4+ complexes, bonding
between the uranous ion and silica leads to theformation of coffinite (Northrop and Goldhaber,1990). It should be noted that organic matter
played an important role during the ore-forming
process, as is evident from the very close association between coffinite and plant debris. In
addition to the fact that organic matter is itself areductant, it supplies a source of nutrient for sulfate-reducing bacteria that in turn supply biogenic
H2S for reduction of U6+ and V4+ to their insoluble
lower valency forms. In the Grants mineral belt of
the Colorado Plateau region uranium mineralization is associated with organic matter in the formof humate that was introduced epigenetically into
the sediments by organic acid rich fluids.
The model illustrated in Figure 3.36b is particularly applicable to the tabular, stacked ores of
the Colorado Plateau region in which there is auranium–vanadium association. It suggests thatthe stacked nature of the tabular ore bodies could
be the product of a two-fluid interface that episodically migrates upwards with time. The dominant
control on ore formation is pH, although adsorption, bacterial mediation, and redox reactions
are also important processes. Evidence for the
migration of the low pH zone is provided by thefact that the footwall zones to each ore horizon
retain evidence of ore-related dolomite crystals
partly dissolved by the acidic fluids as they move
upwards through the sequence.
Roll-front type
Although somewhat similar to tabular deposits,roll-front ores form as a consequence of very

PART 2 HYDROTHERMAL PROCESSES

BrushyBasinMemberMiddle silt
Paleohydrologicalgradient
Salt Wash
MemberTony M.
orebodyMorrisom
Formation(a)Frank M.
orebody

Mudstone

Sandstone
Salt Wash
Member212

Middle silt

15 m

Tidwell
MemberTidwell
MemberTidwell
MemberSalt Wash
Member

(b)

pH
7
t=1
(initial mixing of meteoric/brinewaters)
Meteoricwater+
2–(Ca , HCO3, VO , UO2(CO3)2 )2+

Dolomite

Zone ofmixing
V-Clay

Brine
+
2+
2+
–
2–
(Na , Mg , Ca , Cl , SO4 )t=2(upward migration of mixing
zone)
Meteoricwater
Orezone {
1

pH
7

Organic
debris

Coffinite V-Clay
Brine

Partly dissolved
residual dolomite

V-claypH7t=3(continued upward migration of
mixing zone)

Stackedmineralizedintervals

Scale

Coffinite V-Clay
Brine

Ore
zone 2
Orezone 1

V-clayFigure 3.36 (a) The orebody
geometry and nature of host rocks
for the tabular uranium–vanadium
deposits of the Henry Basin areaof Utah, considered typical of theColorado Plateau region as a whole.
(b) Secular genetic model for thetabular, stacked orebody geometries
of Colorado Plateau type deposits,
involving an upward-migrating
mixing zone between a lower
stagnant brine and an overlying,ore-bearing, aquifer-focusedmeteoric fluid (after Northropand Goldhaber, 1990).

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3(a)Relevant
ionic
speciesin solution
UO2(CO3)22–
UO2(CO3)34–
V4O124–
AsO42–
SeO32–MoO42–

213
f(b)
Oxidized

Shale(imperm
eable)
Sands
tone (aquifer)

Altered(oxidized)
Solution flow
Alterationenvelope

Uraniumore zone

Hematite
Calcite leached
Metals leached
Organic C destroyed
Feldspars altered
Fe-Mg silicates alteredrmeable)Pitchblende/coffinite Mo haloOre-stage pyrite
Unaltered silicates
CalcitePre-ore pyrite, calcite
andGoethite Organic C
carbonaceous mattersiderite V, Mo, Co, Cu
As, Sepreserved
Se, S

Euhedral/framboidal
pyriteer

ReducedPre-oreUnaltered(reduced)

Shale (im
pe

Aquif

tion

Roll-frO

2

H

sulfida2S

ont formationMixingzone
Alteration envelopefFigure 3.37 (a) Cross section through a typical roll front type sandstone-hosted uranium deposit, such as thosein Wyoming, South Dakota, and south Texas. Uranium mineralization occurs at the interface (or redox front)
between the altered/oxidized and unaltered/reduced portions of the same aquifer (after De Voto, 1978; Misra, 2000).
(b) A two-fluid mixing model for roll front deposits. Initially a reduced fluid migrates up a syn-depositional fault (f)and sulfidizes the sandstone (top box). This is followed by down-dip migration of an oxygenated, uranium-bearing
meteoric fluid and the precipitation of uranium ore at the redox front formed in the mixing zone (bottom box).
This model is after Goldhaber et al. (1978).

different chemical processes. In its simplest form
the ore genesis model for roll-front deposits thatbest accommodates their geometry, in both planand section, is one where an oxidized, meteoric
fluid transporting soluble uranyl–carbonate complexes flows along a sandstone aquifer and precipitates uranium ore at a redox front (Figure 3.37a).
Up-dip of the roll-front the sandstones are altered(detrital silicate minerals are altered to clays,
Fe–Ti oxide phases are leached) and oxidized (secondary hematite has formed, organic carbon isbiodegraded). The redox front itself is characterized by an inner alteration envelope (marked bygoethite, siderite, pyrite, or marcasite) and anouter ore zone (comprising pitchblende and/or
coffinite with pyrite and some organic carbon).
Down-dip of the redox front the sandstone isrelatively unaltered and contains a more reduced
assemblage of pre-ore pyrite, calcite, and organic
matter, with detrital mineral phases relatively
intact (Figure 3.37a). In this model the redox front
can be regarded as the position in space and time
where the meteoric fluid has lost its capacity tooxidize the sandstone through which it is per-colating. Eh of the fluid changes dramatically at
the redox front and soluble uranyl–carbonate
complexes are destabilized with precipitation of
uranous oxide or silicate minerals. Other soluble
metal–oxide complexes also present in the fluid(such as V, As, Se, Mo, Cu, Co) are likewise
reduced and precipitate as various minerals,
either before or after uranium depending on theirrelative solubilities as a function of Eh. The rollfront does not remain static and it will migrate
down the paleoslope as the meteoric fluid isrecharged and the process of progressive, downdip oxidation of the sandstone evolves with time.
The down-dip migration of the redox front can beequated to the process of zone-refining and theore body is “frozen in” only once the paleohydrological regime changes and fluid flow ceases.
A modification of this scheme has been proposedby Goldhaber et al. (1978, 1983), who noted thatsome roll-front deposits are spatially associatedwith syn-depositional faults, and that the uraniumores are superseded by paragenetically later sulfide
minerals. A fluid mixing model was proposed toaccommodate these features and it is envisaged

214

PART 2 HYDROTHERMAL PROCESSES

that a reduced fluid migrated up the fault andsulfidized the sandstone aquifer for several hundred meters on either side (Figure 3.37b). This
stage of ground preparation created the reduced
environment responsible for locating the orerelated redox front. Evidence for this additional
stage in the mineralization process is provided byeuhedral pyrite (in deposits that do not containorganic matter) or framboidal pyrite (in those
deposits that do). This stage was followed by, or
was largely coincidental with, down-dip migra-

tion of an oxidized, uranium-bearing meteoric
fluid and the subsequent precipitation of uranium
ore at a redox front that, in this case, is represented by a zone of fluid mixing (Figure 3.37b).
In certain cases, such as the deposits of south
Texas, continued emanation up the fault of sulfidebearing solutions from deep within the basin resulted in post-mineralization sulfidization andthe growth of late-stage pyrite and marcasite that
overprinted even the altered up-dip sections of thesandstone aquifer.

There are many different types of fluids circulating
through the Earth’s crust and these have given rise
to a wide variety of hydrothermal ore deposit types
that exist in virtually every tectonic setting andhave formed over most of Earth history. Juvenile
fluids exsolve from magmas, in particular thosewith felsic compositions, and give rise to granite(sensu lato) related ore deposits that include porphyry Cu–Mo, skarn and greisen related Sn–W,
intrusion linked Fe oxide–Cu–Au, and high sulphidation Au–Ag ores (discussed in Chapter 2).
Metamorphic fluids, derived from volatiles liberated during prograde mineral reactions, are typically aqueo-carbonic in composition and areassociated worldwide with orogenic gold deposits
that are particularly well developed in Archean
and Phanerozoic rocks. Connate fluids formed
during diagenesis interact with either reduced(forming Pb–Zn dominant ores) or oxidized (forming Cu dominant ores) sedimentary environments.
These fluids are implicated in the formation ofMississippi Valley type (Pb–Zn) and stratiform
sediment-hosted copper ores. Circulation of nearsurface meteoric waters can dissolve labile constituents such as the uranyl ion, giving rise to avariety of different sediment-hosted uranium
deposits. Finally, sea water circulating throughthe oceanic crust in the vicinity of (ridge-related)fracturing and volcanic activity is vented onto thesea floor as black smokers, providing a modern
analogue for the formation of Cu–Zn-dominated
VMS deposits. Similar exhalative processes also

occur in different tectonic settings and with different metal assemblages, giving rise to sedimenthosted, Zn–Pb-dominated SEDEX type deposits.
The various hydrothermal deposits describedin Chapters 2 and 3 and their relationship to thedifferent fluid types discussed in these sections
are summarized in Figure 3.38.
Formation of hydrothermal ore deposits islinked not only to the generation of significant
volumes of fluid in the Earth’s crust, but also to
its ability to circulate through rock and be focused
into structural conduits (shear zones, faults, breccias, etc.) created during deformation. The abilityof hydrothermal fluids to dissolve metals provides
the means whereby ore-forming constituents areconcentrated in this medium. Temperature andcomposition of hydrothermal fluids (in particular
the presence and abundance of dissolved ligands
able to complex with different metals), together
with pH and fO2, control the metal-carrying capability of any given fluid. Precipitation of metals
is governed by a reduction in solubility which canbe caused by either compositional changes (interaction between fluid and rock, or mixing withanother fluid), or changes in the physical parameters (P and T) of the fluid itself. Economically
viable hydrothermal ore deposits occur when alarge volume of fluid with a high metal-carrying
capacity is focused into a geological location thatis both localized and accessible, and where efficient precipitation mechanisms can be sustained
for a substantial period of time.

HYDROTHERMAL ORE-FORMING PROCESSES CHAPTER 3WATER

HYDROTHERMAL ORE DEPOSIT TYPES

Groundwater/sea water

VMS/SEDEX
Kuroko, Cyprus
Broken Hill, Red Dog

Sandstone-hosted/surficial
uranium
Colorado plateau, Yeelirrie

Shallow
circulation

Approximate
depth

Me

teo
ri c

Stratiform diagenetic Cu
Zambian Copperbelt,
Kupferschiefer

Deep
circulation

orph
t am
Me aterw

Epigenetic Au
in quartz-pebble conglomerate
Witwatersrand, Sierra
Jacobina
Skarn
King Island,
MacTungiccmati
Mag ertwaEpithermal
Kasuga,Hishikari

Iron oxide-CuAu
Carajas,Olympic Dam

Mississippi Valley
Viburnum trend
Hydrothermal
solution

Connatewater

215

Porphyry Cu/Mo
Escondida,
Henderson
Sn–W greisens
Cornwall

Orogenic gold
Kalgoorlie,
CarlinChapter 2
Chapter 3

Figure 3.38 Diagram illustrating the relationship between different fluid types and various hydrothermal ore
deposit types. The diagram is relevant to both Chapters 2 and 3, and is modified after Skinner (1997).

For those readers wishing to delve further into hydrothermal ore-forming processes, the following is a selection of references to books and journal special issues.
Geochemistry of Hydrothermal Ore Deposits, ed.
H.L. Barnes. 1st edn 1967. Holt, Rinehart and Winston
Inc., pp. 34–76. 2nd edn 1979. John Wiley and Sons,
pp. 71–136. 3rd edn 1997. John Wiley and Sons,
pp. 63–124 and 737–96.
Gill, R. (1996) Chemical Fundamentals of Geology.
London: Chapman & Hall, 290 pp.
Guilbert, J.M. and Park, C.F. (1986) The Geology of
Ore Deposits. London: W.H. Freeman and Co., 985 pp.
Pirajno, F. (1992) Hydrothermal Mineral Deposits. New
York: Springer-Verlag, 709 pp.
Misra, K.C. (2000) Understanding Mineral Deposits.
Dordrecht: Kluwer Academic Publishers, 845 pp.

Reviews in Economic Geology: Volume 8: Barrie, C.T.
and Hannington, M.D. (eds) (1999) Volcanic-associatedMassive Sulfide Deposits: Processes and Examples inModern and Ancient Settings. El Paso, TX: Society of
Economic Geologists, 408 pp.
Reviews in Economic Geology: Volume 10: Richards,
J.P. and Larson, P.B. (eds) (1999) Techniques inHydrothermal Ore Deposits Geology. El Paso, TX:
Society of Economic Geologists, 256 pp.
Reviews in Economic Geology: Volume 13: Hagemann,
S.P. and Brown, P.E. (eds) (2000) Gold in 2000. El Paso,
TX: Society of Economic Geologists, 559 pp.
Wolf, K.H. (ed.) (1976) Handbook of Strata-bound andStratiform Ore Deposits, volumes 1–14. New York:
Elsevier.

Sedimentary/Surficial
Processes

Surficial and supergene
ore-forming processes

PRINCIPLES OF CHEMICAL WEATHERING
dissolution and hydrationhydrolysis and acid hydrolysis
oxidation–reduction
cation exchange
FORMATION OF LATERITIC SOIL/REGOLITH PROFILESBAUXITE ORE FORMATIONNICKEL, GOLD, AND PGE IN LATERITES
CLAY DEPOSITSCALCRETES AND SURFICIAL URANIUM DEPOSITS
SUPERGENE ENRICHMENT OF COPPER AND OTHER METALS

4.1 INTRODUCTION
Once metals have been concentrated in the
Earth’s crust and then exposed at its surface, they
are commonly subjected to further concentration by chemical weathering. The relationshipbetween weathering and ore formation is often
a key ingredient that leads to the creation ofa viable deposit and many ores would not bemineable were it not for the fact that gradeenhancement commonly occurs in the surficial
environment. There are several deposit types
where the final enrichment stage is related tosurficial weathering processes. Some of these
deposit types are economically very importantand contain ores, such as bauxite, that do notoccur in any other form. Ore-forming processes inthe surficial environment are intimately asso-

Box 4.1 Lateritization – the Los Pijiguaos bauxite
deposit, Venezuela
Box 4.2 Supergene and “exotic” copper – the porphyry
copper giants of northern Chileciated with pedogenesis, or soil formation, which
is a complex, multifaceted process reflecting local
climate and the dominantly chemical interactions
between rock and atmosphere. Soil itself is anextremely valuable resource since the world’s
food production is largely dependent on its existence and preservation. This chapter examines
the concentration of metals in soil, as well as in
the associated regolith (i.e. all the unconsolidated
material that rests on top of solid, unalteredrock). Many different metals are enriched in thesurficial environment, the most important of
which, from a metallogenic viewpoint, include
Al, Ni, Mn, Fe, Cu, Au, Pt, and U. Emphasis isplaced here on laterites, with their associatedenrichments of Al, Ni, Au, and platinum groupelements (PGE). In addition, clay deposits, calcretes, and associated uranium mineralization,
as well as the supergene enrichment of Cu in
porphyry-type deposits, are discussed. Surficial
enrichment of metals takes place in many othersettings too, and is a key process, for example, inthe formation of viable iron ores associated withbanded iron-formations. The latter are describedin more detail in Chapter 5 (Box 5.2).

220PART 3 SEDIMENTARY/SURFICIAL PROCESSES

4.2 PRINCIPLES OF CHEMICAL WEATHERING
From a metallogenic point of view chemical
weathering can be subdivided into three processes:• Dissolution of rock material and the transport/
removal of soluble ions and molecules by aqueous
solutions. Some of the principles applicable tothis process at higher temperatures were discussed in Chapters 2 and 3.
• Production of new minerals, in particular clays,oxides and hydroxides, and carbonates. Again,
this topic was discussed briefly in the section onhydrothermal alteration in Chapter 3.
• Accumulation of unaltered (low solubility)
residual material such as silica, alumina, and gold.
The main chemical processes that contributeto weathering include dissolution, oxidation,hydrolysis, and acid hydrolysis (Leeder, 1999).
Once weathering has commenced and fine clay
particles have been produced, cation exchange
further promotes the breakdown of minerals in
the weathering zone. Each of these processes isrelevant to ore formation in the surficial environment and is discussed briefly in turn below.
4.2.1 Dissolution and hydration
Certain natural materials such as halite (NaCl)and other evaporitic minerals, as well as the carbonate minerals (calcite, siderite, dolomite, etc.),
tend to dissolve relatively easily and completely
in normal to acidic groundwaters. This type of
dissolution contrasts with the breakdown of mostrock-forming silicate minerals which dissolve
less easily and do so incongruently (i.e. only certain components of the mineral go into solution).
The relative solubilities of different elements insurface waters depends on a variety of factors, but
can be qualitatively predicted (Figure 4.1) interms of their ionic potential (or the ratio ofionic charge to ionic radius). Cations with low
ionic potentials (<3) are easily hydrated and aremobile under a range of conditions, although they
will precipitate under alkaline conditions andare readily adsorbed by clay particles. Similarly,
anions with high ionic potentials (>10) form soluble complexes and dissolve easily, but will precipitate together with alkali elements. Ions withintermediate values (ionic potentials between 3
and 10) tend to be relatively insoluble and precipitate readily as hydroxides.
Over the pH range at which most groundwaters
exist (5–9), silicon is more soluble than aluminum(Figure 4.2) and consequently chemical weathering will tend to leach Si, leaving behind a residual
concentration of immobile Al and ferric oxides/
hydroxides. This is typical of soil formation processes in tropical, high rainfall areas and yields
lateritic soil profiles, which can also contain concentrations of bauxite (aluminum ore) and Ni.
Lateritic soils will not, however, form underacidic conditions (pH < 5) as Al is more soluble
than Si (Figure 4.2) and the resultant soils (podzols) are silica-enriched and typically depleted in
Al and Fe.
Another process that contributes to the dissolution of minerals in the weathering zone ishydration. In aqueous solutions water molecules
cluster around ionic species as a result of theircharge polarity and this contributes to theefficacy of water as a solvent of ionic compounds.
Hydration of minerals can also occur directly(Bland and Rolls, 1998), good examples of which
include the formation of gypsum (CaSO4.2H2O)
from anhydrite (CaSO4) and the incorporationof water into the structure of clays such asmontmorillonite. Mineral hydration results in
expansion of the lattice structure and assistsin the physical and chemical breakdown of thematerial.
4.2.2 Hydrolysis and acid hydrolysis
Hydrolysis is defined as a chemical reaction inwhich one or both of the O–H bonds in the watermolecule is broken (Gill, 1996). Such reactions
are important in weathering. One example occurs
during the breakdown of aluminosilicate minerals such as feldspar, and also the liberation ofsilicon as silicic acid into solution, as shown by
the reaction:
Si4+ + 4H2O ⇔ H4SiO4 + 4H+[4.1]

Although quartz itself is generally insoluble, it
is feasible to mobilize silica over a wide range ofSURFICIAL AND SUPERGENE ORE-FORMING PROCESSES CHAPTER 4

Soluble oxy-anion complexes(precipitate withalkali elements)

Insoluble hydroxides immobile
in acid and alkali
conditions

Oxy-anions

Cation charge (valency)

5

N

4

C

3

2

=1
0

Cation

Mo

3
l=
tia
n
e
t
po
ic
Ion
Solublecations

P V

Si
B

Sn

Ti Zr Mn

Be

Cu Mn Mg Fe Ni

Li Cu

0.5

UTh

HydrogenMobile as hydrated cations(precipitate under alkalineconditions, adsorbed
on clays)REE

Al Fe V Cr

10

OxygenInsolublehydrolysates

Ion
ic p
ote
ntia
l

S

6

221

Ba

Ca Sr Pb
K

Na

1.0
Ionic radius (Å)

Rb

1.5

2.0

Figure 4.1 Simplified scheme on the basis of ionic potential (ionic charge/ionic radius) showing the relative
mobility of selected ions in aqueous solutions in the surficial environment (modified after Leeder, 1999).

pH (see Figure 4.2b) by hydrolysis reactions that
result in the formation of relatively soluble silicic
acid (H4SiO4). Another example relates to thehydrolysate elements such as Fe and Al (Figure 4.1)
that are relatively soluble in acidic solutions,but will precipitate as a result of hydrolysis. The
hydrolysis of aluminum, yielding an aluminumhydroxide precipitate, is illustrated by reaction[4.2]:(i.e. those with a net charge defect) react withacids (H+ ions) in solution. This process displaces
metallic cations from the crystal lattice (see section 4.4 below), which then either go into solution
or precipitate. The process is most active at surfaces exposed by fractures, cleavages, and latticedefect sites.

Al3+ + 3H2O ⇔ Al(OH)3 + 3H+

Oxidation (and reduction) refers essentially tochemical processes that involve the transfer ofelectrons. In the surficial environment oxygen,present in either water or the air, is the mostcommon oxidizing agent. The element mostcommonly oxidized in the surficial environmentis probably iron, which is converted from theferrous (Fe2+) to the ferric (Fe3+) valence stateby oxidation (loss of electrons). An example[4.2]

It is this type of process that results, for example,
in the concentration of aluminum (as gibbsite
Al(OH)3) and ferric iron (as goethite FeO(OH)) inlateritic soils.
Acid hydrolysis refers to the processes whereby
silicate minerals break down in the weatheringzone. In this process activated mineral surfaces

4.2.3 Oxidation

222

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

(a)

a

Silic

(Insoluble)(Soluble)

(b)

min

–4

Below pH = 4.5
Al more soluble
than Si

A lu

25°C

300

–6

–8

250
SiO2 (amorphous)
200
150
H3SiO4–

Al(OH)4–
soluble

Al(OH)3
insoluble–1

Al3
solubleSi concentration (mg l )a

Log molar concentration

6100
H2SiO4

50

5

6

7

8

9

10

11

pH
5

7

9

pH
Figure 4.2 (a) The solubility of Si and Al as a function of pH (after Raiswell et al., 1980). (b) The solubility ofamorphous silica at 25 °C (after Bland and Rolls, 1998).

of the role of oxidation in chemical weathering is provided by the relative instability ofbiotite compared to muscovite. Biotite has theformula K+[(Mg2+, Fe2+)3(Si3Al)O10(OH)2]− and ismuch more easily weathered than muscovite
K+[(Al2)(Si3Al)O10(OH)2]− because of the ease withwhich the ferrous iron can be oxidized to ferriciron. Weathering and the resulting oxidation of Fe
in biotite leads to a charge inbalance that destabilizes the mineral, a process that is less likely tohappen in muscovite since it contains no iron inits lattice. The presence of iron in minerals suchas olivine and the orthopyroxenes is one of themain reasons why they are so unstable in theweathering zone.

4.2.4 Cation exchange
Clay particles are often colloidal in nature (i.e.
they have diameters < 2 µm) and are characterized
by a net negative surface charge brought about by
the replacement of Si4+ by Al3+ in the clay lattice.
The negative charge is neutralized by adsorptionof cations onto the surface of the colloids (seeFigure 3.16, Chapter 3). The adsorbed cations may
be exchanged for others when water passes throughweathered material containing clay colloids andthis has an effect on mineral stabilities as well as
the nature of leaching and precipitation in regolith
profiles. A simplified scheme (after Bland and Rolls,
1998) illustrating the effect of cation exchange at

SURFICIAL AND SUPERGENE ORE-FORMING PROCESSES CHAPTER 4

the surface of a colloidal particle is shown in reaction [4.3] below. In this example cation exchange
promotes the solution of Ca in groundwaters:
COLLOID – Ca2+ + 2H+(aq) →COLLOID – H+ + Ca2+(aq)
|
H+

[4.3]

4.3 LATERITIC DEPOSITS
4.3.1 Laterite formation

Lag
Soil
Fe laterite Lateritic gravel
Residuum Lateritic duricrust
Mottled zone
(kaolinite matrix)
Plasmic zone, mainly
kaolinite and goethite(Primary fabricdestroyed)

Saprolite>20% weatherable
minerals altered(Primary fabricpreserved)
Saprock<20% weatherable
minerals altered
Unaltered bedrock

Weathering front

Saprolith

Regolith

Pedolith(a) Generalized lateritic
regolith profile

Cementation front

Laterite is defined as the product of intense
weathering in humid, warm, intertropical regionsof the world, and is typically rich in kaolinitic
clay as well as Fe- and Al-oxides/oxyhydroxides.
Laterites are generally well layered, due to alternating downward percolation of rainwater andupward movement of moisture in the regolith
during seasonal dry spells, and are often cappedby some form of duricrust. Laterites are economically important as they represent the principal
environment within which aluminum ores (bauxite) occur. They can also contain significant
concentrations of other metals such as Ni, Mn,
and Au, as well as Cu and PGE. Laterites form on

223

stable continental land masses, over long periodsof time. In parts of Africa, South America, andAustralia laterization has been ongoing for over
100 million years (Butt et al., 2000), and, consequently, they are characterized by thick regolith
profiles (up to 150 m) in which intense leachingof most elements has occurred such that they nolonger reflect the rock compositions from which
they were derived. As with most soil/regolith
profiles the main processes involved in lateritization can be subdivided into those of eluviation,
where clays and solutes are removed from a particular horizon, and illuviation, where materialis accumulated, usually at a lower level. A generalized lateritic regolith profile is shown in Figure
4.3, together with an indication of the pattern ofhorizontally orientated leaching and retention of avariety of different elements in the regolith zones(Butt et al., 2000).
The base of a lateritic regolith profile is characterized by the saprolith zone, which is highly
weathered rock where the primary texture andfabric is still preserved. Given that the fluids
involved in this type of weathering are typicallyoxidizing and slightly acidic, the lowermost
saprock zone (Figure 4.3) is characterized by thedestabilization of sulfides and carbonates and theassociated leaching of most chalcophile metals(b) Element mobility in typicallateritic profiles
Host minerals

LeachedPartly retained (in secondary minerals)

Released in the mottled andferruginous zones
Aluminosilicates (muscovite, kaolinite)iron oxides; gold

K, Rb, Cs
Trace elements; Au

Si, Al (kaolinite)

Released in upper saprolite
Aluminosilicates (muscovite)
Ferromagnesians (chlorite, talc,
amphibole)
Smectite clays

Cs, K, Rb
Mg, LiSi, Al (kaolinite)
Fe, Ni, Co, Cr, Ga, Mn, Ti, V (Fe and Mn oxides)

Ca, Mg, Na

Si, Al (kaolinite)

Released in the lower saprolite
Aluminosilicates
Ferromagnesians (pyroxene, olivine
amphiboles, chlorite, biotite)

Ca, Cs, K, Na, Rb
Ca, MgSi, Al (kaolinite); Ba (barite)
Fe, Ni, Co, Cr, Ga, Mn, Ti, V (Fe and Mn oxides)

As, Au, Cd, Co, Cu, Mo,
Ni, Zn, S
Ca, Mg, Fe, Mn, Sr

As, Cu, Ni, Pb, Sb, Zn, (Fe oxides; sulfates,
arsenates, carbonates, alunite-jarosite)

Released at weathering front
SulfidesCarbonates

Figure 4.3 (a) A generalized lateritic regolith profile showing the different horizons and the terminology used intheir description. (b) Generalized pattern of element mobility in lateritic regoliths (after Butt et al., 2000).

224PART 3 SEDIMENTARY/SURFICIAL PROCESSES

and alkaline/alkaline earth elements. The lower
saprolite zone is characterized by the destructionof feldspars and ferromagnesian minerals, with Si
and Al retained in clay minerals (kaolinite andhalloysite). Fe oxides/oxyhydroxides also form inthis zone with partial retention of some transition metals in phases such as hematite andgoethite. The mid- to upper-saprolite zone sees
alteration of all but the most resistant of minerals
as well as destruction of earlier formed secondary
minerals such as chlorite and smectite. Only minerals such as muscovite and talc tend to survive
intact through this zone. The upper part of theregolith profile, the pedolith zone, is characterized by complete destruction of rock fabric andleaching of all but the most stable elements. This
zone is dominated compositionally by Si, Al, andferric Fe occurring mainly in kaolinite, quartz,
and hematite/goethite. A ferruginous residualzone is best developed over mafic/ultramafic
bedrock, whereas kaolinite is more abundant overfelsic lithologies. The zone is also characterized
by accretionary and dissolution textures such as
pisoliths and nodules, or kaolinite replacement by
gibbsite or amorphous silica. Some metals (Au,Cr, V, Sc, Ga) tend to be associated, by adsorption,with Fe oxides/oxyhydroxides in the ferruginous
residuum, whereas other elements are retained in
the upper zones simply because their mineralhosts are particularly stable (i.e. Zr, Hf in zircon,
Cr in chromite, Ti in rutile, etc.).
Laterites can be subdivided into ferruginous(ferricretes) and aluminous (alcrete or bauxite)
varieties. Ferricretes are characterized by a ferruginous residuum in the upper zone, where Fe2O3
contents can be as high as 20%. They tend to form
under specific climatic conditions where rainfall
is less than about 1700 mm per year but averagetemperatures are high (Tardy, 1992). Under higher
rainfall conditions ferricrete dissolution tends tooccur and aluminum hydroxide (gibbsite Al(OH)3)
accumulates.
4.3.2 Bauxite ore formation
Bauxitic ore, in the form of the minerals gibbsite/
boehmite and diaspore, is the principal source ofaluminum metal, demand for which increased

dramatically in the latter half of the twentiethcentury. The accumulation of an alumina-rich
residuum, as opposed to one enriched in iron, inthe upper zone of a lateritic profile is a function ofhigher rainfall, but also lower average temperatures (around 22 °C rather than 28 °C for ferricretes) and higher humidity (Tardy, 1992). Actual
alumina enrichment in the upper parts of laterite
profiles is due, at least in part, to relatively high Si
mobility compared to Al, and probably reflects
near neutral pH conditions (between 4.5 and 9; see
Figure 4.2a). This results in incongruent dissolution of minerals such as felspar and kaolinite,where Si is leached in preference to Al, yielding
a gibbsite-like residue. This process is shown
schematically (after Bland and Rolls, 1998) as:
feldspar – (loss of Si) → kaolinite – (loss of Si)[4.4]→ gibbsite (Al(OH)3)
Seasonal climatic variations are also considered
important to the formation of bauxitic ores as
the alternation of wet and dry spells promotes
fluctuations in groundwater levels and, hence,
dissolution and mass transfer. Variations inbauxitic profiles, as well as transformation fromhydrated gibbsite to the relatively dehydrated
version, boehmite, or to diaspore (AlO(OH)), result from such fluctuations. The mineralogical
profiles in zones of bauxite mineralization may bequite variable. In humid, equatorial laterite zones,
hydrated minerals such as gibbsite and goethitepredominate, whereas in seasonally contrasted
climates the ores are relatively dehydrated andboehmite–hematite assemblages form (Tardy,1992).
The redistribution of iron, and the segregationof Al and Fe, is a necessary process in bauxite
formation because ferruginous minerals tend tocontaminate the ore. High quality bauxitic ores
require that both Fe and Si be removed, but notalumina, whereas ferricretes and conventional
laterites are characterized by different combinations of element leaching. The interplay of Ehand pH is critical to the formation of high quality
bauxitic ores as discussed in some detail by
Norton (1973). The essential features of Norton’smodel are shown in Figure 4.4. In terms of Eh andSURFICIAL AND SUPERGENE ORE-FORMING PROCESSES CHAPTER 4

1.2

–8

–4

10

0.6

3
2

25°C

O
H

2

2O

4 Eq
1

of
lity
lubi
l so
ua

0.4
Eh (volts)

10–10
10–8
Σ Al = 10–6

10

0.8

ΣFe = 10–6

1.0

Al minerals more soluble
than Fe minerals
Fe minerals more soluble
than Al minerals

5

0.2

6

Fe
nd
Al a

0.0
–0.2

H
2O
H

He
m
Ma atit
e
gn
eti
te

2

–0.4
–0.6
–0.8
2

4

6

8

10

12

pH
Figure 4.4 Eh–pH diagram showing conditions
relevant to the formation of laterites and bauxite ore.
The solubility contours are in mol l−1 and assume
equilibrium of the solution with gibbsite andhematite/goethite (after Norton, 1973).

pH, Si is leachable by hydrolysis (see equation 4.1
above) and is increasingly soluble at higher pH,
whereas Al is most soluble at either very low orhigh pH (Figure 4.2a). Fe is most mobile as ferrous
iron at low Eh and pH. In lateritic environments
where concentrations of both Fe and Al occur, itis the rather special conditions where these two
metals are segregated that provide the means forhigh quality bauxite formation. The plotting of
Fe and Al solubility contours in Figure 4.4 allows
one to draw an isosolubility curve where Al andFe are equally soluble. Above this curve, at highEh and pH, Al minerals such as gibbsite will be
more soluble (and Al preferentially leached) thanbelow this curve where Fe minerals such ashematite or goethite will be more soluble (and Fe
preferentially leached). Using these constraints,

225

several situations can be described that are
relevant to laterite formation. The area of Eh–pH
space encompassing fields 1 and 2 on Figure 4.4
will be characterized by leaching of both Fe and Al
and will not result in laterite formation, but more
likely Si-enriched podzol soils. Fields 3 and 4 arecharacterized by restricted solubility of Fe and Al
and laterites and bauxites will only form if the
bedrock composition itself has high iron or aluminum contents. In field 3 bauxites are unlikely
to form because of preferential leaching of Al
from the soil, although over a protracted time
period and given the right climatic conditions,high Fe laterites will form under these conditions.
Groundwater solutions forming in field 5 willcontain more Fe than Al and laterites will form inthis environment. The optimum conditions forbauxite formation are provided in field 6 where
groundwater solutions will preferentially remove
Fe. In this field Al hydrolysates are stable, especially at pH between 5 and 7, and gibbsite willaccumulate. This model shows that both the initial and subsequent Eh and pH of the groundwater
solutions involved will impose important controls on the relative mobilities of Al and Fe and,
consequently, on the dissolution and precipitation of relevant minerals during laterite and bauxite formation. It should be noted that disequilibria
and the role of organic acids in element mobility
might complicate some of the generalizations madeabove and could account for the variability that issometimes observed in actual weathering profiles.
The huge bauxite deposits of Jamaica represent
a good example of the controls involved in theformation of aluminum ores. Thick deposits ofbentonitic volcanic ash were laid down on alimestone bedrock and the former are considered
to be the ultimate source of residual alumina inthe bauxite deposits. Desilication of volcanic
glass and other silicate minerals by rapid andefficient drainage through the underlying karst
limestone is considered to have been responsible
for the gibbsite-dominated ores (Comer, 1974).
In Venezuela, by contrast, lateritic bauxites arederived by deep weathering of a granitic bedrock(Meyer et al., 2002) during prolonged uplift anderosion of the Guyana Shield starting about 35
million years ago (see Box 4.1).

226PART 3 SEDIMENTARY/SURFICIAL PROCESSES

Lateritization: Bauxite: the Los Pijiguaos
deposit, Venezuela
Approximately 90% of the world’s bauxite, the principal
source of alumina, is obtained from laterites formed bydeep weathering of continental planation surfaces exposed
to hot humid climates. Global production of bauxite
has risen significantly in recent years and is presently ataround 125 million tons per annum (Meyer et al., 2002).
One of the largest bauxite deposits in the world is at Los
Pijiguaos in Venezuela, where annual production is about
5 million tons from a total resource that could be as muchas 6 billion tons.
The Los Pijiguaos bauxite deposit formed from a
Mesoproterozoic, metaluminous Rapakivi-type granite (theParguaza granite) that was deeply eroded and weathered
during Cretaceous–Paleogene times to form the Nuria
planation surface (Figure 1). This profound weathering
event gave rise to the development of thick lateritic soil
cover in which economic concentrations of bauxite(around 50 wt% Al2O3) form a mineable zone that averages 8 m thick. The complete laterite profile comprises
an upper concretionary zone (bauxite), underlain by amottled or saprolitic layer and then the granitic bedrock.
The bauxite layer itself can be subdivided into four zones
comprising various concretionary and pisolitic textures.

The deposit as a whole is subdivided into blocks that are
defined by fluvial channels that cut through the deposit
subsequent to its formation (Figure 1).
The bauxite at Los Pijiguaos is made up principally ofgibbsite with lesser quartz and rare kaolinite. The original
granite textures in the bauxite layer have been almost
entirely destroyed, with feldspars, and to a lesser extent
quartz, being largely replaced by gibbsite (Figure 3). An
increasing degree of lateritization is marked by systematiccompositional changes in the rock. This is demonstrated
in the Al2O3–SiO2–Fe2O3 ternary plot in Figure 2 where
the process of lateritization is characterized by an increasein the alumina content. Quartz and gibbsite are inversely
correlated in the bauxite ore and iron contents remain
relatively constant, with an increase in Fe2O3 in the mosthighly lateritized samples. The bauxite at Los Pijiguaos is
a product of desilication, hydration and residual Al2O3
concentration of the Parguaza granite, and involves a massloss of some 60% (Meyer et al., 2002). These processes
produce a high grade and high purity gibbsitic ore fromwhich alumina can be extracted by low temperature digestion, making Los Pijiguaos one of the most commercially
profitable bauxite deposits in the world.
South

North

1000

Kamarata Surface(Jurassic)

Altitude (m)800
Nuria Surface
600

Los Pijiguaos Bauxite deposit(Cretaceous)

400
200
0

Incised drainage
Los Llanos Surface(Pliocene–Pleistocene)

0

10
km

Figure 1 Schematic profile through the Los Pijiguaos bauxite deposit showing the concentration of laterite on
a dissected late Cretaceous planation surface (after Meyer et al., 2002).

SURFICIAL AND SUPERGENE ORE-FORMING PROCESSES CHAPTER 4

227

SiO2

Kaolinization

PG

Weak lateritization

Moderate
lateritization

Figure 2 (a) Ternary
Al2O3–SiO2–Fe2O3 plot
showing the residual increase
in alumina content with
progressive lateritization in
the Los Pijiguaos bauxite ores.
PG is the composition of theunaltered granite.

Bauxitization50Stronglateritization

Al2O3

50

Fe2O3

Figure 3 Photomicrograph showing the replacementof quartz and feldspar by Kaolinite and gibbsite inbauxite ores from Los Pijiguaos (photograph courtesyof Michael Meyer).

4.3.3 Nickel laterites
Laterites that are developed on ultramafic rocks
containing abundant olivine and orthopyroxene
and, hence, with high Ni contents (i.e. 0.1 to
0.3%; see Chapter 2), commonly develop concentrations of Ni silicate or oxide that are up to 10
times the original concentration. Given the relative ease of mining of this type of ore comparedto hard-rock magmatic sulfide ores, as well as
recent developments in extractive metallurgicalprocesses, Ni laterites are a much sought after
commodity. The principles by which these ores
form are relatively straightforward and basicallyinvolve eluviation of Ni from the uppermost lateritic residuum and concentration in underlying
saprolitic illuvium as nickeliferous talc, serpentine, or smectite, or less commonly together withgoethite.
The primary olivine and orthopyroxene minerals
that are the source of the Ni form the main ingredients of an ultramafic rock that may originally228

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

have been part of an obducted ophiolite complex(in the case of the New Caledonia deposits inthe south Pacific) or a layered mafic intrusion(such as at Niquelandia in Brazil). These minerals
are rapidly serpentinized, in some cases prior tolateritization by interaction with sea water or
during low-grade metamorphism or alteration.
The alteration of olivine is by hydration to amorphous silica, serpentine, and limonite, with theserpentine able to accommodate as much or moreof the Ni as the original olivine. Most nickel
laterites form on bedrock that is already largely
altered to serpentinite (Golightly, 1981). In lateritic regoliths the groundwaters become progressively less acidic with depth (pH increases
downwards from around 5 to 8.5; Golightly, 1981)
and bicarbonate is the main anion in solution.
Olivine breaks down completely under suchconditions and this is followed by orthopyroxene,
serpentine, chlorite, and talc. An important
reaction in this process is given in [4.5] below(after Golightly, 1981) where olivine breaks
down to smectitic clays (saponite–nontronite) andgoethite:4(Fe2,Mg2)SiO4 + 8H+ + 4O2 ⇔olivine(Fe2,Mg3)Si4O10(OH)2 + 6FeO(OH) + 5Mg2+
smectitegoethite[4.5]
4.3.4 Gold in laterites

Once smectite, and also serpentine and talc, arepresent in the regolith, further concentration of
Ni takes place largely by cation exchange, mainlywith Mg2+. This results in the formation of avariety of unusual Ni enriched phyllosilicate
minerals such as kerolite (Ni–talc), nepouite(Ni–serpentine), and pimelite (Ni–smectite). An
example of such an ion exchange reaction, where
nepouite forms from serpentine, is shown in [4.6]
below (after Golightly, 1981):
Mg3Si2O5(OH)4 + 3Ni2+(aq) ⇔serpentine
Ni3Si2O5(OH)4 + 3Mg2+(aq)nepouitenierite,” after the Frenchman Jules Garnier who
discovered the enormous Ni laterite ores of New
Caledonia in the nineteenth century. Some Ni
concentration is also associated with goethite,
although the exact nature of the fixation mechanism is not known. It is possible that adsorptionof Ni onto goethitic colloids might occur atneutral to slightly alkaline pH, as discussed previously in Chapter 3 and shown in Figure 3.15b.
The limonite zone in the upper part of most laterite profiles (Figure 4.5) is, however, generally
depleted in Ni.
Enormous resources of Ni ore are located inultramafic rooted lateritic regoliths in many partsof the world. The thickest laterite and the richest
concentrations of garnierite tend to occur where
the bedrock is characterized by closely spaced
jointing, as this is where maximum groundwater
circulation and fluid–rock interaction takes place.
The best Ni laterite concentrations also appear
to display a topographic control and tend to occur
either beneath a hill or on the edge of a plateau
or terrace (Golightly, 1981). This is because these
deposits are sensitive to surface erosion and tofluctuations in water table levels that in turn control the distribution of zones of eluviation andilluviation.

[4.6]

The collective term applied to Ni–phyllosilicate
ore minerals in lateritic environments is “gar-It is well known that gold occurs in the upper
pedolithic portions of laterite weathering zones
in many parts of the world, including Brazil,
west Africa and Western Australia (Nahon et al.,1992). This gold takes many forms, ranging fromlarge, rounded nugget-like particles and gold
dendrites in cracks and joints, to small crystals
in pore spaces. In the Yilgarn Craton of Western
Australia, for example, the recovery of gold fromlateritic profiles is a major activity and some spectacular nuggets, occasionally over 1 kg in mass,
have been found. Whereas the primary sourceof gold in these environments is high in silver(strictly a Au–Ag alloy with typically 5–10% Ag
as well as minor Hg and other impurities), thegold that is concentrated in lateritic profiles has
been chemically purified. This suggests that differential mobilization and decoupling of Au andSURFICIAL AND SUPERGENE ORE-FORMING PROCESSES CHAPTER 4

229

MgO, Fe2O3, SiO2 (wt%)
Surface

10

20

30

40

50

60

70

80

Zone ofresidual laterite
1–20 meters(eluviation)

Pedolith

Iron crust
Pisolites
Red limonite

Al2O3
Ni
Cr2O3

Yellow limonite
Zone of alteredperidotite andnickel concentration
1–20 meters(illuviation)
Zone offresh peridotite

Saprolith

Fe2O3
SiO2

Ore withboulders(Garnierite,
Amorphous silica)
Rocky oreNi
MgO

Bedrock
1

2

3

4

5

6

7

Ni, Cr2O3, Al2O3 (wt%)
Figure 4.5 Descriptive profile and Ni ore distribution in a lateritic regolith typical of the New Caledonian deposits.
The chemical profile clearly distinguishes the ferruginous/aluminous residual zone where Si, Mg, and Ni areleached, from the saprolith where illuviation has resulted in concentration of Ni (after Troly et al., 1979; Guilbert
and Park, 1986).

Ag takes place as meteoric fluids percolate
through the weathering zone.
It is only under certain very specific conditionsin the near surface environment that Au and Ag
can be substantially mobilized. Actual measurements of near surface groundwaters in lateritic
environments indicate that aqueous solutions
are relatively acidic (pH < 5), oxidizing, and inWestern Australia have low to moderate salinities
because of a marine influence on rainfall (Mann,
1984). With increasing depth, groundwater pH
tends to increase (i.e. become more alkaline) andthe fluids become more reducing as the ferric/ferrous ratio decreases (Figure 4.6). The process oflateritization, described in section 4.3.1 above, iscontrolled essentially by hydrolysis as groundwaters react with the regolith and iron is oxidized.
Two reactions reflect the nature of the process
and indicate why groundwaters become more
acidic upwards in lateritic profiles. Reaction [4.7]describes the initial breakdown, by hydrolysis, of
sulfide phases such as pyrite in the lower sectionsof the weathering zone close to bedrock, and the
production of hydrogen ions:

2FeS2 + 2H2O + 7O2 ⇔ 2Fe2+ + 4SO42− + 4H+ [4.7]
Reaction [4.8] illustrates the oxidation of ferrous
iron in the vicinity of the water table, to form
goethite in the ferruginous pedolith, with further
production of hydrogen ions (after Mann, 1984):
2Fe2+ + 3H2O + O2 ⇔ 2FeOOH + 4H+
goethite[4.8]

These reactions illustrate some of the processes
active during lateritization and explain why
groundwaters are likely to be acidic in such environments. Laterites formed over felsic rocks areprone to be more acidic than those over a mafic
substrate, as the latter will have pH buffered bybicarbonate ions formed during alteration (Mann,
1984).
Experimental work (Cloke and Kelly, 1964;
Mann, 1984) indicates that at low pH and highEh, and in the presence of Cl− ions, gold in thesurficial environment can go into solution as an
AuCl 4− complex (Figure 4.7). Substantial Au dissolution only occurs, however, at relatively high Eh230PART 3 SEDIMENTARY/SURFICIAL PROCESSES(a)
Higher purity gold rim(Ag-depleted)

Peripheral
void

(b)
Iron oxyhydroxide
replacement halo,containing gold
(Ag-depleted)Iron
oxyhydroxide

Primary
Au–Ag
alloy

AgCl
O2

O2

Dissolution of Au, Ag

Water table
Low pHHigh Eh
AgCl

AgCl
Pedolithou
rce

Reprecipitation of Au

Q ua
rtz
v

ein
–

go

ld s

AuCl4–

Fe2+
Mafic

Saprolith

Fe2+

FelsicFe2+
Neutral pH
Low EhBedrock

Figure 4.6 The nature and characteristics of gold and silver redistribution in a typical lateritic profile. The inset
diagram shows two mechanisms whereby a primary Au–Ag alloy particle could be broken down during lateritizationand silver preferentially removed (modified after Mann, 1984; Nahon et al., 1992).

when Fe2+ is completely oxidized in the presenceof free oxygen. By contrast, silver will go into
solution more readily, and under more reducing
conditions, as several complexes, including AgCl(the most stable) and AgCl −2 or AgCl 32− (Figure 4.7).
Accordingly, particles of Au–Ag alloy in thesurficial environment will be more likely tobe leached of their Ag over a broader range ofconditions.
A model for gold concentration in laterites isshown in Figure 4.6 and is discussed below interms of gold– and silver–chloride speciation, as
well as the likely Eh–pH conditions of a typicallateritic profile, shown in Figure 4.7. Using theseSURFICIAL AND SUPERGENE ORE-FORMING PROCESSES CHAPTER 4

231

1.2
–

Figure 4.7 Eh–pH diagram
showing the characteristics ofa lateritic profile in relation tolacustrine and riverine waters. Also
shown are the redox potentials andstabilities of Au, Ag, Pt, and Pd as
chloride complexes, illustrating therange of Eh that controls solubilityof the various metals (informationcompiled from data in Fuchs andRose, 1974; Mann, 1984; Bowles,
1986).

