Geology and Geochemistry of Epithermal Systems - Reviews Vol 2-Search

REVIEWS IN ECONOMIC GEOLOGY Volume 2

GEOLOGY AND GEOCHEMISTRY OF EPITHERMAL SYSTEMS B. R. Berger a P. M. Bethke, Editors CONTENTS THEGEOTHERMAL

R . W. Henley

FRAMEWORK FOR EPITHERMAL DEPOSITS

A PRACTICAL GUIDE TO THE THERMODYNAMICS OF GEOTHERMAL FLUIDS AND HYDROTHERMAL ORE DEPOSITS

THEBEHAVIOR OF SILICA IN HYDROTHERMAL SOLUTIONS CARBONATE TRANSPORT AND DEPOSITION IN THE EPITHERMAL FLUID INCLUSION SYSTEMATICS IN EPITHERMAL SYSTEMS LIGHTSTABLE-ISOTOPE SYSTEMATICS IN THE EPITHERMAL

R . W. Henley & K . L . Brown R . 0 . Fournier ENVIRONMENT R . 0 . Fournier R . J . Bodnar, T. J . Reynolds, & C . A . Kuehn

C. W. Field & R . H . Fifarek

ENVIRONMENT

GEOLOGIC, MINERALOGIC,

AND GEOCHEMICAL CHARACTERISTICS OF VOLCANIC-HOSTED EPITHERMAL PRECIOUS-METAL D . 0 . Hayba, DEPOSITS

P. M . Bethke, P. Heald, & N . K . Foley

GEOLOGIC CHARACTERISTICS

OF SEDIMENT-HOSTED, DISSEMINATED PRECIOUS-METAL DEPOSITS IN THE WESTERNUNITEDSTATES

W. C . Bagby & B . R . Berger

RELATIONSHIP OF TRACE-ELEMENT

PATTERNS TO ALTERATION AND MORPHOLOGY IN EPITHERMAL PRECIOUS-METAL DEPOSITS

RELATIONSHIPS OF TRACE-ELEMENT

M . L . Silberman & B . R . Berger

PATTERNS TO GEOLOGY IN HOT-SPRING

TYPE PRECIOUS-METAL DEPOSITS

B . R . Berger & M . L. Silberman

BOILING,COOLING,

AND OXIDATION IN EPITHERMAL SYSTEMS: A NUMERICAL MODELING APPROACH

M . H . Reed & N. Spycher

USINGGEOLOGICAL

INFORMATION TO DEVELOP EXPLORATION STRATEGIES FOR EPITHERMAL DEPOSITS

Series Editor: James M. Robertson

SOCIETY OF ECONOMIC GEOLOGISTS

S . S . Adams

REVIEWS IN ECONOMIC GEOLOGY (ISSN 0741-0123) Published Annually by the SOCIETY OF E C O N O M I C GEOLOGISTS Printed by BookCrafters, Inc., 140 Buchanan Street, Chelsea, MI Series Editor: James M . Robertson

481 18

Additional copies of this volume may be obtained from: The Economic Geology Publishing Company P.O. B o x 637 University of Texas at E l Paso El Paso, T X 79968-0637 U S A (915) 533-1965

Vol. 1: FLUID-MINERAL EQUILIBRIA IN HYDROTHERMAL SYSTEMS (1984) Vol . 2: GEOLOGY AND GEOCHEMISTRY OF EPITHERMAL SYSTEMS ( 1985)

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ISBN 0-9613074-0-4 ISBN 0-961 3074- 1-2

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REVIEWS IN ECONOMIC GEOLOGY (ISSN 074141123)

Volume 2

GEOLOGY AND GEOCHEMISTRY OF EPITHERMAL SYSTEMS ISBN 0-9613074-1-2

Volume Editors: B. R. BERGER Branch of Exploraton Geochemistry U. S . Geological Survey MS 973 Box 25046, Federal Center Denver, CO 80225-0046

P. M. BETHKE Branch of Resource Analysis U. S . Geological Survey MS 959, National Center Reston, VA 22092

Series Editor: JAMES M . ROBERTSON New Mexico Bureau of Mines & Mineral Resources Campus Station Socorro, NM 87801

SOCIETY OF ECONOMIC GEOLOGISTS

The Authors: Samuel S. Adams 3030 Third Street Boulder, CO 80302 William C. Bagby Branch of Western Mineral Resources U. S . Geological Survey MS 901 345 Middlefield Road Menlo Park, CA 94025 B. R. Berger Branch of Exploration Geochemistry U.S. Geological Survey MS 973 Box 25046, Federal Center Denver, CO 80225-0046 Philip M. Bethke Branch of Resource Analysis U.S. Geological Survey MS 959, National Center Reston, VA 22092 R. J. Bodnar Department of Geological Sciences Virginia Polytechnic Institute and State University Blacksburg, VA 2046 1

Robert 0 . Fournier Branch of Igneous and Geothermal Processes U.S Geological Survey MS 910 345 Middlefield Road Menlo Park, CA 94025 Daniel 0 . Hayba Branch of Resource Analysis U.S. Geological Survey MS 959, National Center Reston, VA 22092 Pamela Heald Branch of Resource Analysis U.S. Geological Survey MS 959, National Center Reston. VA 22092 R. W. Henley Chemistry Divsion D.S.I.R.. Private Bag Taupo New Zealand C. A. Kuehn Department of Geosciences The Pennsylvania State University University Park, PA 16802

K. L. Brown Chemistry Division D.S.I.R., Private Bag Taupo New Zealand

Mark H. Reed Department of Geology University of Oregon Eugene, OR 97403

Cyrus W. Field Department of Geology Oregon State University Corvallis, OR 9733 1-5506

T. J. Reynolds FLUID, Inc. P.O. Box 6873 Denver, CO 80206

Richard H. Fifarek Department of Geology Southern Illinois University Carbondale, IL 62901

M. L. Silberman Branch of Exploration Geochemistry U. S . Geological Survey MS 912 Box 25046, Federal Center Denver, CO 80225-0046

N. K. Foley Branch of Resource Analysis U .S. Geological Survey MS 959, National Center Reston, VA 22092

N. Spycher Department of Geology University of Oregon Eugene, OR 97403

GEOLOGY & GEOCHEMISTRY OF EPITHERMAL SYSTEMS CONTENTS

F O R E W O R D . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

PREFACE

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

x

xi

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . x v i

BIOGRAPHIES

CHAPTER 1 THE GEOTHERMAL FRAMEWORK OF EP1THERMA.L DEPOSITS R . W . Henley

...................................... HYDROTHERMAL SYSTEMS IN GENERAL . . . . . . . . . . . . . . . . . . . . . . . . . . . . . INTRODUCTION.

1 1

Collision-Related Amagmatic Hydrothermal Systems Terrestrial Magma-Related Hydrothermal Systems TERRESTRIAL MAGMATIC-HYDROTHERMAL SYSTEMS

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

4

Laree Scale Structure -. Natural Discharges Hydrothermal Eruption Vents Heat and Mass Flow in Geothermal Systems -

-

-

........................... EPITHERMAL ORE-FORMING SYSTEMS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . CHEMISTRYOFGEOTHERMALDISCHARGES.

11 12

Requirememts for Ore Deposition Chemistry of Systems Responsible for Ore Formation Chemical and Physical Processes in Ore Formation Host-Rock Relations

......................................... EPILOGUE. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ACKNOWLEDGMENTS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . SUMMARY

19 21 21 21

CHAPTER 2 A PRACTICAL GUIDE TO THE THERMODYNAMICS OF GEOTHERMAL FLUIDS AND HYDR0THERMA.L ORE DEPOSITS R . W . Henley and K . L . Brown

..................................... GEOLOGICAL CHARACTERISTICS OF THE BROADLANDS GEOTHERMAL SYSTEM . . . . . . . . . . . . . . FLUID CHEMISTRY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . INTRODUCTION..

25 25 26

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

FLUID-MINERAL EQUILIBRIA: ALTERATION MINERALOGY FLUID-MINERAL EQUILIBRIA: TRACE-METAL CONTENTS Lead . Gold . Other Metals:

28 32

Copper, Silver, and Arsenic

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

MINERAL DEPOSITION

36

Silica Calcite Metal Sulfides and Gold

..................................... REVIEW QUESTIONS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

41

APPENDIX

43

ACKNOWLEDGMENTS

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41 41

CHAPTER 3 TEE BEHAVIOR OF SILICA IN EYDROTEERMAL SOLUTIONS

R

. 0 . Fournier

....................................... SOLUBILITIES OF SILICA MINERALS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . THE BEHAVIOR OF DISSOLVED SILICA IN HOT-SPRING SYSTEMS . . . . . . . . . . . . . . . . . . ALKALINE WATERS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ACIDWATERS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . REACTIONWITHGLASS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . AMORPHOUS SILICA-CHALCEDONY RELATIONS . . . . . . . . . . . . . . . . . . . . . . . . . . SPECULATIONS REGARDING SOME TEXTURES OF QUARTZ . . . . . . . . . . . . . . . . . . . . . . INTRODUCTION

45 45 46 48 50 51 51 51

Jasperoid and Massive Replacement of Limestone by Silica Quartz Solubility at High Temperatures

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55 ACKNOWLEDGMENTS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56 REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56 APPENDIX . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 CONCLUSIONS

CARBONATE TRANSPORT AND DEPOSITION IN TEE EPITBERMAL ENVIRONMENT R. 0 . Foumier

...................................... C02 DISSOLVED IN AQUEOUS SOLUTIONS . . . . . . . . . . . . . . . . . . . . . . . . . . . . THE SOLUBILITY OF CALCITE IN AQUEOUS SOLUTIONS . . . . . . . . . . . . . . . . . . . . . . SUMMARY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . INTRODUCTION.

63 63 67 71

FLUID-INCLUSION SYSTEMATICS IN EPITEERMAL SYSTEMS R. J. Bodnar, T . J . Re,ynoZds, and C . A. Kuehn

................................ ...... INFORMATION AVAILABLE FROM FLUID-INCLUSION PETROGRAPHY . . . . . . . . . . . . . . . . . .

INTRODUCTION.

............ IDENTIFICATION OF GASES IN FLUID INCLUSIONS FROM THE EPITHERMAL ENVIRONMENT . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . INTERPRETATION OF FLUID INCLUSIONS FROM THE EPITHERMAL ENVIRONMENT . . . . . . . . . . . . APPLICATION OF FLUID INCLUSIONS IN EXPLORATION FOR EPITHERMAL PRECIOUS-METAL DEPOSITS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . SUGGESTIONS FOR FUTURE FLUID-INCLUSION RESEARCH . . . . . . . . . . . . . . . . . . . . . REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . IDENTIFICATION OF FLUID INCLUSIONS TRAPPED FROM BOILING SOLUTIONS

73 73 79

83

93

94

95 96

CHAPTER 6 LIGar STABLE-ISOTOPE SYSTEMATICS IN THE EPITHERMAL ENVIRONMENT C . W . Field and R. H . Fifarek

...................................... CONVENTIONS, SYSTEMATICS, AND RATIONALE . . . . . . . . . . . . . . . . . . . . . . . . . INTRODUCTION.

99 99

Fractionation Equilibrium Reaction Applications GEOLOGIC DISTRIBUTIONS

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

110

Hydrogen and Oxygen Carbon Sulfur EPITHERMALDEPOSITS

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 113

Hydrogen and Oxygen

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 124 REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 125 SUMMARY

GEOLOGIC, MINERALOGIC, AND GEOCHEMICAL CHARACTERISTICS OF VOLCANIGEOSTED EPITBERMAL PRECIOUS-METAL DEPOSITS D. 0 . Hayba, P. M . Bethke, P . Heald, and 1. K. Foley

SUMMARY OF THE CHARACTERISTICS OF VOLCANIC-HOSTED EPITHERMAL ORE DEPOSITS

. . . . . . . . 129

Characteristics of Adularia-Sericite-Type Deposits Characteristics of Acid-Sulfate-Type Deposits Summary of Characteristics THE ADULARIA-SERICITE ENVIRONMENT:

CREEDE AS AN EXAMPLE

. . . . . . . . . . . . . . . . . 136

Creede as an Exemplar Summary of Important Studies Geologic and Mineralogic Characteristics Geochemical Environment Hydrologic Environment Boiling and Mixing in the Ore Zone Summary of Creede Mineralization THE ACID-SULFATE ENVIRONMENT:

SUMMITVILLE AS AN EXAMPLE

. . . . . . . . . . . . . . . . . 151

Geologic and Yineralogic Characteristics Geochemical Environment - - - Summary of Summitville Mineralization ~

~

-

GEOTHERMAL INTERPRETATION OF VOLCANIC-HOSTED EPITHERMAL DEPOSITS

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

158

Adularia-Sericite Deposits Acid-Sulfate Deposits

. . . . . . . . . . . . . . . . . . . . . . . . . . 159 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 162

MECHANISMS OF ACID-SULFATE ALTERATION ACKNOWLEDGMENTS

CHAPTER 8 GEOLOGIC CHARACTERISTICS OF SEDIMENT-HOSTED, DISSEMINATED PRECIOUS-METAL DEPOSITS IN THE WESTERN UNITED STATES W. C. Bagby and B. R. Berger

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 169 CLASSIFICATION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 169 REGIONAL GEOLOGIC CHARACTERISTICS OF DEPOSITS IN MINERAL TRENDS AND ISOLATED DEPOSITS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 172 INTRODUCTION.

The Getchell Trend The Carlin Trend

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 228 REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 230

ACKNOWLEDGMENTS

CHAPTER 1 0 RELATIONSHIPS OF TRACE-ELEMENT PATTERNS TO GEOLOGY IN HOT-SPRINGTYPE PRECIOUS-METAL DEPOSITS B . R. Berger and M . L . SiZberman

. . . . . . . . . . . . . . . . . . . . . . . . . . . . 233 TRACE-ELEMENT PATTERNS IN STUDIED DEPOSITS . . . . . . . . . . . . . . . . . . . . . . . . 235 CONTROLS ON TRACE-ELEMENT PATTERNS

Hasbrouck Mountain, Nevada Round Mountain, Nevada DISCUSSION.

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 245 CHAPTER 11 BOILING, COOLING, AND OXIDATION IN EPITAERMAL SYSTEMS: A NIlMERICAL MODELING APPROACH M . H. Reed and N . F . Spycher

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 249 BOILING RESULTS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 252 DISCUSSION OF BOILING AND COOLING . . . . . . . . . . . . . . . . . . . . . . . . . . . . 252 BOILING

Sulfide and Carbonate Mineral Precipitation Precipitation of Silicates Boiling Without Fractionation and Cooling Only

. . . . . . . . . . . . . . . . . . . . . . . . . . . 258 BOILING AND GOLD PRECIPITATION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 261 THE HOT-SPRING ENVIRONMENT . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 262 SUPER- AND SUB-ISOENTHALPIC BOILING

Condensation of the Boiled Gas Oxidation of Gases to Produce Acid-Sulfate Waters Reaction of Gases with Meteoric Ground Water Gold Precipitation from Mixing of Acid-Sulfate Water with Boiled Aqueous Phase Gold Precipitation from Mixing of Oxygenated Ground Water with Boiled Aaueous Phase

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 269 ACKNOWLEDGMENTS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 270 SUMMARY

CHAPTER 12 USING GEOLOGICAL INFORMATION TO DEVELOP EXPLORATION STRATEGIES FOR EPITEIERMAL DEPOSITS S . S. Mums

....................................... SOME CONSIDERATIONS IN THE USE OF GEOLOGICAL INFORMATION IN EXPLORATION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . STRATEGIC FACTORS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

INTRODUCTION

273

273 274

Organizational Objectives Commodity Prices Financial Resources Exploration Organization Regulations and Land Availability Competitor Activity Previous Ex~loration

Risk HUMAN FACTORS

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

279

Personal Objectives Education and Training Problem Solving Intuition and Creativity Uncertainty Aversion to Loss DEVELOPMENT OF MINERAL-DEPOSIT MODELS

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

282

Organization of Geologic Information Model Terminology Level of Model Development DATA-PROCESS-CRITERIA MODEL

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

Definition of a Mineral-Deposit Type Compilation of Analog Deposits Selection of Geologic Data Data-Process Linking Identification of Formation Processes Evaluation of Data-Process Links Selection of Diagnostic Criteria Evaluation of Data-Process-Criteria Model Application of Data-Process-Criteria Model Summary of Data-Process-Criteria Model

286

Exploration

....................................... REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . TABLE OF CONVERSION FACTORS . . . . . . . . . . . . . . . . . . . . . . . . . . . Inside Back Cover CONCLUSIONS

296 297

FOREWORD

Geology and Geochemistry of Epithermal 2 of Reviews Economic Systems--Volume Geology--was c r e a t e d t o accompany a Society of Economic Geologists (SEG) short course of t h e s a m e n a m e t h a t was given in October, 1985, prior t o t h e annual meetings of t h e Geological Society of America and Associated Societies in Orlando, Florida. As was t h e c a s e with Volume 1, t h e final published version of Volume 2 unfortunately postdates t h e short course by s o m e months. Geology and Geochemistry of Epithermal Systems presents a synthesis of t h e current understanding of t h e processes responsible for t h e concentration of m e t a l s (especially gold and silver) in near-surface environments, provides an overview of t h e systematics of t h e most important approaches t o t h e study of epithermal ores and processes, and summarizes t h e geology of both sediment-hosted and volcanic-hosted epithermal precious-metal deposits. A f t e r t h e volume editors, t h e most significant contributors t o t h e production of this voIume were t h e members of t h e Editorial Support Group, Branch of

Exploration Geochemistry, U.S. Geological Survey, Denver, Colorado. These ladies, Marilyn A. Billone, Candace A. Vassalluzzo, and especially P a m e l a S. D e t r a and Dorothy B. Wesson, accomplished t h e long, arduous, and often frustrating job of assembling, editing, and formatting t h e book with a uniformly high level of professionalism and good cheer. Their e f f o r t s a r e gratefully acknowledged. Carol Hjellming of t h e New Mexico Bureau of Mines and Mineral Resources (NMBMMR) editing staff checked, balanced, and helped interpret t h e chemical equations; Lynne McNeil (NMBMMR) f o r m a t t e d t h e cutlines. Lastly, I wish t o express my continuing appreciation t o t h e New Mexico Bureau of Mines and Mineral Resources and i t s Director, Frank Kottlowski, who provide t h e Series Editor with time, space, and encouragement. J a m e s M. Robertson Series Editor Socorro, NM March, 1986

PREFACE

In a speech on May 10, 1911, before t h e Geological Society of Washington, Waldemar Lindgren described his systematic classification of all types of mineral deposits. One of his categories included deposits related t o intrusive and eruptive igneous rocks t h a t form veins a t shallow depths t h a t contain opencavity filling t e x t u r e s and t h a t have been a primary source of "bonanza" grades of gold and silver--the epithermal deposits. Historically, most of t h e ores in epithermal systems have been mined from quartz veins, breccias, or disseminations t h a t a r e associated with non-marine volcanic rocks. Open-space filling textures and structures a r e common--comb structure, crustification, symmetrical banding, and crystal-lined vugs. Ore minerals include native gold, native silver, electrum, argentite, sulfosalts, tellurides, and selenides and often t h e common sulfides sphalerite, galena, and chalcopyrite. Common gangue minerals a r e quartz, adularia, calcite, barite, rhodochrosite, and fluorite. Alteration is commonly widespread in epithermal systems, particularly in t h e upper portions of the vein systems; among the alteration phases a r e quartz, adularia, illite, chlorite, alunite, and kaolinite. Lindgren (1928) recognized the difficulty of developing a rigid subsidiary classification s c h e m e for epithermal deposits; he separated them into six categories:

I. 2. 3. 4. 5. 6.

Cold deposits Argentite-gold deposits Argentite deposits Gold selenide deposits Gold telluride deposits Gold telluride deposits with alunite Nolan (1933) and Ferguson (1929) f e l t t h a t few of these six characteristics were restricted enough t o be diagnostic and proposed only t w o classes of epithermal systems based on t h e weight ratio of gold t o silver, silver-gold, and gold-silver. Based on his experience with deposits in Nevada, Ferguson (1929) found t h a t t h e r e is a bimodal distribution of gold-silver ratios, and Nolan (1933) f e l t t h a t t h e bimodality was due t o genetic processes. For t h e silver-gold deposits, Nolan (1933) noticed t h a t through-going fault fissures control t h e o r e and f e l t t h a t this implies a deep origin for t h e source of t h e metals. Nolan (1933) also noted t h a t t h e preciousm e t a l ores a r e very commonly sharply limited above and below by approximately parallel surfaces referred t o a s t h e o r e "horizon." He suggested t h a t these limits a r e related t o temperature. Base metals tend t o increase a t and below t h e base of t h e lower surface of t h e precious-metal ore. Figure 1 is a longitudinal, vertical projection of t h e Last Chance-Confidence silver-gold vein in t h e Mogollon mining district, New

n

Confidence

Last Chance

700-FT LEVEL 500

1000 FEET

900-FT LEVEL

Figure P.1. Vertical, longitudinal projection of the Confidence-Last Chance vein in the Mogollon mining district, New Mexico (Ferguson, 1927). Banded quartz vein is continuous along strike with ore grade material occurring in specific masses (stippled areas) in the vein. The tops and bttoms of the silver-rich ore bodies describe near parallel surfaces referred to as the "ore horizon."

Mexico (Ferguson, 1927) illustrating t h e o r e horizons, t h e shape of o r e bodies, a n d t h e typical distribution of o r e grades within a continuous banded quartz-adulariaBurbank (1933) reported t h a t base sericite vein. m e t a l s appear t o b e m o r e abundant in silver-gold deposits in regions of sedimentary rocks with overlying volcanic rocks and in thick, volcanic sequences with a long history of volcanic activity. In contrast t o t h e silver-gold deposits, Nolan (1933) noted t h a t goldsilver deposits a r e commonly within or close t o small, shallow intrusive bodies a n d t h a t t h e ore-controlling f r a c t u r e s y s t e m s a r e frequently more discontinuous than those associated with silver-gold deposits. The gold-silver o r e s a r e also m o r e irregular in distribution than t h e silver-gold ores. Nolan f e l t t h a t this irregularity may be r e l a t e d t o t h e complex thermal regimes in t h e s e types of systems due t o t h e shallow intrusive activity. Figure 2a shows a series of plan views of t h e January mine, Goldfield mining distrct,

Nevada and a cross section through t h e January s h a f t (Ransome, 1909) showing t h e relationships of o r e t o quartz-alunite-kaolinite replaced wallrock ("ledge matter") and t h e host rocks. Figure 2b shows t w o cross sections from Ransome (1909,. p. 154) of t h e Combination mine in Goldfield ~ l l u s t r a t i n g t h e irregular vertical distribution of bonanza-grade o r e masses within t h e "ledge matter." Also, t h e ore bodies were not persistent along strike. Although Waldemar Lindgren (1928) recognized t h e correlation between epithermal systems and a c t i v e geothermal systems, i t was Donald E. White (1955, 1981) who championed the detailed study of a c t i v e systems and t h e application of t h e results and concepts derived from t h e s e studies t o epithermal ore deposits. The i m p a c t of White's leadership in t h e study of hydrothermal systems, in general, and epithermal systems, in particular, was recognized by t h e Society of Economic Geologists when i t held a symposium in

JANUARY SHAFT

a

JANUARY SHAFT

109-FT L E V E L 51-FT LEVEL

8 I-FT L E V E L

232-FT LEVEL

283-FT LEVEL CROSS SECTION

160-FT LEVEL

232-FT LEVEL

0

50

I

I

IqO

150 FEET I

283-FT LEVEL

Figure P.2. a). Plan views of the January mine at selected mining levels and a cross section through the January shaft Goldfield mining district, Nevada (Ransome, 1909). Bonanza-grade gold ores occur in replaced dacite referred to as "ledge matter." The ore bodies are not persistent either downdip or along strike, and occur both on the hanging wall side of the ledge and on the foot wall side. b). Two cross sections from Ransome (1909, p. 154) of the Combination mine in the Goldfield district. Sonanza-grade gold ores occur in irregular, discontinuous masses within the ledge. The ledge follows a lithologic contact and flattens with depth.

80-Ft L e v e l 130-Ft L e v e l 180-Ft L e v e l 230-Ft L e v e l 280-Ft L e v e l

380-Ft L e v e l

xii

his honor in February, 1984 entitled: Geothermal Systerns and Ore Deposits. It clearly emphasized t h e value of using active geothermal a r e a s a s models of fossil, ore-forming hydrothermal systems. Thus, the evolution of understanding of t h e geology and genesis of epithermal precious-metal deposits has followed a pathway from t h e early, vividly descriptive studies of mining districts such a s t h e Comstock Lode, Nevada (Becker, 18821, Cripple Creek, Colorado (Lindgren and Ransome, 19061, and Waihi, New Zealand (Bell and Fraser, 1912) t o t h e later, topical studies on s t r u c t u r e (Wisser, 1960), alteration (Hernley and Jones, 19641, stable isotopes (Taylor, 19731, and fluid chemistry (Barton e t al., 1977). The most recent research on epithermal deposits has built on these past studies and has emphasized the thermal and compositional roles of volcanic rock terranes; t h e genesis, significance, and pattern of alteration mineralogies; t h e sources of t h e geothermal fluids and t h e paleohydrology of t h e systems; and, t h e chemical conditions surrounding t h e deposition of t h e o r e minerals. The present volume is a n a t t e m p t t o provide a synthesis of t h e c u r r e n t s t a t e of geological and geochemical knowledge of epithermal precious-metal systems. I t follows on, and should b e used in conjunction with, t h e first volume in this series: Mineral-Fluid Equilibria in Hydrothermal Systems by Henley e t al. (1984). In t h e present volume we have a t t e m p t e d t o provide a framework for understanding t h e systematics of controls on fluid compositions and of m e t a l and gangue transport and deposition. The structure, dynamics, and transport properties of a c t i v e geothermal systems a r e used a s a starting point. With active systems a s a reference, t h e evolution of fluid compositions and t h e constraints on m e t a l and gangue transport and deposition in t h e epitherrnal environment a r e explored. The systematics of fluid inclusion and light stable-isotope applications is developed because these two approaches have been so useful in t h e development of our understanding of epithermal processes. The importance of boiling, cooling, and oxidation in transport properties of epithermal systems is evaluated through a numerical modelling approach. With t h e foregoing a s background, t h e observational base and its interpretation for epithermal ore deposits in continental volcanic and sedimentary terranes is explored through summaries of the geologic, mineralogical, and geochemical characteristis of, and traceelement distributions in, some well-studied epithermal o r e deposits. The final chapter is devoted t o t h e use of our understanding of epithermal systems in t h e development of exploration strategies. This volume does not a t t e m p t t o be t h e final word on epithermal o r e deposits, nor does i t claim comprehensive t r e a t m e n t . The absence of a chapter on t h e hydrology of epithermal systems documents t h e f a c t t h a t our c u r r e n t understanding of this aspect is woefully inadequate. I t does not r e f l e c t a lack of recognition of t h e importance of hydrologic controls. Similarly, this volume focuses on volcanic- and sediment-hosted epithermal deposits in t h e cordillera of western North America, particularly t h e United States. It does not t r e a t aspects of alkaline- or basaltic-rock related deposits such a s Cripple Creek,

Colorado, and Vatacoula, Fiji, nor does i t t r e a t t h e relationship of epithermal systems t o deeper hydrothermal systems responsible for t h e formation of porphyry-type deposits. Again our reason is t h e lack of an a d e q u a t e observational base. Our primary purpose in organizing this volume and t h e r e l a t e d Short Course has been t o stimulate critical studies t o improve our c u r r e n t understanding of epithermal deposits and processes rather than t o document it. Perhaps our omissions will serve this purpose equally a s well a s o u r inclusions.

REFERENCES Barton, P. B., Jr., Bethke, P. M., Roedder, E., 1977, Environment of ore deposition in t h e C r e e d e mining district, San Juan Mountains, Colorado: 111. Progress toward interpretation of t h e chemistry of t h e ore-forming fluid f o r t h e OH vein: Economic Geology, v. 72, p. 1-25. Becker, G. F., 1882, Geology of t h e Comstock lode and U.S. Geological Survey t h e Washoe district: Monograph 3, 442 p. Bell, J. M., and Fraser, C., 1912, The g r e a t Waihi gold Survey, Bulletin mine: New Zealand Geological 15. Burbank, W. S., 1933, Epithermal base-metal deposits in O r e deposits of t h e Western States: American Institute of Mining Metallurgical Engineers, New York, P a r t VI, p. 641-652. Ferguson, H. G., 1927, Geology and o r e deposits of t h e Mogollon mining district, New Mexico: U.S. Geological Survey Bulletin 787, 100 p. Ferguson, H. G., 1929, The mining districts of Nevada: Economic Geology, v. 24, p. 131-141. Hemley, J. J., and Jones, W. R., 1964, Chemical aspects of hydrothermal alteration with emphasis on hydrogen metasomatism: Economic Geology, v. 59, p. 538-569. Henley, R. W., Truesdell, A. H., and Barton, P. B., Jr., 1984, Fluid-mineral equilibria in hydrothermal systems: Society of Economic Geologists, Review in Economic Geology, v. 1, p. 267. Lindgren, W., 1928, Mineral Deposits: Third Edition, McGraw Hill, New York, 1049 p. Lindgren, W., and Ransome, F. L., 1906, Geology and gold deposits of t h e Cripple Creek district, Colorado: U.S. Geological Survey, Professional Paper 54, 516 p. Nolan, T. B., 1933, Epithermal precious-metal deposits in O r e deposits of t h e Western States: American Institute of Mining Metallurgical Engineers, New York, P a r t VI, p. 623-640. Ransome, F. L., 1909, The geology and o r e deposits of Goldfield, Nevada: U.S. Geological Survey Professional Paper 66, 258 p Taylor, H. P., Jr., 1973, 18/616/0 evidence for meteoric-hydrothermal alteration and ore deposition in t h e Tonopah, Comstock Lode, and Goldfield mining districts, Nevada: Economic Geology, v. 68, p. 747-764.

White, D. E., 1955, Thermal springs and epithermal o r e deposits: Economic Geology, F i f t i e t h Anniversary Volume, p. 99-154. White, D. E., 1981, A c t i v e geothermal systems and hydrothermal o r e deposits: Economic Geology, Seventy-fifth Anniversary Volume, p. 392-423.

Wisser, E., 1960, Relation of o r e deposition t o doming in t h e North American Cordillera:. Geological Society of America, Memoir 77.

ACKNOWLEDGMENTS

As is t r u e for any e f f o r t of t h e scope of this volume, many people in addition t o t h e editors played key roles along t h e road t o final publication. The t i m e and e f f o r t expended by each author is greatly appreciated a s a r e t h e contributions of t h e large cadre of individual reviewers who have offered insights and alternative perspectives t o t h e authors. Technical support t o t h e editors including manuscript preparation and revision, final formatting for publication, and badgering of both editors and authors was provided by t h e Editorial Support Group, Branch of Exploration Geochemistry, U.S. Geological Survey. Within this group we would especially like t o thank P a m e l a Detra,

Dorothy Wesson, Marilyn Billone, and Candy Vassalluzzo. An earlier version of this t e x t was assembled for use a t t h e Society of Economic Geologists Short Course by t h e Branch of Exploration Geochemistry Clerical Support Group. Finally, we would like t o express appreciation f o r t h e patience of J a m i e Robertson, Series Editor, Reviews in Economic Geology, and t h e support of t h e Society of Economic Geologists. Byron R. Berger Philip M. Bethke

BIOGRAPHIES

BYRON R. BERGER received a B.A. degree in ~ c o n o m i c ~ e o from l o ~ ~Occidental College in 1966 and a M.S. in Geology from t h e University of California, Los Angeles in 1975. He worked a s a petroleum exploration geologist for Standard Oil Company of California from 1968-1970 and a minerals exploration geologist and research scientist for Continental Oil Company from 1971-1977. He joined t h e U.S. Geological Survey in 1977, and has been involved in research o n epithermal precious-metal deposits and t h e relationship of magma genesis t o o r e genesis. He is currently t h e Chief of t h e Branch of Exploration Geochemistry. He is an adjunct assistant professor of geology in t h e Department of Geological Sciences a t t h e University of Colorado, Boulder, where he has taught courses on t h e geology and geochemistry of epithermal ore deposits and exploration geochemistry. H e is a member of several professional societies including t h e Geological Society of America and t h e American Geophysical Union. PHILIP --

M. BETHKE received a B.A. degree in Geology from Amherst College in 1952 and a Ph.D. in Geology (specialization in Mineralogy and Ore Deposits) from Columbia University in 1957. He was Assistant Professor of Geology a t t h e Missouri School of Mines and Metallurgy (now t h e University of Missouri-Rolla) from 1955 t o 1959. He joined t h e U.S. Geological Survey a s a WAE research geologist in 1957 and transferred t o full t i m e in 1959. His research has combined field and laboratory approaches t o t h e study of hydrothermal ore deposits, particularly t o epithermal vein systems. He has held several administrative positions with t h e U.S.G.S., most recently, Chief of t h e Branch of Experimental Geochemistry and Mineralogy. He is a member of several professional societies and is currently a Councillor of t h e Society of Economic Geologists. He has been active in t h e establishment of t h e SEG Short Course Series, and is currently Chairman of t h e Short Course Committee.

SAMUEL S. ADAMS received B.A. and M.A. degrees from Dartmouth College in 1959 and 1961, and a Ph.D. degree from Harvard University in 1967. From 1964 t o 1977 he served a s mine neolonist, - . exploration geologist, exploration manager, and exploration vice president, employed by International Minerals and ~ h e m i c a i ~ o ; ~ o i a t i o n -and then the Anaconda Company. During this period, his work emphasized sediment-hosted mineral deposits, particularly potash and uranium. Since 1977 he has served a s a lecturer and consultant t o industry, research organizations, and

-

government agencies in t h e a r e a s of mineral deposits, exploration, and resource assessment. His principal research i n t e r e s t is t h e representation of d a t a and concepts f o r all types of mineral deposits in coherent and predictive models for exploration and resource studies. He is currently a Councillor of t h e Society of Economic Geologists and t h e Geological Society of America. WILLIAM C. BABGY received a Ph.D. degree in Earth Science from t h e University of California, Santa Cruz, in 1979 based on petrogenetic research of Tertiary volcanic rocks in t h e Sierra Madre Occidental, Mexico. His industry experience includes geologic uranium in t h e evaluation of volcanic-hosted McDermitt caldera complex, Nevada, and t h e bulk mineability potential of t h e amythest silver vein system a t Creede, Colorado. Industry research included development of a n occurrence model f o r hot spring-related gold deposition based on t h e McLaughlin gold deposit in California. Present research interests a r e focused on t h e genetic aspects of sediment-hosted precious-metal deposits. ROBERT J, BODNAR received a n M.S. degree from t h e University of Arizona and a Ph.D. degree from The Pennsylvania S t a t e University and has been involved in various a s p e c t s of fluid-inclusion research for t h e past 10 years. He worked for 1 year a s a research geochemist in t h e Ore Deposits Group of Chevron Oil Field Research Company and is currently an assistant professor in t h e Department of Geological Sciences a t Virginia Polytechnic Institute and S t a t e University. KEVIN BROWN received a n -

M.S. degree in Chemistry in 1969 and a Ph.D. degree in Chemical Crystallography in 1972 from t h e University of Except for a two-year Auckland, New Zealand. sojourn a t t h e E.T.H. in Zurich, he has worked a t t h e Department of Scientific and Industrial Research, New Zealand. Initially in Wellington, his research i n t e r e s t c e n t e r e d around t h e crystal structures of organic reaction intermediates, but h e gradually c a m e down t o e a r t h with t h e crystal structures of some new epithermal minerals. In 1981, he shifted t o t h e Geothermal Section a t Wairakei, where his present research is concerned with experimental studies of mineral deposition from geothermal fluids.

CYRUS W. ---

FIELD received a B.A. degree in Geology from Dartrnouth College in 1956 and M.S. and Ph.D. degrees in Economic Geology, Geochemistry, and Petrology from Yale University in 1957 and 1961,

respectively. H e worked a s a n exploration geologist during t h e summers of 1955, 1956, and 1957 for t h e Oliver Iron Mining Company and Quebec C a r t i e r Mining Company subsidiaries of the U.S. Steel Corporation, and served a s a research geologist from 1960 t o 1963 with t h e Bear Creek Mining Company division of Kennecott Copper Corporation. In 1963, he joined t h e faculty of Oregon S t a t e University where h e is currently Professor of Geology. His research interests a r e largely concerned with t h e geology and geochemistry of hydrothermal mineral deposits; particularly t h e application of stable isotope and major-minor-trace e l e m e n t investigations t o their genesis. He is a member of several professional societies and was Vice President of t h e Society of Economic Geologists in 1981. RICHARD H, FIFAREK received a B.S. degree in Geology from t h e University of Washington in 1974, and M.S. and Ph.D. degrees in Geology (specialization in Economic Geology) from Oregon S t a t e University in 1982 and 1985, respectively. From 1974 t o 1984, he worked periodically a s a n exploration geologist (4 yrs.) for several mining companies, a s a research assistant/ scientist (1 yr.) a t t h e facilities of t h e Branch of Isotope Geclogy (Denver), U.S. Geological Survey, and a s an instructor f o r Oregon S t a t e University. Presently, he is a n assistant professor in t h e Department of Geology a t Southern Illinois University where he teaches and conducts research in economic geology and isotope geochemistry. His research interests include integrated geologic (field) and geochemical investigation of massive sulfide and epithermal Au-Ag deposits, and modeling t h e isotopic evolution of fluids and rocks in hydrothermal systems. NORA K. ---

FOLEY received a B.S. degree in Geology and Mineralogy from t h e University of Michigan in 1978 and a n M.S. degree in Geological Sciences from Virginia Polytechnic Institute and S t a t e University in 1980. She is currently working towards a doctoral degree in Geology through Virginia Polytechnic Institute and S t a t e University. Since 1980, she has been a research geologist a t t h e U.S. Geological Survey in Reston, Virginia. Her research has included fluid-inclusion and isotopic studies of different types of ore deposits, including Ag- and base-metal-bearing, epithermal deposits, sediment-hosted, stratabound, Pb-Zn deposits, and Kuroko-type massive sulfides.

ROBERT O, FOURNIER received a n A.B. degree in Geology in 1954 from Harvard College and a Ph.D. in Geology (specializing in Economic Geology, in general, and t h e Ely porphyry copper deposit, in particular) from t h e University of California a t Berkeley in 1958. Since then, he has been a research geologist with t h e U.S. Geological Survey. His research interests have ranged from laboratory studies of mineral-water interactions a t hydrothermal conditions appropriate for shallow levels in t h e crust, t o field studies of presently active hydrothermal systems, including Yellowstone National Park, Coso and Long Valley, California, and Zunil, Guatemala. Experimental studies have emphasized solubilities of

silica species in w a t e r and saline solutions. H e has also been a leader in t h e development of several chemical geothermometers a n d mixing models t h a t a r e now widely used in t h e exploration for geothermal resources. His present research focuses mainly on internally consistent chemical, isotopic, and hydrologic models of presently a c t i v e hydrothermal systems. He has served on NATO c o m m i t t e e s t o review geothermal energy development programs in Iceland, France, Greece, Portugal, and Turkey, and other c o m m i t t e e s t o review geothermal exploration programs in Argentina and Thailand. He was Chairman of t h e Organizing C o m m i t t e e f o r t h e 1975 United Nations International Symposium on Geothermal Energy, and Chairman of t h e Technical Program C o m m i t t e e for t h e 1985 G R C International Symposium on Geothermal Energy. He now serves on panels t o oversee geothermal developments in Costa Rica and Panama, and several U.S. Continental Scientific Drilling Committees. He is a member of several societies and has served on t h e Board of Directors of t h e Geochemical Society and t h e Geothermal Resources Council. DANIEL -- 0. HAYBA received a B.A. degree in Geology from t h e College of Wooster in 1976 and a n M.S. degree in Geochemistry and Mineralogy from the Pennsylvania S t a t e University in 1979 following a study of t h e Salton Sea geothermal systern. From 1978 t o 1980, he worked for Exxon Production Research Company on computer modeling of ore deposits. Since t h a t time, h e has been a research geologist with t h e U.S. Geological Survey where his research has been directed towards understanding t h e ore-forming processes in epithermal systems. PAMELA HEALD received a B.A. degree in Geology in 1971 from Vassar College and a n M.S. degree in Geology from George Washington University in 1977. She has been a research geologist a t t h e U.S. Geological Survey since 1972. Her research has included spectral reflectance and structural studies in Nevada, with a focus on ore deposits, and mineralogical and geochemical studies t o evaluate oreforming processes in epithermal precious- and basem e t a l deposits. RICHARD W. HENLEY received a B.S. degree in Geology in 1968 from t h e University of London and a Ph.D. degree in Geochemistry from The University of Manchester in 1971 following experimental studies of gold transport in hydrothermal solutions and t h e genesis of some Precambrian gold deposits. He was Lecturer in Economic Geology Memorial University of Otago, New Zealand, from 1971 t o 1975, and a t Memorial University, Newfoundland, until 1977. Research interests have focused on t h e mode of origin of a number of different types of ore deposits including post-metamorphic gold-tungsten veins, porphyry copper, massive sulfide, and placer gold deposits. He is currently with t h e Geothermal Chemistry Section of t h e Department of Scientific and Industrial Research a t Wairakei, New Zealand, and a visiting lecturer a t t h e Auckland Geothermal

Institute. Through 1983-84, he was a Fulbright Fellow and G u e s t Investigator a t t h e U.S. Geological Survey a n d during t h a t t i m e produced Volume I of this Review series. His present research includes a number of isotope and c h e m i c a l studies relating t o t h e exploration and development of geothermal s y s t e m s a n d g e o t h e r m a l implications f o r t h e origin of o r e C. A. KUEHN received a n M.S. degree Pennsylvania S t a t e University and has 7

from t h e years of experience in exploration f o r sediment-hosted gold deposits. H e i s c u r r e n t l y a n NSF Reseaqch Assistant and Ph.D. c a n d i d a t e a t t h e Pennsylvania S t a t e University a n d part-time employee of t h e U.S. Geological Survey working on t h e Carlin gold deposit.

MARK H. REED received a B.A. d e g r e e in C h e m i s t r y --and in Geology from C a r l e t o n College in 1971 and M.A. a n d Ph.D. d e g r e e s in Geology a t t h e University of California, Berkeley, in 1977. His Ph.D. research was on t h e geology a n d geochemistry of t h e massive sulfide deposits of t h e West Shasta District, California. F r o m 1977 through 1979, he worked f o r t h e Anaconda Minerals Company a t Butte, Montana. Since t h a t time, h e has t a u g h t a n d conducted research a t t h e University of Oregon, where he is currently Associate Professor of Geology. His research has focused on a l t e r a t i o n and m e t a l zoning in t h e porphyry copper a n d l a r g e vein deposits a t B u t t e a n d t h e geochemistry of hydrothermal alteration, m e t a l transport, and o r e deposition in massive sulfide and epithermal systems.

J. REYNOLDS T. -

received a n M.S. degree from t h e University of Arizona a n d has been a n exploration

xviii

geologist specializing in t h e application of fluid inclusions t o mineral exploration f o r t h e past 5 years. MILES --

L. SILBERMAN received a B.S. d e g r e e from t h e C i t v Universitv of New York and M.S. and Ph.D. d e k e e s from t i e University of Rochester, New York. H e is a member of t h e Branch of Exploration Geochemistry of t h e U.S. Geological Survey, with current assignments t o t h e Redding, California (CUSMAP) project, and t o t h e study of t h e geochemistry of volcanic and metamorphic-hosted gold deposits in t h e western U.S. and northern Mexico. Previous work f o r t h e U.S;G.S. included geochronological, geochemical, and regional geological studies of precious- and basem e t a l deposits in t h e G r e a t Basin a n d Alaska, and t e c t o n i c syntheses with particular focus on t h e relationships of hydrothermal precious-metal deposits t o m a g m a t i c and m e t a m o r p h i c evolution. Between tours a t t h e U.S.G.S., he designed a n d supervised exploration programs f o r precious-metal deposits in t h e G r e a t Basin f o r t h e Anaconda Minerals Company.

NICOLAS G S P Y C H E R received a B.S. d e g r e e in E a r t h Sciences in 1979 a n d a Dipl. e s Sc. in Exploration Geophysics in 1980 from t h e University of Geneva, Switzerland. He is now a Ph.D. candidate a n d research assistant a t t h e University of Oregon. His present research includes studies of t h e t r a n s p o r t of arsenic and antimony in hydrothermal solutions, t h e mixing properties of geothermal gases, and t h e geochemical modeling of hot spring systems.

Chapter 1 THE GEOTHERMAL FRAMEWORK OF EPITHERMAL DEPOSITS R. W. Henley

INTRODUCTION In the c o n t e x t of exploration for epithermal deposits, why study geothermal systems a t all? A f t e r all, not one exploited system t o d a t e has been shown by drilling t o harbor any economically significant m e t a l resource--but then until recently not o n e had b e e n drilled for other than geothermal energy exploration.* The l a t t e r involves drilling t o depths of 500-3000 m e t e r s in search of high t e m p e r a t u r e s and zones of high permeability which may sustain fluid flow t o production wells for steam separation and electricity generation. In many cases such exploration wells have discovered disseminated base-metal sulfides with some silver and argillic-propylitic alteration equivalent t o t h a t commonly associated with orebearing epithermal systems (Browne, 1978; Henley and Ellis, 1983; Hayba e t al., 1985, this volume). In general, however, geothermal drilling ignores t h e upper f e w hundred m e t e r s of t h e a c t i v e systems and drill sites a r e situated well away from natural f e a t u r e s such a s hot springs o r geysers, t h e very f e a t u r e s whose characteristics (silica sinter, hydrothermal breccias) a r e recognizable in a number of epithermal preciousm e t a l deposits (see, for example, White, 1955; Henley and Ellis, 1983; White, 1981; Berger and Eimon, 1983; Hedenquist and Henley, 1985a; and earlier workers such a s Lindgren, 1933). Knowledge of t h e upper f e w hundred m e t e r s of active geothermal systems is s c a n t and largely based on interpretation of hot-spring chemistry. Tantalizingly, in a number of hot springs, transitory red-orange precipitates occur which a r e found t o be o r e grade in gold and silver and which c a r r y a suite of elements (As, Sb, Hg, TI) now recogpized a s characteristic of epithermal gold deposits (Weissberg, 1969).

*Kennecott has recently announced significant gold discoveries in still active geothermal fields on Lihir and Simberi Islands, Papua, New Guinea. Today's active geothermal systems occupy t h e s a m e tectono-volcanic niche a s those hydrothermal systems, preserved from the past, which hosted t h e near-surface (0-1000 m) formation of epithermal o r e deposits in t h e Tertiary volcanic t e r r a n e s of t h e Circum-Pacific region and elsewhere--the relatively shallow origin of these deposits resulting in t h e i r loss by erosion from erstwhile similar, but older, terranes. Formed a t deeper levels (2-5 km or so) beneath calcalkaline volcanoes in these s a m e volcanic t e r r a n e s (Sillitoe, 1973; Henley and McNabb, 1978), porphyry-

t y p e copper and molybdenum deposits a r e preserved in both Tertiary and much older hydrothermal systems. The purpose of this chapter is t o review s o m e of the principal chemical and physical characteristics of t h e active geothermal systems which a r e essential t o t h e understanding of t h e origin of epithermal o r e deposits and therefore t o their successful exploration. For more detailed information, t h e reader is referred t o t h e publications cited in t h e text. HYDROTHERMAL SYSTEMS IN GENERAL The t e r m "hydrothermal" encompasses all types of hot-water phenomena in t h e earth's c r u s t although most commonly the t e r m is used in r e f e r e n c e t o those associated with impressive geyser activity, aesthetically a t t r a c t i v e hot pools, etc. These features a r e most common in volcanic a r e a s such a s Yellowstone National Park, U.S.A., Iceland, or in t h e Taupo Volcanic Zone of New Zealand, but other t e r ranes also host hydrothermal activity even though subsurf a c e temperatures m a y b e relatively low and surface f e a t u r e s less impressive. Warm springs in t h e Rocky Mountains, t h e European or New Zealand Alps, or in t h e sedimentary massifs of c e n t r a l Europe a r e examples, and i t is clearly important for mineral exploration t o discriminate these types of systems from those in more favorable geological environments. Geothermal systems a r e extraordinarily abundant in t h e tectonically a c t i v e zones of t h e earth's crust and may be broadly classified according t o their plate t e c t o n i c setting and principal source of heat (Table 1.1). Chemical differences a r i s e from the sources of recharge water and contribution of gases from magmatic or metamorphic sources. Warm springs also occur in t h e tectonically stable crust where t h e deep crustal penetration of groundwater occurs in favorable sedimentary formations such a s limestones and the h e a t supply is t h e ambient continental h e a t flow. Each of these classes of geothermal systems appears t o have some correlative preserved in t h e geologic past and most commonly recognized a s one or another of t h e various families of hydrothermal ore deposits. For magma-related hydrothermal systems, these range from ophiolite-hosted massive sulfides through t h e polymetallic massive sulfides of island a r c s t o t h e porphyry copper and epithermal preciousm e t a l deposits of terrestrial continental terranes, while for a m a g m a t i c systems t h e s e range from t h e Mississippi Valley and r e l a t e d base-metal deposits in sedimentary basins t o t h e post-metamorphic vein deposits associated with orogeny.

T a b l e 1.1--Crustal s e t t i n g of hydrothermal s y s t e m s c l a s s i f i e d a c c o r d i n g t o p r i n c i p a l heat-source and c r u s t a l h o s t . CRUSTAL HOST/ HEAT SOURCE

Oceanic

AMAGMATIC

MAGMATIC

Ridge, h o t s p o t , back-arc b a s i n

-----

Magmatic a r c Continental

Crustal extension (Hot s p o t , r i f t )

Plate collision Plate-interior basins

Collision-Related Amagmatic Hydrothermal Systems

volcanic terranes, i 1 0 0 0 m in andesitic volcanic t e r r a n e s and for t h e mountain belt systems, 2000 m.

Only recently have d a t a become available from geothermal investigations in mountain belts. In t h e Southern Alps of New Zealand, for example, hot Terrestrial Magma-Related Hydrothermal Systems springs occur in t h e central, relatively aseismic region with t h e highest uplift r a t e (10-20 mm/year) where t h e By contrast, systems in volca ' c terranes have combination of uplift and erosion "exposes" a thermal high 3 ~ teo 4 ~ ratios e and t h e 6 0 of alteration anticline with near-surface gradients up t o 150°C/km minerals are, with few exceptions, depleted relative t o (Allis e t al., 1979). A similar environment is proposed primary minerals. Tern e r a t u r e s encountered during for hot springs in other collision-related mountain drilling range up t o 400 gC (Batini e t al., 1983a) and belts. Recent drilling a t Yangbajing (Tibet) and in t h e waters a r e predominantly m e t e o r i c in origin, and P a r b a t i Valley (N. India), for example, has located hot typified by t h e presence of chloride ion with mC1- > waters up t o 1 7 0 ' ~ (Giggenbach e t al., 1983) which a r e predominantly m e t e o r i c in origin, but contain low 3 ~ e >mSy= --they a r e here, for convenience, designated chlorl e waters. Although some highly saline fluids t o 4 ~ ratios e typical of helium of deep crustal origin. a r e evolved in rift zones such a s t h e Imperial Valley The uplift s e t t i n g of t h e s e hydrothermal systems (California), salinities a r e typically low, clustering is perhaps analogous t o t h a t of L a t e Mesozoic postaround 10,000 rng/kg C1 (1.6 wt.-% NaCl equivalent) in metamorphic gold and scheelite veins on t h e South andesitic volcanic terranes, 1000 mg/kg in rhyolitic Island (New Zealand) and, by inference, similar volcanic t e r r a n e s and much lower in basaltic volcanic deposits in much older terranes. Examples a r e t h e terranes. Dissolved gas, always preponderantly C 0 2 , gold veins of the Valdez Group (S. Alaska), Mother a f f e c t s a major contrast between systems and ranges Lode (California), Yellowknife (Northwest Territories) from very low (0.01 wt.-% C 0 2 ) a t Wairakei (New and Kalgoorlie (W. Australia). In e a c h of these, in Zealand) and Ahuachapan (El Salvador) t o several contrast t o t h e epithermal precious-metal depos'ts wt.-% a t Broadlands and Ngawha (New Zealand) (see discussed below, vein quartz is enriched in "0 Table 1.3). Other dissolved components a r e controlled relative t o host rocks. This f e a t u r e has led many by mineral-fluid and gas-gas reactions. Alteration workers (e.g., Henley e t al., 1976; F y f e and Kerrich, assemblages in t h e s e types of geothermal systems 1984) t o suggest a metamorphic origin for t h e correspond closely t o those encountered in epithermal hydrothermal fluid; vein formation occurring from and porphyry-style mineral deposits. fluids of metamorphic dehydration origin in response The deep hydrologic s t r u c t u r e of t h e terrestrial t o post-metamorphic uplift and/or overthrusting. It geothermal systems is controlled by t h e convective may also be possible, however, t o generate these s a m e upflow of chloride w a t e r s (evolved by water-rock i isotope characteristics by interaction of m e t e o r i c magma interaction a t depths of 5 t o 8 km) but above w a t e r and rocks a t a low w a t e r t o rock ratio opening depths of around I km surface topography plays a t h e possibility t h a t such deposits may be much major role in t h e dispersion of the chloride water by shallower* in origin than t h e 10 t o 20 km generally introducing a l a t e r a l flow component toward considered. topographic lows. Boiling occurs a s chloride water rises through t h e system, t h e resultant s t e a m migrating t o t h e surface independently where nearsurface condensation and oxidation of co-transported *In this paper t h e t e r m "shallow" is used rather H2S produces sulfate-dominated steam-heated irreverantly t o r e f e r t o depths less than about 500 waters. These f e a t u r e s a r e incorporated in t h e general meters. In o r e deposit research, depths (estimated perhaps from fluid-inclusion data) a r e generally also model of t h e s t r u c t u r e of a geothermal system reproduced in Figure l . l a (from Henley and Ellis, used irreverantly, taking no account of t h e importance 1983). of topographic relief; of t h e order of *I00 m in silicic

f'i

adv argillic alteration Near Acid neutrat pH Sulphate & Chlor~de Dilute with

Meteoric water

a

alteration Groundwater

-

0

. . ., .

1 km HEAT AND MASS (NaCI, C02, 502, Hz0 . ) TRANSFER FROM MAGMA SYSTEMS. KEY Pre-Volcanic Basement

Steam-heated Acid ~ 0 ~ / ~ 0 ~ - r i c h waters

Intrusive Volcanics

SO;-

C l - waters ( f i g I b )

0 Near neutral Chloride waters (kithin 200" Isotherm approx.)

Low Permeability Stratum e.g. Mudstones

Two Phase Region Water Liquid + Steam ( +Gas) - ~-

Figure l.la. Generalized structure of a typical geothermal system in silicic-volcanic terrane. Notice the overall size of the system relative to the size of the discharge features ( i . e . , hot springs, etc.). The temperature distribution shown is based on the Wairakei system where a west-to-east flow occurs in the upper portion of the system and boiling occurs above about 500 meters. In other systems such as those in Figure 1.2, more or less lateral flow may occur. Boiling may extend to much greater depths if C02 contents are high (see text), and higher temperatures may occur at shallower depths than shown in this figure, as at Mokai (Fig. 1.2d).