Eh (volts)

0.8

AuCl4
Au
PtCl2–
4
Pt

2

H2 O

PdCl2–
4
Pd

Rive

0.4
Late
prof ritic
ile

AgCl
Ag

Pt O
2 .×
Pt (O H2 O
H)
Pt 2

r wa

Lak

ter

Pd (

OH)
Pd 2

ew

ater

0

–0.4

H2 O
H2

2

4

6

8

10

12

14

pH

constraints it is apparent that primary Au–Ag
alloy particles in the near surface regolith above
the water table will be readily dissolved in theacidic, oxygenated, and moderately saline groundwaters prevailing in this environment. Mann
(1984) has shown that the breakdown of the particle and dissolution of Au and Ag actually takes
place by replacement of the alloy by an iron oxyhydroxide phase (such as goethite). The goethiteitself, however, can be seen to contain tiny particles of high purity gold (Figure 4.6 inset b) suggesting that the latter metal is rapidly reprecipitatedas goethite forms, whereas silver remains in solution and is transported away. This is explainedby the fact that high gold solubilities are only
maintained at high Eh and that the metal will
precipitate with only a slight decrease in Eh under
mildly reducing conditions, as described by thefollowing reaction:
AuCl 4− + 3Fe2+ + 6H2O ⇔Au + 3FeOOH + 4Cl− + 9H+

O

Fe3+
Fe2+

[4.9]

A decrease in AuCl −4 solubility and precipitationof gold is, therefore, promoted by interaction of

the fluid with a more reducing regolith in which
Fe2+ and Mn2+ contents are higher with increasing
depth in the profile and below the water table.
Redox reactions, therefore, control the precipitation of gold in this environment and explain its
association with minerals such as goethite andmanganite. It is clear in Figure 4.6 that under thesame conditions in which gold precipitates, silver
as AgCl remains in solution and will be transported away from the site of gold deposition.
Primary Au–Ag alloy particles that are notclosely associated with an iron or manganese oxyhydroxide tend to show a preferential depletionin silver around their edges and the formation ofhigher purity gold halos (Figure 4.6 inset a). Suchfeatures are a product either of silver diffusionfrom the outer margins of the particle, or of complete dissolution of the entire alloy and in situ
reprecipitation of the gold. It should also be notedthat the efficacy of chloride ion ligands in promoting gold and silver dissolution is particularly
appropriate to laterites that presently occur inarid climates, such as Western Australia. In humid,
tropical regions chloride ion activity is reduced byhigh rainfall dilution and other compounds, such

232

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

as organic humic and fulvic acids, may play amore important role in gold dissolution (Nahonet al., 1992).
Finally, it should be noted that microorganisms
play an important role in the concentration ofgold in lateritic soils. Secondary gold, in theform of the nuggets found either in laterites or inother environments where placer gold deposits
occur, are sometimes characterized by sphericalmorphologies consistent with the size and shapeof bacteria (Watterson, 1991). Southam andBeveridge (1994) have shown experimentally that
Bacillus subtilis, a common bacterium in soils,
is capable of accumulating gold by diffusion
across the cell wall and precipitation within thecytoplasm. It is suggested that gold was stabilizedinternally as an organo-gold complex and thenprecipitated in colloidal form once a critical concentration had been reached. Subsequent diagenesis of sediment containing such gold-enrichedmicroorganisms would result in recrystallizationand coalescence of gold to form nugget like shapes.
4.3.5 A note on platinum group element (PGE)
enrichment in laterites
In addition to gold in laterites, it is now widely
appreciated that PGE can also be concentrated
in weathering profiles. Laterites in Sierra Leone,
for example, are known to contain PGE deposits(Bowles, 1986) and significant concentrations ofthese precious metals also occur in non-lateritic
soils such as over the Stillwater Complex inMontana and in Quebec (Fuchs and Rose, 1974;
Wood and Vlassopoulos, 1990). Evidence for themobilization of PGE in the weathering zone isprovided by the existence of well formed, crystalline Pt–Fe or Os–Ir–Ru alloys in pedolith, andin some cases PGE-sulfide minerals, that are compositionally different, and often larger, than thosein the source rock (Bowles, 1986). In the case oflaterite-hosted PGE concentrations, it is considered likely that the controls on Pt and Pd solubility were similar to those affecting Au and Ag andthat these precious metals readily go into solution
as chloride complexes in acidic, oxidized environments (Bowles, 1986). Figure 4.7 shows that PdCl 2−
4
is stable over a broader range of Eh and pH than

PtCl 2−
4 , a feature that is consistent with observations that palladium is more readily dissolved andconcentrated in groundwaters than is platinum(Wood and Vlassopoulos, 1990). Like gold andsilver, platinum and palladium tend to become
decoupled in the weathering environment, with
Pd going into solution and Pt concentrating insoils and sediment. Similarly, redox reactions
control the precipitation of PGE in this environment, with Pt, like Au, precipitating before Pd
as Eh is reduced. In non-lateritic environments,
where groundwaters are near neutral to alkaline
and more reducing, PGE will not be transported
as chloride complexes but more likely as neutral hydroxide complexes (PdOH2 and PtOH2).
Decoupling of Pt and Pd still occurs in nonlateritic weathering profiles, although it is suggested that whereas Pd is dispersed in solution, Pt,
together with Au, is less easily mobilized, perhaps
as colloidal particles (Wood and Vlassopoulos,
1990). Finally, it should be noted that there isevidence pointing to an effective interaction between PGE and organic acids, in particular humic
acid, in weathering zones (Wood, 1990; Bowles
et al., 1994). Pt and Pd are reasonably soluble innatural solutions containing humic acid, although
the exact nature of the complex involved is notknown. Humate is also thought to promote PGE
transport as colloidal suspensions, whereas solidorganic debris may also be responsible for precipitation of precious metals.
A spectacular example of where both Au andPGE have been concentrated into a deeply weathered lateritic profile is the Serra Palada depositin the Carajás mineral province of Brazil. Thisdeposit was discovered and exploited exclusively
by artisinal miners who took advantage of the richand easily extractable laterite-hosted ore. Average
grades are reported to be around 15 g ton−1 Au,
4 g ton−1 Pd, and 2 g ton−1 Pt, with the primary ore
body thought to resemble other Fe oxide–Cu–Au
deposits (such as the Olympic Dam deposit; seeBox 3.1, Chapter 3) found in the Carajás province(Grainger et al., 2002). Weathering has resulted
in a lateritic profile dominated by kaolinite,manganese oxides, and Fe oxides/oxyhydroxides.
Amorphous carbon is present as the weathered
residue of an originally carbonaceous siltstone.

233

SURFICIAL AND SUPERGENE ORE-FORMING PROCESSES CHAPTER 4

(b)

(a)

80

80
Smectite

MAFIC ROCKS
Kaolinite

Relative clay mineral content (%)

Relative clay mineral content (%)

FELSIC ROCKS

60

Vermiculite

40

Illite20

Smectite
60

Kaolinite

40
Vermiculite

20

Gibbsite
Gibbsite
0

60
20
40
Mean annual rainfall (inches)

80

0

20
40
60
Mean annual rainfall (inches)

80

Figure 4.8 The effect of rainfall on clay mineral formation in (a) felsic and (b) mafic rocks (after Barshad, 1966).

Most of the ore constituents have been extensively remobilized during lateritization, with basemetals occurring in Mn oxides and Au–Pt–Pd occurring together with amorphous carbon and Fe–Mnoxides/oxyhydroxides. Some of the ore also occurs
as nuggets of Au and PGE alloy, some of which arereported to weigh as much as 60 kg (Grainger
et al., 2002).

4.4 CLAY DEPOSITS
Clay minerals are volumetrically the most abundant product of weathering and either occur
in situ or have been transported to a site of deposition. They are also economically important andused for a variety of industrial applications including paper, ceramics, filtration, and lubricants.
Clay formed during weathering reflects both thenature of the source material and the weathering conditions. Nevertheless, very similar clay
mineral assemblages ultimately form from both
felsic and mafic rocks (Figure 4.8). Temperatureand rainfall in particular, as well as local Eh–pH
conditions, determine the nature and rates ofclay-forming processes. Some of the more
important clay minerals considered here include
kaolinite, illite and the smectite group (including montmorillonite). Kaolinite is formed underhumid conditions by acid hydrolysis of feldsparbearing rocks (e.g. in laterites), whereas illite
forms under more alkaline conditions by weathering of feldspars and micas. The smectite clays
commonly weather from intermediate to basic
rocks under alkaline conditions and are highly
expandable, containing intracrystalline layers ofwater and exchangeable cations. It should be
noted that clay minerals are not only the result ofweathering processes but can also form as the
products of low temperature hydrothermal alteration (see Chapter 3).
Clay formation is initiated during chemical
weathering by acid hydrolysis (see section 4.2.2
above and Figure 4.9). A typical reaction that
describes the process is shown below:(Ca,Na)Al2Si2O8 + 2H+ + H2O ⇔plagioclase
Al2Si2O5(OH)4 + Na+ + Ca2+
kaolinite[4.10]

234

PART 3 SEDIMENTARY/SURFICIAL PROCESSES(a)

Capillary fringeof H2O (with H+)

K-feldspar
Al
–

Si

H+

K+

O

O

–

+

K

–
–

H+
(b)
Defectlattice site

In the above reaction plagioclase feldspar is broken down during weathering to form kaolinite,with the release into solution of Na and Ca
ions. The actual mechanisms of weathering are
undoubtedly more complex and may involve
several incremental stages during the formationof clays. A schematic diagram illustrating thevarious stages involved during the weatheringand breakdown of K-feldspar by acid hydrolysis isshown in Figure 4.9. Alteration by acid hydrolysis
is most relevant at acidic pH and the process iskinetically less effective under more alkalineconditions, where other processes such as Sihydroxylation may be more efficient.
The mobility of different elements in the surficial
environment varies considerably (Figure 4.1) andthis has led to the recognition of a hierarchy ofmobility, as follows:[4.11]tionCa > Na > Mg > Si > K > Al = Fe

In

so

lu

H+

–

+

–

K

H+

H+
K+
–
–

H+
H2O

H+
Exchange of K+ for H+ at
surface (protonation)

(c)

H 2O
–

Si

–

K+

H+
–

O

H+

The alkaline and alkali earth elements are typically the most soluble in groundwaters, with
Al and ferric Fe being relatively immobile. The
combination of kinetic and thermodynamic constraints on clay mineral formation during weathering results in a well defined sequence of reaction
pathways that can be related, for example, toclimatic conditions. The effects of rainfall on theformation of clay mineral assemblages is shownin Figure 4.8, for both felsic and mafic parentalrock compositions. The general pattern is thatsmectitic clays tend to form under relatively arid
to semi-arid conditions, whereas kaolinite tends
to dominate in wetter climates. This pattern may
also reflect a paragenetic sequence of clay formation, with smectites forming relatively early on
and being superseded, or replaced, by kaolinite orvermiculite.

H+

In solution
as silicic acid– H+

Figure 4.9 (left) Representation of the stages during
the progressive weathering and breakdown of K-feldspar
by acid hydrolysis, leading to the formation of clay
minerals (modified after Leeder, 1999).

SURFICIAL AND SUPERGENE ORE-FORMING PROCESSES CHAPTER 4

Figure 4.10 Strongly kaolinized granite in Cornwall,being mined for “china clay.”

4.4.1 The kaolinite (china clay) deposits
of Cornwall
The Cornubian batholith of southwest England,so well known for its polymetallic magmatichydrothermal mineralization (Box 2.2 in Chapter 2), also hosts some of the largest and best
quality kaolinite (or china clay) deposits in theworld. Of considerable demand in the paperand ceramic industries, the kaolinite resourceof Cornwall surpasses the base metal resource invalue, and considerable reserves of high quality
china clay still exist.
The formation of the Cornish kaolinite deposits
is complex and controversial. Some workers haveargued in favor of a weathering-related origin(Sheppard, 1977), but this has been countered byproponents of a hydrothermal process who maintain that clay formation represents the end stageof a very long-lived paragenetic sequence initiatedsoon after granite emplacement and terminatedwith circulation of low temperature meteoric
fluids during clay formation (Alderton, 1993). The
actual processes of kaolinite formation are essentially the same irrespective of the actual fluid
origin and on balance it seems likely that both alate hydrothermal process and a tertiary weathering event contributed to clay formation.
Zones of intense kaolinization are best developed in areas where greisen altered Sn–W ore
veins and major NW–SE trending faulting haveoccurred. This suggests a genetic link to previ-

235

ous, higher temperature, hydrothermal processes,
but could also point to groundwater circulation
focused into sites previously accessed by hightemperature fluids. Kaolinite itself is preferentially developed from the plagioclase feldspar component of the granite, while K–feldspar remains
relatively stable except in zones of extreme
alteration/weathering. This is consistent withthe mobility hierarchy described in [4.11] above,
where Ca and Na are more easily leached than K
during acid hydrolysis. The process is also demonstrated in reaction [4.10] although, as mentioned
previously, the actual mechanism is likely to be
more complex. Many workers have described anintermediate stage of either muscovite or smectite formation prior to the ultimate formation of
kaolinite (Exley, 1976; Durrance and Bristow,1986). This, too, is consistent with the pattern ofprogressive clay mineral formation illustrated as afunction of rainfall in Figure 4.8. It seems likely,
therefore, that acid hydrolysis type reactions gave
rise to early alteration of plagioclase feldspars
to form sericite (fine grained muscovite), as wellas incipient argillic alteration. Exhumation of
the batholith during Mesozoic times, together
with further meteoric fluid circulation and deep
weathering, continued the alteration process toform zones of intense kaolinization, particularly
in areas where previous alteration had takenplace. It appears likely that kaolinization was also
accompanied by leaching of Fe from the altered
granite, a process that contributed to the very
high purity and “whiteness” of the Cornish china
clay.

4.5 CALCRETE-HOSTED DEPOSITS
Most of the laterally extensive surface or nearsurface calcrete (or caliche) layers that typify
arid environments around the world are referred
to as “pedogenic calcrete” because they represent
calcified soils (Klappa, 1983). Calcrete is defined
simply as an accumulation of fine grained calcite(CaCO3) in the vadose zone (i.e. above the watertable) during a combination of pedogenic (soilforming) and diagenetic (lithification) processes.
The solution and precipitation of calcite in the236PART 3 SEDIMENTARY/SURFICIAL PROCESSES

Process

Phreatic

VadoseH O
H2O
liquid 2
CO2
CO2
Soil moisturezone
Gravitational
water zone
CapillaryfringeCalcrete classification
Non-pedogenic
superficial
calcrete

vapor

Calcretes

Water
table

Pedogenic calcrete
Soilforming
processes
Descendingwater
Capillarytransport

Predominant transport
in solution
Surficial transport(Essentially no uranium)
Verticalredistribution(Minor uranium)

Non-pedogenic
groundwater
calcrete
Valley (channel) calcrete
Deltaic calcrete
Lake margin calcrete

Lateral transport(Largest uranium
potential)

Groundwater

Figure 4.11 Simplified classification scheme for calcretes and a summary of the processes by which they form. An
indication of the suitability of calcretes for hosting surficial uranium deposits is also provided (after Carlisle, 1983).

surficial environment is represented by the following general equation (after Klappa, 1983):
Ca2+ + 2HCO3− ⇔ CaCO3 + CO2 + H2O[4.12]The solubility of CaCO3 increases with decreasing temperature, with lower pH, and with increasing partial pressure of CO2. Thus, extraction of
CO2 from an aqueous solution will precipitate
CaCO3. Precipitation of calcite may also result
when water is removed from a soil profile during
evaporation and evapotranspiration, resulting in
an increase in the concentration of aqueous ionsin the solution. A simplified classification of
calcrete types is shown in Figure 4.11 where it isevident that, in addition to the very extensive
pedogenic calcretes, a more locally distributed
non-pedogenic variant, termed valley or channel
calcrete, also occurs. It is the latter variety thatis particularly important as the host rock forsurficial uranium deposits.
4.5.1 Calcrete-hosted or surficial
uranium deposits
Important uranium resources have been discovered in channelized calcretes from arid regions inAustralia and Namibia. The Yeelirrie (Western
Australia) and Langer Heinrich (Namibia) deposits
represent well known examples of surficial uranium ores formed by accumulations of a brightyellow potassium–uranium vanadate mineral,
carnotite (K2(UO2)2(V2O8).3H2O), within calcretized
fluvial drainage channels (Carlisle, 1983). Since
carnotite is a uranyl (U6+)–vanadate, reduction ofhexavalent uranium is not the main process
involved in the precipitation and concentration of
uranium ores (see Chapter 3). Rather, the formation of this type of uranium deposit is related tohigh rates of groundwater evaporation and theresultant decrease of aqueous carbonate, vanadium, and uranium solubilities within a few
meters of the surface. Calcretized fluvial channels
represent the remnants of rivers from a previous
higher rainfall interval. Such channels occasionally drained a uranium-fertile source region and,
where preserved, represent zones of focusedgroundwater flow within which mineral precipitation resulted in concentration of uranium ore.
In this process uranium and other components
are required to remain in solution until they reacha zone where carnotite can be precipitated byevaporation of the groundwater.
At Yeelirrie in Western Australia, fluvial

SURFICIAL AND SUPERGENE ORE-FORMING PROCESSES CHAPTER 4

237

(a)
Calcrete channel
Archean granite

Watershed

10 km

Yeelirrie
Orebody
Yeelirrie

ChannelLake
Miranda(b)
Alluvial
plainHigh evaporation (CO2 + H2O)
Salinity increase

pH 4.5–7.0

Calcrete channel

Water
table
pH 6.0–7.0

Basement highpH 7.0–8.5

Weathered Archean granite

Groundwater flow

Increasing EhZone of carnotite mineralization

Figure 4.12 (a) Geological setting of the Yeelirrie carnotite deposit hosted in channelized calcrete, Western
Australia (after Mann and Deutscher, 1978; Carlisle, 1983). (b) Model depicting the setting and processes involvedin the formation of carnotite deposits in calcretized channels (after Carlisle, 1983).

channels started to incise the bedrock during
Paleogene times, although the period of aridity
commenced only in the late Pliocene. Active calcrete deposition and carnotite ore formation arethought to have occurred in only the past 500 000
years (Carlisle, 1983). Calcrete crops out along
the axis of a paleochannel that can be traced forover 100 km and within which a carnotite bearing
orebody, some 6 km long, 0.5 km wide, and up to
8 m thick with a resource of about 46 000 tonsof U3O8, is defined (Figure 4.12a). The Langer

Heinrich deposit in Namibia is remarkably similar in most respects to Yeelirrie.
The main ingredients of carnotite, U, K, and V,
are derived locally from weathered granites (forthe uranium and potassium) and possibly more
mafic rocks for the vanadium. In relatively carbonated groundwaters under near neutral conditions, U6+ would be transported as a carbonate
complex (see Figure 3.36), whereas vanadium
was probably in solution as V4+, although theexact species and complexing agent is not known

238

PART 3 SEDIMENTARY/SURFICIAL PROCESSES(Mann and Deutscher, 1978). Focused groundwater flow introduced the ore-forming ingredients
into the channel, with the combination of highevaporation rates and calcite precipitation ensuring that the solution evolved toward higher salinities and pH along the flow path (Figure 4.12b).
In detail, however, the actual precipitation ofcarnotite is a complex process and appears to berelated either to evaporation and decomplexationof uranyl–carbonate complexes, or to oxidation of
V4+ to V5+, or both (Mann and Deutscher, 1978). A
general equation for the precipitation of carnotite
can be written as follows (after Carlisle, 1983):
2UO 22+ + 2H2VO −4 + 2K+ + 3H2O ⇔K2(UO2)2(V2O8).3H2O + 4H+
carnotitein deeper groundwaters flowing below the calcretized channel and that this fluid mixed withthe surficial, oxidized, U- and K-bearing fluid toform carnotite in a mixing zone. The migrationof the lower V-bearing fluid to the near surface
environment is caused by high points in thebasement topography (Figure 4.12b) which force
groundwaters upwards into the evaporative zone
to mix with oxidized fluids, or be directly oxidized by interaction with the atmosphere. This
notion is consistent with the common observation that carnotite mineralization is spatiallyrelated to pinch-outs and constrictions in thefluid flow path.

[4.13]
4.6 SUPERGENE ENRICHMENT OF CU AND OTHER METALSIN NEAR SURFACE DEPOSITS

Evaporation on its own removes CO2 fromsolution and drives equilibrium equations such as[4.12] and [4.13] above to the right, promoting precipitation of calcite and carnotite. Precipitation ofcarnotite, however, produces hydrogen ions (seeequation [4.13]), which lowers pH and increases
the solubility of calcite. Dissolution of calcite
provides additional CO 2−
3 to the solution, which
in turn increases uranium solubility and favors
dissolution of carnotite. The processes describedin terms of equations [4.12] and [4.13] appear,
therefore, to be counter-productive and this isevident because the orebodies themselves provide
textural evidence for repeated dissolution andreprecipitation of ore and gangue minerals. In
reality, however, the very high rates of evaporation increase the Ca2+ and Mg2+ concentrations of
the groundwater to such an extent that calcite(as well as dolomite) is forced to precipitate andcarbonate activities are decreased. Formation ofcarbonate minerals, under conditions in which
they would not otherwise have precipitated, also
promotes the destabilization of uranyl–carbonate
complexes and eventually results in carnotite
formation coeval with calcretization. Mann andDeutscher (1978) have suggested that oxidationof vanadium may also play a role in carnotite
formation, since V4+ is soluble in mildly reducing
waters at near neutral pH, but is precipitated as
V5+ with increasing Eh. They envisage that vanadium might have been transported separately

The processes of weathering can be responsible
for the in situ enrichment of Cu, as well as othermetals such as Zn, Ag, and Au, in many deposits
that occur at or near the surface. The process isgenerally referred to as supergene enrichment
and is a product of oxidation and hydrolysis ofsulfide minerals in the upper portions of weathering profiles. It is an extremely important process
in the formation of low-grade porphyry copper
deposits because the presence of an enriched,easily extractable supergene blanket of secondary
copper ore minerals above the primary or hypogene ore is often the one factor that makes themeconomically viable. The processes involved in
the formation of supergene mineralization are
similar to those already discussed in section 4.3.4
above and applicable to the concentration of gold
nuggets or platinum group elements in lateritic
weathering zones. This section focuses specifically on copper and the formation of supergene
enrichments in porphyry-type deposits, although
the principles also apply to other deposit types, inparticular VMS and SEDEX base metal ores.
4.6.1 Supergene oxidation of copper deposits
Unlike the processes applicable to gold and nickel,
copper enrichments are not specific to lateritic
environments and supergene copper deposits occur
in any surficial environment where oxidized,

SURFICIAL AND SUPERGENE ORE-FORMING PROCESSES CHAPTER 4

Illuvial

Weatheredsurface

Leachedcapping(gossan)
Leachedzone
Oxidized
ores
Water table(redox barrier)Secondarysulfideenrichment

Primary, unalteredhypogene ore

acidic groundwaters are able to destabilize sulfide
minerals and leach copper. The principles involvedin this process are illustrated in Figure 4.13 andalso summarized in terms of the reaction below:
4CuFeS2 + 17O2 + 10H2O ⇒chalcopyrite
4Fe(OH)3 + 4Cu2+ + 8SO42− + 8H+
goethite

Water
seepage[4.14]

In most porphyry copper environments pyrite is
the dominant sulfide mineral and its hydrolysis
and oxidation dictates the production of hydrogenions (i.e. the decrease in pH) in the weatheringzone (see equations [4.7] and [4.8] above). Pyrite
breakdown is also accompanied by the formationof goethite in the regolith and the liberationof SO 2−4 . Chalcopyrite is the major copper–ironsulfide mineral and its breakdown, describedin reaction [4.14], produces soluble cuprous ionsthat are dissolved in groundwater solutions. In
the regolith profile, therefore, oxidizing, acidic
groundwaters leach the primary, hypogene orebody of its metals, leaving behind an eluviatedzone that may, if the intensity of leaching is
not too severe, be residually enriched in iron,

MINERALS

Hydrated ironoxidesCu2+
Cu2+

Supergene enrichment

Figure 4.13 Schematic section
through a copper deposit
showing the typical patternof an upper, oxidized horizon(the leached or eluvial zone)
overlying a more reduced zoneof metal accumulation (thesupergene blanket or illuvial zone).
The uppermost zone of ferruginous
material, often containing theskeletal outlines of original sulfide
minerals, is known as gossan.
The redox barrier may be thewater table or simply a rock buffer(after Webb, 1995).

MineralizedveinReducing oxidizing

Eluvial

WEATHERING
ZONES

239

Coppercarbonates,oxides,silicates, etc.
Chalcocite,
covellite,
borniteChalcopyrite,
pyriteas hematite or goethite/limonite. The upperclay- and Fe–oxyhydroxide-rich capping, which
also contains the skeletal outlines of the original
sulfide minerals, is referred to as a gossan.
Gossans are very useful indicators of the previous
existence of sulfidic ore, not only in porphyrysystems, but in many other ore-forming environments too. The soluble copper ions percolate
downwards in the regolith profile and encounter
progressively more reducing conditions, either asa function of the neutralization of acid solutions
by the host rock, or at the water table. Copper isthen precipitated as various secondary minerals,
the compositions of which reflect the groundwater composition as well as local pH and Eh inthe supergene zone. The so-called “copper-oxide”
minerals that precipitate here are actually compositionally and mineralogically very complex
and can include a variety of copper –carbonate,–silicate, –phosphate, –sulfate, –arsenate, as well as–oxyhydroxide phases (Chávez, 2000). In addition,
Cu also replaces pre-existing sulfide minerals(i.e. pyrite and chalcopyrite) in more reduced
zones where the latter minerals are still stable. Cu
typically replaces the Fe in hypogene mineralssuch as pyrite and chalcopyrite, to form a suite of

240

PART 3 SEDIMENTARY/SURFICIAL PROCESSESLEACHED ZONE
Hematite/Goethite
AluniteCa-Nontronite
Gibbsite
Kaolinite
Muscovite
QuartzSUPERGENE BLANKET
BorniteChalcocite
Covellite
PyriteAluniteCa-Nontronite
Kaolinite
Muscovite
Quartz
PRIMARY HYPOGENE ORE
Chalcopyrite
PyriteAlbiteBiotiteCa-Nontronite
Epidote
K-Feldspar
Muscovite
Na-Nontronite
QuartzStages1

2

3Supergene blanket
formation
5000
Time (years)

Cu-enriched sulfide phases including chalcocite(Cu2S), covellite (CuS), and bornite (Cu5FeS4).
A more detailed illustration of the distributionof stable mineral assemblages in a typical supergene weathering profile over a porphyry copper
deposit is shown in Figure 4.14. This simulation
integrates the effects of time (or more accurately
the evolution of the chemical system in terms of
fluid/rock ratio) in much the same way that alteration processes were described in dynamic terms
as a function of evolving fluid/rock ratios in section 3.6 of Chapter 3. The primary, hypogene ore
is considered to be hosted in a granite and comprises pyrite + chalcopyrite that is weathered by aflux of slightly acidic rainfall. Within about 6000
years (period 1 in Figure 4.14) the primary sulfides
are totally dissolved from the leached zone andcopper is reprecipitated as chalcocite and covel-

10 0004

Figure 4.14 Illustration of thestable mineral assemblages
modeled as a function of time in theweathering zones above a porphyry
copper deposit. The primary,hypogene ore (least altered) is
compared to the overlying
supergene blanket and the oxidized,
leached zone. Note that with time(from stages 1 to 4) the mineral
assemblages evolve as certain
phases are consumed and others
precipitated (after Ague andBrimhall, 1989).

lite by replacement of primary sulfides in the
supergene blanket, as shown in reaction [4.15]:
CuFeS2 + 3Cu2+ ⇒ 2Cu2S + Fe2+
chalcopyrite
chalcocite[4.15]The stable assemblage in the leached (gossanous)
zone during stage 1, and also into stage 2, ishematite/goethite, alunite, and quartz. The dissolution and/or replacement of primary sulfides
is differential and chalcopyrite tends to disappear
from both leached and supergene zones before
pyrite, which is only removed completely fromthe upper zone by stage 2. With time, and as thefluid/rock ratio increases, the leached zone develops gibbsite, muscovite/sericite, and minor clay
minerals (stage 3), eventually evolving into a unit
dominated by sericite and kaolinite (stage 4).

SURFICIAL AND SUPERGENE ORE-FORMING PROCESSES CHAPTER 4

Alunite disappears and SO42− is removed fromthe weathering zone. The pH of groundwaters in
the leached zone starts off low (acidic) during thesulfide dissolution stage and then evolves toslightly more alkaline conditions with time. The
supergene blanket initially develops an assemblage comprising covellite and chalcocite, as
the primary sulfides are replaced, together withquartz and alunite. Once chalcopyrite is consumed from the leached zone, the downward
migrating flux of Cu ions in solution is reduced
and bornite develops later in the paragenetic
sequence (i.e. from stage 2 onwards). Again in thiszone, the clay minerals and muscovite/sericite
are stabilized by progressive hydrolysis of primaryfeldspars and micas and, ultimately, the stablemineral assemblage is not unlike that of theoverlying leached zone except that copper sulfide(bornite) is still stable. In the primary host rockweathering has only limited effects, althoughover time relatively unstable minerals such asplagioclase, biotite, and magnetite will tend todisappear. It is interesting to note that sulfate(SO42−), produced in abundance during the oxidation of primary sulfides (see reaction [4.14] above),is generally transported away from the in situ
weathering zone, and very little sulfate reduction(either organic or inorganic) tends to occur inthe supergene blanket, or below the water table.
This is another reason why low-S copper sulfides
such as covellite and chalcocite form in thisenvironment.
In addition to the formation of high metal/
sulfur copper sulfide minerals at a redox barrierin the supergene blanket, it is apparent that adiverse suite of secondary copper minerals canform at the base of the oxidized zone (Figure 4.13).
This suite includes oxides such as cuprite andtenorite (the former usually associated with nativeCu), carbonates such as malachite and azurite,
sulfates such as brochantite, antlerite, and chalcanthite, chlorides such as atacamite, silicatessuch as chrysocolla, and phosphates such as
libethinite. The precipitation of these oxidized
secondary minerals is generally due to direct precipitation from groundwaters that are saturated
with respect to one or more of the various components that make up this suite of minerals. Figure 4.15 shows an Eh–pH diagram that identifies

241

the stability fields of several of the copper minerals encountered in supergene enrichment zones
above porphyry copper deposits. Under relatively
low Eh conditions, below the water table, the highmetal/sulfur copper sulfides are stable, whereas at
or close to the water table, cuprite and native Cuform. Above the water table the stability fields ofa range of secondary copper minerals (note only a
few are shown) occur and these are controlled
essentially by pH. This is because it is essentiallythe pH that controls the solubilities of complex2−
2−−
ing ligands (i.e. CO 2−
3 , OH , SO 4 , PO 4 , etc.)
and determines which of the relevant copper
complexes are likely to be saturated in any given
environment. The extent to which leached andsupergene zones develop, and the nature of thesecondary minerals that form, depends to a certain extent on the amount of primary, hypogene
sulfides in the host rock and the acidity generated
in the groundwaters as a consequence of theiroxidation/weathering. Weathering of a host rockwith a low sulfide content will typically result inminimal acidity, limited mobility of Fe and othermetals, and formation of secondary copper minerals that are stable in the near neutral to slightly
acidic pH range (Chávez, 2000). Conversely,
weathering of protores with a high sulfide mineralcontent will result in more extreme acidity andthe formation of secondary copper minerals stable
only at lower pH. Local conditions also play animportant role, however, and supergene enrichment in the proximity of a limestone, for example, will result in local groundwaters with ahigh CO 2−
3 content, resulting in the stabilizationof minerals such as malachite or azurite under
neutral to alkaline conditions (Figure 4.15 andequation [4.16] below).
−
2Cu2+ + CO 2−
3 + 2OH ⇔ Cu2(OH)2CO3
malachite[4.16]

Probably the most spectacular examples of supergene enrichment are associated with the giant
porphyry copper deposits of northern Chile, suchas Chuquicamata, El Salvador, and El Abra. In
addition to world-class hypogene orebodies at depth,
these deposits have been subjected to a complex
Paleogene to Neogene geomorphological evolution that resulted in very significant supergene

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

O
H

Chalcanthite
Antlerite
Brochantite
Malachite

2

2O

1.0

Eh (volts)

Cu2 (OH)2 CO3

Cu4SO4 (OH)6

0.5

Cu3SO4 (OH)4

CuSO4 5H2O

Cu2+

Cu

Ch

Co

H
H

2O

llite

(C

ite

(C

rni

te

2 O)

Tenorite
(CuO)
of

wa

ter

(C

tab

le

u

2S

)

Bo

To
p

u

oc

uS

2

te

alc

ve

0.0

pri

10–1
pO2 (atmospheres)

242

)

Na

(C

5

10–40

tiv

uF

eS

4 ),

eC

u

Ch

alc

op

–0.5

yri

te

10–55
(C

uF
eS

2)

25°C
1 atmosphere
–1.0
0

2

4

10–83

6

8

10

pH

enrichment (Sillitoe and McKee, 1996). The supergene and “exotic” ores at Chuquicamata and El
Salvador are described in more detail in Box 4.2.
A note on supergene enrichment of other metals
Zones of supergene enrichment are also foundin the surficial environment above any exposed
metal orebody. In porphyry systems, molybdenum is normally comprehensively removed fromregolith profiles, although it can accumulate
under very oxidizing conditions as ferrimolybdite, and under alkaline conditions as powellite(Ca–molybdate). Relative to Cu, the other basemetals do not form as commonly as secondary
mineral enrichments. Zn is often dispersed in

12

14

Figure 4.15 Eh–pH diagram
showing the stability fields of
selected copper minerals at 25 °C
and 1 atmosphere (after Guilbert
and Park, 1986).

groundwaters, whereas Pb tends to be relatively
immobile. These two metals do, however, form a
range of sulfate and carbonate supergene minerals,
such as anglesite (PbSO4), cerrusite (PbCO3), andsmithsonite (ZnCO3), under certain conditions
With new technologies available in extractive
metallurgy, non-sulfide zinc ores, especially thosecomprising willemite (Zn2SiO4), now represent
an important category of mineralization associated with oxidation of stratiform, sedimenthosted base metal deposits. Examples include the
Skorpion Zn mine in Namibia and Vazante in
Brazil. Silver tends to behave in much the same
way as copper, and acanthite (Ag2S) is readily oxidized in acidic groundwaters, with Ag being reprecipitated as the native metal, or as Ag–halides

SURFICIAL AND SUPERGENE ORE-FORMING PROCESSES CHAPTER 4

243

Supergene processes: supergene and “exotic”
mineralization in the porphyry copper giants ofnorthern Chile
Porphyry coppers are high tonnage–low grade deposits.
The average grade of primary sulfide, or hypogene,mineralization is typically so low that they are only marginally economic. Fortunately, many porphyry copper
deposits are capped by a blanket of enriched ore formedby supergene processes. The supergene blanket, usuallyaccessible during the early stages of mining, contributes
significantly to the overall viability of the mining operation. At Chuquicamata, for example, which is the largest
copper deposit in the world, with a total reserve of some
11.4 billion tons of ore at 0.76 wt% copper, the supergene
blanket comprised a major proportion of the ore bodythat has now largely been mined out. The supergene
blanket is made up of a barren leached zone, an upper
copper “oxide” zone of antlerite, brochantite, atacamite,and chrysocolla, and an underlying copper sulfide zone
made up mainly of chalcocite (Ossandón et al., 2001).
In addition to the in situ supergene ores, some of the
giant porphyry copper deposits of northern Chile, suchas Chuquicamata, El Salvador, and El Abra, also contain“exotic” copper oxide mineralization. This is secondary
ore that has been transported laterally from the leachedportion of the supergene blanket and precipitated in

the drainage network surrounding the deposits. Again
at Chuquicamata, exotic copper oxide mineralization is
located in a gravel filled paleochannel, extending fromthe main pit to the South or Mina Sur deposit, a distanceof about 7 km. Some 300 million tons of exotic ore wasdeposited in these gravels as chrysocolla and copper wad,a Cu-rich, K-bearing Mn oxyhydrate (Mote et al., 2001).
Similar exotic copper mineralization occurs at El Salvador(the Damiana and Quebrada Turquesa deposits) and alsoat El Abra.
Supergene and exotic mineralization is related to theformation of acidic groundwaters that take up Cu into
solution and reprecipitate it elsewhere. Secondary copper
is either reprecipitated in the chalcocite rich supergene
blanket beneath the leached zone (Figure 1a), or transported laterally by groundwaters moving through paleodrainage channels to form the distal exotic ore bodies ofchrysocolla and copper wad (Figure 1b). The high degreeof preservation of secondary copper ores in northern
Chile is related to episodic tectonic uplift in combination
with the pattern of global climatic rainfall and watertable fluctuations (Mote et al., 2001). Accurate dating ofcopper wad and alteration assemblages in supergene(a)Chuquicamata
3600N section
inal topography
Orig
Leachedf

Supergene(oxides)

d

he

ac

Le

e

sic alt

erationChloritic
alteration(propylitic)

Potas

Supergene

Figure 1 (a) Cross section
through the Chuquicamata
mine showing the distributionof the supergene blanket(represented by the leached cap,
the copper “oxide” zone, andsupergene chalcocite ore) in
relation to the primary hypogeneore (after Ossandón et al., 2001).

West fault

Quartz-sericite
alteration(phyllic)(sulfides)

tlin
ou
t
i
P 96
19

f

0

500
m

244

PART 3 SEDIMENTARY/SURFICIAL PROCESSES(b)
W69°37'
El Salvador District

Rio Seco

Quebrada
Turquesa

El Salvador Townsite

Drainagecourses
Open Pit

S26°15'N

Damiana

Exotic deposits in gravels
Gravels
05km

Figure 2 Exotic Cu-Mn oxide mineralization (wad)in gravels exposed in a bench cut through theDamiana ore body, El Salvador (photo courtesy ofGeorge Brimhall).

Figure 1 (cont’d) (b) Simplified
map of the El Salvador district
showing the distribution ofgravel-filled paleochannels
and exotic deposits extending
radially away from the mainpit (after Mote et al., 2001).

mineralized zones of northern Chile indicates that
surficial processes were long-lived and episodic. Some
supergene mineralization is preserved from just a few
million years after the hypogene ores formed at the
Eocene–Oligocene boundary (around 35 Ma). The main
pulses of supergene ore formation, however, occurred
toward the end of the Oligocene (20–25 Ma) and in
the mid-Miocene (12–15 Ma). These pulses coincided
broadly with periods of relatively high rainfall which promoted weathering, leaching of copper, and groundwater
flow, as well as the deposition of gravels in proximal
drainage channels. The subsequent preservation of bothsupergene and exotic styles of mineralization is due to thedrop in erosion rates (i.e. reduction in uplift) and, more
importantly, the onset of hyperaridity and limited groundwater flow, as the Atacama desert formed from the midMiocene onwards (Alpers and Brimhall, 1989).

SURFICIAL AND SUPERGENE ORE-FORMING PROCESSES CHAPTER 4

245(AgCl, AgBr, AgI) in regions characterized by arid
climates. In the Atacama desert of Chile there
have been spectacular concentrations of supergene silver ores discovered, including a 20 ton

aggregate of embolite (Ag(Cl,Br)) and native silver(Guilbert and Park, 1986). The concentration ofgold, with specific reference to lateritic regolith
profiles, is discussed in section 4.3.4 above.

The chemical processes that contribute to weathering include hydration and dissolution, hydrolysis and acid hydrolysis, oxidation and cationexchange. Surficial and supergene ore-forming
processes are related essentially to pedogenesis,
which can be simplified into an upper zone ofeluviation (leaching of labile constituents andresidual concentration of immobile elements) andan underlying zone of illuviation (precipitation oflabile constituents from above). Laterites are theproduct of intense weathering in humid, warm
intertropical regions and are important hosts tobauxite ores, as well as concentrations of metals
such as Ni, Au, and the PGE. Residual concentrations of alumina, and the formation of bauxitic
ores, occur in high rainfall areas where Eh andpH are such that both Si and Fe in laterites are
more soluble than Al. Ni enrichments occur
above ultramafic intrusions in the illuviated laterite zone where the metal is concentrated inphyllosilicate minerals by cation exchange. Au
and Pt enrichments also occur in laterites above
previously mineralized terranes, forming in thepresence of highly oxidized, acidic, and saline

groundwaters. Fixation of the precious metals
occurs by reduction or adsorption, in the presenceof carbonaceous matter or Fe and Mn oxyhydroxides. Oxidized, acidic groundwaters are capable ofleaching metals, not only during laterite formation, but in any environment where such fluids
are present. Supergene enrichment of Cu can bevery important in the surficial environment above
hypogene porphyry styles of mineralization.
Enrichment of copper again occurs in the illuviated zone, although in this case both Cu–sulfide
minerals (in the relatively reduced zone beneaththe water table) and Cu–oxide minerals (above thewater table) form.
The formation of viable clay deposits, such as
kaolinite, is a product of progressive acid hydrolysis, essentially of plagioclase feldspar. Calcrete
is a pedogenic product of high evaporation environments, but can also form in paleo-drainage
channels in now arid climatic regions. In uraniumfertile drainage systems calcrete channels represent zones where concentrations of secondary
uranium minerals precipitate together with calcite
by evaporation-induced processes.

Bland, W. and Rolls, D. (1998) Weathering: An Introduction to the Scientific Principles. London: Arnold,
271 pp.
Guilbert, J.M. and Park, C.F. (1986) The Geology of Ore
Deposits. New York: W.H. Freeman and Co., 985 pp.
(Chapter 17).
Leeder, M. (1999) Sedimentology and Sedimentary
Basins: From Turbulence to Tectonics. Oxford: Blackwell Science, 592 pp. (Chapter 2).

Martini, I.P. and Chesworth, W. (1992) Weathering,
Soils and Paleosols. Developments in Earth Surface
Processes 2. New York: Elsevier, 618 pp.
Williams, P.A. (1990) Oxide Zone Geochemistry. New
York: Ellis Horwood, 286 pp.
Wilson, R.C.L. (1983) Residual Deposits: Surface Related
Weathering Processes and Materials. London: The
Geological Society of London and Blackwell Scientific
Publications, 258 pp.

246

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

Sedimentary ore-forming
processes

SEDIMENTARY BASINS AND THEIR TECTONIC SETTINGS
CLASTIC SEDIMENTATION AND HEAVY MINERALCONCENTRATIONS

sorting mechanisms relevant to placer formation
application of sorting mechanisms to placer
deposits
sediment sorting in beach and aeolian
environments
CHEMICAL SEDIMENTATION AND ORE FORMATION
ironstones and banded iron-formations
bedded manganese deposits
phosphorites
black shales
ocean floor manganese nodules
evaporites
FORMATION OF FOSSIL FUELS – A BRIEF OVERVIEW
oil and gas formation
coalification
oil shales and tar sands
gas hydrates

5.1 INTRODUCTION
Sedimentary rocks host a significant proportion of
the global inventory of mineral deposits and also
contain the world’s fossil fuel resources. Previous
chapters have considered a variety of ore-forming
processes that have resulted in the epigenetic concentration of metals and minerals in sedimentaryrocks. This chapter will concentrate on processes
that are syngenetic with respect to the hostsediments and where the ores are themselves

Box 5.1 Placer processes: the alluvial diamond depositsof the Orange River, southern Africa
Box 5.2 Chemical sedimentation: banded ironformations: the Mount Whaleback iron ore deposit,
Hamersley Province, Western Australia
Box 5.3 Fossil fuels: oil and gas; the Arabian (Persian)Gulf, Middle East

sediments or part of the sedimentary sequence.
Processes to be discussed include the accumulation of heavy, detrital minerals and the formationof placer deposits, the deposition of organic-richblack shales, and the precipitation mechanismsthat give rise to Fe, Mn, and P concentrations inchemical sediments. In addition, there is a brief
overview of the origins of oil and gas deposits,
as well as of coalification processes, because theorganic source material for these deposits reflects
the local depositional environment and ore formation was broadly coeval with sediment accumulation. Although it might be argued that the
migration of oil through sediments involves later
fluid flow through the depository, such processes
are generally early diagenetic in character andtherefore warrant inclusion in the present
chapter. By contrast, syngenetic, clastic sedimenthosted SEDEX Pb–Zn–Ag, sandstone-hosted U
ores, and late diagenetic red-bed Cu–(Ag–Co) typedeposits are discussed in Chapter 3, since they
involve processes that are more hydrothermalin character, or have their metals originatingfrom outside the sedimentary host rocks. TheseSEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

subdivisions are somewhat arbitrary and reflect
more the organizational structure of this bookthan a rigorous genetic classification.
Modern trends in the application of basin analysis techniques to exploration have undoubtedly
been set in the oil industry, where, in recent
decades, major advances have been made inunderstanding the relationships between sedimentology, tectonics, and the maturation oforganic material. The evolution of oil and gas
within this cycle can be effectively assessed interms of organic chemistry and fluid flow patterns
in the rock (Eidel, 1991). This level of understanding requires a high degree of integration ofnumerous earth science disciplines (including
paleontology, stratigraphy, sedimentology, structural geology, plate tectonics, geohydrology, organic
geochemistry, and others), an approach that isincreasingly being applied to all types of exploration in sedimentary basins (Force et al., 1991).
As with the fossil fuels, ore-forming processes
that give rise to placer deposits and metal enrichments associated with chemical sediments anddiagenetic fluid flow are intimately related to theorigin, tectonic setting, and evolution of the sedimentary host rocks. Most major syn-sedimentaryore deposits tend to occur in a limited range ofbasin types (Eidel, 1991). Passive continental
margin chemical sedimentary ores (banded ironformations and ironstones, bedded Mn deposits,
and phosphate ores), as well as shoreline andfluvial placer deposits (gold, cassiterite, diamonds,
and zircon–ilmenite–rutile black sands), reflect
ore-forming processes that prevail in cratonic settings, where features such as protracted stability
and high depositional energies for reworking of
sediment load apply. Sediment-hosted ores inwhich late diagenetic or epigenetic processes are
responsible for metal accumulation (such as
red-bed Cu–(Ag–Co) deposits and some SEDEX
Pb–Zn–Ag ores; see Chapter 3) tend to be associated with rift basins where rates of deposition arehigh and active faulting promotes the circulationof connate and meteoric fluids. It is also noteworthy that many oil and gas fields are also associatedwith rift-related cratonic basins.
The relationships between sedimentation andore formation are multifaceted and only selected247

topics are discussed in this chapter. Additional
information is available from the works listed at
the end of this chapter.

5.2 CLASTIC SEDIMENTATION AND HEAVY MINERALCONCENTRATION – PLACER DEPOSITS

A placer deposit is one in which dense (or“heavy”) detrital minerals are concentrated during sediment deposition. They are an importantclass of deposit type and can contain a wide
variety of minerals and metals, including gold,
uraninite, diamond, cassiterite, ilmenite, rutile,and zircon. Well known examples of placer deposits
include the late Archean Witwatersrand and
Huronian basins (Au and U) in South Africa andCanada respectively, although the origin of theseores is controversial. Geologically more recent
occurrences include the diamond placers of theOrange River system (see Box 5.1) and the western
coastline of southern Africa, the cassiterite (Sn)placers of the west coast of peninsula Malaysia,
and the beach-related “black sand” placers (Ti,Zr, Th) of Western Australia and New South
Wales, South Africa, Florida, and India.
The fluid dynamic processes involved in placer
formation are invariably very complex. It is
exceedingly difficult to predict, for either exploration or mining purposes, where heavy mineral
concentrations occur. This section will only
touch briefly on what has become a highly mathematical subject that has applicability to geomorphology and flood control engineering, in addition
to ore deposit geology. More detailed reviews of
the principles involved in clastic sedimentationand placer formation can be found in Miall (1978),
Slingerland (1984), Slingerland and Smith (1986),
Force (1991), and Pye (1994).
5.2.1 Basic principles
The formation of placer deposits is essentially aprocess of sorting light from heavy minerals
during sedimentation. In nature heavy mineral
concentration occurs at a variety of scales, ranging from regional systems (alluvial fans, beaches,etc.), through intermediate features (the inner

248PART 3 SEDIMENTARY/SURFICIAL PROCESSESRe = ULδf/η

(a)Vortex(b)

Scour

Figure 5.1 Results of a flume experiment simulating(a) fluid flow at the confluence of a tributary with its
trunk river and (b) the resulting distribution of heavy
minerals (after Best and Brayshaw, 1985).

bank of a river bend or a point bar), to small-scalefeatures (bedding laminae or cross-bed foresets).
An experimental simulation of placer processes is
illustrated in Figure 5.1, where a heavy mineral is
fed via a “tributary” into the “trunk river.” The
confluence of two orthogonal flow streams resultsin the creation of a vortex or eddy downstream of
the junction, with the accompanying erosion of
the bed removing light particles and heavy minerals concentrated in the scoured zone.
The transportation and deposition of sedimentin fluvial and related systems is a complex topic(see overviews in Selley, 1988; Friedman et al.,
1992; Pye, 1994; Allen, 1997; Leeder, 1999), thebasic principles of which are essential to theunderstanding of placer processes. To complicate
matters, sediments can also be sorted by theaction of wind, and this process has applicability
to the formation of placers in, for example, beachenvironments (Kocurek, 1996).
One of the parameters used to quantify the conditions of fluid motion is the Reynolds number,which is a dimensionless ratio identifying fluid
flow as either laminar and stable, or turbulent andunstable. The Reynolds equation is expressed as:[5.1]

where Re is the dimensionless Reynolds number;
U is the fluid velocity; L is the length over which
the fluid is flowing; δf is the fluid density; and η isthe fluid (molecular) viscosity.
For low Reynolds numbers flow is laminar andvice versa for turbulent flow (Figure 5.2a). The
behavior of a particle in a river channel will obviously depend to a large extent on the type of flow
it is being subjected to. Flow in natural stream
channels is predominantly turbulent, the detailed
anatomy of which is shown schematically in
Figure 5.2b. Three layers of flow can be identified.
The bottom zone is the non-turbulent viscous
sublayer, which is very thin and and may break
down altogether in cases where the channel floor
is very rough and turbulence is generated by the
upward protrusion of clasts from the bed load.
Above it is the turbulence generation sublayer,
where shear stresses are high and eddies are generated. The remainder of the stream profile is theouter or core layer, which has the highest flow
velocities. A shear stress (τo in Figure 5.2b) is
imposed on the bed load by the moving fluid andis a function of fluid density, the slope of thestream bed, and flow depth. The curved velocity
profile of the channel section is caused by frictional drag of the fluid against the bed.
The types of fluid flow in water (or air) define
the character and efficiency of mass (sediment)
transport. A particle or grain will move through a
fluid as a function of its size, shape, and density, as
well as the velocity and viscosity of the fluid itself.
In water, a particle at any instant will move in oneof three ways: the heaviest particles (boulders,gravel) roll or slide along the channel floor to form
the bedload (or traction carpet); intermediate sizedparticles (sand) effectively bounce along with thecurrent (a process known as saltation); while thefinest or lightest material (silt and clay) will be
carried in suspension by the current (Figure 5.3a).
In air the types of movement are similar, but thelower density and viscosity of air relative to water
dictate that moving particles are smaller, but
their motion is more vigorous (Figure 5.3b).
The type of particles moving in a fluvial
channel by saltation and suspension is partly a

(a)

(b)
Surfa

ce

Chann

el

Me
an
velo
city
prof
ile

Laminar flow

J

τ0

Bed

Turbulent flow

Outer (core) region

Turbulence-generation layer
S
Viscous sublayer

Figure 5.2 (a) Schematic illustration, using streamlines, of the nature of laminar and turbulent fluid flow.
(b) Internal structure of turbulent fluid flow in a natural channel. τo is the boundary shear stress imposed by the fluidon its bed and is a function of J (the flow depth) and S (the bed slope), as well as fluid density (after Slingerland andSmith, 1986).

(a) Water
Rolling
Sliding
Suspension

Saltation(b) Air
Descending saltation

Creep

Figure 5.3 Illustration of thedifferent mechanisms of sedimenttransport in water (a) and air (b)(after Allen, 1994).

Suspension

Saltation

250

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

512

GravelSand

5.2.2 Hydraulic sorting mechanisms relevant toplacer formation

Silt

256
EROSION

128

d

32
N
IO
SIT
PO
DE

Current speed (cm s–1)

64

be

ive
es
h
Co

16
8

Non-cohesive bed

TR

AN

SP

OR
TA
TIO

N

4

Slingerland and Smith (1986) divided the mechanisms of sorting into four types, namely:• free or hindered settling of grains;• entrainment of grains from a granular bed load
by flowing water;• shearing of grains in a moving fluidized bed;
• differential transport of grains by flowing water.
Each of these is discussed in more detail below
with respect to their roles in the formation ofplacer deposits.

2

Settling

1

The settling velocity of a perfectly spherical particle in a low Reynolds number (non-turbulent)
fluid can be determined using Stokes’ Law, whichis expressed as follows:

0.5
64

16

4
1 0.25 0.06 0.01 0.004
Particle diameter (mm)Figure 5.4 A Hjulström diagram showing how
sedimentary processes can be assessed in terms ofhydrodynamic (flow velocity) and physical (grain size)
parameters. Critical conditions for deposition are shown,as well as those for erosion and transportation for two
situations in which cohesive and non-cohesive channel
beds apply (after Sundborg, 1956; Friedman et al., 1992).

function of the nature of the fluid flow (as definedby the Reynolds number), whereas bed load
movement is determined by shear stress at theboundary layer (Figure 5.2b) and the characteristics of the particles themselves. Combining fluid
flow (hydrodynamic) and physical mass transportparameters provides a useful semi-quantitative
indication of the processes of sedimentation, atechnique first presented diagrammatically by
Hjulström (Figure 5.4; after Sundborg, 1956). In
this diagram the conditions under which either
erosion, transportation, or deposition will takeplace are shown as a function of flow velocity andgrain size. Deposition occurs either as flow velocity decreases or grain size increases (or both), andthese parameters are very relevant to the formation of placer deposits.

V = gd2(δp − δf)/18η

[5.2]

where V is the particle settling velocity; g is
the acceleration due to gravity; d is the particle
diameter; δp and δf are the densities of particle andfluid respectively; and η is the fluid (molecular)
viscosity.
The relationship indicates that particle settling
velocities in the same fluid medium are proportional to particle diameters (squared) and densities. Figure 5.5 shows that, in terms of Stokes’
Law, different sized grains of quartz, pyrite, andgold may settle at the same velocity, a condition
that is referred to as hydraulic (or settling) equivalence. This suggests that particles of differing
size and density could conceivably reside together in the same sedimentary layer, to form arock such as a conglomerate. Hydraulic equivalence is sometimes mistakenly taken as an
explanation for the concentration of small, heavy
detrital minerals in a coarse sedimentary rock.
In fact the type of settling equivalence predicted
by Stokes’ Law is a poor approximation of thecomplex and dynamic sorting mechanisms taking
place during actual placer formation and otherprocesses are required to explain the situationin nature.

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

251

Hydraulic equivalence
Water δf = 1
QuartzPyriteFigure 5.5 Illustration showing theprinciple of hydraulic equivalence
for particles settling according to
Stokes’ Law. The settling velocitiesof quartz, pyrite, and gold, withradii and densities as shown, are thesame, indicating that they would
settle out of a non-turbulent
column of water into the same
sedimentary layer.

=

d = 0.5
δ = 17d=1
δ=5

d = 16
δ=2
VQ

Gold

VP

=

VG

Velocity α d2. (δp – δf) (for gold)
Velocity α d1/2. (δp – δf) (for quartz and pyrite)

There are several reasons why a simple
Stokesian settling is inadequate as an explanation
for placer forming process. These include the factthat stream flow is generally turbulent (it has ahigh Reynolds number), particles are not spherical, and particle sizes may be either too big or too
small for the relationship in equation [5.2] to beupheld. In addition, if the concentrations of grainsin the fluid is high (>5%) then settling is no longer
unhindered and settling velocities are retarded bygrain–grain collisions and current counterflow.
The random instability of turbulent flow makes
it virtually impossible to predict particle settling
velocities and there is no completely satisfactorymodel for simulating this condition (Slingerland
and Smith, 1986). Likewise, particle shape has animportant effect on settling velocity. A tabular
biotite grain, for example, will settle between 4
and 12 times slower than a quartz grain of equivalent diameter. Big grains have large coefficientsof drag in a fluid and accordingly their settling velocities vary as a function, not of the square of thediameter (equation [5.2]), but of the square root
(Figure 5.5). Smaller grains are, therefore, much
more effectively sorted by settling than are big
grains. Although the ratios of particle diameters(quartz:pyrite:gold approximately 32:2:1) reflecting hydraulic equivalence in Figure 5.5 appear to

be reasonably consistent with what one might
expect in a gold-bearing conglomerate such as in
the Witwatersrand Basin, the diameter of gold itself
in these deposits is typically much too small to be
explainable by Stokesian-type settling. For thatcomponent of Witwatersrand gold that is detrital,
it is, therefore, likely that some other concentration mechanisms, such as entrainment (see below)
might have applied (Frimmel and Minter, 2002).
Particle settling in nature is a difficult parameter to quantify because of the wide rangeof variables likely to affect it. A more realistic
expression of particle settling velocity, whichtakes into account the frictional drag that arises
from different shapes, is provided by the following
expression (after Slingerland and Smith, 1986):
V = [4(δp – δf)gd/3δfCd] /2
1

[5.3]

where V is the particle settling velocity; g is theacceleration due to gravity; d is the particle diameter; δp and δf are the densities of particle andfluid respectively; and Cd is the coefficient of dragand is defined as 24/Reynolds number.
Settling does play a role in the formation ofplacer deposits, but on its own is of little use inunderstanding the processes by which they form.
The existence of a state of hydraulic equivalence252PART 3 SEDIMENTARY/SURFICIAL PROCESSES

will not explain how heavy detrital minerals are
sorted or concentrated in dynamic river or beachsystems. Hydraulic equivalence describes a condition of equal settling velocities and accounts
for unsorted accumulation of heavy minerals
in coarser grained sediment (conglomerate). By
contrast, if settling velocities are not equal andsettling occurs in a system that is flowing (i.e.
a river), then heavy minerals will be segregatedlaterally downstream as a function of size anddensity (see section on transport sorting below).
In any event it is the movement or flow of thefluid medium that has the dominant role to playin placer formation.
Entrainment
Entrainment sorting refers to the ability of a fluidin contact with bed load particles to dislocate
certain grains from that bed and move them further downstream. As with settling, a considerable
amount of effort has been spent by fluid dynamicists to quantify the criteria for entrainment
in terms of variables such as the hydraulic/flow
conditions, particle size, shape and density. In
the channel cross section shown in Figure 5.2b,
the concept of a fluid force (referred to as τo, theboundary shear stress) acting on the bed load isdemonstrated. It is clear that τo must exceed theforces that keep any given particle in place (i.e.
size, mass, shape, friction) before the particle canstart to move. The critical shear stress required toinitiate movement of any given particle is known
as the Shields entrainment function, or simply
the Shields parameter (θ), a detailed description
and derivation of which is provided in Slingerland
and Smith (1986). The Shields parameter is expressed as:θ = τc /[(δp − δf)gd][5.4]

where θ is the dimensionless Shields parameter; gis the acceleration due to gravity; d is the particle
diameter; δp and δf are the densities of particle andfluid respectively; and τc is the critical boundary
shear stress.
In relatively simple situations the Shields criterion can be shown to simulate the movement of

bed load particles in a channel reasonably well,and Figure 5.6a illustrates the critical conditions
that differentiate between entrainment of a particle into the channel and non-movement of thegrain. For a uniform bed load, particle size is thedominant control and sand, for example, will
have a lower entrainment threshold than gravel.
Again, however, a uniform bed load seldom
applies and there are many factors which complicate entrainment sorting. The bed load roughness,or, put another way, the extent to which a particle
sits proud of the bed, will clearly have an effect on
the entrainment threshold. In Figure 5.6a thisvariable is illustrated and quantified by the ratio
p/D. An increase in this value equates with grains
which jut out and this can be seen to have the logical effect of decreasing the entrainment threshold. Other factors which complicate entrainment
sorting include the clast shape (sphericity tends tomake entrainment easier), bed consolidation, andthe range of grain sizes in any given bed load. The
latter point is particularly important in the situation where sediment is made up of a bimodal
assemblage of large and small grains. In this casesmall grains, even if they are relatively light, will
not be entrained because they rapidly become
entrapped within the much larger particles andare no longer available for entrainment.
The effects of entrainment on sorting andplacer accumulation are shown in Figure 5.6b. In
this diagram (similar to the Shields diagram butwhere θ is replaced by τc and Re* is replaced bygrain diameter) the critical boundary shear stressfor entrainment increases as a function of grain
density. This indicates that, for a uniformly sized
bed and with all other factors being equal, lighter
particles will be effectively entrained at lower
shear stresses and, therefore, winnowed away,
leaving a residual concentration of heavier particles. It is, therefore, quite feasible to entrain
quartz from the bed load of a stream and leavebehind a residual accumulation of heavy minerals
that could be preserved as a placer deposit. Thin
laminae of heavy minerals in a fluvial setting, oralong cross-bed foresets, are likely to be the resultof this type of sorting process. Entrainment processes also apply to wind-blown sedimentationand sorting, but the magnitude of the critical

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

253

(a)
100
p

EntrainmentDθ 10–1p/D = 0
0.2
0.4
0.6

Non-movement
10–2

100

101

102

103

104

Re*
(b)

(δ

=

G

5)

ol
d

(δ

=

19
)

103

2.

en

5)

ite

102

Q

ua

rtz

(δ

=

Ilm

Figure 5.6 (a) Shields diagram
showing the threshold conditions
between entrainment and nonmovement of a grain for a bed
load of uniform size in terms of
the Shields parameter (definedin equation [5.6]) and the grain
Reynolds number (Re* or the
Reynolds number as it pertains
to a single particle). The effects
of increasing the protrusion of aparticle above the bed floor (i.e.
increasing the ratio p/D) will
have the effect of shifting theentrainment threshold: the higher
a particle protrudes above the bed,
the easier it will be to entrain.
(b) Diagram showing the effects
of grain density (δ) on the criticalboundary shear stress (τc). Greater
shear stresses are required toentrain denser particles. Details
after Slingerland and Smith (1986)
and Reid and Frostick (1994).

τc 101

100

10–1
10–3

thresholds will be quite different because of thequite different mechanical conditions that pertain to the aeolian environment.
Shear sorting
Shear sorting of grains is a process that only
applies to the concentrated flow of suspended
particles in a fluidized bed. During the movementof suspended particles in a dense granular dispersion, grain collisions create a net force thatis perpendicular to the plane of shearing and

10–2

10–1

100

101

d (cm)

disperses the granules toward the free surface(i.e. upwards). Counter-intuitively, the dispersive
pressure is greater on large and dense grains
within the same horizon of flow so that thesegrains migrate upwards relative to smaller andlighter particles. The same effect can be rationalized in terms of a concept known as kinetic
sieving, where smaller grains simply fall between
larger ones (Slingerland, 1984).
The effects of shear sorting have been quantified for sediment populations of mixed sizes anddensities by Sallenger (1979), who showed that

Relative height of granular bed

254

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

0.00

0.00
Coarsequartz

0.25

0.25

0.50

0.50

Zone ofmagnetite
enrichment
Fine
quartz

0.75

1.00
2.0

1.0

0.5
0.25 0.125
Grain size (mm)

0.75

1.00

0
10
20
30
Magnetite concentration (%)

two grains of different densities (δ) coming to restin the same horizon would have relative sizes (d)given by:dh = dl(δl/δh)1/2[5.5]

where dh and dl are the diameters of heavy andlight fractions and δh and δl are the densities ofheavy and light particles respectively.
An illustration of shear sorting in a concentrated granular dispersion comprising quartz andmagnetite, in the proportion 90:10, is shownin Figure 5.7. The horizon at a relative heightof between 0.75 and 0.5 units from the surface
is seen to contain magnetite concentrations up to
double the initial concentration. The mechanism
is, therefore, one of the few processes that will
explain heavy mineral concentrations in an elevated horizon of the sediment strata and could
also explain inversely graded particle concentrations. Shear sorting can apply only to concentrated suspended particle loads such as beachswash zones and wind-related dune formation.
The process could be applicable to the concentration of Ti–Zr–Th black sand placers in theseenvironments.
Transport sorting
Transport sorting is by far the most importantsorting process and the one most applicable to thebroadest range of environments in which placer
deposits form. The quantification of transportsorting is a complex process, but conceptually it

Figure 5.7 Quartz grain size
and magnetite concentrations in ashear sorted granular dispersion
comprising quartz and magnetite
in the proportion 90:10. The
relative depth of the granular bed ison the ordinate; original magnetite
concentration (10%) is shown as
the vertical dashed line (afterSlingerland, 1984).

refers simply to the differential transport rates
that exist during movement of particles in a
flowing fluid medium. It is complex because it
incorporates two distinct components, namely
the varying rates of movement of grains both in
the bed load (determined by entrainment), and insuspension (determined by settling).
The previous discussion of settling considered
the concept only in terms of unhindered, nonturbulent flow and it was emphasized that it
had limited use in understanding the dynamics
of placer-forming processes. The concept of suspension sorting, however, is an extension to therather simplistic considerations of the previous
section and refers to the fractionation of grains
with different settling velocities into different
levels above the bed in a turbulent channel
flow system. Subsequent to its deposition downstream, sediment sorted in this fashion can resultin substantial heavy mineral enrichments. It is
discussed here, rather than earlier, because it is
one of the two components that contribute to thedetermination of transport sorting.
The concentrations of heavy (h) and light (l)particles that coexist at any point in the channel
flow system can be quantified in terms of an equation derived in Slingerland (1984):(C/Ca)h = (C/Ca)Vl h /Vl[5.6]

where C is the concentration at a given level inthe channel flow; Ca is a reference concentration;
and Vh and Vl are the settling velocities of heavy
and light particles respectively.

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

The above equation predicts that particles in aturbulent flowing channel will be sorted vertically according to their settling velocities which,in turn according to Stokes’ Law, are determinedby their relative sizes and densities. The concept is pertinent to transport sorting because theeffects of suspension sorting are only relevant
downstream once deposition of a particular sediment horizon, with its possible enrichment of
a hydraulically equivalent mineral, has takenplace. Note that the results of suspension sortingare contrary to those obtained by shear sorting ina concentrated fluidized bed.
The other component of transport sorting
applies to the movement of bed load and is discussed previously in the section on entrainment.
The entrainment threshold of a particle sitting on
the bed is determined by the Shields parameter(equation [5.4]) but is also strongly influenced by
bed roughness. Consequently sediment transport
rates would be expected to decrease as bed roughness increases. Slingerland (1984) has computermodeled transport sorting processes using an initial sediment comprising quartz and magnetite,in the proportion 90:10, and mean grain diametersof about 0.4 and 0.2 mm respectively. The results
confirm that for a given shear velocity (anothermeasure of the boundary shear stress, as shownin Figure 5.2b) all sizes of quartz and magnetite
exhibit a decrease in transport rate as bed roughness increases. In addition, and as shown in Figure
5.8a, the proportion of magnetite in the moving
sediment load increases as a direct function ofshear velocity for a given bed roughness (trenda–a′) but decreases as a function of roughness for agiven shear velocity (trend b–b′).
The effects of transport sorting have also been
demonstrated experimentally in a flume, wheresediment transport is simulated using simplified,scaled down parameters. Figure 5.8b shows thata positive relationship exists between transport
velocity and shear velocity for any particle sizeand bed roughness. However, for any measure of
bed roughness (K) larger grains move faster thansmaller ones, a feature that is explained by thefact that smaller grains progress less easily over arough bed because of trapping and shielding.
Thus, for any given shear velocity and bed rough-255

ness, the largest grains have the fastest transport
rates (trend c–c′–c″ in Figure 5.8b). Conversely,
the slowest transport rates for a given shear velocity are a feature of the smallest particles (i.e.
0.21 mm in this particular experiment) moving
over the roughest bed (K = 0.78 in Figure 5.8b).
Clearly, transport sorting is a complex process
which can have quite different outcomes fromone part of a channel flow system to another. A
bed of sediment that is being deposited at any
given moment may undergo enrichment in heavy
minerals if the combination of shear velocity
and the sizes of heavy and light particles relative
to bed roughness mitigate against entrainment,
thereby resulting in the formation of a lag deposit.
If, on the other hand, the same particles are subsequently entrained into the sediment load undera different flow regime, they might then be subjected to suspension sorting which would resultin heavy mineral enrichment at a completely
different downstream location.
5.2.3 Application of sorting principles toplacer deposits
Enrichment of heavy minerals by grain-sorting
mechanisms occurs on all scales in nature, fromsingle grain laminae on cross-bed foresets, to large
regional scale concentrations that have accumulated in a particular sedimentary environment
such as an alluvial fan or beach system. It is
tempting to suggest that small scale systems
might be more easily explained in terms of only
one of the mechanisms discussed above, and that
regional systems were more complex, involving
several mechanisms acting in concert. This rationale does not hold, however, and even small scale
systems can be a product of complex interactions;the following examples demonstrate the application of sorting mechanisms to different scales ofdeposition.
Small scale
As an example of heavy mineral concentration ata small scale, the mechanisms of grain sorting inand around a dune or ripple migrating along the
bed of a stream are shown in Figure 5.9a. The dune

256PART 3 SEDIMENTARY/SURFICIAL PROCESSES

(a)

its

os

8

g

La

p
de

6

4

Ap

l
ate
m
i
x
p ro

im

it o

n
ag
fm

ite
et

ns
tra

rt

1

b’

2
3

a

2

po

a’

4
5
6
7

b
0

10

20

30

40
U* (cm s–1)

Grain diameter
0.78 mm
0.39 mm0.21 mm
K = Bed roughness, mm

c′′

50

60

Magnetite concentration (%)
in sediment load

Roughness (mm)

10

70

(b)

50

Vg (cm s–1)

40

c′

30

K = 0.21

20

K = 0.39c

K = 0.78

102

4
6
U* (cm s–1)

8

10 12 14crest is characterized by high shear velocities andnon-turbulent flow which increases the Shields
factor and promotes entrainment of larger, lighter
grains and residual concentration of the heavier
particles. The dune foreset, on the other hand,is likely to receive heavy mineral concentrations

Figure 5.8 (a) Computer
simulation showing the effects oftransport sorting in terms of bed
roughness (measured by the heightof clasts protruding from the bed inmillimeters) and shear velocity U*.
The plot shows the approximate
limit for transport of magnetiteand the magnetite concentrations(%) in the sediment load. Maximum
enrichment of magnetite on the bed(i.e. the formation of a lag deposit)
would occur above this limit underconditions of low shear velocity
and high bed roughness (afterSlingerland, 1984). (b) Results ofa flume experiment showing theeffects on grain transport velocities(Vg) of shear velocity (U*), bed
roughness, and grain diameter(after Meland and Norman, 1966).

by shear sorting of high concentration grain
avalanches down the slope of the advancing dune.
The trough or scour forming ahead of the dune
is likely to receive heavy grain concentrations
by settling sorting in a locally turbulent microenvironment. It is, therefore, evident that several

(a) Small scale
020
cm
c Settling sorting(settling sorting ofsuspended particles
in turbulent zone)

Stream flow

a
Bed

b

a Entrainment sorting(high shear velocity
and laminar flow)cb Shear sorting (concentrated
granular avalanches downslope)(b) Intermediate scaleACh

meanderelannDivergent
flow

t

Poin

bar

0

30
metersA’
A

A’Convergent
flow
Older
alluvium(c) Large scaleN

Welkom gold
field

e
Steyn
B a sa l –Alluvial
fan

ntry
fro

nt

Figure 5.9 Heavy mineralsorting processes at various scales
and in different sedimentary
environments: (a) at a small scalealong laminae associated withdunes and ripples forming along a
stream bed (after Slingerland, 1984);(b) at an intermediate scale in an
aggrading point bar forming along
the convex bank of a meander
channel (after Smith and Beukes,
1983); and (c) at a large scale ona large braided, alluvial fan
complex such as the Welkom
goldfield, Witwatersrand Basin(after Minter, 1978).

0

2
km

Braided channels
High Au/U ratio
Low Au/U ratio

258PART 3 SEDIMENTARY/SURFICIAL PROCESSES

sorting mechanisms apply even to the small scale
heavy mineral concentrations that one often seesin and around dune and ripple features in river- or
beach-related sediments.
Intermediate scale
An example of heavy mineral concentration at
an intermediate scale is provided by the development of point bars along the convex bank of ameandering river channel. Such sites are wellknown to miners dredging river channels foraccumulations of minerals such as cassiterite orgold. Meandering river channels are also prime
targets during the exploitation of placer diamonds
along the Orange River in southern Africa (seeBox 5.1). Figure 5.9b shows the geometry and flow
patterns associated with the aggradation of apoint bar in a meander channel and its migration
toward the opposite concave bank, which is being
subjected to erosion. Heavy mineral concentrations actually form in degraded scours along thebottom of the channel itself. Both components oftransport sorting seem to contribute to theenrichment process, namely settling of grains in
the highly turbulent environment formed by convergence of disparate flow orientations, andentrainment, since shear stresses are above thethresholds for most of the bed load. Transportsorting, therefore, results in placer accumulations
which are then amenable to preservation as thepoint bar sediment migrates over the accumulated heavy minerals. This process has been
applied to the concentration of gold and uraninite
in point bars formed on the fluvial fan deltas of
the Witwatersrand basin (Smith and Minter,1980).

channels on the fan delta complex. Close to theentry point channel fill and longitudinal bars
comprise coarse pebble conglomerate, with maximum clast sizes in the range 20–40 mm in diameter. By contrast, the contained gold particles
are much finer and range between about 0.5 and0.005 mm in size, although accurate size distribution patterns are not known. Several kilometers
downstream the fan delta is built of less deeply
channeled quartz–arenite. The Au/U ratio decreases
downstream and this is a function of both adecrease in gold and an increase in uraninite down
the paleoslope (Minter, 1978). This has been interpreted to reflect transport sorting which results in
net deposition of gold in more proximal locationsof the fan, with the less dense uraninite being
more effectively concentrated in distal parts.
Transport sorting ensures that clast sizes (andalso, therefore, bed roughness), as well as shear
velocity, decrease exponentially down slope.
Uraninite is significantly larger, and also lighter,than gold and will exhibit higher grain transport
velocities, and, therefore, be carried a greater distance down slope. In addition, gold will be transported less easily over the proximal, high bed
roughness portions of the fan delta complex andwill, therefore, be trapped more effectively inthese areas compared to uraninite. Many of thelaterally extensive Witwatersrand placers represent regional, low dip angle, unconformity surfaces where the fan delta complexes are being
degraded very slowly and the fluid regime is consequently very consistent over large areas. This
type of setting represents an excellent site for theformation of very large placer deposits becausesignificant volumes of sediment can be subjected
to a consistent set of transport sorting mechanisms (Slingerland and Smith, 1986).

Large scale
Large scale sorting of heavy minerals over areasof tens of square kilometers has been describedfrom deltaic fan conglomerates of the Welkom
goldfield in the Witwatersrand basin, South
Africa. Figure 5.9c shows the relative distributions of gold and uraninite (as the Au/U ratio) in
the composite Basal–Steyn placer with respect tothe sediment entry point and major fluvial braid

5.2.4 A note concerning sediment sorting inbeach and eolian environments
Much of the previous discussion applies to processes pertaining to unidirectional water flow andis, therefore, mainly relevant to fluvial environments. As mentioned previously, however, many
important placer deposits are associated withsediments deposited in shoreline environments

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

259

Placer processes: the alluvial diamond deposits of theOrange River, southern Africa
Diamond is the hardest known natural substance. Highquality (i.e. based on clarity, colour and a lack of flaws)
diamonds are used for jewelry, whereas flawed and poorly
crystalline stones are used as abrasives in a wide rangeof cutting tools. The specific gravity (3.5) and hardness of
diamond ensure that it is concentrated together with otherheavy minerals in both fluvial and marine placer deposits.
Over much of recorded history India was the only country
to produce diamonds from alluvial sources, but in the1860s diamonds were discovered in South Africa, specifically in gravels of the Orange River and its tributaries.

This led to the discovery of the primary kimberlitic sourceof diamonds around the town of Kimberley, and then in1908 of the huge beach placers along the west coast ofsouthern Africa.
The alluvial and beach diamond placers of South
Africa and Namibia are the product of deep erosion ofkimberlites on the Kalahari Craton in Cenozoic times.
The erosion of diamondiferous kimberlite liberated the
diamonds onto the land surface for subsequent redistribution into the very sizable catchment of the Orange
River drainage system and its precursors. This drainage
Zimbabwe

N

Kalahari Craton

bique

Ha
Ri rtz
ve
r
fsh
Of

l

a
Va

ore

Ora

Mozam

n

ri Crato

Namibia

KalahaBotswana

r

mo

dia

nge Rive

Johannesburg
Swaziland

r

ve
Ri

Kimberley

nd

O

tho

so

ra

Le

co

ng

Durban
r

sio

ve

es

Ri

nc

e

ns

South Africa

Atlantic
Ocean

Indian
Ocean

Paleo-drainage
channelsCape Town

0

400
km

Port Elizabeth

Diamondiferous kimberlite
Alluvial diamond deposits
Beach-related
diamond placers

Figure 1 Map showing the Orange River drainage system in relation to the distribution of deeply eroded
kimberlites on the Kalahari Craton. The distribution of alluvial diamond deposits along major river channels
and paleo-channels, as well as the beach placers along the west coast of southern Africa, are also shown(after Lynn et al., 1998).

260

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

V

III
III

Or
an
g
V

V

eR

iver

III

IV

Gravel terraces
I
Upper terrace
II
III
IV
V

V

Intermediate terraces
Lower terrace

Exposed basement
I

I

0

1000
meters

2000

I

V
I
III

NamibiaSouth
Africa

V

e

ng

ra

O

Ri

ve

I

N

r

I

V

ib
Tr
ry

a
ut

I

flows westwards and exits into the southern Atlantic at
Alexander Bay–Oranjemund at the South Africa–Namibia
border. Diamonds are trapped in gravel terraces that
represent preserved sections of river sediment, abandonedby the present river as it migrates laterally or incises downwards (Lynn et al., 1998). Diamonds are also washed out
into the ocean and redistributed by long-shore currents to

Figure 2 Map showing the
distribution of gravel terracesin the vicinity of the Bloeddrift
diamond mine along the lower
reaches of the Orange River(after Van Wyk and Pienaar,1986).

be concentrated in gravels either beneath the wave base,or in remnant beaches reflecting previous fluctuations ofsea level.
In the upper reaches of the Orange River system (which
includes the Vaal River) diamonds are preferentially concentrated in areas where the rivers flow over resistantbedrock, such as the Ventersdorp lavas, where good gullySEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

and pothole trap sites can form. Diamonds are concentrated as residual lag deposits in such traps. Concentrationis enhanced by the extreme hardness of diamond as otherless resistant minerals undergo attrition and are more easily
washed out of irregularities in the bedrock (Lynn et al.,1998). In the lower, more mature reaches of the Orange
River, a number of gravel terraces occur. These containlarge, low grade accumulations of diamonds and were
deposited in lower Miocene to upper Pleistocene times,
occuring at different elevations relative to the present
day channel (Van Wyk and Pienaar, 1986). As many asfive terraces can be identified, each characterized bydiagnostic fossil faunal and floral assemblages.
Diamonds derived from gravels of the lower reachesof the Orange River are 97% gem quality since it is the
unflawed diamonds that typically resist the rigors ofmechanical transport and abrasion. Most of the diamonds
are between 0.85 and 1.30 carats in size, but occasionallyvery large stones (60–100 carats) are also recovered.
The single most important feature in the concentrationof diamonds in the terrace gravels is bedrock irregularity.
The development of large potholes, plunge pools, andbedrock ribbing are all features that impinge on grade
distribution in the gravels. The lowermost sections of thegravel profile are typically the highest grade. Bedrock

where sediment sorting is largely controlled by
the dynamics of waves and by tidal fluctuations.
It is also possible in such environments thatsorting processes could be influenced by wind
action and that the latter might interact withwater-borne sediment dispersal patterns. In fact,
heavy mineral concentrations in some of theenormous beach-related “black sand” (Ti–Zr)
placer deposits of the world, such as Richard’s Bay
in South Africa, are a likely product of both beachand wind-related sorting processes. This section
briefly considers the physical processes involvedin such environments and emphasizes some ofthe differences that exist compared to the fluvial
system.
Beaches
The sorting of sediment in beach environments
occurs at a number of scales, ranging from longshore and orthogonal processes (related to masstransport parallel and at right angles to the shoreline) to swash zone-related features attributed to261

depressions can be extremely productive, with grades ofover 40 carats per 100 tons of ore being reported (Van
Wyk and Pienaar, 1986). Large terraces are generally higher
grade and contain bigger stones in their upstream portions,attesting to efficient trapping and sorting mechanisms
prior to preservation.

Figure 3 Diamondiferous gravels related tothe paleo-Orange River system, depositedunconformably on irregular, potholed bedrock(foreground).

the “to and fro” motion of waves breaking on
the shore. A detailed overview of these various
processes is presented in Hardisty (1994) andAllen (1997). Beach-related placer deposits appear
to be mainly related to processes occurring in theswash zone and it is on this environment that thepresent discussion will focus.
The origin and nature of ocean waves is aquantifiable topic of study and the effects of wave
action on sediment bed forms are well understood. Beaches represent the interplay between
sediment supply, wave energy, and shoreline gradient. Sediment transfer on beaches is influenced
largely by tides and currents, whereas sedimentsorting is dominated by waves and swash processes. Figure 5.10a illustrates the spectrum ofwave-dominated shoreline types, emphasizing
the differences between a low energy, shallow gradient beach comprising mud and silt, and a steep,highly energetic environment in which gravel
beaches are deposited. An intermediate situationreflects the setting in which sand-dominatedbeaches are likely to form.

262

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

Increasing grain size, gradient of shoreline and nearshore wave energy

Increasing abundance of sediment supply(a)Dissipation of wave energy
Absence of wave breaking

Shallow

profileMixed
mud/silt

Dissipation of wave energy
Wave break
TroughBar

Sand

Reflection ofwave energy
Wave break
Steep beach

Gravel6–1

Xs (m), Us and U (m s )

4
2
0
–2
–4
–6
–1

Seaward flow Landward flow(b)

Xs-motion of theswash front
Us-velocity of theswash front
U-water velocity within
swash
0

1

2

3
4
Time (s)5

67

Figure 5.10 (a) Factors controlling the nature of wave
dominated beaches and shorelines (after Reading and
Collinson, 1996). (b) Mathematical model illustrating
the dynamics of the swash zone created for a 0.5 m highbreaker at the shoreline (after Hughes et al., 2000).

Although the type of beach forming is important to whether a placer deposit is likely to develop,
environmental considerations alone are unlikely
to explain the dominant processes in beach placer
formation. It is apparent from the limited amountof research that has been carried out in this fieldthat factors such as source fertility and swash
processes are a more important consideration(Komar and Wang, 1984; Komar, 1989; Hughes
et al., 2000). Source fertility is self-evident anddictates whether a placer deposit is likely to contain ilmenite–zircon or diamond concentrations,or none at all. By contrast, the flow dynamics ofa swash zone are more difficult to evaluate. The
application of these processes to the concentration of heavy minerals on a beach has beenstudied by Hughes et al. (2000). An illustration ofmodeled swash dynamics is presented in Figure
5.10b and shows that motion of the swash front
is symmetric about the surge and backwash, but
the water velocity is asymmetric because local
flow reversal occurs before the front has reached
its final landward advance. The dynamics of theswash zone have been used to evaluate mineralsorting mechanisms in the beach placers of southeast Australia and show that neither settling
nor entrainment is likely to be important in theconcentration of heavy minerals. Settling is discounted because swash zones are typically notdeep enough to allow effective settling of heavy
minerals to occur. Entrainment, likewise, wasineffective as a sorting mechanism because thelarge bed shear stresses prevailing imply that
most, if not all, the minerals would be in motion
with little or no prospect of selective entrainment(Hughes et al., 2000). The dominant process
appears rather to have been shear sorting, where
the dispersive (upward) pressures acting on large,
dense grains moving in a concentrated sheet
flow were larger than those applicable to smaller,
less dense grains. In a study of the heavy mineralplacers of the Oregon, USA, coastline, Komar
and Wang (1984) were also able to discount grain
settling as a sorting mechanism because the settling velocities of light and heavy mineral fractions were found to be similar. In contrast to theAustralian situation, this study found that selective grain entrainment was likely to have been the

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

(b)

(a)

Increasingly
difficultto entrain
Zircon50

Ilmenited = 0.21 mm
δ = 3 g cm–3

30
20

Hornblende

10

GarnetIlmenite

0

0

1000

Augite40

Concentration factor% heavy minerals in sediment

263

10

d = 0.15 mm
δ = 4.7 g cm–3

20
30
40
50
Offshore distance (m)Garnet100
10EpidoteHornblende

1
Quartz

60

0.1
1

2

Hypersthene

AugiteConcentrated
in placerMoves
offshore

3
4
5
6
Threshold stress, τ (dynes cm–2)

7

Figure 5.11 (a) Plot showing the changes in heavy mineral contents along a profile of beach sediment in Oregon as afunction of distance offshore. (b) Plot of concentration factors of the principal minerals in an Oregon beach placer
deposit as a function of calculated threshold stress, illustrating the likelihood of selective entrainment processes
during the formation of the deposit (after Komar and Wang, 1984).

most efficient concentration process and the onethat also accounted for lateral variations observedin the distribution of heavy minerals. Figure 5.11ashows the variable distribution of heavy minerals
as a function of distance offshore, a local characteristic that demands a sorting process capableof laterally sorting the mineral components.
Figure 5.11b illustrates the calculated thresholdstresses for the various minerals examined in
the Oregon beach placers as a function of the measured concentration factors and shows that thedensest and smallest grains were the most effectively concentrated. This may be due largely tothe fact that these grains were the most difficult
to entrain and, consequently, were residuallyconcentrated in the near shore environment. The
fact that entrainment sorting seems to apply to the
Oregon placers, but not those formed along
the Australian coastline, is possibly a function of
the higher wave energy in the latter area and thedevelopment of conditions favoring shear sorting
over entrainment (Hughes et al., 2000). Although
the mechanics of the sorting processes are similar, the detailed patterns of heavy mineral distribution on beaches may be quite different to thosein fluvial systems where transport sorting is thedominant process.

Wind-borne sediment transport
As illustrated in Figure 5.3, the transport of
sediment by wind, although similar in principle
to that by water, differs in detail because of themuch lower viscosity and density of air and thegenerally higher kinetic energies of eolian transport. Grains transported by wind are subjected todifferent entrainment criteria than those moved
subaqueously, and are also subjected to more
dramatic ballistic effects due to their energetic
motion. In equation [5.4], the critical shear stressrequired to initiate entrainment of a given particle is seen to vary linearly as a function of its
density and size. Bagnold (1941) showed thatfor eolian sediment transport the critical shear
stress required to initiate entrainment varied as afunction of the square root of density and size.
This implies that it is easier to move small particles by wind than it is to move them by water,which explains the abundance of widespread,coarse grained lag deposits in desert and beachenvironments. Such deposits form by the deflationof a surface initially comprising multiple sizedparticles, by winnowing away the fine grainedmaterial and leaving behind the gravel residue.
Wind-related sediment transport characteristics

264

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

Sand

Clay

Silt

2.0
Sh

Shear velocity (m s–1)

en

t th

n
sio

inm

Long-termsuspension

en

on

tra

p
sus

Bedload
creep

rm

En

ti
lta
sa

0.5

- te

ied
dif

ort

Mo

1.0

res

hold

No transport
0.1

1

0.1
0.01
Grain diameter (mm)are shown in Figure 5.12 and can be comparedwith the Hjulström diagram of Figure 5.4 whichpertains to subaqueous sediment transport.
The extent to which beach-related placer
deposits are modified by wind-blown sediment
dispersal is a question that has not received muchattention, despite the fact that such interactions
are almost certain to have taken place (Chan and
Kocurek, 1988). At the Richard’s Bay deposits, forexample, a substantial proportion of the Ti–Zr
heavy minerals are extracted from eolian dunes
even though much of the original concentration
took place in the swash zone of the adjacentbeaches.
5.2.5 Numerical simulation of placer processes
Sophisticated computer simulations of complex
natural placer processes have now been developedwith a view to improving the prediction of heavy
mineral distribution patterns in placer deposits.
An example is the MIDAS (Model Investigating
Density And Size sorting) program that predicts
the transport and sorting of heterogeneous size–
density graded sediment under natural conditions(Van Niekerk et al., 1992; Vogel et al., 1992). It
has been applied to the problem of grade control
in the Witwatersrand basin and can predict gold
distribution patterns in the host conglomerates
with remarkable accuracy (Nami and James, 1987).
The conglomerates that host concentrationsof detrital gold in the Witwatersrand basin were
deposited in braided alluvial fan systems, the

0.001

Figure 5.12 Wind blown (eolian)sediment transport characteristics
as a function of shear velocity andgrain size (after Tsoar and Pye,
1987).

main elements of which comprise channel andoverbank flow deposits. The geometry of thisenvironment is shown in Figure 5.13a, which
illustrates that fluid flow can be resolved into two
components, a higher velocity flow contained
within the channel, and a lower velocity flow inthe overbank plain that is transverse to that in thechannel. Hydrodynamic considerations suggestthat the fine grained detrital gold particles in
the Witwatersrand deposits were transported insuspension as the prevailing hydraulic conditions
exceeded those required to entrain and suspend
gold in the channel. The model considers the distribution of gold, therefore, in terms of an interaction between flow in the channel where gold
is being entrained, and flow on the overbank
plain where suspended particles will tend tosettle because of reduced transport capacity. The
MIDAS program integrates the effects of flow
velocity, sediment concentration, and the various
parameters of transport sorting for both the longitudinal flow in the channel and the transverse
flow in the overbank plain. The results of computer simulations are compared to actual gold
grade distributions in two examples from theCarbon Leader Reef (Figure 5.13b) and exhibit aremarkable degree of correlation, even when
the bed morphology is complex. The fact thatthe highest gold grades are often found along theedges of channels has long been known, but has
lacked a feasible explanation. The simulationsuggests that the highest gold grades coincidewith zones of flow interaction and that gold

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5(a)Zones of turbulence

265

Longitudinal flow
direction withsuspension transport

Transverse flow
direction
Water surface

Overbank plain
Channel(b)
Carbon Leader Reef

1000

g t –1

Measured
Simulated

500

Figure 5.13 (a) Geometry andflow characteristics of a braided
fan delta complex used to simulate
gold distribution patterns in
Witwatersrand placer deposits.
(b) Comparisons of computer
simulations and actual gold
distribution patterns in two
different profiles across the
Carbon leader reef. Channel edge
gold concentrations are reasonably
well modeled in terms of transportsorting (after Nami and James,
1987).

g t –1

1000

500

266

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

distribution patterns in the Witwatersrand deposits are a function, at least in part, of transportsorting. It should be emphasized, however, that
this relationship exists despite the fact that it iswell known that most of the presently observedgold in the basin has been remobilized and issecondary in nature. This is one of the reasons
why the origin of Witwatersrand ores is still socontroversial.