The relatively high relief of andesite volcanic t e r r a n e s results in l a t e r a l flows of hot chloride water for up t o 20 km while t h e occurrence of near-surface magmas exsolving gases (HCI, SO2, etc.) often produces high t e m p e r a t u r e fumaroles and/or acid sulfate-chloride c r a t e r lakes, such a s those on Mount Ruapehu, New Zealand ( P l a t e 1.1) and El Chichon, Mexico (Giggenbach, 1974; Kyosu and Kurahashi, 1984; Casadevall e t al., 1984). These l a t t e r features, with their associated intense advanced argillic alteration, a r e possible correlatives of t h e upper portions of t h e t y p e of hydrothermal systems responsible for t h e formation of gold--(enargite) sulfide deposits of t h e "Goldfield type" (Ransome, 1909) such a s Goldfield (Nevada), Summitville (Colorado), Bor (Yugoslavia), and elsewhere. Figure 1.lb provides a general s t r u c t u r a l model for this geothermal environment. They m a y also b e related in some cases t o t h e upper portions of developing porphyry copper deposits (Sillitoe, 1983). The geochemistry and structure of magmar e l a t e d hydrothermal systems have been reviewed in a number of r e c e n t t e x t s t o which t h e reader is r e f e r r e d for background reading and discussion of hydrothermal chemistry--see, for example, Ellis and Mahon, 1977;

Meteoric

Henley and Ellis, 1983; Henley e t al., 1984. A brief summary of hydrothermal chemistry is given in Henley and Brown (1985, this volume). In t h e remainder of this chapter attention is focused on those a s p e c t s of t h e chemistry and structure of geothermal systems relevant t o t h e understanding of t h e formation of epithermal ore deposits. TERRESTRIAL MAGMATIC-HYDROTHERMAL SYSTEMS Large-Scale Structure Early in t h e commercial development of the Wairakei geothermal field in New Zealand, t h e accumulating d a t a from exploration wells showed (a) t h a t t h e fluids present w e r e = directly exsolved from shallow bodies of crystallizing m a g m a and (b) t h a t t h e hydrothermal activity seen a t t h e s u r f a c e was a minor phenomenon associated with t h e discharge of a very large, deeply convecting body of heated groundwater (Elder, 1966). Using analog and numerical modelling, Elder and other research scientists showed t h a t convection, with a depth s c a l e of a t l e a s t 5 km, was

water

Neutral chloride water

F i g u r e l.lb. S t r u c t u r e of a t y p i c a l geothermal system i n andesitic-volcanic t e r r a n e s emphasizing ( 1 ) e x t e n s i v e lateral f l o w and (2) g e n e r a t i o n o f a large a d v a n c e d - a r g i l l i c a l t e r a t i o n zone i n r e s p o n s e t o h i g h - l e v e l volcanism. (Modified and reproduced w i t h p e r m i s s i o n f r o m Henley and E l l i s , 1983.)

Plate 1.1. Oblique, aerial view of the Waiotapu system, New Zealand from the southeast. Topographic features may be related to the system map (Fig. L.2b). Mount Tarawera (on the horizon) is a composite rhyolite dome which, in 1886, violently erupted &salt through an axial rift. Associated phenomena were 6?e destruction of the Pi& and WLite Silica Terraces (Henley et al., 1984, Plate 1.21, and a n~mberof hydrothermal eruptiors in the btomahana-Waimanq geothermal system. The natural discharge dominating the surface expression of the system is *Ae Pool (middle right) which occupies a hydrothermal eruption vent formed 900 years ago and which may overlie some 0.1 million ounces of gold formed 'by boiling in the conduit of the ,ml. Surface antimony-arsenic precipitates xcur which are ore-grade in silver and gold (photo D. L. Homer, N. Z. Geological Survey).

responsible for t h e e x t r e m e thermal gradients and t e m p e r a t u r e p a t t e r n s observed in t h e exploration drilling. A t t h e s a m e tinie, at Wairakei and in other fields, t h e e f f e c t s of near-surface (depths less than 1000 m) stratigraphy a n d s t r u c t u r e and of reliefcontrolled groundwater fiow became evident Iargely through geophysical techniques, especially resistivity surveying (Healy and Hochstein, 1973). Hanaoka (1983) h a s numerically modelled t h e e f f e c t s of topographic r e i i e l on near-surface hot-water flow and i t s dispersion by cold groundwater. This e f f e c t is partly responsible for t h e mushrooming of isotherms shown in many convective models a n d field cross sections. Figure 1.2 provides s o m e examples of t h e iaterai-flow characteristics and distribution of natural discharges in a number of geothermal fields explored by drilling, with perhaps t h e Niokai field in New Zealand (Fig. 1 . 2 ~ ) being a particuiarly good illustration of l a t e r a l flow as shown in cross-section in Figure 1.2d. In geothermal systems hosted by silicic volcanic rocks, s u r f a c e topography is primarily controlled by block-faulting o r caldera collapse providing relief of a f e w hundred m e t e r s and consequent l a t e r a l flow over distances of up t o about 5 km. In t h e higher relief t e r r a n e typical of andesitic volcanism, more e x t r e m e i a t e r a i flow occurs up t o about 20 km, An additional f e a t u r e of a c t i v e andesitic volcanic t e r r a n e s is t h e occurrence of high-level volcanism which allows volcanic gas t o vent t o summit fumaroles or t o summit c r a t e r lakes (Giggenbach, 1974) and t o maintain high-ievei "perched" aquifers containing very acid sulf ate-dominated waters. Exploration wells a t high eievation in such terranes often encounter vapor-dominated geotherma! environments. In t h e majority of systems, liquid water provides t h e continuum for fluid fiow but in other, f a r less common systems, water vapor dominates t h e discharges of d e e p exploration wells. The preexploitation s t a t e s of these "vapor-dominated" systems a r e poorly known and various models have been produced based on production d a t a from exploited fields. For example, f o r t h e Geysers (California) and f o r Eardereilo (Italy), White e t ai. (1971) suggest th.e presence of a very deep convecting brine overlain by a n "alteration-sealed" c a p of vapor. Of particular i n t e r e s t is t h e association of these systems with epithermai mercury and gold mineralization (e.g., McLaughlin, California), but both t h e Geysers and larder el!^ also contain base-metal sulfides and other "ore-related1' minerai phases in drill c o r e (Beikin et al., 1983; Sternfe!d, 1981) which suggest t h a t t h e present system has evo!ved from s o m e previous liquikdominated s t a t e . Others have suggested t h a t eievated gas-content (dominantly C 0 2 ) perhaps coupled with relatively low host-rock porosity, may account f o r t h e vapor-dominated c h a r a c t e r o i well-discharges and post-exploitation pressure data. It is interesting t o n o t e t h a t most of t h e explored "liquid-dominant" geothermal systems, in silicic volcanic t e r r a n e s especially, a r e associated with t e c t o n i c subsidence (about -5 mm per year in t h e Taupo Volcanic Zone in New Zealand), but both t h e Geysers and Larderello occur in regions of high t e c t o n i c uplift associated with volcanism. Quantitative d a t a from t h e Geysers region

a r e not available, although regional topography and erosion a r e suggestive of high uplift rates. At Larderello uplift r a t e s a r e of the order +5 t o 1 3 mm per year (M. Puxeddu, persona: communication^ and a r e evidenced by t h e coastiine migration of t h e Pisa area. The high h e a t ilow and geothermal activity of t h e Larderello region appears t o be r e l a t e d t o t h e emplacement of a post-orogenic batholith into continental c r u s t (Batini e t al., 198%; Puxeddu, 1989). Natural Discharges Hot water convecting into t h e near-surface part of a large hydrothermai system may be dispersed by mixing with iateraliy flowing cold groundwater or discharged directly t o t h e surface. Only a minor amount of h e a t energy is lost by conduction, but most is dispersed a s hot w a t e r and vapor flows a t t h e surface. The processes affecting a deep f!uid penetrating t o t h e surface depend on a variety of factors. Direct discharge depends on t h e availability of a suitable f r a c t u r e system (or hydrothermal eruption vent, s e e below) and gives rise t o a boiling spring, high in chloride and .mantied by silica sinter. Examples a r e t h e Champagne Pool, Waiotapu and the Pink and White T e r r a c e s of Rotomahana, New Zealand (Plate 1.1; and s e e Henley et ai., 19841, Geysers a r e a special class of boiling discharge which have a periodic discharge due t o t h e geometry of t h e conduit (Kieffer, 1984:. Often dilution precedes boiling of t h e mixed fluid a s i t finally moves t o t h e surface a s in t h e Ohaaki Pool a t Broadlands (Ohaaki) or t h e boiling springs of t h e Wairakei and Tauhara systems (Fig. I.4a). Fluids which a r e diluted with respect t o t h e deep chloride w a t e r form where interaction with nearsurface aquifers occurs either due t o high surface relief and groundwater flow or t o t h e proximity of t h e system margin. The natural discharges of the Wairakei-Tauhara and Mokai systems a r e examples (Fig. 1.2). Figure 1.3a shows schematicaliy the pressure distribution associated with various discharge phenomena. Drill-hoie d a t a suggest t h a t pressure gradients in t h e deeper system a r e generally about 10% above hydrostatic pressure with t h e excess pressure due t o t h e buoyancy of hot water relative t o surrounding cold groundwater (Eider, 1966; Cathles, 1977; Grant et al., 19821, and in some cases (e.g,, Mokai) a demonstrable component of hydrostatic head due t o recharge from a r e a s of relatively high relief. An excess pressure gradient is a requirement f o r ilow through permeable media. Below a hot-spring vent, fluid expansion leads t o two-phase flow in t h e highpermeability conduit. Phase separation may occur with t h e vapor discharging independently a t t h e surface a s a fumarole o r interacting with groundwater t o produce a steam-heated water. As suggested i n Figure 1.3a, minor throttling may occur aiong t h e flow path, but pressure drops a r e unlikely t o be greater than I bar. Where silicification isolates t h e conduit from t h e surrounding groundwater system, boiling, deepsystem iluid e x i t s t h e surface; but, where only partial isolation occurs through mineral deposition, the iiquid may itself i n t e r a c t with surficial groundwater before reaching t h e surface a s a hot o r warm spring, In t h e

NNE zs Si s e L =

SSW

3

-

n20i-

ilhlYI,*. E.0.r."

8

,,

..,,,",,.. Depaufi

d

Figure 1.2. DLstrikution of natural discharges in some active geothermal systems. T:?e field boundaries shown are based on the maximum resistivity gradient located by field silrveys reflecting the contrast between unmineralized groundwater and the ehlcride water present in the upper 500 meters of the geothermal systems. Fumaroles, steaming ground, and outflows of steam-heated waters are indicated by the 0 syxilwl an6 hot-water discharges by the g symbol. The Location of the principal convective upflow for each field is indiNotice that geothermal cated by the v exyloration and production wells are situated well away from natural features. Numerals designate features shown in the mixing diagrams of Figure 1.4.

.

Figure 1.2 (cont'd) a). Wairakei-Tauhara, New Zealand. These two fields are interconnected as shown by the resist-ivity and both show the occurrence of vapor discharge in the central region and hot..-wate qe on the margins followincj dilution. There is no evidence that water from take Taup penetrates either field, reaarge being derived from gr to the east and west. b). Waiotapu, New Zealand* This field has an extensive north-tosouth lateral flow originating in the vicinity of the 160,005 years hp. &cite domes to the north. Thermal features are related to major faults and a number of hydrothermal eruption craters have been recognized (major centers shown by the circles)--for full discussion see Fledempist and Wenley, 1985a. c). Mokai, New Zealand. Extensive lateral flow occurs from the vicinity of the caldera wall in the south toward the Waikato River to the north. Dilute hot springs occur north of the "field boundary" in the gorge of a stream following a major fault. d). Cross-section of the Mokai geothermal field showing the effect of lateral flow and dispersior. on the thermal structure of the system and distribution of natural features. (The cross section runs from the tap right-hand corner of Figure 1.2~ to the caldera wall south of well MK6).

Plate 1.2. Crater Lake, Ruapehu, New Zealand. Condensation of volcanic gas into the Crater Lake waters produces a fluid of pW 1.5 at about 55OC. The lake seldom overflows despite the presence of an incised channel (foreground)suggesting that much of the acid fluid drains through the core of the active andesite volcano producing an extensive high-level zone of advainced-argillic alteration. Interaction of this fluid with an underlying near-neutral pH hydrothermal system nay generate a gold d e p i t of the Goldfield type [photo by permission, R B. Glover, DSIR).

PRESSURE 10

20

30

bars

Pressure-depth relations in the Figure 1.3a. upper portion of a geothermal system. The diagram shows the transition between the deep system pressure and the pressure within the high permeability fracture network or conduit below a hot spring. Below the hot springs, the pressure at a specified depth is due to the weight of a standing column of hot water; the pressure-depth relation is here designated ''hot hydrostatic". Deeper in a system pressures exceed hydrostatic so that flow is maintained through the permeable aquifer--this is shown as the ''hot hydrodynamic" curve. Some minor pressure discontinuities are shown to indicate the possible occurrence of minor throttles which may occur due to fracture geometry or silicification, but these are probably rare. Phase separation may occur resulting in the presence of fumaroles or (acid) steam-heated waters in the vicinity of a boiling hot spring (e.g., Norris Geyser Basin, Yellowstone, Champagne Pool, Waiotapu). The effect of raising or lowering the ambient groundwater piezometric surface may be gauged by redrawing the curve for cold-water hydrostatic pressure. For example, if the cold-water piezometric surface is at +20 meters and the hot-spring conduit is not isolated by mineral d e p i -

tion, dilution may occur near surface. Dilution occurs on the margin of a hydrothermal system due to the relative pressure of cold groundwater over that of the hotwater system. example shown (Fig. 1.3a), deep mixing may occur where t h e pressure of cold water exceeds t h a t of t h e hot upflow. Exercise: The e f f e c t of relief, through a higher or lower piezornetric surface, may be gauged by adding cold water pressure curves t o Figure 1.3a corresponding t o higher and lower piezometric surfaces. Try i t by drawing curves parallel t o t h e r e f e r e n c e cold-water c u r v e in t h e figure. Haas (1971) has described t h e limiting hydrostatic conditions f o r temperature a s a function of depth in hydrothermal systems. The limiting condition (Fi.g. 1.3b) is t h e phase change t o vapor; liquid water rislng within a system boils a t t h e phase boundary with consequent formation of a low-density vapor fraction and a decrease in t e m p e r a t u r e (for a disctission of reversible and irreversible boiling in hydrothermal systems, s e e Barton and Toulmin, 1961). As discussed above, hydrodynamic pressures prevail a t d e t h in geothermal systems so t h a t a t , for example, 250 gC t h e boiling-point depth is a t about 400 r a t h e r than 462 meters. The e f f e c t of salinity on t h e boiling pointdepth relation is well known, but m o r e recently t h e e f f e c t of gas pressure has been recognized (Sutton and McNabb, 1977) a s shown in Figure 1.3b. The l a t t e r e f f e c t makes i t particularly difficult t o obtain reliable depth information from e s t i m a t e s of t e m p e r a t u r e (e.g., from fluid inclusions) in fossil hydrothermal systems (Hedenquist and Henley, 1985b; Bodnar e t al., 1985, this volume). The distribution of springs relative t o t h e geothermal system a s a whole is evident from t h e field maps shown in Figure 1.2. Areas occupied by hotwater discharge seldom represent more t h a n about 5% of t h e a r e a of t h e hydrothermal field itself. I t is also evident from these field examples t h a t t h e distribution of discharges is strongly controlled by topography, t h e presence of faults, etc. In general, f e a t u r e s associated with vapor-flow from t h e d e e p system occupy higher ground. They range from fumaroles t o hot springs f e d by steamh e a t e d surficial groundwaters t o steaming ground which results from t h e boiling of steam-heated waters. The l a t t e r originate above two-phase ('boiling') zones in t h e deep convective system from which C 0 2 and H2S-rich vapor escapes, but a r e adsorbed into surficial groundwater or condensate. Where H2S oxidation occurs due t o shallow interaction with t h e atmosphere, low pH steam-heated w a t e r s occur which a r e characterized by t h e presence of sulfate and absence of significant chloride in solution a s well a s t h e lack of significant silica sinter around C02-rich steam-heated waters, t h e hot spring. associated with illitic alteration, a r e also common marginal t o many fields (Mahon e t al., 1980; Hedenquist and Stewart, 1985) t o depths of several hundred meters; pH's a r e around 5 due t o dissolved C 0 2 a n d often result in e x t r e m e corrosion of

100

TEMPERATURE 200

300

OC

discontinuously despite continuous input of volcanic vapor, m e t e o r i c water, and glacial melt. The c r a t e r lake itself has been shown t o b e over 300 m e t e r s deep and is occupied by a n acid sulfate-chloride water with pH = 1.25 a t 55OC. Presumably, t h e bulk of this acid fluid drains downward through t h e flanks of t h e volcano causing advanced-argillic alteration e n route, and m a y encounter a normal hydrothermal system a t depth. Silica- and iron-enriched springs occur on the flank of t h e volcano a t lower elevations. These processes, a s noted, m a y b e responsible for t h e acids u l f a t e t y p e o r e environments (see Hayba et al., 1985, this volume) and raises a l l s o r t s of problems with respect t o terminology like hypogene and supergene! Hydrothermal Eruption Vents

Figure 1.3b. Hydrostatic boiling-pint versus depth relations of hydrothermal fluids, showing the contrasting effects of salinity and gas content. As discussed in the text, observations from active systems suggest that pressure gradients a t depth are about 10% greater than hydrostatic allowing higher temperatures at shallower depth than shown here.

geothermal well casings. C 0 2 exsolution following mixing of hydrothermal fluid w ~ t hcool groundwater, and dissolution of t h e C O into groundwater may be t h e dominant process in t i e i r formation r a t h e r than adiabatic boiling.

The hot springs described above a r e passive f e a t u r e s of t h e topography, but in some cases t h e system itself may g e n e r a t e high-permeability flow paths t o t h e surface. For example, a t Waiotapu (Fig. 1.2, P l a t e 1.1) t h e largest single discharge of liquid from t h e d e e p system--the Champagne Pool--is independent of t h e stratigraphy and original topography and occupies t h e v e n t of a hydrothermal Such eruption c r a t e r formed s o m e 900 years b.p. eruption vents a r e now known t o have formed in almost a l l of t h e New Zealand geothermal systems, but a r e less well known elsewhere due t o t h e frequent confusion of t h e eruptive products with volcaniclastic breccias which m a y also b e common in t h e vicinity. Hydrothermal eruption breccias a r e characterized by a n absence of primary volcanic material and a r e generally polylithic and m a t r i x supported. Clasts have a range of alteration styles and, together with stratigraphic data, indicate a n origin from depths up t o about 300 meters. An origin by gas exsolution has been proposed by Henley e t al. (1984) and by Hedenquist and Henley (1985a). Eruption breccias of shallower origin a r e also common in geothermal a r e a s and result from t h e interaction of vapor with surficial groundwater o r local removal of confining pressure, a s appears t o b e t h e c a s e for eruptions in Yellowstone (Muffler e t al., 1971). H e a t and Mass Flow in Geothermal Systems

Note: The t e r m "solfatara" discharges such a s fumaroles,

encompasses s t e a m but is now most commonly used t o refer t o volcanic g a d s t e a m discharges associated with sulfur deposition and advanced-argillic alteration. It is o f t e n used incorrectly in t h e discussion of m e t a l transport in subs e a floor systems!

As noted above, acid sulfate-chloride waters derived by condensation of volcanic gas occur a t high levels in andesitic t e r r a n e and sometimes t h e y m a y mix with meteoric water and a c c u m u l a t e t o form c r a t e r lakes. Downward movement of such high-level w a t e r s is of special interest and is now well known in explored geothermal fields in t h e Philippines and Taiwan. A t Ruapehu, New Zealand, for example, t h e summit crater lake (Plate 1.2) overflows

Table 1.2 shows h e a t and mass output d a t a from some geothermal systems. These a r e obtained by integration of ground t e m p e r a t u r e d a t a and physical measurements of t h e outflow r a t e s and temperatures of discharging hot springs and fumaroles. In s o m e cases a n independent e s t i m a t e of t h e upflow is obtainable using measurements of t h e chloride content of river water up and downstream of a geothermal field (Ellis and Wilson, 1955; Fournier e t al., 1975). These d a t a may b e r e l a t e d t o t h e convective upflow of high-temperature fluid in t h e system assuming some knowledge of t h e upflow temperature. For example, a t Waiotapu t h e measured surface h a t flow is 600MW(h) (600 MW(h) = 600 x 10% Joules/s). Exploration drilling and geochemical d a t a sug e s t t h a t t h e fluid feeding t h e field is a t about 300 8,C. The enthalpy of steam-saturated w a t e r a t 3 0 0 ' ~ is about 1350 Joules/gm (see Henley and Brown, 1985, this

T a b l e 1.2--Summary o f h e a t and mass f l o w s i n some New Z e a l a n d geothermal f i e l d s Field

T o t a l H e a t Flow

Equivalent upflow

Wairakei Tauhara Waiotapu Ohaaki roadla lands) Mokai

CHEMISTRY O F GEOTHERMAL DISCHARGES

volume, Fig. 2.3) so t h a t t h e mass flux of 300°C fluid is obtained by

Exercise: The Champagne Pool a t Waiotapu (Plate 1.1) has a discharge of about 10 kg/s of 70°C water and, we Calculate t h e estimate, about 7 kg/s of steam. proportion of t h e t o t a l convective upflow of h e a t and mass in t h e system which is discharged by this f e a t u r e alone (the enthalpy of 70°C w a t e r is about 300 Joules/gm and of s t e a m about 2600 Joules/gm). Within t h e heat-flow budget t h e most difficult f a c t o r t o assess is t h e proportion of t h e upflow which may b e dissipated by subsurface groundwater flow. On t h e basis of t h e size of t h e field and i t s deep temperatures, a t Broadlands (Ohaaki), t h e total heat flow is thought t o be greater than 100 MW(h) and equivalent t o >75 kg/s of chloride water, but t h e observed surface h e a t flow is less t h a n a third of this estimate. The principal outflow of hot w a t e r from this field is t h e Ohaaki Pool but, a t 10 kg/s, this accounts for only about 5 MW(h).

T a b l e 1.3--Summary

Table 1.3 compares t h e chemistry of waters from natural f e a t u r e s with t h e chemistry of deep waters encountered by d e e p drilling. The principal discriminating f e a t u r e s with respect t o origin have already been noted, i.e., d e e p fluid characterized by CI>>SO and surficial steam-heated waters by sO~>>C!. In a given field, comparison of chloride contents in hot springs provides information about mixing processes and flow directions. Careful application of chemical geothermometer techniques may also provide some unique insights into temperature patterns in t h e underlying system and processes occurring during outflow (Fournier, 19811, and therefore provide a n important guide for geothermal exploration. Chemical relations between natural discharges and t h e deeper chloride-water system a r e most commonly illustrated by means of "mixing-diagrams." Figure 1.4 provides examples where t w o conservative quantities a r e compared; in this case chloride concentration and h e a t content (enthalpy). The l a t t e r is frequently assumed t o be conservative during

of t h e c h e m i s t r y of h o t s p r i n g s and 5 e o t h e r m a l f l u i d s C o n c e n t r a t i o n s i n mg/kg

Field

Feature

Waio~apu

Champagne P o o l Well 7

Waiorapu

Champagne P o o l Well 80

Tauhara

Crow's Nest Kathleen Spring Well 1

Mokai

Northern Springs Well 3

Tongonan

Banati Spring Well 405

t°C

pH(t)

C1

SO4

H2S

c02

a Tauhara

1000 CHLORl DE

.,

,,

,,

2000 mg/ kg

hydrothermal processes since conductive h e a t transfer is a minor component of t h e overall h e a t budget and s y s t e m s a r e assumed t o b e in a steady-state with respect t o heat and mass transfer (heat released by a l t e r a t i o n reactions is also a minor component). The principal processes occurring a r e dilution (mixing) and a d i a b a t i c boiling during t h e irreversible expansion of t h e deeper system fluid a s i t rises and is subject t o less confining pressure. As discussed above, chemical t r e n d s due t o t h e s e processes a r e clearly shown in t h e diagrams which then allow t h e interpretation of t h e origin of individual t h e r m a l features. (In many cases, local mixing involves a steam- or conductively heated water with a t e m p e r a t u r e in t h e vicinity of 150°c.) Some of t h e Wairakei-Tauhara springs, for example, show evidence of dilution prior t o boiling below a hot spring, while others, as at Mokai, a r e simply derived by dilution (Figs. 1.4a and c). The Champagne Pool a t Waiotapu is a n example of t h e direct discharge t o t h e surface of deep fluid (Fig. 1.4b).

EPITHERMAL ORE-FORMING SYSTEMS Requirements f o r O r e Deposition

I

1000 CHLORl DE

2000 mg/kg

\ 2000. ENTHALPY kJ/kg

1000.

T h e transport chemistry of t h e epithermal group of m e t a l s has been reviewed by Barnes (19791, Weissberg e t al. (19791, and Henley and Brown (1985, this volume). In this chapter, these d a t a a r e built into t h e geothermal system framework t o provide a n understanding of t h e origin of epithermal mineral deposits in general based o n t h e solution chemistry of t h e m e t a l s and their response t o t h e t w o principal processes operating in t h e upper levels of these systems--dilution and adiabatic boiling. It is also i m p o r t a n t t o discuss t w o important interrelated c r i t e r i a for o r e deposition. These a r e (a) t h e availability of m e t a l s in solution and (b) t h e t i m e required f o r t h e o r e depositing system t o operate. Figures 1.5a and b show t h e solubilities of gold, silver, and lead (representing t h e base metals) a s functions of t e m p e r a t u r e and ligand concentration. As shown elsewhere (Ellis, 1970; Giggenbach, 1981; Henley e t al., 1984; Henley and Brown, 1985, this volume), t h e pH of hydrothermal fluids i n active and fossil systems is buffered by fluid + alumino-silicate reactions such a s t h e conversion of plagioclase t o mica and/or clay minerals. For low salinity fluids (CI = 1000 mg/kg), pH's a r e around 6.1 a t 2 5 0 ' ~ (i.e., on t h e alkaline side of neutral pH) but about 1 pH unit more acid (5.1) f o r fluids a n order of magnitude higher in salinity a t t h e s a m e temperature. F o r lead t h e dominant dissolution reaction in chloride solutions is PbS + 2H+

Figure 1.4a, b, and c. Fluid-mixing relations for the geothermal systems shown in Figure 1.2.

+ 2C1-

= PbC12

+ H2S

s o t h a t pH, chloride concentration and H2S content a r e t h e solubility controlling variables a t a given temperature. Clearly, t h e higher t h e salinity and lower t h e pH, t h e more m e t a l is dissolved, whereas high H S contents limit t h e solubility (Fig. 1.5a). (Note & a t in low chloride concentrations a t near

1.0

.

0.1

B

-

0.0 1

I I

.

.

:I ipb j / : I ; :

Ag/

I

/ : I :

i

9

OCEAN I

2

3

4

I

5

Log CI q / k g

I

I

I

200

250

300

Tempemturn

OC

log CI mglkg

1

Figure 1.5. Gold, galena, and argentite solubility (mg/kg) versus (a) temperature and (b) ligand concentration for mineral-buffered hydrothermal fluids (calculated from the data summarizetj. in Henley et al., 1984, Chapter 9). In Figures 1.5a, b, and c, the fluid is considered buffered with respect to pH by the assemblage Kmica-Kfeldsparquartz and with respect to fH2 by the empirical relation for the assemblage writeFe-silicate-quartz derived by Giggenbach (1980). The pH of the fluid decreases to the right and f increases upward. Figures 1.5a, b, ant c refer to examples discussed in the text. Eletal contents of the Broadlands system fluids (66 mg/kg H2S) are shown for reference. As an exercise, convert these data to values representing solutions containing 100 mg/kg H2S. The slopes of the solubility curves shown relate to well-established thermodynamic data for the metal complexes considered, but relative solubilities may be in error due to the absence of reliable solubility constants for PbS and Ag2S., In the low-salinity fluids (see text), b~sulfide complexes of silver and hydroxy-carbonate complexes of lead may allow higher solubilities than calculated on the basis of chloride complexing alone. The stippled region in Figure 1.5a emphasizes the temperature-salinity-metal concentration range of principal interest in epithermal studies. Figure 1 . 5 ~ ~ for 250°c, includes an estimate of the solubility of Ag2S as Ag(HS)2- at low salinities and as chloride complexes at higher salinity. The ordinate in the case shows total reduced sulfur rather than H2S, and this introduces the curvature at low salinities (from Henley, 1986).

CHAPTER 1

+Ag

' 7 Million o u n c e s A u

1000

0.01

b

Initial solution:

Cortez, Gold Acres

efficiency

(%I

Carlln

C I 1OOOmglkg Gold

Ag

:

.-. -.

AU

10

million ounces Ohnakl Pool

.-'---. -. k

B R 2 2 Sulfide Scale

2 Kuroko

log time

4

6

years

unnalua.0

Figure 1.6. Au-Ag as a function of total gold and silver for a number of precious-metal deposits and the examples discussed in the text (modified from Graykal, 1981). Note that the ratio shown for case A is based on chloride complexing of silver, so that a more appropriate estimate for low-salinity fluids may be gauged from the ratio observed in the Broadlands BR 22 precipitate.

The host rocks of t h e Broadlands system a r e really quite unremarkable silicic volcanic rocks and, a t depth, greywackes. Irrespective of t h e host rock composition, a s shown above, significant gold transport and deposition can occur in relatively short t i m e periods. The formation of a n economic deposit is therefore more a function of t h e hydrology and chemistry of t h e system than i t is of t h e availability of unusual host-rock gold contents. Availability of t h e metal(s) t o solution is, however, a f a c t o r which may contribute t o t h e transport and deposition efficiency of t h e system a s a whole. In essence, f o r o r e exploration t h e recognition of the source of o r e metals may pale into insignificance relative t o t h e recognition of a source for t h e metal-transporting ligand. A t Waiotapu, Hedenquist and Henley (1985a) showed, using appropriate e s t i m a t e s of t h e pH, fH2, and m ~ of2 t h~e deep fluid, t h a t t h e size of t h e system was such a s t o be f b l e t o supply a l l t h e r e uired m e t a l for a t least 10 years. In t h a t case in 10 years, with a constant h e a t and mass flux equivalent t o t h e present, about 3.6 x lo7 grams (1.2 million ounces) of gold could have been transported, but since t h e hydrology of t h e near-surface system appears t o allow only one s i t e with t h e focused

8

Figure 1.7. Gold deposition as a function of time, flow rate (10 kg/s) and process efficiency for a fluid initially at 3 0 0 ~(see ~ Fig. 1.5). The flow rate relates to that of a typical hot spring whereas the total flow of most geothermal systems is ten to forty times larger. How does this affect the position of the curves7 flaximum and minimum lifetimes for hydrothermal systems are shown for reference (for discussion, see Henley and Ellis, 1983).

depositional process, t h e overall efficiency of t h e hydrothermal system a s a m e t a l concentrator is quite low, around lo%, t h e remaining 90% being disseminated widely through t h e shallow boiling zones. Chemistry of Systems Responsible f o r Ore Formation I t is immediately obvious from Figure 1.5a, t h a t t h e prime requirement f o r t h e formation of a gold deposit is $ fluid relatively high in H2S but of low salinity ( 10 mg/kg C1). The systern temperature is of By lower significance, a s shown by Figure 1.5b. contrast, silver-rich base-metal deposits require fluids of high salinity ( seawater) and t h e e f f e c t of H2S concentration is secondary. The temperature coefficients of solubility f o r t h e s e metals a r e large, but of less significance t h a n salinity in determining m e t a l transport capability. In reviewing fluid-inclusion d a t a from available studies on epithermal systems, Hedenquist and Henley (1985b) confirmed t h e validity of these salinity criteria. In many cases t h e low salinities of t h e fluids responsible for gold deposition were obscured by the presence in solution of dissolved gas, predominantly C 0 2 , which contributes t o t h e freezing point of t h e

inclusion fluid in t h e s a m e way a s o t h e r solutes. For example, 4 wt.-% C 0 2 in solution depresses t h e freezing point by t h e s a m e amount, -1.7'~, a s 2.8 wt.-% NaCl. Since t h e molar C O /H S r a t i o of a c t i v e geothermal-system fluids range2 f?om 10 t o LOO (Giggenbach, 1980), t h e high C O implies high H2S a s required by t h e solution model. h u e t o t h e scarclty of reliable gas analyses f o r fluid inclusions, t h c correlation of high- salinity, low-gas fluids with base metal-silver deposition has yet t o be fully demonstrated. Although no statistical analysis of fluid compositions is possible, i t appears from t h e d a t a available from well-explored geothermal systems t h a t salinities r e f l e c t host rock a n d crustal setting. This is shown schematically in Figure 1.8, which makes t h e hoc assumption of normal frequency distributions f o r fluid compositions in different crustal environments. Salinities f o r basalt-hosted systems a r e lower t h a n those f o r systems hosted by silicic volcanic and these in turn a r e lower t h a n f o r andesite-hosted systems. Doesn't this broadly r e f l e c t t h e ore-host relationships seen in many districts? Fluid-inclusion d a t a suggest t h a t silver deposits of t h e Creede-type formed from fluids of even higher salinity. As discussed elsewhere (Hedenquist and Henley, 1985b), such high-salinity fluids a r e encountered in s o m e t e r r e s t r i a l geothermal systems. Such fluids, s o m e with e x t r e m e l y high salinity, occur in systems typified by those of t h e Imperial Valley, California, within which evaporites a r e present (Rex, 1983), reflecting both t h e t e c t o n i c

-

I I

setting, c r u s t a l rifts, a n d a m b i e n t climate. Using regional geologic d a t a , therefore, i t m a y b e possible t o discriminate hydrothermal systems, both a n c i e n t and modern, which could host silver-base metal mineralization from t h o s e potentially hosting gold, if their gas c o n t e n t s w e r e high enough. Chemical and Physical Processes in O r e Formation If i t is a c c e p t e d t h a t present-day a c t i v e geothermal s y s t e m s a r e t h e a r c h e t y p e s of those responsible f o r e p i t h e r m a l o r e deposition in t h e past, t h e discussion of t h e physics of n a t u r a l geothermal discharges (above) b e c o m e s immediately relevant t o t h e discussion of ore-forming processes. T h e natural discharge of t h e s e l a r g e g e o t h e r m a l s y s t e m s i s focused on highly localized f e a t u r e s such a s hot springs whose locations r e f l e c t both topography and underlying geological structure. A s shown above, t h e s e f e t u r s a r e usually confined t o a n a r e a of t h e o r d e r 10 m , less than 5% of e s u r f a c e a r e a of t h e p a r e n t system (say 7.5 x 107 m ). T h e f e a t u r e common t o a l l these discharge paths was t h e progressive pressure drop from t h a t of t h e upflow s y s t e m t o t h e a m b i e n t pressure of near-surface aquifers or t o atmospheric pressure. Boiling and dilution by near-surface w a t e r s a r e t h e accompanying processes which lead t o mineral deposition along t h e s e flow paths. Although t h e solubility a n d solution chemistry of a f e w m e t a l s have b e e n outlined experimentally t o an e x t e n t sufficient t h a t t h e i r gross transport in these

4 5

Y

GEOTHERMAL FLUID COMPOSITIONS

I

-

Change in equivalent wt % N o ~due l to add~t~on of C02

I

VOLCANIC HOSTED SYSTEMS EVOLVED CONTINENTAL BRINES Salton Sea 25 wt % Cheleken 26 wt % Cesano 13 wt %

recharged systems

EPITHERMAL FLUID COMPOSITIONS

1

Correction to average equ~valentwt % N a ~ l for C02 Gold-silver ores

Kuroko ores

~

Base rnetol ores

-

NaCl wt %

Figure 1.8. Distribution of fluid salinities in the earth's crust in relation to host-rock and crustal environment. A normal frequency distribution has been assumed for each fluid type in the absence of evidence for a continuity of compositions. For discussion, see text and Hedenquist and Henley, 1985b.

hydrothermal systems may be mapped, the understanding of ore-deposit~onal mechanisms is f a r less advanced; indeed, n o experimental studies have been a t t e m p t e d in view of inherent difficulties. Simple kinetics suggest, however, t h a t optimum deposition r a t e s a r e achieved from t h e most supersaturated solutions. This allows t h a t discussion of o r e deposition rnay concentrate on t h e major changes in solution chemistry consequent on boiling or dilution. The e f f e c t s of adiabatic boiling and dilution o n solution chemistry h a v e been discussed elsewhere (Ellis, 1970; Henley and Brown, 1985, this volume; Henley e t al., 1984; Drurnmond and Ohmoto, 1985). Boiling is a n especially important process because t h e formation of only a f e w percent of vapor allows t h e loss of more than 90% of dissolved C 0 2 , with a concomitant pH increase by more than one pH unit and t h e loss of H2S. Consider t h e e f f e c t s of boiling and dilution o n t h e c a s e history solutions A and B discussed above. In both cases t h e t e m p e r a t u r e change due t o t h e s e processes is taken for illustration a s 50°C. Lead and silver supersaturations, a t t a i n e d by dilution with fresh water (with host-rock pH buffering), a r e 500 and 100, respectively, due t o changes in temperature, pH, and chloride concentration. A t C r e e d e , Colorado, f o r example, Hayba (1984) and Hayba et al. (1985, this volume have shown t h a t t h e mixing of high-salinity (7.2 x 10 m g C1-/kg) fluid with surficial steam-heated water was contemporaneous with o r e deposition. The exsolution of C 0 2 during such mixing cooling may also contribute t o base-metal and silver deposition in s o m e environments. The supersaturations attained due t o , a r e 2000 and adiabatic boiling ( ~ p H = + 1 . 5AH+-80%) 300 for P b and A g (chloride s p e c ~ e s ) ,respectively. Figure 1.5a suggests t h a t t h e mineral-buffered solubility of gold increases with dilution. This occurs in response t o t h e increasing pH of t h e silicatebuffered solution a s t e m p e r a t u r e falls and clearly gold deposition cannot be related t o simple dilution. If acidic fluids a r e t h e dilutant, a s may be t h e c a s e in t h e high-level andesite-hosted systems, fluid mixing c a n cause gold deposition. In this case, a relative pH decrease lowers t h e solubility of gold allowing a paragenesis of gold-kaolinite-alunite t o occur in t h e mixing zone between high-level acid-sulfate-chloride w a t e r s (with associated advanced-argillic alteration) and deeper near-neutral pH chloride waters (with associated propylitic alteration). The high arsenic content of t h e o r e may also r e f l e c t this environment (see below). In many epithermal deposits, gold is clearly associated with mineralogical indicators of boiling (open-space filling calcite, adularia, sulfides). The fluid d a t a from Broadlands (Brown, 1985) also very clearly demonstrate t h e effectiveness of boiling a s a process for t h e deposition of gold, silver, and copper, e a c h in solution a s a bisulfide complex (Henley and Brown, 1985, this volume). In this case, t h e initial pH increase due t o C 0 2 loss may undersaturate t h e fluid but sustained loss of t h e more soluble H2S with t h e formation of only a few percent s t e a m leads t o supersaturation and deposition.

t

The loss of H2 due t o boiling gives a n apparent increase in t h e redox s t a t e of t h e fluid (relative t o t h e pyrite-pyrrhotite stability boundary) and also leads t o a n increase in solubility inside t h e H2S stability field. Quantitative gold deposition may b e e x p e c t e d if t h e final redox s t a t e corresponds t o some point in t h e SO; stability field but this poses some severe headaches with respect t o t h e stabilities of accessory minerals (For and t h e kinetics of H S H20-SO: reactions. f u r t h e r discussion, s e e fIinley and Brown, 1985, this volume, and R e e d and Spycher, 1985, this volume). As discussed by Thorstenson (19841, modelling of t h e redox response of a fluid subject t o such a nonequilibrium process a s boiling is fraught with difficulty both in concept and in dealing with t h e redistribution of electrons over t h e large number of e l e c t r o a c t i v e couples available in natural fluids. In t h e discharge of well BR22 a t Broadlands, gold- and silver-ore deposition is largely c o m p l e t e within a few seconds of boiling a t t h e well-head This observation throttling plate (Brown, 1985). suggests t h a t t h e loss of ligands (e.g., HS- a s H2S) is most important. Reaction equations like

illustrate this and emphasize t h a t in discussing t h e deposition of gold under non-equilibrium conditions i t is necessary t o identify sources of electrons f o r t h e f u r t h e r reduction of aurous ions t o t h e metal. Wgh gold a t concentration of t h e order of Ipg/kg (3 x 10molal), t h e r e a r e a number of possibilities because of t h e extremely low electronegativity of gold. The oxidation of other dissolved metals and of reduced sulfur a r e all possible sources. The deposition of silver in electrum proceeds in t h e s a m e manner. A role for arsenic in gold transport a n d deposition has often been suspected. Some experimental d a t a suggest t h a t arsenic complexing may increase t h e solubility of gold in reduced sulfur solutions, but t h e field d a t a from Broadlands comparing observed gold content with t h e solubility calculated from bisulfide complexes suggest t h a t t h i s e f f e c t is minor. The high solubilities of arsenic sulfides in alkaline-sulfide solutions (Seward, 1984) suggest t h a t thioarsenide complexes occur, but t h e s e may well be supplanted by arsenites in moderate19 acid-neutral pH deep, system waters. If so, thioarsenides may provide a sink for H2S in residual, relatively high-pH w a t e r s derived by boding. Arsenic concentrations in geothermal waters a r e usually in t h e range 1-10 mg/kg and could adequately lock up residual H S. ~ i e l c fexperiments (Brown e t al., 1983) involving geothermal waters suggest t h a t amorphous arsenic sulfide may be precipitated by acidification through reactions of t h e form

Natural examples of this process occur in t h e Tamagawa Hot Springs, Japan (Nakagawa, 1971) and in

a number of springs in t h e New Zealand and Yellowstone geothermal areas. A t Champagne Pool, Waiotapu, for example, t h e large surface a r e a results in h e a t loss t o t h e atmosphere and internal convection t o a depth of about 7 5 meters. With a mass influx of about 20 kg/s, d e e p fluid entering t h e pool is quickly cooled and t h e pH buffered internally t o about 5 by C02-HCO; equilibration (Hedenquist and Henley, 1985a). A similar internally buffered process could be invoked f o r deposits like t h e sediment-hosted deposits a t Getchell, Nevada. Both in natural occurrences in New Zealand and in t h e geothermal field experiments, t h e amorphous arsenic sulfide precipitate is ore grade in gold and personal communication, silver. Seward (D.S.I.R., 1985) and Ellis (1969) have suggested t h a t colloidal arsenic sulfide scavenges gold from solution. Recrystallization t o realgar and orpiment with f r e e gold and associated minerals such a s stibnite and cinnabar may t h e n account for t h e late-stage, lowgrade o r e assemblages occurring a t Getchell and elsewhere. Together with t h e zonation of t r a c e m e t a l s observed in t h e upper p a r t of t h e Broadlands and other geothermal systems (Ewers and Keays, 19771, t h e association of gold with arsenic is commonly indicative of shallow depositional environments. As noted above, t h e elemental association of gold and arsenic is also common in t h e Goldfield-type Goldfield, Nevada; Summitville, deposits, e.g., Colorado. In t h e s e environments mixing of descending high-level acid-sulfate waters with normal hydrothermal chloride-dominant waters may account for t h e deposition of gold and arsenic (as enargite) in association with a n advanced-argillic alteration assemblage, t h e m e t a l s derived from t h e deep system. The association of gold with organic m a t t e r has raised questions concerning gold transport and deposition in deposits such a s Carlin and others in northern Nevada. Interaction of normal hydrothermalsystem waters with organic-rich calcareous host rocks leads t o t h e r m a l maturation of t h e hydrocarbon a s suggested by Ilchik (1984) for t h e Alligator Ridge deposit and documented f o r the C e r r o P r i e t o geothermal system (Barker and Elders, 1979). L i t t l e else distinguishes t h e possible mineralogical response of t h e s e rocks t o hydrothermal alteration and i t is most likely t h a t t h e Nevada deposits result from subhot spring boiling a s discussed above. Finely dispersed carbon may have a secondary role t o play in t h e extraction of gold from solution, although even this interaction is not well supported by mineralogical d a t a (Wells and Mullens, 1973). Host-Rock Reactions Whereas mixing and boiling a r e processes common t o all hydrothermal systems, specific interactions with host rocks which may deposit o r e a r e f a r less common. An exception is t h e interaction of relatively low-pH, high-salinity fluids with carbonate rocks which may, through loss of acidity, deposit massive replacement ores of silver and base m e t a l s (e.g., Taxco, Mexico). Interaction of a high-H2S fluid with a n iron-rich sediment may be a r a r e possibility for t h e formation of a gold-pyrite assemblage.

SUMMARY The original question posed a t t h e beginning of this chapter was "why study geothermal systems in t h e c o n t e x t of t h e origin of epithermal o r e deposits?" I t is clear from t h e foregoing paragraphs t h a t such studies highlight a number of important f a c t o r s necessary for t h e understanding of chemical and physical processes in epithermal systems. These may b e summarized a s follows: 1. The a c t i v e high-temperature geothermal systems in volcanic-rock t e r r a n e s a r e t h e archetypes of those systems responsible for epithermal preciousand base-metal o r e deposits in analogous ancient terranes. Porphyry copper-molybdenum deposits represent crust-magma interactions a t somewhat deeper levels within calc-alkalic volcanic-rock hosted systems and Kuroko-type massive sulfides a r e essentially sea-floor telescoped equivalents of the t e r r e s t r i a l epithermal deposits. 2. Studies of a c t i v e geothermal systems provide insight into t h e physical processes governing flow t o surface discharge f e a t u r e s such a s hot springs and near-surface interaction with ground and other hot waters. Geological s t r u c t u r e provides a principal control, but in many geothermal fields hydrothermal eruptions focused a t depths of about 300 m e t e r s provide focused flow paths t o t h e surface. The c h a r a c t e r i s t i c hydrothermal eruption breccias formed by t h e s e events have been recognized in a number of epithermal precious-metal deposits such a s Round Mountain, Nevada (Tingley and Berger, 1985), and McLaughlin, California. Figure 1.9 provides a n interesting c a s e history for t h e reconstruction of a n epithermal system based oq field observations, fluid inclusions, and isotope techniques. Such reconstructions a r e important t o t h e exploration geologist for targeting drilling and for comparison with o t h e r districts under exploration where complementary but piece-meal d a t a may b e available. 3. Studies of a c t i v e geothermal systems provide insight into t h e range of fluid compositions present in t h e crust. In conjunction with fluid and laboratory studies i t may b e shown t h a t t h e formation of goldsilver (base-metal poor) epithermal deposits requires fluids of low salinity and high gas content (i.e., H2S in association with C O ), whereas silver-rich base-metal deposits (relatively Tow in gold) require high-salinity, low-gas fluids in their respective hydrothermal systems. The salinities of hydrothermal fluids r e f l e c t their volcanic-tectonic environment. For example, andesite-hosted hydrothermal systems a r e , perhaps through high-level interaction with volcanic gases or t h e involvement of deep, connate waters, m o r e saline t h a n those of silicic volcanic-rock terranes, but less saline than those encountered where evaporite sequences occur (grabens, some caldera moats); t h e l a t t e r reflecting also a climatic control. Although acid-sulfate waters a r e locally encountered in silicicrock hosted systems, they a r e quite abundant i n highlevel andesitic-rock terranes, and may b e important components in t h e formation of certain types of gold deposits. Such studies also highlight t h e important role of

GEOLOGY $ ALTERATION

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D

!r&

Figure 1.9a, b, and c. Distribution of alteration types and fluid-inclusion data for the Round Mountain gold deposit, Nevada. Data from Tingley and Berger, 1985. The topographic outline shown is drawn from the viewpoint shown in a). Hypothetical reconstruction of the Round Mountain hydrothermal system (25 million years b.p.) emphasizes the dissection of the system by erosion. Much controversy has centered on the so-called hot-spring origin of this deposit; clearly the deposit is hot-spring related but is not itself a hot spring1 There seems to be little advantage to the use of the term "hotspring-type deposit" over the generally applicable "epithermal" unless very specific evidence is available for a given deposit relating it to a hot-spring environment (e.g., silica sinter, eruption breccia) as at McLaughlin, California. The hot-spring terminology may indeed be far tcx, restrictive for the guidance of exploration programs even though the recognition of hotspring features is a very important tool for the exploration geologist.

gas composition in controlling t h e physical processes within systems and their m e t a l transport capability, a s well a s demonstrating t h a t t w o processes (adiabatic boiling and mixing) provide t h e principal controls on fluid chemistry in t h e epithermal environment. Access t o s o m e gas source ( C 0 2 + H S) becomes, therefore, paramount in determining whe&er a given low-salinity system m a y or may not develop a gold deposit. As discussed elsewhere (Henley, 19861, t h e r e may b e t e c t o n i c controls here (e.g., a c c e s s t o gas from subduction zone metamorphism?) which introduce t h e link between t e c t o n i c setting, fluid chemistry, and o r e formation, which together provide guides t o o r e search. 4. Active geothermal systelns provide t h e opportunity for studying t h e deposition of t r a c e m e t a l s such a s gold and t h e initial concentrations of oreforming elements in t h e deep system. Geothermal systems such a s Broadlands show t h a t m e t a l concentrations may b e regarded a s close t o saturation and demonstrate t h a t host rocks play only a passive role in determining whether systems a r e likely t o deposit o r e near t h e surface. The principal control on t h e l a t t e r is t h e provision of permeable f e a t u r e s within which flow may b e focused and adiabatic boiling and/or dilution may occur. The kinetics of m e t a l deposition during irreversible boiling a r e not well understood and hard t o model realistically, but pH and H S concentration appear t o be more important than reTatively slow redox reactions. More specialized environments involving gas buffering of pH o r mixing with acid waters may be indicated by t h e association of gold with orpiment-realgar or enargite-alunite assemblages.

or subsequent volcanism and tectonism which obscure A useful and dismember t h e original structure. !'thought exercise (or, in modern parlance, experiment") is t o imaginatively modify copies of Figure 1.1 by inclusion of layers of volcanic material and/or f a u l t dissection t o explore s o m e of t h e problems which arise in reconstruction based on geological evidence. In tackling this exercise, remember t h e three-dimensional aspects of t h e systems. (c) Discriminant analysis is a t e r m used in statistics t o describe a method for classifying multivariate observations into groups. In t h e above paragraphs, the variability of surface expression a n d chemistry has been stressed but a s noted by Koch and Link (1970, p.326) this is rather pointless if no ways a r e established t o s o r t o u t t h e s e kinds of variability. This is t h e essence of t h e design of scientific experiments, but exploration for epithermal o r e deposits is seldom e f f e c t e d in this manner, is it? Since t h e coincidence of a suitable flow s t r u c t u r e , f r a c t u r e pattern and system chemistry is responsible for ore-metal deposition rather t h a n dispersion, t h e probability of exploration success is t h e f a c t o r requiring early determination in a n exploration program. F a c t o r s such a s a previous mining history a r e obvious high-score factors a s is recognition of s t r u c t u r a l style. Modern techniques such a s isotope analysis and fluid-inclusion studies should also b e used in a discriminatory manner; e.g., fluid inclusions may very early in a project determine whether a system is relatively dilute and therefore not likely t o host base metal-silver o r e or relatively gassy and t h e r e f o r e quite likely t o have transported and perhaps deposited gold. Understanding of t h e geochemistry of m e t a l transport and deposition rnay b e used in t h e s a m e way.

EPILOGUE If physical and chemical processes alone a r e t h e control on whether or not a system may deposit a n o r e body, what chance does t h e exploration geologist have in targeting exploration effectively? The answer is in t h r e e parts: (a) Since active geothermal systems a r e relatively similar in t e r m s of t h e distribution of alteration minerals, these assemblages may b e used t o provide key d a t a on t h e level of exposure within a system; clays and chlorites, for example, a r e quite sensitive indicators of temperature'of formation (Browne, 1978; Henley and Ellis, 1983). Recognition of other depth indicators such a s hydrothermal eruption or vent breccias is also a powerful targeting tool. Similarly t h e recognition of trace-element enrichments (As, Sb, Hg, TI) provides a clue t o structural level in t h e system. (b) Good qlold-fashioned' structural studies have been much neglected in r e c e n t decades, but they probably still provide t h e most e f f e c t i v e means of targeting drilling since structure, particularly f r a c t u r e analysis provides t h e main clue t o flow s t r u c t u r e in t h e fossil system. Knowledge of t h e hydrodynamic characteristics common t o active geothermal systems also provides a guide t o t h e exploration of fossil systems. Additional f a c t o r s t o remember a r e t h e e f f e c t s of contemporary

ACKNOWLEDGMENTS Discussions with K. L. Brown, R. G. Allis, P. B. Barton, and P. M. Bethke have been valuable in preparing this overview. T h e musical accompaniment of J a m e s Galway was especially soothing t o t h e writer and may well provide solace t o t h e reader! REFERENCES Allis, R. G., Henley, R. W., a n d Carman, A. F., 1979, The thermal regime beneath t h e Southern Alps; in Walcott, R., and Cresswell, M. (eds.), Origin of t h e Southern Alps: Royal Society of New Zealand Bulletin 18, p. 79-85. Barker, C. E., and Elders, W. A., 1979, Vitrinite r e f l e c t a n c e geothermometry in t h e C e r r o P r i e t o geothermal field, Baja California, Mexico: Geothermal Resources Council Transactions, V. 3, p. 27-30. Barnes, H. L., 1979, Solubilities of ore minerals; & Barnes, H. L. (ed.), Geochemistry of Hydrothermal O r e Deposits: 2nd ed., John Wiley and Sons, New York, p. 404-460. Barnes, H. L., and Czamanske, G. K., 1968, Solubilities a n d transport of o r e minerals; & Barnes, H. L.