5.3 CHEMICAL SEDIMENTATION – BANDED IRONFORMATIONS, PHOSPHORITES, AND EVAPORITESIn contrast to the mechanical processes ofsedimentation discussed above, where clastic
sediment is sorted and deposited by water andwind, chemical sedimentation refers to the precipitation of dissolved components from solution,
essentially out of sea water or brine. A wide
variety of rocks are formed by the compaction
and lithification of chemical precipitates andthese include carbonate sediments (limestone
and dolomite), siliceous sediments (chert) andiron-rich sediments (ironstones and banded ironformations), as well as less voluminous accumulations of manganese oxides, phosphates, andbarite. The majority of the world’s Fe, Mn, andphosphate resources, all extremely importantand strategic commodities, are the products ofchemical sedimentation and are hosted in chemical sediments. In addition, rocks such as limestone, comprising essentially CaCO3, have greatvalue as the primary raw material for the manufacture of cement. Furthermore, sediments known
as evaporites, in which chemical precipitation
is promoted by evaporation, contain the maineconomically viable concentrations of elements
such as K, Na, Ca, Mg, Li, I, Br, and Cl, as well ascompounds such as borates, nitrates, and sulfates,
all of which are widely used in the chemical andagricultural industries.
Most chemical sediments form in marine or
marginal marine environments. The continental
shelves, together with intratidal and lagoonalsettings, represent the geological settings where
chemical sediments and associated deposits are
generally located. The chemical processes bywhich ore concentrations form are complex
and controlled by parameters such as oxidation–
reduction and pH, as well as climate, paleolatitude, and biological–atmospheric evolution. Thissection will first consider the processes associatedwith the formation of ironstones, banded ironformations, and bedded manganese oxide ores.
This is followed by a discussion of phosphorites
and also the formation of evaporites, carbonaceous “black shales,” and manganese nodules.
5.3.1 Ironstones and banded iron-formations
In terms of morphology, texture and mineralogy,
there are three main types of iron ores and it
is instructive to consider each of these individually since they illustrate different ore-forming
process. The three types, in order of increasing
importance, are bog iron deposits, ironstone
deposits, and banded iron-formations, the latter
generally referred to by the acronym BIF.
Bog iron ores
Bog ores form principally in lakes and swampsof the glaciated tundra regions of the northernhemisphere, such as northern Canada andScandanavia. The deposits are typically small
and thin, and comprise concentrations of goethiteand limonite (Fe–oxyhydroxides) associated withorganic-rich shale. They formed in the recent
geological past, and in some places are still doing
so at present (Stanton, 1972). The principal concentration mechanism for iron in bog ores, aswell as other iron ore deposits, is related to thefact that the metal occurs in two valence states,
namely Fe2+ (the ferrous ion) which is generally
soluble in surface waters, and Fe3+ (the ferric ion)
which is insoluble and precipitates out of surface
solutions. Concentration of iron occurs when
aqueous solutions containing labile Fe2+, in a
relatively reduced environment, are oxidized,with the subsequent formation and precipitationof Fe3+. Figure 5.14 illustrates the principle withrespect to bog iron ores accumulating at the interface between oxygenated surface waters flowing
along an aquifer and the reduced iron-rich solutions percolating downwards through a swamp.

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

267

Aquifer withoxygenatedgroundwater
Swamp
Fe2+

Fe2+

Fe2+

Reduced carbonaceousshale
Bog iron ore
at redox interface
Fe3+ accumulation(as limonite Fe (OH)3)

Figure 5.14 The development of limonitic concentrations in the formation of a bog iron ore where a reduced
solution transporting ferrous iron interacts with oxidized groundwater flowing along an aquifer (after Stanton, 1972).

At this interface ferrous iron in solution is oxidized to ferric iron which precipitates as limonite
or goethite.
Ironstone deposits
Ironstones are typically Phanerozoic in age andare widespread in occurrence, representing animportant source of iron in the eastern USA
and western Europe, particularly in the early partof the twentieth century. Ironstone deposits arecommonly referred to as minette or Lorraine-typeiron ores, well known occurrences of which were
found in the Jurassic sediments of Englandand the Alsace-Lorraine region of France andGermany. In North America the Silurian Clinton
type iron ores of Kentucky and Alabama are
analogues of the younger European deposits.
Ironstones form in shallow marine and deltaic
environments and typically consist of goethiteand hematite that has been rolled into oolites
or pellets, suggesting the action of mechanicalabrasion. The deposits contain little or no chert,
but are commonly associated with Fe-rich silicate
minerals such as glauconite and chamosite. The
environment of deposition suggests that the iron
was introduced from a continental source via a
fluvial system in which the metal was either insolution as Fe2+ or transported as a colloid (seeChapter 3). Ironstones are often linked stratigraphically to organic-rich black shales and, incertain cases, seem to develop in strata representing periods of major marine transgression andcontinental margin flooding. They do not occur
randomly in time and show distinct peaks in theOrdovician–Silurian and again in the Jurassic
periods. Their formation appears, therefore, tobe related to a pattern of global tectonic cyclicity
and specifically to times of continental dispersal
and sea-level highstand (see Chapter 6), as wellas periods of warmer climate and increased rates
of chemical sedimentation (Van Houten andAuthur, 1989).
The origin of ironstones is complex and controversial and needs to account for both the Feconcentration processes and the formation of theubiquitous oolites that typify these ores (Youngand Taylor, 1989). A model for the formation ofminette-type ironstones, after Siehl and Thein(1989), is presented in Figure 5.15. Iron is thought
to have been concentrated initially on continentsthat were being subjected to deep weathering anderosion in a warm, humid climatic regime. Highly
oxidized lateritic soils forming in such an environment would have been the sites of iron enrichment as insoluble Fe3+ remained in situ whileother components of the regolith were leachedaway (see Chapter 4). In addition, laterites werealso the sites where iron ooids formed in response
to low temperature chemical and biogenicallymediated processes. The lateritic soils and ooids
were then transferred into a shallow marine

268PART 3 SEDIMENTARY/SURFICIAL PROCESSES

Sea-level highstandTransgression

lta

Laterite

De

Black shale.

.

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

.
. ... .
...

.. .

River

. ..

Reworked ironstone(mechanically reworked and sorted)

Fe concentration(pedogenic ooid
formation)

Figure 5.15 Simplified environmental model for the formation of oolitic ironstone ores. The model invokes initial
Fe enrichment and pedogenic ooid formation in lateritic soils on the continental edge and transfer of this “protore”
to a marginal marine setting with subsequent mechanical abrasion and reworking/concentration of ooids (after Siehl
and Thein, 1989).

environment, either by flooding during transgression or by erosion during regression, to be
reworked and concentrated in fluvio-deltaic or
littoral settings. Note that the chemically orbiogenically formed pedogenic ooids are different
to those that arise from the mechanical abrasionof particles in the swash zones of shallow marine environments. Subsequent diagenesis of theironstone accumulations resulted in further
post-depositional modifications to the texture
and mineralogy of the ores. Unlike the relatively
simple processes invoked to explain the formation of bog ores, ironstones appear to require a
combination of specific environmental conditions, as well as a variety of processes (including
oxidation–reduction, diagenesis, mechanical sedimentation, and microbial activity), to form substantial deposits.
Banded iron-formations (BIFs)
The term banded iron-formation is somewhat
controversial because of the connotation it haswith respect to stratigraphic terminology. For
this reason the term is hyphenated, and applies
strictly to a bedded chemical sediment comprising alternating layers rich in iron and chert (Klein
and Beukes, 1993). BIFs represent the most important global source of iron ore and far outweigh theironstone and bog iron ores in terms of reserves
and total production. Whereas the latter occur
predominantly in Phanerozoic rock sequences,
BIFs are much older and formed in essentiallythree periods of Archean and Proterozoic Earthhistory, namely 3500–3000 Ma, 2500–2000 Ma,
and 1000–500 Ma (Figure 5.16). These three classes
equate broadly with different tectonic settings
and are referred to as, respectively, Algoma, Lake
Superior, and Rapitan types. Algoma type BIFs areassociated with volcanic arcs and are typicallyfound in Archean greenstone belts. These deposits
tend to be fairly small but they are mined inplaces such as the Abitibi greenstone belt of
Ontario, Canada. The majority of Lake Superior,or simply Superior, type BIFs are located on stable
continental platforms and were mainly depositedin Paleoproterozoic times. They represent by far
the most important category of iron ore deposits(Figure 5.16) and most of the major currently producing iron ore districts of the world fall into thiscategory. Examples include the Hamersley Basinof Western Australia (see Box 5.2), the Transvaal
Basin of South Africa, the “Quadrilatero Ferrifero”
of Brazil, the Labrador trough of Canada, theKrivoy Rog–Kursk deposits of the Ukraine, theSinghbhum region of India, and the type area inthe Lake Superior region of the USA. Finally, theRapitan type iron ores represent a rather unusualSEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

7.0

269

Superior

Tons (×1014) of BIF

6.0
5.0
4.0
3.0
2.0
1.0

AlgomaRapitan
0.0
4.0 3.5 3.0 2.5 2.0 1.5 1.0 0.5 0.0
Time before present (Ga)eet

Ice shSea level
Continental
crust
Magmaticarc“Blacksmoker”

Continental
platform

Oceanic crust
Figure 5.16 Tectonic and environmental model showing the depositional settings for Algoma, Superior, andRapitan type BIFs (after Clemmey, 1985; Maynard, 1991). The inset histogram illustrates the approximate tonnages
of BIF resource for each of the three major types as a function of time (after Holland, 1984).

occurrence of iron ores associated with glaciogenic sediments formed during the major Neoproterozoic ice ages. The type occurrence is the
Rapitan Group in the McKenzie Mountains ofnorthwest Canada.
In addition to the above tectonic classification,
BIFs have also been categorized in terms ofthe mineralogy of the associated iron phases.
Although in most BIFs the iron mineral is anoxide phase (hematite or magnetite), carbonate(siderite), silicate (greenalite and minnesotaite),and sulfide (pyrite) iron minerals also occur,
together with chert or carbonaceous shale. Thisobservation led James (1954) to suggest a facies
concept for BIF formation whereby the progression from oxide through carbonate to sulfide

phases was considered to reflect precipitation of
the relevant iron minerals in successively more
reducing environments (iron silicate phases arestable over a wide range of Eh and do not, therefore, conform to this simple progression). A consideration of the stability of the main iron phases
as a function of Eh and pH confirms that the redox
state of the depositional environment plays animportant role in determining the mineralogy of
BIFs, although the situation is undoubtedly more
complex than the facies concept would suggest.
Figure 5.17a shows that ferrous iron is stable insolutions that are acidic and reducing, but that fora given pH, oxidation (or an increase in the Eh)will stabilize hematite (Fe2O3 or ferric oxide),which is the principal iron phase over a wide270

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

(a)

(b)
1.0

Fe3+

1.0

Up

pe

0.8

r li

t fo

rw

rs

0.6

2+

Fe

0.6

He
m
Fe atit
2O e

0.4

Lo
wa werter limsta it fo
bil r
ity

–0.4

Siderite
FeCO3
0

2

Mn2+

4

0.2

Ma
n
Mn gani
2 O te3

Rhodochrosite
MnCO3

0.0
Lo

we

r li

mi

t fo

6

8

rw

ate

–0.4
Ma
g
Fe neti
3 O te

rw

2

–0.2

–0.6
–0.8

Eh (volts)

Pyrite
FeS2

0.0

t fo

ate
rs
tab
Py
rol
ility
Mn usit
O e

3

0.2

mi

ite
ann
usm
Ha Mn 3O 4

Eh (volts)

tab

ility

0.4

r li

0.8
ate

–0.2

Up

pe

mi

rs

tab

ility

–0.6

4

10

12

14

pH

–0.8

0

2

4

6

8

10

12

14

pH

Figure 5.17 (a) Eh–pH diagram showing the stabilities of common iron minerals. The conditions that apply to thisparticular phase diagram are: T = 25 °C, Ptotal = 1 bar, molarities of Fe, S, and CO3 are, respectively, 10 −6, 10 −6, and 1.
(b) Eh–pH diagram showing the stabilities of common manganese minerals. Identical conditions apply, but with themolarity of Mn = 10 −6 (diagrams modified after Garrels and Christ, 1965; Krauskopf and Bird, 1995). Note that the
manganese oxides (MnO2 and Mn2O3) are stable at higher Eh than the equivalent ferric oxide (hematite), and would
only form, therefore, under more oxidizing conditions.

range of geologically pertinent conditions. Sideriteand pyrite are only stable under reducing conditions although these fields would obviously
expand if the activities of total carbonate orsulphur in solution were increased. Similarly, thephase diagram indicates that magnetite is only
stable under reducing and alkaline conditions, but
this field expands well into the range of neutral
pH if the activities of carbonate and sulphur
are lower than the prevailing conditions for thisdiagram.
Banded iron-formations are chemical sediments
in which the major components, Fe and Si, appear
to have been derived from the ocean itself, rather
than from a continental source, as in the case forironstones. This is evident from the lack of aluminous and silicate mineral particulate matter in

BIFs. Models for the formation of these rocks arecontroversial and hampered by a lack of modern
analogues. Features which need to be explained
include the origin of the Fe and Si, their transportand precipitation mechanisms, the cause of thedelicate silica- and iron-rich banding at various
scales, and their formation during certain time
periods, in particular between around 2500 and2000 Ma.
The source of iron is dominantly from theocean itself, through either direct introductionof Fe2+ from hydrothermal exhalations (blacksmokers; see Chapter 3) on the sea floor or dissolution of oceanic crust. Rare earth element patterns of BIFs suggest that hydrothermal venting
and volcanic activity are likely to have contributed significantly to the source of Fe2+ in Algoma

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

271

Ultraviolet radiation

Reducing Oxidizing

Surface
Photo-oxidation

Photosynthesis
Photic limit 50 m

Si

Si

Fe3+

Si
White

Fe-oxides Si(hydroxides)

Si
Fe2+

C

Si

O2

Redox interface 100 m
Upwelling andhydrothermal
input
Bottom >150 m

O2
Si

Fe-poor chert

Carbonsupply
Si
C
Black

FeCO3
Si

Fe3+

e

erit

e
Hematite Magnetit
ation
Banded iron-form

Sid

Figure 5.18 Model invoking upwelling and oxidation of ferrous iron from an oceanic source to explain thedepositional environment for BIFs. Oxidation of ferrous iron and precipitation of ferric iron compounds occurs at adiffuse redox interface formed by the production of oxygen in the upper water levels, either by photosynthesizingorganisms or by ultraviolet radiation induced photo-oxidation, or both. The lateral zonation of BIF facies (i.e.
siderite–magnetite–hematite) shown here differs from the simple scheme envisaged by James (1954). Diagram
modified after Klein and Beukes (1993).

type BIFs, but that this influence diminishedwith time such that Rapitan ores reflect a direct
ocean water character (Misra, 2000). However, theenormous volumes of iron required to form theLake Superior type BIFs, given that they appear
to be spatially unrelated to volcanic or exhalative
sea-floor activity, remains something of a problem. The oceans themselves are likely to represent
an adequate source of Si since the solubility ofamorphous silica (as Si(OH)4) is relatively high(about 120 mg l−1). It seems likely that theArchean–Paleoproterozoic oceans would havebeen saturated with respect to silica, ensuring a
steady precipitation and accumulation of siliceous
matter on the ocean floor during this early stage ofEarth history. This pattern is likely to havechanged in later geological times when the organisms that use silica to build an exoskeleton (i.e.
siliceous protozoans such as radiolaria) evolved
and the element was extracted from the water column. The modern oceans, for example, typicallycontain <10 mg l−1 dissolved silica and aremarkedly undersaturated (MacKenzie, 1975). The
decrease of available Si in the oceans as a functionof biological evolution contributes to an explana-tion for the changing character of iron ores, fromBIFs to ironstones, with time.
The transport and precipitation mechanisms
involved in BIF genesis have traditionally been
explained in terms of the Fe2+ upwelling model,
analogous to the observations of phosphorus
upwelling onto the continental shelves (see section 5.3.3 below). Figure 5.18 illustrates the mainelements of the model, after Klein and Beukes(1993). Ferrous iron from deep, reduced oceanlevels is introduced to shallower shelf environments by upwelling currents similiar to thoseseen along the western coastlines of continents
such as Africa and the Americas today. These
currents interact with the shallower waters
and Fe2+ is oxidized, with subsequent hydrolysis
and precipitation of ferric hydroxide (Fe(OH)3).
In marine environments where CO3 activities
are sufficiently high, FeCO3 would precipitate.
Oxidation takes place at shallow water levels
where sunlight can penetrate and is due either
to production of oxygen by photosynthesizingorganisms living in the photic zone, or to photons
of ultraviolet and blue light, which induce photooxidation of ferrous to ferric ions (Cairns-Smith,272PART 3 SEDIMENTARY/SURFICIAL PROCESSES

1978). Precipitation of silica is problematical, but
it has been suggested that evaporation from theocean surface could promote local silica oversaturation and thereby promote precipitation asa gel which, on compaction, is transformed tochert. The explanation of banding in BIFs remains
a problem and suggestions range from diurnal
cycles (where photosynthesis and photo-oxidation
shut down at night) for very thin lamellae, toseasonal cycles (where summer seasons promotebiological activity and enhance iron precipitation)
for mesobands. Other mechanisms that may havecontributed to the formation of banding in BIFs
include differential flocculation of Fe and Si colloids as a function of water chemistry, and someform of biological mediation. Studies of biomineralization processes (see Chapter 3) in modern day
environments have shown that the filamentous
bacterium Chloroflexus is capable of binding
either silica or iron, and occasionally both silica
and iron, to the cell walls (Konhauser and Ferris,
1996). This form of induced biomineralizationis considered to apply to the formation of some
Precambrian BIFs, a notion supported by thepresence of spheroidal structures in BIFs interpreted as microfossils (LaBerge, 1973). During thelate Archean and Paleoproterozoic, early marine
planktonic bacteria that photosynthetically produced oxygen may have been responsible for theoxidation of ferrous iron in the ocean waters andprecipitation of ferric hydroxides. At the sametime oversaturation of the oceans with respect tosilica would have led to continuous biologicallymediated precipitation of chert, where cell walls
acted as the templates for nucleation of silica(Konhauser and Ferris, 1996). In this model, thebanding in iron-formations is attributed either to
episodic depletion of ferrous iron or to periodic(seasonal) diminished activity of the oxygen producing planktonic bacteria.
The formation of BIFs in discrete time periods(Figure 5.16) is another feature that lacks detailed
explanation, but is almost certainly related toatmospheric and biological evolution, as well as
to the pattern of global tectonic cycles (see Chapter 6). The enormous Superior type BIFs formedon stable continental shelves during periods ofmarine transgression and sea-level highstand,conditions which reflect tectonic cycles of continental dispersion and active, elevated mid-ocean
ridges. These conditions, coupled with the relevant states of biological and ocean–atmosphere
evolution, were unique to the Paleoproterozoic
and help to explain the predominance of BIFs
at this time. Subsequent to oxyatmoinversion(i.e. the period in the Paleoproterozoic whenoxygen levels in the atmosphere increased substantially; see Chapter 6) the atmosphere andoceans became more oxidized and iron existed
largely as insoluble Fe3+, such that the likelihoodof Lake Superior type BIFs forming diminished.
Rapitan type ores, on the other hand, appear to
be intimately associated with widely distributed
Neoproterozoic glaciogenic sediments and thenear-global glaciations that characterized thistime period (i.e. the so-called “Snowball Earth”
period). The large-scale covering of the oceanswith an ice-cap would have isolated them fromthe atmosphere and reintroduced the sort of
reducing conditions in much of the ocean watercolumn that only existed prior to oxyatmoinversion. It is widely held that the development of the
Rapitan type iron ores is related to this near global
oceanic anoxia, which, in turn, led to enrichmentof the oceans in ferrous iron. Inter- and postglacial melting of the ice-cap resulted in glaciogenic sedimentation, reoxidation of the oceans
and associated precipitation of ferric iron. Such a
scenario has been used, for example, to explain
the formation of the Braemar BIFs and diamictic
ironstones associated with Sturtian aged (circa
750 Ma) glaciogenic sediments in the Adelaide
geosyncline of South Australia (Lottermoser andAshley, 2000).
5.3.2 Bedded manganese deposits
Most of the world’s sedimentary manganesedeposits formed in environments similar to thosein which BIF and ironstone ores also formed.
Superior type BIFs are sometimes closely associated with manganese ores, the prime example
being the world’s largest exploited Mn deposits
in the Kalahari manganese field of South Africa.
In this case, enormous reserves of bedded
manganese oxide ores were formed within the

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

273

Chemical sedimentation: banded iron-formations:
the Mount Whaleback iron ore deposit, Hamersley
Province, Western Australia
The Archean to Paleoproterozoic Hamersley Province ofWestern Australia contains huge volumes of Superior-type
banded iron-formations (BIF) that host some of the world’s
largest iron deposits. Western Australia is the world’s
largest exporter of iron ore, the bulk of which comes fromenriched BIF in the Hamersley Province. Several large
mines, including Mesa J, Tom Price, Paraburdoo, andMount Whaleback, together with smaller operations, combine to produce over 170 million tons of export qualityiron ore annually, much of which goes to Japan, South
Korea, and China (Kneeshaw, 2002).
The Mount Whaleback deposit is the largest single ironore deposit in Australia, with an original ore resourceof some 1.8 billion tons. The deposit is located in thesoutheastern portion of the Hamersley Province in an areathat was episodically deformed, with the most intense
folding attributed to the Ophthalmian orogeny between
2400 and 2200 Ma (Figure 1). The ore is hosted in BIF andferruginous shales of the Dales Gorge Member of the
Brockman Iron Formation. The latter unit is the middle ofthree major iron-formation sequences in the Hamersley
Province, forming between 2540 and 2450 Ma (Barley
et al., 1997). This period of Paleoproterozoic Earth history
was a time when similar chemical sediments were being
deposited in continental margin or platformal settings elsewhere in the world, such as South Africa, Brazil, and theUSA (see section 5.3.1 of this chapter). These countries also
contain large BIF-hosted iron ore deposits that are very
similar to those of the Hamersley Province. The BIF hostrock in all these deposits was originally a bedded chert–
magnetite/hematite chemical sediment that typically contains around 25–30% Fe. In order to make an economically
viable ore deposit the BIF needs to be upgraded so that
the Fe content is at least double this value, a process thatis achieved by either selectively leaching or replacing thesilica in the rock. The processes whereby BIF is enriched toform iron ore are complex and episodic. There is considerable controversy in the literature regarding the nature of theenrichment process and both supergene and hypogene
processes have been advocated.
The conventional model of ore formation at Mount
Whaleback, and indeed most other deposits of the
Hamersley Province, is that BIF was exposed and subjected to supergene alteration after the main period of
deformation and folding of the host rock. This process
involved in situ replacement of chert by goethite andoxidation of magnetite to hematite (a process calledmartitization). The actual processes involved are some-what akin to those taking place during the formation oflaterites (see section 4.3 in Chapter 4), although rather
than simple leaching of silica and residual enrichmentof iron, the chert is largely replaced by Fe-oxyhydroxides.
After supergene alteration the resulting goethite–martite
rock was then buried and metamorphosed to form microcrystalline hematite–martite ore. This model, envisaging
a combination of supergene and metamorphic processes,
was initially attributed to Morris (1985) and is believed
to be consistent with the characteristics of most of theenriched BIF deposits of the region. These characteristics
include the high Fe grades (65% Fe) of the enriched BIF,
the lumpy character of the ore, and low phosphorous contents (around 0.05%), all features that are highly desirable
in the marketing of the ore (Kneeshaw, 2002).
Exposure of the BIF and supergene alteration apparently
took place prior to 2200 Ma (Martin et al., 1998a), whichis the maximum age of deposition of conglomerates adjacent to the Hamersley Province (i.e. the Wyloo Group;
Figure 1) within which detrital fragments of hematite ore
are found. Fluid inclusion and stable isotope data have alsosuggested that the fluids involved with the early phases ofhematite–martite ore formation were hot (up to 400 °C),
saline hydrothermal solutions with characteristics inconsistent with a supergene or near-surface setting. Theseand other observations, such as the mineralogy of theore assemblage at the Tom Price mine, have resulted in a
number of workers suggesting a hypogene origin for theenriched BIF-hosted hematite–martite ores (Barley et al.,
1999; Powell et al., 1999). One suggestion is that hot,oxidizing basinal fluids were focused into low angle thrustfaults during the main stage of deformation and folding(i.e. “orogenic fluids”; see Chapter 3) and that it was these
fluids that were initially responsible for silica dissolution
and hematite recrystallization in the BIF.
One of the features that appears to favor a hypogene
origin for the Hamersley iron ores is that they can extend
to considerable depths (over 500 m) below the surface.
This feature on its own does not, however, militate against
supergene processes. The Morris et al. (1980) model for
Mount Whaleback envisages supergene alteration byelectrochemical processes (Figure 2). The high electrical conductivity of magnetite-rich layers in BIF enables
oxidation potentials to extend to significant depths, with
an electrical circuit being completed by ionic conduction
through groundwaters circulating along suitably positioned
fault systems. A cathode is set up at surface where electrons
are consumed in oxygenated groundwaters. The anode274

PART 3 SEDIMENTARY/SURFICIAL PROCESSESPort
Hedland
Australia
Western
Australia
Perth
Location map

Port Hedland

+
+

+
+

+

+
+

Pilbara province
+

+

+
+
+ +
+
+
+
+
+
+
+ +
+
+
Dampler+++ + Yarrie+
Archeangranite–
+
+
+
+ +
+
+
greenstone terrace
+
+
+
+
+
Roebourne
+
+
+ +
+ +
+
+
+ + + ++ +
+
+
+
+
+
+
+
+ +
+
+
+
+
+
+ +
+
+
+
+ + +
+
+
+
+
+
Marble
Bar +
+
+
+ +
+ + +
+
+
+
+
Fortescue +
+
+
+
Group
+
+
+
+
+
+
+
+ +
+ +
+
+
+
+
+
Mesa J
+ +
+
+
+ +
+
+
+
+
+
+
+
+
+ + +
+
+
+
+
+
+
+
Nullagine ++Hamersley
+
+
+
+ +
+
Group
+
+
+ +
Ha
me
rsle
y
+
+
+
+
Tom Price
+
+
++
+ ++
pro
+
vin
ce
+
Paraburdoo
Mt Whaleback+ +Wyloo Group

+

+
+ +
+ + + +
+ + + +
+ +
+ +
+
+
+

Figure 1 Simplified geology and location of some of the major iron ore deposits in the Hamersley Province,
Western Australia (after Kneeshaw, 2002).

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

exists at depth where ferrous ions are oxidized to ferricions and then precipitated as an iron-oxyhydroxide such asgoethite. This electrochemical process explains alteration
processes at significant depths, but does not preclude theexistence of normal supergene processes, or lateritization,in the near surface environment (Morris et al., 1980).
Finally, it should be noted that the Hamersley Province
is characterized by a significant component (about 35%)of much younger mineralization that clearly is related toCathode275

supergene processes active during a cycle of Paleogene
erosion and weathering. These iron deposits are referred
to as channel and detrital ores and are located in matureriver courses and colluvial fans, both of which are spatiallylinked to bedded ores (Kneeshaw, 2002). Although theorigin of the enriched BIF ores remains controversial,
the region has been subjected to a long and involved geological history and the ore-forming processes are expected
to be complex and episodic.

Water table

f

Gro

und

wat

er

–

e
–

4e +O2+H2O

–

4OH

Fe2+
Fe3++3H2O

e–Fe3++e–Fe(OH)3+3H+

AnodePresent surface
BIF

Anode

ult

Goethite–martite

Fa

Low grade ore

100 m

zo

ne

Hanging wall shale

High grade ore
Footwall shale

100 m
f

Figure 2 Simplified cross section through the Mount Whaleback iron ore deposit showing the extent ofenriched BIF-hosted ore and the postulated nature of iron enrichment by an electrolytically mediated process(after Morris et al., 1980).

Transvaal basin at stratigraphic levels just above
those where the equally substantial BIF-hosted
iron ores occur. This is not unexpected given theremarkably similar chemical properties of Fe
and Mn. The geochemical behavior of manganese,like iron, is controlled by oxidation potential andit exists as Mn2+, which is soluble under reducing
and acidic conditions, as well as Mn3+ and Mn4+,

which are less soluble and stabilize as manganeseoxides under relatively oxidizing and alkalineconditions (Figure 5.17b). Pyrolusite, or MnO2 , isthe dominant oxide phase at high Eh and over arange of pH. Comparison of the manganese with
the iron phase diagram (Figure 5.17a and b) showsthat higher oxidation potentials are required tostabilize pyrolusite than hematite. The fact that

276

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

(a)

(b)
Concentration (ppb)

0

100
300
Oxidized(O2 zone)

Fe

500
Approximate
shelf break

Detritalsourcematerial
Mn 0.1%
Fe 5%

Oxic facies sediments
Mn ~0.15% Fe ~5%
Zone of MnO2 enrichment(no dilution by clastic
sediment input)

Mn2+

–500
Reduced/euxinic

Black
Sea

(H2S zone)
–1000

nt
rre
Cu

Depth relative to redox interface (m)200

Euxinic-facies sediments
Mn < 0.05% Fe > 5%
line

ore

–1500

Sh

Figure 5.19 (a) Depth concentration profile for Fe and Mn in the Black Sea. Fe is depleted in the water column bypyrite accumulation on the sea floor, whereas Mn occurs to moderately high levels in solution in the dominantly
reduced water column. Maximum enrichment of Mn occurs just below the redox interface where Mn2+ is oxidized to
Mn3+ and precipitates. (b) The distribution of Mn in sea floor sediments of the Black Sea. Accumulation of Mn occurs at
the intersection of the zone of Mn precipitation with the sea floor (diagrams modified after Force and Maynard, 1991).

Fe2+ oxidizes more readily than Mn2+ means thatiron can precipitate while Mn remains in solution, and provides a possible explanation for theobservation that iron ores may be spatially (orstratigraphically) separated from manganese ores,
as in the case of the Transvaal basin. In the latter
case, early precipitation of ferric iron out of solution to form huge volumes of BIF (within which
the Sishen and Thabazimbi deposits occur) would
have depleted the water column of iron so thatlater precipitation, at a higher stratigraphic leveland under more oxidizing conditions, resulted in
the formation of Fe-depleted manganese oxide ores.
Manganese deposits are located in sediments ofvariable age, from the Paleoproterozoic to recent.
Some of the biggest deposits are Phanerozoic
in age, such as the Cretaceous Groote Eylandt
manganese oxide ores in northern Australia, andthe Molango district of Mexico where the ores are
made up dominantly of rhodochrosite (MnCO3).
An interesting modern day analogue for the formation of sedimentary manganese ores is provided
by the Black Sea, where active sedimentation
results in ongoing accumulation of MnO2. The

Black Sea is markedly stratified in terms of water
density and composition and the waters below
about 200 m depth are euxinic (i.e. highly reduced,
with H2S stable). Pyritic muds accumulate on thesea floor and this effectively depletes the entire
water column of iron. Mn by contrast is concentrated in the deeper waters, since it exists in solution as Mn2+, but is depleted in the upper oxidized
200 m where it precipitates and settles out. The
depth–concentration profiles for Mn and Fe in
the Black Sea are illustrated in Figure 5.19a,
which shows the zone of manganese enrichmentthat forms just below the redox interface wheremixing between high-Mn euxinic and oxidized
surface waters occurs. At the interface Mn2+ isoxidized to Mn3+ and particulate Mn forms andthen settles at deeper levels where it can redissolve. However, where the enriched mixing zone
intersects shallow sea floor, accumulation ofeither MnO2 or MnCO3 (depending on the local
oxidation potential; see Figure 5.17b) takes placeon the substrate. Mineralization is further
enhanced if the sites of accumulation are protected from dilution by clastic sediment input, as

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

shown for the north-central portion of the BlackSea (Figure 5.19b).
5.3.3 Phosphorites
Phosphorus is a very important element that isessential for the growth and development of mostliving organisms. It is a basic ingredient of some
proteins but, more importantly, it is a building
block for the bones and teeth of vertebrates, aswell as the shells and chitinous exoskeletonsof many invertebrates. Plants also require it for
growth and it is for this reason that phosphate
is such an important ingredient of fertilizer.
In order to sustain global food requirements the
manufacture of artificial fertilizers has become anenormous industry that requires approximately
150 million tons of phosphate per year, the raw
material for which comes predominantly from
phosphorus-rich sediments or phosphorites (i.e. asediment containing more than 15–20% P2O5).
Phosphorites generally form in much thesame environmental niches as do banded ironformations, ironstones, and bedded manganeseores, namely along continental shelves and inshallow marginal marine settings such as lagoonsand deltas. The processes of formation are, therefore, very similar to those for Fe and Mn deposits
and essentially involve the upwelling of cold
ocean water containing above-normal concentrations of metals, and their subsequent precipitation onto the continental shelves. In fact one of
the earliest published accounts of the upwelling
hypothesis, attributed to the Russian scientist
A.V. Kazakov in the 1930s, applied to the formation of phosphate-rich sediments in a continental
shelf setting, and the concept was only later
applied to iron and manganese deposits. Deep sea
drilling and ocean floor exploration have shownthat phosphate concentrations are presently forming off-shore along the coastlines of Namibia andChile–Peru, and are related to upwelling of deep,
cold currents carrying small amounts of phosphate in solution. The same environments are also
well known as prolific fishing grounds, a feature
which substantiates the relationship between
the supply of phosphatic nutrients and biologicalproductivity.

277

When the idea of oceanic upwelling was first
introduced, it was suggested that precipitationof phosphorus onto the sea floor in areas of
upwelling was inorganic and related to decreasing
solubility in the near shore environment. This
notion is over-simplified and it is now known thatthe phosphorus cycle in the oceans is largely
controlled by organic processes. A role for microorganisms in the deposition of phosphate is also
considered likely (Trudinger, 1976). Phosphorus
dissolves in sea water as either PO 43− (stable undervery alkaline conditions), HPO 42− (the dominantanionic complex), or H2PO 4− (stable under acidic
conditions) complexes. These anionic complexes
are absorbed by the organisms living in shallow
marine environments to form their shells, bones,
and teeth. The subsequent accumulation of phosphorus on the sea floor, therefore, is not related
to a chemical redox reaction, as in the cases ofFe and Mn, but occurs after the host organism
dies and settles to the ocean floor. Phosphorus
is released from the decaying organism to form
a calcium phosphate compound, which thenconverts to an impure, cryptocrystalline formof apatite (Ca5(PO4)3[F,OH]) or collophane (acarbonate–apatite). The concentration of calcium
phosphate on the sea floor is described in terms of
the reaction:
3Ca2+ + H2PO4 ⇔ Ca3(PO4)2 + 2H+[5.7]

which suggests that its precipitation is dependent,
among other things, on pH. Figure 5.20 shows
apatite solubility in sea water as a function of pH
and oxalate content of the solution (oxalic acid is
H2C2O4). It is apparent that apatite solubility issignificantly higher under more acidic conditions
and, on this basis, that apatite formation ismore likely to occur under the relatively alkaline,
warm-water conditions that apply to continental
shelves. Deeper, colder and more acidic waters
will promote phosphate solubility.
The formation of apatite/collophane in phosphatic sediments is, however, not a straightforward process and there has been considerable
debate as to whether it develops as a primary
precipitate at the sea water–sediment interface(i.e. where the organic material is decaying) or as a

Solubility of apatite (moles liter–1)

278

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

10–1

pH 5

10–2

pH 6

10–3

pH 7

10–4

pH 8

10–5
10–6
0.0001

0.001
0.01
0.1
Total oxalate concentration (moles liter–1)

Figure 5.20 Estimates of apatite solubility in sea
water (expressed in terms of phosphate concentration)
as a function of solution pH and oxalate concentration(after Schwartz, 1971).

product of diagenesis during later compaction
and dewatering of the sediment. This debate has
considerable bearing on the origin and formationof phosphorite ores and evidence seems to exist
for both processes. Typical concentrations of
dissolved phosphate in the deep portions of thepresent day oceans are low (about 50–100 ppb;
Bentor, 1980) and the shallow marine environment is even more strongly depleted because of
biological uptake. Given the low present day
concentrations of phosphate in solution, even in
the deep ocean, it seems very unlikely that sea
water is saturated with respect to apatite. This, inturn, suggests that the primary precipitation of
apatite at the sea water–sediment interface is also
unlikely. Although this situation applies generally to the present day oceans, it may not, however, have prevailed in previous geological eras
when factors such as ocean chemistry and biological evolutionary development were different.
It may also not apply to restricted, but biologically productive, environments (such as a lagoon)
where high rates of organic decay can produce
localized saturation of phosphate in sea water,
as in the case of the southern African example
described below.
Sheldon (1980) has suggested that phosphateconcentrations in the Precambrian oceans might

have been higher than at present because thelower-O2, higher-CO2 atmospheric conditions
prevailing at that time would have meant a more
acidic sea-water composition and higher phosphate solubility. It is also apparent that oceansthat existed before the development of phosphatedependent organisms did not witness biologicalphosphorus depletion. By the same token, oceansat this time could not produce the environments
in which biologically mediated deposition andconcentration of apatite or collophane might occur.
Substantial Precambrian phosphorite deposits,
therefore, do not exist, and the few that do occur
are likely to have formed from the direct precipitation of phosphate from sea water. As theatmosphere evolved, however, any event thatcoincided with a substantial proliferation of life,
such as the Cambrian explosion, could have had
the effect of removing CO2 from the oceans andatmosphere, increasing pH in the shallow water
environment and further promoting the direct
precipitation of apatite from a saturated sea watercolumn. A more efficient form of concentration,
however, would have been through the decay ofphosphate-dependent life forms and the formation of phosphorites in association with organicrich sediments. This is probably the reason why
all large phosphorite deposits are Phanerozoic inage. It is pertinent to note that the Precambrian–Cambrian boundary, for example, coincides witha period in Earth history when substantial deposition of phosphoritic sediments took place (Cookand Shergold, 1984), good examples of which
include the deposits of Mongolia, southeast Asia
and Australia. By contrast, formation of Mesozoic
and Cenozoic phosphorite deposits must havetaken place essentially during the diagenetic stageof sediment evolution, since the phosphate concentration levels in the water column were too
low to allow for direct apatite or collophane precipitation. In this scenario, the concentrationof phosphate in the organic-rich host sedimentsbuilds up progressively with time such that theinterstitial waters would eventually become
highly enriched in phosphate (measurements of
up to several thousands ppb HPO 2−
4 in solution
have been recorded; Bentor, 1980). This represents
an ideal environment for the formation of diagen-SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

etic apatite as the sediments are compacted andlithified. Such processes would appear to pertain
to deposits such as the Cretaceous–Paleogene deposits of Morocco and possibly also the oolitic phosphorites of Florida. In the latter case, however,subsequent mechanical abrasion and sedimentarytransport of diagenetic aggregations resulted indeposits that are reworked and allochthonous(Riggs, 1979).
Other factors, such as global climatic patterns
and eustatic sea-level changes, are also known tohave played a role in the formation of phosphorites (Cook and McElhinny, 1979; Sheldon, 1980).
Figure 5.21a shows that there is a broad correlation between the number of phosphate deposits
that have formed and periods of sea-level highstand. The explanation offered is that elevated sea
levels and flooding of the continental shelves
enhanced circulation and upwelling. If enhanced
upwelling occurred in an equatorially aligned sea
way (Figure 5.21b) then phosphorite formation
would have been promoted by the biogenic
concentration and/or solubility decrease that
accompanied the existence of such a biologically
productive environment. The late Mesozoic to
Cenozoic phosphorite deposits of North Africa,
deposited in the Tethyan sea way, may well represent examples of this control mechanism. Figure
5.21a shows that there is also a correlation between glacial events and phosphogenesis. In thiscase the lack of vertical mixing in the oceans
during stable periods results in a phosphate buildup in the deep water sink. The development ofstrong trade wind systems between glacial andnon-glacial events in mature, longitudinally orientated oceans would promote the establishmentof major oceanic gyrals which would, in turn,result in strong upwelling along the western edgesof continental land masses (Figure 5.21c). Thismechanism could explain the development of
the Permian Phosphoria Formation, probably theworld’s largest accumulation of phosphate ores,
along the western margin of continental USA in
Pangean times. This scenario also obviously
applies to the present day Atlantic ocean andaccounts for the pattern of upwelling and phosphate deposition along the western margin ofAfrica (see below).

279

A model for phosphogenesis based on present
day deposition
Birch (1980) studied the nature and origin of phosphorites forming off the western and southern
continental margins of South Africa and Namibia.
Exploration of the continental shelf in this region
reveals two types of phosphorite (Figure 5.22a).
One variety developed mainly along the Namibiancoastline and comprises oolitic apatite/collophaneores derived by accretionary growth arising fromthe direct precipitation of phosphate from sea
water. The other type, formed mainly along theSouth African coastline, consists of phosphatic
replacement of fossiliferous limestone and isdiagenetic in origin. Both types of phosphorite
are geologically recent (Eocene to Miocene) in age
and are believed to have formed at the same time.
A model for this type of penecontemporaneousore formation is shown in Figure 5.22b. Diffuse
upwelling along the outer reaches of the shelfresults in reduced nutrient supply and biologicalproductivity, and consequently the rate of phosphate release during decay of organisms on the sea
floor is also low. Apatite is undersaturated andcannot precipitate directly from the sea water, but
subsequent diagenesis can lead to replacement oflimestone by calcium phosphate. By contrast, inthe near shore environment very high biologicalproductivity in restricted shallow embayments
and lagoons is maintained by wind-induced
nutrient upwelling. Mass mortality of life forms
in these restricted environments (especially during summer low tides) leads to the accumulationof abundant organic matter on the substrate withlocal increases in dissolved phosphate content tosaturation levels. Varve-like layers of pure apatite
as well as accretionary oolitic growth provide thetextural support for a mode of formation by direct
precipitation of phosphate at the sea water–
sediment interface.
5.3.4 Black shales
Shales that are rich in organic matter are economically important because they are often enrichedin a large variety of metals, including V, Cr, Co,
Ni, Ti, Cu, Pb, Zn, Mo, U, Ag, Sb, Tl, Se, and Cd.

280PART 3 SEDIMENTARY/SURFICIAL PROCESSES(a)
Sea level (m)200 0 –200

0

GlacialeventsPhosphate
deposits
20

10

0

Cenozoic
100
Mesozoic

Million years

200

300

400

Paleozoic

500

600

(b)

Latitudinal model(c)Longitudinal model

Phosphogenesis

Atlanticlike
ocean

nt
tine
Co
n

Tethyslike
seawayNorthern
gyre
de
Tra

d
e win
Trad elling
upw

Continent

Continent

tine
n
Con

wi
nd
Southern s
gyre

Tra
up de w
we inllin dgtEquatorial upwelling

Figure 5.21 (a) Generalized correlations between the development of phosphate deposits, glacial events andeustatic sea level changes in the Phanerozoic Eon. (b) Model for phosphate deposition related to equatorial upwellingin a latitudinally orientated seaway such as Mesozoic Tethys. (c) Model for phosphate deposition related to tradewind induced circulation patterns in a mature longitudinally orientated ocean (after Sheldon, 1980).

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

281

(a)
Directly precipitated phosphate-rich rocks
Diagenetic phosphate-rich rocks
Shelf break

20°

Walvis Bay
25°
Namibia
Orange River30°

DurbanSouth Africa

0

100 200 300

Cape Town

km
35°

Port
Elizabeth

Agulhas
Bank

15°

20°

25°

30°

(b)

Figure 5.22 (a) Map showing the
distribution of phosphate-richsediments along the continental
margins of South Africa andNamibia. (b) Model showing
contrasting modes of phosphate
deposition in the outer shelf(where diagenetic accumulationof apatite is taking place) and in
restricted near-shore environmentscharacterized by high biologicalproductivity and seasonal massmortality (where precipitation of
apatite directly from sea water isoccurring). Diagrams after Birch(1980).