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(ed.), Geochemistry of Hydrothermal O r e Deposits: 1st ed., Holt, Rinehart, and Winston, New York, p. 334-381. Barton, P. B., Jr., a n d Toulmin, P., 1961, Some mechanisms f o r cooling hydrothermal fluids: U.S. Geological Survey, Professional Paper, 424-D, p. 348-352. Batini, F., Bertini, G., Gianelli, G., Pandeli, E., and Puxeddu, M., 1983a, San Pompeo 2 Deep Well--A high-temperature and high-pressure geothermal system: Proceedings 3rd International Seminar, Results of E C R e s e a r c h and Demonstration P r o j e c t s i n t h e Field of Geothermal Energy: European Geothermal Update, Extended Summaries, European P a t e n t Office, Munich, p. 341-353. Batini, F., Bertini, G., Gianelli, G., Pandeli, E., and Puxeddu, M., 1983b, Deep s t r u c t u r e of t h e Larderello Field--contribution from r e c e n t geophysical a n d geological data: Memoir, Society of Geology, Italy, v. 27 (in press). Belkin, H., Vivo, B. D., Gianelli, G., and Lattanzi, P., 1983, Fluid inclusion reconnaissance study of hydrothermal minerals from geothermal fields of Tuscany (Italy): Extended Abstracts, 4th International Symposium on Water Rock Interaction, Misasa, Japan, p. 43-47. Berger, B. R., a n d Eimon, P., 1983, Conceptual models of epithermal precious-metals deposits; in Shanks, W. C., 111 (ed.), Cameron Volume Unconventional Mineral Deposits: Society of Mining Engineer, A.I.M.E., p. 191-205. Bodnar, R. J., Reynolds, T. J., and Kuehn, C. A., 1985, Fluid inclusion s y s t e m a t i c s in epithermal systems; Berger, B. R., and Bethke, P. M. (eds.), Geology and Geochemistry of Epithermal Systems: Society of Economic Geologists, Reviews in Economic Geology, v. 2. Brown, K. L., 1985, Gold deposition from New Zealand geothermal wells: Economic Geology (in press). Brown, K. L., McDowell, G. D., Lichti, K. A., and Bijnen, H. J. C., 1983, pH control of silica scaling: Proceedings 5th New Zealand Geothermal Workshop, p. 157-162. Browne, P. R. L., 1978, Hydrothermal alteration in a c t i v e geothermal fields: Annual Review of E a r t h and Planetary Sciences, v. 6, p. 229-250. Casadevall, T. J., dela Cruz-Reyna, S., Rose, W. I., Bagley, S., Finnegan, D. L., and Zoller, W. H., 1984, Crater Lake and post-eruption hydrothermal activity: Journal of Volcanology a n d Geothermal Research, v. 23, p. 169-191. Cathles, L. M., 1977, An analysis of t h e cooling of intrusives by groundwater convection which includes boiling: Economic Geology, v. 72, p. 804-826. Drummond, S. E., and Ohmoto, H., 1985, Chemical evolution a n d mineral deposition in boiling hydrothermal systems: Economic Geology, v. 80, p. 126-147. Elder, J. W., 1966, Hydrothermal systems, h e a t and mass transfer in t h e earth: New Zealand D.S.I.R. Bulletin 169, 115 p. Ellis, A. J., 1969, Present-day hydrothermal systems a n d mineral deposition: Ninth Common Mining

on

and Metallurgy Congress, Institute of Mining and Metallurgy (London) 1-30. Ellis, A. J., 1970, Quantitative i n t e r p r e t a t i o n of chemical characteristics of hydrothermal systems: Proceedings from t h e United Nations Symposium o n t h e Development a n d Utilization of Geothermal Resources, v. 2, p. 516-528. Ellis, A. J., and Mahon, W. A. J., 1977, C h e m i s t r y and Geothermal Systems: Academic Press, New York, 392 p. Ellis, A. J., and Wilson, S. H., 1955, The h e a t f r o m t h e Wairakei-Taupo t h e r m a l region calculated from t h e chloride output: New Zealand Journal of S c i e n c e and Technology, Section B., v. 36, (61, p. 622-631. Ewers, G. R., and Keays, R. R., 1977, Volatile and precious-metal zoning in t h e Broadlands Economic geothermal field, New Zealand: Geology, v. 72, p. 1337-1354. 1981, Application of water Fournier, R. O., geochemistry t o geothermal exploration a n d reservoir engineering; in Rybach, L., and Muffler, L. P. J. (eds.1,-~eothermal Systems: Principles and C a s e Histories: John Wiley and Sons, New York, p. 109-143. Fournier, R. O., White, D. E., and Truesdell, A. H., 1975, Convective h e a t flow in Yellowstone National Park: Second United Nations Symposium on t h e Development a n d Use of Geothermal Resources, San Francisco, v. I , p. 731-749. Fyfe, W. S., and Kerrich, R., 1984, Gold: natural concentration processes; & Foster, R. P. (ed.), Gold '82, The Geology, Geochemistry, and Genesis of Gold Deposits, Geological Society of Zimbabwe Special Publication No. 1, p. 99-128. Giggenbach, W. F., 1974, T h e chemistry of C r a t e r Lake, Mt. Ruapehu (New Zealand) during and a f t e r t h e 1971 a c t i v e period: New Zealand Journal of Science, v. 17, p. 33-45. Giggenbach, W. F., 1980, Geothermal-gas equilibria: Geochimica e t Cosmochimica Acta, v. 44, p. 2021-2032. Giggenbach, W. F., 1981, Geothermal-mineral equilibria: Geochimica e t Cosmochimica A c t a , V. 45, p. 393-410. Giggenbach, W. F., Geonfiantini R., Jangi, B. L., and Truesdell, A. H., 1983, Isotopic a n d chemical composition of P a r b a t i Valley geothermal discharges, N.W. Himalaya, India: Geothermics, V. 12, p. 199-222. Grant, M. A., Donaldson, I. A., and Bixley, P. F., 1982, Geothermal Reservoir Engineering: Academic Press, New York, 369 p. Graybeal, F. T., 1981, C h a r a c t e r i s t i c s of disseminated silver deposits in t h e Western United States; in Dickinson, W. R., a n d Payne, W. D. ( e d s x Relations of Tectonics t o O r e Deposits in t h e Southern Cordillera: Arizona Geological Society Digest, v. 14, p. 271-281. Haas, J. L., Jr., 1971, T h e e f f e c t of salinity on t h e maximum t h e r m a l gradients of a hydrothermal system a t hydrostatic pressure: Economic Geology, v. 66, p. 940-946. Hanaoka, N., 1980, Numerical model experiment of

hydrothermal system-topographic effects: Geological Survey Bulletin, Japan, v. 31 (7), p. 321-332. Hayba, D. O., 1984, Documentation of t h e r m a l a n d salinity gradients a n d interpretation of t h e hydrologic conditions in t h e OH vein, C r e e d e , Colorado: Geological Society of America, Abstracts with programs, v. 16, p. 534. 1984, Isotopic Kyosu, Y., and Kurahashi, M., geochemistry of acid t h e r m a l waters a n d volcanic gases from Zao volcano, Japan: Journal of Volcanology and Geothermal Research, v. 21, p. 31 3-332. Koch, G. S., a n d Link, R. F., 1970, Statistical Analysis of Geological Data: John Wiley and Sons, New York, 438 p. Lindgren, W., 1933, Mineral Deposits: IvlcGraw-Hill Book Company, New York, F o u r t h Edition, 930 p. Mahon, W. A. J., Klyen, L. E., and Rhode, M., 1980, Natural sodium bicarbonate sulphate h o t w a t e r s in geothermal systems: Chinetsu (Journal J a p a n Geothermal Energy Association), v. 17, p. 11-24. Muffler, L. J. P., White, D. E., and Truesdell, A. H., 1971, Hydrothermal explosion c r a t e r s in Yellowstone National Park: Geological Society of A m e r i c a Bulletin 82, p. 723-740. Nakagawa, R., 1971, Solubility of orpiment (As2S3) in T a m a g a w a Hot Springs, Akita Prefecture: Nippon Kagaku Zasshi, v. 92, p. 154-159. Ozima, M., Takay na i M., Zashu, S., and Amari, S., r a t i o in o c e a n sediments, 1984, High 'HJ'He Nature, v. 31 1, p. 448-450. Puxeddu, M., 1984, S t r u c t u r e and L a t e Cenozoic evolution of t h e upper lithosphere in southwest Tuscany (Italy): Tectonophysics, v. 101, p. 357-382. Ransome, F. L., 1909, The geology and o r e deposits of Goldfield, Nevada: U.S. Geological Survey, Professional Paper 66, 258 p. Reed, M. H., a n d Spycher, N. F., 1985, Boiling, cooling, and oxidation in epithermal systems: Numerical modeling approach; Berger, B. R., and Bethke, P. M. (eds.), Geology and Geochemistry of Epithermal Systems: Society of Economic Geologists, Reviews in Economic Geology, v. 2. Rex, R. W., 1983, Origin of t h e brines of t h e Imperial Valley, California: Annual Meeting, Geological Society of America, Program with Abstracts, Indianapolis. Roedder, E., 1984, Fluid inclusion evidence bearing on t h e environments of gold deposition; & Foster, R. P. (ed.), Gold '82, The Geology, Geochemistry and Genesis of Gold Deposits: Geological Society of Zimbabwe Special Publication, no I , p. 129-164. Seward, T. M., 1984, The transport and deposition of gold in hydrothermal systems; Foster, R. P. (ed.), Gold '82, The Geology, Geochemistry and Genesis of Gold Deposits: Geological Society of Zimbabwe Special Publication, no. 1, p. 165-181. Sillitoe, R. H., 1973, The tops and b o t t o m s of porphyry copper deposits: Economic Geology, v. 68, p. 799-815. Sillitoe, R. H., 1983, Enargite-bearing massive sulfide deposits high in porphyry copper systems:

Economic Geology, v. 78, p. 348-352. Sternfeld, J. N., 1981, Hydrothermal petrology a n d s t a b l e isotope geochemistry of t w o wells in t h e Geysers geothermal field, Sonoma County, California: Unpublished M.S. thesis, University of California (Riverside), 202 p. Hayba, D. O., Bethke, P. M., Heald, P., and Foley, N. K., 1985, Geologic, mineralogic, and geochemical characteristics of volcanic-hosted epithermal precious-metal deposits; & Berger, B. R., and Bethke, P. M. (eds.), Geology and Geochemistry of Epithermal Systems: Society of Economic Geologists, Reviews in Economic Geology, v. 2. Healy, J., and Hochstein, M. P., 1973, Horizontal flows in geothermal systems: Journal of Hydrology (New Zealand), v. 21, p. 71-82. Hedenquist, J. W., and Henley, R. W., 1985a, Hydrothermal eruptions in t h e Waiotapu Origin, geothermal system, New Zealand. breccia deposits and e f f e c t on precious-metal mineralization: Economic Geology, v. 80, p. 1640-1668. Hedenquist, J. W., and Henley, R. W., 1985b, E f f e c t of C O on freezing-point depression measurements of f3uid inclusions--evidence from a c t i v e s y s t e m s a n d application t o epithermal studies: Economic Geology, v. 80, p. 1379-1406. Hedenquist, J. W., and S t e w a r t , M. K., 1985, Natural COZ-rich steam-heated waters in the Broadlands-Ohaaki geothermal system, New Zealand. Their chemistry, distribution and corrosive nature: Geothermal Resources Council, Transactions, v. 9 (21, p. 247-250. Henley, R. W., 1986, O r e transport and deposition in epithermal environments; & Herbert, H. (ed.), S t a b l e Isotopes and Fluid Processes in Mineralization: Geological Society of Australia Special Issue (in press). Henley, R. W., and Brown, K. L., 1985, A practical guide t o t h e chemistry of geothermal a n d epithermal systems; Berger, B. R., and Bethke, P. M., Geology and Geochemistry of Epithermal Systems: Society of Economic Geology, Reviews in Economic Geology, v. 2. Henley, R. W., and Ellis, A. J., 1983, Geothermal systems, a n c i e n t and modern: E a r t h Science Reviews, v. 19, p. 1-50. Henley, R. W., and McNabb, A,, 1978, Magmatic vapor plumes and ground water interaction in porphyry copper emplacement: Economic Geology, v. 73, p. 1-20. Henley, R. W., Norris, R. J., and Paterson, C. J., 1976, Multistage o r e genesis in t h e New Zealand geosyncline--a history of post-metamorphic lode Mineralium Deposita, v. 1 I, emplacement: p. 180-196. Henley, R. W., Truesdell, A. H., and Barton, P. B., Jr., 1984, Fluid-Mineral Equilibria in Hydrothermal Systems: Society of Economic Geologists, Reviews in Economic Geology, v. 1, 267 p. Ilchik, R. P., 1984, Hydrothermal maturation of indigenous organic m a t t e r a t t h e Alligator Ridge gold deposits, Nevada: Unpublished M.S. thesis, University of California (Berkeley), 77 p.

Kieffer, S. W., 1984, Seismicity of Old Faithful Geyser--an isolated source of geothermal noise and possible analogue t o volcanic seismicity: Journal of Volcanology and Geothermal Research, v. 22, p. 59-96. Sutton, F. M., and McNabb, A., 1977, Boiling curves a t Broadlands geothermal field, New Zealand: New Zealand Journal of Science, no. 20, p. 333-337. Thorstenson, D. G., 1984, The concept of electron activity and its relation t o redox potentials in aqueous geochemical systems: U.S. Geological Survey, Open-File Report 84-072, 67 p. Tingley, J. V., and Berger, B. R., 1985, Lode gold deposits of Round Mountain, Nevada: Nevada Bureau of Mines and Geology Bulletin 100, 62 p. Weissberg, B. G., 1969, Gold-silver ore-grade precipitates from New Zealand thermal waters: Economic Geology, v. 64, p. 95-108. Weissberg, B. G., Browne, P. R. L., and Seward, T. M., 1979, Ore metals in active geothermal systems; in Barnes, H. L. (ed.), Geochemistry of

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Hydrothermal Ore Deposits, 2nd ed.: John Wiley and Sons, New York, p. 738-780. Wells, J. D., and Mullens, T. E., 1973, Gold-bearing arsenian pyrite determined by microprobe analysis, C o r t e z and Carlin gold mines, Nevada: Economic Geology, v. 68, p. 187-201. White, D. E., 1981, Active geothermal systems and hydrothermal o r e deposits: Economic Geology, 75th anniversary volume, p. 392-423. White, D. E., Muffler, L. J. P., and Trusdell, A. H., 1971, Vapor-dominated hydrothermal systems compared with hot-water systems: Economic Geology, v. 66, p. 75-97. White, D. E., Thompson, G. A., and Sandberg, C. H., 1964, The rocks, structure, and geologic history of t h e Steamboat Springs thermal area, Washoe County, Nevada: U.S. Geological Survey, Professional Paper 458-B. White, D. E., 1955, Thermal springs and epithermal ore deposits; & Bateman, A. M. (ed.): Economic Geology, 50th anniverary volume, p. 99-154.

Chapter 2 A PRACTICAL GUIDE TO THE THERMODYNAMICS OF GEOTHERMAL FLUIDS AND HYDROTHERMAL ORE DEPOSITS R. W. Henley and K. L. Brown

INTRODUCTION In trying t o understand t h e depositional processes which led t o o r e deposition in fossil hydrothermal systems, we a t t e m p t t o reconstruct t h e chemistry of t h e fluid phase from observation of i t s relics (e.g., alteration minerals, fluid inclusions). We may also a t t e m p t t o thermodynamically model t h e chemical changes experienced by this fluid a s i t passes upward through a vein, vents t o t h e seafloor, boils o r mixes with other waters, etc. A number of important assumptions a r e made; one is t h e assumption of equilibrium and another is t h a t t h e thermodynamic d a t a base is sound. Analyses of fluids discharged from geothermal wells, together with drill-core data, allow t h e opportunity t o independently check t h e validity of t h e thermodynamic d a t a base and t o observe directly, chemical processes leading t o t h e deposition of gold, base-metal sulfides and common gangue minerals like quartz and calcite. The calculations involved a r e not trivial, but a r e essential t o t h e understanding of epithermal or any other t y p e of hydrothermal o r e deposit. To illustrate these procedures, we shall examine production well in t h e Broadlands t h e discharge of geothermal field in New Zealand. Through t h e use of thermodynamics, t h e amount of information we shall retrieve about t h e reservoir and depositional processes is quite astonishing. We shall then turn t o s o m e review questions t o consider implications f o r t h e formation of some hydrothermal o r e deposits. In this chapter we have tried t o follow a pragmatic course, avoiding the temptation t o overindulge in the (essential) nuances of thermochemistry a t t h e expense of t h e proscribed goal--a practical understanding of hydrothermal processes. A wider discussion of geothermal chemistry, a broader d a t a base, and references t o t h e current literature a r e presented in Henley et al. (1984).

originally containing q u a r t z and andesine with a glassy or fine grained groundmass a n d minor hornblende, biotite, magnetite and other common accessories. The geology and alteration mineralogy of cores from several wells a r e summarized in Table 2.1. As well a s ubiquitous pyrite and lesser pyrrhotite, copper, lead and zinc sulfides have been recognized in veins and vugs in core frorn s o m e wells, a n d a r e commonly associated with adularia and calcite. A t t h e surface in t h e Ohaaki Pool, i n t e r m i t t e n t deposition of a n amorphous sulfide precipitate has oc'curred on the silica sinter formed a t t h e e d g e of t h e pool. This orange precipitate is o r e g r a d e with r e s p e c t t o gold and silver and rich in As, Sb, TI, and Hg. A similar

b

one

BAY OF PLENTY

B"0~OLANOS .A?: .'"PO

0 10

27

The geology and alteration mineralogy of t h e Broadlands field have been recently reviewed by Weissberg e t al. (1979). Figure 2.1 shows t h e location of exploration and production wells in t h e Broadlands field and Figure 2.2 shows in cross-section t h e structure and stratigraphy of t h e field in relation t o measured subsurface temperatures. The host rocks for t h e hydrothermal system a r e mostly silicic ash-flow

t Ctj

P,

THERMAL AREA

-

1

GEOLOGICAL CHARACTERISTICS O F THE BROADLANDS GEOTHERMAL SYSTEM

I

,

0

500

l000rn

Location of geothermal wells and Figure 2.1. hot springs at Broadlands New Zealand. Wells marked, 8 , intersected base-metal sulfide minerals at depth; wells marked, 8 , deposited precious-metal-bearing, antimony-rich precipitates at the surf ace; , have both surface and wells marked, precipitates and base-metal sulfides at depth (reproduced with permission from Weissberg et al., 1977).

Table 2.1--Distribution of base-metal sulfides in Broadlands drill holes (reproduced from Weissberg et al., 1979) Base-metal sulfide zone

Well No.

Depth of well (m)

Depth range (m) Minimum

Temperature range (OC)

Maximum Minimum

precipitate formed in t h e discharge disposal channels of a number of exploration wells. The depth zonation indicated by these observations is also reflected in chemical analyses of sulfide separates (Ewers and Keays, 1977) and recalls t h e general pattern of m e t a l zonation encountered in many epithermal o r e deposits.

FLUID CHEMISTRY Table 2.2 contains analyses of t h e liquid water and s t e a m separated from the discharge of BR22 a t Broadlands. The discharge from this well is fairly typical, although higher temperature zones a r e encountered a t depth elsewhere in t h e field.

Rock types in base-metal sulfide zone

Associated hydrothermal minerals

rhyolite tuff, greywacke

quartz, adularia, albite, calcite, pyrite, illite, chloriLe

conglomerate, greywacke

quartz, adularia, chlorite, illite

argillite, conglomerate

quartz, adularia, chlorite, illite, pyrite

ignimbrite, tuff

calcite, quartz, chlorite, adularia, albite, pyrite

Broadlands Dacite, Rautawiri Breccia, greywacke

quartz, illite, chlorite, calcite, pyrite, pyrrhotite

Upper Waiora pumiceous tuff breccia

calcite, illite, wairakite, quartz, pyrite

tuff

quartz, pyrite, sphene, zoisite

tuff, tuffaceous sandstone

illite, pyrite, chlorite, sphene

Maximum

Exercise

1.

For later convenience, complete t h e conversion of t h e liquid sample analysis t o molal units. (n.b. Cindicates analytical totals for dissociated species; e.g.

In order t o examine chemical relationships between t h e reservoir mineral assemblage and t h e fluid phage, t h e d a t a must first b e recombined t o t h e single high-temperature and pressure phase which f e d t h e well. For this purpose a h e a t balance equation is used (with t h e assumption t h a t negligible h e a t loss occurs) t o calculate t h e mass of liquid water converted t o vapor a s t h e fluid rises, a t several m/sec t o t h e wellhead.

R. W. HENLEY & K. L. BROWN Table 2.2--Analytical

data for fluids discharged from well BR22, Broadlands, New Zealand

Liquid sample Collection Pressure

11 b.a.

pH20

7.39

Na

K

Ca

870

188

Mg

C1

B

SO4

0.1

1432

42

2

1.2

Sio2 CNH3 XHC03

705

6.1

216 10 3 x moles kg-'

11.75

Steam sample Collection Pressure

C02

H2S

NH3

H2 6.6

lo3 x moles kg-'

5.0

Digression No. 1

Figure 2.2. Geologic cross-section of the Broadlands, New Zealand, geothermal field (reproduced with permission from Weissberg et al., 1 9 7 7 ) .

For convenience in this chapter, Figure 2.3 shows steam fractions for pure water a s a function of enthalpy, pressure and temperature. To use this diagram l o c a t e t h e discharge enthalpy on t h e twophase coexistence curve--this point represents conditions in t h e reservoir. Then project downward a t constant enthalpy t o t h e sampling pressure and read off t h e s t e a m fraction present under t h e reduced pressure conditions. This procedure, or more precisely t h e h e a t balance equation, is t h e basis for t h e calculation of changes in chemistry and t e m p e r a t u r e during boiling (adiabatic, closed system) of hydrothermal fluids in a n ore zone (for example, s e e Fournier (1985, this volume) and Reed and Spycher (1985, this volume)). H e a t and mass balance equations a r e also t h e basis for similar calculations involving t h e mixing of fluids. A mass balance equation is required t o recombine t h e analytical data. For a component, X

XTD = (1

- y)Xl + yXv

(2)

so t h a t , f o r example S i 0 2 TD = 0.81 x 705 mg/kg where H is t h e enthalpy of t h e t o t a l discharge (T.D.), and of liquid water (1) and vapor (v) a t t h e separation pressure; y is t h e s t e a m fraction a t t h e sampling pressure. Note t h a t 100 y is t h e percent of s t e a m by mass in t h e two-phase mixture.] HTD is determined by physical measurement and Substituting for this well is 1160 (1.20) J/g. appropriate values of HI and Hv from tables of thermodynamic d a t a for water (see Henley et al., 1984, Appendix 1111, equation (1) gives y = 0.19

.

and converting t o molality molecular weight of S i 0 2

by

Si02,TD = 9.52 x

dividing by

the

moiesikg

where a volatile component such a s C 0 2 (534.3 x 10- 3 moleslkg in t h e vapor in this example) is considered

''~,TD = 0.81

X

~ H C O+ 0.19 3

X

534.3

X

moleslkg

Figure 2.3. Enthalpy-temperaturepressure-density relations for ~ 600 bars water to 5 0 0 ~and (from Henley et al., 1984).

ENTHALPY ABOVE ' 0 C, Joules/grn

Notice t h a t for C 0 2 we combine all t h e analytical c a r b o n a t e carbon (i.e., C 0 2 , H C O 3 , C 0 5 , and C 0 2 ), whereas f o r silica we may reasonably assume t h a t tke silica content of t h e s t e a m is negligible. Exercise 2.

C o m p l e t e Table 2.3 t o obtain t h e t o t a l discharge composition of t h e well. FLUID-MINERAL EQUILIBRIA: ALTERATION MINERALOGY

What controls t h e concentration of e a c h component represented in t h e analysis? First we examine silica. It is generally assumed t h a t since q u a r t z is ubiquitous in t h e reservoir, i t s solubility controls m . 0 2 in t h e fluid. We can check this through availazfe experimental d a t a f o r t h e solubility of q u a r t z in water (reaction 3)

Figure 2.4 shows original solubility d a t a f o r q u a r t z (Kennedy, 1950; Morey et al., 1962). The observed concentration of silica in t h e aquifer fluid is 570 mg/kg (Table 2.3) s o t h a t inspection of Figure 2.4 yields an aquifer t e m p e r a t u r e of 2 6 5 ' ~ . Exercise 3.

Now check t h e measured enthalpy of t h e discharge against s t e a m t a b l e s or Figure 2.3 t o find a t w h a t t e m p e r a t u r e i t corresponds t o t h e enthalpy of vapors a t u r a t e d liquid. You have probably guessed t h e answer anyway; t = 2 6 5 ' ~ .

C a n you now infer t h a t msio2 t h e solubility of quartz?

is controlled by

Equilibrium constants f o r this and other useful reactions a r e given a s t e m p e r a t u r e dependent regression equations in Appendix Table 2.AI. These equations apply over limited t e m p e r a t u r e ranges. For example, for quartz, t h e equation given assumes a linear relationship between log K and 1/T from 200 t o 2 8 0 ' ~ and conforms t o t h e solubility d a t a recommended by Fournier and P o t t e r (1982). Various other equations have been published in t h e past and r e l a t e t o t h e e x t e n t of experimental d a t a available, t h e authors' bias, t h e regression technique used, and t h e assessment of errors. In this case, t h e equation yields a slightly higher aquifer t e m p e r a t u r e which m a y or may n o t be real. T h e enthalpy measurement and experience in a large number of wells confirm t h a t q u a r t z controls t h e silica c o n t e n t of t h e aquifer fluid above about 200°C (see Fournier, 1985, this volume). Other components m a y be examined in a similar way. For example, f o r t h e hydrolysis reaction 3KA1Si3O8 Kspar

+

2w

which is here written with conservation of aluminum in t h e solid phases. Appendix Table 2.A1 gives log K265 = 7.897 a t 2 6 5 ' ~

R. W. HENLEY& K. L. Table 2.3--Total

BROWN

d i s c h a r g e composition f o r w e l l BR22 moles kg-'

Na

and

K

Ca

Mg

C1

B

log K = 21og a ~ + +2 pH

Si02

(5)

CNH3

ZCO2

H2S

x lo3 H2

CH4

N2

A simplistic calculation o f t e n provides a useful approximation of pH. Consider t h e equilibrium

where a is the thermodynamically e f f e c t i v e concentration of, in this case, potassium ion. * undissociated dissolved where H ~ C O ~represents

co2.

Digression No. 2 We need for a moment t o consider activities, ai, and activity coefficients, Y i , for solution species designated i. Yi provides t h e link between t h e real messy world and t h a t of ideal thermodynamics.

A t 2 5 ' ~ log K = - 6.36 so t h a t in t h e liquid analysis we may calculate (with yHCO3- '0.8 a t laboratory temperature) t h a t mHco3-

For t h e silica problem, we adopted Si(OH)4 = I; but this is allowable only for neutrally charged species in relatively dilute solutions up t o say 1 molal. Individual ion activity coefficients a r e tricky t o calculate, but a r e usually approximated in dilute solutions up t o say 3 molal using t h e extended Debye-Hiickel equation. For t h e high-temperature calculations discussed here, the following approximate values a r e quite satisfactory

= 3.29 x 10- 3 and mH2C03* = 0.25;

Recalculating t h e discharge composition thro gh (21, a t 26i°C, mHco3= 2.67 x 10-Y and mH2C03* = 101.2 x

Log K f o r reaction (8)

is -7.84 a t 2 6 5 ' ~ and since log K = log m H c 0 3 - + log YHco3-

+ log a ~ +

- log mH2C03+ we may substitute values t o obtain pH. From these d a t a our pH e s t i m a t e is 6.1! (There a r e a number of hidden assumptions in this approach so t h a t i t may

Equation (5) becomes log K265 = 2 log m K + + 2 log y K + + 2 pH

(7)

Digression No. 3 Equation (5) contains a pH t e r m (pH = -log aH+). Because t h e liquid sample taken a t t h e s u r f a c e is greatly depleted in C 0 2 and H2S relative t o t h e reservoir fluid, t h e pH measured on t h e laboratory sample is really of no use t o us and we need t o calculate a new value, pHt, for t h e reservoir fluid a t 2 6 5 ' ~ . This is not a trivlal task and i t requires an iterative calculation best performed on a computer or programable calculator. The usual procedure involves using an ion or proton balance f o r all t h e pH-sensitive ions in solution. For more details, s e e Henley e t al. (1984). Through these methods, pH265 for t h e BR22 discharge is found t o be 6.1.

TEMPERATURE

OC

Figure 2.4. Experimental solubility data for quartz in water as a function of temperature.

b e quite erroneous for s o m e wells, particularly those discharging from a heavily e x p l o ~ t e d reservoir.) Giggenbach (1980) has developed a n approach t o mineral-fluid equilibria which avoids t h e need t o conserve A1 a s required in reaction (4) or itself. In t h i s approach t h e log ratio A1+'Ye / a p+H designated AIH, is used a s a principal variable. #i: makes f o r fun reading a s you can imagine but t h e method is extremely valuable!

-

,.

2.

3.

Experimental determination of f r e e energies for t h e r e a c t a n t s and products, particularly ions, is very difficult. Direct determination of t h e equilibrium constant is equally difficult due t o unfavorable kinetics. The reaction coefficients required in hydrolysis reactions also require e x t r e m e precision in t h e thermodynamic data. A t 2 6 5 O ~ ,sericite is b e t t e r described a s illite and t h e activity of the KA13Si301, (OH)2 component should b e determined, if possfile, by x-ray and this value added into equation 7. The pHt calculation is dependent upon precise thermodynamic d a t a for ion association reactions and reliable e s t i m a t e s of activity coefficients.

I

-

+QUARTZ

aNo;/aH+

5.

L o c a t e the BR22 fluid composition point in Figures 2.5 and 2.6 using t h e d a t a of Table 2.3 and t h e e s t i m a t e d reservoir fluid pHt.

The calcium aluminosilicates a r e a particular problem because they occur in solid solution with iron and because their thermodynamic d a t a a r e poorly known. For this reason Figure 2.5 has been constructed using field d a t a for t h e calcium aluminosilicate phase boundaries. Exercise

6.

Scan t h e introductory section and decide which is t h e dominant calcic phase in t h e alteration assemblage.

Digression No. 4. The mineral is, of course, calcite, which, due t o t h e high C 0 2 c o n t e n t of t h e fluid, fixes most of t h e To assess i t s solubility in t e r m s of available Ca". aCa++/a2H+ you will need a n appropriate log K d e r ~ v e dfrom Appendix Table 2.A1.

1

250" C

ALBITE

LOG

I

4. Is t h e reservoir fluid close t o equilibrium with respect t o Kmica and Kspar a t q u a r t z saturation o r is t h e pHt t o o high or low?

Other hydrolysis reactions may b e assessed similarly and t h e results a r e usually expressed on activity r a t i o diagrams such a s those shown in Figures 2.5 and 2.6. Exercise

The log analytical quotient a K + / a . ,Orvalue, the reservoir fluid is -2.56 + 6.1 = 3.54. his compared t o t h e equilibrium quotient (3.94), suggests t h a t t h e reservoir fluid lies in t h e 'mica' stability field, 0.4 log units away from t h e mica-Kspar boundary. This is compatible with t h e petrography since sericitization is widespread, but a r e t h e r e s o m e d a t a checks t o b e made? 1.

Exercise

'-

MONT.

a+,

(L

-

a a.

LOG aR+

/

V)

a

2-

-I

KAOLlNlTE

K-MICA

W

.

L I

x

2

4

6

LOG aK/ + a+, F i g u r e 2.5. A c t i v i t y - a c t i v i t y r e l a t i o n s f o r the s y s t e m Na20-K20-A1203Si02-H20.

F i g u r e 2.6. A c t i v i t y - a c t i v i t y r e l a t i o n s f o r the s y s t e m Ca0-K20-A1203Si02-H20.

R. W. HENLEY & K. L.

31

BROWN

CaC03

Te Kopla(O.OOl)

+ 2H+ = ~ a + ++ H 2 c 0 3 *

1

The coefficients of t h e regression equations (Appendix Table 2.AI) m a y b e summed t o give

7-

I

loe" K = -235.56 + 10676.81/T

P

+ 89.084 log T - 0.0392 T

A t 2 6 5 ' ~ this equation gives log K = 6 46. The analytical quotient, ac ++mH2C03*/a2H', for t h e BR22 fluid is 6.1, suggests t h a t t h e fluid is close t o saturation with calcite. The difference between t h e analytical a n d equilibrium quotients is 0.36 log units, a small value compared with t h e sum + , if underground of possible e r r o r s on p ~ , ~ C a +and, boiling has occurred, mH2C03*.

whit%

\

5

.-g

eKawerau(O.03)

m 6.

~ ~ h i ~ ( 0 . 1 )

\

Cerro Prleto(0.25)

1

200

D i g r e s s i o n No. 5.

300 Temperature "C

Control of fluid pH through silicate reactions has a n important implication. Taking t h e Kspar-Kmica reaction, f o r example

Relationship between fluid pH and Figure 2.7. salinity for fluids in equilibrium with Kspar-Kmica-quartz.

i t is self-evident t h a t m K + is a function of pH. Potassium and sodium a r e t h e predominant cations in these high-temperature solutions and their r a t i o is limited by t h e reaction of plagioclase t o Kspar and Kmica, and, below about 2 7 0 ' ~ t~h e formation of secondary albite. T h e l a t t e r , t h e coprecipitation of albite and Kspar, affords a limiting condition for t h e ratio of sodium and potassium, which is used a s a geothermometer. Combining t h e two reaction constants, a t e m p e r a t u r e dependent function of pH and m ( ~ K) is found, and since (mNa + m ) mcl-, is related t o salinity. Flgure shows thls function together with some d a t a points for other a c t i v e geothermal systems.

Another important group of constituents (the gases ammonia, methane, nitrogen, hydrogen, etc.) a r e included in Tables 2.2 and 2.3. These d a t a provide a g r e a t deal of information about sources of various components, processes in t h e aquifer, etc.--see, for example, Giggenbach (1977, 1980). Although in t h e laboratory gas + gas reactions a r e normally slow, they appear t o b e a s f a s t a s silica precipitation in t h e fluid + mineral system of a geothermal aquifer. This provides a useful independent check on aquifer t e m p e r a t u r e s through reactions like

PA

&'

From a n assessment of fluid compositions from a number of wells, Giggenbach (1981) showed t h a t t h e Kspar-Kmica reaction was important in t h e upper two-phase (i.e., boiling ) zone of geothermal systems, but t h a t in general another equilibrium prevailed plagioclase

+ H 2 c 0 3 * = 'clay' + calcite

(9)

D a t a from a number of New Zealand fields a r e summarized in this context in Figure 2.8. Giggenbach (1980) has also shown t h a t i t may be more appropriate t o r e f e r t o geothermal mineralfluid interactions with reference t o a steady s t a t e rather t h a n t o equilibrium, but this does not a f f e c t t h e validity of t h e general relations discussed here.

although if t h e fluid boils e n r o u t e t o t h e well vapor loss o r gain may occur in t h e discharge relative t o t h e aquifer fluid. For this well, BR22, these e f f e c t s a r e minor and t h e analysis quotients for these reactions a r e close t o those expected for equilibrium a t 265'C. Reviewing d a t a from Broadlands and other fields, Giggenbach (1980) showed t h a t t h e ratio H2/H2S was a function of t e m p e r a t u r e (Fig. 2.9) reflecting equilibrium between pyrite and Fe-silicates represented by chlorite and/or epidote. Only near 300°C do t h e observed hydrogen contents approach those for t h e equilibrium between pyrite and pyrrhotite, or pyrrhotite + Fe-silicate.

o Woirokei o Kawerau Broodlands o other oreas

Figure 2.8. Boiling point vs depth (a) and fluid pressure P (b) curves for a series of mineral assemblages. The data points for Wairakei, Kawerau, Broadlands, El Tatio, and Iceland wells correspond to the point at which measured temperature/depth and temperature/pressure curves changed from (nearly) constant temperature with depth behaviour to that indicating boiling point vs depth conditions (reproduced with permission from Giggenbach, 1981).

FLUID-MINERAL EQUILIBRIA: TRACE-METAL CONTENTS Table 2.4a summarizes base- and precious-metal analysis from waters discharged a t atmospheric pressure from t w o Broadlands wells (Weissberg e t al., 1979). As noted earlier, base-metal sulfides a r e t o b e found in drill c o r e and cuttings from many of t h e wells a t Broadlands, but gold has been reported only in a complex antimony-arsenic-mercury-thallium-sulfide precipitate in a hot spring. F r e e gold (as e l e c t r u m )

Figure 2.9. The ratio H2/H2s as a function of temperature for selected mineral pairs in relation to data from active geothermal systems, plotted with respect to quartz and alkali geothermometer temperatures. (FeO) represents an iron-aluminium silicate (chlorite?, epidote?). CH4/Co2 contours are also shown. Note the differences between silica and alkali geothermometer temperatures (T is the NaKCa empirical geothermometer oYK$ournier and Truesdell, 1973). These differences reflect aquifer processes consequent on exploration and have high significance for monitoring geothermal field development (reproduced with permission from Giggenbach, 1980).

does occur in c o r e from nearby systems, e.g, Kawerau (B. Christenson, personal communication) and, s e e below, is t o b e anticipated in Broadlands core--but no one has yet looked for it. R e c e n t work a t BR22 has shown t h a t the concentrations of t r a c e metals discharged a t atmospheric conditions do not always represent t h e concentrations of those m e t a l s in t h e d e e p geothermal fluid (Brown, 1985). Considerable deposition occurs in t h e piping so t h a t corrected concentrations (Table 2.4b) a r e in f a c t much higher than would be judged from Table 2.4a.

Table 2.4a--Concentration of t r a c e m e t a l s i n w a t e r d i s c h a r g e d a t t h e s u r f a c e from e x p l o r a t i o n w e l l BR2 a t Broadlands

Table 2.4b--Minimum c o n c e n t r a t i o n s of t r a c e and b a s e m e t a l s i n w a t e r d i s c h a r g e d from BR22 a t Broadlands ( ~ r o w n , 1985). These d a t a a r e b a s e d on e m p i r i c a l d a t a f o r t h e m e t a l c o n t e n t s of t h e t o t a l d i s c h a r g e , b u t f o r e a s e of comparison w i t h Table 2.4a, t h e s e d a t a a r e c o r r e c t e d t o 1 b.a. s e p a r a t i o n p r e s s u r e , a l t h o u g h a s d i s c u s s e d below, a l a r g e p e r c e n t a g e o f t h e c o p p e r , s i l v e r and gold i s p r e c i p i t a t e d p r i o r t o discharge a t atmospheric pressure. Separation Pressure

Exercise 7.

Fe

Cu

Pb

Calculate t h e aquifer concentrations of these m e t a l s a t BR22. What percentage of t h e gold and silver i s precipitated by boiling during discharge to the surface? Water from BR22 flashed t o 1 bar a. contains similar Au and Ag t o those shown in Table 2.4a f o r BR2.

Zn

Ag

Au

t h a t in t h e s e low salinity fluids, s o m e o t h e r complex species m u s t be present; however, t h e r e a r e a s y e t no experimental d a t a available f o r possible species like P ~ ( o H ) + , ~ b H C 0 3 , etc., to c a l c u l a t e which predominates under t h e s e conditions. Digression No. 6

The concentration of. these m e t a l s m a y b e controlled by mineral deposition o r dissolution in t h e aquifer, so f i r s t w e must examine, a s with c a l c i t e and t h e various silicates, whether t h e fluids a r e close t o saturation with respect t o common sulfides, or t h e native m e t a l (for gold). Lead -

Figure 2.11 shows t h e calculated solubility of galena a s a function of t e m p e r a t u r e f o r a 1.0 m chloride solution (Q6 wt% NaCI). What is t h e increase in solubility t o b e e x p e c t e d a t 2 6 5 ' ~ if species like PbHCO; a r e present? First w r i t e a simple reaction f o r t h e formation of P~HCO;. PbS + 2H+

Figure 2.10 shows t h e distribution of t h e d i f f e r e n t lead chloride species a s a function of m based on r e c e n t experimental d a t a (Seward, 19&( F r o m t h e s e d a t a w e s e e t h a t PbCl constitutes about 50% of t h e lead in solution a s a c t l o r i d e complex a t 2 6 5 ' ~ in a low salinity fluid. The solubility of galena m a y then be w r i t t e n PbS

+ 2Hf + X I - = PbC12 + H2Saq

Inserting analytical d a t a m p ,i 0.015 ug/kg (as s o u b l l ~ t y of galena is, species, 0.03 ug/kg. This 1-10 kg actually found

(10)

for t h e reservoir fluid yields PbC12), s o t h a t t h e a c t u a l including a l l t h e chloride value is much lower t h a n t h e in these w a t e r s and indicates

+ HCO? = P ~ H C O ; + H2S

(11)

Increasing chloride relative t o t h e Broadlands fluid does not a f f e c t t h e reaction a s w r i t t e n e x c e p t through changes in activity coefficients, s o t h a t t h e observed contribution due t o this o r o t h e r species m a y be plotted directly o n t o Figure 2.10. It is c l e a r t h a t t h e chloride complexes dominate a t t h e s e higher salinities. Notice also t h a t in a hydrothermal system with fluid a t this salinity, t h e pHt of t h e fluid would b e about 1 t o 1.5 units less than f o r Broadlands (Fig. 2.7). The increased solubility due t o this e f f e c t , reaction (101, is, however, partly offset by t h e decrease of m H C O j Incidentally, using d a t a from Table 2.A1 and t h e observed lead c o n t e n t of t h e Broadlands fluid, a n e s t i m a t e of log K265 f o r this reaction is 7.58.

F i g u r e 2.10. T h e p e r c e n t a g e d i s t r i b u t i o n o f lead as c h l o r o l e a d ( I 1 ) complexes as a f u n c t i o n o f t o t a l chloride c o n c e n t r a t i o n , CIT, u p t o 3 0 0 ~a t ~ t h e s a t u r a t e d vapor pressure o f t h e s y s t e m ; the c u r v e f o r each c o m p l e x i s labelled a c c o r d i n g t o l i g a n d n u m b e r , e.g., c u r v e s l a b e l l e d '2' r e f e r to pbclZO (reproduced w i t h permission f r o m S e w a r d , 1984).

Gold

P i = X i s , where

s ti

In c o n t r a s t t o lead (and zinc), m e t a l s such a s gold, arsenic a n d antimony h a v e very stable bisulfide complexes. F o r gold t h e principle dissolution reaction (Seward, 1973) is

a n d a regression equation f o r t h e equilibrium constants of this reaction i s given in Appendix Table 2.A1.

Digression No. 7 Determination of t h e solubility of gold introduces a new solution p a r a m e t e r , t h e relative redox s t a t e , which is traditionally represented by t h e value of log fO2 f o r t h e solution, where f i s t h e fugacity. The choice of fO2 is unfortunate since its fugacity is immeasurably small, whereas fH2 can be m o r e easily directly measured. Since g a s concentrations a r e relatively low, i t is permissible t o substitute t h e partial pressure, Pi, f o r fugacity. T h e concentration of H 'n t h e reservoir fluid from Table 2.3 is 1.07 x 1 6 ' molesikg. In order t o calculate fy2,. a knowledge of Henry's Law is required. Thls IS given by

9

= Henry's Constant for i

Xi = m o l e f r a c t i o n of i P i = P a r t i a l P r e s s u r e of i The m o l e f r a c t i o n o f hydrogen i n t h e r e s e r v o i r fluid is

Henry's constant f o r H2 a t 2 6 5 ' ~ is 26,000 bars/mole-fraction, and t h e r e f o r e f H 2 = 0.182

X

X

26000 = 0.0473

A t 2 6 5 ' ~ log K = -19.46 f o r t h e reaction H20(1) =

+

1/202(g)

Therefore, log PO2 = -36.27. L o c a t e t h e BR22 reservoir fluid on t h e f02/pH diagram provided in Figure 2.12. A t t h e t e m p e r a t u r e s under discussion for Broadlands (E 3000

Q u a r t z Solubility a t High Temperatures I t was previously noted t h a t below about 300°c, pressure has a moderate a f f e c t and added salt has little a f f e c t upon t h e solubility of quartz. Above 300°c, both pressure and added s a l t a r e very important. Calculated solubilities of quartz in pure water over a wide range of t e m p e r a t u r e s and pressures, using t h e equation of Fournier and P o t t e r (1982a), a r e shown in Figure 3.5. In t h a t figure t h e r e is a solubility maximum (reported by Kennedy, 1950) t h a t extends from about 3 4 0 ' ~ a t t h e vapor pressure of solution t o 520°C close t o 900 bars. The shaded a r e a in Figure 3.5 shows a region of retrograde solubility in pressure-temperature space. Where cold water is h e a t e d a t constant pressure less t h a n about 900 bars, i t will dissolve more and more silica until either t h e solution s t a r t s t o boil (at pressures below about 165 bars) or t h e solubility maximum is reached. With further heating t h a t water will precipitate quartz. The precipitation of quartz in d e e p parts of a hydrothermal system may decrease t h e permeability t o such a n e x t e n t t h a t little convecting m e t e o r i c w a t e r can a t t a i n temperatures much greater than those shown by t h e quartz solubility maximum in Figure 3.5 (Fournier, 1977, 1983a, 1983b). However, computer modeling by L. A. Keith and P. T. Delaney (written communication, 1984) shows t h a t a completely impermeable seal is not likely t o result in realistic times solely by deposition of quartz (or a mixture of q u a r t z and other minerals) from a solution t h a t is h e a t e d a s i t flows toward a h e a t source. This is because t h e process is self-limiting: a s permeability is decreased by quartz deposition, the r a t e of flow decreases, which, in turn, decreases t h e r a t e of transport of silica t o t h e place where deposition c a n occur. However, other factors also may contribute t o t h e a t t a i n m e n t of a completely impermaeble seal, such a s quasi-plastic flow of rock t h a t takes place a t increasingly rapid r a t e s a s temperature increases. T e m p e r a t u r e profiles calculated from heat-flow d a t a for several localities in t h e western United States, and earthquake focal depths a t those s a m e locations, show t h a t t h e temperature a t which deformation changes from frictional (brittle fracture) t o quasi-plastic flow ranges from about 300' t o 4 5 0 ' ~ (Smith and Bruhn, 1984). This overlaps t h e 350' t o 5 0 0 ' ~ t e m p e r a t u r e range in which self sealing by precipitation of q u a r t z a n d o t h e r minerals is likely t o occur when solutions a r e h e a t e d a t constant pressure (Fournier, 1977, 1983a, 1983b; Sleep, 1983). Therefore, because of t h e above mentioned permeability reduction processes, t h e t i m e interval over which meteoric water a t hydrostatic pressure may i n t e r a c t directly with a shallow intruded

Temperature

,

C

F i g u r e 3.5. C a l c u l a t e d s o l u b i l i t i e s o f q u a r t z i n w a t e r u p t o 9 0 0 ~a t~ the i n d i c a t e d p r e s s u r e s . The s h a d e d area e m p h a s i z e s a r e g i o n of retrograde solubility.

body of magma (or still very hot rock) may be limited t o t h e early s t a g e of development of t h e hydrothermal system, or episodically t h e r e a f t e r with creation of new f r a c t u r e s by t e c t o n i c activity or thermal o r hydraulic cracking (Secor, 1965; Phillips, 1973; Henley and McNabb, 1978). This model has some interesting implications from t h e point of view of o r e genesis. With t h e circulation of meteoric water through shallow, intrusive rocks c u t off a t a n early s t a g e in t h e cooling process, those intrusive rocks will cool a t a slower r a t e , and hydrothermal activity will continue for a longer time. This is because h e a t must be transferred by conduction from t h e remaining very hot rock t o t h e cooler convecting hot water. Conductive transfer of thermal energy is much less efficient than convective transfer. Hydrothermal explosion activity is another possible consequence of t h e deposition of a n impermeable q u a r t z seal (Henley and McNabb, 1978). Large and s t e e p t e m p e r a t u r e and pore-pressure gradients a r e likely t o evolve where a n impermeable zone becomes established about a h e a t source. Even though convective flow of m e t e o r i c w a t e r is c u t off from t h e outside, t h e pore spaces within t h e zone between t h e quartz-sealed barrier and t h e remaining very hot rock a r e likely t o contain w a t e r or brine. This fluid may b e all or part m e t e o r i c o r connate water, l e f t over from before t h e silica sealing became complete. However, some or a l l of t h a t fluid could be volatiles evolved from a crystallizing magma (Burnham, 1967, 1979). If volatiles continue t o be evolved from a crystallizing magma, i t is easy t o envision a situation in which t h e fluid pressure on t h e high-temperature side of t h e q u a r t z seal becomes very

large (Phillips, 1973); sufficiently large t o cause formation of a b r e c c i a pipe o r even a conduit f o r a volcanic eruption (Morey, 1922). Hydraulic f r a c t u r i n g will occur when t h e porefluid pressure e x c e e d s t h e confining pressure (the l e a s t principle stress) by a n a m o u n t equal t o t h e tensile s t r e n g t h of t h e rock. T h e confining pressure m a y range from less t h a n normal hydrostatic t o lithostatic, depending on whether open fissures a r e present, t h e n a t u r e of t h e fluid in those fissures, and permeability relations. Propogation of e i t h e r a hydraulic o r t e c t o n i c f r a c t u r e through i m p e r m e a b l e rock from a region of high fluid pressure into a region of lower fluid pressure m a y c a u s e a significant decompression of t h e high-pressure fluid. If t h e t h e r m a l energy in t h e decompressing liquid a n d surrounding rock is large, massive flashing of w a t e r t o s t e a m may result. The expanding s t e a m may explosively propel rock f r a g m e n t s i n t o t h e alr where flashing occurs a t relatively shallow levels, a n d i n t o cavities and open fissures a t deeper levels. Even without a m a g m a t i c contribution t o t h e trapped fluids, pore pressures of those fluids could increase sufficiently t o rupture t h e enclosing rock a s a result of conductive heating. Hydrothermal explosive activity may be a n i m p o r t a n t contributing f a c t o r t o o r e deposition f o r various reasons, both physical and chemical. Brecciation greatly increases t h e permeability, providing easy a c c e s s f o r l a t e r hydrothermal fluids t h a t m a y deposit ores. When a hydrothermal explosion occurs, a lot of w a t e r is converted t o steam. Other volatile constituents, such a s H2S and C O , initially dissolved in t h e liquid phase, a r e partitioned into t h a t steam. This partitioning of volatiles, in turn, m a y i n c r e a s e t h e pH of t h e residual liquid. A t t h e s a m e t i m e t h e concentrations of t h e non-volatile constituents remaining in t h e liquid phase increase a s a result of t h e separation of s t e a m , while t h e solubilities of minerals generally d e c r e a s e a s a result of t h e d e c r e a s e in pressure. In Figure 3.5, n o t e t h e large decrease in q u a r t z solubility with decreasing In pressure a t t e m p e r a t u r e s a b o v e about 3 4 0 ' ~ . addition, where boiling occurs, t h e t e m p e r a t u r e of t h e system should d e c r e a s e because t h e r m a l energy is required t o convert liquid w a t e r t o steam. The above f a c t o r s generally favor deposition of silica, sulfides, and noble metals. Whether o r not o r e is deposited will depend in p a r t on t h e m e t a l and sulfur content of t h e initial fluid. However, when initial t e m p e r a t u r e s a r e above about 340°c, q u a r t z and o t h e r minerals should precipitate when and where t h e r e is a sudden drop in pore pressure. Also, any K-feldspar t h a t precipitates along with t h e q u a r t z a s a result of a sudden drop in pressure is likely t o be more potassium-rich t h a n t h a t which was in equilibrium with t h e fluid prior t o t h e drop in pressure (Fournier, 1976). A t Wairakei and Broadlands, quartz, adularia, a n d generally pyrite a r e t h e phases observed in t h e hydrothermal breccias t h a t ~~ and Browne, formed a t about 2 0 0 ~ - 3 0 0(Grindley 1976). Where exceptionally high degrees of silica supersaturation occur, particularly at lower temperatures, amorphous silica m a y precipitate and t h e n a l t e r t o q u a r t z o r chalcedonic silica. Many of t h e conclusions in t h e above discussion w e r e based on t h e solubility behavior of quartz in pure

water. The e f f e c t s of added salts c a n be modeled using NaCl solutions. C a l c u l a t e d solubilities of q u a r t z in aqueous NaCl, using t h e method of Fournier (1983b), show t h a t adding dissolved s a l t s should change t h e positions of t h e q u a r t z solubility maxima and t h e e x t e n t of t h e field of r e t r o g r a d e q u a r t z solubility t h a t a r e shown in Figure 3.5. However, t h e conclusion t h a t q u a r t z deposition should contribute significantly t o t h e formation of a n impermeable barrier t h a t prevents fluids a t hydrostatic pressure f r o m interacting directly with very hot rock o r m a g m a , is not changed by adding s a l t t o t h e system. Figure 3.6 shows t h e e f f e c t of 5 and 18 weight percent NaCl upon q u a r t z solubility a t 500 bars pressure and high temperatures. Added NaCl greatly increases t h e solubility of q u a r t z a t t e m p e r a t u r e s above about 300°c, and shifts t h e solubility maximum toward higher temperatures. A t very high concentrations of salt, t h e vapor-pressure curve m a y b e i n t e r s e c t e d before a solubility maximum is a t t a i n e d , such a s a t point A in Figure 3.6. In t h a t event, t h e solution will boil if t h e t e m p e r a t u r e is increased f u r t h e r without increasing t h e pressure. Wherever a solution boils, t h e concentration of dissolved silica in t h e residual solution should increase a t a more rapid r a t e t h a n c a n b e accommodated by t h e increasing salinity. T h e a m o u n t of silica t h a t can dissolve in t h e newly f o r m e d gas o r s t e a m phase is generally insufficient t o o f f s e t t h e supersaturation generated in t h e residual brine, and q u a r t z should precipitate because t h e t e m p e r a t u r e is high. Thus, q u a r t z veining should occur either when a dilute solution is h e a t e d t o a t e m p e r a t u r e above t h e q u a r t z

Quartz solubility i n NaCl solutions

500 bars

/'

/-. A

/

/

/

/

.

/

8

/

240

I

I

I

300

400

500

Temperature,

OC

Figure 3.6. Comparison of calculated quartz, solubilities (Fournier, 198313) in water and 5 and 18 weight percent aqueous NaCl at 500 bars and the indicated temperatures.

solubility maximum o r when a saline solution exceeds t h e t e m p e r a t u r e of t h e vapor-pressure c u r v e a t a given pressure. T e m p e r a t u r e s and approximate depths a t which boiling will occur in w a t e r a n d 5 and 18 weight p e r c e n t NaCl solutions a r e shown in Figure 3.7. Hydrostatic conditions a r e assumed in Figure 3.7, with pressure fixed by a n overlying liquid with a n a v e r a g e density of I throughout a vertical column up t o t h e ground surface. R e l a t i v e t o t h e depth scale, lower assumed a v e r a g e densities will move t h e boiling point curves down, a n d higher assumed densities will move t h e curves up. The approximate positions of t h e q u a r t z solubility maxima f o r w a t e r a n d 5 weight percent NaCl also a r e shown in Figure 3.7. Because t h e initial permeability several kilometers d e e p i n a hydrothermal system is likely t o be limited t o a f e w widely spaced f r a c t u r e s o r f r a c t u r e d zones of rock, a n irnpermeable zone resulting in g r e a t part from q u a r t z deposition in those f e w f r a c t u r e s m a y go unrecognized a s a significant feature. Also, in fossil hydrothermal systems where e s t i m a t e d t e m p e r a t u r e s a t t h e t i m e of vein formation a r e g r e a t e r t h a n 340°c, i t m a y be difficult t o d e t e r m i n e whether a given q u a r t z vein deposited a s a result of increasing o r decreasing temperature. If t h e r e is other hydrothermal alteration associated with t h e q u a r t z deposition, t h a t a l t e r a t i o n m a y give a n indication of t h e t h e r m a l history: albite is likely t o form in veins and a f t e r K-feldspar where a solution is heating and K-feldspar or ~ n u s c o v i t e would be

Figure 3.7. Temperature vs. depth (pressure) diagrams showing boiling point curves for pure water (curve A), 5 weight percent aqueous NaCl (curve B), and 18 weight percent aqueous NaCl (curve C). Also shown are the positions of quartz solubility maxima in water (curve D) and 5 weight percent aqueous NaCl (curve E).

deposited in veins and a f t e r plagioclase where a solution is cooling (Hemley e t al., 1971). CONCLUSIONS In well-established hydrothermal systems, where w a t e r remains in c o n t a c t with t h e surrounding rock a t a given high ternperature f o r m o r e than a f e w days or weeks, q u a r t z controls aqueous silica (Rimstidt and Barnes, 1980). Slow cooling of a hydrothermal solution generally will result in t h e deposition of q u a r t z if initial t e m p e r a t u r e s a r e between about 200' and 340°C. Rapid cooling allows supersaturated silica solutions t o form, particularly when t h e cooling is predominantly t h e result of decompressional boiling. Supersaturated silica solutions also m a y evolve where hot w a t e r dissolves glass-bearing rocks, and where rocks a r e a l t e r e d by very acid solutions. High alkalinity (high pH) i s generally not important in m o s t natural hydrothermal systems, but might be a f a c t o r in a few places f o r short periods of time. Slight silica supersaturation in respect t o q u a r t z is required f o r chalcedony t o precipitate directly from solution. Chalcedony a p p e a r s t o form and persist only a t t e m p e r a t u r e s below about 180°C. Large degrees of silica supersaturation a r e required f o r amorphous silica t o precipitate. Voluminous deposits of siliceous sinter generally indicate deposition f r o m neutral t o slightly alkaline (by loss of C 0 2 ) , chloride-rich w a t e r s t h a t flowed quickly t o t h e s u r f a c e from a reservoir with a Waters t e m p e r a t u r e in t h e range 200' t o 270°C. flowing from reservoirs with lower t e m p e r a t u r e s contain t o o l i t t l e silica t o form thick, hard, sinter deposits. Waters flowing from reservoirs with t e m p e r a t u r e s above 2 7 0 ' ~ contain s o much dissolved silica t h a t significant amounts precipitate in t h e channelways leading t o t h e surface, stopping hotspring activity b e f o r e large sinter deposits can form. Very l i t t l e silica precipitates from waters with pH values below about 3 t o 4. Amorphous silica i s relatively unstable and transforms t o poorly crystalline cristobalite, opal C T and chalcedony o r quartz. T h e t i m e required f o r t h e s e transformations depends upon t e m p e r a t u r e a n d t h e composition of fluid in c o n t a c t with t h e amorphous material. Morphologic f e a t u r e s show t h a t s o m e chalcedonies have formed a f t e r amorphous silica. Where such morphologic f e a t u r e s a r e absent, t h e origin of a given chalcedony is in doubt. I t is important t o know if chalcedony f o r m e d a s a primary precipitate o r by transformation of amorphous silica because of t h e implications about conditions required t o precipitate one o r t h e other. A t t e m p e r a t u r e s above about 300°c, increased pressure a n d added s a l t greatly increase t h e solubility of quartz. However, when a solution is h e a t e d a t constant pressure, eventually i t will either boil or e n t e r a field of r e t r o g r a d e q u a r t z solubility. In either event, q u a r t z should precipitate, decreasing t h e permeability in t h e h o t t e s t part of a convecting hydrothermal system. This precipitation of q u a r t z is likely t o occur a t 300' t o 550°c, depending on t h e d e p t h of circulation a n d t h e salinity of t h e system. An impermeable barrier m a y form by a combination of

q u a r t z deposition and quasi-plastic flow t h a t prevents subsequent d i r e c t c o n t a c t of circulating w a t e r a t hydrostatic pressure with very hot rock, o r magma. This, in turn, will i n c r e a s e t h e t i m e required t o cool a shallow, intrusive magma, a n d might lead t o very s t e e p t e m p e r a t u r e and pore-pressure gradients a c r o s s t h e impermeable barrier. If t h e impermeable barrier is then ruptured by seismic activity or increasing pore pressure in t h e confined, high-pressure side of t h e system, a hydrothermal explosion may occur. This may b e a mechanism by which s o m e breccia pipes form. T h e sudden d e c r e a s e in density of t h e pore fluids, formation of s t e a m , separation of gases, and decrease in temperature accompanying the hydrothermal explosion should cause silica a n d K-rich feldspar t o precipitate along with other minerals. ACKNOWLEDGMENTS Portions of this manuscript were reviewed by B. R. Berger, 3. W. Hedenquist, and D. E. White. I t has benefited greatly from their c o m m e n t s and suggestions. REFERENCES Allen, E. T., a n d Day, A. L., 1935, Hot springs of t h e Yellowstone National Park: Carnegie Institute of Washington Publication 466, 525 p. Arnorsson, S., 1975, Application of t h e silica geothermometer in low-temperature hydrothermal a r e a s in Iceland: American Journal of Science, v. 275, p. 763-784. Arnorsson, S., and Sigurdsson, S., 1974, The utility of w a t e r from t h e high-temperature a r e a in Iceland f o r s p a c e heating a s determined by their chemical composition: Iceland National Energy Authority, D e p a r t m e n t of National I-leat, Report OSJHD 7426, 34 p. Barton, P. B., and Toulmin, P., 111, 1961, Some mechanisms f o r cooling hydrothermal fluids: U.S. Geological Survey, Professional P a p e r 424-B, p. 348-352. Bodvarsson, G., 1964, Utilization of geothermal energy f o r heating purposes and combined s c h e m e s involving power generation, heating and/or byproducts: Geothermal Energy 11: United Nations Conference on New Energy, 1961, Proceedings, V. 3, p. 429-436. Briner, E., and Roth, P., 1948, Recherches sur I'hydrolyse par l a vapeur d'eau d e chlorures alcalins seuls ou additions d e divers adjuvants: Helvetica Chimica A c t a , v. 31, p. 1352-1360. Burnham, C. W., 1967, Hydrothermal fluids a t t h e m a g m a t i c stage; & Barnes, H. L. (ed.), Geochemistry of Hydrothermal O r e Deposits: Holt, Rinehart and Winston, New York, p. 34-76. Burnham, C. W., 1979, Magmas and hydrothermal fluids; & Barnes, H. L. (ed.), Geochemistry of Hydrothermal O r e Deposits, 2nd Edition: John Wiley and Sons, New York, p. 71-1 36. Calarnai, A., Cataldi, R., Dall'Aglio, M., and F e r r a r a , G. A., 1976, Preliminary report on t h e C e s a n o hot brine deposit (northern Latium, Italy):

Proceedings, 2nd United Nations Symposium of Development and Use of Geothermal Resources, V. 1, p. 305-313. C a r r , R. M., and Fyfe, W. S., 1958, Some observations on t h e crystallization of amorphous silica: American Mineralogist, v. 43, p. 908-916. Chen, C. H., 1970, Geology a n d geothermal power potential of t h e T a t u n volcanic region: Proceedings from t h e United Nations Symposiu~n on t h e Development and Utilization of Geothermal Resources: Geothermics, Special Issue 2, v. 2, pt. 2, p. 1134-1 143. Chen, C. H., 1975, Thermal w a t e r s in Taiwan, a preliminary study: Proceedings of the International Assocication of Hydrological Sciences, Grenoble, Publication 199, p. 79-88. Chen, C. T. A., and Marshall, W. L., 1982, Amorphous Behavior in pure w a t e r silica solubilities--1V. and aqueous sodium chloride, sodium sulfate, magnesium chloride, and magnesium s u l f a t e up t o 350'~: Geochimica e t Cosmochimica A c t a , V. 46, p. 279-287. Coplen, T. B., Combs, J., Elders, W. A., Rex, R. W., Burckhatter, G. C., and Laird, R., 1973, Preliminary findings of an investigation of t h e Dunes t h e r m a l anomaly, Imperial Valley, California: California D e p a r t m e n t of Water Resources, 48 p. C r e r a r , D. A., Axtmann, E. V., and Axtmann, R. C., 1981, Growth a n d ripening of silica polymers in aqueous solutions: Geochimica e t Cosmochimica A c t a , v. 45, p. 1259-1266. Dickson, F. W., and P o t t e r , J. M., 1982, Rock-brine chemical interactions: Final Report, E l e c t r i c Power Research Institute P r o j e c t 653-2, AP-2258, 8 9 p. Ellis, A. J., 1959, The solubility of c a l c i t e in carbon dioxide solutions: American Journal of Science, V. 257, p. 354-365. Ellis, A. J., 1963, The solubility of c a l c i t e in sodium chloride solutions a t high temperatures: American Journal of Science, v. 261, p. 259-267. Ellis, A. J., 1977, Chemical and isotopic techniques in geothermal investigations: Geothermics, v. 5, p. 3-12. Ellis, A. J., 1979, Explored geothermal systems; & Barnes, H. L. (ed.), Geochemistry of Hydrothermal O r e Deposits, 2nd Edition: John Wiley and Sons, New York, p. 632-683. Ellis, A. J., and Mahon, W. A. J., 1977, C h e m i s t r y and Geothermal Systems: Academic Press, New York, 392 p. Ernst, W. G., and Calvert, S. E., 1969, An experimental study of t h e recrystallization of porcelanite and i t s bearing on t h e origin of s o m e bedded cherts: American Journal of Science, v. 267, p. 114-133. F a c c a , G., and Tonani F., 1967, The self-sealing geothermal field: Bulletin of Volcanology, v. 30, p. 271-273. ~ l e m i n ~B. , A., and C r e r a r , D. A., 1982, Silicic and ionization a n d calculation of silica solubility a t e l e v a t e d t e m p e r a t u r e and pH--Application t o geothermal fluid processing a n d reinjection: Geothermics, v. I I, p. 15-29.

R. 0. FOURNIER Fournier, R. O., 1973, Silica in t h e r m a l waters: k Laboratory and Field investigations; Proceedings of International Symposium on Hydrogeochemistry and Biogeochemistry, J a p a n 1970, Volume I, Hydrogeochemistry: J. W. Clark (publisher), Washington, D.C., p. 122-1 39. Fournier, R. O., 1976, Exchange of Na' and K+ b e t w e e n water vapor and feldspar phases a t high temperature and low-vapor pressure: Geochimica e t Cosmochimica Acta, v. 40, p. 1553-1561. Fournier, R. O., 1977, Constraints on t h e circulation of m e t e o r i c water in hydrothermal systems imposed Geological by t h e solubility of quartz (abs.): Society of America, Abstracts with Programs, v. 9, p. 979. Fournier, R. O., 1979, Geochemical and hydrological considerations a n d t h e use of enthalpy-chloride diagrams in t h e prediction of underground conditions in hot-spring systems: Journal Volcanology Geothermal Research, v. 5, p. 1-16. 1981, Application of w a t e r Fournier, R. O., geochemistry t o geothermal exploration and reservoir engineering; in Rybach, L., a n d Muffler, L. J. P. (eds.), - ~ e o t h e r m a l Systems: Principles and C a s e Histories: John Wiley a n d Sons, New York, p. 109-143. Fournier, R. O., 1983a, Self-sealing a n d brecciation resulting from quartz deposition within hydrothermal systems: Extended abstracts, F o u r t h International Symposium on Water-Rock Interaction, Misasa, Japan, p. 137-140. Fournier, R. O., 1983b, A method of calculating q u a r t z solubilities in aqueous sodium chloride solutions: Geochimica et Cosmochimica A c t a , v. 47, p. 579-586. Fournier, R. O., 1985, Silica minerals a s indicators of conditions during gold deposition; & Tooker, E. W. (ed.), Geologic Characteristics of t h e Sediment- a n d Volcanic-hosted Types of Gold Deposits--Search f o r a n Occurrence Model: U.S. Geological Survey, Bulletin 1646, p. 15-26. Fournier, R. O., and Marshall, W. L., 1983, Calculation of amorphous silica solubilities a t 25' t o 300°C and apparent hydration numbers in aqueous s a l t solutions using t h e concept of e f f e c t i v e density of water: Geochimica e t Cosmochimica Acta, v. 47, p. 587-596. Fournier, R. O., and P o t t e r , R. W. 11, 1982a, An equation correlating t h e solubility of q u a r t z in w a t e r from 25' t o 9 0 0 ' ~ a t pressures up t o 10,000 bars: Geochimica et Cosmochimica A c t a , v. 46, p. 1969-1973. Fournier, R. O., and P o t t e r , R. W. 11, 1982b, A revised a n d expanded silica q u a r t z geothermometer: Bulletin Geothermal Resources Council, v. 11, no. 10, p. 3-12. Fournier, R. O., and Rowe, J. J., 1966, Estimation of underground t e m p e r a t u r e s from t h e silica c o n t e n t of water from hot springs and wet-steam wells: American Journal of Science, v. 264, p. 685-697. Fournier, R. O., and Rowe, J. J., 1977, The solubility of amorphous silica in w a t e r a t high t e m p e r a t u r e s and high pressures: American

57

Mineralogist, v. 62, p. 1052-1056. Fournier, R. O., and Truesdell, A. H., 1970, Chemical indicators of subsurface t e m p e r a t u r e applied t o hot-spring w a t e r s of Yellowstone National Park, Wyoming; & Proceedings from t h e United Nations Symposium on t h e Development and Utilization of Geothermal Resources: Geothermics (Special Issue 21, v. 2, P a r t I, p. 529-535. Frondel, C., 1962, The System of Mineralogy. Volume 111, Silica Minerals: John Wiley a n d Sons, New York, 334 p. Fyfe, W. S., and McKay, D. S., 1962, Hydroxyl ion catalysis of t h e crystallization of amorphous silica a t 3 3 0 ' ~ and s o m e observations on t h e hydrolysis of a l b i t e solutions: American Mineralogist, v. 47, p. 83-89. Galobardes, D. R., Van Hare, D. R., a n d Rogers, L. B., 1981, Solubility of sodium chloride in dry steam: Journal of C h e m i c a l Engineering Data, V. 26, p. 363-366. Gooch, F. A., and Whitfield, J. E., 1888, Analyses of w a t e r s of t h e Yellowstone National Park, with a n account of t h e methods of analysis employed: U.S. Geological Survey, Bulletin 47, 8 4 p. Grindley, G. W., and Browne, P. R. L., 1976, Structural and hydrologic factors controlling the permeabilities of s o m e hot-water geothermal fields; & Proceedings of t h e Second United Nations Symposium on t h e Development and Use of Geothermal Resources, San Francisco, 1975: v. I, p. 377-386, U.S. Government Printing Office. Harper, R. T., and Arevalo, E. M., 1982, A geoscientific evaluation of t h e Basley-Danin prospect, Negros Oriental, Philippines: Proceedings of the Pacific Geothermal Conference, 1982, 4th New Zealand Geothermal Workshop, pt. I , p. 235-240. Hawkins, R. B., 1982, Discovery of t h e Bell MineJ e r r i t t Canyon District, Elko County, Nevada: Mining Congress Journal, v. 68, p. 28-32. Hein, J. R., and Scholl, D. W., 1978, Diagenesis a n d distribution of l a t e Cenozoic volcanic sediments in t h e southern Bering Sea: Geological Society of America Bulletin, v. 89, p. 197-210. Heming, R. F., Hochstein, M. P., a n d McKenzie, W. F., 1982, S u r e t i m e a t geothermal system: An example of a volcanic geothermal system: Proceedings of the Pacific Geothermal Conference, 1982, 4th New Zealand Geothermal Workshop, pt. I, p. 247-250. Hemley, J. J., Montoya, J. W., Nigrini, A., a n d Vincent, H. A., 1971, Some a l t e r a t i o n reactions in t h e system Ca0-A1203-Si0 -H20: Society of Mining Geology Japan, specicaT Issue 2, p. 58-63. Henley, R. W., and Ellis, A. J., 1983, Geothermal systems ancient and modern: A geochemical review: Earth-Science Reviews, v. 19, p. 1-50. Henley, R. W., and McNabb, A., 1978, Magmatic vapor plumes and ground-water interaction in porphyry copper emplacement: Economic Geology, v. 73, p. 1-19. Heydemann, A., 1964, Untersuchungen uber die bildungsbeidugungen von quartz im

t e m p e r a t u r b e r e i c h zwischen 1 0 0 and ~ ~ 250'~: Beitrage zur Mineralogie und Petrographie, v. 10, p. 242-259. Holland, H. D., and Malinin, S. D. 1979, On t h e solubility a n d o c c u r r e n c e of non-ore minerals; & Barnes, H. L. (ed.), Geochemistry of Hydrothermal O r e Deposits, 2nd Edition: John Wiley and Sons, New York, p. 461-508. Iwasaki, I., and Ozawa, T., 1960, Genesis of s u l f a t e in acid hot springs: Bulletin Chemical Society of Japan, v. 33, p. 1018-1019. Jones, J. B., and Segnit, E. R., 1971, The n a t u r e of opal. I. Nomenclature a n d constituent phases: Journal of Geological Society of Australia, v. 18, p. 57-68. Kamiya, H., Ozaki, A., and Imahashi, M., 1974, Dissolution r a t e of powdered q u a r t z in acid solution: Geochemical Journal, v. 8, p. 21-26. Kano, K., and Taguchi, K., 1982, Experimental study on t h e ordering of opal-CT: Geochemical Journal, v. 16, p. 33-41. Kastner, M., Keene, J. B., and Gieskes, J. M., 1977, Diagenesis of siliceous oozes. I. Chemical controls on t h e r a t e of opal-A t o opal-CT transformation--an experimental study: Geochimica e t Cosmochimica Acta, v. 41, p. 1041-1059. Keenan, J. H., Keyes, F. G., Hill, P. G., and Moore, J. G., 1969, S t e a m Tables (international editionm e t r i c units): John Wiley and Sons, New York, 162 p. Keith, T. E. C., and Muffler, L. J. P., 1978, Minerals produced during cooling and hydrothermal a l t e r a t i o n of ash flow tuff from Yellowstone drill hole Y-5: Journal of Volcanology Geothermal Research, v. 3, p. 373-402. Kennedy, G. C., 1950, A portion of t h e system silicawater: Economic Geology, v. 45, p. 629-653. Lawless, J. V., a n d Gonzales,.R. C., 1982, Geothermal geology and review of exploration, Biliran Island: Proceedings of t h e P a c i f i c Geothermal Conference, 1982, 4th New Zealand Geothermal Workshop, pt. 1, p. 161-166. Leach, T. M., and Bogie, I., 1982, Overprinting of hydrothermal regimes in t h e Palimpinon Geothermal Field, Southern Negros, Philippines: Proceedings of the Pacific Geothermal Conference, 1982, 4th New Zealand Geothermal Workshop, pt. 1, p. 179-184. Lovering, T. G., 1972, Jasperoid in t h e United States. Its characteristics, origin, and economic significance: U.S. Geological Survey, Professional P a p e r 71 0, 164 p. Mahon, W. A. J., 1966, Silica in hot water discharged from drill holes a t Wairakei, New Zealand: New Zealand Journal of Science, v. 9, p. 135-144. Mahon, W. A. J., Klyen, L. E., a n d Rhode, M., 1980, Neutral sodium/bicarbonate/sulphate hot waters in geothermal systems: Chinetsu (Journal Japanese Geothermal Energy Association), v. 17, p. 11-24. Makrides, A. C., Turner, M., and Slaughter, J., 1980, Study of silica scaling from geothermal brines: Journal of Colloid and I n t e r f a c e Science, v. 73, p. 345-367.

Marshall, W. L., 1980, Amorphous silica solubilities. 11. Activity c o e f f i c i e n t relations and predictions of solubility behavior in s a l t solutions, 0-300'~: Geochimica e t Cosmochimica A c t a , v. 44, p. 925-931. Marshall, W. L., and Chen, C. T. A., 1982a, Amorphous silica solubilities. V. Predictions of solubility behavior in aqueous mixed e l e c t r o l y t e solutions t o 3 0 0 ' ~ : Geochimica e t Cosmochimica Acta, v. 46, p. 289-291. Marshall, W. L., and Chen, C. T. A., 1982b, Amorphous silica solubilities. VI. Postulated sulfate-silicic acid solution complex: Geochimica e t Cosmochimica A c t a , v. 46, p. 367-370. Marshall W. L., and Warakomski, J. M., 1980, 11. E f f e c t of Amorphous silica solubilities. aqueous s a l t solutions a t 2 5 ' ~ : Geochirnica e t Cosmochimica A c t a , v. 44, p. 915-924. Martynova, 0. I., and Samoilov, Yu. F., 1959, Dissolution of sodium chloride in a n atmosphere of water vapor of high parameters (Transactions): Zhurnal Neorganisheskoi Khimii, v. 11, no. 12, p. 2829-2833. Mizutani, S., 1970, Silica minerals in t h e early s t a g e of diagenesis: Sedimentology, v. 15, p. 419-436. Morey, G. W., 1922, T h e development of pressure in magmas a s a result of crystallization: Washington Academy of Science Journal, v. 12, p. 219-230. Morey, G. W., Fournier, R. O., Hemley, J. J., and Rowe, J. J., 1961, Field measurements of silica in w a t e r from hot springs and geysers in Yellowstone National Park; & Short Papers in t h e Geologic and Hydrologic Sciences: U.S. Geological Survey, Professional Paper 424-C, p. C333-336. Murata, K. J., and Larson, R. R., 1975, Diagenesis of Miocene siliceous shales, Temblor Range, California: U.S. Geological Survey, Journal of Research, v. 3, p. 553-566. Murata, K. J., and Nakata, J. K., 1974, Cristobalite s t a g e in t h e diagenesis of diatomaceous shales, Temblor Range, California: Science, v. 184, p. 567-568. Murata, K. J., and Randall, R. G., 1975, Silica mineralogy and s t r u c t u r e of t h e Monterey Shale, Temblor Range, California: U.S. Geological Survey, Journal of Research, v. 30, p. 567-572. Nakamura, H., Sumi, K., Katagiri, K., and Iwata, T., 1970, T h e geological environment of Matsukawa geothermal a r e a ; & Proceedings from t h e United Nations Symposium on t h e Development a n d Utilization of Geothermal Resources: Geothermics (Special Issue 2), v. 2, pt. I , p. 221-231. Oki, Y., and Hirano, T., 1970, The geothermal system a t t h e Hakone Volcano; & Proceedings from t h e United Nations Symposium on t h e Development and Utilization of Geothermal Resources: Geothermics (Special Issue 21, v. 2, pt. 2, p. 1157-1 156. Phillips, W. J., 1973, Mechanical e f f e c t s of retrograde boiling and i t s probable importance in t h e formation of s o m e porphyry o r e deposits: Institute Mining Metallurgy Transcripts, sec. B,

v. 82, p. B90-98. R a d t k e , A. S., Rye, R. O., and Dickson, F. W., 1980, Geology and s t a b l e isotope studies of t h e Carlin gold deposit, Nevada: Economic Geology, v. 75, p. 641-672. Rimstidt, J. D., and Barnes, H. L., 1980, The kinetics of silica-water reactions: Geochimica e t Cosmochimica A c t a , v. 44, p. 1683-1699. Rothbaum, H. P., Anderton, B. H., Harrison, R. F., Rohde, A. G., and S l a t t e r , A., 1979, E f f e c t of silica polymerization and pH on geothermal scaling: Geothermics, v. 8, p. 1-20. Rowe, J. J., Fournier, R. O., and Morey, G. W., 1967, The system water-sodium oxide-silicon dioxide a t Inorganic Chemistry, v. 6, 200, 250, a n d 300': p. 1183-1 188. Rowe, J. J., Fournier, R. O., and Morey, G. W., 1973, Chemical analysis of t h e r m a l waters in Yellowstone National Park, Wyoming, 1960-65: U.S. Geological Survey, Bulletin 1303, 31 p. Rye, R. O., 1985, A model f o r t h e formation of carbonate hosted disseminated gold deposits a s indicated by geologic, fluid inclusion, geochemical, a n d s t a b l e isotope studies of t h e Carlin and C o r t e z deposits, Nevada; & Tooker, E. W. (ed.), Geologic C h a r a c t e r i s t i c s of t h e Sediment- and Volcanic-hosted Types of Gold Deposits--Search f o r a n O c c u r r e n c e Model: U.S. Geological Survey, Bulletin 1646, p. 35-42. Saki, H., and Matsubaya, O., 1977, Stable isotopic studies of J a p a n e s e geothermal systems: Geothermics, v. 5, p. 97-124. Seastres, J . S., Jr., 1982, Subsurface geology of t h e Nasuji-Sogongon sector, Southern Negros Geothermal Field, Philippines: Proceedings of t h e Pacific Geothermal Conference, 1982, 4th New Zealand Geothermal Workshop, pt. 1, p. 173-178. Secor, D. T., Jr., 1965, Role of fluid pressure in jointing: American Journal of Science, v. 263, p. 633-646. Shettel, D. L., 1974, The solubility of q u a r t z in supercritical H 2 0 - C 0 2 fluids: Unpublished M.S. thesis, The P e n n s y l v a n ~ aS t a t e University, 52 p. Sleep, N. H., 1983, Hydrothermal convection a t ridge axes; & Rona, P. R., and Lowell, R. L. (eds.), Hydrothermal Processes a t Seafloor-Spreading Centers: Plenum Press, New York, p. 71-82. Smith, R. B., a n d Bruhn, R. L., 1984, I n t e r p l a t e extensional t e c t o n i c s of t h e e a s t e r n BasinRange: Inferences on s t r u c t u r a l style from seismic reflection data, regional tectonics, and

t h e r m a l mechanical models of brittle-ductile deformation: Journal of Geophysical Research, V. 89, p. 5733-5762. Sweeton, F. H., Mesmer, R. E., and Baes, C. F., Jr., 1974, Acidity measurements a t elevated temperatures. VII. Dissociation of water: Journal of Solution Chemistry, v. 3, p. 191-214. Toulmin, P., 111, and Clark, S. P., 1967, Thermal a s p e c t s of o r e formation; k Barnes, H. L. (ed.), Geochemistry of Hydrothermal O r e Deposits: Holt, R i n e h a r t a n d Winston, New York, p. 437-464. Truesdell, A. H., 1976, Summary of section 111 geochemical techniques in exploration; & Proceedings of t h e Second United Nations Symposium on t h e Development and Use of Geothermal Resources, San Francisco, 1975, v. 1: Washington, D.C., U.S. Government Printing Office, p. liii-lxxxix. Truesdell, A. H., a n d Fournier, R. O., 1977, Procedure f o r e s t i m a t i n g t h e t e m p e r a t u r e of a hot-water component in a mixed w a t e r by using a plot of dissolved silica versus enthalpy: U.S. Geological Survey, Journal of Research, v. 5, p. 49-52. Weres, O., Yee, A., and Tsao, L., 1982, Equations and t y p e curves f o r predicting t h e polymerization of amorphous silica in geothermal brines: Society of Petrological Engineering Journal (Feb. 19821, p. 9-16. White, D. E., 1965, Saline w a t e r s of sedimentary rock: American Association of Petroleum Geologists, Memoir 4, p. 342-366. White, D. E., Brannock, W. W., and Murata, K. J., 1956, Silica in hot-spring waters: Geochimica e t Cosmochimica A c t a , v. 10, p 27-59. White, D. E., Fournier, R. O., Muffler, L. J. P., and Truesdell, A. H., 1975, Physical results of research drilling in t h e r m a l a r e a s of Yellowstone National Park, Wyoming: U.S. Geological Survey, Professional P a p e r 892. White, D. E., Muffler, L. J. P., and Truesdell, A. H., 1971, Vapor-dominated hydrothermal systems compared t o hot-water systems: Economic Geology, v. 66, p. 75-97. Yamada, E., 1976, Geological development of t h e Onikobe caldera a n d i t s hydrothermal system; k Proceedings of t h e Second United Nations Symposium on t h e Development and Use of Geothermal Resources, San Francisco, 1975: v. I , p. 665-672, U.S. Government Printing Office.

APPENDIX

h r p h o u s silica Log S = [-731/(t+273.15)] +

Information For Use In Calculating Silica Solubilities Approximate solubilities of selected silica species in liquid water a t the vapor pressure of the solution can be calculated using equations (a)-(f) below (after Fournier, 1981; Fournier and Potter, 1982b). Concentrations of dissolved silica (S) are in rng/kg, t is temperature in degrees Celsius and the temperature range of application is 0' t o 250 b C, except as noted. Quartz

Log S = [-1309/(t+273.15)] + 5.19

Chalcedony

(a)

Log S = [-1032/(t+273.15)] + 4.69 (b)

4.52

(el

Q u a r t z (20'-330'~) t = -42.196

s2

+ 0.28831 S - 3.6685 x

+ 3.1665

s3 +

x

77.034 log S

(f

More precise solubilities of amorphous silica in the ternperature range 90' t o 3 4 0 ' ~ a t t h e vapor pressure of the solution and a t 1000 bars can be calculated using equations (g) and (h) (from Fournier and Marshall, 1983). T is temperature in Kelvin and m is t h e molality of dissolved silica. Vapor Pressure of solution

a -Cristobalite Log S = [-1000/(t+273.15)]

+ 4.78

(c)

log m = -6.1 16 + (0.01625 T) 41.758 x

+ (5.257 x

6 -Cristobalite Log S = [-781/(t+273.15)]

+ 4.51

(dl

T2)

T3)

1000 bars

Appendix Table 3.Al--Temperatures, enthalpies (Keenan et al., 1969), and quartz solubilities (Fournier and Potter, 1982b) in liquid and gaseous water (steam) at the vapor pressure of the solution

No.

T OC

H

Si02

J/g

mg/kg

No.

T OC

H J/g

Si02 %/kg

No.

T OC

H J/g

Si02 mg/kg

log m = -7.010 + (0.02285 T) 43.262 x

T2)

The solubility of quartz in water in t h e t e m p e r a t u r e range 25' t o 900°C a t specific volume (V) of t h e solvent ranging from about I t o 10 and from 300° t o 6 0 0 ' ~ at specific volume of t h e solvent ranging from about 10 t o 100 can be calculated using equations (i)-(I) (from Fournier and Potter, 1982a).

Appendix Table 3.A gives t e m p e r a t u r e , enthalpy, and quartz solubility d a t a t h a t a r e useful for calculating silica concentrations a f t e r mixing of w a t e r s with different initial t e m p e r a t u r e s and a f t e r boiling, a s discussed by Fournier and P o t t e r (1982b). The dissociation of silicic acid a s a function of pH can be determined a s follows, where square brackets denote activities of t h e indicated species, K I is t h e first dissociation constant, m is molality,Y is t h e activity coefficient, and T is t e m p e r a t u r e in Kelvin

where

The solubility of quartz in saline solutions c a n be calculated using equation (i) when (-log p F) is substituted f o r (log V), where P is t h e density of the solution and F is t h e weight fraction of water in t h a t solution (Fournier, 1983b). The molal solubility of amorphous silica in saline solutions (m,) a t temperatures ranging from 100° t o 340°c, and pressures ranging from t h e vapor pressure of t h e solution t o about 1000 bars can b e calculated using t h e following equations where ps is t h e density of t h e saline solution,pO is t h e density of pure w a t e r a t any given temperature and t h e indicated pressure (obtained from steam tables) and m0 is t h e molal solubility in pure water obtained either from equation (g) for t h e vapor pressure of t h e solution, or equation (h) for 1000 bars (Fournier and Marshall, 1983). logm, =

-

n l o g psF

+ n l o g pO(v.p.) + l o g m0 (v.p.1 n =

(m)

(n) [ l o g pO(lOOO b a r )

-

log P O ( V . ~ . ) ]

can b e calculated using t h e Debye Values of YHjSi0; Hiickel equation. When m H 4 s i 0 4 is determined using equation (i) and substituting -log p F f o r log V, t h e is unity. value of YHQSi Values o? K in t h e t e m p e r a t u r e range 1 t o 350°C can b e calcuiated using t h e following equation

+ .000133266 T~ + 267.6478 l o g T

(r)

For most natural thermal w a t e r s values of pH a r e too low for t h e second dissociation of silicic acid t o b e important, and

Chapter 4 CARBONATE TRANSPORT AND DEPOSITION IN THE EPITHERMAL ENVIRONMENT Robert 0. Fournier

INTRODUCTION The factors affecting the transport and deposition of carbonate in hydrothermal systems have been discussed in detail by Holland and Malinin (1979). Solubilities of carbonates a r e strongly temperature, and t h e influenced by pH, PCOIl presence of other disso ved salts. The alkali carbonates, Na, K, and Li, a r e relatively soluble a t all temperatures and generally precipitate only where t h e r e is e x t r e m e evaporation. In contrast, t h e alkaline e a r t h carbonates, C a , Mg, Sr, and Ba, a r e moderately t o sparingly soluble and commonly precipitate in hydrothermal systems. C a l c i t e is by f a r t h e most abundant and important carbonate found in t h e epithermal environment, and m o r e solubility d a t a a t hydrothermal conditions a r e available for i t t h a n for any of t h e other carbonates. Therefore, a f t e r briefly reviewing t h e system C02-water, t h e discussion will focus on t h e transport and deposition of calcite in hydrothermal solutions. The behaviors of other moderately t o sparingly soluble carbonates in hydrothermal solutions a r e similar t o t h a t of calcite.

given in Table 4.3. Adding s a l t t o t h e system C 0 2 H 2 0 increases t h e Henry's Law coefficient and decreases t h e solubility of carbon dioxide in t h e solution (Fig. 4.1). Salting-out coefficients for carbon dioxide in sodium chloride solutions generally a r e of t h e Steschenow t y p e

where k is t h e salting-out coefficient, m is t h e molality of NaCl, and KOH and KH a r e respectively t h e Henry's Law coefficients for pure water a s solvent and f o r t h e saline solution (Ellis and Golding, 1963). Approximate salting-out coefficients can b e obtained from equation (b) in Table 4.1, and a t selected t e m p e r a t u r e s from Table 4.3. Some dissolved carbon dioxide r e a c t s with water (hydrates) t o form carbonic acid

and s o m e of t h e carbonic acid, in turn, dissociates according t o t h e reactions H p 3 = HDj

C 0 2 DISSOLVED IN AQUEOUS SOLUTIONS

+ Hf

(4)

There is extensive l i t e r a t u r e on pressure-volumet e m p e r a t u r e measurements for t h e system C02-water, with and without additional dissolved salts (Bowers and Helgeson, 1983; and references therein). The experimental work t h a t is most applicable t o conditions appropriate for t h e formation of epithermal o r e deposits was carried o u t by Ellis and Golding (1963), who used solubility d a t a t o calculate Henry's Law coefficients, KH, f o r carbon dioxide in water and NaCl solutions (Fig. 4.1). According t o Henry's Law

where fCO2 is t h e fugacity of carbon dioxide and X is t h e mole fraction of carbon dioxide dissolved in t h e liquid phase. Because fugacity coefficients for carbon dioxide in dilute aqueous solutions a t temperatures below about 3 3 0 ' ~ a r e near unity (Ellis and Golding, 19631, fC02 in equation (I) c a n be replaced by t h e partial pressure of carbon dioxide, PEOZ, with l i t t l e error. An equation expressing Henry's aw coefficient for t h e system carbon dioxide-water a s a function of temperature is given in Table 4.1, equation (a). The constants used in t h e equations of Table 4.1 a r e given in Table 4.2. Values of t h e Henry's Law coefficient for t h e system C 0 2 - H 2 0 a t s e l e c t e d temperatures a r e

0 0

100

200

300

400

Temperature , OC

Figure 4.1. Values of Henry's Law constant, %, for the solution of carbon dioxide in water and sodium chloride solutions. (Redrawn from Ellis and Golding, 1963).

T a b l e 4.1--Equations expressing t h e temperature dependence of v a r i o u s e q u i l i b r i u m c o n s t a n t s and o t h e r c o e f f i c i e n t s t h a t a r e described i n the text. T i s t e m p e r a t u r e i n k e l v i n o r OK, and t i s t e m p e r a t u r e i n C e l s i u s o r O C . The respective constants f o r equations ( a ) through (k) a r e given i n Table 4.2. The d a t a used i n t h e d e r i v a t i o n of t h e s e e q u a t i o n s a r e shown i n T a b l e 4.3.

KH

=

a + bT + CT' + dT-I + e l o g T

(a)

k

=

a + bT + cT2 + d~~ + e l o g T

(b)

-

l o g K~ = a + b~ + CT-I + d l o g T

(c)

-

l o g K~ = a + b~ + c ~ '+ d ~ - l+ e l o g T

(dl

-

l o g K~ = a + bT + cT2 + d ~ - ' + e l o g T

(el

HL

gt-l

+ ht-'

+ i log t

l o g Kc = a + bT + cT2 + dT3 + eT-l A

= a + b~ +

log

+

= H2033

(8

(f)

(g) (h)

+ d ~ '

C T ~

+ e~~ + fT-'

-

+ i log t

= a + b t + c t 2 + d t 3 + e t4

+ ft5 +

-

032(gas)

= a + bt + ct2 + dt3 + et4

+ f t 5 + g t - l + ht-'

HG

where square brackets indicate activities of t h e enclosed species, and KI, and K 2 a r e respectively t h e first and second dissociat~onconstants of carbonic acid (Table 4.3). Equations (c) and (dl in Table 4.1 express t h e t e m p e r a t u r e dependence of K1 and K2. The reaction shown by equation (3) takes place relatively slowly, while t h e reaction shown by equation (4) is almost instantaneous (Kern, 1960). This information will b e of use later when t h e consequences of boiling a r e discussed. By tradition a distinction is not made between aqueous C 0 2 and H 2 C 0 3 , and t o t a l dissolved C 0 2 is A n e t reaction is generally reported a s H2C03. written

+ g log T

(i)

= a + b~ + c ~ '

+ dT3 + eT-l + fT-'

(k)

where KO is t h e equilibrium constant for t h e reaction shown by equation (8). Values of KO a t t e m p e r a t u r e s ranging from loo0 t o 3 0 0 ' ~ (Table 4.3) can be The calculated using equation (e) in Table 4.1. [H2C03] t e r m in equations (5) and (9) includes t h e activity of dissolved, nonhydrated C 0 2 . H 2 C 0 3 is less ionized a t high t e m p e r a t u r e s compared t o I w temperatures (value K1 range from about 10'g*57 a t OOC t o about 10-8.44at 3 0 0 ~ ~ ) . Therefore, a s a hydrothermal solution cools, bicarbonate dissociates (equation (4)), liberating hydrogen ions t h a t a t t a c k t h e minerals in t h e wall rock. Hydrolysis reactions involving feldspars generally buffer t h e pH a t near neutral t o slightly acidic conditions and cooling solutions become richer in cations (Na a t higher t e m p e r a t u r e s and C a a t lower temperatures) a s hydrogen ions a r e consumed by t h e formation of micas or clays. Where boiling occurs, generally a s a result of decreasing hydrostatic pressure e x e r t e d upon an ascending hydrothermal solution, C 0 2 is strongly partitioned into t h e gas (steam-rich) phase. The t o t a l gas pressure is equal t o t h e sum of t h e partial pressures of all t h e constituent gases, and these partial pressures a r e proportional t o t h e respective mole fractions. By Raoult's Law, for a mixture of t w o components t h a t exhibit ideal behavior (may b e closely approximated when t h e components a r e similar in rnolecular structure, or when o n e of t h e components is present in g r e a t excess)

Table 4.2--Coefficients

-

log K1

-

log K2

-

log KO

for use with equations listed in Table 4.1

HL

HG

-

log Kc

A B

-

log

AO

K~

k

-

log K 1

-

log K 2

-

log KO

H~ HG

-

log Kc A B

-

log

AO

where PYa and Pxb a r e t h e vapor pressures of pure components A and B respectively a t t h e given temperature, and Xa and X b a r e t h e respective mole fractions of A and B ~n t h e m ~ x t u r e . When boiling is f i r s t initiated, t h e r a t i o of C 0 2 t o w a t e r in t h e gas phase tends t o b e relatively l a r g e because most of t h e carbon dioxide initially dissolved

in t h e liquid exsolves quickly into t h e gas phase, while only a small amount of water changes t o steam. With continued boiling, t h e mole fraction of C 0 2 in t h e gas phase steadily decreases because l i t t l e a d d ~ t ~ o n C a l0 2 is available t o partition into t h e gas phase while t h e fraction of water t h a t is converted t o s t e a m increases a t a relatively constant rate. As t h e t e m p e r a t u r e of

T a b l e 4.3--Values

o f d i s s o c i a t i o n c o n s t a n t s , e n t h a l p i e s of l i q u i d w a t e r HL and s t e a m

..

HG, and Debye-Huckel C o e f f i c i e n t s , A and B , f o r t h e i n d i c a t e d t e m p e r a t u r e s . Enthalpy u n i t s a r e J/g. kg1/2mole-1crn-1 ~ ( O C ) 100

* (1) (2) (3)

125

U n i t s f o r A a r e kg1/2mole-1/2

and f o r B a r e

x l o 8 ( t h e product

x B cancels t h e lo8 f a c t o r )

150

225

175

200

250

275

300

Reference

Extrapolated. Henley e t a l . ( 1 9 8 4 ) . E l l i s and G o l d i n g ( 1 9 6 3 ) . Keenan e t a l . (1969).

t h e ascending gas-water mixture decreases, t h e volume of t h e gas phase i n c r e a s e s due t o t h e d e c r e a s e in hydrostatic load. The n e t e f f e c t is a d r a s t i c d e c r e a s e in t h e partial pressure of C 0 2 a s a boiling fluid ascends toward t h e earth's surface. Procedures f o r calculating t h e partitioning of relatively volatile constituents between coexisting liquids and gases, using hand-held, programmable calculators, a r e described in Henley e t al. (1984). F o r relatively dilute s y s t e m s a n d low initial dissolved gas concentrations, a distribution coefficient, B, is defined a s t h e concentration of gas in t h e vapor divided by t h e concentration of gas in t h e liquid. Ciggenbach (1980) derived t h e following equation t h a t expresses t h e t e m p e r a t u r e dependence of B f o r carbon dioxide in dilute aqueous solutions l o g B = 4.7593

-

.01092t

( 11 )

where t is t e m p e r a t u r e in d e r e e s Celsius. Equation (11) is valid f r o m 100' t o 340 C. Henley e t al. (1984) give t h e following equation for calculating t h e concentration of C 0 2 remaining in t h e liquid phase (CI) f o r t h e situation in which all t h e evolved gas remalns in c o n t a c t with a boiling fluid during a d i a b a t i c

8

decompression (single-step s t e a m separation)

where C o is t h e initial concentration of dissolved C 0 2 b e f o r e boiling, B is t h e distribution coefficient, a n d y is t h e f r a c t i o n of s e p a r a t e d steam. The corresponding equation t h a t gives t h e concentration of C 0 2 in t h e coexisting s t e a m ICv) is

Values of y a r e generally c a l c u l a t e d using enthalpy d a t a f o r pure boiling w a t e r and t h e relationship

Y =

Ho- HL

Hc - %

(14)

where Ho is t h e enthalpy of t h e initial liquid prior t o boiling, and HL and HG a r e t h e enthalpies of

coexisting liquid water and steam a f t e r boiling (Table 4.3). Enthalpies of liquid water and s t e a m a r e generally obtained from s t e a m tables (Keenan e t al., 1969) or they can be calculated using equations (f) and (g) in Table 4.1. Equation (131, however, yields values of y t h a t a r e slightly in error because t h e enthalpy of s t e a m containing C 0 2 is different from t h e enthalpy of pure steam. Other f a c t o r s also may cause t h e calculated concentration of C 0 2 in t h e liquid and steam fractions of a boiling solution t o b e in error. Assulnptions implicit in t h e use of equations (12) and (13) a r e t h a t dissolved C 0 2 does not become supersaturated in t h e liquid phase a s pressure is released, and t h a t little or no H C 0 3 converts t o H C O j a s t h e boiling solution cools. $he rapid transfer of most of t h e dissolved C 0 2 into t h e s t e a m fraction a t a n early s t a g e of boiling and t h e relatively slow conversion of dissolved C 0 2 t o H 2 C 0 3 (previously discussed) will tend t o limit t h e amount of H C O t h a t can form, but some non-equilibrium partitioning of C 0 2 between the liquid and gas phase is likely, particularly when t h e first boiling is initiated a t a ternperature below about 2 0 0 ' ~ . Another f a c t o r t h a t must b e considered is physical removal of t h e steam fraction from c o n t a c t with t h e residual liquid a s t h e boiling process proceeds. Compared t o single-step s t e a m separation, multistep and continuous s t e a m contact. in separation t h e lastHenley result liquidine and much t al. s t(1984) lower e a m fractions concentrations present methods t h a t ofa r C eand 0in2 equations for dealing with multistep and continuous steam separation. Figure 4.2 shows values of C,/Co for single-step s t e a m separation for a variety of inltlal and final temperatures, calculated using equations (12) and (14).

o

-1

-

-2

-

-3

.

0" 1

0

-a

&&

loo

I

Temperature 200 I

IOC

300 I

I

-

400

Figure 4.2. The C02 remaining in the residual liquid (Cl) after single-step steam separation at various temperatures to C02 ,in the initial liquid (Co) before boiling.

THE SOLUBILITY O F CALCITE IN AQUEOUS SOLUTIONS Ellis (1959, 1963) determined experimentally the solubility of c a l c i t e in aqueous solutions a t conditions appropriate for t h e formation of epithermal ore deposits. A t a given partial pressure of C 0 2 , t h e solubility of calcite decreases with increasing t e m p e r a t u r e (Fig. 4.3). Adding NaCl t o t h e solution increases t h e solubility of calcite (Fig. 4.4). A t any given t e m p e r a t u r e t h e solubility of calcite in solutions in equilibrium with a vapor phase increases with increasing C O pressure until mc-2 = 1 molelkg (Miller, 1952; ?egnit e t al., 1962). In solutions held a t a constant t o t a l pressure, t h e solubility increases with increasing C 0 2 concentration until rnc02 2 1 mole/kg and then decreases a t higher C 0 2 concentrations (Sharp and Kennedy, 1965; Malinin and Kanukov, 1971). The simplest equation representing t h e reaction by which c a l c i t e dissolves in aqueous solutions can be written

and t h e equilibrium constant (Kc) for reaction (15) is

Equation (16) is useful mainly for testing whether a solution of given cornposition is unsaturated, saturated, or supersaturated in respect t o calcite. Values of Kc in t h e t e m p e r a t u r e range 100' t o 300°C (Table 4.3) c a n b e calculated using equation (h), Table 4.1. Because very l i t t l e CO? is present in most natural hydrothermal solutions, t h e solubility of c a l c i t e is commonly expressed in t e r m s of reactions using equations (5), (71, involving H', HCO, and f (9), and (16). C02

100

150

200 Temperature.

250

300

OC

The s o l u b i l i t y o f calcite i n w a t e r F i g u r e 4.3. u p t o 300% a t v a r i o u s partial p r e s s u r e s o f c a r b o n d i o x i d e . (Redrawn f r o m E l l i s , 1959).

E .-.-

=

2

n Temperature,

In equations (16), (181, (20), and (221, t h e activity of c a l c i t e is unity if t h e r e is no significant substitution of other cations f o r calcium in solid solution, such a s Mg, Fe, or Mn. In order t o evaluate equations (16)-(221, activities of t h e indicated aqueous species must b e used. In dilute solutions, activities of dissolved constituents a r e about equal t o t h e corresponding molalities. In saline solutions, however, t h e molality of e a c h species i (mi) must be multiplied by i t s activity coefficient (Y.) t o obtain t h e activity (ai = Y im.). Activity coefficients for solutions with ionic strengths less than about 2 molal can be calculated using a n extended form of t h e Debye-Hiickel equation

where gi is t h e ionic charge, I t h e ionic strength, and A, B, ai and b a r e constants (Henley e t al., 1984). However, A and B vary with temperature. Their values from 100' t o 350°c, in 2 5 ' ~ increments, a r e given in Table 4.3. The coefficients A and B also can b e calculated t o t h r e e decimal places using equations (i) and (j), Table 4.1. The ionic strength is defined a s

F o r most hydrothermal waters I is approximately equal Values of Si and zi a r e t o t h e sum of m~ + and m listed in Table 4.C Up to%>O°C, b has values in t h e range 0.03 t o 0.05 when concentrations a r e up t o 3 molal (Helgeson, 1969). In natural hydrothermal solutions many dissolved constituents and a variety of chemical reactions involving solids, liquids and gases influence t h e

.

OC

F i g u r e 44. The s o l u b i l i t y o f calcite i n w a t e r and sodium c h l o r i d e s o l u t i o n s a t a carbon d i o x i d e p r e s s u r e o f 1 2 a t m o s p h e r e s (12.2 bars). (Redrawn f r o m E l l i s , 1963).

dissolution and deposition of calcite. The situation is particularly complex when boiling occurs. Computer programs can be used t o e v a l u a t e t h e s e complex reactions (Truesdell and Singers, 1971, 1974; Morel and Morgan, 1972; Kharaka and Barnes, 1973; Truesdell and Jones, 1974; Plummer e t al., 1975; Wolery, 1979; Reed, 1982; R e e d and Spycher, 1984). These programs, however, generally require main-frame computers for their execution. Arnorsson (1978) specifically calculated t h e amount and location of c a l c i t e deposition in geothermal wells in Iceland where natural thermal waters flash t o s t e a m during production of t h e resource. The results of his calculations, showing t h e d e g r e e of supersaturation with respect t o c a l c i t e t h a t occurs during single-step adiabatic flashing (boiling) in t h e wells, a r e shown in Figure 4.5 (supersaturated solutions t h a t plot below t h e thick solid line). Programs t h a t c a n b e used with programmable hand-held calculators, and t h a t a r e applicable t o calcite transport and deposition in natural waters, a r e given in Henley e t al. (1984). These programs a r e very useful even though a few simplifying assumptions a r e required, and only t h e most important dissolved species a r e included. The ensuing discussion follows t h e general procedures presented in Henley e t al. (1984). From a consideration of t h e cation-anion charge balance t h a t must b e maintained in a l l solutions (and neglecting e f f e c t s of oxidation-reduction reactions and pH-dependent ions t h a t a r e likely t o b e present only in very small amounts in most natural hydrothermal solutions, such a s CO;, H2SiOi, and OH-) a constant,A,

0

Table 4.4--Values of i o n i c c h a r g e , z , and i o n s i z e p a r a m e t e r , a , f o r t h e common i o n i c s p e c i e s i n geothermal f l u i d s (from Henley e t a l . , 1984) H+

~ a + HCO;

HS-

OH-

a-

CO;

H3Sio;

~ i +

H~BO;

K+

F-

N H ~

SO;

HSO;

~a++ M~++

can be defined t h a t is independent of t e m p e r a t u r e and equal t o t h e sum of t h e principal pH-dependent ions

The concentrations of t h e ionic species indicated in equation (25) can be calculated using t h e following relationships, where K 1 is t h e f i r s t dissociation constant for e a c h of t h e respective weak acids

"k"03

%2( =

total )

[H+l Y m j

(26)

150 2 0 0 250 150 200 250 Temperature O C T h e -computed a c t i v i t y p r o d u c t o f F i g u r e 4.5. ~ a + a+n d Cog i n g e o t h e r m a l w a t e r s d u r i n g s i n g l e - s t e p adiabatic f l a s h i n g i n r e l a t i o n t o the c a l c i t e s o l u b i l i t y c u r v e ( t h i c k solid l i n e ) . The s o l i d l i n e s assume maximum d e g a s s i n g a n d t h e d a s h e d l i n e s 1 / 5 o f maximum degassing. (From Arnorsson, 1978).

The reader is referred t o Henley e t al. (1984) for derivation of equations (261429). If t h e pH, ionic strength, and chemical composition (particularly t o t a l dissolved carbon, silica, boron, and ammonia) of a solution a r e known a t a given temperature, equations (25)-(29) can be used t o

e s t i m a t e t h e indicated species concentrations a t any other t e m p e r a t u r e up t o t h e limit of t h e available thermodynamic d a t a (now about 300' t o 350°C). Note, however, t h a t equations (25)-(29) do not t a k e account of changing concentrations and partitioning of constituents between t h e liquid and gas phase during boiling. For adiabatic boiling resulting from decompression, t h e value of 'changes a s a function of the fraction of s t e a m (y) t h a t forms.

'

(before boiling) =

-

Y) '(after

boiling) (30)

Equation (35) is expressed in t e r m s of mole ratios, so molal concentration units can b e substituted for t h e number of moles of t h e given species in t h e liquid fraction. Combining equations (51, (71, and (351, multiplying molalities by activity coefficients t o obtain activities where required, and rearranging gives

The e f f e c t of partitioning of C 0 2 between liquid and gas can b e accounted f o r by using t h e relationship

where n and nv a r e t h e number of moles of t h e indicated species in t h e liquid and gas phases respectively, and values of AQ a t various ternperatures with boiling water and b o ~ l i n g NaCl solutions a s solvents a r e given in Ellis and Golding (1963). A, f o r dilute solutions also can be calculated using equatlon (k) in Table 4.1. Rearranging equation (31) and substituting

Account can b e taken of t h e partitioning of other volatile constituents in a similar manner. The resulting equations can be used t o e s t i m a t e by

Hveragerdi Namafall For one-step s t e a m separation without carbonate precipitation (supersaturation allowed in the calculation) t h e t o t a l number of moles of C02-bearing species remains constant during the boiling process, even though C 0 2 partitions between t h e gas and liquid phases

p Nesjavellir

The concentration of CO; in equation (33) cannot be neglected when solutions boil because t h e pH may increase significantly (Fig. 4.6). Combining equations (32) and (33)

and dividing equation (34) by n H C O j

I

Reykjanes

Leira

Figure 4.6. The variation in pH in geothermal waters during one-step adiabatic flashing in relation to the calcite solubility curve (thick solld line). The solid lines assume maximum degassing and the dashed lines 1/5 of maximum degassing. (Redrawn from Arnorsson, 1978).

R. 0 . FOURNIER iteration t h e pH and distribution of pH-dependent species a t given t e m p e r a t u r e s for boiling conditions. The equations and procedures a r e cumbersome using a hand-held calculator, but a r e easily dealt with using table-top micro o r personal computers. Where cooling occurs adiabatically (generally too quickly f o r solution-mineral reactions involving silicates t o buffer pH) t h e pH of t h e residual liquid usually rises a s a result of partitioning of C O and other acid-forming gases into t h e steam phase. A g u r e 4.6 shows calculated changes in pH t h a t accompany single-step adiabatic flashing of natural thermal waters in geothermal wells in Iceland (Arnorsson, 1978). These a r e t h e s a m e well waters used t o illustrate t h e d e g r e e of supersaturation with respect t o calcite t h a t o c c u r s during single-step flashing (Fig. 4.5). SUMMARY In most natural waters heating will cause calcite and other carbonates t o precipitate, whereas cooling without boiling will cause them t o dissolve (Fig. 4.4). However, where a n ascending solution boils a s a result of decompression (cooling adiabatically), carbonate is likely t o precipitate a s a result of t h e boiling (Fig. 4.5). The cooling t h a t tends t o move a solution toward a condition of undersaturation with respect t o t h e various carbonate minerals is generally more than offset by t h e s t r o n g partitioning of C 0 2 into t h e vapor phase (and concomitant d e c r e a s e in partial pressure of C 0 2 ) t h a t decreases t h e solubility of carbonates. A t present, calculations t h a t t a k e account of all t h e physical processes (mainly boiling) and chemical reactions t h a t influence t h e transport and deposition of carbonate minerals can b e carried o u t only with t h e aid of large computers. However, if only t h e most abundant dissolved species in natural waters a r e considered, and simplifying assumptions a r e made about enthalpies of coexisting liquids and gases (mainly steam), small table-top cornputers and hand-held, programmable calculators can be used effectively t o calculate t h e approximate conditions f o r transport and deposition of c a l c i t e in hydrothermal solutions.

REFERENCES Arnorsson, S., 1978, Precipitation of calcite from flashed geothermal waters in Iceland: Contributions t o Mineralogy and Petrology, v. 66, p. 21-28. Bowers, T. S., and Helgeson, H. C., 1983, Calculation of the thermodynamic and geochemical consequences of nonideal mixing in t h e system H20-CO NaCl on phase relations in geologic s y s t e m s : 2 - ~ q u a t i o n of s t a t e f o r H20-C02-NaCI fluids a t high pressures and temperatures: Geochimica et Cosmochimica Acta, v. 47, p. 1247-1275. Ellis, A. J., 1959, The solubility of c a l c i t e in carbon dioxide solutions: American Journal of Science,

71

Ellis, A. J., 1963, The solubility of calcite in sodium high temperatures: chloride solutions at American Journal of Science, v. 261, p. 259-267. Ellis, A. J., a n d Golding, R. M., 1963, The solubility of ~ ~water and in carbon dioxide above 1 0 0 in sodium chloride solutions: American Journal of Science, v. 261, p. 47-60. Giggenbach, W. F., 1980, Geothermal gas equilibria: Geochimica e t Cosmochimica Acta, v. 44, p. 2021-2032. Henley, R. W., Truesdell, A. H., and Barton, P. B., Jr., 1984, Fluid-mineral equilibria in hydrothermal systems: Society of Economic Geologists, Reviews in Economic Geology, Volume 1, 267 p. Holland, H. D., and Malinin, S. D., 1979, On t h e solubility and occurrence of non-ore minerals; & Barnes, H. L. (ed.), Geochemistry of Hydrothermal O r e Deposits (2d ed.): John Wiley and Sons, New York, p. 461-508. Keenan, J. H., Keyes, F. G., Hill, P.G., and Moore, 3. G., 1969, S t e a m Tables (international editionm e t r i c units): John Wiley and Sons, New York, 162 p. Kern, D. M., 1960, The hydration of carbon dioxide: Journal of Chemistry Education, v. 37, p. 14-23. Kharaka, Y. K., and Barnes, I., 1973, SOLMNEQ: Solution-mineral equilibrium computations: National Technical Information System Technical R e p o r t PB 214-899, 82 p. Malinin, S. D., and Kanukov, A. B., 1971, The solubility of c a l c i t e in homogeneous H20-NaCI-C02 systems in the 2 0 0 ~ - 6 0 0 ~t ~e m p e r a t u r e interval: Geochemistry International, v. 9, p. 410-418. Miller, J. P., 1952, A portion of the system calcium carbonate-carbon dioxide-water, with geologic implications: American Journal of Science, v. 250, p. 161-203. Morel, F., and Morgan, J., 1972, A numerical method for computing equilibria in aqueous systems: Environmental Science and Technology, v. 6, p. 58-67. Plummer, L. N., Parkhurst, D. L., and Kosiur, D., 1975, MIX2: A computer program for modeling chemical reactions in natural waters: U.S. Geological Survey, Water-Resources Investigations, p. 75-61. Reed, M. H., 1982, Calculation of multicomponent chemical equilibria and reaction processes in systems involving minerals, gases and a n aqueous phase: Geochimica e t Cosmochimica Acta, v. 46, p. 513-528. Reed, M. H., and Spycher, N., 1984, Calculation of pH and mineral equilibria in hydrothermal waters with application t o geothermometry and studies of boiling and dilution: Geochimica e t Cosmochimica Acta, v. 48, p. 1479-1492. Segnit, E. R., Holland, H. D., and Biscardi, C. J., 1962, The solubility of c a l c i t e in aqueous solutions-I. The solubility of calcite in water between 75' and 200' a t COC? pressures up t o 60 atm: Geochimica e t osmochimica Acta, v. 26, p. 1301-1331.