WindHigh biological productivity,
Dilute upwellinghigh rates of mass
low biological productivity mortality, apatite saturation
low dissolved PO4 content in sea water“Pavement” of diagenetic
phosphate-rich rocks

Phosphatic
varieties
formed bydirectprecipitation

282

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

Although few metalliferous black shales arepresently mined (with the exception of China
where several small deposits are exploited) they
represent a significant metal resource and probably also a source of ore components for youngerhydrothermal deposits. As mentioned previously,
many black shales are spatially associated withironstones (Figure 5.15) and form in the sameenvironments. At a global scale, they formed
penecontemporaneously with ironstones, andtheir major periods of development can be linked
to the first order eustatic sea-level highstands in
the Ordovician–Devonian and Jurassic–Paleogene
intervals (see Chapter 6). In this respect they arealso linked to the Ocean Anoxic Events (OAEs)
that are now well documented in the Phanerozoic
Eon. Sea-level highstand resulted in continental
margin flooding and transgression of muddy
sediment across the shelf. The deeper water conditions that generally applied resulted in poorer
circulation, a reduction in oxygen levels, andwidespread oceanic stagnation (Van Houten andAuthur, 1989), providing the kind of depositional
environment favorable for metalliferous blackshale formation.
Well known examples of metalliferous blackshales include the Cambro–Ordovician Alum shaleof Scandanavia, which is particularly enriched inuranium, and the Devonian New Albany shale ofIndiana, USA, which has significant concentrations of Pb as well as V, Cu, Mo, and Ni. The
euxinic shaley sediments forming on the floor of
the Black Sea represent an example of manganese,
as well as Co, Cu, Ni, and V, enrichment, in apresent day sedimentary basin (Holland, 1984).
Black shales are characterized by significant
quantities of organic carbon which is preserved
from degradation/oxidation by the almost totallack of free oxygen in the immediate environmentof deposition. This environment is not only anoxic
but also euxinic (where reduced sulfur species arestable) and it is this combination, together with alack of clastic dilution, that provides the optimal
conditions for development of metalliferous blackshales (Leventhal, 1993). A sediment organic content of even 1% will bacterially and chemicallydeplete the oxygen content of the immediate
environment to virtually zero. This microenvir-

onment promotes the growth of sulfate-reducing
bacteria which in turn produce HS− or H2S(depending on the local pH) and the developmentof euxinic conditions. The biologically mediateddevelopment of euxinic waters can be describedin terms of the following reaction:−
R(CH2O)2 + SO2−
4 ⇒ R + 2HCO3 + H2S[5.8]

where R(CH2O)2 is an abbreviated representationof a complex (organic) carbohydrate molecule.
The subsequent concentration of metals in thisenvironment is brought about by the affinity ofpositively charged metals in solution in the sea
water for the thio complex, as demonstrated in
the following simplified reaction:HS− + Me2+ ⇒ MeS + H+

[5.9]

where Me is a metal. The above reactions imply acorrelation between organic carbon, sulfur, andmetal contents in black shales since the higher
the content of sulfate-reducing bacteria, the more
sulfide is produced and the greater the degree ofextraction of metal from the sea water column
would be. Such a correlation is well demonstrated
in Figure 5.23a, for Devonian black shales in theeastern USA, and lends support to the processes
outlined above.
The source and metal content of black shales
will obviously vary from place to place andthis accounts for the observation that different
horizons have characteristic metal signatures.
In oceanic settings submarine exhalative vents, or
black smokers, represent likely metal sources,
whereas in continental margin settings (such as
the Black Sea) metals are derived from terrigenous
clastic input. Figure 5.23b summarizes the geological and environmental conditions necessary forthe optimal development of a metalliferous blackshale. This type of sedimentary rock represents
a potentially important source of exploitable metals in the future.
5.3.5 A note on ocean floor manganese nodules
Since their discovery during the 1872–6 oceanographic expedition by HMS Challenger, it has

283

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

(b)

(a)
6

Sulfur (wt%)

ce

O

ur

o
lS

5
a
et

rg

Felsic(Th, U, REE, Pb, Ba)

4

M

3

Mafic
(Ni, Cr, Co, Y, Cu,
Fe, Mn, Ag, Zn)

Terrestrial source

an

ic

M

at

te

2

r

Marine source

1
Marine Basin/Sink
0

Source of SO42–
Stratified water column
Restricted basin
Slow sedimentation rate
(less clastic dilution)

Molybdenum (ppm)

70
60
50

Anoxic Environment

40

Low or no O2 in water column
sulfate reduction in sediment
Good

30
20

rvation of odesd sulfier an

rga nic matt

Euxinic Environment
Sulfate reduction in watercolumn and sediments
H2S present
High S/C
Buria
l and time

10
0
40

Metal richblack shale

35
Uranium (ppm)

prese

30
25
20
15
10
5
0

0

1

2

3 4 5 6 7 8
Elemental carbon (wt%)

910

Figure 5.23 (a) Plots of carbon content versus sulfur, molybdenum, and uranium for Devonian black shales fromthe eastern USA (after compilations by Holland, 1984). (b) Schematic outline of the geological and environmental
conditions required for the optimal formation of metalliferous black shales (after Leventhal, 1993).

become apparent that a vast resource of Mn, Fe,Cu, and Ni (as well as lesser amounts of Co, Zn,
and other metals) is contained within the (ferro)manganese nodules that occur in the deep, pelagic
portions of all the major oceans. Manganese nodules typically occur in parts of the ocean basins
where sediment accumulation rates are very low(less than 7 meters per million years; Heath, 1981)

and are absent from areas of rapid sedimentation,
as well as from equatorial and high-latitude
regions of high biological productivity and chemical sedimentation. They appear to be best developed in the Pacific Ocean, where exploration has
identified large areas of prolific nodule formation(up to 100 nodules per square meter). These
areas also contain nodules that have higher than

284

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

60°N
RussiaNorth
America

0.4

Wt%
Mn 25
Fe 7
Ni 1.1Cu 1.1
Co 0.2
Zn 0.1
1.0

0.5

40°

20°

1.0
ClarionClipperton
regionEast Pacific

1.0

Australia

0°

Rise

Equator

20°

0.5

40°S

100°E

140°

180°

140°

100°

normal metal contents, with up to 2 wt% combined Cu + Ni and substantial Co and Zn (Figure 5.24). One such area is the Clarion–Clipperton
region, to the west of Mexico.
Manganese nodules are typically ovoid, a few
centimeters in diameter, and have the appearance
on the sea floor of a burnt potato. In cross section
they are complexly layered, often partly concentric, and represent concretions that haveapparently nucleated around a clastic or biogenic
fragment. They often exhibit internal radiating
shrinkage cracks reflecting a change in mineralogy from hydrated Fe and Mn oxyhydroxides to avariety of more compact, but chemically complex, Fe and Mn oxides such as vernadite, birnessite, todorokite, goethite, and hematite (Heath,1981). The formation of manganese nodules is acomplex process and still not fully understood.
Metals are undoubtedly derived from the sea
water itself, but ultimately come from various
sources including submarine exhalative vents,
clastic and volcanic input, and diagenesis of

60°W

Figure 5.24 Map showing subequatorial zones of base metalenriched manganese nodule
concentration in the Pacific Ocean,contoured with respect to thecopper content (in wt%) of nodules.
Nodule-free areas occur close to thecontinents, where sedimentation
rates are high, as well as along theequator and East Pacific Rise,
where biological productivity ishigh. The Clarion–Clipperton
region, a particularly enriched
zone, is boxed (after Heath, 1981).

pelagic sediments. The formation of the nodules
and the concentration of metals within them
have been attributed to one or more of the following mechanisms:• Settling of clay and biogenic (fecal) debris that contained either absorbed or ingested metals obtained
from the sea water, with subsequent release of
metals during dissolution and decomposition.
• Direct precipitation of metals from sea water
onto a suitable substrate which forms the nucleus
around which concretionary growth takes place.
• Upward diffusion of metals in ocean bottom
sediment pore waters.
• Authigenic reactions in ocean bottom sediments during alteration and compaction.
• Bacterial activity and the oxidation of transition metals. In this regard, certain types of bacteria (such as Metallogenium) are known to beparticularly prevalent in the oxidation and subsequent precipitation of labile Mn2+ and have been
detected in manganese nodules from the Atlantic
Ocean (Trudinger, 1976).

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5• Preferential uptake of dissolved transition
metals into specific Fe and Mn oxyhydroxideminerals, such as the affinity of Cu and Ni for
todorokite (a multi-element Mn–oxyhydroxide).
The current level of uncertainty about theprocesses involved in the formation of deep ocean
manganese nodules is reflected in the variety ofsuggested mechanisms listed above. Many ofthese suggestions are not mutually exclusive andit is likely that nodule formation is a long-lived
and complex process. The biggest problem withthe exploitation of manganese nodules at present
is not so much technical as it is political and economic. The technologies required for exploiting
manganese nodules at depth in several kilometersof ocean water could be overcome, but the formulation of an internationally acceptable Law of the
Sea is a much bigger and more contentious issue.
Manganese nodules, nevertheless, represent anenormous metal resource for the future.
5.3.6 Evaporites
The bulk of the world’s production of rock salt(halite), as well as of potash, borates, and nitrates
for agricultural fertilizers, comes from rocks
known as evaporites. These are chemical precipitates that form as a result of evaporation ofa brine, usually derived from sea water. There
are two main environments in which evaporites
form. The major one, in which the so-called“saline giants” of the world formed, is marginal
marine in setting and represented by large lagoonsor restricted embayments into which periodic orcontinual sea water recharge occurs. This type ofdeposit may be both laterally extensive and thick,but is characterized by a relatively limited rangeof mineral precipitants derived from sea water,which itself has remained fairly constant in termsof bulk composition over much of geologicaltime. The second setting is intracontinental
and lacustrine and results in much smaller,thinner deposits which are characterized by amore diverse range of mineral precipitants sincecontinental fluvial input introduces a broader
range of dissolved ingredients to the lake waters.
Examples of deposits representing the former
include the Permian Zechstein Formation, which

285

formed in a shallow sea covering large areas of
the United Kingdom, The Netherlands, Germany,and Poland, and also the Silurian Salina Formationof New York and Michigan states in the USA.
Typical products mined from these deposits
include halite, potash, and sulfates. Probably thelargest known evaporite sequence occurs in late
Miocene strata beneath the floor of the Mediterranean Sea, where deposits extend laterally forover 2000 km and may be up to 2 km thick. Muchof this resource is obviously unexploitable, but
is so large that its formation must have caused
a significant reduction in oceanic salinity at thetime (Kendall and Harwood, 1996). Examples of
intracontinental lacustrine deposits include thedry lakes of Chile, Death Valley in California, and
the Great Salt Lake of Utah. In addition to the
products mentioned above, these deposits are also
important producers of borates (from California)
and nitrates (from Chile), as well as Mg, Br, and Li(from Utah).
The principles behind the formation of evaporite deposits are relatively simple in theory,
although actual deposits can be complex andexhibit a large variety of different mineral salts
and paragenetic sequences. As sea water or brine
evaporates and water vapor is removed into theatmosphere, the salinity (i.e. the total content
of dissolved salts in solution) of the residual
solution increases and individual salts precipitate
as their respective solubility limits are reached.
The order or sequence of precipitation reflects thescale of increasing solubility at a given temperature, such that the salts with the lower solubilities precipitate first, followed consecutively bysalts with progressively higher solubilities. The
relative quantities of precipitated products in an
evaporite deposit are also constrained by thesolubility limits of the various salts dissolvedin the brine solution. These considerations seemto suggest a relatively straightforward precipitation process and uniform paragenetic sequence. Infact the chemistry of brine solutions is very complex and factors such as convection dynamics,
post-precipitation diagenesis, and equilibrium
as opposed to fractional crystallization combine
to ensure considerable diversity in the nature ofthese deposits.

286PART 3 SEDIMENTARY/SURFICIAL PROCESSES(a)

Sea water density (in g cm–3) increases

Open
ocean

1.05–1.1

1.1–1.2Land

1.2–1.3

KCl and bitterns
CaCO3

Algal
mound

CaSO4

NaCl

(b)
Mineral precipitate

Sequence of precipitation

Sea water g kg–1

First

Last

CaCO3

0.11

1000

Calcite (CaCO3)

CaSO4

1.20

900

Gypsum/anhydrite(CaSO4/CaSO4.2H2O)

NaCl

26.90

MgSO4

2.25

KCl

0.70

MgCl2

3.18

borates,sulfates

0.01
0.02

Total

34.48

Volume (liters)

0.03
0.09

Halite (NaCl)

700
600
500
400
300
200

Epsomite (MgSO4.7H2O)

a
se
of
me
Volu

NaBr
MgBr2

800

wa

Kainite (KMgClSO4.3H2O)
Sylvite (KCl)Bitterns

Carnallite (KMgCl3.6H2O)
ter

Borates andcelestite (SrSO4)

100

Density of sea water 1.0

1.1

1.2
–3
g cm

1.3

1.4

Figure 5.25 (a) Schematic cross section showing the important features necessary for the formation of large marine
evaporite sequences. (b) Paragenetic sequence for an evaporite assemblage from typical sea water containing theingredients shown in the left hand column. The amount of sea water (per 1000 liter volume) that has to evaporate
in order to consecutively precipitate the observed sequence of mineral salts is shown by the curve adjacent to theparagenetic sequence (diagrams modified after Guilbert and Park, 1986).

Marine evaporites that formed from sea waterthat was well mixed and had a constant composition over much of geological time tend to be dominated by the same assemblage of major mineral
precipitates, namely halite (NaCl), gypsum/
anhydrite (CaSO4/CaSO4.2H2O), and sylvite (KCl).
Evaporite deposits can, however, also containminor accumulations of other salts, especiallythose representing the late stage precipitation ofhigh solubility compounds. The latter typicallycomprise hydrated, multi-element, Mg, Br, Sr, K
chlorides and borates, and are called bitterns.
The preservation of bitterns in evaporite deposits
appears to be controlled largely by geological factors. Figure 5.25a shows a schematic cross section
through a shallow, marginal marine setting inSEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

which an evaporite sequence is being deposited.
Sea water recharge ensures replenishment of thesolution and sustained mineral precipitation fromthe brine. Intense evaporation results in marked
density stratification and accumulation of highly
saline, dense brines in the bottom waters. In
environments where there is a constriction at theentrance of the basin, there is little opportunity
for this dense brine to discharge back into thedeep parts of the open ocean and, hence, the highsolubility components of the solution remain
within the basin to eventually precipitate out asbitterns once the solubility limits are exceeded.
Bitterns do not accumulate if reflux of the dense
bottom water brines to the open ocean occurs
because the high solubility components of thesolution are then flushed out of the basin. In thelatter case, if the bitterns do not precipitate,
the evaporite mineral assemblage may be truncated or incomplete and the deposit less diverse interms of the commodities it could exploit. Figure
5.25b illustrates the typical mineral paragenetic
sequence for evaporite deposits as a function of
precipitation from sea water, the composition ofwhich is shown on the left hand side of the figure.
The proportion of sea water remaining after evaporation (as a function of a 1000 liter volume) and
its increasing density are also shown in relation to
the precipitation sequence.
The formation of the giant Phanerozoic saline
deposits is a contentious issue, mainly because
modern analogues do not exist and ancientevaporite sequences are generally altered andtectonically modified. The evaporite deposits
on the floor of the Mediterranean Sea are now
believed to have formed in a sabkha environment,
implying that much of the Mediterranean basin
was subaerial and desiccated during late Miocene
times (Kendall and Harwood, 1996). Evaporitic
minerals which form in a sabkha environment
precipitate from groundwater brines that evaporate in the capillary zone immediately above thewater table. Recharge of the brine occurs either by
periodic flooding or by sea water seepage into thesabkha basin. A good example of a complete evaporate deposit formed as part of a Phanerozoic
saline giant is the Boulby Mine in the Zechsteinof northeast England.

287

5.4 FOSSIL FUELS – OIL /GAS FORMATION AND
COALIFICATION

A fossil fuel is formed from the altered remains oforganic matter (plant and animal) trapped in sedimentary rock. The hydrocarbon compounds thatform during the burial and degradation of organic
material retain a significant proportion of thechemical energy imparted into the original livingorganism by the sun. This energy is harnessedin many different ways by burning or combusting the fuel. The main fossil fuels are coal andpetroleum, the latter being a general term thatencompasses crude oil and natural gas (mainly
methane), as well as solid hydrocarbons.
Although most ore deposit books do not include
descriptions of fossil fuel occurrences, it is feltthat any text dealing with ore-forming processes
would be incomplete without a brief discussion
on the accumulation of hydrocarbons in the
Earth’s crust. There is an enormous volume of
published literature on the origin and nature offossil fuel deposits and, consequently, the following section presents only brief overviews relevant
to oil/gas formation and coalification processes.
Mention is also made of tar sands, oil shales, andgas-hydrates, all of which represent potential fuelresources in the future.
5.4.1 Basic principles
The global carbon cycle is complex, involving
both circulation, mainly in the form of gaseous
CO2 and CH4, and storage, as reduced carbon orcarbonate minerals, in different reservoirs such as
the deep ocean and sedimentary rocks. Organic
processes (mainly photosynthesis) ensure that asmall proportion of the carbon in the global cycle
is combined with hydrogen and oxygen to form
the molecular building blocks of the various biota
that inhabit the Earth’s surface. The areas ofgreatest organic productivity on Earth are represented on the continents by tropical vegetative
zones, and in the oceans by areas of cold current
upwelling. As a first approximation, therefore,
it is logical to expect that fossil fuels are most
likely to have formed in these two environments.
Indeed, coals develop mainly in continental

288PART 3 SEDIMENTARY/SURFICIAL PROCESSES

Mainly landderived
plants

Mainlyphytoplankton

Peat
Lignite
1

30

Diagenesis

Sub-bituminous
coal

Catagenesis

Bituminous
coal

Kerogen

Depth (km)

90

3
Oil

120

4

5

T (°C)

60

2

Gas

150
Metamorphism

settings, from accumulation of vegetation in either
humid, tropical swamp environments or, as in thecase of younger coal seams, more temperate topolar regions. Oil and gas deposits, by contrast,
develop mainly from the accumulation of phytoand zooplankton in marine settings.
When organisms die they decompose by bacterial decay and/or oxidation, and rapidly break
down to relatively simple molecular constituents
such as CO2 and CH4. Such a process would not be
conducive to the formation of fossil fuels, which
requires preservation of much of the complex
organic hydrocarbon material. Suitable source
rocks from which fossil fuels accumulate must,
therefore, have formed in reducing environments
where sedimentation rates are neither too slow
(in which case oxidation and molecular breakdown might occur) nor too rapid (in which casedilution of the total organic matter content of therock would occur).
The prevailing theory for the origin of both
petroleum and coal is based on the notion thatorganisms are buried during the sedimentary
process and then subjected to a series of alteration
stages as pressure and temperature increase. As
shown in Figure 5.26, progressive burial of phytoplankton results in the early liberation of CO2 andAnthracite

Figure 5.26 Simplified scheme
illustrating the formation of oil, gas,and coal by the progressive burial ofdifferent types of mainly vegetative
matter (after Hunt, 1979).

H2O and formation of kerogen (a collective namefor sedimentary organic matter that is not soluble
in organic solvents and has a polymer like structure; Bjørlykke, 1989). As the kerogen “matures”
the long-chain covalent bonds that typify organic
molecules are progressively broken to form lower
molecular weight compounds. At around 100–
120 °C and burial depths of 3–4 km (depending on
the geothermal gradient), a liquid hydrocarbon
fraction develops which can then migrate fromthe source rock. This interval is known as the“oil window.” With further burial and crackingof molecular bonds, significant volumes of gas(mainly methane, CH4) develop, which are alsoamenable to migration. The solid residue remaining in the sediment is referred to as kerogen, butwith progressive burial it devolatilizes further
and has a composition approaching pure carbon(graphite). By contrast, when humic, land-derived
vegetation is buried, little or no liquid oil isformed, although significant volumes of gas
will, again, be generated. The solid residue thatremains in this case is more voluminous and compacts to form coal seams. The nature of the solid
coal residue changes with depth, from peat andlignite at shallow burial depths (less than about
500 m), to bituminous coal and subsequently

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

Table 5.1 Average chemical compositions (in wt%) of
the main organic building blocks compared to those of
petroleum and a typical coal

Carbohydrates
LigninProteins
LipidsPetroleum
Coal

C

H

O

N

S

44
63
53
76
85
70

6
5
7
12
13
5

50
31.6
22
12
0.5
23

–
0.3
17
–
0.5
1

–
0.1
1
–
1
1

Source: after Hunt (1979).

anthracitic coal (at depths of around 5000 m).
The calorific value (i.e. the amount of energy produced on combustion) of the coal increases withmaturity or burial depth.
5.4.2 Oil and gas formation
All living organisms are made up of relatively few
molecular building blocks that have apparently
changed little over geologic time. The main
molecular building blocks involved in petroleum
formation include:• Carbohydrates and related substances – theseare mono- and polysaccharide polymers and include the sugars, chitin and lignin, the latter beinga major precursor to coal.
• Proteins – these are high molecular weightamino acid polymers and represent one of themost important constituents of life processes.
• Lipids, which are represented in animal fats andvegetable oils and are abundant in crude oils.
• Other substances such as resins and pigments
are of lesser importance but also occur variably inboth plant and animal matter.
Plants comprise mainly carbohydrates (40–70%),
although the higher, woody forms also contain substantial lignin for strength. Phytoplankton contains
around 20% protein. Animals, by contrast, aremade up mainly of proteins (55–70%) with lesser
carbohydrates and lipids but no lignin (Hunt, 1979).
The average chemical composition of these various organic building blocks is shown in Table 5.1.
As the original organic constituents are buried,they are subjected to increases in pressure and289

temperature, resulting in systematic changes
which are divided into three stages, termed
diagenesis, catagenesis, and metamorphism. This
progression, and its application to the burial oforganic matter in sediments, together with theorganic processes and changes that occur in each,are summarized in Figure 5.27. Diagenesis refers
to the early biological and chemical changes that
occur in organic-rich sediments prior to the pronounced temperature-dependent effects of later
reactions (i.e. less than about 50 °C). During thisstage the biopolymers of living organisms undergo
a wide variety of complex low-temperature reactions. Put simply, they are either oxidized or
attacked by microbes and converted into lesscomplex molecules which may, in turn, react andcondense to form more complex, high molecularweight geopolymers that are the precursors to
kerogen. Biogenic reactions during diagenesis
also produce significant amounts of gas, termed
biogenic gas, which typically escapes into theatmosphere or water column and is not retainedin the sediment. Catagenesis occurs between
about 50 and 150 °C and is the most important stage as far as petroleum generation is concerned. During this stage temperature plays animportant role in catalyzing reactions, the majority of which result in the formation of light hydrocarbons from high molecular weight kerogenby the breaking of carbon–carbon bonds. In thisprocess, known as thermal cracking, a complex
organic molecule such as a paraffin or alkane will
split and form two smaller molecules (a different
paraffin and an alkene or olefin). Each of the products contains a carbon atom, with an outerorbital electron from the original covalent pairingthat is now shared. The shared electron meansthat the resultant alkane and alkene each containa “free radical” that makes them reactive andamenable to further breakdown. Alternatively,
reactions may occur by catalytic cracking, wherea carbon atom in one of the resultant molecules
takes both electrons, thereby becoming a “Lewis
acid” or double electron acceptor. The moleculethat loses the electron pair is called a carbonium
ion and with its net positive charge is alsoamenable to further decomposition. For example,
alkanes subjected to catalytic cracking could

290PART 3 SEDIMENTARY/SURFICIAL PROCESSESORGANIC MATTERBIOPOLYMERSPROCESSES

MATURATION RANGE

Carbohydrates, proteins,
lipids, lignin

MICROBIAL ALTERATION, HYDROLYSISBIOMONOMERSSugars, amino acids,
fatty acids, phenols

DIAGENESISCONDENSATION,
REDUCTION,POLYMERIZATION

Nitrogenous andhumic complexes

50 °C

THERMOCATALYTIC CRACKINGCATAGENESISGEOPOLYMERSGEOMONOMERSPetroleum hydrocarbons
and low molecular weightorganic compounds

150 °CTHERMAL CRACKING

METAMORPHISM

Gas and pyrobitumens

yield gaseous products such as ethane or butane(Hunt, 1979). Catalytic cracking tends to be thedominant process in petroleum generation up to
about 120 °C, but at higher temperatures thermal
cracking becomes increasingly important.
The integrated effects of time and temperature
have important implications for petroleum generation. Temperature and time are inversely related
in terms of petroleum productivity so that, forexample, a given quantity of oil formed at 110 °C
over 25 million years would require 100 million
years to form at 90 °C (Hunt, 1979). This relationship is quantified on a plot of time versus temper-

Figure 5.27 Summary of the stages
and processes involved in thetransformation of organic matter
during burial to form oil and gas(after Hunt, 1979).

ature in Figure 5.28, where data for selected
petroleum-bearing basins are also shown. The
graph can be very useful in petroleum exploration
as it predicts, for example, that oil would not be
formed in a young basin with low geothermal
gradient because the threshold for efficient hydrocarbon generation might not yet have been reached.
Likewise, little oil could be preserved in an old,
hot basin as it would all have been destroyed during metamorphism. The time–temperature evolution of sedimentary basins is, therefore, a criticalfactor in petroleum exploration. As a generalrule, oil resources are sought either in young, hot

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

600

Time–temperature limits
for petroleum
prospective basins
Oil
Gas

1

400
300
200

2
8

Time (Myr)150

Figure 5.28 Graph showing therelationships between temperatureand time with respect to thegeneration of oil and gas insedimentary basins. The shaded
area refers to the optimum range ofconditions for petroleum generationand the points are actual examplesof petroleum-bearing basins fromseveral locations around the world(after Connan, 1974).

3

100
4
60
No oil
40

291

Main
oil
and
gas
zone

Oil
phase
out
5

1 Amazon Basin, Brazil
2 Paris Basin, France
3 Rio de Oro,
Gas
West Africa
phase 4 Douala Basin,
out
Cameroon
5 Camargue Basin, France
6 Offshore Taranaki
Basin, New Zealand
7 Los Angeles Basin,
California
8 Aquitaine Basin, France

6
20
7
20

40

basins or in old, cold ones, as identified by the relevant field in Figure 5.28.
Source considerations
Most of the organic matter on Earth can beclassified into two major types, namely sapropelic, which refers to the decomposition productsof microscopic plants such as phytoplankton, andhumic, which refers essentially to the maturation
products of macroscopic land plants. Sapropelicorganic matter has H/C ratios in the range
1.3–1.7, whereas humic matter has a lower H/Cratio around 0.9. These compositional differences
have led to a more rigorous classification of
kerogen types that has relevance to the generationof fossil fuels. The classification was originallydevised by Tissot et al. (1974) in terms of theso-called Van Krevelen diagram, a plot of H/Cagainst O/C (see inset in Figure 5.29). A more
recent version of the Tissot classification is
described in Cornford (1998) using a plot of hydrogen index versus oxygen index (Figure 5.29), twoparameters which are analogous to H/C and O/C,

60

80 110 140 180 230 280
Temperature (°C)

respectively, but can be obtained directly by spectrometric analysis.
The classification scheme recognizes threemain kerogen types, with a fourth type having
little or no relevance to petroleum generation.
Type I kerogen is derived from algal material and,on maturation, yields mainly oil. Type II kerogen
is obtained from the breakdown of plant spores,
pollens, exines, and resins, but it may also comprise bacterially degraded algal (Type I) organic
matter. The Jurassic source rocks of the North Sea
oil deposits represent an example of hydrocarbon
derivation from the Type II kerogen. Type III
kerogen is derived from humic land plants richin lignin and cellulosic tissue and on maturation
yields only gas. As mentioned previously, thisis the material from which coals are formed.
Type IV kerogen is obtained from oxidized woody
material and has no hydrocarbon potential. The
arrows in Figure 5.29 indicate the compositional
trends of different kerogens with progressive
burial and maturation, and all coalesce toward theorigin, represented by pure carbon (graphite). In
considerations of source it should be noted that

292

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

2.0

1.5

Principal products ofkerogen evolution
CO2, H2O
I
Oil
Gas
s II
i
s
ene

Catagenesis

H/C

g

Dia

1.0

III

0.5
1000
0.1

0

0.2
O/C

0.3

Type I
Algae(plankton)

800

Hydrogen index

Oilprone
Bacteriallydegradedalgae ormixtures

600

Resin
cuticle,sporesType II

400

200
Type III

0

0

Land
plants:cellulosic
tissue,wood

Altered(oxidized)
Type IV
wood
50
100
150
Oxygen index

Waxy oilandcondensate
prone

Gasprone“Dead
carbon”
No potential

oil-prone Type I and II kerogens will yield not
only oil, but also both “wet” (i.e. with a high condensable fraction) and “dry” gas (mainly methane).
Gas-prone kerogen (usually Type III) will produce
mainly dry gas.
Modern analytical and modeling techniques
allow the quantitative estimation of gas andoil yields to be calculated from a source rockundergoing progressive burial. A summary ofhydrocarbon yields from two different source
rocks is shown in Figure 5.30. Figure 5.30a shows
the quantities of petroleum products derived froma source rock comprising Type II kerogen and differentiates between the condensate, comprising

Figure 5.29 Classification scheme
for kerogen types based on thehydrogen index and oxygen index(after Cornford, 1998). Inset diagramshows the original classification of
kerogen on the basis of H/C and
O/C atomic ratios (after Tissot andWelte, 1984).

high (C15+) and low (C2–C8) molecular weight
oils, and a dominantly methane gas component.
Figure 5.30b is the equivalent scenario for asource rock sediment comprising kerogen that is
only gas-prone, which on maturation initiallyyields only a minor amount of wet gas (ethane,
propane, butane, etc.) but mainly dry gas (methane),leaving a coal residue. The latter process isrelevant to the generation of coal-bed methane
resources.
Source rock considerations are important during petroleum exploration. The origin of organic
matter is obviously a pointer to the depositional
environment of the sedimentary rock and, there-

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

(a)

293

(b)
3

3

8

6

Billion m3 CH4/km3 seam coal

Million m oil/km rock

12

10

4

8

6

2

100

200

300

4

Billion m3 CH4/km3 rock00
Early biogenic gas1

Early matureoils fromwell drained
source

3
C

8 –C
15

C15+
4

Depth (km)2

Temperature (°C)

50

Main phase ofoil generationfrom thicksource rock

CH4

Biogenic gas

Early, wetter gas
expulsion from thincoal seams

100

Later expulsionof dry gas fromthicker coal

150

C2–C8

CH4
5
200

Unexpelled oil
cracks to gas
Figure 5.30 Maturity curves showing (a) the cumulative generation of oil types and methane gas from sedimentsource rock originally containing a 1% total organic carbon (TOC) content of Type II oil-prone kerogen; and (b) thegeneration of wet and dry gas only from a 1 km3 sediment volume of gas-prone kerogen which ultimately yields a
coal deposit (after Cornford, 1998).

fore, also useful in placing constraints on thenature of oil/gas migration and entrapment.
Petroleum migration and reservoir
considerations
One of the features of oil and gas formation is
the fact that petroleum originates in a fine grained
source rock and then migrates into more
permeable, coarser grained, reservoir sediments.
Knowledge of petroleum migration and entrapment processes is obviously important to both the
exploration and exploitation stages of the industry. Although the processes are complex, they can

now be accurately evaluated in terms of fluid
hydraulic principles and sophisticated computermodeling techniques.
Petroleum engineers differentiate between
primary migration, which accounts for the movement of oil and gas out of the source rock during
compaction, and secondary migration, which
describes flow within the permeable reservoir, as
well as the segregation of oil and gas. As theorganic source rock is lithified during diagenesis,
water (usually a low to moderate salinity brine) is
expelled from the sediment to form a connate
fluid (see Chapter 3). In the early catagenic stagesof hydrocarbon maturation, therefore, oil and gas

294PART 3 SEDIMENTARY/SURFICIAL PROCESSES

migrate in the presence of water. The actual
mechanisms of hydrocarbon migration and therole of water in this process are not completely
understood. Oil is unlikely to migrate in aqueous
solution since hydrocarbons generally have low
solubilities in water (e.g. 3 ppm for pentane,
24 ppm for methane, 1800 ppm for benzene, allat room temperature; Bjørlykke, 1989). However,small hydrocarbon molecules dissolve more readily
than large ones and methane (CH4), for example,
does become fairly soluble in water as pressure
increases (about 7500 ppm at 6000 m depth; Hunt,
1979). Methane and other low molecular weightcompounds can, therefore, be dissolved at depth
and then released as the aqueous solution rises tothe surface. Major amounts of hydrocarbon liquids
and high molecular weight compounds, however,are unlikely to flow by this mechanism. At catagenic pressures and temperatures, hydrocarbon
gases themselves will increasingly dissolve oils andin some source rocks there is evidence that primary migration is achieved by gas phase dissolution.
Although dissolution mechanisms are likely
to play a role in petroleum flow, they cannot
account for the huge oil accumulations observedin the major oil producing regions of the world.
Two other mechanisms are considered to be more
important during the primary flow of hydrocarbons and these are oil phase migration and diffusion. The movement of crude oil out of the sourcerock is initiated once compaction has driven off
much of the bulk pore water in the sediment.
Initially a source rock will usually contain a relatively high proportion of pore water relative to oil
and in such a case the hydrocarbon will be unable
to move. This is mainly because the flow of oil is
impeded by the presence of water which, becauseof its lower surface tension, tends to “water-wet”
the sediment pore spaces. In such a situation,
oil globules floating in the water-dominated porespaces cannot generate the capillary forces required
to initiate migration (Figure 5.31a). As compaction
progresses, however, the bulk of the pore water isdriven off and the little remaining water is structurally bound on clay particles. Oil now occupies
enough of the pore volume to be subjected tocapillary forces and start flowing through the rock
(Figure 5.31b). Once “oil-saturated” pathways are

established in the rock, oil migration becomesfeasible because permeability is not impeded by
the presence of water. In the less common situation when pore spaces are initially “oil-wet”
because of very high initial organic/hydrocarbon
contents, then porosity reduction is not necessary
to invoke the unimpeded flow of oil through thesource rock.
Another way in which hydrocarbon molecules
could migrate in the primary environment isby diffusion along an activity or free energy gradient (Hunt, 1979). Figure 5.31c illustrates, at a
nanometer scale, the arrangement of structured
water tetrahedra along the edge of a clay particle.
Hydrocarbon molecules or aggregates in the sediment pore space will typically be encased bya jacket of water molecules to form hydrate
or clathrate compounds. Interaction of the hydrocarbon clathrate with structured water would,
from a thermodynamic viewpoint, require energy
to facilitate the breakdown of one or both
molecules. A diffusional energy gradient is, therefore, set up that promotes migration of the hydrocarbon clathrate towards lower free energy and,
therefore, away from the structured water clay
mineral edge. This type of diffusion translates, ata geological scale, to movement of hydrocarbons
from finer grained shales towards coarser sands,
which accords with natural observations. Although
not directly involved as the transporting agent
in primary migration, it is evident that water is
implicated in all the mechanisms involved. In
the case of low organic content source rocks
generating mainly gas, hydrocarbon migration
would occur by diffusion, in aqueous solution, or
directly as the gas phase. In high organic content,
oil-prone, source rocks, however, hydrocarbons
will migrate predominantly as the oil phase.
Intermediate situations might well exhibit theentire range of migration processes (Hunt, 1979).
At a geological scale, where oil and gas are
migrating over tens to hundreds of kilometers,
the main factor that controls hydrocarbon migration pathways is the existence of a pressure differential in the host rock. Particularly relevant topetroleum migration is the concept of fluid overpressure, described in more detail in section 3.3.2
of Chapter 3.

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

(a)

295

Bulk pore water
Sediment particle

Water flows outof source rocks

Oil globule prohibitedfrom moving(b)Structured pore water

Oil flows throughto reservoir rocks

Oil

(c)

Figure 5.31 Explanation of
oil-phase primary migration.
(a) “Water saturated” pore spaces ina sedimentary source rock in whichlittle or no oil migration takes placebecause oil capillarity is impededby the presence of water-wet pores
which block the pore throats.
(b) “Oil saturated” pore spaces
formed after compaction of thesediment has driven off much of thebulk pore water, which facilitates
capillary-related oil migration byremoving pore throat blocking
free water (after Bjørlykke, 1989).
(c) Simplified arrangement, at a
nanometer scale, of a hypotheticalpore space and the setting inwhich diffusional migration ofhydrocarbons takes place awayfrom structural water-boundmineral grains (after Hunt, 1979).

Structured
pore
water

Bulk
pore
water
Migration ofhydrocarbon clathrate
along a diffusional
gradient

Sed

ime

Water molecule
Hydrocarbon
complex

Clathrate

nt p

artic

le (c

lay)

296

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

Orinoco basin

Shale

Sandstone

Sandstone
Shale

0

0

220

440

0
220
440
Excess hydrostaticpressure (KPa m–1)

If the removal of pore water is impeded by low
permeability in, for example a fine grained shale,then compaction will be retarded and fluid pressures increased to values above normal hydrostatic pressures. Fine and coarse sediment will
expel pore waters at different rates during burial
and will, therefore, compact along different pressure gradients. Overpressured fluids will tend toexist in the finer grained sediments. Figure 5.32
shows actual measurements of fluid pressures in aseries of drill wells from the Orinoco Basin andthe hydrostatic overpressures that occur in the
poorly drained shale units relative to the welldrained sandstone. This information allows the
migration pathways to be identified and showsthat the main sandstone aquifer can be fed by
fluids from both above and below. Several factors
actually control the stresses that cause hydrostatic
fluid overpressures in rocks and these include
rapid sediment deposition, thermal expansion of
fluids, tectonic compression, and the actual generation of low density oil and gas in the sourcerock.
Many of the factors discussed above also apply
to secondary petroleum migration. In the advanced

220

Figure 5.32 Actual measurementsof excess hydrostatic pressures(i.e. overpressure) in three drill
wells in the Orinoco basin and theanticipated fluid flow lines in thesequence of alternating shale andsandstone (after Hunt, 1979).

stages of catagenesis, however, temperatures andpressures are such that many hydrocarbons exist
close to their critical points such that density contrasts between liquid and vapor are minimized.
The segregation of oil from gas, therefore, mainlyreflects the decline in temperatures that accompanies migration of petroleum products into
reservoirs that are well away from the sites ofoil/gas generation. The progressive separation ofliquid and vapor, into a low density, buoyant gas
phase and a more dense, viscous oil, is a processthat generally accompanies secondary migration.
Entrapment of oil and gas
Primary migration of hydrocarbons typicallyoccurs over short distances (hundreds of metersor less) and is constrained by the proximity of
the first available aquifer to the source rock.
Secondary migration, by contrast, occurs overtens, and possibly even hundreds, of kilometers
and is only constrained by the presence of a trap
which prevents the petroleum from further flow
and allows it to accumulate. Hydrocarbon traps
are critical to the formation of viable oil andSEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

gas deposits and take the form of any geologicalfeature that either reduces permeability of thereservoir or provides a physical barrier that impedes the further migration of fluid. Evaporite
layers, for example, may be good hydrocarbon
traps as they are laterally extensive and have virtually zero permeability because of the ability of
halite to flow plastically at elevated temperatures. Many of the giant oil fields of the Arabian(Persian) Gulf are capped by evaporite sequences(Box 5.3), as are some of the North Sea gas deposits.
In general, however, structural features such as
faults and anticlines tend to be the dominanthydrocarbon traps. This section briefly outlines
some of the geological scenarios that represent
potential traps for hydrocarbon deposits in various parts of the world.
Figure 5.33 shows that there are basicallythree categories of trap site, namely stratigraphic,
structural, and a less common miscellaneousclass that includes hydrodynamic and asphalt
traps. Stratigraphic traps arise entirely from sedimentological features and can be represented
by features such as sediment pinch-out zones,
where permeability is reduced as one sedimentfacies grades into another, or unconformities
where the reservoir rock sub-outcrops against aunit of reduced flow capability (Figure 5.33a). The
Athabasca tar sands (see section 5.4.4 below) are
believed to have resulted from migration andentrapment of hydrocarbons along a major unconformity in the early Paleogene period. In
addition to sandstones, carbonate rocks represent
important reservoirs because the high solubility of minerals such as calcite promotes secondary porosity and migration of hydrocarbons.
Sedimentary features such as limestone “pinnacle”
reefs (Figure 5.33a), are important trap sites in carbonate sequences, especially in the giant Middle
East deposits.
Structural traps result from some form of
sediment deformation and generally provide physical barriers which prevent the continuation of
fluid flow along an aquifer. A fault, for example,
might juxtapose a reservoir sediment against a
shaley unit and, as long as the fault itself remains
impervious, will act as a barrage behind whichhydrocarbons will accumulate. Since most hydro-

297

carbons flow up-dip, the folding of a reservoir rockinto an anticline or dome-like feature also provides a very efficient trap site. The huge oilfieldat Prudhoe Bay in Alaska represents an example
where petroleum is trapped by a combination ofboth folding and faulting of the reservoir sediments. The scale of hydrocarbon accumulation inthis setting can be estimated by considering theclosure volume of the structure and also the spill
point, beyond which oil will continue its migration into another aquifer system (Figure 5.33b).
Other important structurally related trap sites
are linked to the deformation that accompanies
salt diapirism. When salt beds underlie more
dense rock strata the resultant gravitational instability causes the diapir to pierce the overlying
beds and terminate reservoir sediments creating
ideal trap sites for hydrocarbons (Figure 5.33b).
Such trap sites are of great importance in many
parts of the Arabian Gulf. Other trap sites of lesser
importance include sites where oil migration isdiverted by a strong flow of groundwater which,as explained previously, will water-wet the sediment pore spaces and impede the flow of hydrocarbons. These are known as hydrodynamic traps
and result from the shear stresses set up at theoil–water contact in aquifers where strong waterflow is occurring. Groundwater will also interact
chemically with reservoir oil, resulting in oxidation, biodegradation, and the formation of a bituminous asphalt layer at the interface. This layer,
known as an asphalt trap, will itself act as abarrier beneath which hydrocarbons may accumulate (Figure 5.33c).
5.4.3 Coalification processes
The majority of the world’s energy requirements
is still obtained from the burning of coal andlignite. The two largest producers of coal in theworld are China and the USA, although Australia
ranks top as an exporter of high quality coal.
As mentioned previously, most coal is derived
from Type III kerogen and generally represents thein situ accumulation of land-derived vegetation
subjected to alteration and compaction. Figure 5.26
shows that coal is derived from a peat precursor
and, with burial or coalification, is progressively

298PART 3 SEDIMENTARY/SURFICIAL PROCESSES(a) Stratigraphic traps

Carbonate “pinnacle” reef
(i)

(ii)

Unconformity
Shale

Dolomite
Oil

Pinch out

Calcareous shale
Sandstone
Shaleylimestone
Limestone(b) Structural traps
(i)

(ii)
Diapir

Fault

Shale

one

dst

San

Salt diapir

le

Sha

(iii)
Closure volume
Spill
point

Anticline
r
oi
rv ck
e
o
s r

Re

(c) Other traps

Oil

(i)

(ii)

Hydrodynamic

Asphalt
Hydrostatic
head

Water

Meteoric
water

Biodegradedoil/asphalt

Figure 5.33 Geological scenarios for hydrocarbon trap sites. (a) Stratigraphic traps represented by unconformities,
pinch-outs, and carbonate “pinnacle” reefs. (b) Structural traps represented by faults, diapiric features, andanticlinal or dome like structures. (c) Other features, such as hydrodynamic and asphalt traps (after Hunt, 1979;
Bjørlykke, 1989).

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

299

Fossil fuels: oil and gas – the Arabian (Persian) Gulf,
Middle East
The Arabian (or Persian) Gulf Basin contains well overhalf of the conventionally recoverable oil reserves of theworld, and also huge reserves of natural gas. This extraordinarily rich basin extends over a length of more than2000 km from Oman in the south to Syria and southeast
Turkey in the north (Figure 1). Production has been
derived from some 250 reservoirs in the region, of whichmore than 80% are in Jurassic–Cretaceous sediments.
The Middle East is unique for the size of its individual
deposits, with 14 fields having recoverable reserves inexcess of 10 billion barrels (Shannon and Naylor, 1989).
These include such giants as the Ghawar field in Saudi
Arabia and Burgan field in Kuwait (both with more than
70 billion barrels of oil) and the Rumaila and Kirkuk fields
in Iraq, each with around 15–20 billion barrels of oil(Tiratsoo, 1984). Many of the oil fields in the Persian Gulf
are also associated with huge gas resources. In addition,
however, some very large unassociated gas fields also
occur in the region, such as the North Dome field off
Qatar. These gas deposits occur in the Permian Khuff
limestone and are, therefore, older and underlie the main
oil deposits. Unassociated gas deposits in many otherparts of the world appear to be related to coal measuresof Carboniferous age. There are several reasons why the
Middle East is so well endowed in hydrocarbon resources
and these are briefly discussed below.
45°

The oil fields of the Arabian Gulf are located in late
Mesozoic to early Paleogene reservoir sediments that were
deposited on a continental shelf forming off the northeastern margin of the Arabian portion of Gondwana. A
significant proportion of these sediments are limestones,
indicating deposition at equatorial paleolatitudes in the
Tethyan seaway (North, 1985). As Pangea fragmented
the Arabian land mass became part of a downgoing slab
involved in subduction to the north and east. Fortunately,
subduction ceased in Pliocene times with the formationof the Zagros mountains in present day Iran, and thishydrocarbon-rich basin was preserved largely intact. It
is likely that other parts of the Tethyan seaway, probably
equally well endowed, were consumed by subductionin Cenozoic times, perhaps accounting for the fact thatthe western and eastern extremities of the seaway containlesser oil and gas deposits.
The tectonic and sedimentological regimes active in thisenvironment produced at least three major source rocksediments in the Arabian foreland, along the axis of thepresent day basin and in the orogenic belt of Iran. Source
rocks are all marine in origin and developed over the entire
extent of the Arabian platform. Oil and gas is hosted in avariety of different lithologies and the principal reservoirs
become progressively younger from southwest to northeast. The thick carbonate and sandstone reservoirs that

35°

55°

25°

N

TURKEY
KirkukIRAN
Ahwaz

Kharg/Darius

Rumaila

IRAQ

BurganKUWAIT

SYRIA
LEBA

Q
AT
A
Qatif R
DukhanGhawarUAE

Maghrib

OMAN

SAUDI
IRAN
ARABIA

NON

35°

Dividedzone

North dome

ISRAELJORDAN

0

200
km

35°

25°

45°

55°

Figure 1 Location and distribution of some of the major oil and gas fields in the Arabian Gulf (after Shannon
and Naylor, 1989).

300

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

formed in the Persian Gulf Basin are also characterized bya remarkable lack of intergranular cements and a high degree of primary porosity, making them very effective forthe migration and ultimate concentration of hydrocarbons.
This also contributes to very good recovery rates and thefact that very few wells need to be be drilled in order toexploit a given field. Periodic uplift and subaerial exposure also resulted in the deposition of numerous evaporitic
horizons that make good seals for oil and gas entrapment.
Sedimentation in the Jurassic and Cretaceous periods wascharacterized in general by a marine transgression over
the Arabian platform. Initially widespread carbonate sedimentation occurred and this was followed by increasingaridity and evaporate formation. In detail, sea level fluctuations and climate changes gave rise to complex interfingering of lithologies (Shannon and Naylor, 1989), as illustrated
in the environmental interpretations for the region during
mid-Jurassic and mid-Cretaceous times (Figure 2).
The structural development of the region was also
conducive to the formation of very large and effective
trap sites. Three types of structural traps are recognizedin the Arabian Gulf Basin. These include large, but gently
dipping, anticlinal warps which caused oil entrapment(a)

Mid-Jurassic

mainly in Jurassic limestones (e.g. Ghawar in Saudi
Arabia); diapiric salt dome structures arising from basement evaporitic sequences which deform Cretaceous andJurassic limestones (e.g. Burgan in Kuwait); and steeply
dipping anticlines in the foothills of the Pliocene mountain building episode in Iraq and Iran (e.g. Kirkuk in Iraq).
The presence of thick evaporitic sequences is considered
to be one of the most important features of the region
because the salt horizons are believed to have absorbed
much of the shortening that accompanied subduction
to the north and east of the basin. This had the effect ofpreserving the large gently anticlinal structural traps thatcharacterize the southwestern portions of the region andaccounts for the huge sizes (up to 100 km long) of many of
the oil fields (Shannon and Naylor, 1989). In addition to
the presence of salt, some of the other geological featuresof the Middle East that have made it so productive for oil
and gas include a long history of marine sedimentation
at relatively low paleolatitudes, numerous transgressive
and regressive sedimentary cycles resulting in a close
association between source and reservoir rocks, and theoccurrence of widespread seals, in the form of evaporitic
sequences, that prevent leakage of the oil and gas.

(b)

Mid-Cretaceous
Shallow carbonate
shelf

N
Intrashelf shale
basin

N

Deep carbonate
shelf

Shallow carbonate
shelfbasin Continental
slope
Zagros
crush
zone

Deep
clastic
shelf

Zagros
crush
zone
Mixed
carbonate
and clastic
shelf

Deep
marine0200
km

Mixed carbonate
and shale shelf
Alluvial
plain
0

200
km

Shallow clastic
shelf

Lower coastal plain

Figure 2 Interpretations of the evolving sedimentary environment in the region of the Arabian Gulf during(a) mid-Jurassic and (b) mid-Cretaceous times (after Shannon and Naylor, 1989).