Sharp, W. E., and Kennedy, G. C., 1965, The system C a 0 - C 0 2 - H 2 0 in t h e two-phase region c a l c i t e and aqueous solution: Journal of Geology, v. 73, p. 391-403. Truesdell, A. H., and Jones, B. F., 1974, WATEQ, a c o m p u t e r program f o r calculating chemical equilibria of natural waters: U.S. Geological Survey, Journal of Research, v. 2, p. 233-248. Truesdell, A. H., and Singers, W. A., 1971, C o m p u t e r calculation of downhole chemistry in geothermal

areas: New Zealand DSIR C h e m i s t r y Division R e p o r t CD2136, 145 p. Truesdell, A. H., and Singers, W. A., 1974, Calculation of aquifer chemistry in hot-water geothermal systems: U.S. Geological Survey, Journal of Research, v. 2, p; 271-278. Wolery, A. T., 1979, Calculation of chemical equilibrium between aqueous solutions a n d minerals; the EQ3/6 software package: UCRL-52658, L a w r e n c e Livermore Laboratory.

Chapter 5 FLUID-INCLUSION SYSTEMATICS IN EPITHERMAL SYSTEMS R. J. Bodnar, T. J. Reynolds, and C. A. Kuehn

INTRODUCTION Fluid-inclusion analyses have provided s o m e of the most useful information for determining t h e physical and chemical environments of mineral formation. The purpose of this c h a p t e r is t o describe those fluid-inclusion characteristics which s e r v e t o distinguish relatively near-surface, epithermal formation conditions from deeper and, potentially, higher temperature formation conditions, and t o discuss several techniques and problems which a r e specific t o fluid inclusions trapped in t h e epithermal A detailed summary and critique of environment. fluid-inclusion literature related t o epithermal systems has not been attempted. For this information t h e reader is referred t o t h e recent compilations of Buchanan (19811, Heald-Wetlaufer et al. (1983), Roedder (19841, and Hedenquist a n d Henley (19 85). Moreover, we have not a t t e m p t e d t o r e l a t e any particular fluid-inclusion c h a r a c t e r i s t i c t o a specific type o r s t a g e of mineralization, because a n adequate d a t a base t o do so does not presently exist. This presentation is limited t o t w o subjects--the petrography and petrology of fluid inclusions from t h e epithermal environment--and is intended t o provide t h e explorationist with a basic understanding of t h e criteria for recognizing and interpreting inclusions trapped in this environment. Two i m p o r t a n t topics will be discussed in detail: (1) t h e identification and interpretation of fluid inclusions trapped from boiling fluids, and (2) t h e identification of gases (mainly C 0 2 ) in fluid inclusions and t h e e f f e c t of volatiles on calculated pressures and depths of trapping. We will not, however, discuss t h e important chemical consequences of boiling and dissolved volatiles, a s these subjects a r e covered in detail in o t h e r c h a p t e r s in this volume (see Henley, 1985, this volume; Henley and Brown, 1985, this volume; Fournier, 1985, this volume; and Reed and Spycher, 1985, this volume). INFORMATION AVAILABLE FROM FLUID-INCLUSION PETROGRAPHY Characteristics of fluid inclusions trapped in the epithermal environment c o n t r a s t markedly with those of inclusions formed in deeper environments. This section considers those f e a t u r e s which a r e diagnostic of shallow crustal environments and which a r e readily observable by anyone with access t o a standard petrographic microscope. In addition, owing t o t h e n a t u r e of fluid inclusions trapped in t h e epithermal environment, particular c a r e in t h e selection of fluid inclusions for detailed microthermometric analysis

must be practiced, and such precautions will be discussed. Diagnostic information common t o all minerals containing fluid inclusions formed in the epithermal environment will b e presented first, followed by a detailed discussion of information available from petrographic observations of fluid inclusions in quartz. Fluid inclusions in t h e epitherrnal environment typically contain only t w o phases a t room temperature--a low-salinity H 2 0 liquid phase and a vapor bubble. Daughter minerals of halite and sylvite a r e notably absent in t h e epithermal environment. Readily observable evidence for gases is iacking also: gases rarely occur a s condensed phases in fluid inclusions, and evidence of gases is normally not found by sirnple crushing tests, f o r reasons described below. However, low concentrations of gases have been identified by capacitance manometer and mass spectrometric techniques (Sommer e t al., 1985; Hedenquist and Henley, 1985). Exceptions t o t h e s e generalities may occur when an epithermal system overprints a n earlier, higher t e m p e r a t u r e system, or vice versa. For example, a t Summitville, Colorado, which is a rather high-level, fossil hydrothermal system within a volcanic dome (Perkins and Nieman, 1982), s o m e early quartz rarely contains a f e w isolated healed microfractures defined by vapor-rich H 2 0 + C 0 2 (270 mole % C O ) inclusions and/or healed microfractures defined by hajite-bearing inclusions with small vapor bubbles. These inclusions presumably contain samples of early magmatic fluids trapped before t h e development of the near-surface epithermal system a t this locality. Similarly, highvapor-rich H20 + C02 t e m p e r a t u r e ( 360°C), (270 mole % C 0 2 ) inclusions found in some active geothermal systems (e.g., in c e r t a i n deep portions of t h e Geysers geothermal field) could be magmatic fluids. Also, t h e high concentrations of C 0 2 and CHq present in fluid inclusions from sediment-hosted gold deposits may not be temporally related t o t h e hydrothermal system attending gold mineralization. Certainly, some epithermal systems must have been subsequently buried and subjected t o fluids of deeper origins, and in t h e s e cases COZ-bearing and/or salts a t u r a t e d inclusions may postdate t h e epithermal inclusions. Fluid inclusions a r e trapped in many minerals formed in t h e epithermal environment; quartz, sphalerite, calcite, and fluorite have yielded useful t h e r m o m e t r i c data. Of these minerals, quartz usually provides t h e most f e r t i l e opportunities for collection of interpretable fluid-inclusion data. Just as the megascopic crustiform banding of quartz is

c h a r a c t e r i s t i c o f v o i d filling in t h e epithermal environment, s o a r e t h e microscopic t e x t u r e s in quartz diagnostic. F u r t h e r m o r e , the fluid-inclusion textures in q u a r t z vary s y s t e m a t i c a l l y in a manner t h a t permits general t h e r m a l c o n d i t i o n s t o b e predicted from f e a t u r e s observed u n d e r t h e microscope. These s y s t e m a t i c v a r i a t i o n s i n epithermal-quartz fluidinclusion t e x t u r e s m a y r e f l e c t t h e temperature dependence of t h e kinetics of dissolution and reprecipitation o f q u a r t z during "maturationn* of individual inclusions. T h u s , the following discussion of t h e microscopic t e x t u r e s diagnostic of t h e epithermal environment a r e l i m i t e d to those found in quartz.

*The t e r m "maturation" is used t o describe t h e process of dissolution a n d r e p r e c i p i t a t i o n of a host mineral surrounding a t r a p p e d fluid. The initial "immature" inclusion is generally l a r g e and very irregularly shaped. With t i m e , t h i s l a r g e inclusion will neck down t o form numerous s m a l l e r , m o r e regularly shaped inclusions, with t h e m o s t m a t u r e inclusion obtaining t h e negative c r y s t a l s h a p e of t h e host mineral. (See Roedder, 1984; figs. 2-15.) A t t h e o u t s e t , i t should be noted t h a t the microscopic t e x t u r e s displayed by q u a r t z discussed below a r e not unique t o t h e epithermal environment. Similar t e x t u r e s m a y o c c u r in q u a r t z from deeper, higher-temperature hydrothermal systems. To mitigate possible ambiguities in the final interpretation of t h e t e x t u r e s observed, many samples must b e collected a n d observed t o gain a broad perspective of a l l f e a t u r e s present, and these d a t a must b e combined with fundamental geologic knowledge. As is t r u e f o r most geologic studies, concIusions should n o t b e formulated from a single observation, and this is t r u e f o r any study involving fluid inclusions. From a fluid-inclusionistls viewpoint, quartz formed in t h e epithermal environment can be divided into t w o groups--quartz t h a t contains fluid inclusions large enough t o study bl.5pm) a n d quartz t h a t contains very f e w o r no inclusions large enough t o study with a standard petrographic microscope equipped for t o t a l magnifications a t least a s high a s 480X. Figure 5.IA shows t h r e e different types of quartz commonly found in t h e epithermal environment t h a t contain few, if any, inclusions large enough t o study. The lower part of Figure 5.1A shows finely crystalline, euhedral q u a r t z crystals t h a t grew contemporaneously from many nucleation sites. This t e x t u r e is often noted a s l a t e vug-fillings, but may also be a result of replacement (silicification) of original wallrock. The finely crystalline quartz seldom contains fluid inclusions large enough t o study, but on t h e r a r e occasion when appropriate inclusions a r e found (see below), homogenization t e m p e r a t u r e s a r e typically 600

Reference Ohmoto and Lasaga (1982)

350-1050

Ohmoto and Rye (1979)

0.45(106/T2)

uncertain

Ohmoto and Rye (1979)

0.40(106/T2)

200-700

Ohmoto and Rye (1979)

50-705

Ohmoto and Rye (1979)

-0.05(10~/~~)

200-600

Ohmoto and Rye (1979)

-0.16(106/T2)

200-400

Ohmoto and Rye (1979)

uncertain

Ohmoto and Rye (1979)

-0.63(106/T2)

50-700

Ohmoto and Rye (1979)

-0.70(106/T2)

uncertain

Ohmoto and Rye (1979)

-0.75(106/T2)

uncertain

Ohmoto and Rye (1979)

uncertain

Ohmoto and Rye (1979)

uncertain

Ohmoto and Rye (1979)

4.70(106/T2)

0.5

o.~o(~o~/T~)

-0.25(10~/~~)

eq. 11 -0.80(10~/~~) Comment

1. Mineral abbreviations and compound states are as follows: ag, argentite; bn, bornite; cb, cinnabar; cc, chalcocite; cp, chalcopyrite; gn, galena; H2S, gaseous; mo, molybdenite; pyl pyrite; SO, native sulfur; sb, stibnite; sl, sphalerite; SO2, gaseous; SO;, either aqueous or solid sulfate.

107

C. W. FIELD & R. H.FIFAREK referenced t o silicate-H20 reactions and t h a t of 3 4 in ~ sulfides t o sulfide-H2S reactions, i t follows t h a t more geologically useful equations f o r t h e fractionation of these isotopes between mineral phases, such a s between silicate-silicate and sulfide-sulfide pairs, may derived from algebraic identities. For example, the "0 fractionation equations given for t h e quartz-H20 and barite-H20 equilibrations (Table 6.3, eqs. (3) and (5)) a r e likely t o b e of interest only t o some geothermalists, whereas their algebraic difference

qtz-H20

1000 l n a

-

1000 Ina

bar-H20

yields t h e fractionation equation for t h e quartz-barite reaction, which is an assemblage of in r e s t t o many explorationists. By having analyses of $0 performed on t h e members of this pair, and applying t h e approximation given by equation (4)

'OoO

qtz-bar " Aqtz-bar

lna =

alsoqtz OIOO -

~ 1 8 %O/oo ~ ~

(11)

t h e depositional t e m p e r a t u r e (in OK) of t h e quartzbarite assemblage is readily calculated by substitution and solving f o r T in equation (10). Listed in Table 6.5 a r e t h e equilibrium fractionation equations, and their analogues rearranged in t e r m s of T, for isotopic mineral pairs of oxygen and sulfur in potentially useful assemblages. Variations of their respective 1000 l n a (=A) values over t h e t e m p e r a t u r e range from O°C t o 6 0 0 ' ~ a r e illustrated in Figure 6.5. The equations listed have either been calculated from those in Tables 6.1-6.4 by t h e method described (eq. 101, o r they have been taken from sources previously noted such a s Friedman and OfNeil (19771, Matsuhisa e t al. (1979), Ohmoto and Rye (19791, etc. Many other equations for mineral-mineral reactions can b e derived from those provided in Tables 6.1-6.4, but those given in Table 6.5 a r e restricted t o mineral pairs commonly found in hydrothermal deposits. The use of t h e s e and other fractionation equations in isotope geothermometry requires further elaboration. Because of problems with disequilibrium t h a t a r e especially prevalent a t low reaction temperatures, not a l l of t h e equations in Tables 6.1-6.5 a r e equally reliable f o r isotope geothermometry. For example, t h e quartz-H20 reaction has been repeatedly investigated (see Friedman and O'Neil, 1977; O'Neil, 1977; Matsuhisa e t al., 19791, because t h e exchange of oxygen isotopes between q u a r t z and water is particularly slow, and probably most other systems will be reexamined eventually. For similar reasons, isotopic equilibrium is unlikely t o be attained with sulfate-sulfide reactions below 350°c, although i t may be enhanced by conditions such a s long residence t i m e and(or) low r a t e of cooling, low pH, and high sulfur content of t h e fluids (Ohmoto and Rye, 1979; Ohmoto 1982). Pyrite-chalcopyrite pairs and Lasaga, commonly give unreasonable isotopic temperatures

TEMPERATURE OC

Figure 6.5. Fractionation curves for isotopic mineral pairs of oxygen and sulfur.

indicative of disequilibrium (Field and Gustafson, 19761, whereas sphalerite-galena pairs generally do not (Ohmoto and Rye, 1979). Furthermore, t h e carbonbearing minerals a r e not only few in number, but t h e f l a t slope of t h e dolomite-calcite fractionation curve with respect t o t e m p e r a t u r e renders this geologically common mineral pair of l i t t l e value a s a geothermometer; in contrast t o t h e steeper slope of t h e calcite-graphite fractionation curve (Fig. 6.1, curves 8 and 3). Application of hydrogen isotopes t o geothermometry is severely limited because of uncertainties caused by compositional variations of t h e minerals, post-depositional re-equilibration, and for s o m e reactions t h e insensitivity of fractionation t o changes in t e m p e r a t u r e (Fig. 6.2). The 1 8 0 compositions of t h e silicate minerals a r e subject t o reequilibration by aqueous fluids, and this postdepositional e f f e c t is largest for t h e feldspars and biotite, and l e a s t for quartz and magnetite (Brigham, 1984, and references cited therein). Finally, t h e geologist must exercise c a r e in the selection of samples f o r analysis. Although minerals t h a t exhibit t e x t u r a l evidence of noncontemporaneity a r e unlikely t o be in isotopic equilibrium, t h e converse is not always true. Provenance--Geochemical investigations commonly have a s their objective t h e identification of t h e source or provenance of fluids and(or) minerals by means of chemical tracers. Stable isotopes have been successfully applied t o many of these endeavors. In particular, t h e isotopes of hydrogen and oxygen a r e used routinely t o identify t h e source of aqueous fluids,

108

CHAPTER 6

T a b l e 6.5 F r a c t i o n a t i o n and t e m p e r a t u r e e q u a t i o n s f o r i s o t o p i c m i n e r a l p a i r s of oxygen and s u l f u r Mineral P a i r

1000 l n a

A

TOK

+ + + +

0.39 (250-500°C) 2.56 (500-80o0c)

2.19(lo3)/(n 1.88(103)/(n

0.54 (250-500°c)

l.13(lo3)/(~- 0.54)l/~

+

3.48 (250-500°C) 5.65 (500-80o0c)

0.58(103>/(~ 3 . 4 ~ ) ~ ~ ~ -0.97(103)/(n - 5.65)lI2

0.96(lo6/T2) -0.33(lO6/T2)

+ +

0.58 (250-50o0c) 2.75 (500-650'~)

0.98(lo3)/(n -o.57(lo3)/(~

anh-qtz

0.54(106/T2) 1.83(lO6/T2)

+

0.41 (250-500°C) 1.76 (500-800'~)

o . 7 3 ( i o 3 > / ( ~- 0 . 4 1 ) ~ ~ ~ 1.35(lo3)/(n + 1.76)ll2

qtz-cal

0.56(lo6/T2) -0.73(lO6/T2)

0.42 (250-500°C) 1.75 (500-800°C)

0.74(io3)/(~+ 0.42)l/~ -0.85(lo3)/(~ - 1 . 7 5 ) ~ ~ ~

1

qtz-mag

4.81 ( l o 6 / T 2 ) 3.52(lo6/T2)

2

qtz-kal

1.29(106/T2)

3

qtz-bar

0.34(lO6/T2) 0.95(lO6/T2)

4

qtz-mus

6

7

-

+

10

PY-ag

1.20(10~/~~)

12

sl-gn

0.73(lO6/T2)

(uncertain)

(50-70o0c)

-

0.39)ll2 2.56)ll2

-

-

0.58)l/~ 2.75)ll2

l . l o ( l ~ ~ ) / ( ~ ) ~ / ~

0.85(10~)/(~)~/~

Comments 1. These e q u a t i o n s a r e d e r i v e d from t h o s e l i s t e d i n T a b l e s 6.3 and 6.4 by methods described i n the text. 2. M i n e r a l and compound a b b r e v i a t i o n s a r e a s f o l l o w s : Ab, a l b i t e ; a g , a r g e n t i t e ; a n h , a n h y d r i t e ; b a r , b a r i t e ; c a l , c a l c i t e ; c p , c h a l c o p y r i t e ; gn, g a l e n a ; k a l , k a o l i n i t e , K f , K - f e l d s p a r ; mag2 m a g n e t i t e ; mus, m u s c o v i t e ; p y , p y r i t e ; q t z , q u a r t z ; s l , s p h a l e r i t e ; and SO4 e i t h e r aqueous o r s o l i d s u l f a t e .

whether o r n o t now present, in many rock- a n d oreforming environments. The viability of this m e t h o d s t e m s from t h e f a c t t h a t t h e isotopic domains of m a g m a t i c , m e t e o r i c , and o c e a n w a t e r s a r e generally distinct, and t h a t t h e composition of fluids (hydrothermal, metamorphic, etc.) m a y b e c a l c u l a t e d from t h e analytical d a t a for o n e or m o r e minerals using t h e appropriate reaction equations (Tables 6.2 a n d 6.3). F o r example, given t h e AD and 6180 values f o r a sample of hydrothermal s e r i c i t e , and assuming equilibrium and a t e m p e r a t u r e of f o r m a t i o n based on independent criteria (fluid inclusions, mineral

assemblages, sulfur isotopes, etc.), t h e compositions of t h e fluid may be d e t e r m i n e d by substitution of t h e d a t a and assumed t e m p e r a t u r e (in OK) in t h e muscovite-H20 r e a c t i o n equations ( a b l e 6 . 2 , eq. ( 5 ) ; Table 6.3, eq. (8)). Similarly, t h e 6 I 0 composition of t h e fluid may b e c h e c k e d f r o m a n additional analysis of coexisting q u a r t z a n d use of t h e quartz-H 0 r c t i o n (Table 6.3, eq. (3)). In addition, having t i e 6"O values of both q u a r t z and s e r i c i t e allows use of t h e quartz-muscovite isotopic g e o t h e r m o m e t e r (Table 6.5, eq. (411, which t h e r e b y s e r v e s t o c h e c k t h e validity of t h e assumed t e m p e r a t u r e .

6

C. W. FIELD & R. H.FIFAREK

109

There a r e variants t o t h e foregoing procedures, and where applicable t h e y may offer advantages in t e r m s of t i m e and expense. Provided i t can b e assumed that the hydrothermal fluid was predominantly of m e t e o r i c origin and t h a t t h e system was characterized by large water-to-rock ratios, i t is possible t o e s t i m a t e both t h e hydrogen- and oxy enisotopic composition of t h e fluid from a single "0 analysis of q u a r t z or other oxygen-bearing mineral. This approximation is feasible, t o t h e e x t e n t t h a t a l l pfgceding assumptions a r e valid, because t h e D and meteoric waters, although 6 0 contents of extraordinarily variable, change sympathetically and linearly according t o t h e equation

c i t e d by Craig (1966) and Taylor (1974a). However, because of t h e qualification with respect t water:rock ratios, which if low may perturb t h e 61'0 compositions of fluids and minerals precipitated t erefrom, i t is normally customary t o calculate t h e 61'0 values of t h e fluids from t h e 6 D values of minerals using equation (12), or from measured 6 D values of fluids e x t r a c t e d from mineral inclusions. Me o r i c waters a r e characteristically depleted in D and "0 by values t h a t may e x c e e d -150 and -20 permil, respectively, r a t i v e t o their oceanic source (0 O/oo for 6 D and Sf4 0, by definition). These depletions t a k e place a s a result of equilibrium fractionations at low temperatures that accompany t h e progressive condensation and crystallization of rain and snow from a finite quantity of atmospheric water vapor (see Figs. 6.2 and 6.3, curves 1 and 12, respectively). Their magnitudes vary directly with altitude, latitude, and t h e relative amount of water vapor removed from t h e system. Because isotopic studies have demonstrated t h a t waters of meteoric origin dominate geothermal and most hydrothermal systems, especially those of epithermal character, t h e present-day compositions of these waters over much of North America is shown in Figure 6.6 (after Taylor, 1979, p. 243). Note t h a t precipitation over Nevada, and most of the Basin Range province, is characterized by 6 D and 6 0 values of -130 t o -80 and -80 t o -11 permil, respectively. These depletions a r e similar t o those of "fossil" waters calculated from t h e mineral d a t a for Tertiary hydrothermal systems of this region. However, i t is appropriate t o conclude this discussion of "calculated" fluid compositions with a n o t e of caution. Studies by Truesdell (1974) and o t h e r s and t h e d a t a and discussions of Friedman and O'Neil (1977) and Taylor (1967, 1979) indicate t h a t t h e calculated compositions of water from saline hydrothermal l u ~ d smay b e in with respect error, and usually depleted in D and t o t h e t r u e values. This solute e f f e c t is caused by t h e tendency of s o m e cations, particularly those of l a r g e ionic potential, t o hydrate in solution with t h e isotopically heavier molecules of water, which thus leaves t h e f r e e water t h a t equilibrat with minerals Although proportionately depleted in D and "0. difficult t o evaluate, this solute-controlled isotopic e f f e c t is probably small in epithermal systems having fluids of low salinity.

?dld

"0

Figure 6.6. Distributions of 6 D and 6180 in meteoric waters over part of North America (after Taylor, 1979).

Other applications of t h e stable isotopes t o investigations of provenance warrant brief mention. It is well known t h a t authigenic minerals formed in sedimen ry environments a r e conspicuously enriched in '$0 relative t o their m a g m a t i c counterparts because of t h e larger fractionations permitted a t low te eratures. Smaller, yet significant, enrichments of7$ found in s o m e igneous rocks and minerals have been used in conjunction with other geochemical and geologic d a t a t o document examples and sources of rnagma contamination and t h e sedimentary component of S-type granites (Magaritz e t al., 1978, and references therein). The isotopes of carbon and sulfur also undergo appreciable fractionations a t low t o intermediate temperatures, and these may b e enhanced when t h e reaction pairs involve both oxidized and reduced forms of these elements (Fig. 6.1, curves As a 1, 2, and 3; Fig. 6.4, curves 1 and 2). nsequence, t h e carbonates a r e commonly enriched in I3C relative t o graphite, organic carbon, hydrocarbons, and marine sulfates a r e enriched in relative t o m a g m a t i c sulfur ( " 0 O/oo) and in contrast These t o 3 4 ~ - d e p l e t e d sedimentary sulfides. enrichments and depletions originally served a s a qualitatively convenient means by which t o interpret t h e isotopic d a t a in t e r m s of oceanic, magmatic, and biogenic sources and processes. However, Sakai (1968) and Ohmoto (1972) demonstrated t h e fallibility of subjective interpretations by showing t h a t t h e isotopes of carbon and sulfur could undergo large fractionations a t relatively high t e m p e r a t u r e s by way of inorganic redox reactions controlled by Eh and pH. In spite of

errors inherent with isotopic generalizations, Thode et al. (1954) and Feely and Kulp (1957) deduced from carbon- and sulfur-isotopic d a t a t h e biogenic origin of c a l c i t e and native sulfur in t h e c a p rock of Gulf Coast s a l t domes from sources of evaporitic anhydrite a t depth. In addition t o t h e scientific merits of their contribution, t h e practical corollary was t h a t s a l t domes lacking a n anhydrite-calcite c a p rock a r e unlikely t o contain economic accumulations of native sulfur and petroleum! GEOLOGIC DISTRIBUTIONS Abundances of t h e s t a b l e isotopes in geologically important environments a r e now summarized and briefly described. The purposes a r e twofold: first, these d a t a o f f e r background and perspective for discussions of t h e epithermal environment t h a t follow; and second, t h e y exhibit trends t h a t a r e largely consistent with those derived from theory and experiment a s previously described. For some readers i t may b e disconcerting t o observe t h a t t h e isotopic range f o r any of t h e four e l e m e n t s may be broad and overlapping from one environment t o another. However, this a p p a r e n t lack of isotopic definition is principally a n a r t i f a c t of these compilations. The isotopic signature of a n e l e m e n t for a particular member of a n environment, such a s a single mineral deposit, plutonic phase, sedimentary formation, etc., is usually narrow and well defined. Nonetheless, t h e environment m a y exhibit a larger isotopic range because i t s members may b e compositionally dissimilar owing t o differences of provenance, depositional conditions, and age. Hydrogen and Oxygen I t is evident from previous discussions t h a t D and ''0 form a n isotopic couplet t h a t is uniquely suited t o t h e study of fluids and minerals in aqueous systems. However, because oxygen is a major constituent of t h e c r u s t ("46.6 wt%), whereas hydrogen (20.1 wt%) is not, t h e ''0 c o n t e n t of fluids is more likely t o b e perturbed by water-rock reactions than is t h e D content. Accor ngly, i t is useful t o first review t h e distributions of "0 in various geologic environments. These d a t a , a s portrayed in Figure 6.7, a r e taken largely from t h e summaries of Taylor (1967, 1974a, 19791, Garlick (1972 F a u r e (l977), Hoefs (19801, and other sources. T h e 6 0 values for ultramafics ( > 5 and < 7 O/oo) a r e similar t o those of meteorites (-3 t o 7'/00) and a r e consistent with t h e presumed m a n t l e origin of t h e former. Mor siliceous igneous clans a r e progressively enriched in "0 for t h e sequence from basalts and gabbros (-5.5 t o I2 O/oo), t o rhyolites and granites ("6 t o 1 3 O/oo). This generalized isotopic trend is compatible with observed fractionations between t h e common rock-forming minerals wherein 6180 values a r e largest in quartz, carbonates, and alkali feldspars; intermediate in plagioclase feldspars, micas, and ferromagnesian minerals; and smalles in t h e Fe-Ti oxides. A similar trend of diminishing values is evident among the authigenic minerals of

1%

5'0

marine sedimentary rocks t h a t include c h e r t s (-20 t o 39 O/oo: Garlick, 1972; Kolodny and Epstein, 19761, carbonates (-15 t o 36 O/oo: Garlick 1972; Veizer and Hoefs, 19761, shales (- 11 t o 29 b loo: Savin and Epstein, 1970b), and ferromanganese dules ("10 t o I 4 O/oo; Field et al., 1983). The largel'O enrichment of sedimentary rocks and minerals a s compared t o those of m a g m a t i c origin results from t h e larger fractionations permitted a t t h e low t e m p e r a t u r e s prevailing in t h e hydrosphere, and in spite of t h e f a c t t h a t magmas a r e enriched in ''0 (-6 t o 10 O/OO) relative t o s e a water ("0 O/oo). This distinction between isotopically heavy authigenic sedimentary minerals and their lighter magmatic counterparts gives support t o hypotheses of assimilation or a n a t e c t i c melting of sedimentary r ks t o account for t h e f e w documented examples of "0-enriched (up t o 16 O/oo) plutonic and volcanic rocks (Taylor and Turi, 1976; Magaritz e t al., 1978). By analogy, t h e sources of fluids and mineral constituents in hydrothermal systems m a y also b e constrained by similar isotopic differences. The d a t a f o r hydrothermal deposits in Figure 6.7 a r e subdivided between Cordilleran and volcanogenic massive-sulfide types, with t h e principal differences being t h a t t h e former a r e associated with epizonal plutons whereas t h e l a t t e r a r e deposited in a submarine environment on or short distances below t h e s e a floor (Sawkins, 1972). For purposes of comparison, we have excluded t h e d a t a for epithermal deposits from t h e Cordilleran subgroup a s su marized in Figures 6.7, 6.9, and 6.10. Values of 6 18"0 for q u a r t z (2-4 t o 1 3 O/oo) carbonates ( < 6 t o 14O/oo), and sulfates (-6 t o 20 d/oo) of t h e Cordilleran subgroup a r e chiefly from Sheppard e t al. (19711, Fuex and Baker (19731, and Watanabe and Sakai (1983). Although t h e r e a r e abundant d a t a f o r hydrothermal silicates other than q u a r t z (feldspars, micas, clays, oxides, etc.), they a r e not illustrated in Figure 6.7 t o preserve clarity among t h e dominant mineral phases. Their isotopic distributions would largely mim' those of quartz, but they would be more depleted in '$0 because of smaller fractionation f a c t o r s a t any given t e m p a t u r e (Table 6.3 a n d Fig. 6.3). Distributions of 6 0 values in q u a r t z (-7 t o 14 O/oo), carbonates (=9 t o 20°/oo), and sulfates ("5 t o 15 O/oo and more) of t h e volcanogenic massive-sulfide deposits a r e taken from Kusakabe and Chiba (19831, Watanabe and Sakai (1983), and Fifarek (1985). Although t h e d a t a base for massive-sulfide deposits is less extensive than t h a t f o r t h e Cordilleran types, their minerals appear be isotopically less varlable and more enriched in '0. Such differences, if real, probably r e l a t e t o t h e larger range of depositional t e m p e r a t u r e s and differing sources of t h e aqueous component (magmatic versus meteoric) in Cordilleran hydrothermal systems. D a t a relevant t o t h e foregoing iscussion a r e given by t h e distributions of dD and 61%0 in various waters and minerals, a s displayed in Figure 6.8, provide d a t a relevant t o t h e foregoing discussion, and a r e illustrative of several useful applications of t h e hydrogen-oxygen isotope pair. The isotopic locations of standard mean ocean water (SMOW) and t h e m e t e o r i c w a t e r line (MWL) a s derived from equation 12 (Craig, 1966; Taylor, 1974a), plotted a t t h e t o p c e n t e r and diagonally down t h e left-hand margin,

f5

Environment

nereorlres 1gneovs R o c k s ulrramafr~s basaits/gabbros andeaiLes/granod~orires rhyol,tes/graniten

Sedimentary Rocks

chert= Carbonares

shales sulfates Fe-"n nodules H y d r o t h e m a l Deposrtr Cordilleran

quartz carbonates sulfates Vols. Hass. Sulfide

quartz carbonates Sulfates

Figure 6.7. Distributions of 6180 in geologic environments.

Figure 6.8. Distributions of 6 D and 6 180 in various waters, minerals, and hydrothermal fluids of epithermal deposits.

respectively, a r e useful points of reference. The diagonal a t t h e right-hand margin represents t h e kaolinite line (KL) c i t e d by Sheppard e t al. (1969) and based on t h e work of Savin a n l f p s t e i n (1970a), which marks a continuum of 6 D and d 0 values in kaolinites from modern soils. Parallelism of lines MWL and KL documents equilibrium isotope exchange of D and 1 8 0 between m e t e o r i c waters and kaolinites formed during t h e weathering and transformation of rocks t o clayrich soils. The compositional domain of m a g m a t i c waters, given by t h e rectangular box a t center-right of Figure 6.7, has been calculated from isotopic analyses of hydrous m a g m a t i c silicates. Most m a g m a t i c waters have a relatively confined range of values e t w e e n -85 and -40 permil D and 5.5 t o 9 permil 6 & 0 (Taylor, 1974a; 1979). Not illustrated, for reasons of clarity and possible lack of relevance, a r e t h e broad isotopic fields of metamorphic waters and t h e saline brines of sedimentary basins. According t o Taylor (1979), t h e metamorphic waters inherit their compositio a1 variability ('-65 t o -20 O/oo AD, and 5 t o 25 '100 6I 60 ) from dehydration and fluid-mineral reactions with isotopically variable igneous and sedimentary rocks, whereas t h e d a t a f o r t h e brines a r e broadly s c a t t e r e d t o t h e right of t h e MWL and suggest varying mixtures of both connate and meteoric waters t h a t have been modified by other fluid-sediment interactions a t depth within t h e basins. Compositional variations of waters associated with a number of well-known geothermal systems, a s modified from t h e d a t a of Craig (19661, White e t al. (19731, and Truesdell and Hulston (19801, a r e plotted on Figure 6.8 Surface waters (large closed circles) have 6 D and 6180 values located on or near t h e MWL, whereas t h e values for related subsurface w a t e r s (small closed circles) trend variably and horizontally t o t h e right (dashed lines) from the MWL. Such trends, known a s t h e "oxygen isotope shift", a r e c h a r a c t e r i z e d by increasing values of 6180 a t nearly constant 6 D and a r e common t o t h e subsurface fluids of many

geothermal systems. The shift t o larger 6180 values is caused b equilibrium isotope-exchange r e a c i ns between r80-depleted m e t e o r i c waters and "0enriched rocks during I%ater-rock reactions. In general, t h e size of t h e 0-shift correlates directly with t e m p e r a t u r e and salinity of t h e fluids, and inversely with t h e mass ratios of w a t e r t o rock. Because t h e e f f e c t s of t h e 180-shift and water:rock ratios are interrelated and may influence interpretations of t h e analytical d a t a for fluids and minerals, these phenomena will b e discussed a t g r e a t e r length in forthcoming considerations of t h e epithermal deposits. However, t h e reader should n o t e t h a t m e t e o r i c waters similar t o those within t h e compositional range from Wairakei/Broadlands t o t h e Sal n Sea (6D 2-40 t o -80 O/oo) could b e driven by a n '%-shift into t h e isotopic domain of m a g m a t i c w a t e r s by exchange reactions involving small t o subequal amounts of water relative t o t h a t of rock (low t o intermediate water:rock mass ratios). Values of AD, in contrast t o those of 6180, a r e not significantly a f f e c t e d by water:rock exchange because a t t h e s e water:rock mass ratios t h e principal source of hydrogen is in t h e aqueous fluids. Thus, t h e geothermal fluids retain t h e 6D value of their m e t e o r i c source, and water-rock isotope-exchange reactions a r e manifest a s t h e subhorizontal trend lines of Figure 6.8. The slight positive slope t o a f e w trend lines, such a s f o r t h e Salton S e a and Yellowstone geothermal areas, may result from deuterium enrichment t h a t is unrelated t o water-rock reactions. Concentration of deuterium is caused by evaporation of surface waters prior t o recharge a t t h e Salton S e a (Craig, 1966) and by t h e boiling of subsurface w a t e r s a t Yellowstone (Truesdell e t al., 1977) and perhaps elsewhere. As previously noted, t h e isotopic compositions of hydrothermal and other fl i s m a y b e calculated from t h e measured 6 D and 6 0 values of associated minerals by means of t h e appropriate mineral-water

1.f

fractionation curves (Tables 6.2 and 6.3) and using a n e s t i m a t e d t e m p e r a t u r e of deposition. The "calculated" fluid compositions derived from many biotites and a f e w sericites of porphyry-type deposits plot within t h e m a g m a t i c water box (Sheppard e t al., 1971; Osatenko and Jones, 1976; Taylor, 1979). However, t h e d a t a for a f e w biotite, most sericites, and nearly all other hydrous minerals of h y d r ~ t h e r m a l ~ g r i g ianr e located within t h e broad range of 6 D and 6 0 values between t h e MWL and K L boundaries. Compositions of t h e fluids calculated from t h e isotopic d a t a for these minerals, or from t h e fluids of inclusions contained therein, suggest w a t e r s of either m e t e o r i c origin, or of mixed meteoric-magmatic or magmatic-oceanic parentage. According t o Sheppard e t al. (1969) and Taylor (1974a), minerals such a s t h e kaolinites t h a t may form either by hypogene o r supergene processes may be distinguished on t h e basis of temperature-controlled fractionations, which d e t e r m i n e their isotopic positions relative t o t h e hypogene/supergene kaolinite line of Figure 6.8. Supergene kaolinites formed a t low temperatures and in equilibration with m e t e o r i c waters plot t o t h e right of this line and up t o t h e KL, whereas t h e hypogene kaolinites formed a t higher temperatures plot t o t h e left. Also illustrated on Figure 6.8 a r e t h e l'calculated" compositions of hydrothermal fluids responsible for t h e deposition of many epithermal precious-metal deposits (closed triangles) of the western U.S. and a f e w elsewhere. These d a t a a r e from t h e results of other investigators and will be c i t e d subsequently. They a r e based on t h e 6 D values obtained from ?Ifid inclusions and(or) hydrous minerals, and on 6 0 values determined from q u a r t z and(or) other associated minerals. The broad isotopic distribution of these "calculated" compositions largely precludes a significant input of m a g m a t i c water t o t h e s e hydrothermal systems, but i t does record a major contribution from m e t e o r i c sources. Further detail and elaboration about these fluids will be deferred t o our concluding discussion of t h e epithermal deposits. Carbon I 3 c in various geologic Abundances of environments a r e illustrated in Figure 6.9, and they a r e based largely on t h e d a t a c i t e d by C r a i g (19531, Bender (1972), Fuex and Baker (1973), Ohmoto nd R y e (1979), and Hoefs (1980). Values of 614C a r e surprisingly variable in t h e products of hight e m ~ e r a t u r eenvironments such a s m e t e o r i t e s (--I2 t o 9 Ojoo), igneous rocks ('-10 t o 3 O/oo), diamonds ( -6 t o 3 O/oo), and carbonatites (2-10 t o 2 O/oo); particularly because t h e compositional e x t r e m e s have been o m i t t e d from these ranges. The causes of such isotopic variability a r e uncertain, but possibly r e l a t e redox reactions, to equilibriurn or kinetic contamination, inhomogeneities of source, and(or) differing proportions of compositionally distinct carbon in t h e samples. For example, t h e d a t a given for meteorites and igneous rocks ' s t h a t of t o t a l and carbon, which consists both of "C-enriched oxidized (carbonate, up t o 66 '100) and 13c-depleted and reduced (graphite, "organic," carbonyl, etc.; up t o

-30 O/oo) forms of carbon. In accordance with fractionation theory (Table 6.1 and F' 6.1), similar relative enrichments and depletions of "C a r e present between t h e oxidized (CO -10 t o 2 '100) and reduced (CHb = -31 t o -16 ~ / o $ ' c o m p o n e n t s of volcanic/ geot ermal gases. Narrower compositional variations c h a r a c t e r i z e marine limestones (2-5 t o 4 O/oo), s e a water H C O (-5 t o -2 O/oo), a atmospheric C 0 2 ("-8 t o -6 O/oo), and t h e relative "C enrichments among t h e s e compounds ( C a C o 3 > HCO?COZ) a r e consistent with experimental-theoretical f r a c t ~ o n a t i o n trends. T h e cluster of values around and near 0 permil is expectable because t h e carbon PDB standard is calcite of marine derivation. Fractionations t h a t accompany kinetic photosynthetic reactions of atmospheric and hydrospheric C 0 2 t o form organic carbon lead t o marked depletions of 1 3 c in marine and land plants ('-34 t o -12 O/oo), and t h i s isotopic record of biogenic processes is preserved in carbonaceous sediments, coal, and petroleum ('-35 t o -10 '100). Although t h e pronounced isotopic differences between organic carbon and inorganic carbonates seem academic in view of obvious visible differences between these compounds, they do s e r v e a s a t r a c e r of biogenic precursors where reduced forms of carbon become oxidized and remobilized in some magmatic, metamorphic, and hydrothermal (?) environments. The d a t a for hydrothermal carbonates provided by Sheppard e t al. (19711, Fuex and Baker (19731, Ohmoto and R y e (19791, and Fifarek (1985) show remarkably little isotopic variability (--I 1 t o 1 O/oo), regardless of textural variety or g e n e t i c occurrence. According t o Ohmoto and Rye (19791, the 6 1 3 c value of carbon in mantle-derived igneous rocks is -5 2 permil, and this value is not demonstrably different from t h a t of carbon in average sedimentary or crustal rocks based on considerations of mass balance. Thus, carbon isotopes do not offer promise a s a means for distinguishing between mantle and crustal sources of magma and igneous rock. The carbon in minerals and fluids of hydrothermal systems may be derived from mantle sources, a s is permissive from t h e isotopic evidence, o r from diverse sources in country rocks either by oxidation of reduced forms o r by dissolution and(or) decarbonation reactions of carbonates (Ohmoto and Rye, 1979).

*

Sulfur Isotopic abundances of sulfur portrayed in Figure 6.10 a r e taken principally from t h e d a t a and summaries of Field (19721, Field e t al. (1976, 1983, 19841, Ohmoto and Rye (1979), Claypool e t al. (1980), Hoefs (19801, and Sakai e t al. (1984). Compositions of orthomagmatic t o t a l sulfur in igneous rocks (2-3 t o 3 O/oo) a r e predictably close t o t h a t of meteoritic However, those of t h e sulfide-sulfur (20 O/oo). component oxidized and 3 4 ~ - e n r i c h e d sulfate (up t o 10 '100) and reduced and 3 4 ~ - d e p l e t e dsulfide (up t o -10 '100) f o r m s a r e more variable, because of redox reactions. The 6 3 4 ~values of m a g m a t i c Cu-Fe-Ni sulfides in layered m a f i c intrusions (--6 t o 14 '100) exhibit larger variations attributable both t o contamination from sedimentary sources in nearby country rocks and t o redox reactions within t h e host

Environment ueteorites:

total-c

total-C

Igneous ROC*.:

carbonatites diamonds Volcanic/Geothemal

carbonate*

'%

=O2

CH4 Graphite A-spheric

~ t o their counterparts generally depleted in 3 4 relative in volcanogenic massive-sulfide deposits. The Cordilleran subgroup includes a large spectrum of porphyry, vein, and replacement types of deposits t h a t apparently a r e devoid of a distinctive isotopic signature regardless of differences in host rock, metals, minera and depositional textures of o r e and d a t a for sulfates and sulfides of a gangue. The few deposits have values suggestive of country rock contamination, but t h e majority a r e consistent with derivation from a -0 permil source of deep-seated "magmatic" sulfur. In contrast, sulfates and sulfides of t h e v lcanogenic massive-sulfide subgroup a r e normally "S-enriched because they derive their sulfur largely o r entirely from isotopically heavy sea-water sulfate. D a t a f o r t h e sulfates (= 12 t o 39 O/oo) a r e essentially equivalent t o those of temporally similar marine evaporitic sulfates (-10 t o 30 O/oo), and those for associated sulfides (28 t o 22 O/oo) a r e correlative ~ about 15 t o t o t h e age-trend but depleted in 3 4 by 18 O/oo relative t o oceanic sulfate-sulfur a s a consequence of t e m p e r a t u r e dependent fractionation (Sangster, 1968; Franklin e t al., 1981; Fifarek, 1985). The d a t a for a single mid-ocean hydrothermal vent (21' North, E a s t Pacific Rise, Baja, California) a r e illustrative of such isotopic effects under contemporary oceanic conditions (Styrt e t al., 1981).

C02

sea water "CO;

~ilneetones l m r i n e ) organic Carbon marine plants coal, p e t r o l e m , and ~arbon.CeOYB matter n i s ~ i s s i p pvalley: ~ caco3 Hydrothermal:

C.C03

PorPhYry-tyPe vein-type replacement-type "01C. ma*..

sulfide

Figure 69. Distributions of 613c in geologic environments.

magma chambers. Sulfur-bearing products of volcanic/geothermal emanations are broadly isotopically similar to, but more variable, than 3 O/oo), and they a r e m a g m a t i c sulfur (=-3 t increasingly depleted in 3'S with some overlap in t h e redox sequence from SO (=-8 t o 18 O/oo), through native sulfur (2-15 t o 3 6 O/oo), t o H S (--9 t o EPITHERMAL DEPOSITS 6 O/oo). Sea w a t e r s u l f a t e has a value of 2' 0 *-I O/oo a t present, but has ranged from about 10 t o 30 O/oo Distributions of t h e stable isotopes in epithermal over Phanerozoic t i m e a s deduced from studies of deposits a r e now considered. Also included a r e t h e m rine evaporites (Claypool et al., 1980). Thus, t h e d a t a for several geothermal systems because they a r e 3eS-age curve serves within broad limits t o d a t e regarded by many t o b e contemporary analogues of t h e marine sedimentary s t r a t (and volcanogenic massiveepithermal environment (White, 1981). For t h e sulfide deposits). The "S-enriched sulfate (=20 t o purposes of comparison and discussion, we have 35 O/oo) and sulfide ('-6 t o 25 O/oo) minerals of t h e subdivided t h e epithermal deposits into sedimentand(or) evaporites. Sedimentary sulfides of hosted, volcanic-hosted, and zoned polymetallic vein diagenetic-syngenetic or l a t e r epi e n e t i c origin have occurrences. Our subdivisions differ partly from those a n extraordinarily large range of 6 4~ values (2-50 t o of Hayba e t al. (1985, this volume) in t h a t most o r a l l 50 O/oo). Such large variations, especially t h e 3 4 ~ deposits of our zoned polymetallic vein subtype a r e , or depletions, have been documented by experimental may be, grouped in their Adularia-Sericite subtype of investigations, and they a r e caused by kinetic fractionations of -20 t o -50 permil and more t h a t may accompany t h e biogenic reduction of SO; t o H S 5 /w -10 -10 O +lo 2 0 t10 +,P (Ohmoto a n g p y e , 1979). Although sulfides variably environmenr 4 Meteorire. S a r e considered t o be typical of t h e depleted in rgneous R O C ~ ~ r:o t a ~ - $ formed by biogenic processes, others may have "S Cu-Fe-Ri aulfides of layered mafic 1ntr. enrichments t h a t r e l a t e t o f a c t o r s such a s source, "olsanis Em.n.ci.ns reservoir of sulfate, and other conditions within t h e native sulfur system. Regardless of environment or locale, sea Wakar =zS 30; hydr thermal sulfates of hypogene origin a r e enriched Sedimentary Rocks relative t o associated sulfides, which is in $S' sulfates sulfides -50 ---------.+so consistent with fractionation theory and t h e inferred missi..lppi valley Deposits range of depositional t e m p e r a t u r e s from about 2 0 0 ' ~ sulfates sulfides t o 6 0 0 ' ~ (Figs. 6.4 and 6.5). Contrary t o t h e HydIOrheFD.1 Deposit. impressi n given by t h e d a t a in Figure 6.10, variations Cordillar.": ."Ifate* avlfides in t h e a 4 S values f o r either sulfates o r sulfides of "010. msa: sulfates individual deposits rarely exceed 5 t o 7 O/oo, although sulfides Ocean I ~ d g e : sulfatea s o m e deposits may exhibit a compositionally distinct sulfldell cluster of absolute values t h a t may r e l a t e t o source, age, and(or) unique conditions within t h e system (Field Figure 6.10. Distributions of 6 3 4 ~ in geologic e t al., 1983; Fifarek, 1985). Both t h e sulfates and environments. sulfides of Cordilleran-type hydrothermal deposits a r e

5

0

----------

-

-

-

-

volcanic-hosted epithermal deposits. These deposits may exhibit characteristics suggestive of a deeper level of hydrothermal mineralization, such a s fluid inclusions having higher salinities and homogenization temperatures; a not uncommon magmatic component t o t h e hydrothermal fluids a s deduced from hydrogenand oxygen-isotope data; host rocks t h a t include plutonic and sedimentary lithologies; alteration t h a t lacks widespread and pervasive zones of advanced argillic and alunitic assemblages; and o r e s t h a t contain relatively abundant sulfides and sulfosalts of t h e base metals. Presumably such differences a r e not genetic, but instead represent variations in style and content of mineralization t h a t have developed from chemical and t h e r m a l gradients in deeply convecting hydrothermal systems. Carbon D a t a for 1 3 c in various geothermal systems and e ' t h e r m a l deposits a r e illustrated in Figure 6.1 1. The (PJC values for all occurrences, e x c e p t t h e Geysers and Pueblo Viejo, a r e remarkably uniform within t h e narrow range of -10 t o 1 permil, and suggest t h a t t h e carbon was derived either from magmafjic or marine limestone sources (Fig. 6.9). Extreme C depletions (-25 t o -24 O/oo) of carbonaceous material in volcaniclastic sedimentary rocks a t Pueblo Viejo (Kesler e t al., 1981) a r e typical of reduced f o r m s of organic carbon. Oxidized carbon cornpounds such a s C 0 2 , aqueous HCO, host-rock carbonate, and vein c a l c i t e a t t h e Geysers geothermal a r e a exhibit large variations i n 6 1 3 c (-19 t o 1 O/oo) according t o t h e work of White e t al. (1973) and Sternfeld (19811, and other references cited by these authors. Detailed fluidinclusion, isotopic, and mineralogical Sternfeld (1981) suggest t h a t much of t h e variability may be a t t r i b u t e d t o multiple sources of biogenic, marine, and m a g m a t i c carbon in t h e igneous and sedimentary host rocks t h a t were rernobilized and redeposited during subsequent and temporally s e p a r a t e metamorphic and geothermal events. Vein calcites diminish in 613c ( ~ 4O/oo/lOOO m) with increasing depth, as a result of temperature-ind fractionation, and a distinct population of depleted calcites is a t t r i b u t e d t o late-stage reequilibration with C02-rich steam. In contrast t o t h e Geysers, isotopic compositions of carbon in carbonate clasts of host rocks a t t h e C e r r o Prieto (Williams and Elders, 1984) and Salton Sea (Clayton et al., 1968) geothermal a r e a s and in vein carbonates from t h e s e and t h e Broadlands (Blattner, 1975) and Wairakei (Clayton and Steiner, 1975) reas of New Zealand a r e less variable. Values of 6I C in host rock and vein carbonates decrease with increasing depth a t Broadlands and C e r r o Prieto. The cause of these isotopic trends is uncertain. They may possibly r e l a t e to diminishing equilibrium fractionations with increasing temperatures a t depth, and (or) t o other complexities such a s boiling, d carbonation and (or) dissolution reactions, influx of 'JC-depleted organic carbon, and possibly other processes (see Blattner, 1975; Williarns and Elders, 1984). D a t a for carbonates from Wairakei (Clayton and Steiner, 1975) and t h e Salton Sea (Clayton e t al., 1968) do not show isotopic

4

trends re1 t e d t o depth o r reservoir temperatures. However, I3C-depleted carbonates from t h e Salton S e a field correlate inversely with t h e t o t a l carbonate (wt.%) content of t h e host, which suggests they a r e , t h e residuals of decarbonation reactions accompanied by t h e loss of 13c-enriched C 0 2 (Clayton e t al., 1968). Carbonates from unaltered and a l t e r e d host rocks and veins of t h e sediment-hosted epithermal deposits a t Carlin (Radtke e t al., 1980) and C o r t e z (Rye e t a1 1974) exhibit a narrow and overlapping range of 61'3C values (Fig. 6.11; %-6 t o 1 O/oo). These and other d a t a suggest t h a t most of t h e carbon in hydrothermal calcites was probably e x t r a c t e d by dissolution reactions from t h e carbonate host rocks a t depth. However, t h e fluids a t Carlin a1 must have contained a component of oxidized '%-depleted organic carbon from t h e host rocks, provided isotopic equilibrium prevailed, t o account for t h e relatively light compositions of one main-stage and several l a t e low-temperature vein calcites (Radtke e t al., 1980). The various carbonate minerals formed in zo polyrnetallic veins also show a narrow range of 6PPCd values ("J-10 t o 0.1 O/oo), but a s a group they a r e slightly depleted in 1 3 c relative t o those of sedimenthosted deposits. Data for rhodochrosites, manganosiderites, and siderites from Creede, Colorado (-8.2 t o -4.0 O/oo; Bethke and Rye, 19791, calcites and other carbonates (?) from Tui, New Zealand (-7.8 t o O.lO/oo; Robinson, 19741, and c a l c i t e s from Casapalca, Peru (-10.0 t o -2.6 O/oo; R y e and Sawkins, 19741, a r e collectively interpreted a s being indicative of m a g m a t i c carbon (%-5 2 2 O/oo; Ohmoto and Rye, 1979). This conclusion is also supported by hydrogen and oxygen-isotope d a t a of inclusion fluids and host minerals a t C r e e d e and Casapalca. However, t h e source of carbon in calcites and rhodocrosites a t t h e Sunnyside mine (-7.9 t o -1.8 O/oo), near Creede, is not definitive, and may have been derived in p a r t from dissolution of marine limestones o r m a g m a t i c sources a t depth, or from atmospheric C O dissolved in circulating m e t e o r i c water ( ~ a s a d e v a l ?and Ohmoto, 1977).

Geothermal Systems Geysers:

::;-

who?e rock calcite salton Sea/Cerro Prieto: carbonate host altered carb. host carbonates Broadlands/wairakei: calcite Sediment-Hosted Cortez: carbonate host altered carb. host calcite carlin: carbonate host altered carl). host calcite Volcanic-Hosted Pueblo Viejo: carbonaceous sed. Zoned Polymetallic Veins Creede: carbonates Sunnyside: carbonates Tui: carbonates Casapalca: calcite

Figure 6.11. Distributions of ma1 deposits.

613c in epither-

C. W. FIELD & R. H.FIFAREK Thus, in t h e absence of contributing information from other isotopic elements, fluid inclusions, nd assemblages of vein a n d alteration minerals, t h e C' d a t a alone a r e unlikely t o provide unique i n t e r p r e t tions t o geochemical problems. Moreover, vertical I 3 C gradients present in some geothermal systems have not been reported in t h e epithermal deposits. However, e a r l y calcites (-10.1 t o -6.1 O/oo) a r e depleted in 1 3 c relative t o l a t e calcites (-6.4 t o -2.6 O/oo) a t Casapalca (Rye and Sawkins, 19741, and rhodochrosites (-7.9 t o -6.3 O/oo) a r e depleted in 1 3 c relative t o calcites (-3.8 t o -2.8 O/oo) a t t h e Sunnyside mine (Casadevall and Ohmoto, 1977). Sulfur Abundances of 3 4 ~in geothermal systems and epithermal deposits a r e portrayed in Figure 6.12. As expected, sulfide-sulfur from whole-rock samples of post-glacial basalts near t h e geothermal fields in Iceland vield 6 3 4 ~values (-1.8 t o 0.4 O/oo) similar t o O magmatic sulfur ( ~ a k a ei t al., 1980). those O ~ permil The d a t a for aqueous SO: and sulfate minerals occupy t w o distinct isotopic populations. One group represented by samples from Iceland (Sakai e t al., 19801, Yellowstone (Schoen and Rye, 1970; Truesdell et al., 1977, 1978), and Wairakei (Steiner and R a f t e r , 1966) is enriched in 3 4 ~('15 t o 2 3 O/oo) and consists of hypogene sulfates t h a t isotopically equilibrated with H2S a t depth and a t relatively high t e m p e r a t u r e s (?300°C). The other group, also from Iceland, Yellowstone, and Wairakei, is isotopically m o r e ~ t o 12 O/oo). variable and relatively depleted in 3 4 (-6 t consists of mixed proportions of t h e deep S4S-enriched hypogene sulfate and shallow 3 4 ~ depleted "supergene sulfate formed by near-surf c e , non-equilibrium, and quantitative oxidation of S' 1977, 1978). depleted H2S (Truesdell e t al., Isotopically l ~ g h taqueous and mineral sulfates a r e not only common t o t h e acid-sulfate springs of geothermal areas, but they a r e also typical of sorne alunites and barites of volcanic-hosted epithermal deposits; particularly those associated with advanced-argillic alteration (the acid-sulfate type of Hayba e t al., 1985, this volume). Native sulfur a t Yellowstone has formed chiefly by near-surface inorganic oxidation of H2S, a s is consistent with t h e com ositional similarities between elemental (-8.4 t o 3.2 /oo) and reduced (-5.0 t o 4.0 O/oo) sulfur compounds (Schoen and R e, 1970). However, t h e relatively broad spread of d S values (8.4 t o 4.0 O/oo for so and H2S @ 0 O/oo) under e s slight fractionation upon separation into discrete "Sdepleted vapor and 34~-enriched aqueous phases (Truesdell e t al., 1978). The sulfide minerals normally exhibit narrow compositional ranges t h a t a r e similar t o those of associated H2S, although t h e r e may b e s o m e variability attributable t o differing sources of sulfur both within and between t h e geothermal fields. For example, the pyrites from Iceland consist of t w o distinct isotopic populations: one is enriched in 3 4 ~ (2.2.9 t o 7.9 O/oo) and has been derived through reduction of isotopically heavy s e a w a t e r sulfate, and ~ t o 0.9 O/oo) and has t h e other is depleted in 3 4 (-4.6 originated from a magmatic source (Sakai e t al., 1980).

B

4

115

Geothermal Systems Iceland:

magmatrc sulfur SO4 sulfates H2S pyrite Yellowstone: SO4 sulfur H2S Salton Sea: s u l f ~ d e s Broadlands: sulfides Wairakei: SO4 sulfates Hz5 sulfides Sediment-Hosted

-

Cortez:

barite diagenetic py sulfides Carlin: barite diagenetic py sulfldes Volcanic-Hosted Tolfa: sulfates sulfides Pueblo Vlelo: sulfates sulfur sulfides Goldfleld: alunrte pyrite Zoned Polymetallic Veins Sari Juan Mountains: Creede: barite sulfides Sunnyside: sulfates sulfides RICO: sulfides Ouray: sulfldes TUI: barlte sulfides Guanaluato: sulfides ~n country rock s u l f ~ d e s in volc. sulfides in ore Casapalca: sulfides Flnlandia: barite sulfides Western Cascades: sulfides Golden Sunlight: barlte sulf~des

-

45

-

---------

Figure 6.12. Distributions of ma1 deposits.

6 3 4 ~ in epither-

Ranges of 6 3 4 ~ a r e narrow for pyrite and Ag-Cu sulfides (-1.4 t o 3.0 O/oo) of t h e Salton Sea; pyrite, pyrrhotite, galena, and sphalerite (1.4 t o 5.1 O/oo) of Broadlands; and pyrite and pyrrhotite (2.7 t o 6.8 O/oo) of Wairakei. The sulfide compositions a t the Salton Sea a r e a a r e compatible with, but not proof of, a magmatic source of sulfur (White, 1968, 1974), whereas those a t t h e Dro dlands and Wairakei a r e slightly more enriched in 94S and suggest a crustal provenance either by leaching and partial reduction of sulfate from basement rocks (Drowne et al., 1975) or by magma generation in the upper c r u s t (Steiner and R a f t e r , 1966). Although t h e d a t a base is meager, and e x c e p t f o r t w o sphalerite-galena pairs from Broadlands, t h e r e is little evidence of complete isotopic equilibrium between sulfate-sulfide and sulfide-sulfide assemblages in the geothermal environment. All sulfide and most sulfate minerals in and near t h e sediment-hosted gold deposits a t Carlin (Radtke e t a1 1980) and C o r t e z (Rye e t al., 1974) a r e enriched in 34\ relative t o their counterparts in geothermal systems and other epithermal deposits. The evidence, particularly from Carlin, suggests t h a t most of t h e

116

CHAPTER 6

sulfide-sulfur in pyrite-galena-realgar-sphaleritestibnite o r e (4.2 t o 16.1 O/oo) was derived from t h e hydrothermal remobilization of diagenetic pyrite (1 1.7 t o 14.3 O/oo a t Carlin; 5.1 t o 11.4 O/oo a t C o r t e z ) in t h e sedimentary host rocks. However, t h e source of sulfate-sulfur in barites (27.8 t o 31.7 O/oo a t Carlin) is uncertain, and m a y h a v e originated either by hydrothermal solution of sedimentary barite in t h e country rocks, or by fractionation t h a t may have accompanied partial oxidation of sedimentary sulfidesulfur incorporated in t h e fluids (Radtke e t al., 1980). Regardless of origin of t h e sulfate-sulfur, i t is remarkable t h a t isotopic t e m p e r a t u r e s calculated from barite-pyrite values ( 2 5 0 ' ~ t o 3 0 5 ' ~ ) a r e reasonably consistent with those derived from t h e homogenization of fluid inclusions ( 1 8 0 ' ~ t o 365'~). The volcanic-hosted deposits of Tolfa (Italy), Pueblo Viejo and Goldfield have geologic f e a t u r e s common t o many geothermal systems in t h a t their volcanic host rocks have been pervasively a l t e r e d t o + pyrophyllite assemblages of alunite-kaolinite advanced argillic alteration. According t o Field and Lombardi (1972) and C o r t e c c i e t al. (19811, isotopic similarities between alunite and barite (1.9 t o 9.6 O/oo) and hypogene pyrite, cinnabar, and galena (3.4 t o 10.3 O/oo) a t Tolfa suggest t h a t these sulfates, and possibly marcasite (-1.5 and -0.6 O/oo), a r e of supergene origin. However, t h e sulfate-sulfide assemblages of Tolfa may consist of two isotopically 34~-depleted distinct populations because the marcasites a r e associated with t h e lightest alunites (1.9 and 2.5 O/oo). Perhaps t h e s minerals formed in surficial acid-sulfate pools from g4S-depplted H2S t h a t had separated with boiling of reservoir fluids a t depth, a s proposed by Truesdell et al. (1978) for some hot springs of t h e Yellowstone area. In contrast, most sulfates a t Pueblo Viejo a r e considered t o b e of hypogene origin by Kesler e t al. (198 because t h e relative t o barite and alunite a r e enriched in pyrite and sphalerite (18.8 t o 21.6 O/oo versus -10.1 t o However, Jensen e t al. (1971) have -3.5 O/oo). documented alunites of both hypo e n e (11.6 t o 23.3 O/oo) and supergene (-2.5 t o 1.7 8/oo) origin a t Goldfield; t h e l a t t e r group bein8 isotopically similar t o hypogene pyrite (-2.8 t o 2.4 loo) from which they derived their sulfur. Thus, isotopic and geologic evidence suggest t h a t t h e sulfates (aqueous and mineral) of many geothermal and volcanic-hosted environments a r e o supergene origin, and acquired sulfate-sulfur (as their distinc ive '4.S-depleted contrasted t o 14S-enri hed hypogene sulfates) by nearsurface oxidation of 54S-depleted hypogene sulfidesulfur (H2S and mineral; s e e Field, 1966; Schoen and Rye, 1970; Jensen et al., 1971; Field and Lombardi, 1972). The source of hydrothermal sulfur a t Goldfield is considered t o be m a g m a t i c (Jensen e t al., 1971). In contrast, t h e source a t Pueblo Viejo is thought t o b e a mixture of pristine and biogenically reduced sulfate (15 O/oo) from C r e t a c e o u s s e a water (Kesler e t al., 19811, and t h a t a t Tolfa may have been derived from gypsiferous Miocene-Pliocene mudstones t h a t underlie this young Pliocene-Pleistocene volcanic complex (Field and Lombardi, 1972). This l a t t e r interpretation is supported by t h e work of Taylor nd Turi (1976) who a t t r i b u t e t h e exceedingly high 61 60 values (15.3 t o

"ki

16.4 O/oo) of Tolfa quartz l a t i t e s and rhyolites t o m a g m a t i c assimilation of 180-enriched argillaceous sedimentary rocks a t depth. Although fractionation e f f e c t s between hypogene sulfide and sulfate-sulfide minerals of t h e volcanic-hosted group appear t o be consistent with t h e a t t a i n y g n t of a t least partial isotopic equlibrium (with 6 S values of sulfates > sulfides and those of sulfides mostly in t h e order py >sl > gn), a more precise evaluation of these apparent trends is difficult because the data for contemporaneous o r spatially associated mineral pairs and triplets a r e not available. The isotopic d a t a base for zoned polymetallic vein deposits (Fig. 6.12) is large and detailed, particularly f o r those of t h e C r e e d e district (Bethke e t al., 1973; Bethke and Rye, 1979; Foley et al., 1982; Hayba et al., 1985, this volume) and t h e Sunnyside mine of t h e San Juan Mountains, Colorado (Casadevall and Ohmoto, 1977); Tui, New Zealand (Robinson, 1974); Casapalca, Peru (Rye and Sawkins, 1974); Finlandia vein, Colqui district, P e r u (Kamilli and Ohmoto, 1977); mining districts of t h e Western Cascades, Oregon (Taylor, 1971, 1974b; Power, 1985; Field and Power, 1985); and Golden Sunlight mine, Montana (Porter and R,jfley, 1985). Sulfides i n most of these deposits have 6 S values e a r 0 permil t h a t contrast markedly with hypogene sulfates, such a s a t associated 3'S-enriched C r e e d e (-4.1 t o 1.7 O/oo versus 19.8 t o 45 O/oo), Sunnyside (-6.3 t o 2.7 O/oo versus 15.3 t o 22.9 O/oo), Tui (-2.4 t o 4.9 O/oo versus 16.0 t o 19.5 O/oo) and Finlandia (-4.0 t o 1.5 O/oo versus 14.0 t o 14.1 dloo). Moreover, compositions of sulfides t h a t a r e unassociated with sulfates, such a s those given by Jensen et al. (1960) for Rico (-0.6 t o 4.0 O/oo) and Ouray (-2.0 t o 1.9 O/oo) in t h e San Juan Mountains of Colorado, and by R y e and Sawkins (1974) for Casapalca, a r e also closely grouped near 0 permil. Because of t h e overall isotopic similarity of most sulfides in t h e s e deposits t o t h e 0 permil value of m a g m a t i c sulfur, and other geologic and geochemical considerations, a m a g m a t i c source of sulfur is advocated by t h e authors of investigations a t Rico, Ouray, Casapalca, and t h e Western Cascades, and possibly a t C r e e d e and Finlandia. However, the commonly assumed "genetic" equivalence of 0 permil sulfides and m a g m a t i c (0 '100) sources of sulfur is a hazardous generalization when applied t o hydrothermal environments. It is valid for sulfides only t o t h e e x t e n t i t m a y b e assumed t h a t concentrations of reduced sulfur (H2S) a r e approximately equal t o those of t o t a l sulfur in these systems. However, under conditions of high fO2 and (or) low pH, a s may be inferred from t h e presence of oxide and sulfate minerals in vein and alteration assemblages, part of t h e H2S becomes oxidized t o SO;, and t o t a l sulfur must then consist of both o x i d ~ z e d and redu components. Because of t h e large fractionation of between sulfates and sulfides and of mass balance considerations between oxidized and reduced forms of aqueous sulfur in t h e fluids, sulfide minerals b c o m e increasingly depleted in 3 4 ~ relative t o t h e 6f 4S of t o t a l sulfur in t h e system, a s ratios of SO5:H S increase with increasing f 2 and (or) decreasing Redox changes a s d e s c r i b e 2 such a s increasing s t a t e s of oxidation with evolution of hydrothermal systems,

have been proposed by Robinson (1974) t o account for t h e progressive 3 4 ~ depletion observed in t h e paragenetic order of sulfide depos' ion a t Tui. This socalled Eh-pH control of mineral "S compositions was first defined by Sakai (1968) and Ohmoto (1972), and has been subsequently refined by R y e and Ohmoto (1974) and Ohmoto and R y e (1979). It is for reasons of this Eh-pH control, a s deduced from mineralogical and geochemical evidence, t h a t isotopically heavy marine sulfates have been proposed a s sources of sulfur in t h e hydrothermal deposits a t Sunnyside (upper Paleozoic evaporites of % I 2 O/oo; Casadevall and Ohmoto, 1977) and Tui (Jurassic s e a w a t e r of ~ 1 '1 600; Robinson, 19741, and in spite of t h e f a c t t h a t compositions of associated sulfides (-6.3 t o 2.7 O/oo and -2.4 t o 4.9 O/oo, respectively) bracket t h a t of 0 O/oo magmatic sulfur. Moreover, this control is implicit t o t h e interpretation of a m a g m a t i c source of sulfur a t t h e Golden Sunlight deposit (Porter and Ripley, 19851, although both sulfates (1.2 t o 5.9 O/oo) and sulfi e s (-15.8 t o -4.0 O/oo) a r e appreciably depleted in '4S relative t o t h e majority of d a t a for hypogene equivalents (Fig. 6.12; also s e e Field and Gustafson, 1976, Fig. 3, for a graphical portrayal of this control). In contrast, Gross (1975) has proposed a crustal source f o r sulfur and metals contained in vein sulfides (-19.5 t o -3.4 O/oo) of pyrite-sphaleritegalena-argentite o r e from Guanajuato, Mexico, which a r e considered t o have been derived from metal-rich and sulfide-bearing (-16.6 t o 6.3 O/oo) Mesozoic sedimentary country rocks by heated ground waters during Oligocene volcanic activity. Fractionation trends f o r sulfate-sulfide and sulfide-sulfide mineral pairs from t h e zoned polymetallic veins a r e largely consistent with those predicted from theory and experiment (Tables 6.4 and 6.5, and Figs. 6.4 and 6.5). However, with t h e exception of a few sphaleritegalena pairs, from Sunnyside, Tui, Casapalca, Finlandia, and t h e Western Cascades, t h e 3 4 ~ t e m p e r a t u r e e s t i m a t e s a r e rarely consistent with those obtained by fluid-inclusion homogenization methods. This disparity implies t h a t isotopic equilibrium largely did not prevail in these systems for reasons t h a t may include low depositional temperatures, relatively rapid a c c e n t of fluids and deposition of minerals, and possibly abrupt changes in SOG:H2S ratios with t h e ascent of fluids into more oxidizing environments. The lack of equilibrium between sulfates and sulfides is clearly expectable based on t h e discussions and d a t a presented by Ohmoto and R y e (1979) and Ohrnoto and Lasaga (1982). Moroever, biogenic processes (Hayba e t al., 1985, this volume), a s previously su gested by Kesler e t al. (1981) t o account for t h e J'S-depleted compositions of bedded sulfides a t Pueblo Viejo, in addition t o t h e e f f e c t s of chemical and (or) isotopic disequilibrium (Bethke et al., 1973) a r e not considered ~ (19 t o t o b e responsible for t h e large 3 4 enrichments 45 'loo) of sulfates a t Creede, Colorado. Hydrogen and Oxygen Distributions of ''0 in t h e host rocks and minerals of many geothermal and epithermal occurrences previously described a r e portrayed in

Figure 6.13. Because t h e s e d a t a a r e voluminous and based on numerous investigations of variable detail, our discussion will focus principally on t h e major isotopic trends and reasons thereof. More detailed information may b e obtained a s needed from t h e references cited. Although t h e s e d a t a exhibit a large

Geothermal Systems Geysers:

steam whole rock quartz calcite salton Sea/Cerro Prieto: carbonate host altered carb. host quartz carbonates Broadlanddwairakei: volcanic host altered volc. host quartz sinter calcite adularia Sediment-Hosted Cortez: carbonate host altered carb. host quartz calcite hematite Carlin: carbonate host altered carb. host sedimentary chert jaeperoid iuaitz barite calcite Others: adularia Vol~anic-Hosted Tolfa: volcanic host altered volc. host chalcedony Comstock: altered volcanics quartz Goldfield: altered volcanics quartz Tonopah: volcanic host altered volcanic host quartz calcite K-feldspar Bodie: volcanic host altered volcanic host quartz K-feldspar calcite quartz Others: adularia

Zoned Polymetallic Veins San Juan Mountains: altered plutonics altered volcanics Creede: quartz carbonates illite chlorite Sunnyside: quartz carbonates Tui: carbonates barite Casapalca: quartz calcite Finlandia: quartz Yankee Fork: intermediate volcanics ash-flow tuffs rhyolitic intrusions vein quartz Western Cascades: volcanics altered volcanics altered intrisions Golden Sunlight: clastlc seds. altered host breccias quartz seiicite

Figure 6.13. Distributions of 6 180 in epitherm a 1 deposits.

amount of compositional s c a t t e r , they reveal a number of s y s t e m a t i c trends when examined on t h e basis of individual occurrences. T e large isotopic variability results from (I) differing lk0 compositions of t h e host compositions of t h e rocks, (2) differing ''0 hydrothermal fluids, and (3) interactions between t h e host rocks and fluids over a range of low temperatures ( 1 0 0 ~t o ~3 0 0 ' ~ ) and a t t a i n m e n t s of equilibrium t h a t produced hydrotherma minerals of variable and generally intermediate "0 compositions. As previously noted, most of t h e common marine sedimentary ocks a r e isoto ically m o r e variable and enriched in ''0 (=LO t o 40 loo) than a r e t h e common igneous rocks ('5 t o 12 O/oo; s e e Fig. 6.7). These petrologic distinctions accou for t h e obvious isotopic differences between t h e '0-enriched "unalteredQ' sedimentary host rocks of t h e Geysers and Salton Sea geothermal areas and t h e C o r t e z and C lin epithermal deposits, a s compared t o t h e less "0enriched "unaltered" volcanic host rocks of nearly all o t h e r epithermal deposits. The derivation of nearly all geothermal and epithermal fluids from m e t e o r i c sources of w a t e r has also been mentioned reviously. This conclusion is compositions of these fluids, based on t h e D and which although variable plot in close proximity t o or by lateral "oxygen isotope shift" away from t e m e t e o r i c water line (Fig. 6.8). The 6D and 61 60 compositions of hydrothermal fluids t h a t formed t h e epithermal deposits of this discussion and those for many of their nearby present-day m e t e o r i c w a t e r s a r e listed in Table 6.6. Sources of d a t a for most deposits of t h e Basin and Range a r e from O'Neil and Silberman (19741, e x c e p t for those of Carlin (Radtke e t al., 1980) and C o r t e z (Rye et al., 1974). D a t a for epithermal deposits of other geographic localities include Tolfa (Lombardi and Sheppard, 19771, Yankee Fork-Idaho Batholith (Criss and Taylor, 1983; Criss e t al., 19851, and others previously mentioned. On t h e basis of isotopic similarities t o pgesent-day m e t e o r i c waters and of widespread D and 0 depletions, t h e sources of most epithermal fluids must have been local m e t e o r i c w a t e r s (Table 6.6). Although magrnatic w a t e r s may constitute a small proportion of t h e s e hydrothermal fluids, t h e y have been d e t e c t e d with confidence from D analyses of fluid inclusions only a t Casapalca (Rye and Sawkins, 19741, for t h e early and intermediate s t a g e s of carbonate mineralization a t C r e e d e (Bethke and Rye, 19791, and a t Golden Sunlight (Porter and Ripley, 1985). Based on isotopic and other considerations, t h e fluids of Tui (Robinson, 19741, Finlandia (Kamilli and Ohmoto, 19771, and possibly t h e Comstock Lode (Taylor, 1973) may be partly of m a g m a t i c origin. Thus, the compositions of hydrothermal fluids depicted by closed triangles on Figure 6.8, with t h e exception of those of Casapalca and Golden Sunlight, a r e illustrative of m e t e o r i c water-dominated epithermal systems. Compositionally complex fluids such a s those of C r e e d e (Bethke and Rye, 19791, Sunnyside (Casadevall and Ohmoto, 19771, Finlandia (Kamilli and Ohmoto, 1977) and p a r t s of several other hydrothermal systems a r e not portrayed in Figure 6.8. Fluids f r o r 8 t h e s e deposits have exceedingly variable 6D and 6 0 values (Figs. 6.6 t o

8

'6

6.8) and thus enclose large isotopic domains a s summarized by Taylor (1979). This variability is attributable either t o paragenetically distinct and (or) mixed sources of m a g m a t i c and m e t e o r i c waters, a s deduced from t h e analyses of many samples, or t o subsequent contamination by deuterium-depleted ground waters t h a t were trapped in pseudosecondary inclusions (Foley e t al., 1982). Distributions of t h e isotopic d a t a (Fig. 6.13) show t h a t a l t e r e d host rocks, regardless of igneous or sedimentary parentage, a r e variably depleted in ''0 relative t o their unaltered counterparts. hese trends a r e evident from comparisons of t h e "0 d a t a for unaltered and a l t e r e d host rocks a t Salton Sea-Cerro Prieto (Clayton e t al., 1968; Williams and Elders, 19841, Broadlands-Wairakei (Blattner, 1975; Clayton and Steiner, 19751, C o r t e z (Rye e t al., 19741, Carlin (Radtke et al., 19801, Tolfa (Taylor and Turi, 1976; Lombardi and Sheppard, 19771, Tonopah (Taylor, 1973), Bodie (OQNeil e t al., 1973), San Juan Mountains, Colorado (Taylor, 1974b; L e a e t al., 1984), Western Cascades, Oregon (Taylor, 1971; 1974a), and Golden Sunlight (Porter and Ripley, 1985). Although comparisons t o unaltered equivalents a r e lacking, a l t e r e d sedimentary host rocks a t t h e Geysers (Sternfeld, 1981) and volcanic host rocks a t t h e Comstock Lode and Goldfield (Taylor, 1973; O'Neil and Silberman, 1974), and a t Yankee Fork (Criss and Tay r, 1983; Criss et al., 1985) a r e similarly depleted in "0 relative t o normal m a g m a t i c compositions. These ''0 depletions of volcanic and sedimentary host rocks a r e a corollary of t h e "oxygen isotope shift" previously noted f o r many geothermal fluids (Fig. 6.8). They arise during geothermal-hydrothermal a c t ' i t y from isotope-exchange reactions between "0depleted meteori waters and "0-enriched host r o s and result in theC1'O enrichment of t h e fluids and depletion of t h e rocks a s a consequence of ter-rock ' with reactions. Progressive depletions of 0 increasing depth in host rocks of t h e C e r r o Prieto (Williams and Elders, 19841, Geysers (Sternfeld, 1981, and references therein), and Salton S e a (Clayton e t al., 1968) a r e a s cannot b e related t o temperature-induced fractionation trends, and thus m u s t be caused by changes in fluid compositions with depth t h a t result f r o m recharge and (or) water-rock reactions. Isotopic trends and permutations resulting from these phenomena range from subtle t o dramatic, and they will b e discussed extensively in t h e section t h a t follows. Compositions of the fracture-controlled hydrothermal gangue minerals from geothermal systems and epithermal deposits (Fig. 6.13) a r e largely compatible with isotopic e f f e c t s described in t h e foregoing discussion and from previous considerations of fractionation. The d a t a a r e from sources previously cited, except for t h e "other" categories of sedimentand volcanic-hosted deposits t h a t a r e from O'Neil and Silberman (1974; also s e e Table 6.6) and represent miscellaneous mines and prosp t s from t h e Basin and Range province. Values of 6 0 for these minerals generally occupy a range t h a t is intermediate between t h e compositions of associated altered-unaltered host rocks (Fig. 6.13) and those of t h e geothermal or

'$6

fS

Table 6.6 Compositions of 6D and 6180 in hydrothermal fluids of Tertiary epithermal deposits and some local meteoric waters Epithermal F1yQds 6D 6 0

Meteoric W

Basin and Range Volcanic-Hosted Deposits Bullfrog (BU) Aurora (A) Trade Dollar (TD) Wonder (W) Jarbidge (J) Rawhide ( R ) Gilbert(G) Tonopah (T) Bodie (B) Cornstock Lode (CL) Sediment-Hosted Deposits Tenmile (TE) Humboldt (H) Cortez (CO) Carlin (CA) Manhattan (M) Elsewhere Volcanic-Hosted Deposits Tolfa (TF) Zoned Polymetallic Vein Creede Sunnyside Tui (TU) Casapalca Finlandia Yankee Fork (Idaho Batholith, IB) Western Cascades (WC) Golden Sunlight (GS)

complex complex 0 7 complex -1 6

*calculated from equation for the meteoric water line (eq. 12)

hydrothermal fluids (Table 6.6 and Fig. 6.8). The 6180 values of hydrothermal minerals formed by mineralH 2 0 exchange reactions a r e determined by t h e t e m p e r a t u r e of deposition, which controls t h e e x t e n t of isotopic fractionation, and fluid compositions, which in turn a r e controlled by t h e source of fluids, types of host rock, and water-host rock reactions. Isotopic e f f e c t s related t o one or m o r e of these determinants a r e qualitatively apparent from t h e d a t a of Figure 6.13. The t e m p e r a t u r e control, with mineral-H20 fractionation increasing with decreasing t e m p e r a t u r e ,

must account f o r t h e l a r g e 1 8 0 enrichments of siliceous hot-spring sinters (22.2 and 23.6 O/oo) relative t o vein q u a r t z (mostly 3.9 t o 12.3 O/oo) formed a t higher t e m p e r a t u r e s and depths in the Broadlands- Wairakei a r e a s (Clayton and Steiner, 1975; Blattner, 1975). Moreover, this t e m p e r a t u r e contro of fractionation is largely t h e cause of progressive j80 depletions of vein q u a r t z and (or) c a l c i t e with increasing depth a t Broadlands (Blattner, 1975)) C e r r o P r i e t o (Williams and Elders, 19841, Geysers (Sternfeld, 1981), Salton Sea (Clayton e t al., 1968)) and Wairakei

(Clayton and Steiner, b975). However, t h e parallel trend of decreasing d. 0 values with depth in host rocks of several geothermal areas, a s previously noted, cannot b e a t t r i b u t e d t o t e m p e r a t u r e effects, but instead must result from changes in fluid composition. In addition, compositions of t h e gangue minerals appear t o b e influenced by those of their host rocks. For example, hydrothermal c a l c i t e and q u a r t z associated with "0-enriched sedimentary host rocks of Carlin and C o r t e z a r e largely isotopiffly heavier than their counterparts i n relatively 0-depleted volcanic host rocks of t h e other epithermal deposits; particularly in t h e Basin and Range province where compositions of t h e m e t e o r i c waters a r e l e a s t variable (Table 6.6). T h e isotopic composition of hydrothermal fluids, a s partly determined by source, may also leave i t s imprint o n t h e minerals. Quartz and c a l c i t e magmatic waters deposited from ''0-enriched (Casapalca, Golden Sunlight, and possibly carbonates a t Creede) or from relatively undepleted m e t e o r i c w a t e r s (Tolfa and Tui) a r e isotopically heavier t h a n their equivalents formed in hydrothermal systems having more depleted fluids of m e t e o r i c origin such as those of t h e Basin and Range (see T le 6.6 and Fig. 6.13). The sequence of relative enrichments among t h e vein minerals ( q t z > c a l > b a r > K f ) is generally consistent with that obtained from experimentally derived fractionation d a t a (Fig. 6.3). However, analyses for coexisting mineral pairs and triplets a r e conspicuously few, and they a r e largely suggestive of isotopic disequilibrium. This result is not surprising because of t h e relatively low t e m p e r a t u r e s (IOO°C t o 300°c), changes in fluid chemistry, and varied sequences of mineral paragenesis t h a t may c h a r a c t e r i z e geothermal-epithermal systems, and which collectively render isotopic equilibrium unlikely. Investigations of fluid-mineral isotopic equilibria in geothermal systems by Clayton e t al. (19681, Blattner (19751, and Clayton and Steiner (1975) have demonstrated t h a t q u a r t z is most resistent t o isotopic exchange, whereas c a l c i t e and alkali feldspars m a y rapidly undergo re-equilibration and thus be susceptible t o compositional change during postdepositional s t a g e s of hydrother al activity. ' in host rocks and The principal trends for 0 hydrothermal gangue minerals associated with epithermal activity a r e summarized by t h e composite illustration given in Figure 6.14 (after Taylor, 1971; 1973). Volcanic rocks of t h e Western Cascades in Oregon host numerous zoned polymetallic vein deposits of t h e base and precious metals. These deposits a r e mostly clustered within larger district-sized a r e a s of hydrothermal alteration (Field and Power, 1985) t h a t a r e cored by small granodiorite intrusions of Tertiary age. According t o Taylor (19711, t h e volcanic country rocks (origin l y 5.5 t o 8 O/oo) a r e progressively . . depleted in 0 wlth Increasing proximity t o t h e intrusions (Fig. 6.14A; a f t e r Taylor, 19711, a s a consequence of hydrothermal alteration imposed by reactions between heated m e t e o r i c ground w a t e r s ("-9 O/oo, Table 6.6) and t h e vol nic host rocks (now ' depletion, which - 5 . 5 t o 5.5 O / o o . This trend of 0 increases with intensity of alteration, is analogous t o t h e "oxygen isotope shift" of geothermal waters, but

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opposite in direction. Volcanic host rocks of t h e Tonopah district exhibit similar depletions of ''0 with progressive alteration (Fig. 6.148; a f t e r Taylor, 1973), a n d evidence for this having been a m e t e o r i c waterdominated hydrothermal syste is additionally depletions (= -7 t o s t r ngthened by t h e pronounced I ~ ~ o ofo associated ) quartz, calcite, and adularia vein minerals. The a r e a l distribution of 1 8 ~ - d e p l e t e d country rocks a r e large a t Bohemia in t h e Western Cascades (Taylor, 19711, Tonopah (Taylor, 19731, and a t Yankee Fork and other hydrothermally a l t e r e d a r e a s of t h e Idaho Batholith (Criss and Taylor, 1983; Criss et al., 1985), and these "negative" isotopic anomalies form well-defined t a r g e t s appropriate t o t h e reconnaissance s t a g e of mineral exploration. Waterxock ratios--In this section we examine t h e systematics of oxygen- and hydrogen-isotopic exchange between fluid and rock and present a n

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8d8 (%.I

Figure 6.14. Variations of 6 180 i n (A) volcanic host rocks adjacent t o T e r t i a r y intrusions i n t h e W e s t e r n Cascades o f Oregon ( a f t e r Taylor, 1971, p. 7867, Fig. 8) and ( B ) volcanic host rocks and m a g m a t i c and hydrothermal minerals i n t h e Tonopah d i s t r i c t of Nevada ( a f t e r Taylor, 1973, p. 755, Fig. 6)-

application of these principles t o epithermal systems. The application serves not only t o impose constraints on t h e hydrothermal environment, but also a s a n illustration of t h e assumptions, problems, and uncertainties t h a t m a y be involved in modeling isotopic data. The development of t h e principles of isotopic exchange in water-rock systems generally follows t h a t of Ohmoto and R y e (1974). The final isotopic composition of water (6 a f t e r equilibration with rock is a function of: (a) t h e initial (unexchanged) composition of t h e water (6 &) and rock (6: (b) t h e temperature of equilibration, which determines t h e fractionation f a c t o r between rock and water (Ar-,), and (c) t h e ratio of exchanged oxygen and hydrogen a t o m s in t h e water t o those in rock (wit-). This relationship, a f t e r Ohmoto and R y e (1974), is expressed a s

!,

Epithermal precious-metal deposits are commonly hosted either by volcanic rocks of i n t e r m e d i a t e t o felsic composition or by c l a s t i c or chemically precipitated sedimentary rocks. Unaltered andesites, dacites, and rhyolites typically have 6180 and 6D values of about 7 '100 (Fig. 6.7) and -70 '100, respectively. Rocks of this compositional range contain approximately 50 weight-percent oxygen and a s much a s 0.11 weight-percent hydrogen (1 wt.-% H20), according t o t h e average analyses reported by Thus, t h e oxygen- and Nockoids et al. (1978). hydrogen-isotopic composition of a fluid t h a t has equilibrated with volcanic rock of these characteristics can be determined from

and

where t h e coefficients 1.8 and 100 represent ratios of t h e weight-percent oxygen in water (88.8%) t o t h a t in rock (50%) and t h e weight-percent hydrogen in water (1 1.2%) t o t h a t in rock (0.112%), respectively, and R is Because values of R t h e water:rock mass ratio. represent proportions of water and rock t h a t have isotopically equilibrated, they a r e easier t o r e l a t e t o natural systems than a r e values of t h e a t o m i c ratio w/r. Isotopic compositions of sedimentary rocks a r e more variable than those of igneous rocks (Fig. 6.7). However, t h e stratigraphic sequences t h a t host epithermal deposits typically consist of silty argillaceous limestone, dolomite, quartzite, and minor shale. If a "typical" host sequence contains subequal amounts of siliciclastic and carbonate com then i t would average approximately 16 '100 6 0 and

wents,

-60 '100 6D and contain about 50 weight-percent oxygen and perhaps 0.28 weight-percent hydrogen (2.5 wt-% H2.0). The final oxygen and hydrogen isotopic c o m p o s i t ~ o n of a fluid t h a t has equilibrated with a sedimentary rock of this composition c a n b e calculated from

and

The fractionation of oxygen isotopes between fluids and rock has been variously assumed t o be similar t o those of smectite-H 0 (Cathles, 1983), plagioclase feldspar (An 0)-H20 ( f a y l o r , 1974a, 1979; Ohmoto and Rye, 1972; Green e t al., 1983) and muscovite-H20 (Spooner e t al., 1977). The assumptions a r e based on comparisons of these mineral-H20 fractionations with experimental rockH 2 0 fractionations and with t h e oxygen-isotopic compositions of naturally a l t e r e d rocks and secondary minerals. For t h e illustrations t h a t follow, we have used fractionation factors derived from t h e ~ l a ~ i o c l a s e feldspar (An30)-H20 equation of OtNeil andu Taylor (1967) I000 I n a

= 2.68 ( 1 0 6 / ~ * )

-

3.29

(18)

Use of this equation r a t h e r than t h e more r e c e n t one of Matsuhisa e t al., 1979 (from eqs. 7 and 10 in Table 6.3) was done t o maintain continuity between our results and those of other investigators (see Taylor, 1979, and above). These fractionations a r e intermediate between those derived from the muscovite-H20 and smectite-H 0 curves over t h e t e m p e r a t u r e range from 1 0 0 ~ 8 t o 3 0 0 ~ ~ .The fractionation of hydrogen isotopes between fluids and rock is similarly assumed t o b e equivalent t o t h a t of Our biotite or chlorite-H20 (Taylor, 1974a). computations a r e based on chlorite-H 0 fractionation f a c t o r s which were taken from Taylor f1979, Fig. 6.2). Using t h e equations, rock compositions, and mineral-H20 systems described above, t h final isotopic composition of m e t e o r i c water ( 678 1 --1 '/PO and 6% = -120 .O/OO) and m a g m a t i c waters ( = 7.3 '/oo and 6% = -60 '/oo) were computed a t t e m p e r a t u r e s of 1 0 0 t~o ~300°C and R values of 0.01 t o 10. The initial isotopic composition of t h e m e t e o r i c water is typical of present-day m e t e o r i c w a t e r in western Nevada (Table 6.6) and t h e selected t e m p e r a t u r e s and R values a r e considered relevant t o epithermal systems. The results of t h e com putations, portrayed in Figure 6.15, illustrate t h e s y s t e m a t i c s of isotopic exchange between m e t e o r i c or m a g m a t i c waters and rocks in the Basin and ange province. ' and D under First, m e t e o r i c water is enriched in 0 most geologic conditions, whereas m a g m a t i c w a t e r is enriched in deuterium but may b e either depleted

Jgq:

in ''0 thro h exchange with igneous rocks or 0 via exchange with sedimentary rocks. enriched in " Second, the magnitude of t h e isotopic enrichment or depletion varies inversely with t h e water:rock mass ratio. Third, the change in fluid 6180 values exceeds t h a t of 6D values a t high water:rock mass ratios (?,I) but t h e reverse is t r u e a t low ratios ($1). This e f f e c t results from t h e small H:O mass r a t i o in rocks relative t o t h a t in water and from t h e generally larger fractionation f a c t o r s for hydrogen than for oxygen. Accordingly, t h e isotopic influence of rock hydrogen is significant only a t low water:rock mass ratios, but is A potentially larger than t h a t of rock oxygen. c ollary t o this e f f e c t is t h a t t h e generally small '"0 shiftf' observed for geothermal waters (Fig. 6.8) must be a product of isotopic exchange a t relatively high water:rock mass ratios. Fourth, a t equivalent t e m p e r a t u r e s and w a t e r x o c k mass ratios, fluids t h a t have equilibrated with sedimentary rock a r e isotopically heavier than those t h a t have equilibrated wi igneous rock because t h e former is enriched in"0 and D and contains a g r e a t e r quantity of hydrogen relative t o t h e latter. The calculated and analyzed isotopic compositions of t h e hydrothermal fluids associated with epithermal deposits f a l l in t h e range -15 t o 5 O/oo 6180 and -150 t o -90 O/oo 6D (Table 6.6 and Fig. 6.8). Consequently, these fluids a r e isotopically depleted relative t o m a g m a t i c w a t e r or evolved m a g m a t i c water, but a r e similar t o m e t e o r i c water and i t s evolved counterparts (Fig. 6.15). Because i t is unlikely t h a t magmatic fluid will be isotopically depleted in deuterium through equilibration with rock (most mineral-H20 fractionation f a c t o r s a r e negative f o r hydrogen), i t is concluded t h a t m e t e o r i c w a t e r is t h e predominant fluid in epithermal systems. However, if fluid mixing is commonplace, then a minor component of magmatic water cannot be precluded o n t h e basis of isotopic considerations. The isotopic composition of a convected fluid may differ from t h a t of a s t a t i c fluid because, among other factors, the convected fluid has equilibrated over a range of t e m p e r a t u r e s and water:rock mass ratios. Therefore, t o model the evolution of western Nevada m e t e o r i c w a t e r during convection, t h e isotopic composition was calculated a t 20°C intervals from l0fI0C t o 300°C. The final composition of t h e water (6,) computed a t e a c h t e m p e r a t u r e b e c a m e t h e initial composition of t h e water (6;) for t h e calculations a t t h e next higher temperature. Since w a t e r x o c k mass ratios may vary according t o t h e convection path, t h e series of calculations were performed a t R values of 0.01, 0.1, 1, and 10. The results a r e presented in Figure 6.16 and compared t o compositions of t h e fluids responsible f o r those volcanic-hosted (Fig. 6.16A) and sediment-hosted (Fig. 6.16B) epithermal deposits listed in Table 6.6. The computations imply t h a t c o vecting m e t e o r i c w a t e r is progressively enriched in "0 a n d D during heating and exchange with unaltered wall rock. This w a t e r will be isotopically heavier t h a n non-convecting m e t e o r i c water (e.g. pore fluid), under identical conditions of equilibration, because of i t s "history" of exchange (compare Figs. 6.15 and 6.16). Moreover,

Magmatic

3000~ Water

Figure 6.15. Variations of 6~ and 6 180 in fluids that equilibrate with (A) volcanic and (B). sedimentary rock as a function of the initial fluid composition, temperature, and water: rock ratio.

isotopically light m e t e o r i c water may a t t a i n compositions similar t o those of magmatic or evolved magmatic w a t e r through exchange a t low water:rock mass ratios (L0.1). Therefore, t h e distinction between a m a g m a t i c or m e t e o r i c source of fluids and t h e demonstration of mixing between magmatic and m e t e o r i c waters cannot b e made solely on t h e basis of t h e calculated composition of a hydrothermal fluid. The oxygen-isotopic composition of meteoric w a t e r in t h e Basin a n d Range province has remained essentially unchanged since t h e Early Tertiary, whereas t h e hydrogen-isoto i c composition may have decreased slightly (10-20 gloo) in response t o t h e climatic cooling (Sheppard e t al., 1969; Taylor, 1973). Thus, t h e isotopic composition of m e t e o r i c water in this region may b e regarded a s a n approximation of t h e initial composition of Tertiary m e t e o r i c hydrothermal fluids. Meteoric w a t e r compositions near t h e Aurora, Rawhide, Gilbert, Tenmile, Humboldt, C o r t e z and Manhattan epithermal deposits have 6D values between Accordingly, t h e -130 and -120 O/oo (Table 6.6). isotopic compositions of t h e hydrothermal fluids for

these deposits can b e interpreted in t e r m s of t h e calculated curves displayed in Figure 6.16. Isotopic d a t a for t h e remaining deposits in t h e Basin and Range province (Table 6.6) a r e more appropriately compared t o calculated curves t h a t have been shifted, relative t o those in Figure 6.16, in t h e direction a n d t o t h e e x t e n t t h a t t h e initial compositions of t h e fluids differed 0 -120 O/oo 6D. Assuming t h a t from -16 0 / o o 6 ~ ~and temperatures of water-rock equilibration were typically between 150°C and 3 0 0 ' ~ (hachured fields in Fig. 6.16), a s indicated by fluid-inclusion studies, then most of t h e epithermal systems were characterized by However, t h e high water:rock mass ratios (>0.5). model calculations imply t h a t t h e Bodie, Tonopah, and Tenmile fluids evolved a t unusually low water:rock mass ratios (W.01-0.2) and t e m p e r a t u r e s (510O0C). Furthermore, t h e model cannot account for t h e apparent decrease in t h e 6D composition of t h e Carlin and Comstock Lode fluids. Most of t h e water:rock ratios determined above should b e regarded a s t e n t a t i v e (particularly t h e anomalously low values) because processes other t h a n fluid-rock exchange may have influenced t h e isotopic

Variations of CD and 6180 in Figure 6.16. fluids during convection through (A) volcanic and (B) sedimentary rock as a function of temperature and waterzrock ratio.

composition of t h e fluids. F o r example, boiling in t h e convectio system would enrich t h e evolved meteoric fluid in 0 ' and D along a trend subparallel t o t h e MWL (Truesdell et al., 1977; s e e R a d t k e et al., 1980). Unexchanged m e t e o r i c w a t e r m a y become involved a t t h e s i t e of mineralization through mixing or with t h e complete equilibration (alteration) of t h e host rocks. This e f f e c t ould drive t h e fluid composition towards t h e initial 6150 and dD values. Inaccurate water:rock ratios may also result if t h e isotopic composition of modern m e t e o r i c w a t e r is used a s a n approximation of Tertiary m e t e o r i c w a t e r in a r e a s t h a t have undergone considerable uplift o r subsidence. Similarly, 6 D analyses of e x t r a c t e d inclusion fluids t h a t contained a significant fraction of secondary fluids could lead t o erroneous conclusions (see Foley e t al., 1982). The final isotopic composition of t h e rock (6 f) t h a t equilibrated with t h e m e t e o r i c fluid a t e a c h interval in t h e convection model was calculated from t h e relationship

and portrayed a s curves in Figure 6.17. These results ind' a t e t h a t rocks become progressively depleted in '$0 and D with a n in e a s e in t h e water:rock mass ratio+ The depletion in "0 is negligible a t low ratios because of t h e overwhelming abundance of rock oxygen. With increasing temperature, both t h e 6180 and 6D values increase a t water:rock mass r tios of 1 or less, whereas a t a r a t i o of 10 t h e 6 1'80 value decreases and t h e 6D value increases. Decreases in t h e 6180 values of hydrothermally a l t e r e d rocks n e a r c e n t e r s of hydrothermal activity have been documented for t h e Tonopah epithermal Au-Ag deposit (Taylor, 19731, zoned polymetallic veins of t h e Bohemia district in t h e Western Cascades (Taylor, 19711, Yankee Fork and other a l t e r e d and mineralized localities of t h e Idaho Batholith (Criss and Taylor, 1983; Criss e t al., 1985), and several volcanogenic massive sulfide depo i t s (see Franklin e t al., 1981). The whole rock 61 20 compositions a t Tonopah range from values typical of unaltered igneous rocks (5.5 t o 10.0 O/oo) t o nearly -6 permil and whole rock 6D compositions range from -150 t o -135 permil. The field represent72 by six Tonopah samples for which both 6D and 6 0 analyses have been reported is shown in Figure 6.17. Compositions of these samples a r e most consistent with isotopic exchange a t high water:rock mass ratios ( > I ) , although t h e 6 D values a r e somewhat larger t h a n predicted. On t h e basis of whole-rock 1 8 0 depletions observed a t Tonopah and elsewhere, water:rock mass ratios f o r many epithermal districts throughout t h e western U.S. range from about 0.2 t o 2 (Taylor, 1974a) and t h e s e ratios a r e broadly similar t o those of geothermal systems t h a t range from about 0.15 a t t h e Geysers (Sternfeld, 1981) through 0.45 and 1.3 a t Salton Sea (Clayton et al., 1968) and C e r r o Prieto (Williams and Elders, 19841, t o a s large a s 4.3 a t Wair ei (Clayton and Steiner, 1975). The o u t e r haloes of depletion extend beyond t h e megascopically identifiable e f f e c t s of alteration and anomalies of t r a c e and minor 1983). e l e m e n t s (Taylor, 1971; Green e t al.,

%

y~. I

Figure 6.17. Variations of 6~ and 6180 in (A) volcanic and (B) sedimentary rocks that have equilibrated with convected meteoric water over a range of temperatures and water :rock ratios.

Accordingly, these cones of 1 8 0 depletion may b e potentially useful in t h e exploration for hidden mineral deposits (see Criss and Taylor, 1983; Criss e t al., 1985). Two main conclusions pertaining t o epithermal systems may be summarized from this discussion on t h e exchange of oxygen and hydrogen isotopes between fluids and rock. First, meteoric w a t e r is t h e predominant, and possibly only, source of fluid in most of t h e epithermal gold-silver deposits studied t o date. Second, a t reasonable t e m p e r a t u r e s of 1 5 0 ' ~ t o 300°c, t h e isotopic d a t a for hydrothermal fluids and their associated a l t e r e d rocks both imply t h a t high water:rock mass ratios prevailed during mineral deposition. Because typical porosities limit bulk water:rock mass ratios t o generally less than 0.1 (e.g. 10% porosity is equivalent t o a r a t i o of 0.04), then fluid-rock equilibration and mineral deposition must have occurred in open systems through which masses of fluid circulated t h a t were equivalent t o or larger than those of the rocks.

SUMMARY This overview of light-stable isotopes in the epithermal environment is prefaced by a review of t h e general principles of equilibrium isotope-exchange fractionation, including t h e equations and graphical portrayal of fractionation e f f e c t s b e t w e e com on isotop' compounds, and a summary of D,73C, '0, and "S distributions in geologically important habitats. The available d a t a for these isotopes in geothermal and epithermal systems a r e largely consistent with trends t h a t might be inferred from experimental-theoretical considerations and known distributions in common rock types and mineral deposits. Nonetheless, t h e epithermal d a t a exhibit isotopic characteristics t h a t a r e intermediate between those of near-surface geothermal systems a d deeper hydrothermal deposits. Depletions of D and ' 0 in t h e host rocks, minerals, and inclusion fluids of the epithermal deposits, and compositional similarities of these fluids t o present-day m e t e o r i c waters, suggest t h a t t h e hydrothermal fluids were predominantly ground waters of m e t e o r i c origin; although several deposits (Casapalca, Golden Sunlight, and possibly Creede) may have had varying amounts of a magmatic component. Sources of carbon in t h e carbonate minerals a r e not uniquely defined, and probably have originated from nearby sedimentary host rocks (Carlin and ~ o r t e z ) ,magmas, and possibly from biogenic and other provenances. Those of sulfur a r e considered t o have been derived from magmas or igneous rocks, sulfates ( S u n n ~ s i d eand Tolfa) and sulfides (Carlin, Cortez, and Guanajuato) of sedimentary rocks, and s e a water (Pueblo Viejo and Tui). Temperature-controlled fractionations over l a t e r a l or vertical gradients ranging from 300°C t o IOOOC outward and upward in hydrothermal systems should result in heavy-isotope 1 3 c for calcite, 14 '100 richments of about 12 O "0 for quartz, 12 Oioo for calcite, 26 O/oo 3 4 ~ for sulfates, and 1.7 O/oo 3 4 ~for pyrite. Although these "hypothetical" fractionation trends have not been reported from any of t h e epitherm 1 depos'ts ' and "0 studied t o date, they have been noted for C in c a l c i t e and quartz of geothermal systems (Broadlands, C e r r o Prieto, Geysers, Salton Sea, and Wairakei). Their absence from epithermal deposits may possibly be t h e result of inadequate sample representation, or of isotopic changes in fluid composition caused by boiling, redox and (or) waterrock reactions, disequilibrium, and different sources of t h e isotopic elements. Near-surface occurrences of supergene sulfates, which have formed by oxidation of lfides in t h e host ascending H2S or of preexisting relative t o t h e rocks, a r e markedly depleted in "S deeper hypogene sulfates. The a l t e r e d sedimentary and volcanic host rocks of both epithermal deposits and geothermal systems a r e conspicuously depleted in 1 8 0 relative t o lbheir peripheral and unaltered 0 depletions a r e a corollary of equivalents. These t h e well-known "oxygen isotope shift" of geothermal fluids, and they form in hydrothermal systems characterized by relatively high wa r:rock mass ratios (21). The resultant "negative" I g O anomalies may serve a s a useful guide t o mineral exploration because they a r e isotopically unambiguous and areally

C . W. FIELD& R. H.FIFAREK extensive. Isotopic investigations of epithermal deposits should b e continued, especially in conjunction with geologic and o t h e r topical studies, and particular emphasis should b e given t o s y s t e m a t i c and threedimensional sampling of host rocks beyond ore-bearing structures. REFERENCES Bachinski, D. J., 1969, Bond strength and sulfur isotope fractionation in coexisting sulfides: Economic Geology, v. 64, p. 56-65. Bender, M. L., 1972, Carbon isotope fractionation; b Fairbridge, R. W. (ed.), The Encyclopedia of Geochemistry and Environmental Sciences: Van Nostrand Reinhold Company, New York, p. 133-136. Bethke, P. M., Barton, P. B., and Rye, R. O., 1973, Hydrogen, oxygen, and sulfur isotopic compositions of o r e fluids in t h e C r e e d e district, Economic Mineral County, Colorado abs. : Geology, v. 68, p. 1205. Bethke, P. M., and Rye, R. O., 1979, Environment of o r e deposition in t h e C r e e d e mining district, San Juan Mountains, Colorado--Part IV, Source of fluids from oxygen, hydrogen, and carbon isotope studies: Economic Geology, v. 74, p. 1832-1851. Blattner, P., 1975, Oxygen isotopic composition of fissure-grown quartz, adularia, and c a l c i t e f r o m Broadlands geothermal field, New Zealand: American Journal of Science, v. 275, p. 785-800. Bottinga, Y., and Javoy, M., 1973, C o m m e n t s on oxygen isotope geothermometry: E a r t h and Planetary Science Letters, v. 20, p. 250-265. Brigham, R. H., 1984, K-feldspar genesis and s t a b l e isotope relations. of t h e Papoose F l a t Pluton, Inyo Mountains, California: Unpublished Ph.D. dissertation, Stanford University, 172 p. Browne, P. R. L., R a f t e r , T. A., and Robinson, B. W., 1975, Sulphur isotope ratios of sulphides from t h e Broadlands geothermal field, New Zealand: New Zealand Journal of Science, v. 18, p. 35-40. Casadevall, T., and Ohmoto, H., 1977, Sunnyside mine, Eureka mining district, San J u a n County, Colorado--Geochemistry of gold and base-metal ore deposition in a volcanic environment: Economic Geology, v. 72, p. 1285-1320. Cathles, L. M., 1983, An analysis of t h e hydrothermal system responsible for massive sulfide deposition in t h e Hokuroku Basin of Japan; Ohmoto, H., and Skinner, B. J. (eds.), The Kuroko and R e l a t e d Volcanogenic Massive Sulfide Deposits: Economic Geology, Monograph 5, p. 439-487. Claypool, G. E., Holser, W. T., Kaplan, I. R., Sakai, H., and Zak, I., 1980, The a g e curves of sulfur and oxygen isotopes in marine sulfate and their mutual interpretation: Chemical Geology, v. 28, p. 199-260. Clayton, R. N., Muffler, L. J. P., and White, D. E., 1968, Oxygen isotope study of c a l c i t e and silicates of t h e River Ranch No. 1 well, Salton S e a geothermal field, California: American Journal of Science, v. 266, p. 968-979.

125

Clayton, R. N., and Steiner, A., 1975, Oxygen isotope studies of t h e geothermal system a t Wairakei, Geochimica et Cosmochimica New Zealand: Acta, v. 39, p. 1179-1 186. Cortecci, G., Lombardi, G., Reyes, E., and Turi, B., 1981, A sulfur isotopic study of alunites from Latium and Tuscany, central Italy: Mineralium Deposita, v. 16, p. 147-156. Craig, H., 1953, The geochemistry of stable carbon isotopes: Geochimica e t Cosmochimica Acta, v. 3, p. 53-92. Craig, H., 1966, Isotopic composition and origin of t h e R e d S e a and Salton Sea geothermal brines: Science, v. 154, p. 1544-1548. Criss, R. E., Champion, D. E., and McIntyre, D. H., 1985, Oxygen isotope, aeromagnetic, and gravity anomalies associated with hydrothermally a l t e r e d zones in t h e Yankee Fork mining district, Custer County, Idaho: Economic Geology, v. 80, p. 1277-1296. Criss, R. E., and Taylor, H. P., Jr., 1983, An 180/160 and D/H study of Tertiary hydrothermal systems in t h e southern half of t h e Idaho Batholith: Geological Society of America Bulletin, v. 94, p. 640-663. Dean, J. A. (ed.), 1979, Lango's Handbook of Chemistry, Twelfth Edition, McGraw-Hill Book Company. Faure, G., 1977, Principles of Isotope Geology: John Wiley and Sons, New York, 464 p. Feely, H. W., and Kulp, J. L., 1957, The origin of Gulf C o a s t s a l t dome sulfur deposits: Bulletin of t h e American Association of Petroleum Geologists, v. 41, p. 1802-1853. Field, C. W., 1966, Sulfur isotopic method for discriminating between sulfates of hypogene and supergene origin: Economic Geology, v. 61, p. 1428-1435. Field, C. W., 1972, Isotope geology--stable; Fairbridge, R. W. (ed.), T h e Encyclopedia of Geochemistry and Environmental Sciences: Van Nostrand Reinhold Company, p. 618-622. Field, C. W., Dymond, J. R., Heath, G. R., Corliss, J. B., and Dasch, E. J., 1976, Sulfur isotope reconnaissance of epigenetic pyrite in ocean-floor basalts, Leg 34 and elsewhere; & Yeats, R. S., Hart, S. R. et al. (eds.), 1976: Initial R e p o r t s of t h e Deep S e a Drilling Project, v. XXXIV, Washington, D.C. (U.S. Government Printing Office), p. 381-384. Field, C. W., and Gustafson, L. B., 1976, Sulfur isotopes in t h e porphyry copper deposit a t El Salvador, Chile: Economic Geology, v. 71, p. 1533-1548. Field, C. W., and Lombardi, G., 1972, Sulfur isotopic evidence for t h e supergene origin of alunite deposits, Tolfa district, Italy: Mineralium Deposita, v. 7, p. 113-1 25. Field, C. W., and Power, S. G., 1985, Metallization in t h e Western Cascades, Oregon and southern Washington abs. : Geological Society of America, Abstracts with Programs, v. 17, no. 4, p. 218. Field, C. W., Rye, R. O., Dymond, J. R., Whelan, J. F., and Senechal, R. G., 1983, Metalliferous

CHAPTER 6

sediments of t h e E a s t Pacific; b Shanks, W. C., I11 (ed.), Cameron Volume on Unconventional Mineral Deposits: Society of Mining Engineers, New York, p. 133-156. Field, C. W., Sakai, H., a n d Ueda, A., 1984, Isotopic constraints on t h e origin of sulfur in oceanic igneous rocks; b Wauschkuhn, A., Kluth, C., and Zimmerman, R. A. (eds.), Syngenesis and Epigenesis in t h e Formation of Mineral Deposits: Springer-Verlag, Berlin-Heidelberg, p. 573-589. Fifarek, R. H., 1985, Alteration geochemistry, fluid inclusion, and s t a b l e isotope study of t h e Red Ledge volcanogenic massive sulfide deposit, Idaho: Unpublished Ph.D. dissertation, Oregon S t a t e University, 187 p. Foley, N. K., Bethke, P. M., and Rye, R. O., 1982, A reinterpretation of 6DH20 values of inclusion fluids in q u a r t z from shallow o r e bodies abs. : Geological Society of America, Abstracts with Programs, v. 14, no. 7, p. 489-490. Franklin, J. M., Sangster, D. M., and Lydon, J. W., 1981, Volcanic-associated massive sulfide deposits: Econornic Geology, Seventy-Fif t h Anniversary Volume, p. 485-627. Friedman, I., and O'Neil, J. R., 1977, Compilation of stable isotope fractionation factors of geochemical interest; & Fleischer, M. (ed.), D a t a of Geochemistry, Sixth Edition: U.S. Geological Survey, Professional Paper 440-KK, p. KK1-KK12. Fuex, A. N., and Baker, D. R., 1973, Stable carbon isotopes in s e l e c t e d granitic, mafic, and ultramafic rocks: Geochimica e t Cosmochimica Acta, v. 37, p. 2509-2521. Garlick, G. D., 1972, Oxygen isotope geochemistry; b Fairbridge, R. W. (ed.), The Encyclopedia of Geochemistry and Environmental Sciences: Van Nostrand Reinhold Company, New York, p. 864-874. Graham, C. M., Harmon, R. S., and Sheppard, S. M. F., 1984, Experimental hydrogen isotope studies-hydrogen isotope exchange between amphibole and water: American Mineralogist, v. 69, p. 128-1 38. Graham, C. M., Sheppard, S. M. F., and Heaton, T. H. E., 1980, Experimental hydrogen isotope studies--I. systematics of hydrogen isotope fractionation in t h e systems epidote-H20, zoisiteH 2 0 , and AlO(0H)-H20: Geochimica e t Cosmochimica Acta, v. 44, p. 353-364. Green, G. R., Ohmoto, H., Date, J., and Takahashi, T., 1983, Whole-rock oxygen isotope distribution in t h e Fukazawa-Kosaka area, Hokuroku district, Japan, and i t s potential application t o mineral exploration: Econornic Geology, Seventy-Fifth Anniversary Volume, p. 395-41 1. Gross, W. H., 1975, New o r e discovery and source of silver-gold veins, Guanajuato, Mexico: Economic Geology, v. 70, p. 1175-1189. Hayba, D. O., Bethke, P. M., Heald, P., and Foley, N. K., 1985, Geologic, mineralogic, and geochemical characteristics of volcanic-hosted epithermal precious-metal deposits; Berger, B. R., and Bethke, P. M. (eds.), Geology and Geochemistry of Epithermal Systems: Society of Economic Geologists, Reviews in Economic Geology, v. 2.