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

WatersVegetation

MarineRhizophora

Brackish

301

Fresh

Rhizophora/
Mariscus Mariscus Waterlily

Marinetrangression
Figure 5.34 Generalized scheme
illustrating the nature of peat
stratigraphy in terms of vegetation
type and transgressive/regressive
effects in a swamp environment(after Spackman et al., 1976).

transformed to a series of products ranging fromlignite and sub-bituminous coal (known as “browncoals”), to bituminous coal and finally anthracite(the “hard coals”). The formation of peat is an
essential initial step in the process and involves
restrained biochemical degradation of plant matter without rampant oxidative and bacterial decay.
Coalification follows once peat is covered byoverburden and subjected to an increase in temperature with progressive burial.
Coal is a sedimentary rock that contains morethan 50% by weight of carbonaceous material andcan readily be burnt (McCabe, 1984). It forms
by the compaction of peat in an environment represented by a well vegetated land surface that issaturated with water. This environment is loosely
termed a “swamp” but should not necessarily
imply a tropical or equatorial climate as it is now
well known that many coals have accumulated incold climates at mid- to high latitudes. The greatest concentrations of peat are in fact accumulating in Russia (with about 60% of the world’s
resources; McCabe, 1984) at latitudes of around50–70° N. High rainfall is not an environmental
prerequisite, although abundant free-standing
water promotes plant growth and also submerges
dead vegetation, retarding the rate of decomposition. Furthermore, swampy environments often

Peatstratigraphy

contain waters that are both anaerobic and acidic,
which promotes the preservation of organic material by minimizing oxidation and destroying
bacteria. Tropical rain forests seldom form peat
accumulations because the high temperatures
promote rapid oxidation and decay of organic
material. The best conditions for peat accumulation reflect a balance between organic productivity and decay, with the added prerequisite of low
sediment input necessary to keep dilution of thepeat by clastic material to a minimum.
Coal characteristics
Coal is a heterogeneous sedimentary rock thatreflects both the different sedimentological regimes
within which peat forms and the varied vegetation types from which it is derived. Peat deposits
are known to have formed in many settings
including lacustrine, fluvial, deltaic, and beachrelated (Galloway and Hobday, 1983; McCabe,
1984; Selley, 1988). Peat and coal deposits, therefore, exhibit a wide range of shapes, and are alsocharacterized by the facies variations expected ofany sedimentary rock. Figure 5.34 illustrates thestratigraphic variations one might expect to see in
peat, as a function of vegetation type and fluctuating water levels.

302

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

Table 5.2 Coal classification in terms of maceral groups and lithotypes
Maceral group

Lithotype

Vitrinite (or huminite)

Cell walls of vascular plants

Exinite (or liptinite)

Spores, cuticles, waxes,
resins and algae
Oxidized and burnt (fusinite)
plant material

Inertinite

VitrainClarainDurainFusainCannel coal
Boghead coal

The vegetation shown in Figure 5.34 applies
specifically to the Everglades in Florida, wheremarine transgression results in lateral facies
changes in swamp environment and depositionof different peats that originate from a variety ofplant types. Coals that form by compaction of peat
will also, therefore, exhibit marked variations
in composition and vegetative make-up. This
characteristic has led to a rigorous classificationof coals, the main elements of which are shownin Table 5.2. In the same way that a rock is madeup of minerals, a coal is made up of “macerals,”
which are components of plant tissue, or theirdegraded products, altered during compaction anddiagenesis. The main maceral groups are vitrinite(also called huminite in brown coals), made up
woody material (branches, roots, leaves, etc.), andexinite (also called liptinite), which comprises
spores, cuticles, waxes, resins, and algae. A thirdgrouping is also recognized, termed inertinite,which comprises mainly the remains of oxidized
plant material and fungal remains. Each maceralgroup is further subdivided into individual macerals, but listing of these is beyond the scope of thisbook and can be obtained in more detailed texts,
such as Stach et al. (1982). Different combinations of maceral groups identify the coal lithotypes (i.e. analogous to a rock type) and these arealso listed in Table 5.2.
The majority of coals are humic (i.e. made upof macroscopic plant parts), are peat derived, andcan be subdivided into four lithotypes. These are
vitrain, clarain, durain, and fusain, whose maceralcompositions are shown in Table 5.2. The coallithotypes are important in determining its properties, in particular factors of economic import-

Vitrinite + <20% exinite
54
Layers of all three maceral groups
6 Humic coals
Mainly inertinite and exinite
4
Fusinite
7
Spore-dominated exinite# Sapropelic$ coals
Algae-dominated exinite

ance such as calorific value and liquefaction
potential. In general, the highly reflective, morecombustive coals are made up of vitrain, whereas
the duller, less combustive coals comprise essentially durains. It should be noted, however, thatcoal lithotypes should not be confused with coalrank, which separates low calorific value browncoals from high heat capacity hard coals essentially on the basis of degree of compaction orburial depth.
A small proportion of coals are sapropelic andformed largely from microscopic plant remains
such as spores and algae (exinite). The sapropelic
lithotypes, also referred to as cannel and bogheadcoals (Table 5.2), did not form from peat. They
can be found overlying humic coals, forming in
muddy environments as the swamp was flooded
and buried, and occasionally are linked to the formation of oil shales (see section 5.4.4 below).
In addition to the maceral-based classification,coals can also be subdivided in terms of theirchemical composition, specifically their carbon,hydrogen, and oxygen contents. Figure 5.35 is aternary plot of normalized C, H, and O in which
the compositional fields of peat and the coals areshown in relation to those for oil and oil shales.
As peat is progressively compacted and coalification commences, the carbon content increases
noticeably, whereas hydrogen remains fairly constant and oxygen decreases. Hydrogen only starts
to diminish noticeably once the hard coals start todevelop. The chemical changes that accompany
coalification are accompanied by the productionof CO2 and CH4 gas. The source rocks for some of
the gas fields in the North Sea are believed to beunderlying Westphalian coal beds.

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

303

C(100%)

Anthracite
Bituminous
Oil

Subbituminous
a
Co
ca
lifi
tio

le

ha

s
Oil

n

LigniteFigure 5.35 Ternary C–H–O plot
showing the compositional trend
accompanying coalification and thecomparison between coals and oil
and oil shale compositions (after
Forsman and Hunt, 1958).

Peat
H (30%)

Another important property of coal, but one
which is unrelated to its organic or maceralcharacteristics, is ash content. As peat forms it
inevitably incorporates inorganic, clastic detritus
from a variety of sources. During coalification
this detritus is diagenetically altered to form avariety of authigenic mineral components suchas quartz, carbonate, sulfide, and clay minerals.
These components remain behind to make up
its ash content once the coal is combusted orprocessed. Low ash content coals are obviously
economically advantageous as this parametertranslates favorably into higher calorific values,
better grindability, and less deleterious environmental effects such as sulfur emissions.
A note concerning formation of economically
viable coals
The formation of economically viable coals isdetermined by processes active during both initialpeat accumulation and later coalification. In thelatter case bituminous coals and anthracite are
the high rank products of thermal maturation andburial, and are characterized by the best calorificvalues and lowest volatile contents. Progressive

O (30%)

Figure 5.36 Bituminous coal seam from the Witbank
area, South Africa, hosted in carbonaceous shales anddeltaic sandstones of the Permian Ecca Group (photo iscourtesy of Bruce Cairncross).

compaction will transform most humic coals tothese products, which, if extensive and thickenough, will likely be economically viable. In
addition simply to burial, some anthracites arealso known to be products of tectonic deformation or devolatilization by nearby igneous intrusions, a good example of the former being thePennsylvanian coals of the USA, which grade304

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

laterally into anthracites as they approach theAppalachian fold belt. Controls on the formationof economically viable coals during initial peat
accumulation are, however, more difficult toconstrain. Deltas have traditionally been regarded
as the optimal environment for development of
viable coal seams, although this concept is now
questioned because the high clastic sediment
input in this setting generally results in peat
dilution and formation of less attractive, highash coals (McCabe, 1984). Environmental models
now favor the notion that optimal peat accumulation represents a discrete event that was
temporally distinct from those of underlying
and overlying clastic deposition. The notion ofan environment where peat accumulation is
optimized, but not diluted by clastic input, is best
accommodated by the concept of raised and/or
floating swamps. Floating swamps refer to buoyant peats that develop on the surface of shallow
lakes to eventually accumulate and accrete on
the lake margins (Figure 5.37). Raised swamps
are seen as the final product in a continuum of
processes which extend from initial developmentof shallow lakes with floating peat, followed byformation of well vegetated, low-lying swamps,
and eventually to raised swamps as accretion oforganic material eventually sculpts a landscapethat is elevated with respect to the sediment substrate. The latter environment is obviously one inwhich clastic sediment input is minimized. Thisevolutionary sequence, illustrated in Figure 5.37,
serves as a model for the depositional environment in which optimal formation of thick peat,and subsequently economic coal horizons, mightform. Raised swamps can form in deltaic environments, but the above model applies specificallyto those settings in which sediment input isminimized and the resultant coals are low-ash incharacter. Examples of raised swamps include
the Klang-Langat delta in Malaysia and areas ofSarawak where peat accumulation of more than
2 mm yr−1 is occurring (McCabe, 1984).
5.4.4 Oil shales and tar sands
Although organic matter accumulates in mostsediments, it is the shales or mud rocks whichcontain the highest hydrocarbon contents. Mudrocks that are rich in organic matter are called
sapropelites, but those that yield free oil on heating are specifically referred to as “oil shales.” Oilshales typically form in lacustrine settings where
algae settle out of the aerated upper layers andaccumulate in the anaerobic bottom muds where
they are preserved from rapid oxidative decay. In
this sense they are akin to the sapropelic cannel orboghead coals discussed above. Since algal (TypeI) kerogen is oil prone, these shales represent a
resource of crude oil that has not as yet matured,or passed through the oil window. A major challenge for the future is to develop the technology toheat these rocks up to the temperatures required
to efficiently extract oil, and to do so in an economically viable and environmentally responsible way. The best known oil shale is perhaps
the Green River Formation which developed inan Eocene, shallow-lacustrine or sabkha setting,in parts of Utah, Colorado, and Wyoming, USA.
Despite an enormous potential capacity to produce
oil (grades of up to 240 liters of oil per ton of rockare reported), production from the Green River
shale resource has, thus far, proven non-viable.
Successful processing of oil shale has neverthelessbeen achieved in countries such as Estonia,
Russia, and China. Present day oil shale formation is taking place in Australian saline lakes
where the alga Botryococcus braunii is accumulating to form a type of sapropelic peat known as
coorongite. Likewise, the torbanites of Scotland
formed in lacustrine settings from accumulationof a Carboniferous variety of Botryococcus, as did
the Permian oil shales of Tasmania (tasmanites).
Tar sands are often confused with oil shales,
although they are in fact quite different. Tar sands
are sandstones within which bitumen or asphalt
occurs. Bitumen is a solid or semi-solid hydrocarbon either derived by normal degradation of crudeoils, or formed directly as a high molecular weighthydrocarbon without ever having been a light oil.
In certain instances, tar sands that are near thesurface can be extracted and processed to produce
oil and a range of other useful hydrocarbon byproducts. The best known of these is the Athabasca
tar sand of northern Alberta, Canada. As with oil
shales, though, technological and environmental

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5(I)

Floating swamp on lake margins(II)

Low-lying swamp,
diverse and luxuriant flora

305(III) Slightly raised swamp,
restricted flora(IV) Extensive raised swamp,
restricted and stunted flora
Figure 5.37 Environmental
model for optimal peat
accumulation, in terms of anevolutionary development ofswamp types which leads
ultimately to the formation ofraised swamps in which extensive,
thick peats, devoid of major
sediment input, form (after
Romanov, 1968; McCabe, 1984).

difficulties provide major obstacles to more widespread production from this resource at present.
Both tar sands and oil shales, however, clearly
represent significant petroleum resources for the
future.

Lakes

Vertically zoned peat

5.4.5 Gas hydrates
A fairly recent discovery in the permafrost regionsof the world, as well as in ocean sediments sampled during the Deep Sea Drilling Program, is thePART 3 SEDIMENTARY/SURFICIAL PROCESSES

in sediment under about 500 m of water and theprogressive increase in thickness of the hydratestability zone with depth. Under 3 km of water amethane hydrate zone almost 1000 m thick could
form if the ingredients were available. This particular profile was calculated for a geothermal
gradient of 27 °C km−1 but would be thinner if
the sediment profile was subjected to higher geothermal gradients. These considerations, even if
viewed conservatively, point to the enormousvolumes of gaseous hydrocarbons that potentiallycould be locked up in the ocean sediments.

50

Methanegas + ice

100

1

5
10
Methanegas + water
ra

50

te

500

–g
as

–10

0
10
20
Temperature (°C)

a ry

10 000nd

Methanehydrate +water + gas

5000

100

ou

Methanehydrate +ice + gas

eb
as
ph

1000

Pressure (atmospheres)

10

Ice–water phase boundary

(a)

d
Hy

presence of vast resources of hydrocarbon locked
up as frozen gas hydrates. Gas hydrates are crystalline aqueous compounds that form at low
temperatures when the ice lattice expands toaccommodate a variety of gaseous molecules, themost important of which, for the purposes of this
discussion, is methane (CH4). A methane hydrate
forms when sufficient CH4 is present in water, at
the appropriate pressure and temperature, to form
a solid compound with the ideal composition
CH4.53/4H2O (i.e. the unit cell comprises 46 watermolecules with up to 8 methane molecules).
Other gases, including CO2, H2S, and C2H6, also
form gas hydrates with a similar structure. On a
global scale the volumes of natural gas present ingas hydrate reservoirs are estimated to be a factorof two times greater than the total fossil fuelreserve to present (Kennicutt et al., 1993). Methane
hydrates clearly, therefore, represent an enormous energy resource for the future. They are
also, however, an object of concern in that global
warming might accelerate the natural release ofCH4 by melting gas hydrates and further accentuate the influences of greenhouse gas accumulation on climate change.
It is of interest to consider the conditions
under which methane hydrates form in the ocean
basins of the world. Figure 5.38a shows a simplified phase diagram with the H2O ice–water andCH4 hydrate–gas boundaries marked. If sufficientmethane and pore water are available methane
hydrates will stabilize in ocean floor sediment
where temperatures are around 3–4 °C at pressuresof about 50 atmospheres (equivalent to about 500
meters water depth). Gas hydrates will melt ifeither the temperature increases or the pressure
decreases. The thickness of the methane hydratestability zone in oceanic sediments will dependon the pressure (an increase in which has the effectof stabilizing hydrates to lower temperatures) andthe geothermal gradient, the latter dictating therate at which sediment is heated with progressive
burial. For a constant geothermal gradient the
thickness of the methane hydrate layer will
increase directly as a function of water depth.
Figure 5.38b shows this graphically for a typicalocean-sediment profile on the continental shelf. It
illustrates the initiation of gas hydrate formation

Depth (m)

306

500

30

1000
40

Figure 5.38 (a) Phase diagram illustrating the regionsof gas hydrate stability under most natural conditionsin the near-surface (after Kvenvolden and McMenamin,1980). (b) Profile across a typical ocean–sedimentinterface in a continental margin setting, showing the
progressive increase in the width of the gas hydratestability zone in the ocean sediment with increasing
depth of sea water (after Kvenvolden, 1988).

SEDIMENTARY ORE-FORMING PROCESSES CHAPTER 5

307

(b)

om

wa

ter
s

Sea level

7°C

fb

ott

500

18°C
13°C

re
o

0

ra
tu

4°C
10°

Te
m

pe

1000

Depth (m)1500

Sediment–
water interface

2000
3°C
2500

2°C
1.5°C

3000

0

0
10°

Depth of gas
hydrate stability
zone

20°

10°

30°

20°

40°C

3°C

0

20°

10°

30°

20°

40°C

30°
40°C

Temperature
profiles in
sediment(geothermal
gradient =27°C km–1)

3500

4000

4500

Gas hydrate zone under the ocean

30°
40°C

Figure 5.38 (cont’d)

Sedimentation is a fundamental geological process that takes place over much of the Earth’s surface, in response to the pattern of global tectonic
cycles. A number of very important mineral andfossil fuel commodities are concentrated during
the formation of sedimentary rocks. Placer processes are important during clastic sedimentationand form by the sorting of light from heavy particles. A number of hydrodynamic mechanisms,
such as settling, entrainment, shear sorting, andtransport sorting, are responsible for the concentration of different commodities in a variety of
sedimentary micro- and meso-environments. At alarger scale, placer deposits form mainly in fluvial
and beach-related environments and include concentrations of gold, diamonds, tin, zirconium, andtitanium. Chemical sedimentation, where dissolved components precipitate out of solution
from brines, also gives rise to a wide range ofimportant natural resources. Precipitation of ferric iron from sea water is one of the principal
mechanisms for the formation of iron ore deposits.
Banded iron-formations formed in a variety ofoceanic settings mainly in the Paleoproterozoic
and prior to the onset of significant levels of oxygen in the atmosphere. Ironstone deposits tend to308

PART 3 SEDIMENTARY/SURFICIAL PROCESSES

form mainly in Phanerozoic times. Other types ofmetal concentration linked to syn-sedimentaryand early diagenetic chemical processes occur incarbonaceous black shales (Ni, Co, Cr, Cu, Mn,
Zn, Ag, etc.) and in Mn nodules on the ocean floor.
Accumulation of phosphorus and the developmentof phosphorites in ocean settings is a process thatis linked both to direct chemical precipitationfrom sea water and to biological mediation. The
upwelling of cold currents onto the continental
shelf and biological productivity are processes
implicated in the formation of most of the world’s
important phosphate deposits, so important for
fertilizer production and global food production.
Chemical sediments formed by high rates of
evaporation of sea water and lacustrine brines
host most of the world’s Na, K, borate, nitrate,and sulfate resources, also important to the agricultural industry.
Fossil fuels, in the form of oil, natural gas, andcoal, are the most important and valuable of all

natural resources as they provide the world withmost of its combustion-derived energy. Oil andnatural gas are derived in the initial stages bybiodegradation of largely planktonic organisms inmarine environments to form oil-prone kerogenand biogenic gas. Further burial results in subsequent chemical and thermal modification of kerogen to produce wet gas, and a liquid condensate
in the “oil window,” typically at burial depths of
3000–4000 m and temperatures of 100–150 °C.
Further burial results only in the production of
more dry gas. Burial of vascular land plants gives
rise initially to the formation of peat and the
production of mainly dry gas. Further compactionof peat gives rise to coals, the rank of which is afunction of progressive burial. Other forms offossil fuel, important for the future but largely
unexploited at present for technological and economic reasons, include oil shales and tar sands, aswell as methane hydrates in permafrost regions
and ocean floor sediments.

A huge volume of literature is available in the fields of
sedimentology and fossil fuels, but less with respect tochemical sedimentation and metallogeny. The following books and reviews are subdivided in accordancewith the main topics covered in this chapter.

Parnell, J., Ye, L. and Chen, C. (1990) Sediment-hosted
Mineral Deposits. Special Publication, 11, International
Association of Sedimentologists. Oxford: Blackwell
Scientific Publications, 227 pp.
Young, T.P. and Gordon Taylor, W.E. (1989) Phanerozoic
Ironstones. Special Publication 46. London: The Geological Society of London, 257 pp.

Sedimentological and placer processes
Allen, P.A. (1997) Earth Surface Processes. Oxford:
Blackwell Science, 404 pp.
Leeder, M. (1999) Sedimentology and Sedimentary
Basins: From Turbulence to Tectonics. Oxford: Blackwell Science, 592 pp.
Slingerland, R. and Smith, N.D. (1986) Occurrence andformation of water-laid placers. Annual Reviews ofEarth and Planetary Sciences, 14, 113–47.
Chemical sedimentation and ore formation
Mel’nik, Y.P. (1982) Precambrian Banded Iron Formations. New York: Elsevier, 282 pp.
Melvin, J.L. (1991) Evaporites, Petroleum and Mineral
Resources. Developments in Sedimentology, 50. New
York: Elsevier, Chapter 4.

Fossil fuels
Engel, M.H. and Mack, S.A. (1993) Organic Geochemistry: Principles and Applications. New York: Plenum
Press, 861 pp.
Glennie, K.W. (1998) Petroleum Geology of the North
Sea: Basic Concepts and Recent Advances. Oxford:
Blackwell Science, 636 pp.
North, F.K. (1985) Petroleum Geology. London: Allen &Unwin, 607 pp.
Selley, R.C. (1998) Elements of Petroleum Geology, 2ndedn. New York: Academic Press, 470 pp.
Thomas, L. (1992) Handbook of Practical Coal Geology.
New York: John Wiley.
Ward, C.R. (1984) Coal Geology and Coal Technology.
Oxford: Blackwell Science, 345 pp.

Global Tectonics andMetallogeny

Ore deposits in a global
tectonic context

PATTERNS IN THE DISTRIBUTION OF MINERAL DEPOSITSCONTINENTAL GROWTH RATESCRUSTAL EVOLUTION AND METALLOGENESIS
evolution of the hydrosphere and atmosphereglobal heat production
mantle convection
continental freeboard and eustatic sea-levelchanges
crustal composition and the Archean–Proterozoictransition
METALLOGENY THROUGH TIME
the Archean Eon
the Proterozoic Eon
the Phanerozoic Eon
PLATE TECTONIC SETTINGS AND ORE DEPOSITS – A SUMMARY

6.1 INTRODUCTION
Most of the world’s great mineral deposits are
the products of a fortuitous superposition ofgeological processes that resulted in anomalous concentration of ores, often over an extended
period of time. The preceding chapters have shownthat the formation of ore deposits is, nevertheless, related to much the same sort of processes
that gave rise to the formation of normal igneous,sedimentary, and metamorphic rocks in the
Earth’s crust. The fact that a close relationshipexists between rock-forming and ore-forming processes means that metallogeny must be relevant
to understanding the nature of crustal evolution

through geological time (Barley and Groves, 1992;
Windley, 1995). Conversely, crust-forming processes and the global plate tectonic paradigm havebecome indispensable to the broader understanding of how ore deposits form and no modern economic geologist can practice successfully without
an appreciation of the history and evolution of
the continents and oceans. Consequently, thischapter summarizes some of the recent thinking
that relates the formation of ore deposits to global
tectonics and the evolution of the continents.
There are essentially two ways of doing this. One
is to chart continental evolution and place ore
deposits into a secular and tectonic framework(e.g. Meyer, 1981, 1988; Veizer et al., 1989; Barley
and Groves, 1992). The other is to empiricallydescribe ore deposits in the context of the tectonic
environment and host rocks in which they occur(e.g. Mitchell and Garson, 1981; Hutchison, 1983;
Sawkins, 1990). The latter approach has been
covered in considerable detail in Sawkins’s (1990)
Metal Deposits in Relation to Plate Tectonics,
and there is little point in summarizing this extensive compilation here. The former approach isa more difficult undertaking and requires a thorough knowledge of the evolution of continents
with time, a topic that becomes progressively lesswell constrained – and more controversial – thefurther back in Earth history one goes. There are,however, significant benefits to be gained by understanding the intricate processes of ore formationin terms of the secular evolution of the Earth’s
crust and atmosphere. This is the approach thathas been adopted below, even though it is necessarily cursory and in places speculative.

312PART 4 GLOBAL TECTONICS AND METALLOGENY

6.2 PATTERNS IN THE DISTRIBUTION OF MINERAL DEPOSITS
It has long been known that the different stylesof mineralization are not randomly distributed,
either in time or in space, and that broad patterns
exist when relating deposit types to crustal
evolution and global tectonic setting. The originalcompilations of ore deposit type as a function ofgeological time by Meyer (1981, 1988) have been
re-examined by Veizer et al. (1989) and Barley andGroves (1992) and reveal consistent patterns.
Figure 6.1 distinguishes between the secular distribution of metal deposits formed in orogenic
settings and those formed in anorogenic environments and in continental basins. Although thiscompilation is a “first-order” generalization, it isnevertheless clear that deposit types classified
as orogenic (including lode-gold or orogenic goldores, volcanogenic massive base metal sulfides,
and the porphyry-epithermal family of base andprecious metal deposits) occur either in the latestages of the Archean Eon (between 3000 and2500 Ma) or in the Phanerozoic Eon (between
540 Ma and present). The Proterozoic Eon, between
2500 and 540 Ma, contains fewer of these typesof deposits. By contrast, the majority of metal
deposits that are associated with so-called “anorogenic” magmatism (such as the anorthositehosted Ti deposits and the Olympic Dam andKiruna type Fe oxide–Cu–Au ores), as well as sediment hosted ores (such as the SEDEX type Pb–Zn
deposits and the stratiform “red-bed” type Cu
deposits), are dominantly hosted in Proterozoic
rocks (Figure 6.1). An explanation for this pattern,
as discussed below (see sections 6.4.3 and 6.5 inparticular), is that ore formation might be closely
linked to the so-called “supercontinent cycle,”which describes the broad scale amalgamationand dispersal of the major continental fragmentswith time. It is, for example, evident that peaksin the production of anorogenic and continental
sediment-hosted metal deposits (Figure 6.1, lower
histogram) coincide with the periods of crustal
stability and the existence of large, stable continental amalgamations such as Nena in the
Mesoproterozoic, Rodinia in the Neoproterozoic
and Pangea in the early Mesozoic. Periods oflarge-scale continental fragmentation, by contrast,

appear to be orogenically more active and give rise
to a different suite of ore deposit types.
In a study designed to examine the rates atwhich ore deposits are recycled or dispersed (either
by erosive or tectonic dismembering) in the Earth’s
crust, Veizer et al. (1989) have produced a very
similar pattern of ore deposit distribution withtime. In this compilation hydrothermal andvolcano-sedimentary base–precious metal deposits
formed mainly in late Archean and Phanerozoic
times, whereas chemical–sedimentary and ultramafic deposits reflect concentration mechanisms that took place in the mid-Proterozoic.
The reasons behind this pattern are likely to becomplex and multifaceted and are discussed in
more detail below. It should be emphasized at thisstage, however, that the relationships observed in
Figure 6.1 could also be due, at least in part, to apreservation factor. The existence of numerous
arc-related orogenic deposits in the late Archean,for example, reflects the preservation of greenstone belts in stable shield areas of the world. In
Proterozoic times, by contrast, the preponderanceof collisional orogenies might have uplifted anderoded similar mineralized arcs so that many of
the related near-surface ores were destroyed. The
products of young Cenozoic orogenies are alsopreserved because many of them have not yet
suffered collision and uplift and are, therefore,
still preserved.

6.3 CONTINENTAL GROWTH RATESThe relationships illustrated in Figure 6.1 indicatethat estimates of the rates of continental growthwith time are important to any consideration of
the patterns of continental assembly and dispersal
and, therefore, metallogeny. Such estimates areproblematic because of the difficulties in calculating the rates of recycling (or destruction) of continental material relative to new crust formation.
Growth rate estimates vary enormously, and range
from one scenario, where the majority of continental crust formed rapidly in the early Archean (curve A in Figure 6.2a), to another wherecontinents grew exponentially with time, withthe most substantial contributions coming afterORE DEPOSITS IN A GLOBAL TECTONIC CONTEXT CHAPTER 6

Orogenic

Ma
3000

Volcanogenic massive sulfide
(VMS)

313

2000

1000

Present

1000

Present

Cyprus-type

Abitibi-type

Kuroko-type

Lode-gold veins

Porphyry
deposits

Gold and uranium in
conglomerates

Porphyry copper

Porphyry molybdenum
Anorogenic and continental sediment-hosted
3000
2000

Ma

Kiruna-type
Olympic Dam-type
Ilmenite-anorthosite
Lead-zinc in clastic
sediments

Sediment-hosted

Figure 6.1 Distribution of ore
deposit styles, classified according
to orogenic (upper set of histograms)
and anorogenic/continental
sediment-hosted metal deposits(lower set of histograms) as afunction of time. The length ofeach histogram bar is an estimateof the proportion of ore formed overa 50 million year interval relative tothe global resource for that deposit
type. Periods of supercontinent
amalgamation are shown as Pangea,
Rodinia, and Nena, with decreasinglevels of confidence further back intime. The diagram is originally after
Meyer (1988) and modified afterBarley and Groves (1992).

Anorogenic
intrusions

Uranium inweathered profiles

Lead-zinc in carbonates(Mississippi Valley type)

Copper in clastic
sediments

Nena

Rodinia Pangea

314PART 4 GLOBAL TECTONICS AND METALLOGENY

(b)

(a)
0.15
Fraction of crust

0.8

M

0.6

&

B

0.4
B
0.2
0.0

Fraction of crust produced per 200 Myr

A

1.0

0

1000
2000
3000
Time before present (Ma)

4000

0.10

0.05

0.00
0

1000

2000

3000

4000

Time before present (Ma)
Figure 6.2 (a) Different models for the rates of continental growth. Curves A and B contrast considerations of earlyArchean continental growth with an exponential growth rate model in which the bulk of the continents grew
subsequent to about 1500 Ma. The curve M&B illustrates the broadly linear but episodic continental growth rate
curve suggested by McCulloch and Bennett (1994). (b) Histogram showing the estimated rates of continental growth
averaged over 200 million year intervals based on a global compilation of Sm–Nd model ages (after McCulloch andBennett, 1994).

1.5 Ga (curve B in Figure 6.2a). The most likely
situation probably reflects an intermediate casewhere approximately linear, but episodic, crustal
growth took place. McCulloch and Bennett (1994)
supported this scenario on the basis of Sm–Ndmodel and U–Pb zircon age distribution data forthe continental crust, showing that the propor-

tion of crustal recycling relative to growth isapproximately constant. The model, illustrated in
Figure 6.2b, suggests discrete episodes of enhanced
crustal growth. It envisages that 50–60% of the
Earth’s continental crust had been produced by
the end of the Archean, but that this occurredin two major episodes at 3600–3500 Ma and at

ORE DEPOSITS IN A GLOBAL TECTONIC CONTEXT CHAPTER 6

2800–2600 Ma, with the latter being the mostimportant period of crustal growth in Earth history. Two important episodes of crust formation
were also considered to have taken place in theProterozoic, one at 2000–1800 Ma and the otherat around 1100–1000 Ma.
It is interesting to note that the McCulloch
and Bennett model suggests that only a relatively
small proportion of new crust was added to the
Earth’s surface in the Phanerozoic Eon. This appears
to be somewhat at odds with the observation that
this period of Earth history was particularly productive in terms of ore deposit formation. This
supports the notion of preferential preservationof younger ore-forming environments comparedto older sequences which become either progressively more eroded, or buried and metamorphosed,with time. The episodic continental growthmodel described in Figure 6.2 is reasonably consistent with the pattern of metallogenic evolution
described here and is used as the basis for a closer
examination of the relationships between crustal
evolution and ore deposit formation.

6.4 CRUSTAL EVOLUTION AND METALLOGENESIS
The review by Barley and Groves (1992) suggested
that there were three principal factors which
affected the pattern of global metallogeny. These
are the evolution of the hydrosphere–atmosphere,
the secular decrease in global heat production, andlong-term global tectonic trends. These major factors, together with a number of related parameters,
are discussed in more detail below, with an indication of how and why they influence the temporaland geographic distribution of ore deposit types.
6.4.1 Evolution of the hydrosphere andatmosphere
The changing budget of O2, and to a lesser degree
CO2, in the Earth’s atmosphere with time hada significant role to play in the formation ofore deposits, especially those related to redox
processes and to the weathering and erosion ofcontinental crust. During the Archean it is generally held that the atmosphere contained very

315

little free molecular oxygen (although the actual
amounts are still strongly debated) and what little
did exist was the result of inorganic photodissociation of water vapor. A reduced atmosphere in
the Archean helps to explain many of the featuresof ore formation at that time, including thewidespread mobility of Fe2+ and development of
Algoma and Superior type banded iron-formations,and the preservation of detrital grains of uraninite
and pyrite in sedimentary sequences such as thoseof the Witwatersrand and Huronian basins. The
transition from the Archean to the Proterozoic
Eons, at 2500 Ma, broadly coincides with the firstmajor increase in atmospheric oxygen (Figure 6.3).
This event is possibly related to the evolution ofprimitive life and the development of cyanobacteria capable of photosynthetically producing oxygen. The partial pressures of O2 in the atmospherecan be estimated from the mobility of Fe2+ inpaleosols (the remnants of weathering horizons
or soils preserved in the rock record) and suggestthat a sharp increase in the partial pressure of oxygen (pO2) occurred at around 2200 Ma (Rye andHolland, 1998). The interval between about 2500
and 2000 Ma is known as the period of “oxyatmoinversion” and is generally believed to coincide with a significant rise in atmospheric oxygen
levels. It is also the period after which development of Superior type banded iron-formations andrelated bedded manganese deposits ceased. Another significant increase in atmospheric oxygen
levels coincided with the end of the Proterozoic
Eon, when macroscopic, multicellular life forms(the Metazoan fauna which required oxygenatedrespiration) proliferated. In addition, the earlyPhanerozoic is also the time when vascular plants
evolved and this, too, would have contributed
to the increase in atmospheric oxygen levels(Figure 6.3). In contrast with the net addition of
O2 in the atmosphere, CO2 has progressively
decreased with time. The initially high levels of
CO2 in the atmosphere were a product of extensive volcanism and outgassing during theArchean, but levels have dropped steadily, but
episodically, as a result of carbon and carbonate
deposition in sediments. The evolution of the
Earth’s atmosphere and, in particular, the global
O2 and CO2 trends are essentially non-recurrent

316PART 4 GLOBAL TECTONICS AND METALLOGENY

1.0

Metazoa
Oxyatmoinversion

Atmospheric pO2 (atm)

0.1

0.01

0.001

0.00010.00001

0.000001
2.8

2.7

2.6

2.5

2.4

2.3
2.2
2.1
2.0
Time before present (Ga)

1.9

1.8

1.1

0.45

Figure 6.3 Estimates of the episodic rise of pO2 over the past 3 Gyr, constrained by data derived from the mobilityof Fe in paleosols (fossilized soil horizons). Bars and arrows show actual estimates of the ages and pO2 for a numberof paleosol horizons worldwide, whereas the solid curve is the best estimate of the atmospheric pO2 trajectory withtime (after Rye and Holland, 1998).

and irreversible and contribute to the time-specificoccurrence of ore deposit types such as bandediron-formations and Witwatersrand type placers.
6.4.2 Secular decrease in global heat production
The production of heat from the Earth, mainly
from the decay of long-lived isotopes such as U,
Th, and K, was significantly greater (by a factorof two to three times) in the Archean than it istoday. This pattern, like that for atmosphericevolution, is also non-recurrent and irreversible.
Figure 6.4 shows the decline in global heatproduction with time and a corresponding crustal
growth curve similar to that discussed above.
There are several consequences for crust formation related to the secular decrease in heat production. One of the most obvious is that the Archeanmantle was hotter (by about 200–300 °C; Windley,
1995) than at present and this could have led to anincrease in the volume of melting at this time. A
higher degree of partial melting could account for

production of the high-Mg komatiites that typify
many Archean greenstone belts and which are
relatively rare in younger oceanic settings (seeFigure 6.4), although other explanations are also
feasible. Higher heat production thus provides an
explanation for the preferential occurrence, in theArchean, of komatiite-related magmatic Ni–Cu
sulfide deposits, such as those at Kambalda,
Western Australia and in Zimbabwe. Anotherconsequence of higher heat flow in the Archean isthe suggestion that convective overturn in the
mantle would have been more rapid, leading toshorter ocean crust residence times and more
rapid plate motion (Windley, 1995). Archean
oceanic crust is also considered to have been
thicker than in later periods, which is supported
by suggestions that ophiolites greater than 1000
years old are more than double the thickness of
their younger equivalents (Moores, 1993). As
another consequence, Archean oceans may havebeen more shallow than those at present, resulting in a reduced continental freeboard.

ORE DEPOSITS IN A GLOBAL TECTONIC CONTEXT CHAPTER 6

317

Basalts

75

Cr u
sta
lg

ro
wt

h

3

50

m
Ko
iit
at

e

s
Norites

2

1

4

3

2

25

1

% Continental growth

4

at
He

Figure 6.4 Schematic diagram
showing the exponential decline
in global heat production withtime and the matching decline inproduction of komatiite volcanism.
Also shown is a crustal growth
curve and the period of enhanced
mafic or “norite” magmatism in thelate Archean and early Proterozoic(after Hall and Hughes, 1993).

Relative heat production (no units)

50Time before present (Ga)6.4.3 Long-term global tectonic trends

Rates of continental growth

The broad-scale amalgamation and dispersal ofcontinental fragments with time, also referred
to as the “supercontinent cycle,” is an allencompassing feature of Earth history that has
implications for virtually every aspect of global
evolution, including metallogeny. As seen in
Figure 6.1, broad temporal trends in ore formationreflect the contrast between episodes of orogeny
and periods of stability. In detail, however, it is
the supercontinent cycle that controls the manyother factors more directly related to secular metallogenic trends. Unlike the patterns imposed byatmospheric evolution and heat production, thesupercontinent cycle is recurrent and cyclical inthat continental fragments have amalgamatedand dispersed several times through Earth history.
This cyclicity helps explain why some ore deposit
types, or at least the processes that give rise tothem, are recognizably repetitive despite the preservation bias of ore deposits in younger rocks. A
detailed understanding of how the supercontinent
cycle can be used in the study of ore deposits is
a complex issue and likely to remain a profitablearea of future research. Some of the concepts thatare relevant to both global tectonics and the pattern of crustal/metallogenic evolution are briefly
mentioned below.

The growth of continental crust, whether it beepisodic or linear, early or late, will affect global
metallogeny simply because the majority of
viable ore deposits are located in or on continental
crust. Periods of rapid continental growth reflect
episodes of more intense magmatic activity andtectonism, and this, in turn, promotes the formation of magmatic, magmatic-hydrothermal, andhydrothermal ores. The majority of the world’s
gold deposits, for example, are related to events ineither the late Archean or Mesozoic–Cenozoic(Figure 6.1), both periods of major orogenesis andcrust formation (Figure 6.2b).
The nature of mantle convection
Models for mantle convection suggest two modesof occurrence (Figure 6.5), which may not be
mutually exclusive and could operate at the sametime. One model suggests overturn of material in
two discrete layers with little or no mass transfer
between them, and the other envisages that theentire mantle overturns due to the major downwelling of cold lithospheric slab material withconcomitant up-rise of large magmatic plumes(Stein and Hofmann, 1994). The first scheme(Figure 6.5a) is considered to represent the situation

318

PART 4 GLOBAL TECTONICS AND METALLOGENY(a)

WILSON CYCLES(< 500 Ma)
Arc

(b)

Mid-ocean ridge

MOMO EPISODES(500–1000 Ma?)Arc
Plume heads

Upper mantle
660 kmPlumesLower mantle

Deep penetrating
subducted slab
2900 kmFigure 6.5 Two contrasting models for mantle circulation (after Stein and Hofmann, 1994). (a) Mantle convection
in two discrete layers with little or no mass transfer from one to the other, corresponding to the conventional
plate tectonic regime in which crustal growth occurs in accordance with the Wilson cycle. (b) Deeply penetrative
mantle overturn involving the downwelling of cold lithosphere into the lower mantle and the associated upwellingof major plumes. These MOMO (mantle overturn and major orogeny) episodes are believed to replenish thedepleted upper mantle chemically and also to coincide with the major periods of orogenic activity and enhanced
crustal growth.

during which continental growth occurs mainly
by ocean subduction, arc construction, and accretion. In this model global tectonic patterns arereflected in terms of the Wilson cycle, in which
oceans open and close in ≤ 500 million year periods. The other mode, called the MOMO (Mantle
Overturn and Major Orogeny) model, involves
significant exchange of material between lower
and upper mantle and, accordingly, higher rates ofcrust formation (Figure 6.5b). This mode is considered to operate at longer periodicities than the
Wilson cycle and produces magmatic activity thatcoincides with major orogenic events, such as
those at circa 2.8, 2.0, 1.0, and 0.5 Ga (Figure 6.2b).
From a metallogenic viewpoint, the proliferationof plume activity attributed to MOMO eventsmight coincide with periods of alkaline magmatism and is, therefore, relevant to the ore deposits
associated with such rocks.

Eustatic sea-level changes and“continental freeboard”
Continental freeboard, or the relative elevationof the continental land masses with respect to sea
level or the geoid, is effectively a measure of thearea of the exposed continents compared to thatcovered by the oceans. Continental freeboard isreduced when the average ocean depth is reduced,
since this causes marine transgression and flooding of the continental shelf. This situation islinked to events of large-scale continental dispersal, which in turn reflects enhanced creation ofoceanic crust, and active magmatism and uplift
along mid-ocean ridges. By contrast, an increasein the area of the continents occurs when theoceans are deepened and the continental shelf
exposed, a situation which tends to be associated with periods of tectonic stability and muted

ORE DEPOSITS IN A GLOBAL TECTONIC CONTEXT CHAPTER 6

Wilson cycle

Wilson cycle

100

200

Ma
400 440

300

100

200

300

400 440

SHALLOWOCEANS

4

The future

3
2
1
0
–1

DEEP
OCEANSToday–2

Fragmentation

Dispersal

Assembly

Stasis

ift

Open

0

100
Close

Fragmentation

Dispersal

Assembly

200

Ma
ift

100

R

Close

200

R

300
ea

Open

400

ng

500

R

ift

600

S
co up
nt erin
en
t

–3

Pa

Water depth at shelf break (×100 m)

5

0

319

StasisFigure 6.6 Pattern of Wilson cycles, reflecting Pangean assembly and fragmentation during the Phanerozoic Eon(and into the future), and showing water depth of the continental shelf (after Nance et al., 1986).

mid-ocean ridge magmatic activity (Worsley
et al., 1984; Nance et al., 1986). The cyclicity ofcontinental freeboard and ocean depth in the
Phanerozoic Eon is well documented (Figure 6.6)
and shows that a period of maximum continental
exposure and ocean low-stand (i.e. deep oceans)coincides with the amalgamation of the Pangean
supercontinent in Permian–Triassic times. The
nature and duration of Wilson cycles in the
Phanerozoic, and their projection into the future,
are also shown in Figure 6.6. A similar periodof ocean lowstand may also have applied withrespect to the amalgamation of an earlier supercontinent, Rodinia, at approximately 1000 Ma(see section 6.5.2 below). From a metallogenic
viewpoint, continental freeboard has its biggest
influence on the nature and preservation of
sediments forming on the continental shelves.
Marine transgressions flood the continental shelf,preserving the sediments within which deposits
such as heavy mineral beach placers, SEDEX
Pb–Zn ores, banded iron-formations, bedded manganese deposits, and phosphorites might haveformed. Oceanic lowstand and maximal continental area, on the other hand, would tend to resultin shelf exposure and erosion, with possible destruction of any ore deposits present in the sedimentary sequence.

6.5 METALLOGENY THROUGH TIME
The pattern of crustal evolution in the Phanerozoic
Eon, involving the progressive amalgamation oflarge continental fragments to form the supercontinent of Pangea during Permian and Triassic
times, and its subsequent dispersal to form
the present day continental geography, is well

320

3.0

PART 4 GLOBAL TECTONICS AND METALLOGENY

2.5

2.0

Age in Ga (not linear)
1.5
1.0

0.5

Ur

0.3

E. Gondwana

Yilgarn

N. India, central Aust.

Arctica (also called Laurentia)
E. Antarctica

Nena

W. Gondwana

Rodinia

Baltica
Atlantica

Gondwana
PangeaLaurentia

Present
day
configuration

Kazakhstan,
China,
S. China, etc.

Figure 6.7 Schematic diagram showing the approximate timing of amalgamation and dispersal of the major
continents through time (after Rogers, 1996).

constrained. A relatively high degree of confidence marks the reconstruction of continental
configuration during the Phanerozoic Eon andgeoscientists generally agree on the paleogeography of Gondwana, Laurentia, and Pangea. As
one goes back in time, however, even into the late
Precambrian, the situation becomes progressively
more uncertain and disagreement often accompanies the reconstruction of continental geometry.
There are many reasons for this, including thedifficulties in acquiring and accurately dating
apparent polar wander paths, the high rates ofrecycling (or destruction) of continental andoceanic crust, and progressive deformation andburial of crust with time.
Despite these difficulties, considerable progresshas been made in reconstructing continental
paleogeography during the latter stages of theProterozoic Eon. There are indications that a
Neoproterozoic supercontinent, generally referred
to as Rodinia, was assembled from about 1000 Ma,
and then dispersed again by about 700 Ma. Thereis also some evidence for the existence of a
Mesoproterozoic supercontinent, although at thisstage there is still considerable uncertainty as tothe nature of its constituents, as well as its shapeand position. It is also likely that substantial continents existed in the Paleoproterozoic and eveninto the late Archean, although at this stage theconfiguration of these continents is speculative.
Windley (1995) has compiled a detailed accountof the many ideas regarding the evolution of thecrust over the entire span of Earth history, andthis work provides an excellent platform uponwhich to base the following discussion. Rogers(1996) has also attempted to reconstruct a historyof the continents and their configuration over the
past 3000 million years. He has also chosen toname the larger continental fragments over thistime period and charted their evolution in termsof major periods of amalgamation and dispersal.
In the sections that follow this scheme is summarized and complemented with more recent data
pertaining to Neoproterozoic (Rodinia) and Phanerozoic (Pangea) continental evolution. Althoughstill in its early stages of development and, therefore, speculative, the Rogers scheme neverthelessprovides a useful framework upon which to base a
discussion of metallogeny through time. It is alsoa working hypothesis that can be used to test
ideas relating to crustal evolution and the link to
global metallogeny. Figure 6.7 presents a schematic
history of continental amalgamation and dispersal
between 3000 and 300 Ma, together with names
for suggested early continental fragments.
In brief, the Rogers model commences withan inferred Archean continent known as Ur, comprising the ancient Kaapvaal and Pilbara cratonsof southern Africa and Western Australia, respectively, as well as parts of India and Antarctica. Ur
is believed to have coalesced from about 3000 Ma
and to have existed as a separate block throughout
much of Earth history until it contributed to theassembly of a supercontinent at around 1000 Ma.

ORE DEPOSITS IN A GLOBAL TECTONIC CONTEXT CHAPTER 6

Another major continent, called Arctica, now
largely preserved in parts of Canada, Greenland,
and Siberia, is believed to have amalgamatedtoward the end of the Archean or in the earlyProterozoic. At least two other substantial continental fragments are believed to have formed
during the Paleoproterozoic, namely Baltica andAtlantica. Arctica (also commonly referred to as
Laurentia in earlier reconstructions; Hoffman,1988) and Baltica are suggested to have amalgamated at around 1500 Ma to form the first substantially consolidated mass of continental material(but not yet a supercontinent), called Nena (Figure6.7). Nena, Ur, and Atlantica are considered tohave been the founding blocks for what wasarguably the first supercontinent, Rodinia, which
existed largely in the Neoproterozoic and comprised an amalgamation of most continental massat that time. Thereafter, the pattern of continental evolution is far better constrained. Rodinia
broke up into essentially three large fragments
known as Laurentia (i.e. an enlarged version of
the Paleoproterozoic Laurentia) and east andwest Gondwana, although it is likely that othersmaller continental entities also existed at thistime. The two sections of Gondwana amalgamated into a single entity during the Panafrican
and Brasiliano orogenies at around 500 Ma.
Laurentia collided with Baltica after the closureof the Iapetus Ocean in the early Paleozoic, atabout 400 Ma, to form a transient entity known
as Laurasia. It was the subsequent collision of
Laurasia and Gondwana that gave rise to the second global supercontinent, Pangea, which formedat around 300 Ma. Since then continental fragments have dispersed to many parts of the globe,
forming the present day geographic configuration.
Some areas are currently undergoing reamalgamation, with collision of India with Asia, andthe partial amalgamation of Africa and Europe,
having been particularly active over the past
20 million years.
This pattern of continental evolution is usedbelow as the framework around which to discuss
global metallogenic trends. The discussion progresses from oldest to youngest, even thoughmuch more is known about processes in younger
periods of time.