Hoefs, J., 1980, Stable Isotope Geochemistry, Second Edition: Springer-Verlag, Berlin-Heidelberg, New York, 208 p. Javoy, M., 1977, Stable isotopes and geothermometry: Journal of t h e Geological Society of London, v. 133, p. 609-636. Jensen, M. L., Ashley, R. P., and Albers, J. P., 1971, Primary and secondary sulfates a t Goldfield, Nevada: Economic Geology, v. 66, p. 618-626. Jensen, M. L., Field, C. W., and Nakai, N., 1960, Sulfur isotopes and t h e origin of sandstone-type uranium deposits: Biennial Progress Report f o r 1959-1960, U.S. Atomic Energy Commission C o n t r a c t AT(30-11-2261, 281 p. Kamilli, R. J., and Ohmoto, H., 1977, Paragenesis, zoning, fluid inclusion, and isotope studies of t h e Finlandia vein, Colqui district, c e n t r a l Peru: Economic Geology, v. 72, p. 950-982. Kesler, S. E., Russell, N., Seaward, M., Rivera, J., McCurdy, K., Cumming, G. L., and Sutter, J. F., 1981, Geology and geochemistry of sulfide mineralization underlying t h e Pueblo Viejo goldsilver oxide deposit, Dominican Republic: Economic Geology, v. 76, p. 1096-1 117. Kolodny, Y., and Epstein, S., 1976, Stable isotope geochemistry of d e e p s e a cherts: Geochimica e t Cosmochimica Acta, v. 40, p. 1195-1209. Kulla, J. B., and Anderson, T. F., 1978, Experimental oxygen isotope fractionation between kaolinite and water; & Zartman, R. E. (ed.), Short Papers of ~nternational Conference, the Fourth Geochronology, Cosmochronology, Isotope Geology 1978: U.S. Geological Survey, Open-File Report 78-701, p. 234-235. Kusakabe, M., and Chiba, H., 1983, Oxygen and sulfur isotope composition of barite and anhydrite from t h e Fukazawa deposit, Japan; Ohmoto, H., and Skinner, B. J. (eds.), The Kuroko and R e l a t e d Volcanogenic Massive Sulfide Deposits: Economic Geology, Monograph 5, p. 292-301. Lea, D. W., Larson, P. B., and Taylor, H. P., Jr., 1984, Oxygen isotope and fluid inclusion study of veins and wallrocks in t h e Eureka Graben, San Juan Mountains, Colorado abs. : Geological Society of America, Abstracts with Programs, v. 16, no. 6, D. 572. ~ o m b a r d i , G., and Sheppard, S. M. F., 1977, Petrographic and isotopic studies of t h e a l t e r e d acid volcanics of t h e Tolfa-Cerite area, Italy--the genesis of t h e clays: Clay Minerals, v. 12, p. 147-162. Magaritz, M., Whitford, D. J., and James, D. E 1 78, Oxygen isotopes a n d t h e origin of high 87;r/z6Sr andesites: E a r t h and Planetary Science Letters, v. 40, p. 220-230. Matsuhisa, Y., Goldsmith, J. R., and Clayton, R. N., 1979, Oxygen isotopic fractionation in t h e system quartz-albite-anorthite-water: Geochimica e t Cosmochimica Acta, v. 43, p. 1131-1 140. Nockolds, S. R., Knox, R. W. O'B., and Chinner, G. A., 1978, Petrology: Cambridge University Press, New York, 435 p. Ohmoto, H., 1972, Systematics of sulfur and carbon isotopes in hydrothermal o r e deposits: Economic Geology, v. 67, p. 551-578.

C . W. FIELD& R. H. FIFAREK Ohmoto, H., and Lasaga, A. C., 1982, Kinetics of reactions between aqueous sulfates and sulfides in hydrothermal systems: Geochimica et Cosmochimica A c t a , v. 46, p. 1727-1745. Ohmoto, H., and R y e , R. O., 1974, Hydrogen and oxygen isotopic compositions of fluid inclusions in t h e Kuroko deposits, Japan: Economic Geology, v. 69, p. 947-953. Ohmoto, H., and Rye, R. O., 1979, Isotopes of sulfur and carbon; & Barnes, H. L. (ed.), Geochernistry of Hydrothermal Ore Deposits, Second Edition: John Wiley and Sons, New York, p. 509-567. O'Neil, J. R., 1977, Stable isotopes in mineralogy: Physics and Chemistry of Minerals, v. 2, p. 105-123. O'Neil, J. R., and Kharaka, Y. K., 1976, Hydrogen and oxygen isotope exchange reactions between clay minerals and water: Geochimica et Cosmochimica A c t a , v. 40, p. 241-246. O'Neil, J. R., and Silberman, M. L., 1974, Stable isotope relations in epithermal Au-Ag deposits: Ecoornic Geology, v. 69, p. 902-909. O'Neil, J. R., Silberman, M. L., Fabbi, B. P., and Chesterman, C. W., 1973, Stable isotope and chemical relations during mineralization in t h e Bodie mining district, Mono County, California: Economic Geology, v. 68, p. 765-784. O'Neil, J . R., and Taylor, H. P., Jr., 1967, The oxygen isotope and cation exchange chemistry of feldspars: American Mineralogist, v. 52, p. 1414-1437. Osatenko, M. J., and Jones, M. B., 1976, Valley Copper; & Brown, A. S. (ed.), Porphyry Deposits of the Canadian Cordillera: The Canadian Institute of Mining and Metallurgy, Special Volume 15, p. 130-143. Porter, E. W., and Ripley, E., 1985, Petrologic and stable isotope study of t h e gold-bearing breccia pipe a t t h e Golden Sunlight deposit, Montana: Economic Geology, v. 80, p. 1689-1706. Power, S. G., 1985, The "tops" of porphyry copper deposits--mineralization and plutonism in t h e Western Cascades, Oregon: Unpublished Ph.D. dissertation, Oregon S t a t e University, 234 p. Radtke, A. S., Rye, R. O., and Dickson, F. W., 1980, Geology and stable isotope studies of t h e Carlin gold deposit, Nevada: Economic Geology, v. 75, p. 641-672. Robinson, B. W., 1974, The origin of mineralization a t t h e Tui mine, T e Aroha, New Zealand, in t h e light of stable isotope studies: Economic Geology, v. 69, p. 910-925. Rye, R. O., Doe, B. R., and Wells, J . D., 1974, Stable isotope and lead isotope study of t h e Cortez, Nevada, gold deposit and surrounding area: U.S. Geological Survey, Journal of Research, v. 2, p. 13-23. Rye, R. O., and Ohmoto, H., 1974, Sulfur and carbon isotopes and o r e genesis--a review: Economic Geology, v. 69, p. 826-842. Rye, R. O., and Sawkins, F. J., 1974, Fluid inclusion and stable isotope studies on t h e Casapalca AgPb-Zn-Cu deposit, central Andes, Peru: Economic Geology, v. 69, p. 181-205.

127

Sakai, H., 1968, Isotopic properties of sulfur compounds in hydrothermal processes: Geochemical Journal (Japan), v. 2, p. 29-49. Sakai, H., Des Marais, D. J., Ueda, A., and Moore, J. G., 1984, Concentrations and isotope ratios of carbon, nitrogen, and sulfur in ocean-floor basalts: Geochimica et Cosmochimica Acta, v. 48, p. 2433-2441. Sakai, H., Gunnlaugsson, E., Tomasson, J., and Rouse, J . E., 1980, Sulfur isotope systematics in Icelandic geothermal systems and influence of seawater Geochimica e t circulation a t Reykjanes: Cosmochimica Acta, v. 44, p. 1223-1231. Sangster, D. F., 1968, Relative sulphur isotope abundances of ancient seas and stratabound sulphide deposits: Proceedings of t h e Geological Association of Canada, v. 17, p. 79-91. Savin, S. M., and Epstein, S., 1970a, The oxygen and hydrogen isotope geochemistry of clay minerals: Geochimica et Cosmochimica Acta, v. 34, p. 25-42. Savin, S. M., and Epstein, S., 1970b, The oxygen and hydrogen isotope geochemistry of ocean sediments and shales: Geochimica e t Cosmochimica Acta, v. 34, p. 43-63. Sawkins, F. J., 1972, Sulfide o r e deposits in relation t o p l a t e tectonics: Journal of Geology, v. 80, p. 377-397. Schoen, R., and Rye, R. O., 1970, Sulfur isotope distribution in solfataras, Yellowstone National Park: Science, v. 170, p. 1082-1084. Sheppard, S. M. F., Nielsen, R. L., and Taylor, H. P., Jr., 1969, Oxygen and hydrogen isotope ratios of clay minerals from porphyry copper deposits: Economic Geology, v. 64, p. 755-777. Sheppard, S. M. F., Nielsen, R. L., and Taylor, H. P., Jr., 1971, Hydrogen and oxygen isotope ratios in minerals from porphyry copper deposits: Econornic Geology, v. 66, p. 515-542. Spooner, E. T. C., Beckinsale, R. D., England, P. C., and Senior, A., 1977, Hydration, 1 8 0 enrichment and oxidation during ocean floor hydrothermal metamorphism of ophiolite metabasic rocks from E. Liguria, Italy: Geochimica e t Cosmochimica Acta, v. 41, p. 857-871. Steiner, A., and R a f t e r , T. A., 1966, Sulfur isotopes in pyrite, pyrrhotite, alunite, and anhydrite from s t e a m wells in t h e Taupo Volcanic Zone, New Zealand: Economic Geology, v. 61, p. 1115-1129. Sternfeld, J. N., 1981, The hydrothermal petrology and s t a b l e isotope geochemistry of two wells in t h e Geysers geothermal field, Sonoma County, California: Unpublished M.S. thesis, University of California (Riverside), 202 p. Styrt, M. M., Brackman, A. J., Holland, H. D., Clark, B. C., Pisutha-Arnond, V., Eldridge, C. S., and Ohmoto, H., 1981, The mineralogy and t h e isotopic composition of sulfur in hydrothermal sulfide/sulfate deposits of t h e E a s t Pacific Rise, 2 1 ' ~ latitude: E a r t h and Planetary Science L e t t e r s , v. 53, p. 382-390. Suzuoki, T., and Epstein, S., 1976, Hydrogen isotope fractionation between OH-bearing minerals and waters: Geochimica e t Cosmochimica Acta, v. 40, p. 1229-1240.

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Chapter 7 GEOLOGIC, MINERALOGIC, AND GEOCHEMICAL CHARACTERISTICS OF VOLCANIC-HOSTED EPITHERMAL PRECIOUS-METAL DEPOSITS Daniel 0. Hayba, Philip M. Bethke Pamela Heald, and Nora K. Foley

INTRODUCTION In Chapter I, R. W. Henley summarized our understanding of t h e chemical and hydrodynamic s t r u c t u r e and t h e transport properties of active hydrothermal systems, with particular emphasis on terrestrial magmatic-hydrothermal systems. Such a n overview is especially valuable because active geothermal systems a r e modern "archetypes" of t h e ancient systems which concentrated m e t a l s in their upper portions t o form epithermal ore deposits. More than any other factor, t h e study of a c t i v e systems has provided t h e framework on which t h e observations on epithermal deposits have been arranged in t h e relatively recent development of comprehensive models of epithermal ore formation. The Principle of Uniformitarianism has served us well in this instance. In this chapter, we focus on observations on epithermal ore deposits in continental silicic t o andesitic volcanic terranes. Volcanic-hosted deposits offer t h e most direct comparison with many of the well-studied modern geothermal systems. We first compare the attributes from a number of epithermal o r e deposits and show how they may b e used t o identify two important, and distinct volcanic-related hydrothermal environments. We then examine t h e best-studied deposit of e a c h type: C r e e d e and Summitville, both of which a r e located in t h e San Juan Mountains in southwest Colorado. In so doing, we a r e able t o examine epithermal deposits for evidences of processes t h a t a r e now occurring in geothermal systems. Finally, we use t h e observational base and interpretations derived from e a c h deposit t y p e t o develop generalized "geothermal" models of mineralization. The models have been taken, in large part, from t h e excellent synthesis by Henley and Ellis (1983). We feel t h a t their models a r e soundly based on a myriad of direct observations on a c t i v e geothermal systems, and a r e consistent, for t h e most part, with t h e observations made a t C r e e d e and Summitville.

SUMMARY O F THE CHARACTERISTICS O F VOLCANIC-HOSTED EPITHERMAL ORE DEPOSITS In 1981, Buchanan published a valuable compilation of selected observations from over 60 gold-silver vein deposits in unmetamorphosed volcanicto-subvolcanic environments. These d a t a and t h e integrated model derived from them have formed a useful basis for numerous subsequent analyses of t h e characteristics of epithermal deposits, s o m e of which include Giles and Nelson (19821, Ashley (19821, Bonham

and Giles (1983), Heald-Wetlaufer e t al. (19831, Berger and Eimon (19831, Sillitoe and Bonham (1984), Heald e t al. (19861, and Bonham (1986). Because of the large number of deposits in Buchanan's 1981 study, a detailed evaluation of t h e d a t a base was not possible. Heald e t al. (1986) undertook such a detailed study of 16 carefully selected, well-studied, Tertiary volcanichosted epithermal deposits. Using this d a t a base, Heald e t al. (1986) showed t h a t t w o types of epithermal deposits could b e distinguished and t h a t Buchanan's d a t a base supplemented and supported this conclusion. The t w o types, distinguished primarily on t h e basis of vein and a l t e r a t i o n mineralogies, a r e t h e Adularia-Sericite t y p e and t h e Acid-Sulfate type. The characteristic f e a t u r e s of t h e s e two main types a r e shown in Table 7.1 a n d a r e discussed below. AdulariaSericite-type deposits a r e f a r more numerous t h a n t h e Acid-Sulfate-type deposits (Table 7.21, and so predictably, t h e r e is more information on t h e former. Major telluride deposits (e.g., Cripple Creek, Colorado) were excluded in t h e study by Heald e t al. (1986) because their unique mineral assemblage suggests t h a t t h e s e deposits make up a distinct class (Heald-Wetlaufer et al., 1983; Bonham and Giles, 1983; Hot-spring-type deposits (see Bonham, 1986). Silberman and Berger, 1985, this volume; Berger and Silberman, 1985, this volume) were not well represented in their compilation because of a paucity of published descriptive data. Characteristics of Adularia-Sericite-type Deposits Structural setting--The most common regional structural s e t t i n g f o r Adularia-Sericite-type deposits is along t h e margins of calderas, although other t e c t o n i c settings (typically structurally complex volcanic environments) a r e not uncommon. The importance of t h e caldera s e t t i n g lies in t h e excellent plumbing system i t provides for hydrothermal circulation (Lipman et al., 1976; Steven and Lipman, 1976). It is important t o note, however, t h a t relatively f e w calderas in t h e western United S t a t e s have been mineralized (McKee, 1979; Rytuba, 19811, and in t h e San Juan Mountains, Colorado, only about 113 of t h e known calderas have had significant mineral production (Steven and Lipman, 1976). Size of deposit--There is a large range in t h e size of t h e Adularia-Sericite-type deposits. Guanajuato, a silver- and base-metal-rich district, covers a surface a r e a of approximately 190 s q km. Districts with relatively low base-metal contents, such a s Oatman, tend t o be considerably smaller (12 sq km); Comstock is a n important exception. The length-to-width r a t i o

Table 7.1--Characteristics of the adularia-sericite type and acid-sulfate type deposits (compiled from Heald et al., 1986). Acid-Sulfate

Adularia-Sericite

Structural setting

Intrusive centers, 4 out of the 5 studied related to the margins of calderas

Structurally complex volcanic environments, commonly in calderas

Size 1ength:width ratio

relatively small equidimensional

variable; some very large usually 3:l or greater

Host rocks

rhyodacite typical

silicic to intermediate volcanics

Timing of ore and host

similar ages of host and ore ( I m.y.)

Mineralogy

enargite, pyrite, native gold, electrum, and basemetal sulfides Chlorite rare no selenides Mn-minerals rare sometimes bismuthinite

argentite, tetrahedrite, tennantite, native silver and gold, and base-metal sulfides chlorite common selenides present Mn gangue present no bismuthinite

Production data

Both gold- and silverrich deposits noteworthy Cu production production

Both gold- and silverrich deposits variable base-metal

Alteration

Advanced argillic to argillic ( * - sericitic)

Sericitic to argillic

Extensive hypogene alunite Major hypogene kaolinite No adularia

supergene alunite occasional kaolinite Abundant adularia

Salinity

1 to 24 wt% NaCl eq.3

0 to 13 wt% NaCl eq.

Source of fluids

Dominantly meteoric, possibly significant magmatic component

Dominantly meteoric

Source of sulfide sulfur

Deep-seated, probably magmatic

Deep-seated, probably derived by leaching wallrocks deep in system

Source of lead

Volcanic rocks or magmatic fluids

Precambrian or Phanerozoic rocks under volcanics

Temperature

'could be secondary in some districts. 2~imiteddata, possibly unrelated to ore. 3~alinitiesof 5 to 24 wt% NaCl eq. are probably related to the intense acid-sulfate alteration which preceded ore deposition.

of t h e surface projection of mineralized veins in this t y p e of deposit is generally on t h e order of 3:l. The vertical range of mineralization is usually on t h e order of 400 t o 700 m e t e r s compared t o strike lengths of several kilometers. Host rocks--The composition of t h e host rocks of Adularia-Sericite-type deposits ranges from rhyolitic t o andesitic, and o r e is commonly hosted by several different compositional units within a district. In a few of t h e districts, t h e ore fluids also mineralized associated sediments (Creede, Guanajuato) o r intrusive rocks (Silver City, Idaho), but o r e is mainly confined t o t h e volcanic rocks. T h e occurrence of o r e in several lithologies in Adularia-Sericite deposits implies t h a t composition of t h e host rock(s) is not a controlling factor. The lack of a genetic t i e t o t h e host is further implied by t h e f a c t t h a t ore deposition in AdulariaSericite-type deposits almost always occurred more than 1 m.y. subsequent t o the formation of t h e host rock. Mineralogy--The mineralogy of t h e AdulariaSericite-type deposit is characterized by t h e presence of vein adularia and sericite and by t h e absence of both hypogene alunite and t h e assemblage: enargite + pyrite +/- covellite (Table 7.2). Chlorite is also characteristically present. Although alunite is found in some of t h e s e deposits, in e a c h c a s e i t appears t o b e a near-surface supergene occurrence, unrelated t o t h e primary ore-forming system. Metal ratios--The high silver-to-gold production ratios of most of t h e Adularia-Sericite-type deposits r e f l e c t t h e abundance of native silver, silver sulfides, and sulfosalts. Districts such a s Round Mountain, Nevada, and Oatman, Arizona, have low silver-to-gold ratios. The precious metals in these t w o districts a r e present mainly a s native gold, native silver, and electrum; silver sulfides and silver sulfosalts a r e r a r e (Table 7.1). Base-metal production is also usually low for t h e m o r e gold-rich deposits. In most cases this r e f l e c t s t h e lack of base-metal sulfides and sulfosalts, but base-metals a r e sometimes not recovered in milling operations and therefore production d a t a may not always accurately reflect t h e o r e mineralogy. The fairly large sample size of t h e Adularia-Sericite-type deposits in Table 7.2 permits speculation t h a t t h e r e may b e a continuum from t h e base-metal-rich, silverrich districts, such a s Colqui, Peru t o t h e base-metalpoor, gold-rich districts, such a s Oatman. Wallrock alteration--The alteration p a t t e r n s of Adularia-Sericite-type deposits a r e not yet well defined, due in part t o t h e lack of detailed alteration studies. In general, Adularia-Sericite-type deposits a r e c h a r a c t e r i z e d by the predominance of s e r i c i t i c alteration* t h a t often borders a silicified zone near t h e vein. Also near t h e vein, fine-grained potassium feldspar and/or chlorite a r e often disseminated in t h e wallrock. The sericitic zone typically grades outward into a propylitic zone. An argillic zone between t h e sericitic and propylitic zones is sometimes present. A t Creede, Pachuca (Mexico) and Oatman, t h e alteration over t h e o r e body has been described a s a sericitic "cap", interpreted t o be t h e result of t h e condensation of acid volatiles released a t depth during boiling (Barton et al., 1977; Buchanan, 1981). In many, if not most districts, t h e outermost propylitic alteration

zone (typical of both Acid-Sulfate- and AdulariaSericite-type deposits) appears t o have formed prior t o o r e deposition and may b e unrelated t o t h e oreforming hydrothermal system. *The t e r m "sericitic" is used in this chapter in the sense of alteration consisting of a mica-type mineral (e..g., illite) + q u a r t z + pyrite, including mixed-layer i l l ~ t e - s m e c t i t ein which illite layers a r e predominant. Thermal histor --For t h e Adularia-Sericite-type d e p o s ~ t s , iluid-inclu5on studies tied t o a detailed pa;agenktic sequence (compiled in Hayba, 1983) show t h a t most ore deposition occurred a t t e m p e r a t u r e s between 2 0 0 ' ~ and 300°C, with late-stage fluids typically depositing only gangue minerals between 140°C and 200°C. Of t h e 11 Adularia-Sericite deposits evaluated in Heald e t al. (19861, boiling has unequivocally been demonstrated t o b e associated with precious-metal deposition only a t Colqui (Kamilli and Precious-metal deposition due t o Ohmoto, 1977). boiling has been proposed for other districts based on less definitive evidence (cf. Guanajuato (Buchanan, 1979), Tonopah (Fahley, 1981)). Roedder (1984, p. 426) notes t h a t some of t h e evidence for boiling a t Guanajuato and Tonopah is ambiguous and t h a t t h e occurrence of boiling in these deposits rnay have been overstated. Boiling also has been noted a t Eureka (Casadevall and Ohmoto, 19771, Pachuca (Drier, 19761, and C r e e d e (Barton e t al., 1977) during t i m e s other t h a n precious-metal deposition. Although boiling is a n e f f e c t i v e mechanism for depositing o r e s (Henley e t al., 1984; Henley, 1985, this volume; Drummond and Ohmoto, 1985; R e e d and Spycher, 1985, this volume), i t is not t h e only mechanism. In t h e study of any o r e deposit, t h e evidence for boiling should b e critically evaluated before assuming i t contributed t o preciousm e t a l deposition. Bodnar e t al. (1985, this volume) discuss problems in using fluid-inclusion evidence t o document boiling in ore deposits. In addition t o boiling, t h e mixing of fluids from t w o o r more sources has played a role in t h e thermal history of AdulariaSericite-type deposits a s has been documented by fluid-inclusion and stable-isotope studies a t C r e e d e (discussed in detail later). Composition of fluids-Salinities, deterrnined from freezing point depression measurements, range from 0 t o 13 wt.-% NaCl equivalent for fluid inclusions from those Adularia-Sericite-type deposits evaluated. Most of t h e deposits have consistently low salinities, usually less than 3 wt.-% (Hayba, 19831, but C r e e d e and Colqui have unusually high salinity fluids, ranging from 5 t o 12 wt.-% NaCl equivalent (Woods et al., 1982; Kamilli and Ohmoto, 1977). The high salinities in these t w o deposits may have a bearing on their relatively high base-metal contents because chloride complexes a r e a n e f f e c t i v e means of transporting base m e t a l s (Barnes, 1979; Henley, 1985, this volume). On t h e other hand, fluid inclusions from t h e Sunnyside (nine, Eureka mining district, Colorado, which is also very base-metal rich, have low salinities (0-3.6 percent). Hedenquist and Henley (1985) and Henley

132

CHAPTER 7

Table 7.2--Selected mineralogical and production data for 54 epithermal districts (from Heald et al., 1986). Both types of districts are ordered from higher to lower base-metal production. The last 14 districts may represent a distinctively low base-metal subtype of the adularia-sericite-type district. AD, adularia; AL, alunite; SS, sulfosalts; AGS, silver sulfides; SP, sphalerite; GN, galena; EN, enargite; BA, barite; RC, rhodochrosite; EL - fluorite; X - denotes presence.

District

Ag :Au

Base ~etal' percent

AD

AL

SS

X X X X X

X X X

X X X

AGS

SP

GN

EN

BA

RC

FL

X X X X

X X X X X

X X X X X

X

X

X

X X X X X X X X

X X

X X

X X

X X

X

X

X

ACID-SULFATE TYPE *Red Mtn., Colo. *Julcani, Peru *Lake City 11, Colo. 5 *Summitville, Colo. *Goldfield, Nev.

68 467 23 1.2 0.3

high mod mod 5? 1

259 26 45 150 6 Ag only 47 400 200 100-500 74 1200 64 200 162 4-30 243 2 80-110 Ag only 51

high 17.5 10 4-20 9 8.5 7.5 5 0-10 6-12 mod mod 4 3.5 3 3 1-4 2 2 1.4 1 1ow minor minor v low v low v low rare

X X X X X

X

ADULARIA-SERICITE TYPE *Lake City (12, Colo. 5 *Colqui, Peru Tovar, Mex. Parral, Mex. Bohemia, Oreg. Namiquipa, Me 7. *Eureka, Colo. *Creede, Colo. "Guanajuato, Mex. Guanacevi, Mex. Yoquivo, Mex. Calico, Calif. Fresnillo, Mex. *Pachuca, Mex. El Tigre, Mex. Great Barrier, New Zea. Silver Peak, Nev. Piz Piz, Nicaragua, "Tonopah, Nev. Zalcualpan, Mex. Tayoltita, Mex. Temalscaltepec, Mex. Golden Plateau, Aus. Mohave, Calif. Divide, Nev. Bodie, Calif. Stateline, Utah Guadalupe, Mex. Mogollon, N. Mex. *Corns Lock, Nev. *Silver City, Idaho *De Lamar, Idaho Gold Circle, Nev. National, Nev. Aurora, Nev. Ramsey-Tala, Nev. Republic, Wash. Searchlight, Nev. Rawhide, Nev. Ocampo, Mex.

1 1 1 1 1 1 1 0.3 0.2 0.03

X X X X X X X X X

X

X X X X X

X X X

X

X X X

X X

X X

X X X X

X X X

X

X

X

Table 7.2--Selected mineralogical and production data for 54 epithermal districts (from Heald et al., 1986)--(continued)

District

Ag :Au

Base Metal 1 percent

AD

AL

SS

AGS

SP

GN

EN

BA

RC

FL

ADULARIA-SERICITE TYPE (continued) Rochester, Nev. Tuscarora, Nev. *Round Mtn. Nev. Cornucopia, Nev. Wonder, Nev. Seven Troughs, Nev. El Oro, Mex. Jarbidge, Nev. Gilbert, Nev. Hayden Hill, Calif. Katherine, Ar iz. *Oatman, Ariz.

113 44-100 0.2 68 94 5.4 7 3 Au only 1.5 3 0.4

*Denotes the 16 districts studied in detail in Heald eL al., 1986; the others were modified from Buchanan (1981). '~ase-metal production, usually as sulfides, rarely as metals, as a percent of the total tonnage (Buchanan, 1981). 2~luniteyounger than mineralization. 3~luniteshallow only, may be supergene. 4~luniteolder than mineraliza~ion. 5 ~ l a c k(1980) grouped the Galenea (Henson Creek) and Lake (Lake San Cristobal) districts, Colo. into the Lake City district; the earlier ores (Lake City I ) occur primarily in the Galena district and the later ones (Lake City 11) in the Lake dis~rict. 6 ~ a ~ for a the Finlandia vein. 7 ~ a t amainly for the Sunnyside vein. (1985, this volume) have noted t h a t t h e interpretation of salinity from measurements of fluid inclusions in s o m e deposits can b e skewed t o higher values if t h e fluids have relatively high gas contents, because C 0 2 and H2S contribute t o t h e freezing point depression. However, t h e presence of C 0 2 in these epithermal deposits has not been documented. Although high concentrations of C 0 2 can easily be d e t e c t e d by t h e crushing t e s t techniques described by Roedder (1970), Bodnar e t al. (1985, this volume) show t h a t inclusions trapped under "epithermal" conditions may contain significant amounts of C 0 2 which go undetected. A f e w analyses of inclusion fluids from C r e e d e (Roedder, compiled in Hayba, 19831, Eureka (Casadevall and Ohmoto, 19771, and Colqui (Kamilli and Ohmoto, 1977) indicate C 0 2 concentrations of less than 0.5 molal in almost all samples.

Paleodepth--Paleodepth e s t i m a t e s from both geologic reconstructions and pressures estimated from fluid-inclusion studies show t h a t most of t h e AdulariaSericite-type deposits appear t o have formed a t paleodepths of 300-600 m, although both methods have rather large uncertainties (Roedder and Bodnar, 1980). Colqui, and possibly Eureka, a r e thought t o have formed a t g r e a t e r depths, approximately 1000 and 1400 m, respectively. Based only on geologic reconstruction estimates, Round Mountain, Oatman, and DeLamar s e e m t o have formed a t shallower depths, possibly a s shallow a s 100 m. The shallow paleodepths and high gold/silver ratios of Round Mountain and O a t m a n a r e consistent with models for gold deposition presented by Henley (1985, this volume), Silberman and Berger (1985, this volume), and Hedenquist and Henley (1985).

CHAPTER 7

Sources of fluids--The predominance of m e t e o r i c w a t e r has been documented in several epithermal deposits, but only C r e e d e (Bethke and Rye, 19791, Colaui (Kamilli and Ohmoto. 1977). and Sunnyside ( ~ a & d e v a l l and Ohmoto, 1977) habe had isotopic studies related t o paragenetic sequences. However, even with these more comprehensive studies, t h e s y s t e m a t i c s of hydrogen and oxygen isotopes may b e m o r e complex than previously thought. Very detailed sampling of quartz crystals a t C r e e d e by Foley e t al. (1982) has shown t h a t t h e previously reported 6 D values f o r quartz (Bethke and Rye, 1979) do not truly represent primary fluids, but rather represent a mixture, during extraction and analysis, of hydrothermal fluids trapped in primary fluid inclusions, and isotopically lighter, overlying ground waters, trapped in pseudosecondary inclusions (discussed more thoroughly later). Without such detailed data, i t is difficult t o draw any more specific conclusions other than t h a t t h e mineralizing solutions were deeply circulating, dominantly m e t o r i c waters. I t should be noted, however, t h a t a magmatic component of up t o ten percent could b e hidden in t h e uncertainties and cannot b e ruled out. Source of sulfur and lead--For t h e 16 deposits considered by Heald e t al. (19861, t h e sulfur-isotopic d a t a f o r t h e Adularia-Sericite deposits a r e limited t o measurements made on C r e e d e (R. 0. Rye, U.S.G.S., personal communication, 1985), Sunnyside (Casadevall and Ohmoto, 19771, and Colqui (Kamilli and Ohrnoto, 1977). Although t h e 6 3 4 ~values for t h e sulfides f o r all t h r e e deposits cluster very close t o 0 permil, quite different interpretations of t h e sources of t h e sulfur have been proposed for t h e different deposits. This is due i p a r t t o geologic considerations, but mainly t o t h e 6 4~ values for t h e sulfate minerals; if equilibrium conditions a r e assumed, t h e r e must be a material balance between t h e isotopically light sulfides and heavy sulfates. For Sunnyside and Colqui, t h e sulfates range from + I 0 t o +25 permil, but a t C r e e d e t h e y vary The significance of t h e from +17 t o +45 permil. unusually heavy sulfate sulfur a t C r e e d e will b e discussed later. A t Sunnyside, Casadevall and Ohmoto (1977) suggest t h a t t h e sulfur was derived from evaporitebearing s e 'mentary rocks which have a bulk sulfur value of 6$4 S -- +12 permil a s sulfate and a r e located outside of t h e San Juan and Silverton calderas. They base their conclusion on t h e assumptions of sulfidesulfate isotopic equilib ium (based on a pyriteanhydrite pair having a 6j4S difference of 22.3 permil) and t h a t t o t a l sulfate concentration in t h e fluid was g r e a t e r t h a n or equal t o t o t a l sulfide concentration (indicated from mineralogical data). In order t o produce t h e 0 perrnil sulfides in equilibrium with t h e heavier sulfates a t t h e presumed equilibration t e m p e r a t u r e of 300°c, a source of 6 3 4 ~= +12 permil is required. Kamilli and Ohrnoto (1977) also prefer a sedimentary sulfate source for t h e sulfur a t Colqui, but suggest t h a t a n igneous origin is possible. A t Creede, i t is clear t h a t sulfate-sulfide isotope relationships were not governed by equilibrium fractionation, and t h a t aqueous sulfide and s u l f a t e were essentially independent systems (Bethke e t al., 1973). Although t h e uniquely heavy sulfates a t C r e e d e

'I

require extensive isotopic evolution, Bethke et al.'s proposal for non-equilibrium, essentially independent sulfide-sulfate systems is also tenable f o r t h e other deposits. I t is probable t h a t non-equilibrium sulfidesulfate conditions existed for most epithermal deposits (Ohmoto and Lasaga, 1982). R a t h e r than assuming sulfide-sulfate equilibrium and relying on s c a r c e sedimentary units, sulfate/sulfide ratios, and material balance considerations t o produce 0 permil sulfides, i t is perhaps b e t t e r t o assume t h a t t h e isotopic composition of t h e sulfide sulfur was controlled by t h e thick volcanic piles associated with e a c h of these deposits. For those Adularia-Sericite districts where leadisotopic studies have been done, Tonopah (Zartman, 1974), Pachuca (Cumming e t al., 1979) Creede, Lake City, Colorado, and Sunnyside (Doe e t al., 19791, t h e relatively radiogenic initial-lead ratios of galena suggest t h a t a significant component of t h e lead may have been derived from Precambrian o r Phanerozoic rocks underlying t h e volcanic rocks. This indicates t h a t t h e o r e components of Adularia-Sericite deposits may have been derived, in large part, by leaching of wallrocks deep in t h e system. Characteristics of Acid-Sulfate-type Deposits Structural setting--Four of t h e five Acid-Sulfate deposits listed in Table 7.2 a r e spatially r e l a t e d t o t h e margins of calderas. The other Acid-Sulfate deposit, Julcani, is associated with silicic domes a t t h e intersection of major faults. The association of these deposits with intrusive centers, particularly ringf r a c t u r e volcanic domes on t h e margins of calderas, appears t o be a critical genetic f a c t o r (Heald e t al., 1986) in contrast t o t h e Adularia-Sericite-type deposit where t h e role of calderas is o n e of ground preparation for l a t e r hydrothermal fluids (Lipman e t al., 1976; Steven and Lipman, 1976). Size of deposit--On t h e whole, Acid-Sulf a t e deposits a r e smaller in t e r m s of tonnage than Adularia-Sericite-type deposits, although this may b e a result of a limited d a t a base. In addition, t h e surface projection of t h e productive a r e a s is relatively equidimensional rather than elongated. For example, t h e a r e a l e x t e n t of t h e most productive part of Goldfield, Nevada, is only approximately 2 km long by 1.5 km wide. For all t h e Acid-Sulfate-type deposits, t h e vertical e x t e n t of t h e mineralized a r e a is much smaller than t h e horizontal e x t e n t , usually less than 500 meters. Host rock--The primary host rock for t h e AcidSulfate-type deposit is almost exclusively rhyodacite which is commonly porphyritic. A t Julcani, s o m e o r e also occurs in dacite, and a t Goldfield, trachyandesite and rhyolite, in addition t o rhyodacite, host t h e ore. The timing of o r e deposition relative t o t h e emplacement of t h e host rock is of particular importance. Age d a t e s show t h a t o r e deposition very closely ( 0.5 m.y.1 followed emplacement of t h e host rock (Table 7.3) indicating a possible genetic relationship. An exception is t h e Lake C i t y I1 deposit (i.e., t h e Lake district; s e e f o o t n o t e 5 on Table 7.2) where t h e ores occur in a quartz l a t i t e ash-flow tuff which is considerably older than t h e ore. However, t h e

D. 0 . HAYBA, P. M. BETHKE, P. HEALD, & N. K. FOLEY Table 7.3--Age and type of host rock and age of mineralization for five Acid-Sulfate districts (modified from Heald et al., 1986). Age of Host Rock (m-y.

Age of Mineralization

Reference

District

Host Rock

Red Mtn., Colo.

rhyodacitedome

21.3t023.6

23.1f0.6

Lipman et al., 1976 Mehnert et al., 1973 Gilzean et al., 1984

Julcani, Peru

rhyodacite dome dacite dome

9.67 f 0.05 to 10.13 f 0.03

9.83 f 0.08

Petersen et al., 1977 Noble and Silberman, 1984

Lake City 11, quartz latite Colo. ash-f low

(28.4,

22.5l

Slack, 1980 Lipman et al.,

Summitville, Colo.

rhyodacite dome

22.8 f 0.6

22.4 f 0.5

Mehnert et al.,

Goldfield, Nev

rhyodacite dome trachyandesite rhyolite

21.3 f 0.3 21.5f 0.5 28-33

21.0 f 0.4

Silberman and Ashley, 1970

.

>27.8

1976 1973

'slack (1330) proposes that these later ores were generated during emplacement of the nearby Red Mountain rhyodacite dome, approximately 22.5 m.y. ago.

o r e a t Lake C i t y I1 fills f r a c t u r e s t h a t form a radial p a t t e r n around a rhyodacite flow dome t h a t Slack (1980) suggested is genetically related t o t h e Lake C i t y I1 o r e s and has a similar a g e t o t h e ores. Hon e t al. (1985) have shown t h a t s o m e veins included in t h e Lake C i t y I1 district a r e older than t h e rhyodacite flow dome and probably formed from an earlier hydrothermal system(s). Mineralogy--Acid-Sulfate-type deposits are characterized by t h e occurrence of t h e vein mineral assemblage enargite + pyrite +/- covellite. Adularia and chlorite a r e absent or rarely present. O r e occurs primarily a s native gold and electrum with sulfides, sulfosalts, and tellurides. Bismuthinite has been identified in 3 of t h e 5 Acid-Sulfate deposits in Table 7.2, but selenides, rhodochrosite, and fluorite a r e rare. Metal ratios--Summitville and Goldfield have low silver-to-gold ratios ( 2 : l ) which reflect t h e high proportion of f r e e gold and gold-bearing minerals. Julcani, Red Mountain and Lake C i t y I1 h a v e high silver-to-gold ratios ( 10:l) and a r e c h a r a c t e r i z e d by more abundant silver mineralization, primarily in t h e form of silver sulfides and sulfosalts. Copper constitutes a major proportion of t h e base-metal production for Acid- Sulfate-type deposits, especially a t Goldfield and Summitville, where copper accounts for more than 85% of t h e base-metal production. The silver-rich districts have proportionally m o r e base metals, but copper production is secondary t o lead (1:2 Cu:Pb ratio). Wallrock alteration--A definitive c h a r a c t e r i s t i c of Acid-Sulfate-type deposits is t h e association of advanced-argillic alteration with t h e ore. Kaolinite, usually accompanied by alunite, occurs close t o t h e

vein and is often coextensive with silicification. F a r t h e r from t h e vein, argillic alteration, sometimes intermixed with sericitic alteration, surrounds t h e zone of advanced-argillic alteration. The argillic alteration zone is commonly mineralogically zoned, with kaolinite nearer t h e vein and s m e c t i t e f a r t h e r from t h e vein. The outermost a l t e r a t i o n zone consists of propylitic alteration. Thermal history--The limited fluid-inclusion d a t a for Acid-Sulfate-type deposits indicate t h a t o r e deposition occurred primarily a t t e m p e r a t u r e s similar t o those of Adularia-Sericite-type deposits ( 2 0 0 ' ~ t o 3 0 0 ~ ~ )The . salinities of Acid-Sulfate-type deposits, which a r e not yet well documented, show a wide range. Secondary inclusions in q u a r t z phenocrysts which a r e thought t o contain t h e fluids responsible for t h e intense quartz + alunite +/- pyrite alteration which preceded o r e deposition have salinity ranges of 5-24, 7-21 and 5-18 wt.-% NaCl equivalent a t Julcani, Summitville and Goldfield, respectively (Bruha and Noble, 1983). However, limited d a t a from primary inclusions in quartz and sphalerite associated with main s t a g e o r e deposition from Lake C i t y I1 (Slack, 1980), R e d Mountain (Nash, 19751, and Summitville (Stoffregen, 1985) indicate salinities o n t h e order of 1 t o 6 wt.-% NaCl equivalent. These limited d a t a a r e examined more closely in a later section, but clearly, much remains t o be learned from fluid-inclusion studies of Acid-Sulfate-type deposits. Paleodepth--Paleodepth e s t i m a t e s for AcidSulfate-type deposits appear t o be similar t o AdulariaSericite deposits (300-600 m), although Nash (1975) suggests t h a t t h e R e d Mountain deposit may have formed a t depths g r e a t e r than 1200 m. Based solely on

geologic reconstruction, t h e gold-rich deposits, Goldfield and Summitville, may have formed a t somewhat shallower depths than t h e silver-rich deposits. Source of constituents--The stable-isotope a on Acid-Sulfate deposits is limited t o Xqme 6 0 numbers f o r Goldfield (Taylor, 1973) and 6"s d a t a for Julcani (Goodell. 1970). Goldfield (Jenson e t al.. 1971). and ~ u k m i t v i l i e (R.'o. Rye, U.S.G.S., personai communication, 1985). Without any deuterium data, t h e source of t h e hydrothermal fluid is equivocal. The possibility of a significant magmatic component is discussed later. The paucity of sulfur-isotope d a t a permits only generalities. The sulfides from these t h r e e districts plot between -7 and + 3 permil suggesting a m a g m a t i c source for t h e sulfur. Similarly, lead-isotopic studies a t Summitville and R e d Mountain show t h a t t h e galena is very similar isotopically t o t h e enclosing volcanic rock, implying t h a t e i t h e r t h e adjacent rocks or related magmatic fluids were t h e primary source of the lead (Doe e t al., 1979).

ciagt

Summary of --

closed-basin lake prior t o incorporation into t h e hydrothermal system is a n aspect t h a t may be uncommon, although a similar geologic setting may have existed a t Calico, California. The presence of a n evaporative lake is obviously not a requirement for o r e formation in these deposits, although t h e high salinities developed by evaporation probably made t h e C r e e d e system highly efficient in t h e transport of base-metals (cf. Henley, 1985, this volume). We also recognize t h a t within this group of deposits t h e r e is considerable diversity in specific characteristics. For example, Round Mountain and Oatman a r e both distinctively gold-rich and sulfide-poor relative t o t h e other districts discussed in t h e previous section. Other

1

I

COLORADO

(

I

Characteristics

The primary characteristics which distinguish t h e Adularia-Sericite-type deposits from t h e Acid-Sulfatetype deposits a r e t h e vein mineralogy and t h e alteration assemblages (Heald e t al., 1986). As shown in Table 7.1, t h e r e a r e also many other important, but less definitive, characteristics of e a c h type which need t o be considered in developing genetic models. A comparative study, such a s t h e one done by Heald e t al. (1986) helps d e t e r m i n e which characteristics a r e t h e salient f e a t u r e s of a deposit-type model and which f e a t u r e s a r e just local variations. In order t o examine e a c h of t h e s e t w o types of epithermal deposits in greater detail, a characteristic deposit from e a c h group will b e reviewed in t h e next sections. THE ADULARIA-SERICITE ENVIRONMENT: CREEDE AS AN EXAMPLE In t h e foregoing section t h e characteristics of 16 well-documented Tertiary volcanic-hosted epithermal ore deposits were summarized based on t h e comparative study done by Heald e t al. (1986). I t was concluded t h a t t w o distinctive types of deposits could be distinguished: I ) those deposits characterized by a n alteration assemblage dominated by adularia and sericite, and 2) those deposits characterized by a n alteration assemblage containing kaolinite and alunite. The most thoroughly studied of t h e AdulariaSericite-type deposits is t h e C r e e d e mining district, Colorado. Therefore, C r e e d e is useful a s t h e exemplar for t h e Adularia-Sericite-type deposits. C r e e d e a s a n Exemplar In basing our discussion of t h e Adularia-Sericite environment on C r e e d e we recognize t h a t some f e a t u r e s of t h e C r e e d e district a r e not representative of t h e group a s a whole. For example, t h e chemical and isotopic evolution of t h e C r e e d e o r e fluids in a

Known o r readily inferred calderas

Calderas w i t h major associated mineral deposits

Figure 7.1. Calderas of the San Juan volcanic field: S, Silverton; LC, Lake City; CP, kchetopa Park; Bz, Bonanza; U;, La Garita; SL, San Luis; B, Bachelor; C, Creede; MH, Mount Hope; PI Platoro; Stl, Summitville; L, Lost Lake; U, Ute Creek; SJ, San Juan; UN, Uncompahgre; MI general location of the Elammoth Mountain caldera. Af ter Steven and Eaton ( 1 9 7 5 ) .

D. 0. HAYBA,P. M. BETHKE,P. HEALD, & N. K. FOLEY

137

Figure 7.2. Generalized geology of the Creede and San Luis calderas in relation to remnants of the Bachelor (B) and La Garita calderas and to the Creede mining district (shown in box). Control is moderate to good where boundaries are shown by solid symbols, and conjectural where shown by open symbols. Pr, Point of rocks volcano; S, Spar City. After Steven and Lipman (1976).

differences could b e enumerated. In spite of t h e diversity in specific characteristics, we would argue t h a t t h e differences between t h e various districts represent variations on a theme, whose main characteristics a r e illustrated by t h e similarities. We a r e a w a r e t h a t t h e "Hot-Spring-type" California, and deposits such a s McLaughlin, Hasbrouck Mountain and Sulphur, Nevada, a r e not specifically t r e a t e d when C r e e d e is used a s a n exemplar because t h e r e is no evidence of surface discharge of t h e o r e fluids a t Creede. It is our current opinion, however, t h a t such hot-spring-type deposits form in t h e surficial p a r t s of hydrothermal systems similar t o those which deposited t h e C r e e d e ores. Berger and Eimon (1983), Henley (1985, this volume), Silberman and Berger (1985, this volume), and Berger and Silberman (1985, this volume) t r e a t hot-springtype deposits separately because of important structural a t t r i b u t e s t h a t , in t e r m s of o r e controls, somewhat distinguish these deposits from t h e deeper classical vein deposits. However, Silberman and Berger (1985, this volume) also present evidence from Bodie, California, t h a t links hot-spring-type deposits directly t o t h e Creede-type veins. Summary of Important Studies The C r e e d e mining district has been t h e focus of both extensive and intensive study f o r t h e past 30 years. Studies by T. A. Steven and co-workers of t h e U.S.Geologica1 Survey unravelled t h e volcanic history of t h e district, related t h e o r e deposits t o t h a t history, and documented t h e structural control of o r e

deposition (Steven and R a t t 6 , 1965; RattC and Steven, 1967; Steven and ~ a t t 6 , 1973; Steven and Liprnan, 1973; Steven, 1967; Steven and Eaton, 1975). Continued studies by T. A. Steven and P. W. Lipman and co-workers have worked out t h e volcanic stratigraphy and volcano-tectonic evolution of t h e e n t i r e San Juan volcanic field, providing a particularly well-developed regional context for t h e C r e e d e district (Steven and Lipman, 1976; Lipman e t al., 1970; Lipman et al., 1978; and Doe et al., 1979). Detailed mineralogical-geochemical studies have served t o develop a well-documented, comprehensive model of ore deposition (Roedder, 1960; Roedder, 1965; Roedder, 1977; Bethke e t al., 1976; Barton e t al., 1977; Wetlaufer, 1977; Bethke and Rye, 1979; Woods e t al., 1982; Foley e t al., 1982; Heald-Wetlaufer and Plumlee, 1984; Hayba, 1984; Robinson and Norman, 1984; Plumlee and Hayba, 1985). These studies have been complemented, and t h e models developed from them improved, by detailed, unpublished geologic studies by t h e Homestake Mining Company, Pioneer Nuclear Corporation, Chevron Resources, Freeport Mining Company, Minerals Engineering Company, and by a number of theses (Cannaday, 1955; Chaffee, 1967; Hull, 1970; Giudice, 1980; Robinson, 1981; Battory, 1982; McCrink, 1982; Wason, 1983; Rice, 1984; Horton, 1983; Vergo, 1984; and Misantoni, 1985). Geologic and Mineralogic Characteristics Volcanic-hydrothermal history--The San Juan Mountains a r e t h e main erosional remnant of a volcanic field which covered most of t h e southern

Rocky Mountains in mid-Tertiary time. Approximately 213 of t h e volume of t h e present field is composed of a series of early lavas and volcaniclastic aprons, mainly of andesitic composition, related t o a number of stratovolcanic centers. These "Early Intermediate" composition lavas a r e overlain by a series of q u a r t z l a t i t i c t o rhyolitic ash-flow sheets. F i f t e e n calderas have been identified a s sources of these ash-flows (Fig. 7.1). Many of these calderas a r e nested t o form complexes such a s t h e c e n t r a l San J u a n caldera complex which hosts t h e C r e e d e mining district. The C r e e d e mining district is located along t h e western e d g e of t h e central San Juan caldera complex, a series of 7 nested calderas from which quartz-latitic t o rhyolitic ash-flow s h e e t s erupted over t h e brief interval of 27.6 t o 26.2 m.y. (Fig. 7.2). Only 5 of t h e 7 calderas can b e accurately located, t h e remaining t w o having caved into the younger C r e e d e caldera. The C r e e d e o r e s a r e hosted by t h e intracaldera fill of t h e resurgent Bachelor caldera, t h e second caldera in t h e series. The o r e s a r e contained in a s e t of f r a c t u r e s comprising a graben running between t h e C r e e d e caldera, t h e youngest of t h e series, and t h e slightly older San Luis caldera. Radiometric dating of vein adularia and mixedlayer clay alteration minerals (Bethke et al., 1976; Vergo, 1984) indicates t h a t t h e ores were deposited approximately 1 million years a f t e r t h e youngest dated volcanic event in the district. These d a t e s a r e consistent with t h e observation of Steven and R a t t e (1960b) t h a t mineralization was confined t o s t r u c t u r e s young enough t o c u t t h e sediments of t h e C r e e d e formation which fill t h e m o a t of t h e resurgent C r e e d e caldera. Studies of the C r e e d e Formation by Steven and R a t & (19651, Steven and Van Lonen (19701, McCrink (1982), and Battory (1982) have shown t h a t t h e sediments filling t h e m o a t of t h e C r e e d e caldera consist of landslide debris and s t r e a m channel fill on t h e margins of t h e moat, and lacustrine deposits of airfall and water-lain tuffs, interbedded with a f e w thin ash-flow tuffs, in t h e c e n t e r of t h e moat. These sediments accumulated in a shallow, playa lake environment, and t h e lacustrine sediments were strongly zeolitized during diagenesis. Steven and Eaton (1975) proposed t h a t t h e convecting hydrothermal system was driven by a n unexposed pluton beneath t h e C r e e d e district. They observed t h a t t h e maximum displacement along t h e Amethyst and Bulldog fault systems occurs in t h e central, most heavily mineralized p a r t of t h e district, where subtle magnetic and gravity anomalies a r e suggestive of a buried intrusion. C r e e d e o r e deposits--The base- and preciousm e t a l mineralization a t C r e e d e fills open fractures. The o r e zone, which occupies a narrow vertical range from 250 t o 400 m, has been mined nearly continuously for approximately 3 km along strike on t h e AmethystOH vein system, and for over 2 km on t h e Bulldog Mountain vein system (Fig. 7.3). This horizontallydominated aspect ratio is characteristic of deposits of t h e Adularia-Sericite-type. Mineralization is confined t o t h e f r a c t u r e system comprising t h e C r e e d e graben bounded on t h e e a s t by t h e Solomon-Holy Moses fault system and on t h e west by t h e Alpha-Corsair system

(Fig. 7.3). This graben system follows closely t h e structures of a n older graben interpreted t o b e t h e keystone graben of t h e resurgent Bachelor caldera. Detailed o r e petrology studies of material from t h e major producing s t r u c t u r e s (Amethyst, OH, P and Bulldog Mountain vein systems) indicate t h a t a l l were part of a single hydrologic system and were filled during t h e s a m e mineralizing e v e n t (Barton e t al., 1977; Bethke and Rye, 1979; Heald-Wetlaufer and Plumlee, 1984). A significant, low-grade, bulk-tonnage silver resource also exists a s disseminated replacements in t h e wallrocks of t h e upper parts of the Amethyst vein system near i t s intersection with t h e OH vein (Giudice, 1980). Additional bulk-tonnage resources a r e present a s mineralized s t r e a m sediments in t h e s t r e a m channel facies of t h e C r e e d e Formation adjacent t o i t s

Figure 7.3. Generalized geology of the Creede mining district, modified from Steven and Eaton (1975). Area of map is shown in Faults are dashed where uncerFigure 7.2. tain or inferred; bar and ball show the downdropped side. Tc, Creede Formation; Tf, Fisher quartz latite flow; Tfi, volcanic neck of Fisher quartz latite.

D.0 . HAYBA, P. M.

BETHKE,

P.HEALD, & N.K.