321

6.5.1 The Archean Eon
The Hadean (>4000 Ma) and Eoarchean(>3600 Ma) stages
The Hadean Era refers to that period of Earth history for which there is very little evidence in therock record and which is nominally pre-Archean.
It was a time of global differentiation and accretion, as well as intense meteorite bombardment.
The Hadean was previously regarded as existingprior to 3800 Ma (Harland, 1989) although thegrowing evidence (from U–Pb zircon dating) for
crustal remnants at close to 4000 Ma suggeststhat the latter date is perhaps a more accurate
reflection of the Hadean boundary (Windley,
1995). The lack of any meaningful preservationof Hadean crust anywhere on the face of the Earthis a feature generally attributed to widespread
destruction of this ancient material, either byintense meteorite bombardment or by subduction
associated with a turbulent, rapidly convecting
mantle, or both. There is also evidence to suggestthat the early atmosphere and ocean formed onlyat the end of the Hadean era, once the main periodof accretion and meteorite bombardment had
terminated (Kasting, 1993). De Wit and Hynes(1995) have suggested that the Hadean Earth was
also characterized by loss of heat direct to theatmosphere, in contrast to later periods of time
when heat loss is largely buffered by a liquidhydrosphere. The implications for metallogenesis
are that sedimentation and hydrothermal process
are likely to have been inconsequential in the
Hadean, and any ore deposits that did form at thattime were, therefore, probably igneous in character. It is conceivable, for example, that oxide andsulfide mineral segregations accumulated fromanorthositic and basaltic magmas at this time.
The only preserved record of such rocks withinreach of humankind at present is, however, likely
to be on the Moon.
The Eoarchean refers to the dawn of Archeantime and to rocks formed prior to 3600 Ma,
although for the purposes of this discussion it isconsidered to extend between 4000 and 3600 Ma.
The best preserved section of Archean crust that
falls into this time bracket is the 3800 Ma Isua

322PART 4 GLOBAL TECTONICS AND METALLOGENY(a)
Ur at 3.0 Ga

Dharwar Bhandara

India

Singhbhum

PilbaraNapier Vestfold
Kaapvaal

Antarctica

Dronning MaudLand
Ur at ~1.5 Ga

Aravalli

UR boundary duringcollision with Nena
at ~1 Ga
Yilgarn
Kimberley

Madagascar
INDIA

AFRICA
Zimbabwe

GawlerBundelkhand

East
Australia
IA

AL

ANTARCTICA

AUST R(b)

Ald

Anabar

nd
nla

ee

Gr

Slave

The

lon

-Ak

itka

n tr

an
end

Arctica

ThelonNain

Wyoming

Superior

Figure 6.8 Suggested configurationand paleogeography of the two
major continental blocks, Ur andArctica, that may have been in
existence by the end of the Archean
Eon at 2500 Ma (after Rogers, 1996).
(a) The configuration of Ur at about
3.0 Ga relative to the continental
outlines of southern Africa, India,
Antarctica, and Australia in a
Gondwana make-up and an
enlarged configuration for Ur atabout 1.5 Ga that includes
accretion of continental crust(such as the Zimbabwe craton)
formed largely at 2800–2600 Ma.
(b) The configuration of Arctica
at around 2500 Ma relative tocontinental outlines for North
America and Greenland.

ORE DEPOSITS IN A GLOBAL TECTONIC CONTEXT CHAPTER 6

supracrustal belt and associated Itsaq (previously
called Amitsoq) gneisses of western Greenland.
The Isua belt comprises mafic and felsic metavolcanics, as well as metasediments, and resembles
younger greenstone belts from elsewhere in theworld. Although only 4 × 30 km in dimension,
the Isua belt contains a major chert–magnetite
banded iron-formation component as well as minor
occurrences of copper–iron sulfides in banded
amphibolites and in iron-formation (Appel, 1983).
The largest iron-formation contains an estimated
2 billion tons of ore at a grade of 32% Fe. Scheelitemineralization has also been found in both
amphibolite and calc–silicate rocks of the Isuabelt, an association which suggests a submarineexhalative origin. The coexistence of banded ironformations and incipient volcanogenic massive
sulfide style mineralization points to sea-floor
processes not unlike those active throughout
much of subsequent Earth history. Although thezones of known mineralization in the Isua belt
are sub-economic, at 3800 years old they clearly
represent the oldest known ore deposits on Earth.
The Paleo-, Meso-, and Neoarchean stages(3600 to 2500 Ma)
The main stage of Archean crustal evolution
took place over an extended duration of morethan 1000 million years, during which time geological processes were probably not too dissimiliar to those of today – provided that allowances
are made for features such as higher heat flow,
thicker oceanic crust, and an anoxic atmosphere.
De Wit et al. (1992) described the processes thattook place during this period of time in terms oftwo principal stages, termed “intra-oceanic shieldformation” (between about 3600 and 3100 Ma)
and “intra- and inter-continental craton formation” (between about 3100 and 2500 Ma). The
latter stage terminated, in the Neoarchean, in
what might have been the most intense periodof crustal development in Earth history (Figure6.2b). The Neoarchean was also characterized
by extremely active ore-forming processes representing igneous, hydrothermal, and sedimentarydeposit types.
The Rogers model (Figure 6.7) suggests that by323

the end of the Archean (2500 Ma) there might
have been two major continental blocks in existence, Ur and Arctica. The two blocks would
have comprised segments of ancient crust now
preserved in various parts of different present day
continents and it is, therefore, doubtful that theirconfiguration in the Archean can be known withany real certainty. A suggested configuration for
Ur at around 3000 Ma is shown in Figure 6.8a
and comprises the Kaapvaal and Pilbara cratonsof southern Africa and Western Australia repectively, linked via a corridor made up of Archeansegments from India and Antarctica. The existenceof Ur receives some support from the similiarities
that exist in the nature and ages of Archean
greenstone belts and supracrustal sequences on
the Pilbara and Kaapvaal cratons (Cheney, 1996;
Martin et al., 1998a), a feature that is especiallystriking in the similarities between the hugeSuperior-type banded iron-formations of the tworegions. The unique occurrence of the Archeanaged Witwatersrand basin on the Kaapvaal Craton,and its apparent absence in Western Australia,
on the other hand, would tend to question therelevance of comparative geology that far backin Earth history. Likewise, paleomagnetic data
contradict a close fit of Pilbara and Kaapvaal(Wingate, 1998; Evans et al., 2000) so that thereal paleogeography, and perhaps even the very
existence of Ur at all, must remain questionable. Nevertheless, as a working hypothesis, apossible 1500 Ma configuration for Ur is shownin Figure 6.8a, in which components of Neoarchean crustal growth, such as the highly mineralized Zimbabwe craton, are also shown. Arctica(Figure 6.8b) is considered to have been made up
of several ancient continental fragments (including west Greenland and the Slave province with
their remnants of circa 4000–3800 million year
old crust), although the actual amalgamationand of this continent is considered to have postdated Ur, at around 2500 Ma. The fit between the
Siberian and Canadian portions of Arctica wasoriginally suggested on the basis of colinearity inthe trends of the Paleoproterozoic Akitkan and
Thelon magmatic arcs.
In a metallogenic context, Archean crustal
evolution can be viewed in terms of a two stage

324

PART 4 GLOBAL TECTONICS AND METALLOGENY

model which suggests that early “shield” formation, in which amalgamation of oceanic basaltic
terranes and emplacement of early tonalite–
tronhjemite–granodiorite (TTG) magmas wasfollowed by “cratonization,” where modern plate
tectonic processes such as subduction and continent collision occurred (De Wit et al., 1992). The
characteristics and relevant ore-forming processesof these two stages are illustrated in Figure 6.9.
Shield formation (circa 3600–3100 Ma)
This period conforms approximately with the
Paleoarchean era and is typified by amalgamation of oceanic and arc-formed crust and theincipient stages of continental crust formationby emplacement of early, tonalite–trondhjemite–granodiorite (TTG) plutons (De Wit et al., 1992;
Choukroune et al., 1997). This stage of Archeancrustal evolution is characteized by the development of early continental shield areas comprisinghighly deformed (oceanic) greenstone remnants
occurring as megaxenoliths within extensive TTG
terranes (Figure 6.9a).
The styles of mineralization that formed during
this stage of Earth evolution are limited and best
exemplified by the deposits previously described
for the Isua belt of west Greenland. Algoma type
banded iron-formations are a common component of early greenstone belt assemblages and reflect
the low oxygen levels of the atmosphere and theabundance of ferrous iron, sourced from exhalative activity at the mid-ocean ridges. The existence of oceans at this time and the likelihood of
exhalative hydrothermal processes on the sea floor
would have resulted in the formation of volcanogenic massive base metal sulfide deposits, although
examples are rare. An exception is the well preserved Big Stubby VMS deposit, in the 3460 Ma
Warrawoona Group metavolcanics of the Pilbara
Craton in Western Australia (Barley, 1992).
Cratonization (circa 3100–2500 Ma)
This stage of Archean crust formation coincides
broadly with the Meso- and Neoarchean eras andis illustrated in Figure 6.9b. It is suggested to be
the most prolific period of crustal productionin Earth history (Figure 6.2b) and is also a time ofmajor global mineralization. The processes activeat this time were not unlike those taking place
later on in Earth history and involved widespread
plate subduction, arc magmatism, continent collision and rifting, and cratonic sedimentation.
In Figure 6.9b two sketches are presented. An
earlier sub-stage illustrates consecutive accretionof island arcs onto a previously formed continental
shield and stabilization of the latter by intrusionof large granite batholiths. A later sub-stage envisages the existence of Archean cratons consistingof numerous terranes, each bordered by major
suture zones (possibly the fossilized sites of subduction or arc collision) and flanked by bothactive and passive margins. Sites of intracratonic
sedimentation are also envisaged in this stage ofdevelopment. This scenario is consistent with thepattern of Archean crustal evolution envisaged,
for example, in the Superior Province of Canada
by Choukroune et al. (1997).
From a metallogenic viewpoint this stage ofArchean crustal evolution gave rise to a wide
variety of ore-forming processes (Figure 6.9b).
Arc-related volcanism associated with plate subduction contributed large volumes of magma to
the accreting terranes of the time. Well mineralized examples of continental crust formed in the
period 3100–2500 Ma are represented by the granite–greenstone terranes of the Superior Provinceof Canada, as well as the Yilgarn and Zimbabwe
cratons. Greenstone belts are hosts to numerous
important volcanogenic massive sulfide (VMS)
Cu–Zn ore bodies, such as those at Kidd Creek
and Noranda in the Abitibi greenstone belt of theSuperior Province. Off-shore, in more distal environments, chemical sedimentation gave rise to
Algoma type banded iron-formations, examples ofwhich include the Adams and Sherman deposits,
also in the Abitibi greenstone belt. Greenstone
belts formed at this time also often containkomatiitic basalts that, under conditions favorable for magma mixing and contamination, exsolved immiscible Ni–Cu–Fe sulfide fractions toform deposits such as Kambalda in Western Australia and Trojan in Zimbabwe. During periodsof compressive deformation, major suture zones
became the focus of hydrothermal fluid flow

(a)SHIELD FORMATION (3600–3100 Ma)
Isuatype
VMS
base–metalmineralization

Calc–alkaline
volcanism

Algoma-style
banded-ironformations
ion

Diapiric
gnelssdomes

uct

Sea level
Intr

Shallow
oceans

ean

c

a-o

bd
ic o

MOR
Thick
crust
Tonalite–trondhjemite magmatism
Basaltic
magmatism(b)

CRATONIZATION (3100–2500 Ma)
1 EARLY STAGES:
Greenstonebelts

Major granite
emplacement

Islandarc

Algoma-type
banded-ironformation
Sea level

ShieldMOR
Superior-type
banded-ironformations

Komatite-hosted
Ni–Cu deposits

Archean VMS
deposits

2 LATER STAGES:
Archean“lode-gold” ororogenic gold
PassivemarginMajorsutureWitwatersrandtype
Au–U placer
Forelanddepositsbasin

Islandarc

CratonFigure 6.9 (a) Early stage of Archean crustal evolution, termed intra-oceanic shield formation, occurring in theinterval 3600–3100 Ma. Mineralization in this period seems to have been restricted mainly to ocean floor exhalative
processes and formation of Algoma-type banded iron-formations. (b) Later stage of Archean crustal evolution, termed
intra- and inter-continental craton formation, occurring between 3100 and 2500 Ma. Mineralization processes werevaried and resulted in the formation of many important ore deposit types, including lode- or orogenic gold, VMS
Cu–Zn, paleoplacer Au–U, Algoma- and Superior-type iron ores and komatiite-hosted Ni–Cu deposits (after De Wit
et al., 1992).

326

PART 4 GLOBAL TECTONICS AND METALLOGENY

derived from either metamorphic devolatilizationor late-orogenic magmatism. This resulted in theformation of the voluminous and characteristicstyles of orogenic gold mineralization which are
typical of most late Archean granite–greenstoneterranes worldwide. Examples include important
deposits such as the Golden Mile, Kalgoorlie
district of Western Australia, the Hollinger–
McIntyre deposits in the Abitibi greenstone belt,
the Sheba–Fairview deposits of the Barberton
greenstone belt, and the Freda–Rebecca mine in
Zimbabwe. Early intracratonic styles of sedimentation, often in a foreland basinal setting, gave rise
to concentrations of gold and uraninite represented by the Witwatersrand basin in South Africa.
At least some of this mineralization is placer
in origin and was derived by eroding a fertile
Archean hinterland. The passive margins to theseearly continents would have developed stable
platformal settings onto which laterally extensive
Superior type banded iron-formations could havebeen deposited. A very significant period for deposition of iron ores such as those of the Hamersley
and Transvaal basins of Western Australia andSouth Africa respectively, as well as the Mesabi
range of Minnesota, seems to have been aroundthe Archean–Proterozoic boundary at 2500 Ma.
6.5.2 The Proterozoic Eon
The period of time around 2500 Ma represents amajor transition in the nature of crustal evolution, involving changes in the volume and composition of the continents, tectonic regimes, andatmospheric make-up. It is also clear from secular
metallogenic patterns (Figure 6.1) that these evolutionary changes affected ore-forming processes
and characteristics. The Proterozoic Eon spansa vast period of geological time, from 2500 to
540 Ma, including the period between 2000 and1000 Ma that was marked by a relative paucity oforogenic and/or magmatic-hydrothermal deposit
types, but abundant ores hosted in intracontinental sedimentary basins and anorogenic igneouscomplexes (Figure 6.1). The reasons for this pattern are multifaceted and complex, but, as afirst order approximation, are related to a higher
degree of continental stability and the existenceof major land masses which amalgamated anddispersed in relatively long-lived Wilson cycles(Windley, 1995). A substantial volume of continental crust must have been in existence by the
beginning of the Proterozoic Eon (Figure 6.2) and
it is widely held that the first supercontinent
came into existence during this time, although its
form and evolution are still largely speculative(Piper, 1976; Hoffman, 1988; Park, 1995).
The Rogers model (Figure 6.7) envisages that,in addition to Ur and Arctica, the Paleoproterozoic witnessed continuing amalgamation of landmasses to form Atlantica (from about 2000 Ma
and comprising parts of west Africa and SouthAmerica; Figure 6.10a) and Baltica (the Paleoproterozoic basement to what is now western Europe). Further amalgamation of two (i.e. Arctica
and Baltica) of the four continental land masses inexistence at this time is believed to have occurredat about 1500 Ma, to form a new continent that
Gower (1990) and Rogers (1996) called Nena (anacronym for Northern Europe and North America).
Figure 6.10b shows a possible reconstructionof Nena, at about 1500 Ma, and is very similar
to a scheme proposed by Park (1995) in which
Laurentia and Baltica were joined together
between 1900 and 1500 Ma. The process of continental amalgamation continued through the
Mesoproterozoic, until approximately 1000 Ma,
when virtually complete consolidation gave riseto the formation of a single supercontinent, now
widely referred to as Rodinia (McMenamin andMcMenamin, 1990). Rodinia formed largely byaccretion of Ur and Atlantica to Nena (Figure 6.7)
along a major, almost continuous (in some reconstructions) suture zone known, at least in North
America, as the Grenville orogeny. In the Neoproterozoic, and certainly by about 700 Ma, Rodinia
was broken apart again, mainly along two major
rifts. One of these separated Atlantica from Nena,along a previous suture, to form two fragments
known thereafter as Laurentia and west Gondwana. The other split Nena internally to separate
Laurentia from east Gondwana. Figure 6.11a showsa configuration for Rodinia at 1000–700 Ma andFigure 6.11b illustrates a possible break-up scheme
in which Laurentia has extricated itself, leaving a
recombined west and east Gondwana at the end of

ORE DEPOSITS IN A GLOBAL TECTONIC CONTEXT CHAPTER 6

327

(a)
Atlantica: 2.0 Ga
West
AfricaGuyanaWest Nile

CongoKasai

Brazil~Sao
Francisco

Enlarged
Atlantica
Tanzania

Rio de la Plata(b)
Nena: 1.5 Ga
Siberia
Figure 6.10 (a) Map showing a
possible configuration for Atlantica
at 2000 Ma, comprising mainly
remnants of Archean crust in
Brazil and west-central Africa.
(b) Map showing the configuration
envisaged for the possible
amalgamation of Arctica, Baltica,
and east Antarctica to form Nenaat about 1500 Ma. Both maps areafter Rogers (1996).

Baltica

East
Antarctica

the Proterozoic Eon. This configuration, althoughstill largely conjectural and the subject of continual modification and further research (Dalziel
et al., 2000), sets the scene for the subsequent
pattern of crustal evolution and global tectonics
in the Phanerozoic Eon.
In summary, crustal evolution through theProterozoic Eon was characterized by intermittent continental growth (primarily at around 2000
and 1000 Ma; Figure 6.2) and amalgamation which
ultimately gave rise to the existence of a stable
and long-lived supercontinent in the Neoprotero-North
AmericaGreenland

zoic Era. Although geological processes were notsignificantly different from those of other periodsof Earth history, the interval between 2000 and1000 Ma is typified by ore deposits related toanorogenic magmatism and intracontinental basin
deposition (Figure 6.1). These ore-forming processes point to a pattern in which a substantial
part of Proterozoic crustal evolution was characterized by relatively long-lived periods of continental stability. Continental stability at this time
may have been promoted by an asymmetry inthe distribution of crustal type, where a largely

(a)~750 Ma

UR
India

Rift

Australia

ARCTICA

East Antarctica
Siberia

Ri

ft

KalahariN

Laurentia

NENA

0°

Congo
Amazonia

Phanerozoic belts

1000 Ma belts (Grenville)

CA

TI

AN

L
AT

800–500 Ma belts
BalticaWest
Africapre-1000 Ma
cratons(b)∼500 Ma
E AS
T GO
NDW
AN
A

Proto Pacific Ocean0°

East Antarctica

Laurentia

Plate

Kalahari
India

N
Congo
Amazonia

West
AfricaWES

A
AN
T GONDW

Iapetus Ocean

BalticaORE DEPOSITS IN A GLOBAL TECTONIC CONTEXT CHAPTER 6

oceanic hemisphere was antipodal to a static continental domain. This feature might have beenfurther accentuated, at around 1000–700 Ma, by
the presence of a Rodinia. A similar situation issuggested to have applied during the existence of
the Pangean supercontinent in Permian–Triassic
times (Nance et al., 1986). The metallogenic contexts of the three Proterozoic eras are discussed in
more detail below.
The Paleoproterozoic Era (2500–1600 Ma)
From a metallogenic viewpoint, the period ofearth history between 2500 and 1600 Ma is verysignificant because of the major changes that
occurred to the atmosphere, especially the risein atmospheric oxygen levels at around 2200 Ma(Figure 6.3). Prior to this time, the major oxygen
sink was the reduced deep ocean where any photosynthetically produced free oxygen was consumed by the oxidation of volcanic gases, carbon,
and ferrous iron. In this environment bandediron-formations, as well as bedded manganeseores, developed, as is evident from the widespreadpreservation of both Algoma and Superior type irondeposits. The increase in ferric/ferrous iron ratio
in the surface environment that accompanied
oxyatmoinversion at 2200 Ma, and the accompanying depletion in the soluble iron content of theoceans at around this time, resulted in few BIFs
forming after this time. The stability of easily oxidizable minerals such as uraninite and pyrite isalso to a certain extent dependent on atmospheric
oxygen levels and it is, therefore, relevant thatmajor Witwatersrand-type placer deposits did notform after about 2000 Ma. Besides these nonFigure 6.11 (Opposite) (a) Simplified map showing
a possible configuration for Rodinia, comprising
the amalgamated continental land-masses of Nena
(previously Arctica and Baltica), Atlantica, and Ur,
from about 1000 Ma. Rift lines and arrows show the
way in which Rodinia might have started to breakapart, from about 750 Ma onwards (details from
Hoffman, 1991). (b) Simplified map showing how
fragments of Rodinia might have reconsolidated by
the end of the Proterozoic and early Phanerozoic(at about 500 Ma) to form an early Gondwana
configuration. Both maps are after Rogers (1996).

329

recurrent changes, which only affected oxygensensitive ore-forming processes, the pattern ofmetallogeny in the Paleoproterozoic followed
normal global tectonic constraints.
The continent of Ur remained tectonicallyquiescent throughout much of the Proterozoic
period although it witnessed episodic growth ataround 2000 Ma (possible amalgamation with the
Yilgarn and Zimbabwe cratons) and again at about
1500 Ma (amalgamation with parts of northern
India and central Australia; Figure 6.7). This earlystability is reflected in the widespread depositionand preservation of banded iron-formations along
shallow continental platforms at the Archean–
Proterozoic boundary, already mentioned. On the
Zimbabwe craton, rifting at around 2500 Ma gave
rise to intrusion of the Great Dyke, with its significant Cr and PGE reserves. At 2060 Ma on the
Kaapvaal craton, the enormous Bushveld complex with its world-class PGE, Cr, and Fe–Ti–Vreserves was emplaced, as was the Phalaborwa
alkaline complex with its contained Cu–P–Fe–REE
mineralization. The period between 2000 and1800 Ma, however, was characterized by a global
orogeny, largely accretionary in nature (Windley,
1995), which also affected Ur. The Australian
Barramundi and southern African Kheis orogenies
are events which contributed to the growth of Ur
and, in the latter region, for example, gave rise to
the Haib porphyry Cu deposit and small MVTtype Pb–Zn deposits along the western edge of the
Kaapvaal craton. A further long period of cratonic
stability ensued and Ur was subjected to rifting
and intracontinental sedimentation. The period
between 1700 and 1600 Ma saw the depositionof large dominantly clastic sedimentary basinsthat host the world-class SEDEX Pb–Zn ores ofeastern Australia (Mount Isa, Broken Hill, andMcArthur River) and South Africa (Aggeneys andGamsberg). Elsewhere on Ur at this time, anorogenic magmatism was also occurring, importantlyin the form of the 1600 Ma Roxby Downs granite–rhyolite complex, host to the enormous magmatichydrothermal Olympic Dam iron oxide–Cu–Au
deposit in South Australia.
A slightly different pattern of Paleoproterozoic
metallogeny is evident with respect to the continental fragments of Arctica and Baltica, which

330

PART 4 GLOBAL TECTONICS AND METALLOGENY

had combined by about 1500 Ma to form Nena
(Figure 6.7). Early cratonic stability of Arctica isevident in the deposition of the Huronian Supergroup at 2450 Ma which contains the paleoplacer
uranium ores of the Eliot Lake–Blind River regionsof the Superior province, Canada. Both Arctica
and Baltica were, however, subjected to extensive accretionary orogenies between 2000 and
1700 Ma. The Trans-Hudson, Yavapai–Mazatzal
and Svecofennian orogenies, for example, produced
significant new crust within which volcanogenic
massive sulfide Cu–Zn deposits such as Flin Flonin Canada, Jerome, Arizona, and the Skellefte(Sweden)–Lokken (Norway) ores of Scandanavia
are preserved. Relative stability followed thisperiod of orogenesis, during which time large
intracontinental sedimentary sequences such as
the Athabasca basin formed at around 1700 Ma.
It should be noted that the very rich uranium oresin the latter are epigenetic and probably formed
during several later episodes of fluid flow between
1500 and 1000 Ma (Hecht and Cuney, 2000).
The continent of Atlantica (Figure 6.10a) was
consolidated only after 2000 Ma, subsequent toa major compressional event reflected in west
Africa as the Birimian orogeny at 2100–2000 Ma.
This major crust-forming event gave rise to theimportant orogenic or lode-gold deposits of Ghana,such as Ashanti. Atlantica was relatively stable
for the remainder of the Paleoproterozoic and saw
the emplacement of anorogenic type granite magmatism at around 1900–1800 Ma, with which thelarge iron oxide–Cu–Au deposits of the Carajas
region, Brazil, are associated.
The Mesoproterozoic Era (1600–1000 Ma)
The period of geological time after the formationof Nena (i.e. the amalgamation of Arctica andBaltica) at around 1500 Ma appears to have been
one of tectonic quiescence and continental stability
which lasted for several hundred million years.
It culminated at around 1000 Ma in an episodeof widespread orogenic activity (the Grenvilleorogeny and its many analogues worldwide)
which resulted in the final amalgamation of theRodinia supercontinent. Although this period of
orogenesis affected the continental margins of Ur

it appears to have contributed little to the formation of mineral deposits. An exception is provided
in South Africa by the magmatic Cu–sulfide ores
associated with mafic intrusions of the Okiep
copper district in the 1060–1030 Ma Namaqualand belt. In Nena, by contrast, enormous volumes of anorogenic magmatism, especially in the
period 1500–1300 Ma, provided the host rocks
to a number of very important deposits. In abelt stretching from southern California throughLabrador into Scandinavia, numerous intrusionsof gabbro–anorthosite host the large magmatic
Fe–Ti (ilmenite) ore bodies of the Marcy massif
in the Adirondacks and Lac Allard in Quebec.
The same belt also contains alkali granite–rhyolite complexes which give rise to Fe–Au–REE
resources such as those of the St Francois mountains of Missouri. The granite–rhyolite magmatic
complexes were also eroded to form the sediments of the 1440 Ma Belt basin in the northwest
USA, host to the Sullivan Pb–Zn SEDEX deposit
in British Columbia in Canada. In addition, intracontinental rifting at around 1100 Ma in Nena
gave rise to the formation of the 2000 km long
Keweenawan mid-continental rift, stretchingfrom Michigan to Kansas and filled with a thicksequence of bimodal basalt–rhyolite volcanics
overlain by rift sediments. The latter form thehost rocks to the stratiform Cu–Ag White Pine
deposit in Michigan.
The Neoproterozoic Era (1000–540 Ma)
The Neoproterozoic commenced with the formation, at around 1000 Ma, of the supercontinent
Rodinia, arguably the first substantially consolidated land-mass in Earth history. Rodinia waslong-lived and only started to partially fragment
after a static period of more than 250 million
years. Substantial parts of what had been Rodinia
then reconvened toward the end of the Proterozoic (at 540 Ma) to form the very substantial
Gondwanan land-mass during the Pan-African
orogeny (Figure 6.11b). Break-up of Rodinia
started at around 750 Ma when east Gondwana(previously referred to as Ur) rifted away from thewestern edge of Laurentia to initiate the openingof the proto-Pacific ocean. The remaining portion

ORE DEPOSITS IN A GLOBAL TECTONIC CONTEXT CHAPTER 6

of Rodinia (i.e. Laurentia, west Gondwana, Siberia,
and Baltica) drifted southwards, remaining intact
while at the same time rotating clockwise(Torsvik et al., 1996). At about 650–600 Ma both
Baltica and Siberia started to break away from
Laurentia and the Iapetus Ocean formed. By theend of the Proterozoic Eon, the amalgamation ofeast and west Gondwana had taken place andthis large continental mass was situated at polar
latitudes and across the Iapetus seaway from anequatorially located Laurentia (Figure 6.11b). In
detail, the assembly of Gondwana was long-lived
and polyphase and occurred progressively fromabout 750 to 550 Ma. The Neoproterozoic is,therefore, sometimes referred to as the periodof two supercontinents and was notable for its
protracted periods of continental amalgamation,enhanced freeboard, and tectonic stability. The
time between about 750 and 550 Ma was alsocharacterized by the development of at least two,and in places possibly four, major ice ages, one ormore of which was near global in cover andextended to equatorial latitudes. The concept of
the Neoproterozoic “Snowball Earth” (Harland,
1965; Hoffman et al., 1998) has important implications for understanding climate change and,especially, for the proliferation and diversification of organic life at the Precambrian–Cambrianboundary. Global glaciations also have implications for the nature and formation of ore deposits
in the Neoproterozoic Era.
The major ore deposits of the Neoproterozoic
reflect the conditions of continental stability, as
well as the periods of near global ice cover andattendant anoxia, that prevailed at this time. The
extensively developed ironstone ores of northwest Canada and South Australia, associated withthe 750–725 Ma Rapitan and Sturtian glaciogenic
rocks respectively, are considered to be the resultof the build-up of ferrous iron derived from offshore hydrothermal vents in the reduced ocean
waters that accompanied the development of vast
continental and oceanic ice sheets at this time.
Receding glaciers and a return to more oxidizing
conditions would have resulted in conversion oflabile ferrous iron to insoluble ferric iron andprecipitation of the latter from the ocean watercolumn, together with clastic and glaciomarine

331

detritus to form the ironstones (Lottermoser andAshley, 2000). Even more oxidizing conditions
would also have resulted in the precipitation ofmanganese oxides or carbonates in the succession.
In a similar vein, the Precambrian–Cambrianboundary at around 540 Ma is also characterized
by the first major global phosphogenic event that
resulted in the development of vast deposits ofphosphatic sedimentary rock (phosphorites) inseveral parts of the world (Cook and Shergold,
1984). As with sedimentary iron ores, phosphorites
reflect the upwelling of deep, nutrient-rich ocean
waters onto shallow continental shelves withthe syn-sedimentary precipitation of carbonate–apatite (or collophane) onto the shelf floor.
Although the actual formation of phosphate ores
is a complex process (see Chapter 5), there is compelling evidence to suggest that the onset of phosphogenesis at 540 Ma was related to conditions
prevailing toward the end of the Neoproterozoic.
Many phosphorite deposits worldwide immediately overlie glaciogenic sediments, suggestingthat upwelling of phosphorus-rich ocean waters
was promoted by overturn of a stagnant ocean
during the widespread blanketing of sea-ice
associated with a Snowball Earth scenario. The
Precambrian–Cambrian phosphogenic event also
coincides with the proliferation and diversification of organic life and it is pertinent that asignificant proportion of organisms that evolved
at this time developed calcium phosphate skeletal
structures. Phosphorus is, in addition, a universal
nutrient (unlike oxygen) and its concentrationin the oceans at the end of the Proterozoic, and in
the Cambrian, may also be linked to the evolutionof organic life.
The formation of the vast, stratiform Cu–Co
clastic sediment hosted ores of the Central African
Copperbelt is also considered to have formed in anenvironment influenced by the Snowball Earth.
The host Katangan sediments were deposited ona fertile Paleoproterozoic basement in an intracontinental rift, the development of which overlapped with both the Sturtian and Marinoan
glacial events. The Grand and Petit Conglomeratsof the Katangan sequence, for example, represent
glaciogenic sediments capped by carbonates which
are correlated, respectively, with the Sturtian332

PART 4 GLOBAL TECTONICS AND METALLOGENY

and Marinoan events (Windley, 1995). Influx ofoxide-soluble Cu and Co, perhaps derived fromthe local basement, might have occurred as diagenetic fluids migrated along growth faults andthrough the basin during the postglacial stages ofdeposition. Precipitation of ore sulfides would
have occurred when the metal-charged oxidized
fluids encountered reduced sediments or fluids.
The enormous proliferation and diversificationof organic life in the late Neoproterozoic also
resulted in the first substantial development ofhighly carbonaceous sediments. A good exampleof these are the Hormuz sediments that flank theeastern Arabian shield and which could represent
some of the source rocks for the vast Mesozoic
oil and gas fields of the Arabian Gulf (Chapter 5,
Box 5.3; Windley, 1995). The Neoproterozoic is
notable for the paucity of orogenic type deposits(Figure 6.1). Even the extensive Pan-African orogenic belts representing the suturing of continental fragments during Gondwanan assembly
are strangely devoid of, for example, world-class
volcanogenic massive sulfide deposits, the latterbeing relatively abundant in the Paleoproterozoic
orogenic belts of the world. A few small examples
do nevertheless occur, such as the deposits of the
Matchless amphibolite belt in Namibia, Bleïda in
Morrocco, and the Ducktown, Tennessee deposit(Titley, 1993). The global shortage of these ores in
the Neoproterozoic perhaps reflects either a preservation factor or simply lack of exploration success.
6.5.3 The Phanerozoic Eon
By comparison with earlier eons, crustal evolution during the Phanerozoic is well understoodand there is a large measure of confidence that
accompanies the interpretation of tectonic,
chemical, and biological processes over the past
540 million years. The latter half of this period,
the Mesozoic and Cenozoic eras, has been accompanied by a prolific development of mineralization, the likes of which has probably not been
seen since the late Archean bonanza, despite theinfluence of a preservation factor. Ore-forming
processes and their relationships to the patternof Earth evolution in the Phanerozoic are also
reasonably well understood and additional details

regarding some of the topics discussed below canbe found in Nance et al. (1986), Larson (1991),
Barley and Groves (1992), Titley (1993), Kerrich
and Cassidy (1994), Windley (1995), and Barley
et al. (1998).
A significant proportion of the Phanerozoic wascharacterized by geological processes that reflect
a Wilson cycle, namely the sequence of events
that saw the dispersal of Gondwana in the early
Paleozoic, followed by reamalgamation of continental material to form the Pangean supercontinent, by the early Mesozoic. The remainder of thePhanerozoic has witnessed the start of anotherWilson cycle, involving the dispersal of Pangea
to form the present day continental geography.
It seems likely that we are presently about half
way through this Wilson cycle (Nance et al., 1986)
and that it will be responsible for reassemblinga significant proportion of the continents overthe next 100–200 million years (Figure 6.6). Continued continental amalgamation in the future
will extend processes currently taking place,such as the collision of the Indo-Australian plate(previously part of east Gondwana) with the combined Baltica–Siberia (now called the Eurasianplate) and closure of the Mediterranean Sea.
As previously mentioned, break-out of Laurentia
from Rodinia and rotation of Ur at around 725 Ma(Figure 6.11) led to the formation of a Gondwanan
land-mass by the end of the Proterozoic. Dispersalof Gondwana in the Cambrian–Ordovician may
have been facilitated by the earlier developmentof a superplume (Larson, 1991), evidence for which
is seen in the formation of major dyke swarms
in different parts of Gondwana at 650–580 Ma(Torsvik et al., 1996). The Iapetus Ocean formed
as Laurentia drifted away from Baltica and
Gondwana from the late Precambrian onwards.
The pattern of continental fragmentation is illustrated in Figure 6.12a and was accompanied by a
Figure 6.12 (Opposite) The paleogeographic
patterns of Gondwanan and Pangean dispersal in the
Phanerozoic Eon. (a) and (b) Two reconstructions in theearly Ordovician and Silurian (modified after Torsvik
et al., 1996). (c) A reconstruction of Pangea at its peakamalgamation in the Permian–Triassic. (d) and (e) Two
instants of Pangean dispersal in the mid-Cretaceousand Eocene (after Windley, 1995).

(a)Ordovician (490 Ma)
G O N D WA N A N D I S P E R S A L

Equ

Siberia

at

30

or

(b)

S

Baltica

Silurian (420 Ma)

60

S

Ia
pe
tu
s

SCB
Siberia

SCB

Gondwana

Baltica
Laurentia

ia Bohemian
Massiflon

aAv
30°SArmorica-Iberian
Massifs(c)Permian–Triassic (250 Ma)

Panthalassan
Ocean

(d)

Gondwana

Tethys
Ocean

Mid-Cretaceous (100 Ma)

0

(e)

0

PA N G E A

PA N G E A N D I S P E R S A L
Eocene (50 Ma)

0

334

PART 4 GLOBAL TECTONICS AND METALLOGENY

decrease in continental freeboard, sea-level highstand and marine transgression. This was followedby progressive basin closure, terrane accretion,and granitoid magmatism during the period
extending from the Silurian through to the early
Carboniferous (about 420–300 Ma). This collisional phase commenced with the rapid consumption of Iapetus as Baltica and Avalonia(i.e. a small terrane comprising England, Wales,
southern Ireland, and eastern Newfoundland)
moved toward lower latitudes and collided withLaurentia in Silurian times (Figure 6.12b). The
orogenies that reflect this collision phase arereferred to as the Caledonian in Scandanavia and
Scotland/Ireland and Appalachian in the eastern
USA. Continental amalgamation continued as
Gondwana drifted northwards toward lower latitudes, consuming the Rheic Ocean during the
Devonian and Carboniferous, and eventually
colliding with Laurentia–Avalonia–Baltica in thelate Carboniferous–Permian. These complex andpolyphase collisions are referred to as the Variscan
and Hercynian orogenies (the terms are essentially synonomous) in present day Europe and theAlleghanian orogeny in the eastern USA. Continental amalgamation continued into the Permianwith the accretion of Siberia to Baltica along theUrals suture, after which the Pangean supercontinent had essentially formed (Figure 6.12c).
The break-up of Pangea commenced soon after
the time of its maximum coalescence in the
Triassic at around 230 ± 5 Ma (Veevers, 1989). It
was, therefore, a very short-lived supercontinentcompared to Rodinia, which appears to haveendured during much of the early Neoproterozoic.
One reason for this might have been the development of a superplume, or plumes, that existed
for some 70 million years, in Carboniferous andPermian times (320–250 Ma), an interval that also
coincided with a protracted period of constantreversed magnetic polarity (Larson, 1991). The
effects are seen geologically by increased production of oceanic crust and global volcanism (e.g.
the outpouring of the Siberian continental floodbasalt province and also the central Atlantic magmatic province), eustatic sea-level rise, increased
atmospheric greenhouse gas production (withassociated warming and organic proliferation),and enhanced deposition of organic-rich blackshales. The first land-masses to actually drift apart,more than 70 million years after initial plumerelated rifting, were Laurentia from Gondwanain the Jurassic at around 180 Ma. West and east
Gondwana also started to split at this time, withthe development of the proto-Indian Ocean andoutpouring of the Karoo igneous province. The
opening of the Atlantic Ocean was particularly
prevalent in the early Cretaceous (around 130 Ma)
and this extensional phase was also marked
by continental flood basalt volcanism of theEtendeka–Parana provinces in Namibia and Brazil
respectively. Continued dispersal of Pangea wasstimulated by a second superplume event in
the mid-Cretaceous between 120 and 80 Ma(Figure 6.12d). The evidence for this event isagain seen in a 40 million year period of constant,normal magnetic polarity, as well as the expected
increases in oceanic crust production, eustatic
sea level, atmospheric temperatures, organic productivity, and black shale deposition. The verylarge oceanic plateaux of the western Pacific (e.g.
Ontong–Java) are a magmatic reflection of thismid-Cretaceous event, as is the development ofvery significant global oil reserves related to thesurges in nutrient supply and organic productivity
on the voluminous continental shelves formed by
the rise in sea level (Larson, 1991). In post-plumetimes the late Cretaceous saw the final separationof Indo-Australia from Antarctica (Figure 6.13e),
propelling the former northwards and resulting incontinent–continent collision with Eurasia andthe formation of the Himalayan orogeny in the
Cenozoic. The Alpine orogeny of southern Europe
was more or less coeval with its Himalayancounterpart and resulted from the collision of theAfrican and Arabian plates with Eurasia after
consumption of the Tethyan Ocean.
A very important component of Pangean breakup in the Mesozoic–Cenozoic is reflected in thedevelopment of new crust that accompanied
the Cordilleran and Andean orogenies along thewestern margins of North and South America(Windley, 1995). This huge and complex orogeny,
extending along the ocean–continent interface
of western Pangea (Figure 6.12c), occurred in
response to subduction and translocation of theORE DEPOSITS IN A GLOBAL TECTONIC CONTEXT CHAPTER 6

original Panthalassan Ocean (now the eastern
Pacific) beneath Laurentia (now North America),
the Cocos plate (now central America), and segmented west Gondwana (now South America).
An island arc lay to the west of North America
during Pangean times and this was accreted onto
the continental margin, together with numerous
other exotic terranes, during the consumptionof several thousand kilometers of the Pacificbeneath North America during the Mesozoic Era.
The enormous continent-floored magmatic arc
that resulted from this subduction gave rise to the
130–80 Ma granite batholiths and felsic volcanic
rocks (such as the Sierra Nevada batholith) of
the Cordillera. Continued subduction, albeit ata shallower angle than previously, and crustal
thickening during the early Cenozoic, gave rise toin-board extension and exhumation of metamorphic core complexes. Ongoing subduction during
the Laramide orogeny (80–40 Ma) continued tofeed the magmatic arc and numerous, metallogenically important I-type granite batholiths were
emplaced at this time. As subduction waned
in the late Eocene–Oligocene the magmatic
arc migrated westwards, resulting in continued
calc–alkaline magmatism. Thermal collapse commenced in the Neogene and the resulting crustal
extension ultimately led to the formation of theBasin and Range province. In South America,
Andean evolution was analogous, but less complex,
than further north, and was also characterized bya lesser degree of accretionary and extensional
processes. Subduction of the Nazca plate beneatha Paleozoic sedimentary margin, in the late
Cretaceous, gave rise to a volcanic arc, behind
which back-arc sedimentation took place (Lambet al., 1997). This arc was the site of protracted
magmatism that built the Western Cordillera andgave rise to the volcanic edifices which host thevery important porphyry and epithermal styles ofpolymetallic mineralization in Peru, Bolivia, andChile. Continued subduction transferred compressional stresses in-board and created a thinskin fold and thrust belt in what is now referred toas the Eastern Cordillera. In the late Oligocene,
crust was shortened even more and uplift gave
rise to rapid erosion of the mountain belt toform a thick sedimentary plateau of dominantly

335

Miocene-aged gravel and red-bed sequences known
as the Altiplano. Subduction-related volcanism
continued throughout this period in the Western
Cordillera, which was also elevated to form thehigh Andes and then eroded to contribute sediment into the Altiplano basin as well as westwards onto the coastal plain. Further to the east,
magmatism commenced in the early Miocene as aresult of convective removal of the basal lithosphere and crustal melting beneath the Altiplano(Lamb et al., 1997). The buoyancy and elevatedprofile of the Andes is maintained by present day
underthrusting of the Brazilian shield beneath the
Eastern Cordillera.
Tectonic cycles and metallogeny
The relatively well defined global tectonic cyclesof the Phanerozoic Eon, summarized above, arealso clearly reflected in secular metallogenic
trends. Titley (1993), for example, noted that the
distribution of stratabound ores (i.e. volcanogenic
massive sulfide, clastic sediment hosted Pb–Zn(SEDEX), and Cu (red-bed) deposit types) could be
related to the tectonic cycles of Pangean amalgamation and break-up (Figure 6.13). Preferential
development of VMS deposits appears to be associated with periods in the Wilson cycle of elevatedsea levels (highstand) associated with continental
dispersal, namely, after Gondwana break-up in
the early Paleozoic and in post-Pangean Mesozoic
times. This association was considered to reflect
the processes of rifting, enhanced ocean crust
production, and hydrothermal exhalation that
accompany continental dispersal. Similar patterns
are also evident in the accumulation of organicrich shales and oolitic ironstone ores, which preferentially occur in the same two intervals, namely
after Gondwana break-up in the Ordovician–Devonian and in post-Pangean Jurassic–Cretaceous
times (Figure 6.13). In these cases, increased
exhalative activity, carbon dioxide production,
global warming, and organic productivity areinterrelated processes which result in suitableconditions for black shale and ironstone precipitation in the oceans. In contrast, clastic sediment
hosted base metal ores of the SEDEX Pb–Zn–Ag
and red-bed Cu types tend to form at different

WILSON CYCLE
400(meters)

SEA LEVEL

0

600

500

400

300

200

100

Ma

Wilson cycle

DISPERSAL ASSEMSTASISBLYFRAGMENTATIONFRAGMENTATION

Fossil fuels
Petroleum
Coal

P
A
N
G

Volcanogenic
massivesulfidesE
A

Clastichosted(Red-bed) Cu(Ag)(SEDEX) Pb-Zn-Ag

Marine black
shale

Oolitic Fe

Superplumes
CAMB
600

ORD

500

SIL DEV CARB PER TR
300

400

JUR
200

CRET
100

CEN
0

Ma
Figure 6.13 Occurrences (in terms of relative abundances) of a variety of different ore types with time and withrespect to major Wilson tectonic cycles in the Phanerozoic (modified after Titley, 1993; with information from
Nance et al., 1986; Larson, 1991).

ORE DEPOSITS IN A GLOBAL TECTONIC CONTEXT CHAPTER 6

stages of the Wilson tectonic cycle. These ores
are preferentially developed at the time of maximum continental amalgamation and stasis,
namely in Gondwana during the Precambrian–Cambrian transition and in Pangea during the latePaleozoic–early Mesozoic (Figure 6.13). Theseore types appear to have evolved preferentiallyduring intracratonic rifting and sea-level lowstand (Titley, 1993).
The two Phanerozoic superplume events described above are also related in time to periodsof enhanced organic productivity and formationof fossil fuels (Figure 6.13). The mid-Cretaceous
superplume, active between 120 and 80 Ma, coincides with the formation of voluminous organicrich shales deposited largely in the low-latitude
Tethyan seaway. These shales are thought to begenetically linked to some 60% of the world’s oil
reserves (Larson, 1991). The late Carboniferous–
Permian superplume is similarly associated withdeposition of a large proportion of the world’s coalreserves, formed between 320 and 250 Ma, again
at a time when organic productivity increased
and tropical, swampy conditions prevailed along
flooded continental margins associated with thecorresponding sea-level rise. Continental hotspots,
or beheaded plumes, are also possibly related tothe preferential development of alkaline andkimberlitic magmatism associated with enhanced
igneous activity in the mid-Cretaceous and again
in the Cenozoic. Many of these intrusions, important for Cu–REE–P–Fe–Au mineralization and also
diamonds, are found on old, stable cratons andlocated along ancient lineaments that might havebeen reactivated during extension and crustal
thinning associated with hotspot activity.
Time-bound and regional aspects of
Phanerozoic metallogeny
A summary of ore deposit trends as a functionof the major Phanerozoic tectonic events reveals
that the metallogenic inventory, especially of
convergent margin lithosphere, increases withtime and with each tectonic cycle (Barley et al.,1998). The distribution of ore deposit types also
very often has a regional character to it and, as
mentioned previously, may reflect the extent of337

erosion in different regions and the preservation
potential of ores.
The large-scale development of new crust along
the convergent Mesozoic–Cenozoic plate marginsof the Americas did not occur on the same scale
during the Paleozoic dispersal stage. Hence, acyclic pattern in the development of orogenic
magmatic-hydrothermal ores is not as evident
as it is for statabound deposits, where controls onore formation are different. There are, for example, few known porphyry Cu–Mo type deposits
of Paleozoic age, although sub-economic examples
are recognized along convergent margins suchas the Caledonides of Scotland. Orogenic gold
mineralization is also related to the late stagesof collisional orogens. Cretaceous examples ofthis style of gold mineralization are evident in
California, British Columbia, China, and New
Zealand and could conceivably be related to theenhanced thermal regime related to the superplume break-out at this time. It has also been suggested that these gold provinces might rather
have been linked to areas where mid-ocean ridge,
as opposed to normal oceanic crust, was subducted.
Scenarios in which hot crust was subducted, compared to periods when normal subduction of cold
oceanic crust prevailed, have been suggested as an
explanation for the anomalous thermal conditions required for widespread gold mineralization(Haeussler et al., 1995).
The amalgamation stage of Pangea commenced
with the collision of Avalonia–Baltica withLaurentia during the polyphase Paleozoic
Caledonian–Appalachian orogenies. VMS deposits
formed in response to these processes in, forexample, the Bathurst–Newcastle district of New
Brunswick, the Buchans area of Newfoundland,
and Scandanavia. New Brunswick also has examples of Devonian aged Sn–W–Mo–F granites, someof which are mineralised. The carbonate-hosted
MVT Pb–Zn deposits of the southeastern USA are
also related to the Appalachian orogeny and formed
when the thrust front expelled basinal fluids into
the carbonate platform on the continental margin, giving rise to epigenetic hydrothermal mineralization. The Variscan orogeny, formed when
Gondwana collided with an already amalgamatedLaurentia–Baltica–Avalonia, is also characterized338PART 4 GLOBAL TECTONICS AND METALLOGENY

by carbonate-hosted ore bodies. Some of these,
particularly the Irish deposits at Navan, are transitional in style between MVT and SEDEX-type
Pb–Zn deposits, but are also related to compression-related, gravity- and tectonic-driven, basinal
fluid flow. The Variscan orogeny of Europe isparticularly well known for its association withwidespread, 300–275 Myr old, S-type Sn–W–U
granites. These are well mineralized in many of
the historic European mining districts, such as
Cornwall in southwest England (South Crofty,
St Just, etc.), the Massif Central of France(Bellezane), the Bohemian massif of the Czech
Republic, and Portugal (Panasqueira). Other deposits associated with the Variscan of Europe
include the Devonian aged Rammelsberg and
Meggen SEDEX-type deposits of Germany and thenumerous VMS deposits of the Iberian Pyrite Belt
(Rio Tinto, Neves Corvo, etc.). The brief intervalin the Permian–Triassic of relative stability that
accompanied the existence of the Pangean supercontinent was one in which little mineralization
formed. An exception to this was the stratiform red-bed hosted Cu–Ag ores of the Permian
Kupferschiefer in Poland and Germany.
Mineralization associated with post-Pangean
break-up and orogenesis is both spectacular andwidespread. Exceptions to this would appear tobe the products of continent–continent collision,
such as regions affected by the Himalayan orogeny,
which appear to be devoid of any world-class ore
deposits. The Karakoram mountains of Pakistan,for example, contain rich gemstone deposits (suchas the rubies of the Hunza area), but otherwise
only minor MVT and vein-related precious andbase metal type ores, as well as a few smallhydrothermal uranium deposits, occur (Windley,
1995). More significant stratabound and ophiolitehosted styles of mineralization are containedin rocks caught up in the Himalayan orogeny,
but the ore-forming processes probably predate
the latter. Regions affected by the Alpine orogeny
are, likewise, not particularly well mineralized,
although exceptions do occur. MVT-type Pb–Zn
deposits occur in mid-Triassic limestones of theeastern Alps and north Africa. The Carpathian arcof southeastern Europe contains porphyry Cu–Mo
styles of mineralization, such as at Sar Cheshmeh

in Iran and Bor in the former Yugoslavia, and also
contains significant potential for epithermal gold
mineralization. Important ophiolite-hosted VMS
Cu–Zn and chromite mineralization occurs innumerous obducted remnants of ocean floor, inplaces such as Troodos, Cyprus and in Albania
and Turkey. Again, however, these deposits arelikely related to ore-forming processes that predate the Alpine orogeny.
The Andean and Cordilleran orogens of thewestern Americas contain the greatest concentration of metals on Earth, and are pervasively
mineralized from one end to the other. During theJurassic–Cretaceous period of terrane accretion in
North America, small VMS-type Cu–Zn deposits
were formed, as were the “mother-lode” hydrothermal gold systems, hosted in obducted oceanic
material in California, British Columbia, andMexico. Mesozoic and early Cenozoic plutonism
gave rise to the numerous world-class porphyry
Cu–Mo deposits of the USA and Canada, as wellas the related skarn and polymetallic epithermalvein deposits, such as the Bingham system in
Utah. Thermal collapse and the development of
Basin and Range extension gave rise to the enormous Eocene (40 Ma), hydrothermal, sedimenthosted Carlin-type Au deposits of northeast
Nevada, followed by Miocene volcanism andwidespread development of epithermal Au–Ag
mineralization such as the Comstock Lode of
Virginia City, Nevada. Similarly, in South America
a wide variety of mineralization styles are associated with Mesozoic–Cenozoic orogeny. During
the Jurassic–Cretaceous period island arc volcanism and associated hydrothermal activity gave
rise to the Fe oxide–phosphate ores now preservedin the western, coastal portions of the belt.
The axis of subduction related magmatism commenced in the west during the late Triassic andearly Jurassic and then migrated eastwards withtime through the Cretaceous and early Cenozoic(Sillitoe, 1976). In the Eastern Cordillera, however, minor magmatism occurred early in the
Mesozoic, but most of the magmatism took placein the Miocene. The enormous porphyry copper
deposits of Chile (such as Chuquicamata and El
Teniente) and Peru (Morococha and Toquepala)formed in response to diachronous episodes ofORE DEPOSITS IN A GLOBAL TECTONIC CONTEXT CHAPTER 6

subduction-related magmatism. Many of theimportant porphyry copper deposits in Peru are
Paleocene in age, whereas in northern Chile
they are Eocene–Oligocene and, further south,
Miocene–Pliocene (Sillitoe, 1976). East of the porphyry copper belt a broad zone of vein-related andskarn type Cu–Pb–Zn–Ag mineralization occurs,
spatially related to Eocene–Oligocene calc–
alkaline plutons. The deposits are most prolific in
Peru (Antamina and Cerro de Pasco) but the belt
continues southwards into Bolivia and Argentina.
Finally, a distinct zone of Sn–W–Ag–Bi mineralization occurs mainly in the Eastern Cordillera of
Bolivia and southern Peru, in the elbow of thecentral Andes. These ores occur as both veintypes associated mainly with Mesozoic intrusions, and disseminated porphyry type deposits(such as Llallagua and Potosi in Bolivia) associated with Oligocene–Miocene granites.
Also extremely important with respect tomineralization associated with post-Pangean
orogenesis is the extensive chain of island arcs inthe northern and western Pacific. The northern
and western Pacific is predominantly characterized by collisions of oceanic crust (as opposed tothe ocean–continent collisions of the western
Americas) which form island arcs that build
new crust on oceanic (simatic) basement rather
than on continental (sialic) basement. There areseveral types of island arc, each with its particular
metallogenic character. These include intraoceanic arcs (such as Tonga, New Hebrides, andthe Solomons), island arcs separated from continental crust by a narrow back-arc sea (such as
Japan), and those built directly against a continent(such as Java–Sumatra). Soon after initiation ofocean–ocean collision in the western circumPacific region, early, relatively mafic (andesitic–
dacitic), stages of calc–alkaline magmatism
resulted in porphyry Cu–Au deposit formation,
examples of which include Grasberg, Indonesia,
and Bougainville in Papua New Guinea. Besshitype VMS deposits also occur in back arc settings
where andesite–dacite volcanism and ocean floor
exhalative activity occurs synchronously withdeep water sedimentation. During the main stagesof arc construction and calc–alkaline magmatism,
dacite–rhyolite volcanism occurred and is asso-339

ciated with the development of Kuroko-type VMS
deposits, examples of which are known from the
Miocene-aged Green Tuff belt of Japan. Also very
important, in geologically more recent times, isthe formation of large epithermal Au–Ag deposits.
Several world-class ore deposits of this type occur
in the western circum-Pacific region, such as
Baguio and Lepanto in the Phillipines, Hishikari
in Japan, Ladolam and Porghera in Papua New
Guinea, and Emperor in Fiji. Porphyry Cu–Au
styles of mineralization are commonly known tooccur beneath such epithermal deposits, such as
at Baguio and Ladolam.
As a final comment, it is interesting to notethat the Central Asian orogenic belts are characterized by many of the features, both temporaland tectonic, of the Andean and Cordilleranorogens. Still somewhat underexplored, these terranes probably represent particularly attractive
metallotects for future mineral exploration.

6.6 PLATE TECTONICS AND ORE DEPOSITS – A SUMMARY
The description of ore deposits in the context
of their plate tectonic setting has been carriedout in considerable detail in numerous publications (Mitchell and Garson, 1981; Tarling, 1981;
Hutchison, 1983; Sawkins, 1990). A useful diagrammatic summary is presented in Mitchell andGarson (1981), from which Figures 6.14 and 6.15
are extracted. The latter two diagrams serve as a
useful overview of the major plate-related tectonic settings and the ore deposit types associatedwith each.
6.6.1 Extensional settings
Incipient rifting of stable continental crust isrepresented in Figure 6.14a, where thinning andextension may be related to hotspot activity.
Magmatism is often localized along old suturesand is alkaline or ultrapotassic (kimberlites andlamproites) in character. Anorogenic granites suchas those of the Bushveld Complex (Sn, W, Mo, Cu,
F, etc.), pyroxenite–carbonatite intrusions such as
Phalaborwa (Cu–Fe–P–U–REE etc.), and kimberlites(diamonds) represent ore deposit types formed in

340

PART 4 GLOBAL TECTONICS AND METALLOGENY

EXTENSIONAL SETTINGS(a) INTRA-CONTINENTAL HOTSPOT/INCIPIENT RIFTSEDEX Pb-Zn-Ag

Carbonatite Nb, P2O5, F, Nb, REE

Kimberlite
diamondAnorogenic granite Sn, F, Nb

km

0
Ocean

Continent

30

200

400

(b) INTER-CONTINENTAL RIFT ZONE
Evaporitic layer
Red Sea Mn

Metal-rich brinesand muds

km

0

30

200

400

(c) MID-OCEAN RIDGE AND HAWAII-TYPE HOT SPOT CHAIN
LIL-richbasaltCyprus-type
VMS deposits

km

Fe, Mn, Zn, Pb
0chemical sediments

Potassicbasalt30

Alkali-depleted
mantlePodiform
Cr, NiS, Pt

Plume
Primitive mantle
Plume

Figure 6.14 Simplified illustrations of the major extensional tectonic settings and the ore deposit types associatedwith each (modified after Mitchell and Garson, 1981).

ORE DEPOSITS IN A GLOBAL TECTONIC CONTEXT CHAPTER 6

341

COLLISIONAL SETTINGS(a) ANDEAN TYPE OCEAN–CONTINENT COLLISION

Mn nodules withCu, Ni, Co

Epithermal
goldPorphyry Cu, Mo

Polymetallic
skarn

0
Ocean
km

Continent
30

60

200

400

(b) ISLAND ARC AND INTER-ARC BASIN (ocean–ocean collision)

Kuroko Zn, Cu

Incipient
rift

Porphyry Cu-Au

Inactiveinter-arc
basin

km

0

30

NiS,Ptpodiform Cr

200

400

Figure 6.15 Simplified illustrations of the major compressional tectonic settings and the ore deposit types
associated with each (modified after Mitchell and Garson, 1981).

this setting. Intracontinental rifts can host
SEDEX-type Pb–Zn–Ba–Ag deposits. As continental rifting extends to the point that incipient
oceans begin to open (such as the Red Sea; Figure
6.14b), basaltic volcanism marks the site of a midocean ridge and this site is also accompanied byexhalative hydrothermal activity and plentiful
VMS deposit formation. Such settings also provide
the environments for chemical sedimentationand precipitation of banded iron-formations andmanganiferous sediments. Continental platforms
commonly host organic accumulations that on

catagenesis give rise to oil deposits. Carbonate
sedimentation ultimately provides the rocks whichhost MVT deposits, although the hydrothermal
processes that give rise to these epigenetic Pb–Zn
ores are typically associated with circulation during compressional stages of orogeny. Mid-ocean
ridges are the culmination of extensional processes (Figure 6.14c). Exhalative activity at thesesites gives rise to “black-smoker” vents (such as
those at 21° N on the East Pacific Rise) that provide the environments for the fomation of Cyprus
type VMS deposits. The basalts which form at

342PART 4 GLOBAL TECTONICS AND METALLOGENYCOLLISIONAL SETTINGS(c) ISLAND ARC WITH BACK-ARC BASIN
Epithermal gold

Podiform Cr

km

Sn, W, Bi, Mo, F
Cyprus-type
Cu-Fe

Porphyry
Cu-Au
0Besshi-type
Cu-Fe

30 Auriferous quartz veins(orogenic gold)
60
200

400

(d) CONTINENT–CONTINENT COLLISION (Himalayan type)
Sn, W skarn
U in molasse

km

0

30

60

Pb-Zn-Ag on underthrusting
continent

Fe-Ti
anorthosite
?

200

400
km

Figure 6.15 (cont’d)

mid-ocean ridges also undergo fractional crystallization at sub-volcanic depths to form podiform
chromite deposits as well as Cu–Ni–PGE sulfide
segregations.
6.6.2 Compressional settings
The highly significant Andean type collisional
margins are represented in Figure 6.15a. These are
the sites of the great porphyry Cu–Mo provincesof the world, while inboard of the arc significant
Sn–W granitoid-hosted mineralization also occurs.
The volcanic regions above the porphyry systems
are also the sites of epithermal precious metalmineralization. A similar tectonic setting canexist between two slabs of oceanic crust, as represented by the island arc environment in Figure
6.15b. Porphyry Cu–Au deposits occasionallyoccur associated with the early stages of magmatism in these settings, whereas the later, more
evolved calc–alkaline magmatism gives rise to
Kuroko-type VMS deposits. The back-arc basinsrepresent the sites of Besshi type VMS deposition.
Arc–arc collision in the back-arc environment canalso result in the preservation of obducted oceanic
spreading centers within which podiform Cr andsulfide segregations might be preserved. In theJapanese setting, where the island arc develops
fairly close to a continent (Figure 6.15c), marginal
sedimentary basins floored by oceanic crust occur.

ORE DEPOSITS IN A GLOBAL TECTONIC CONTEXT CHAPTER 6

343

This setting typically hosts Besshi- and Cyprustype VMS deposits. As the arc and continent accrete, ophiolite obduction can occur, andfelsic magmatism may give rise to large-ion lithophile element mineralization. Ultimately, theoceanic crust is totally consumed to form a zoneof continent–continent collision (Figure 6.15d).
Modern examples such as the Himalayas andAlps do not appear to be significantly mineralized,
but this may be an expression of insufficient exhumation of mineralized zones. Older examples
preserve Sn–W–U mineralization in S-type granites, whereas orogeny-driven fluids give rise toorogenic vein-related lode Au systems and MVT
Pb–Zn deposits in suitably preserved platformal
sediments.

Broad trends emerge when relating ore deposit
types to geological time and global tectonic setting. The mechanisms by which the crust has
evolved, and the amalgamation and dispersal ofcontinents over time, have played an overridingrole in the formation of all deposit types.
Although the processes are obscure in the early
periods of Earth history, it appears that primitive
continents started to form after the late heavy
meteorite bombardment, from about 4000 Ma.
Only a limited range of ore deposit types, mainly
related to ocean floor hydrothermal processes,
appears to have developed at this stage. As more
evolved continents formed in the Mesoarchean
and conventional plate tectonic processes wereentrenched, the range of ore deposit types broadened, although many of the near-surface variants
may not be preserved due to erosion. The
Neoarchean Era was a period of intense global orogenesis and ore deposits formed at this time tend
to be arc-related and magmatic-hydrothermal
to hydrothermal in nature, not unlike thosetypifying the latter part of the Phanerozoic Eon.
Although global tectonic processes continued
unabated in the Proterozoic Eon, and a number ofmajor crust-forming orogenies occurred, thisextensive period of time appears to have beencharacterized by longer periods of tectonic quiescence and continental stability. Consequently,
deposits reflect hydrothermal and sedimentaryore-forming processes in intracratonic settings or
passive margins. Likewise, magmatism is morecommonly anorogenic or post-tectonic in natureand is associated with an entirely different suiteof metal deposits than those found in arc-related
provinces. Global tectonic processes are muchbetter understood in the Phanerozoic Eon and the
distribution of continental land masses is well
documented in terms of Wilson cycles. Certain
deposit types (such as VMS Cu–Zn ores) are preferentially associated with break-up and dispersal
stages of the cycle, such as Gondwana in the early
Paleozoic and Pangea in the Mesozoic. Other ores
(such as SEDEX Pb–Zn and red-bed Cu deposit
types) exhibit a complementary association interms of the Wilson cycle and are linked withpeak continental amalgamation and stasis. The
biggest concentration of metals on the Earth’s
surface is, however, linked to subduction along
the western margins of Pangea, commencing in
the Mesozoic and extending through the Cenozoic
Era. The present day Pacific Ocean has been, andstill is being, consumed beneath North and SouthAmerica and has given rise to a large variety ofworld-class deposit types, of which the porphyryand epithermal ores are the most prolific. The
skew in the distribution of these and other deposit
types is, however, also related to the enhanced
preservation of young ore bodies that have not yet
been eroded or consumed at a plate margin.

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IndexPage numbers in italics refer to figures;
those in bold refer to tables.
Abitibi greenstone belt, Canada 3,324, 326
abundant metals 5, 13–14
acid hydrolysis 220–1, 233–4,
234
see also hydrolysis
adsorption 159–63, 161, 162
advanced argillic alteration 120–1,
168, 173
Aggeneys–Gamsberg SEDEX deposits,
South Africa 184, 329
Algoma-type banded iron-formation
see banded iron-formation
alkaline magmas 24, 27–8, 28
alluvial diamond 259–61, 259–61
Alpine orogeny, Europe 338
alteration see hydrothermal
alteration
Alum shale, Scandinavia 282
Andean orogeny, South America
334–5, 338–9
Andean-type plate margin 26
andesite 24, 24, 25, 26
anorogenic, ore forming environment
312, 313
anorthosite hosted Ti–Fe deposits
53–4, 330, 342
anthracite 288, 288, 301, 303
Appalachian orogeny, USA 337
aqueo-carbonic hydrothermal fluids
136–8, 137, 138, 189–90
Arabian (Persian) Gulf, Middle East
297, 299–300, 299–300, 332
Archean Eon, metallogeny 321–6,
325
Archean granite–greenstone terranes
190
Arctica 320, 321, 322, 326–30
argillic alteration 172–3
Argyle diamond mine, Western
Australia 32–3

Ashanti–Obuasi gold deposits, Ghana
191, 330
Athabasca tar sands, Canada 297, 304
Atlantica 320, 321, 326–30, 327
atmophile 7–8, 7
atmosphere, evolution 315–16, 316
Au–Ag alloy 228–31
Bajo de la Alumbrera porphyry deposit,
Argentina 89–90, 90
Baltica 320, 321banded iron-formation 268–72, 269,
271, 273–5, 275, 315, 324, 325,
329
basalt 23–4, 24, 24, 26, 28
batch melting see partial melting
processes
Bathurst VMS deposit, Canada 178,
337
bauxite see laterite, bauxite
formation
Bendigo–Ballarat gold deposits,
Victoria 192
BIF see banded iron-formation
Big Stubby VMS deposit, Western
Australia 324
Bingham deposits, Utah 115, 338
biomineralization 163–6, 232, 272,
284
bitterns see evaporite deposits
bitumen 304
bituminous coal 288, 288, 301, 303“black sand” Ti/Zr deposits see
sorting mechanisms, beach/eolian
environments
Black Sea, manganese 276–7, 276,
282
black shales 279–82, 283, 335–7, 336
black smokers 131, 178–84, 179,
180, 183, 185, 341
see also exhalative vents
Bloeddrift diamond mine, South Africa
260
bog iron ore 266–7, 267

boiling 77–8, 78, 82–5, 91–3, 92,
98–100, 99, 103, 121–2, 121,
155–6
borderline metals/ligands 148,
149–51, 206
see also Pearson’s principle
Boulby mine, England 287
bound water 133–5, 134, 136
boundary layer differentiation 46,
47–8, 48
Broken Hill Pb–Zn deposit, Australia
184, 329
Buchans VMS deposits, Canada 178,
337
Burgan oil field, Kuwait 299–300,
299
Burnham model 81–5, 82, 85
Bushveld Complex, South Africa 3,
26, 47–8, 49, 51–2, 57, 59, 61,
63–5, 68, 124–5, 329
calcrete-hosted uranium deposits
235–8, 237
Caledonian orogeny, Scotland 337
Carajas Fe–Cu–Au deposits, Brazil
232, 330
carbon dioxide
magmatic fluid 91, 123phase diagrams 136–8, 137
carbon dioxide–water mixtures
136–8, 137, 138, 189
Carbon Leader Reef, South Africa
264–6, 265
carbonado diamonds 30
carbonates see oxy-salts
carbonatite 27, 340
carbonatization 173
Carlin-type gold deposits 192–5, 193,
194, 338
Carpathian orogeny, Europe 338
catagenesis 288, 290
cation exchange 222–3
cation metasomatism 167
see also metasomatism

INDEX

Central African Copperbelt 187,
198–9, 198, 203–4, 203–4, 331
chalcophile 7–8, 7
chemical sediments 266–87, 340
chemical weathering 220–3, 221
china clay deposits see clay deposits
chromite and PGE concentration
69–71
chromitite layers 49–53
Chuquicamata mine, Chile 104, 241,
243–4, 243–4, 338
Clarion–Clipperton region, Pacific284, 284
classification, ore deposits 2–4, 3
clathrate see gas hydrate
clay deposits 233–5, 235
Cligga Head granite, Cornwall 112
cluster chemistry 68
coal characteristics 301–3, 302, 303
coal deposits 297, 301–4, 303, 335–7,
336
economic viability 303–4, 305
coalification 287–9, 288, 297, 301,
301
Colorado Plateau type uranium
deposits see sandstone-hosted
uranium deposits
common gangue minerals 8–11
common ore minerals 8–11
compressional settings, ore deposits
341–2, 342–3
Comstock lode, Nevada 118, 338
concentration factors 4–5, 4
connate fluids, ore deposits 197–8
connate water 133–5, 131, 134
Conrad discontinuity 21–2
continental crust 22
continental freeboard 279, 318–19,
319
continental growth rates 312–15,
314, 317“copper-oxide” minerals 239, 241–2,
242
Cordilleran orogeny, western USA
36, 334–5, 338–9
Cornish china clay deposits see clay
deposits
Cornubian batholith, Cornwall
110–12, 111, 112, 338
cratonization, metallogeny 324–6,
325
critical point, water 77–8, 91–3
Critical Zone, Bushveld Complex 51,
68
crustal abundances 4–5, 4, 5, 24, 24,
26, 28crustal architecture 20–2, 21, 22, 31
crustal evolution and metallogenesis
4, 315–19
crystal fractionation see
crystallization processes

crystallization processes 37, 40 –6,
44, 47, 48, 49–52, 50, 53–4
crystallization sequence, granites
83–4
Cu–(Mo) porphyry deposits 106–8,
107
Cyprus-type VMS deposits 181–2,
181, 182, 340
Death Valley evaporite deposits,
California 285
degassing see boilingdehydration melting 79–83, 101–4
densityisochore 78, 91–3
variation in magma chambers
44–5
water phases 77–8
depletion of mineral resources 13–14
diadochy see trace element
substitution
diagenetic processes 133–5, 134,
287–90, 288, 290
diamond 28, 30–4, 31
see also alluvial diamond
diatreme–maar volcano 32–3
dilatancy 142–5, 140, 143–4
see also fluid flow
dissolution 146, 146, 220
Domeyko fault 104
Duluth Complex, Minnesota 54, 62
Dwars River location, Bushveld
Complex 52
earth systems science 1
East Pacific Rise/21°N 178, 186, 341
Ecca coal seam, South Africa 303
eclogite 27
eclogite (E-type) diamonds 30–2
economic geology 1
effervescence 155–6
Eh (fO2)–pH diagrams 160, 200, 206,
210, 225, 231, 242
Ehrenfriedersdorf pegmatite, Germany
95–6, 95El Laco volcano, Chile 55
El Salvador mine, Chile 241, 243–4,
244electronegativity 7
eluviation 223, 229, 239
Emmons “reconstructed vein”
174–5, 175
entrainment sorting 252–3, 253
entrapment, oil–gas 296–7, 298
Eoarchean Era see Archean Eon,
metallogeny
epigenetic 6
epithermal 7, 174
epithermal Au–Ag deposits 103,
117–22, 118, 119, 123, 338–9,
341–2

369

equations of state 77–9
eustatic sea-level change see
continental freeboard
evaporite deposits 285–7, 286, 340
exhalative vents 186–9, 188
exinite see coal characteristics“exotic” copper mineralization
241–2, 243–4, 243–4
extensional settings, ore deposits
339–42, 340“fault valve” model 143–5, 144
Fe oxide–Cu–Au deposits 157–8, 158
ferric/ferrous ratio 35–6, 36, 229, 271
filter pressing 53–4
first boiling 93, 98–100
see also boiling
Fletcher mine see Viburnum trend
fluid-absent melting see dehydration
melting
fluid dynamics see placer deposits
fluid flow 138–47, 140
deformation 142–5
evidence of 145–7, 146
fluid inclusion compositions 89–91
fluid inclusions 90, 138
fluid–melt partitioning see partition
coefficient, fluid–melt
fluid mixing 138, 156–9, 158
fluid/rock interaction 159, 166–70,
169
see also alteration
fluid/rock ratio 168–70, 169
FMQ oxygen buffer 31
fossil fuels (coal and petroleum)
287–308
fractional crystallization see
crystallization processes
fractional melting see partial melting
processes
fracture filling processes 146–7
garnet lherzolite 27
gas hydrate 295, 305–7, 306–7
Gas Law 77
geological time scale 11, 12
Ghawar oil field, Saudi Arabia
299–300, 299
Gibbs free energy 175–6
global population trends 11
global production relative to crustal
abundance 5, 13–14, 13
global tectonics and metallogeny 4,
311–44gold laterites see laterites, gold
concentrations
Golden Mile, Kalgoorlie, Western
Australia 171–2, 171, 326
Gondwana 320, 321, 332–5, 333
gossan 239, 239
Gouy layer see membrane filtration

370 INDEX

granitecompositions and metal specificity
35–7
depth of emplacement 101–6, 103,107
fluid flow in and around 108–12,
109, 112
metallogeny 35–7, 36, 37, 101–7,
103, 107water content 80–3, 81, 84, 101–6,
103
granophile deposits 103, 340
Grasberg porphyry deposit, Papua
115, 339
Great Dyke, Zimbabwe 23, 71–3, 329
Great Salt Lake evaporite deposits,
Utah 285
Green River Formation, USA 304
greisenization 111–12, 112, 173
Groote Eylandt manganese deposits,
Australia 276
ground water see meteoric water
Gulf Basin, connate waters 133–5,
134
H2O fluid 82
H2O saturation 93
see also boiling
Hadean Era see Archean Eon,
metallogeny
Haile gold deposit, Carolina 192
halides 9
Hamersley basin, Western Australia
268, 274, 326
Hamersley province see Mount
Whaleback mine
hard acid see Lewis acid/basehard metals/ligands 148, 148–9
see also Pearson’s principle
heat production 316–17, 317
heavy mineral deposits see placer
deposits
Hekla volcano, Iceland 25
hematitization 174
hiatus model, layered mafic intrusions
71
high sulfidation epithermal deposits
see epithermal Au–Ag deposits
Himalayan-type plate margin/orogeny
26, 338, 342–3, 342
Hishikari mine, Japan 120–1, 120,
121, 339
Holland, H. D., experiment 96–8
hotspot 340
see also plume
humic coal see coal characteristics
Huronian basin, placer U deposits
247, 330
hydration 220
hydraulic (settling) equivalence
251–2, 252

hydraulic sorting mechanisms 250–6hydrocarbons 287–9
chemical composition 289
hydrofracturing 83–5hydrogenbonding 76
ion metasomatism see
metasomatism and hydrolysis
hydrolysis 166–7, 220–1, 221
see also acid hydrolysis
hydrostatic pressure 133, 134, 141–2,
141, 296, 296
see also lithostatic pressurehydrothermal alteration 166–74,
169hydrothermal fluids and PGE
mineralization 71, 122–5, 125hydrothermal ore-forming processes
4, 129–215
hydrothermal ore types 3hydroxides 10
hypogene 6, 238, 240
hypothermal 7, 174
I- and S-type granites 27, 35–7, 36,
79–80, 88, 102–4, 111, 113
igneous ore-forming processes 4,
19–74
igneous ore types 3
illuviation 223, 229, 239
immiscibility
magmatic aqueous solution 91–3
pegmatite melt–aqueous fluid
95–6
silicate–oxide 54–5
silicate–sulfide 55–6, 60–8
inertinite see coal characteristics
inheritance factor 28–37
internal zonation
felsic intrusions 45–6
mafic intrusions 43–5
intra/inter-continental craton
formation 323
intra-oceanic shield formation 323
Ir subgroup (Os, Ir, Ru) 58
Irish-type Zn–Pb–Ba deposits 184,
338
ironstone deposits 267–8, 268, 331,
335–7, 336
Irvine model 49–52, 50, 71–3
Isua belt, west Greenland 323–4
Jacobina Basin, Brazil 195–7, 196
Jahns–Burnham model 94–5
see also pegmatite formation
Jamaica, bauxite deposits 225
Jerome VMS deposit, Arizona 178,
330
J.-M. Reef, Stillwater Complex 61,
68, 72–3
João Belo Reef, Jacobina 196, 197

Johan Beetz uranium deposit, Quebec
42
juvenile water see magmatic water
Kaapvaal craton 320, 322, 323
Kalahari manganese field, South Africa
272, 276
Kambalda Ni–Cu mines, Western
Australia 25, 62, 62, 316, 324
kaolinite (china clay) deposits see
clay deposits
Karakoram mountains, Pakistan 338
Kasuga mine, Japan 120–1, 120
kerogen 288, 291–3, 292
Kidd Creek VMS deposit, Canada
178, 324
Kilauea volcano, Hawaii 53, 54, 56
kimberlite 24, 27–8, 28, 30–4, 31, 33,340
King Island mine, Tasmania 113
Kiruna Fe–P deposits, Sweden 55
komatiite 23, 62, 62, 316, 317
Kupferschiefer, central Europe 187,
198–202, 198, 201, 338
Kuroko-type VMS deposits 178, 180,
339, 341
La Escondida, Chile 104–5, 104–5
Lac Tio Ti–Fe deposit, Quebec 54
Lachlan Fold Belt, southeast Australia
35–7, 36
Ladolam Cu–Au deposit 34–5, 34,339
lamproite 30, 32–3
Langer Heinrich uranium deposit,
Namibia 236–7
Larder Lake–Cadillac fault zone,
Canada 191
late veneer hypothesis 28–30, 29
laterite 219, 223–4, 268
bauxite formation 224–7, 225,
227
gold concentrations 228–32, 230,
231
nickel concentrations 227–8, 229
PGE concentrations 231, 232–3
Laurentia 320, 321
layered mafic intrusions,mineralization model 71–3, 73
Lepanto–Far Southeast deposits,
Phillipines 119
Lewis acid/base 147–8, 148ligand 69, 87
lignite 288, 288, 301, 303
Lihir Island see Ladolamliquid immiscibility 54–6, 55, 56,
60–8, 65
lithophile 3, 7–8, 7, 37
lithostatic pressure 133, 134, 141–2,
141
see also hydrostatic pressure

INDEX

lode-gold deposits see orogenic gold
deposits
Lokken VMS deposit, Norway 178,
330
Lorraine-type iron ores see ironstone
deposits
Los Pijiguaos bauxite deposit,
Venezuela 226–7, 226
low sulfidation epithermal deposits
see epithermal Au–Ag deposits
MacTung tungsten skarn, Yukon
114–15, 114–15
magma
contamination 65–8
fertility see inheritance factor
mixing 61–5
types and metal contents 22–8, 24,
24, 26, 28
magmatic-hydrothermal fluid
composition 85–91, 88, 89,90
magmatic-hydrothermal ore-forming
processes 4, 75–126
magmatic water 79–80, 80, 82
magnetite- and ilmenite-series
granites 35, 80, 115
manganese deposits
bedded 272, 276–7, 276
nodular 282–5, 284, 341
mantle convection 317–18, 318
mantle metasomatism 31–2, 31,
34–5, 34
Martha Hill/Waihi gold deposits,
New Zealand 145
Mediterranean Sea, evaporite deposits
285
Meggen–Rammelsberg SEDEX
deposits, Germany 185, 338
membrane filtration 134, 135
Merensky Reef, Bushveld Complex
3, 63–5, 63–4, 68, 69, 71–3,
124–5, 125
Mesoarchean Era see Archean Eon,
metallogeny
Mesoproterozoic Era see Proterozoic
Eon, metallogeny
mesothermal 7, 174, 190
metal–chloride complexes 147–8,
175–6
metal–ligand complexation 96–8, 97,
148, 148–53
metal–organic complexes 152
metal precipitation
biologically mediated 163–6
see also precipitation mechanisms
metal solubility see solubility
metal specificity 35–7, 36, 37
see also I- and S-type granites
metal–sulfide complexes 147–8,
175–6

metal zoning 110–12, 112, 174–7,
182–4, 183, 199–202
see also paragenetic sequence
metallogenic epoch 7
metallogenic province 7
metallogeny 6
through time 313, 319–21, 337–9
metalloids 7
metallotect 6
metals 7
metamorphic water 135–8, 131, 136
metamorphism 2–3, 41, 115–16, 116
metasomatism 31, 64, 71, 115–16,
116, 166–7
see also cation metasomatism
metasomatized mantle see mantle
metasomatism
meteoric fluids, ore deposits 116,
117, 209
meteoric water 132–3, 131, 132, 132
stable isotope compositions 133,
135, 131
meteorite impacts 29, 30, 66–8
methane hydrate see gas hydratemigration, oil–gas see petroleum
migrationmineral production
igneous rocks 20
sedimentary rocks 20
mineral resources 5–6, 6
see also ore reserves
mineralogical barrier 14
mining depths 14
mining and environment 14–15
Mississippi valley type Pb–Zn deposits
202–9
Mo–(Cu) porphyry deposits 106–8,
107
Mohorovicic discontinuity 21–2
Molango manganese deposits, Mexico276
molar volumes and replacement 177
MOMO (mantle overturn major
orogeny) episodes 318, 318
mother-lode gold deposits, California
192, 338
Mount Isa Pb–Zn deposit, Australia
178, 184, 329
Mount Whaleback iron mine, Western
Australia 273–5, 274–5
Muruntau gold mine, Uzbekistan
192
Muskox intrusion, Cananda 60
MVT see Mississippi valley type
Pb–Zn deposits
Naldrett and von Grünewaldt model
71–3
native elements 8–9
nelsonite 54–5
Nena 313, 320, 326–30, 327

371

Neoarchean Era see Archean Eon,
metallogeny
Neoproterozoic Era see Proterozoic
Eon, metallogeny
nephelenite 27
Neves Corvo VMS deposit, Portugal
178, 338
New Albany shale, USA 282
New Caledonia, nickel laterites 228,
229
nickel laterite see laterite, nickel
concentrations
non-metals 7
Noril’sk–Talnakh Cu–Ni–PGE deposit
62, 72
North Sea, gas fields 297, 302
oceanic crust 21
oil shales 304–5
oil window 288, 288, 291
oil–gas formation 288, 289–91, 290,
291
oil–gas traps see entrapment, oil–gas
Olympic Dam, South Australia
157–8, 157, 158, 329
oolitic ironstones see ironstone
deposits
Orange River diamond deposits
259–61, 259–61
Orapa diamond mine, Botswana
32–3, 33
ore 6
ore deposition see precipitation
mechanisms
ore deposits
classification see classification, ore
deposits
viability 4–5, 4
ore reserves 5–6, 6
see also mineral resources
Oregon, heavy mineral beach placers
262–3
Orinoco basin, overpressure and fluid
flow 296, 296
orogenic gold deposits 3, 190–7, 190,
312, 342
orogenic ore-forming environment
312, 313
orogeny driven fluid flow 140, 204–5,
205
see also fluid flow
overpressure see hydrostatic pressureoxidation/reduction 159, 160,
199–201, 206–7, 209–14, 221–2,
275, 229–31, 239
oxides 7, 9–10
oxyatmoinversion see atmosphere,
evolutionoxyhydroxides 10
see also hydroxides
oxy-salts 10

372 INDEX

Paleoarchean Era see Archean Eon
Paleoproterozoic Era see Proterozoic
Eon, metallogeny
Pangea 312, 313, 319, 320, 321,
332–9, 333, 336
paragenesis 174
paragenetic sequence 174–7, 175
partial melting processes 37–40, 39,
41–2, 41
partition coefficient
crystal–melt 40–2, 46–9, 100, 107
fluid–melt 96–101, 98, 99, 101,
106–8
sulfide–silicate 58, 58, 60
partitioning behavior, Cu, Mo, W
100–1
Pd subgroup (Pt, Pd, Rh) 58
Pearson’s principle 147
peat formation 288, 288, 297, 301,
301, 303
pedogenesis 219, 223–4, 223, 235–6,
236, 268
pegmatiteclassification 93–4
equilibrium with aqueous solution
87–9
formation 93–6
percolation threshold 145
see also permeability
peridotite (P-type) diamonds 30–2
periodic table 7–8, 7
permeability 139, 145–7, 145
Persian Gulf see Arabian Gulf
petroleumdeposits 335–7, 336
migration 293–6, 295–6
petroleum–coal compositions 289,
303
PGE see platinum group elements
PGE clusters 68–71, 70
Phalaborwa Complex, South Africa
27, 329
Phanerozoic Eon, metallogeny
332–9, 336
phase separation 155–6
see also boiling; effervescencephosphates see oxy-salts
phosphogenesis, modern day 279,
281
Phosphoria Formation, USA 279
phosphorites 277–9, 280, 281, 331
phyllic alteration 172
Pilbara craton 320, 322, 323–4
Pine Point Pb–Zn mine, Canada 202,
207
placer deposits 247– 66
beaches 261–3, 262, 263
numerical simulation 264–6, 265
wind-borne 263–4, 264
plagioclase, as a cumulus phase 72–3
Plat Reef, Bushveld Complex 65

plate tectonics, ore deposits 339–43,
340–2
Platinova Reefs, Skaergaard 44–5, 49
platinum group elements (PGE) 29,
34, 58, 231, 232
plumes 23, 31, 332, 334, 336, 337
podiform chromitite 53, 340, 341
pore fluid 133, 136, 295
pressure see hydrostatic pressure
porphyry deposits 103, 104–5,
104–5, 106–8, 107, 119, 338, 341,
342
porphyry–epithermal link 118–19,
119
potassic alteration 170, 173
potholes, formation andcharacteristics 64–5, 64, 124–5,
125precipitation mechanisms 153–66,
199–202, 206–7, 211–14
gold 119–22
primary migration, oil–gas see
petroleum migration
production trends with time (oil,bauxite, Cu, Au) 13
prograde skarn reactions 115–17
propylitic alteration 168, 172
Proterozoic Eon, metallogeny
326–32
Quadrilatero Ferrifero, Brazil 268
quartz pebble conglomerate hosted
gold deposits 195–7, 196
quartz veins 85–7
R factor 59–60, 60
Rapitan-type banded iron-formation
see banded iron-formation
rare earth elements 8
Rayleigh fractionation see
crystallization processes
red-bed copper deposits see
stratiform sediment-hosted copper
deposits
Red Dog mine, Alaska 188–9, 188,
189
Red Sea geothermal system 187–9,
340, 341
redox barrier/front/interface 213,
239, 267
reduction see oxidation/reduction
replacement processes 176–7
retrograde skarn reactions 117
Reynolds number 248–50, 253
rhyolite 24, 25–7, 24, 26
Richard’s Bay Ti–Zr deposits, South
Africa 261, 264
Rio Tinto VMS deposit, Spain 178,
338
Rodinia 312, 313, 319, 320, 321,
326–32, 328

roll-front type uranium deposits see
sandstone-hosted uranium
deposits
Rössing uranium mine, Namibia
41–2, 41
S- and I-type granites 27, 35–7,
79–80, 94, 100–8, 111, 113, 115,
338
Sabie–Pilgrim’s Rest gold deposits,
South Africa 192
St Agnes granite, Cornwall 112
St Francois mountains, Missouri 207,
208, 330
salt filtration see membrane
filtration
Salton Sea geothermal system,
California 184–7, 186
sandstone-hosted uranium deposits
210–14, 212, 213sapropelic coal see coalcharacteristics
saturation vapor pressure 77
scarce metals 5, 13–14
sea water 130–2, 131, 132, 132, 185second boiling 93, 98
see also boilingsecondary migration, oil–gas see
petroleum migration
SEDEX see sedimentary exhalative
deposits
sediment transport 248–50, 249
sedimentary exhalative deposits 159,
177–89, 335–7, 336, 340
sedimentary ore-forming processes 4,
246–308
sedimentary ore types 3
seismic pumping model 142–3,
143
Serra Palada deposit, Brazil 232–3
settling sorting 250–2, 251
shear sorting 253–4, 254
shield formation, metallogeny
324–6, 325
Shield’s parameter 252–3, 253
siderophile 7–8, 7, 29
sidewall boundary layer differentiation
45–6
see also boundary layer
differentiation
silicate–oxide immiscibility see
immiscibility
silicate–sulfide immiscibility see
immiscibility
silicate/sulfide liquid mass ratio see
R factorsilicates 10–11
silication 173
silicification 173, 121
Skaergaard Complex, Greenland
43–5, 44, 49, 58

INDEXskarn deposits 113–17, 113–16,
341–2, 339
Skorpion mine, Namibia 242
SMOW see Standard Mean Ocean
Water“Snowball Earth” 331soft acid see Lewis acid/basesoft metals/ligands 148, 151
see also Pearson’s principle
solidification see crystallization
processes
solubility
apatite in sea water 277–8, 278
metals in aqueous solutions 96–8,
97, 147–53, 154, 182–3, 183
PGE in aqueous solutions 123–4
quartz in aqueous solutions 85–7,
146
Si and Al in aqueous solutions 222,
224–5
sulfur/sulfide in magma 56–8, 57,
60–8, 72–3
water in magma 79–81
sorting mechanisms
application to placer deposits
255–8, 257
beach/eolian environments
258–64, 262, 263, 264
source considerations, oil–gas 291–3,
293
Sr isotope variations, layered mafic
intrusions 63–5
SSC see stratiform sediment-hosted
copper deposits
Standard Mean Ocean Water (SMOW)
131, 135
Stillwater Complex, Montana 53, 61,
71–3
Stoke’s Law 250–1, 251
stratiform sediment-hosted copper
deposits 159, 197–202, 198,
335–7, 336
Strong’s model 102–3
structural water see bound water
sublayer, Sudbury 66–8
“suction pump” model 144–5, 144
Sudbury Ni–Cu deposits 66–8, 66–7
sulfates see oxy-salts
sulfide–silicate partition coefficients
58–9, 58, 60
sulfides 7, 9
sulfo-salts 9superchron event 32
supercontinent cycle 317, 319,
319–21, 320

supercritical fluid 77–8
supergene 6, 238
blanket 238, 240, 243
copper 238 –45, 239, 240, 242
Fe 273–5, 275
processes see surficial ore-forming
processes
Zn, Ag 242–5
Superior craton/province 322, 323–4
Superior-type banded iron-formation
see banded iron-formation
superplumes see plumes
surficial ore-forming processes 4,
219–45
surficial uranium deposits see
calcrete-hosted uranium deposits
sustainable development 11–14
Sverdlovsk skarn deposit, Russia 115
syngenetic 6, 246
Tanco pegmatite, Manitoba 95
tar sands 304–5
Tarkwa conglomerates 197
Taupo volcanic zone, New Zealand117
tectonic cycles, metallogeny 335–7
Telfer gold deposit, Western Australia
192
Tellnes Ti–Fe deposit, Norway 54
Tethys ocean, metallogeny 279–80,
280, 299, 337
thermally driven fluid flow see fluid
flow
torbanite see oil shales
trace element distribution
fractional crystallization 46–9
partial melting 40–2
trace element substitution 23
transition zone, mantle 30–1transport sorting 254–5, 256
Transvaal basin, manganese deposits
see Kalahari manganese field,
South Africa
triple point, water 78
Troodos complex, Cyprus 178,
181–2, 181–2, 338
tuffaceous kimberlite 33
tungstates see oxy-salts
tungsten-type porphyry deposits 107,108, 339
Twin Creeks mine, Nevada 194–5,
194
Udokan–Dzhezhkazgan copper
deposits, Kazakhstan 198–9

373

UG1 chromitite seam, Bushveld
Complex 51–2, 51–2
UG2 chromitite seam, Bushveld
Complex 69, 124
upwelling model (iron, phosphorus)
271–2, 271, 277–9, 280, 281
Ur 320, 322, 329
vanadates see oxy-salts
vapor phase 82
vapor saturation 93
see also boiling
Variscan orogeny, Europe 337–8
Vazante mine, Brazil 242
viable ore deposit see ore deposits,
viability
Viburnum Trend, Missouri 202,
207–9, 208, 209
vitrinite see coal characteristics
VMS see volcanogenic massive
sulfide deposits
VMS–SEDEX continuum model 178,
187–9
volatile phase 82
volcanogenic massive sulfide deposits
177–89, 335–7, 336, 340
waterphase diagram 78
physical and chemical properties
76–9, 76, 78
water–carbon dioxide 136–8, 137,
189
weathering see chemical weathering
Welkom goldfield, South Africa 257
Western Deep Levels gold mine, South
Africa 20
White Pine copper deposit, Michigan
198, 330
white smokers 182
see also exhalative vents
Wilson cycle 21, 318, 319, 332,
335–7, 336
Witwatersrand Basin, South Africa 3,
195–7, 196, 257, 258, 315, 264–6,
265, 325, 326–9
Yeelirrie uranium deposit, Western
Australia 236–8, 237
Yilgarn craton 322, 323–4
gold laterites 228–9
Zaaiplaats tin mine, South Africa
47–8, 48
Zimbabwe craton 322, 323–4