FOLEY

NORTH

SOUTH

present

surface

? OH section

Amethyst section

OH assemblege F i g u r e 7.4. G e n e r a l i z e d l o n g i t u d i n a l p r o j e c t i o n o n t o a v e r t i c a l p l a n e o f t h e OH v e i n a n d t h e s o u t h e r n e n d o f the Amethyst vein. The main workings a n d the g e n e r a l d i s t r i b u t i o n o f the OH a n d B u l l d o g a s s e m b l a g e s are shown. T h e " s t r u c t u r a l d i s c o n t i n u i t y " marks the p o s i t i o n t h a t would h a v e b e e n the i n t e r s e c t i o n o f the OH and Amethyst v e i n s had f r a c t u r i n g b e e n more c o n t i n uous. T h e A m e t h y s t v e i n has b e e n m i n e d n o r t h w a r d f o r a b o u t 2 km b e y o n d the " s t r u c t u r a l d i s c o n t i n u i t y " but is n o t shown here. Adapted from B a r t o n et al. (1977).

truncation by t h e Amethyst vein (Wason, 1983; Rice, 1984). The metals have been introduced into t h e clastic facies of t h e C r e e d e F o r ~ n a t i o nby leakage of t h e o r e fluid from underlying or adjacent veins. Mineralogic characteristics--Mineralization in t h e C r e e d e veins is strongly zoned from a n association in t h e north (OH Assemblage) of chlorite + h e m a t i t e + quartz + adularia + sphalerite + galena + chalcopyrite + pyrite +I- fluorite and t e t r a h e d r i t e t o a barite + rhodochrosite + q u a r t z + adularia + galena + sphalerite + fluorite + t e t r a h e d r i t e + silver sulfosalt + native silver association t o t h e south (Bulldog Assemblage) (Fig. 7.4). Heald-Wetlaufer and Plumlee (1984) have dernonstrated t h a t t h e Bulldog Mountain and OH-PAmethyst vein systems e a c h contain both t h e OH and Bulldog assemblages, and t h a t t h e two a r e contemporaneous and related t o e a c h other through facies changes along t h e fluid-flow path. The mineralization in e a c h of these f a c i e s may be divided into 5 stages on t h e basis of mineralogy and texture, and e a c h of t h e stages can be correlated between t h e t w o assemblages. Details of t h e mineralogy and t h e correlation of stages between t h e t w o assemblages a r e given in Table 7.4. A l a t e association of covellite + chalcocite + a c a n t h i t e has been described by Giudice (1980) and Robinson and Norman (1984) in t h e disseminated ore from t h e upper and southern Amethyst vein system. This assemblage is a volumetrically minor, but economically important, component of t h e Bulldog Assemblage. It has not yet been incorporated into t h e geochemical model for C r e e d e ore deposition, because its limited occurrence and fine-grained nature makes i t difficult t o study. Many minerals of t h e C r e e d e o r e s (sphalerite, rhodochrosite, siderite, Fe-chlorite, tetrahedritetennantite, proustite-pyrargyrite, l a t e gel-pyrite, and, probably, other sulfosalt minerals) exhibit marked compositional variations in both t i m e and space. Of these minerals, only sphalerite (Barton, e t al., 19771, t h e Mn-Fe-carbonates (Wetlaufer, 1977) and Fe-

chlorite (Horton, 1983) have been studied extensively. The co~npositionalvariations in t h e Mn-Fe carbonates and Fe-chlorite have proven too complex t o be very useful t o developing a g e n e t i c model of Creede, but those in t h e sphalerite (primarily iron content) have provided t h e main basis for documenting t h e evolution of t h e o r e fluid in t i m e a n d space, and for constraining t h e chemical environment. The sphalerite is beautifully banded, reflecting, mainly, t h e variations in iron content. This banding records a wealth of information about crystal growth and dissolution. Even more importantly, t h e growth zones can be correlated along t h e OH vein and between t h e OH and other veins. Figure 7.5 shows a composite microprobe tracing of samples of sphalerite from t h e OH vein t h a t is representative of most of t h e B and D s t a g e paragenesis. Hydrothermal leaching (intense for barite, fluorite and rhodochrosite, and less intense for sphalerite and galena) demonstrates t h a t a t times t h e solutions entering t h e o r e zone were undersaturated with respect t o s e l e c t e d components. The increased intensity of leaching and number of leaching horizons a t t h e north end of t h e OH vein suggest t h a t t h e solutions e n t e r e d t h e o r e zone in t h e north and traversed t h e vein system from north t o south. Supergene oxidation has a f f e c t e d t h e C r e e d e ores t o a limited e x t e n t producing extensive manganese oxide (presumably a f t e r rhodochrosite) in t h e southern Amethyst system and local occurrences of kaolinite and halloysite. Some of t h e oxidized rocks contain distinctive f r a c t u r e coatings of dendritic native silver. Veinlets of fine-grained alunite which crosscut mineralization in t h e upper levels of t h e Amethyst vein system have been dated by potassiumargon techniques and range in a g e from 3.5 t o 5.0 million years (M. Lanphere, U.S.G.S., personal communication, 1981). This age corresponds closely t o t h e regional uplift of t h e southern Rocky Mountains which, presumably, raised t h e C r e e d e orebodies above t h e water table. I t is n o t clear whether any or all of

T a b l e 7.4--Mineralogy and p a r a g e n e t i c s t a g e s of t h e OH and Bulldog a s s e m h l a g e s of t h e Creede d i s t r i c t based on t h e work of P. B. Barton and P. Y. Bethke a s r e p o r t e d i n Bethke and Rye ( 1 9 7 9 ) f o r t h e OH, P, and Amethyst v e i n s y s t e m s , on t h a t of Robinson (1981) and G i u d i c e ( 1 9 8 0 ) f o r t h e upper and s o u t h e r n Amethyst v e i n , and on t h e d e t a i l e d s t u d y b y G. S. Plumlee and P. Heald on t h e Bulldog Mountain v e i n s y s t e m a s r e p o r t e d i n Heald-Wetlaufer and Plumlee (1984).

Stage

OH Assemblage (OH, P n o r t h e r n Amethyst v e i n s and n o r t h e r n 1/3 Bulldog v e i n system)

E F i b r o u s p y r i t e w i t h some ( y o u n g e s t ) m a r c a s i t e and s t i b n i t e

Bulldog Assemblage ( s o u t h e r n 2 / 3 Bulldog v e i n system, s o u t h e r n m o s t OH v e i n and s o u t h e r n Amethyst v e i n )

Stage

F i b r o u s p y r i t e w i t h minor m a r c a s i t e , s t i h n i t e and s u l f o s a l t s . L a t e Ag (may b e supergene )

V

D

R e l a t i v e l y coarse-grained s p h a l e r i t e , galena, chalcop y r i t e , and q u a r t z , some hematite; s i l v e r minerals notably a b s e n t ; subdivided i n t o t h r e e s u b s t a g e s on b a s i s of c o l o r banding i n s p h a l e r i t e : i n n e r l i g h t , m i d d l e d a r k , and o u t e r light

Coarse-grained s p h a l e r i t e and g a l e n a w i t h minor l a t e c h a l c o p y r i t e , t e t r a h e d r i t e and Ag-Cu s u l f i d e s and s u l f o s a l t s

IV

C

V o l u m e t r i c a l l y minor s i d e r i t e m a n g a n o s i d e r i t e and f l u o r i t e on q u a r t z ; s i t s on deep e t c h of e a r l i e r B s t a g e ; most f l u o r i t e deeply etched commonly c o m p l e t e l y removed

V o l u m e t r i c a l l y minor (Mn,Fe)c a r b o n a t e and f l u o r i t e on w h i t e and a m e t h y s t i n e q u a r t z ; most f l u o r i t e etched commonly c o m p l e t e l y removed

111

R e l a t i v e l y fine-grained sphale r i t e , galena, chalcopyrite, c h l o r i t e , hematite, p y r i t e , and some t e t r a h e d r i t e tennantite

I n t e r b a n d e d b a r i t e and f i n e g r a i n e d s p h a l e r i t e , g a l e n a , and tetrahedrite

II

P r i m a r i l y q u a r t z w i t h minor c h l o r i t e and s u l f i d e

Early massive rhodochrosite, l a t e r b a r i t e ; minor d i s s e m i n a t e d s p h a l e r i t e and g a l e n a

I

-

B

A (oldest)

t h e wire and leaf native silver, which is c h a r a c t e r i s t i c of t h e Bulldog assemblage, is of supergene origin. Our current interpretation is t h a t i t is primary. wallrock a l t e r a t i o n - - ~ a l l r b c k alteration a t C r e e d e consists of a feldspar-destructive, mixed-layer illite/smectite alteration present along t h e top of t h e orebodies." Alteration intensity ranges from weak, where only t h e pumice fragments a r e altered, t o intense, where the entire rock has been a l t e r e d and where all evidence of primary volcanic t e x t u r e has been obliterated (Horton, 1983). The alteration is strongly fracture-controlled, and t h e intensity of argillization decreases away from t h e veins.

-

*Heald e t al. (1986) included such mixed-layer illites m e c t i t e alteration in t h e sericite category, a n d we will use t h e t e r m "sericite" in their sense t o include both illite and illite-smectite mixed-layer clay minerals.

The most intense alteration forms a "clay cap" t h a t marks t h e upper limit of mining. The vertical e x t e n t of the clay c a p has not been well established, but surface outcrop above highly a l t e r e d a r e a s o f t e n shows no alteration effects, and drilling above t h e OH

D. 0. HAYBA, P.M. BETHKE,P.HEALD,& N. K . FOLEY

]

-I

1

SECTION CONTINUES TO THE RIGHT

70

-

1

Youngest

I

--

Oldest

SECTION CONTINUED FROM THE LEFT

MOLE PERCENT FeS in SPHALERITE

Figure 7.5. Coarse-scale plot of the iron content across a growth-banded aggregate of sphalerite crystals, modified from Barton et al. (1977). Data are taken from composite microprobe tracings of samples from the north end of the OH vein and are replotted to eliminate instrumental scatter. A). tracing of the iron content across finely banded, early (B-stage), undifferentiated sphalerite: B). tracing of the iron content across coarse, late (D-stage) sphalerite that is subdivided on the basis of color and iron and cadmium contents.

vein is reported t o have encountered only fresh rock approximately 100 m above t h e base of alteration. Along t h e OH vein, t h e clay cap is continuous and regularly distributed (Fig. 7.6). In other parts of t h e district, t h e distribution of t h e clay cap is much more irregular; t h e base of alteration undulates over a vertical interval of several hundred f e e t , and p a r t s of t h e veins show no alteration a t all a t any exposed level. T h e continuity and regularity of t h e base of t h e clay c a p above t h e OH vein is interpreted t o r e f l e c t a m o r e uniform hydrologic regime in t h e vertical, simple tension f r a c t u r e t h a t contains t h e vein, in contrast t o t h e hydrology in t h e more complex s t r u c t u r e s of t h e other vein systems in t h e district. Barton e t al. (1977) interpreted t h e clay or sericitic alteration t o mark a

region of recondensation of acid-rich s t e a m derived by boiling of t h e o r e fluid. In t h e upper parts of t h e vein systems, particularly near t h e junction of t h e Amethyst and OH veins, t h e wallrock adjacent t o t h e veins is highly silicified and sometimes bleached. The silicification appears t o pre-date t h e sericitic alteration and t o have protected t h e a f f e c t e d wallrocks from it. In a r e a s of intense silicification, t h e sericitic alteration o f t e n borders t h e silicified zone, although t h e silicification may pass into fresh, unsilicified wallrock, particularly in a r e a s lying beneath t h e sericitic cap. I t is suggested, but not demonstrated, t h a t t h e silicification predated t h e sericitic alteration and protected t h e walls from t h e altering fluids.

142

CHAPTER 7

*

Present surface

-

-

OH section

0

NORTH

0

200

800

-

400 METERS

EXPLANATION lllite alteration Boiling, fluid inclusions Adularia localities

Amethyst sectionSOUTH

1600 FEET

G e n e r a l i z e d l o n g i t u d i n a l p r o j e c t i o n o n t o a v e r t i c a l p l a n e o f t h e OH v e i n a n d t h e F i g u r e 7.6. s o u t h e r n e n d o f the Amethyst vein. The base o f the i n t e n s e sericitic a l t e r a t i o n is i n d i c a t e d : t h e p o s i t i o n o f the top is unknown. Detailed studies of t h e variation in structure and composition of t h e mixed-layer clay minerals by Horton (1983) and Vergo (1984) show t h a t t h e illite component of t h e clays ranges from greater than 95% t o less than 60%. In t h e a r e a just north of t h e Amethyst-OH intersection, where t h e ground between t h e t w o is c u t by many smaller veins, t h e proportion of illite in t h e mixed-layer clays increases toward t h e Amethyst vein. The d e g r e e of ordering of t h e stacking sequence of t h e illite and s m e c t i t e layers, which ranges from random t o long-range ordered (Reichweite = 0 t o 31, also increases toward t h e Amethyst vein. Using empirical relationships between t e m p e r a t u r e and t h e structure and composition of mixed-layer illites m e c t i t e clays, Horton (1983) e s t i m a t e d t h e position of isotherms surrounding t h e upper portions of t h e Amethyst vein near i t s intersection with t h e OH vein. Vergo (1984) found t h e mixed-layer clays along t h e Bulldog vein system t o b e m o r e smectite-rich than most of Horton's samples (presumably because they formed a t lower temperatures). Vergo did not find any systematic variation in t h e illite-smectite r a t i o with distance from t h e vein for distances up t o 70 m from t h e vein, and calculated t h a t t h e system must have been active for a period of a t l e a s t 10,000 years t o produce such a n apparently f l a t t h e r m a l gradient. Below t h e zone of intense sericitic alteration, t h e wallrocks enclosing t h e veins a r e essentially fresh, but were enriched in potassium by a period of potassium feldspar-stable, metal-barren hydrothermal alteration t h a t occurred approximately 2 million years prior t o ore deposition ( R a t t i and Steven, 1967; Bethke e t al., 1985). This intense potassium metasomatism was presumably related t o hydrothermal circulation in t h e keystone graben of t h e Bachelor caldera, and is similar in most respects t o t h e potassic alteration a t Bodie (OINeil e t al., 1973; Silberman and Berger, 1985, this volume). I t does not appear t o have been associated with any earlier period of mineralization. Geochemical Environment An

extremely

detailed

knowledge

of

the

geochemical environment of t h e C r e e d e ore-forming system has been developed through fluid-inclusion, light-stable isotope, and lead-isotope studies, and through thermochemical analysis of t h e o r e and gangue mineral assemblages. A summary of t h e general characteristics of t h e geochemical environment a r e shown in Table 7.5. Temperature-salinity ranges--Much of the evidence used in defining t h e depositional p a r a m e t e r s of t h e C r e e d e system has c o m e from t h e study of fluid inclusions. Numerous homogenization measurements have defined a t e m p e r a t u r e range of 120' t o 2 8 0 ' ~ for t h e C r e e d e ore-forming fluids. Most of t h e measurements have been on sphalerite, q u a r t z and fluorite from t h e OH-Amethyst vein system (Woods e t al., 1982; Robinson and Norman, 1984; J. Goss, U.S.G.S., personal communication, 1985) and on barite from t h e mineralized s t r e a m channel of t h e C r e e d e Formation near i t s truncation by t h e Amethyst vein (Rice, 1984). There is a clear trend from higher t e m p e r a t u r e s in t h e north (OH data, a s summarized by Woods et al., 1982) t o lower t e m p e r a t u r e s in t h e south (southern Amethyst d a t a , Robinson and Norman, 1984, and C r e e d e Formation d a t a , Rice, 1984) although t h e t e m p e r a t u r e ranges f o r minerals and veins overlap (Fig. 7.7). Fluid-inclusion studies have also shown t h a t t h e C r e e d e ores were deposited from relatively concentrated NaCl brines and t h a t these brines mixed with overlying ground w a t e r in t h e o r e zone (discussed below). From freezing and crushing and leaching studies, Roedder (1963, 1965) showed t h a t t h e o r e fluid ranged in salinity from about 4 t o 12 wt.-% NaCl equivalent a n d had a v e r a g e a t o m i c ratios of Na:K:Ca of 9:1:2. These ratios a r e in excellent agreement with those calculated f o r a t e m p e r a t u r e of 2 6 0 ' ~ using t h e alkali geothermometer of Fournier and Truesdell (1973). As noted in t h e previous section, with t h e exception of Colqui, Peru, the salinity of t h e C r e e d e ore fluids is much hinher than all t h e other AdulariaSericite-ty pe deposits-evaluated by Heald e t al. (1 986). Depth-pressure ranges--In addition t o t h e t e m p e r a t u r e and salinity d a t a , fluid-inclusion studies

D. 0. HAYBA,P. M. BETHKE,P. HEALD, & N. K. FOLEY T a b l e 7.5--General e n v i r o n m e n t a l p a r a m e t e r s f o r t h e OH v e i n , C r e e d e , Colorado ( m o d i f i e d from B a r t o n e t a l . , 1977) Range observed

Reference environment

Temperature

190 - 2 8 5 ' ~ 40 - 50 b a r s

250'~ 50 b a r s

Depth

450

600 m

500 m

Salinity Na:K

4 - 12 wt% 7.4 - 9.9

6 wt% 9

Total S

0.018-0.30

Parameter

-

molal

0.02 mola12

S o u r c e of information Fluid inclusions1 E v i d e n c e of b o i l i n g i n fluid inclusions E s t i m a t e d from p r e s s u r e and g e o l o g i c r e c o n s t r . Fluid inclusions A n a l y s e s of f l u i d inclusions Analyses of f l u i d inclusions Calculated

' ~ o s t o f t h e f l u i d i n c l u s i o n e v i d e n c e i s from t h e l a t e r h a l f of t h e m i n e r a l i z a t i o n which i s much c o a r s e r g r a i n e d . 2 ~ e c a u s eo f t h e problem of s u l f u r c o n t r i b u t e d by o x i d a t i o n o f s u l f i d e s d u r i n g sample h a n d l i n g , t h e lower t o t a l s u l f u r v a l u e s a r e c o n s i d e r e d more r e l i a b l e .

have shown t h a t t h e o r e fluid was, a t t i m e s , boiling n e a r t h e t o p of t h e OH vein ( t h e evidence f o r boiling is discussed later). Using t h e t a b l e s f r o m H a a s (1971), t h e pressure a t t h e t o p of t h e o r e body was approxim a t e l y 40 bars, based o n t h e boiling point of a 250°c, 1 molal NaCl fluid. F o r m o s t of o r e deposition, t h e pressure must h a v e been slightly g r e a t e r t h a n this, because boiling a p p e a r s t o h a v e o c c u r r e d infrequently. A h y d r o s t a t i c h e a d of approximately 450 m is necessary t o p r o d u c e t h e requisite pressure, which is in e x c e l l e n t a g r e e m e n t with t h e geologic r e c o n s t r u c t i o n by S t e v e n a n d E a t o n (1975). C h e m i c a l parameters--In 1977, Barton e t al. did a t h e r m o c h e m i c a l analysis of t h e C r e e d e s y s t e m a n d w e r e a b l e t o put l i m i t s on such p a r a m e t e r s a s a c t i v i t i e s of S2 a n d 0 2 , pH, a n d t o t a l sulfur. In t e r m s of t h e activities of S 2 and O 2 (Fig. 7.81, t h e C r e e d e environment is l o c a t e d n e a r t h e junctures of t h e h e m a t i t e , p y r i t e a n d Fe-chlorite stability fields based on t h e c o m m o n o c c u r r e n c e of t h a t mineral assemblage and on t h e iron c o n t e n t of s p h a l e r i t e which usually ranged b e t w e e n 1 a n d 4 mole-% F e S (Fig. 7.5). Because of t h e c o m m o n o c c u r r e n c e of a d u l a r i a associated with minor a m o u n t s of s e r i c i t e in t h e wall rock near t h e vein, Barton e t al. (1977) e s t i m a t e d t h e pH of t h e fluid during mineral deposition in t h e C r e e d e s y s t e m based on t h e feldspar hydrolysis r e a c t i o n

A pH of 5.4 is e s t i m a t e d f r o m t h e t h e r m o d y n a m i c d a t a of Montoya a n d Hemley (1975) f o r a 1 molal solution with a Na/K r a t i o of 9. This i s a nearly n e u t r a l pH a t 250'~. Roedder e t al. (1963) m e a s u r e d t o t a l sulfur c o n c e n t r a t i o n s of about 0.02 molal in inclusion fluids

QUARTZ & BARITE

Figure 7.7. Histograms of homogenization temperatures for fluid inclusions in quartz and sphalerite from the OH vein (Woods et al., 1982), quartz from the southern Amethyst vein (Robinson and Norman, 1984), and barite from the Creede Formation (Rice, 1984).

In order t o determine if t h e s y s t e m a t i c variation of Th with iron content is due t o a mineralogically reasonable buffer, we will examine t h e following reaction which could buffer t h e activity of sulfur in t h e ore-forming solution 3 daphnite

+ 3 K - f e l d s p a r + 5 S2

= 5 hmtite

+ 5 pyrite + 3 h i c a

+ 9 q u a r t z + 9 H20

~ F i g u r e 7.8. Log aS2-%2 d i a g r a m a t 2 5 0 ~showi n g the m i n e r a l s t a b i l i t y f i e l d s f o r the s i g n i f i c a n t m i n e r a l s i n the C r e e d e ores. The shaded f i e l d o f m a g n e t i t e is c o m p l e t e l y p r e e m p t e d by chlorite. The c o n t o u r f o r 20 m o l e p e r c e n t F e S i n sphalerite c o i n c i d e s w i t h the pyrrhotite f i e l d boundary. Quartz is p r e s e n t t h r o u g h c u t the diagram. Abbrev i a t i o n s : py = pyrite, ccp = chalcopyrite, b n = b o r n i t e , h e m = h e m a t i t e , chl = chlor i t e . The s t a n d a r d s t a t e f o r S 2 a n d O2 i s the i d e a l d i a t o m i c g a s a t 1 a t m o s p h e r e and 2 5 0 ~ ~ The . d a t a f o r the i r o n - c o p p e r s u l f i d e s a n d o x i d e s are s u m m a r i z e d i n B a r t o n a n d S k i n n e r ( T a b l e 7.2, 1 9 7 9 ) ; t h i c k s o l i d lines: boundaries b e t w e e n p y r i t e , p y r r h o t i t e , chlorite, a n d hematite; l o n g l i g h t dashes: i r o n c o n t e n t o f s p h a l e r i t e . After B a r t o n e t a l . (1977).

from Creede, but other concentrations in different samples a r e possible. However, a t this t o t a l sulfur concentration (0.02 m), the pH of t h e pyrite + h e m a t i t e + Fe-chlorite triple point is near 5.4, which lends credence t o t h e thermodynamic e s t i m a t e of pH (Fin. " 7.9). Chemical buffering of t h e o r e fluids--Barton et al. (1977) proposed t h a t reactions between iron-rich chlorite, hematite, pyrite, quartz and water controlled t h e redox conditions for t h e OH vein. A means of examining t h a t suggestion arises from Woods et al.'s (1982) observation t h a t t h e iron c o n t e n t s of growthbanded sphalerite from t h e OH vein show a positive correlation with t h e homogenization t e m p e r a t u r e s (Th) of primary fluid inclusions contained in t h e growth bands (Fig. 7.10). Previous investigators (Barton and Toulmin, 1964; S c o t t and Barnes, 1971) have shown t h a t t h e iron content of sphalerite in a pyrites a t u r a t e d system is a function solely of t e m p e r a t u r e and t h e activity of sulfur (the role of pressure may safely b e neglected for these shallow deposits), based on t h e definitive equation f o r t h e iron in sphalerite in pyrite-saturated system given below FeS ( i n sph) + 112 S2 = FeS2 ( p y )

(2)

(3)

This reaction differs from t h a t of t h e proposed buffer of Barton et al. (1977) in t h a t t h e Fe-chlorite in this reaction (3) is daphnite (Fe5AI Si 0 (OH)g) rather than a n aluminum-free end rnemse?. '8aphnite has a more realistic c o m ~ o s i t i o nwhich closelv a ~ ~ r o x i m a t e s t h a t of t h e c r e e h e chlorite ( ~ m m o n s:Ad Larsen, 1923). However, using daphnite in t h e reaction requires t h a t other aluminum-bearing phases be considered t o balance t h e reaction. The t w o most logical choices for C r e e d e a r e K-feldspar and sericite, t h e t w o phases which Barton et al. (1977) suggest controlled t h e pH of t h e system (discussed above). In order t o s e e if reaction (3) predicts t h e observed correlation of t e m p e r a t u r e with sphalerite iron content, i t is necessary t o (1) know t h e change in t h e f r e e energy of t h e reaction with t e m p e r a t u r e so t h a t t h e activity of sulfur can b e estimated, and (2) use t h a t sulfur activity t o calculate t h e iron content of sphalerite a t t h a t temperature. Estimating t h e change in t h e f r e e energy of t h e reaction is limited by t h e lack of f r e e energy d a t a on t h e daphnite component. However, since we a r e only trying t o predict t h e change in t h e activity of sulfur with t e m p e r a t u r e for this buffer reaction, t h e accuracy of t h e f r e e energy d a t a is not a s important a s knowing how i t changes with temperature; Hemingway et- al. (1984) have recently measured t h e h e a t capacity of two natural chlorites (one Fe-rich, t h e other Mg-rich), and their d a t a a g r e e well with Helgeson e t al.'s (1978) h e a t capacity d a t a for daphnite between 2 0 0 ~ - 3 0 0 ~ ~ . Although t h e heat-capacity d a t a allow us t o calculate t h e change in t h e f r e e energy with temperature, in order t o use those d a t a i t is still necessary t o e s t i m a t e t h e chemical potential of daphnite a t o n e temperature. Therefore, we have used t h e e s t i m a t e of Barton e t al. (1977) of t h e chlorite + pyrite + h e m a t i t e triple point a t 2 5 0 ' ~ a t approximately -11.0 f o r log S2 activity and -34.2 for log O2 activity. Using these values, the chemical potential a t 2 5 0 ' ~ for t h e daphnite component of t h e C r e e d e chlorite is e s t i m a t e d t o be -1462 kcal/mole. Using t h e activity of sulfur e s t i m a t e d from t h e above thermodynamic calculation a t a given temperature, the iron c o n t e n t of the sphalerite can be calculated using t h e following equation l o g X ( p y ) = 6.809

-

7340/T

-

0.5 l o g a s

2

(4

This equation which is a numerical expression of reaction (2) has t h e s a m e slope a s S c o t t and Barnes (1971) equation, but t h e i n t e r c e p t has been changed t o a g r e e with Czarnanske's (1974) measurements on t h e iron content of sphalerite in equilibrium with pyrite +

D.0. HAYBA, P.M. BETHKE,P.HEALD,& N.K.

FOLEY

145

important t o n o t e t h a t t h e line on Figure 7.10 was forced t o go through t h e d a t a a t 2 5 0 ' ~ by our e s t i m a t e of t h e f r e e energy of daphnite, but t h a t t h e slope of t h e line was determined by t h e heat-capacity data. The s c a t t e r in t h e d a t a in Figure 7.10, other than can be a t t r i b u t e d t o analysis and correlation error, is presumbably due t o perturbations of t h e chemical system away from t h e buffered environmenf, especially a t t h e higher iron concentrations, a s discussed by Barton et al. (1977).

Figure 7.9. Log aO2 - pH diagram at 2 5 0 and ~ ~ total sulfur = 0.02 molal showing the mineral stability fields pertinent to the Creede ores (modified from Barton et al., 1977). The salinity is 1 molal, with Na+/K+ = 9. Dotted lines: boundaries between aqueous sulfur species; thick solid lines: boundaries between pyrite, pyrrhotite, chlorite, and hematite; short dashes: limit of stability of chalcopyrite; long light dashes: iron content of sphalerite; medium solid lines: boundaries between kaolinite, muscovite, and potassium feldspar. Abbreviations: py = pyrite, bn = bornite, ccp = chalcopyrite, chl = chlorite. bornite + chalcopyrite, a s discussed by Barton et al. (1977, p. 10). In order t o directly r e l a t e t h e iron c o n t e n t of t h e sphalerite t o t h e buffer, we can combine equations (2) and (3) t o g e t t h e following reaction 3 daphnite

+ 3 K-feldspar + 5 p y r i t e

+ 9 q u a r t z + 9 H20

(5)

Figure 7.10 shows t h a t t h e predicted iron content fairly closely m a t c h e s t h e between 200-280°C measured data, indicating t h a t t h e ore-forming system may have been indeed buffered by reaction 13). I t is

Sources of metals--Lead-isotope studies reported by Doe e t al. (1979) and unpublished d a t a of Foley (U.S.G.S., 1985) show t h a t t h e lead-isotopic composition of galena from t h e OH, Amethyst and Bulldog Mountain vein systerns is remarkably uniform and is more radiogenic than t h a t of any volcanic rock in t h e San Juan Mountains, or for t h a t m a t t e r , than any Mesozoic or Cenozoic rock from t h e e n t i r e Rocky Mountain region measured t o date. Doe e t al. (1979) conclude t h a t this requires t h e bulk of t h e lead in t h e C r e e d e system t o b e derived by leaching of 1.4 t o 1.7 billion-year-old Precambrian rock underlying t h e district. Surprisingly, t h e lead-isotopic composition of galenas from t h e Alpha-Corsair f a u l t system and from t h e mineralized C r e e d e Formation a t Monon Hill, nearly on strike of t h e Alpha-Corsair f a u l t system, a r e less radiogenic than t h e galenas from t h e main part of t h e district. These galenas appear t o have a much larger component of lead from t h e volcanic rocks, suggesting t h a t t h e lead deposited along t h e AlphaCorsair system was derived from a different, possibly more shallow, lead reservoir than t h a t of t h e Amethyst-OH-Bulldog Mountain system. Sources of sulfur--Sulfur-isotope studies a t Creede, only partially reported t o d a t e (Bethke e t al., 19731, indicate a complex sulfur history not y e t fully understood. Sulfur-isotopic equilibrium between sulfide minerals and reduced aqueous sulfur species was apparently closely approached. I t is clear, however, t h a t t h e r e was little, if any, sulfur-isotopic exchange between oxidized and reduced aqueous sulfur species in t h e o r e zone, and t h e y appear t o have operated a s s e p a r a t e isotopic (and probably chemical) systems in t h e upper p a r t of t h e hydrothermal system. The narrow range of sulfur-isotopic composition of t h e sulfide minerals is i n t e r p r e t e d t o r e f l e c t a deep, nearly 0 permil, sulfide reservoir, whereas t h e extremely heavy values of 6180 and 6 3 4 ~from sulfate minerals (up t o 19 and 45 permil, respectively) require t h a t a t least a large part of t h e sulfate underwent biogenic reduction in t h e playa lake, t h e presumed reservoir for t h e o r e fluids. Some of t h e sulfate was carried deep into t h e roots of t h e C r e e d e hydrothermal system where t e m p e r a t u r e s were high enough M 5 0 ° c ) f o r i t t o equilibrate with t h e wallrock silicates and t h e 0 permil sulfur reservoir. Most of t h e sulfate did not p e n e t r a t e deeply enough t o a t t a i n a high enough t e m p e r a t u r e t o undergo substantial isotopic exchange. This partially-to-unexchanged s u l f a t e did, however, mix into t h e hydrothermal system, presumably along t h e margins of t h e upwelling brine plume, t o produce t h e exceptionally wide range of sulfur- and oxygenisotopic compositions measured on t h e barites.

Hydrologic Environment The hydrologic environment is a ~ n a j o rf a c t o r in t h e localization of epithermal ore bodies. Unfortunately, i t is also t h e environment which we a r e a t present l e a s t able t o t r e a t quantitatively. However, a number of geologic, isotopic, and fluid-inclusion evidences can b e used t o place some restraints on t h e hydrologic environment of t h e C r e e d e system, and t o develop a qualitative model which can be cornpared

Homoganization Temperature

OC

Figure 7.10. Diagram showing the temperature of homogenization of fluid inclusions vs. the iron content of the host sphalerite growth zone for sample locality NJP-X on the OH vein. The line shows the predicted iron content of the sphalerite if the sulfur fugacity of the system had been buffered by the triple point - Fe-chlorite (daphnite), pyrite, hematite.

with observed hydrologies of active geothermal systems. Figure 7.1 1 shows a generalized hydrologic rnodel for t h e C r e e d e mining district a s i t is presently understood. It is similar in most aspects t o t h e model orginally proposed by Steven and Eaton (1975) based upon geologic grounds, and t o t h a t proposed by Barton e t al. (1977) based upon mineralogical and geochemical evidence. It is also similar in overall aspects t o models of geothermal systems in t h e Taupo Volcanic Zone, New Zealand (cf. Henley and Ellis, 1983). The C r e e d e ores were deposited along t h e top of a saline, deeply circulating hydrothermal system, charged primarily with m e t e o r i c waters, t h a t displaced t h e regional ground water regime. An intrusion underlying t h e district a t a depth of 3 t o 5 km is speculated t o have been t h e h e a t source which provided t h e buoyancy of t h e brine plume. In t h e upper p a r t of t h e system, fluid movement was fracture-controlled and nearly horizontal, from north t o south. T h e o r e zone was overlain by a zone of fresh ground w a t e r , approximately 500 m e t e r s thick, , which flowed southward down t h e regional hydraulic gradient. The base of this ground water zone was heated t o temperatures of about 160°C by h e a t transfer from t h e underlying brine. Precipitation of t h e o r e s a t t h e interface between t h e d e e p brine and t h e overlying ground water appears t o have resulted from t h e dual processes of boiling and mixing. The various evidences supporting this general hydrologic model a r e presented below. Geologic constraints--Steven and Eaton (1975) suggested t h a t t h e circulating hydrothermal system responsible for mineralization was influenced by t w o major lithologic factors. The first was t h e location of soft, non-welded to poorly-welded, relatively impermeable, tuffs along t h e top of t h e ore zone. This

San Luis Caldera

NORTH

SOUTH

Figure 7.11. Schematic representation of the Creede hydrothermal system. Upwelling plume (stip~ ~ displaces regional ground water flow from pled pattern), outlined by the 2 0 0 isotherm, highlands in the north to the low area of the Creede caldera moat to the south. Heat source responsible for buoyancy of plume is shown as stock beneath district (hatched pattern).

D. 0. HAYBA, P.M. BETHKE, P. HEALD, & N. K. FOLEY aquitard largely blocked the upward movement of t h e hydrothermal solutions and forced t h e solutions t o flow laterally t o t h e south. The second lithologic f a c t o r was t h e location of t h e permeable talus-regolith and fanglomerate deposits of the C r e e d e Formation a t t h e southern end of t h e vein system (Figs. 7.3 and 7.11). Where c u t by t h e flow-controlling fractures, these coarse clastics provided a n outlet t o t h e south for t h e hydrothermal solutions. The mineralogically and geochemically based hydrologic model of Barton e t al. (1977) is nearly identical t o t h a t of Steven and Eaton (1975). Barton e t al. describe t h e system a s a freely convecting hydrothermal cell t h a t e x t r a c t e d metals and sulfur from sources a t depth and deposited gangue and o r e minerals near t h e t o p of t h e system. They a t t r i b u t e o r e deposition t o cooling and a slight pH rise due t o boiling of t h e hydrothermal fluid. Recondensation of t h e C 0 2 and H2S, which were strongly fractionated into t h e vapor phase during boiling, in t h e cooler, overlying rocks led t o t h e formation of t h e intensely sericitized c a p above t h e ore (Fig. 7.6). The development of this zone of intense alteration above t h e orebodies undoubtedly served t o increase t h e efficiency of t h e aquitard along t h e top of t h e system. Influence of topography--In addition t o t h e influence of t h e soft-tuff aquitard and sericitic alteration zone, i t is probable t h a t t h e topography a t t h e t i m e of ore-formation played a n important role in maintaining a "cap" of cooler ground water above t h e deeply circulating ore-forming brine. According t o Steven and Eaton (19751, a t t h e t i m e of ore-formation t h e a r e a of t h e San Juan Mountains comprised a widespread volcanic plateau punctuated by regions of rough topography in caldera areas. Local relief in t h e vicinity of C r e e d e approached 1.8 km over horizontal distances of 10 km. The low point was t h e playa lake in t h e m o a t of t h e C r e e d e caldera, and maximum elevations were a t t a i n e d along a string of volcanoes located along t h e present Continental Divide, about 10 km north of t h e geographic center of t h e mining district. This high relief imposed a strong regional hydraulic gradient on t h e ground water from north t o south across t h e o r e zone. As noted by Henley (1985, this volume), such a regional hydraulic gradient will tend t o establish a lateral ground-water flow across t h e top of a geothermal system. Such a n e f f e c t has been evaluated by Hanaoka (1980) using numerical modeling techniques. Hanaokals analysis is simplified in that: ( I ) i t is calculated for pure water; (2) t h e conditions were chosen t o obviate boiling; (3) t h e permeability and thermal conduction-dispersion coefficient were both uniform; and (4) t h e e f f e c t of pressure on viscosity was ignored. Changes in t h e s e variables will modify the quantitative aspects of Hanaoka's results, but, within reasonable bounds, a r e not likely t o a f f e c t t h e general topology of his models. His calculations showed t h a t in a r e a s of m o d e r a t e relief, a two-tiered flow regime may b e produced, consisting of a deep, hot, convecting cell overlain by a cooler region of ground-water flow forming a "hydrologic" c a p t o t h e hydrothermal system. Flow is parallel and l a t e r a l along t h e interface between t h e t w o regimes, and h e a t is transferred from t h e deep t o t h e shallow system by conduction and dispersion.

147

Hanaoka's model is for moderate relief consistent with t h e conditions observed in a number of geothermal systems (Healy and Hochstein, 1973; Healy, 1975; Ellis and Mahon, 1977; Hedenquist, 1983). It is also tantalizingly consistent with t h e evidence for t h e interaction between a hot, deep, saline fluid, and a cooler, overlying fresh ground water a t Creede. Isotopic constraints--Light-stable-isotope studies have shown t h a t in addition t o t h e deep, saline fluids and t h e shallow groundwaters discussed above, fluids from a third, isotopically distinct, probably magmatic, source were involved in t h e C r e e d e mineralization. The evidence for t h e episodic introduction of m a g m a t i c waters into t h e C r e e d e hydrothermal system comes from carbon- and oxygen-isotopic studies of t h e vein carbonates (Bethke and Rye, 1979). These studies indicate t h a t both t h e early (A-stage) rhodochrosite and t h e younger (C-stage) siderite-manganosiderite were deposited from fluids t h a t equilibrated with silicates a t very high t e m p e r a t u r e s (presumably m a g m a t i c fluids). This supposition was based primarily on t h e large 6180 values obtained, but i t is also consistent with t h e carbon-isotopic composition and with t h e hydrogen-isotopic composition of t h e inclusion fluids in a few rhodochrosite samples. The interpretation remains open t o question, however, because i t requires either t h a t t h e vein system was occupied by m a g m a t i c fluids a t t w o different times, separated by a period when i t was occupied by t h e m e t e o r i c w a t e r s from t h e lake sediment reservoir, or t h a t C 0 2 . g i v e n off from t h e magma did not exchange oxygen w ~ t ht h e m e t e o r i c waters t h a t filled t h e vein system. Neither s e e m s a reasonable proposition. Carbonate minerals in many hydrothermal o r e deposits appear t o have been deposited from fluids substantially heavier o r lighter in 6180 than those from which other apparently coeval minerals were deposited (R.O. Rye, U.S.G.S., personal communication, 1985). Perhaps we do not fully understand t h e isotope systematics of C 0 2 in hydrothermal systems. Based on t h e hydrogen- and/or oxygen-isotopic composition of t h e alteration minerals, and chlorite and t h e inclusion fluids in sphalerite, Bethke and R y e (1979) postulated t h a t t h e deep, saline fluids responsible for ore deposition originated a s porew a t e r s in t h e C r e e d e Formation which accumulated in t h e evaporative, closed-basin playa lake in t h e m o a t of t h e C r e e d e caldera. Such a n interpretation is consistent with t h e unusually high salinities of t h e C r e e d e o r e fluids, and with sulfur- and oxygen-isotopic studies on barite (R. Rye, P. Barton and P. Bethke, U.S.G.S., unpublished data, 1985) which suggest t h a t most of t h e sulfate in t h e C r e e d e system must have undergone considerable isotopic evolution in t h e playa lake sediments. The recent work of Foley e t al. (1982) demonstrated t h e existence of a zone of heated, fresh m e t e o r i c water overlying t h e o r e zone. They showed t h a t fluids in primary inclusions in q u a r t z were similar in salinity, homogenization temperature, and hydrogen-isotopic composition (and therefore, presumably, source) t o primary inclusions in sphalerite. This interpretation was contrary t o t h a t originally proposed by Bethke and R y e (1979), which suggested t h a t t h e quartz was deposited from fluids

t h a t originated a s m e t e o r i c waters in t h e high country north of t h e district. Very painstaking sampling and analysis showed t h a t the hydrogen-isotopic analyses of fluid inclusions in quartz reported by Bethke and R y e (1979) were biased by fluids released from pseudosecondary inclusions during analysis. These pseudosecondary inclusions homogenized over a similar but slightly lower t e m p e r a t u r e range a s did t h e primaries, but contained fresh waters whose hydrogen isotopic composition was a t l e a s t 30 permil lighter than t h a t of t h e brine in t h e primary inclusions. Foley et al. (1982) proposed t h a t these light, fresh waters w e r e unevolved m e t e o r i c waters t h a t constituted t h e regional ground waters. These fluids overlay t h e o r e zone, but episodically e n t e r e d i t due t o hydrologic fluctuations during vein filling. Presumably, t h e thermal shock of these heated, but somewhat cooler ground waters, caused f r a c t u r e s in growth-strained q u a r t z crystals which, on rehealing, trapped some of t h e fluids. The lead-isotope studies by Doe e t al. (1979) and unpublished d a t a of Foley (U.S.G.S., 1985), discussed in t h e section on sources of metals, also provide evidence t h a t t h e C r e e d e hydrothermal system circulated t o depths of several kilometers. Their d a t a suggest t h a t t h e bulk of t h e lead in t h e C r e e d e system was derived by selective leaching of radiogenic lead from t h e 1.4 t o 1.7 billiowyear-old Precambrian rock underlying t h e district. The depth t o t h e Precambrian a t C r e e d e (and hence t h e minimum depth of hydrothermal circulation) can only b e estimated, but must be a t l e a s t 2.5-3.0 km written below t h e ore zone (T. Steven, U.S.G.S., communication, 1982). Boiling and Mixing in t h e Ore Zone The most recent modification to the interpretation of t h e hydrology is t h e evidence for mixing of t h e hydrothermal fluid with overlying ground water developed by Hayba (1984) for t h e OH vein and by Robinson and Norman (1984) f o r t h e southern Amethyst vein. Although boiling has been documented in t h e upper levels of t h e OH vein (Roedder, 1970; Woods et al., 19821, i t appears t h a t mixing was responsible for a t l e a s t s o m e of t h e vein mineral deposition. These t w o processes, boiling and mixing, a r e e a c h important mechanisms of mineral deposition; evidences for their roles in t h e C r e e d e hydrothermal system will be discussed below. Boiling--One of the major f a c t o r s influencing t h e hydrologic interpretation of t h e district is t h e f a c t t h a t boiling occurred during a t l e a s t s o m e of t h e depositional history. As noted earlier, t h e r e a r e very few deposits exhibiting unequivocal evidence t h a t boiling occurred during precious-metal deposition. A t Creede, t h e r e is good evidence for boiling, but i t appears t o have taken place during s t a g e D, a basemetal stage, rather than during the precious-metal s t a g e B (Table 7.4). The c a s e for boiling a t C r e e d e is based primarily on t w o lines of evidence: (1) fluid inclusions, and (2) t h e distributions of vein adularia and sericitic alteration. Over 25 years ago, Roedder (1960) discovered some coeval fluid inclusions in D-stage q u a r t z t h a t had widely varying IiquidJvapor ratios, and t h u s a wide range of homogenization temperatures, many over

280°C (about t h e maximum t e m p e r a t u r e for C r e e d e mineral deposition based on d a t a from over 2,500 fluid inclusion measurements). H e interpreted these inclusions a s having resulted from t h e trapping of varying proportions of liquid and vapor from a heterogeneous, two-phase system. This was t h e first such demonstration of boiling in any ore deposit. His interpretation is f u r t h e r bolstered by t h e presence of "empty" or "steam" inclusions in s o m e of these s a m e q u a r t z samples (e.g., Roedder 1970, Fig. 7.71, and by t h e occasional occurrence of o r e s t a l a c t i t e s (indicative of growth in a two-phase regime) protruding into open vugs along t h e OH and Bulldog Mountain vein systems. T o date, evidence for boiling has been found a t six localities along t h e t o p of t h e OH vein (Fig. 7.6) in D-stage quartz. It is notable t h a t Robinson and Norman (1984) found no d i r e c t evidence of boiling in q u a r t z samples from t h e southern Amethyst vein although they carefully looked for indications of boiling. No such evidence has y e t been found in t h e limited number of fluid-inclusion measurements on material from t h e Bulldog Mountain vein system, but t h e material so f a r examined has c o m e from a r e a s on t h a t vein which lie t o t h e south of t h e a r e a on t h e OH from which boiling has been documented. The implication is t h a t boiling was primarily confined t o t h e northern part of t h e district, a t least in t h e later stages of o r e deposition during which most of t h e material studied t o d a t e was deposited. This surmise is consistent with t h e extensive evidence for mixing (discussed below) of t h e o r e fluids with cooler overlying fresh w a t e r s along t h e top of t h e system. To date, fluid-inclusion evidence f o r boiling a t C r e e d e has been found only in q u a r t z crystals, even though i t has been sought carefully in sphalerite and fluorite. It is possible t h a t boiling also occurred during t h e "B-stage" deposition of t h e precious-metal ores, but evidence for boiling during this staged has been carefully looked for and has not been found. The fine-grained nature of B-Stage material s t a g e m a k e s recognition of fluidinclusion evidence for boiling difficult, but t h e lower homogenization t e m p e r a t u r e s for B-Stage inclusions than for D-Stage inclusions is consistent with a lack of boiling. Also shown on Figure 7.6 is t h e distribution of vein adularia and intense sericitic alteration. Adularia occurs mainly below t h e zone of boiling documented by fluid inclusions, while t h e zone of intense sericitic alteration, which caps t h e OH vein, is above t h e zone of documented boiling. The sericitic cap, which is spectacular in i t s contrast and sharpness of c o n t a c t with t h e unaltered wallrock below (the productive p a r t of t h e vein), may have resulted from t h e recondensation of volatiles in t h e overlying, probably fresh waters and from t h e subsequent hydrolysis reactions. At Broadlands, adularia and c a l c i t e deposition has been related t o t h e rise in pH of t h e fluid on boiling, due t o t h e strong fractionation of acid-forming volatiles such a s C 0 2 and H2S into t h e vapor phase (Browne and Ellis, 1970). This p a t t e r n of intense sericitic alteration overlying a zone of vein adularia (+ carbonate in some mining districts) provides indirect, but potentially useful evidence for boiling in epithermal systems. I t m a y b e a particularly useful indicator of boiling in deposits where no material

200

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280

HOMOGENIZATION TEMPERATURE PC)

W

0 K

g

20,000

I

U

water

HEAT CONTENT Jlg

Figure 7.12. a). Plot of homogenization temperature (Th) vs. freezing temperature (Tf) for 221 primary inclusions in a 5-cm band of zoned sphalerite from Creede, Colorado, after Reodder (1977). The numbered areas include all data points from each of the 20 zones sampled, numbered in sequence from w n e 1 (earliest) to zone 20 (latest). The number of inclusions in each of the areas outlined are as follows: 1(2), 2(18), 3(1), 4(21), 5(27), 6(9), 7(4), 8(9), 9(8), 10(4), 11(15), 12(2), 13(7), 14(12), 15(14), 16(11), 17(4), 18(32), 19(13), 20(8). b). Plot of heat content vs. chloride content for primary inclusions (solid triangles) replotted from Figure 7.12a. Data points for surface water and steam are also shown. Tick marks for upper and lower temperature scale are offset because of the effect of chloride content.

a d e q u a t e f o r fluid-inclusion s t u d i e s i s available, o r w h e r e t h e fluid-inclusion e v i d e n c e f o r boiling i s ambiguous. T h e v i r t u a l l a c k of c a l c i t e in t h e C r e e d e d i s t r i c t c a n b e explained by t h e low-calcium c o n t e n t of host potassiuln rocks resulting from the earlier m e t a s o m a t i s m e v e n t , a n d t o t h e l o w - C 0 2 c o n t e n t of t h e fluids during m o s t of t h e depositional history. W e t l a u f e r (1977) h a s shown t h a t t h e fluids responsible

f o r deposition of t h e e a r l y (A-Stage = S t a g e I) Mn-Fe c a r b o n a t e s i n t h e l o w e r p a r t s of t h e s o u t h e r n portions of t h e vein s y s t e m s w e r e similar in t h e i r t h e r m a l history a n d c h e m i s t r y t o t h e l a t e r fluids which deposited t h e bulk of t h e base- a n d precious-metal mineralization e x c e p t t h a t t h e carbonate-depositing fluids had a higher CO-, c o n t e n t . I t i s n o t unreasonable t o suppose t h a t t h e e a r l y Mn-Fe c a r b o n a t e s in t h e C r e e d e vein systerns play t h e r o l e of c a l c i t e in t h e Broadlands g e o t h e r m a l s y s t e m a n d r e c o r d a period of boiling of relatively gas-rich fluids e a r l y in t h e history of vein filling. Mixing--Roedder (1977) showed t h a t t h e r e was a s y s t e m a t i c r e l a t i o n b e t w e e n t e m p e r a t u r e a n d salinity f o r fluid inclusions f r o m t h e O H vein. H e d o c u m e n t e d t h i s relationship b y a d e t a i l e d growth-zone by growthz o n e study of a single, l a r g e s p h a l e r i t e c r y s t a l f r o m a Twenty single locality (NJP-X) o n t h e OH vein. d i s t i n c t g r o w t h z o n e s w e r e d e f i n e d in t h i s c r y s t a l , and homogenization a n d f r e e z i n g t e m p e r a t u r e s w e r e m e a s u r e d o n s e t s of inclusions within e a c h zone. T h e r e s u l t s of his painstaking s t u d y a r e shown in F i g u r e 7.12a. Similar r e s u l t s w e r e o b t a i n e d f o r s e v e r a l d i f f e r e n t l o c a l i t i e s o n t h e O H vein (Woods e t al., 1982). T h e r e a r e m a n y implications of t h e s e s y s t e m a t i c s , p e r h a p s t h e rnost i m p o r t a n t f o r our p r e s e n t purposes being t h e mixing of fluids of d i f f e r e n t t e m p e r a t u r e s a n d salinities. Truesdell a n d Fournier (1975) h a v e shown t h a t plots of h e a t c o n t e n t (enthalpy) a g a i n s t chloride c o n t e n t a r e v e r y useful in e v a l u a t i n g t h e relationship between fluids of different t e m p e r a t u r e s a n d composition in g e o t h e r m a l areas. Both e n t h a l p y a n d c h l o r i d e c o n t e n t a r e additive q u a n t i t i e s s o t h a t t r a j e c t o r i e s f o r boiling a n d mixing a r e linear o n s u c h plots. Roedder's d a t a a r e r e p l o t t e d a s enthalpy-chloride diagrarns in F i g u r e 7.12b. I t c a n be seen from this figure t h a t t h e systernatic relationship b e t w e e n t e m p e r a t u r e a n d salinity c a n b e explained by mixing of saline, h i g h - t e m p e r a t u r e w a t e r s (from z o n e s 8 a n d 9 o n F i g u r e 7.12a) with f r e s h w a t e r s h e a t e d t o a t e m p e r a t u r e of a b o u t 1 6 0 ' ~ . I t i s possible t h a t waters from zones 8 and 9 a r e related t o each o t h e r in t h a t a s m a l l a m o u n t of boiling of z o n e 8 w a t e r s would yield t h e slightly m o r e saline, lower t e m p e r a t u r e w a t e r of z o n e 9. Enthalpy- o r chlorideconservation c a l c u l a t i o n s i n d i c a t e t h a t if t h e s t e a m i s s e p a r a t e d f r o m t h e fluid, only 6 wt.-% of t h e fluid needs t o b e c o n v e r t e d t o s t e a m t o produce w a t e r 9 f r o m w a t e r 8. O n t h e o t h e r hand, in o r d e r t o produce t h e l o w e s t - t e m p e r a t u r e , least-saline fluid ( w a t e r 15), w a t e r 8 would h a v e t o b e m i x e d with m o r e t h a t i t s equivalent w e i g h t of f r e s h w a t e r at 1 6 0 ' ~ . Roedder's d o c u m e n t a t i o n of t h e a b r u p t t e m p o r a l variations i n t h e o r e fluid a l o n g t h e OH s t r u c t u r e w a s a n i m p o r t a n t t i m e c o n s t r a i n t f o r a l a t e r s t u d y on t h e s p a t i a l v a r i a t i o n s i n t h e o r e fluid. Using a d i s t i n c t i v e g r o w t h z o n e i n s p h a l e r i t e a s a time-line t h r o u g h o u t t h e O H vein, H a y b a (1984) d o c u m e n t e d a progressive d e c r e a s e in b o t h t e m p e r a t u r e and salinity f r o m t h e northern, basal e n d of t h e vein t o localities 200 m e t e r s These higher and 1000 m e t e r s f u r t h e r south. t e m p e r a t u r e a n d s a l i n i t y g r a d i e n t s a r e i n t e r p r e t e d as t h e progressive mixing of d e e p e r , saline h y d r o t h e r m a l fluids w i t h overlying, d i l u t e ground w a t e r s t h a t h a v e b e e n p r e h e a t e d t o a p p r o x i m a t e l y 1 6 0 ' ~ . Independent,

isotopic evidence for t h e presence of a dilute ground water in t h e OH vein has been documented by Foley e t al. (1982) (discussed earlier). I t is interesting t o n o t e t h a t t h e e s t i m a t e d t e m p e r a t u r e of 1 6 0 ' ~ for t h e dilute ground w a t e r s is within t h e 100' t o 1 8 0 ~ ~ t e m p e r a t u r e range e s t i m a t e d for t h e diluent in most New Zealand geothermal systems (Hedenquist and Reid, 1984). The fluid-inclusion studies of Robinson (1981; Robinson and Norman 1984) also indicate t h a t t h e d e e p hydrothermal solutions mixed with shallow ground water in t h e southern Amethyst vein. Due t o t h e f a c t t h a t t h e i r fluid-inclusion study was done on quartz, i t was impossible for them t o distinguish growth zones or make detailed t i m e correlations. Instead, a more general approach was t a k e n and t h e inclusions from one s t a g e of q u a r t z deposition were measured on samples covering a vertical range of 336 meters. 100 140 180 220 260 300 Although a t e a c h sampled elevation t h e r e is a large TEMPERATURE ("C) range in both t h e t e m p e r a t u r e and salinity in t h a t Figure 7.13. Diagram showing the solubility of stage, t h e r e is a general decrease in both t e m p e r a t u r e barite contoured on a temperature vs. saand salinity with increasing elevation, which they linity plot. Barite solubility was calcua t t r i b u t e t o mixing. lated at pH = 5.5, mS = 0.02, sQ,/Sred = Additional evidence f o r mixing comes from t h e 10, .mNa/my = 9. Arrows mark boiling and district-wide mineral zonation of t h e sulfide-rich OH mlxing trajectories. vein in t h e north t o t h e barite-rich Bulldog vein in t h e south. In Figure 7.13 t h e solubility of barite is contoured on a temperature-salinity diagram and a l l o r e depostion a t Creede. This evidence is mixing and boiling trajectories relevant t o C r e e d e a r e summarized below: superimposed. It can b e seen t h a t barite solubility I. Fluid-inclusion evidence for boiling has been found changes relatively l i t t l e a t t h e high temperatures and only on t h e northern half of t h e OH vein, a n d only salinities appropriate for input fluids in t h e northern in q u a r t z deposited in t h e latest, silver-poor, s t a g e OH vein, and drops significantly only a t salinities of mineral deposition on t h e OH. below about 6 wt.-% NaCl. equivalent (Plumlee and 2. Fluid-inclusion evidence for substantial amounts of Hayba, 1985). (Note: in sulfate-rich solutions above mixing, increasing upward and t o t h e south, have pH 5, changes in pH have no e f f e c t on barite been found along t h e OH and southern Amethyst solubility). The topology of Figure 7.13 suggests t h a t vein systems. fluid mixing was t h e depositional mechanism for barite 3. The deposition of large amounts of barite in t h e a t Creede. Most mixing paths (decreasing t e m p e r a t u r e southern parts of t h e district and i t s absence in t h e and salinity) cross solubility contours while boiling northern parts is consistent with a mixing model, paths a r e parallel t o them. Thus, only a f t e r t h e hot, n o t with a boiling model of barite deposition. saline fluids, which rose in t h e northern parts of t h e 4. On t h e OH vein, t h e maximum t e m p e r a t u r e district, were diluted significantly did they deposit t h e measured on B-Stage material (the principal silverlarge quantities of barite seen in t h e southern and bearing s t a g e a t Creede) is 241°C, at l e a s t 30' less upper parts of t h e district. The silver content of t h e than t h e maximum t e m p e r a t u r e measured from DC r e e d e o r e s is also higher in t h e southern and upper S t a g e material. To produce boiling a t t h a t parts of t h e district (Barton et al., 1977), suggesting t e m p e r a t u r e would require t h a t t h e water t a b l e be t h a t mixing was an important mechanism of silver lower by about 200 m e t e r s during B-Stage than i t deposition a t Creede. was during D-Stage. 5. In a system dominated by lateral flow (i.e., o n e t h a t allows for only minor changes in pressure), i t Relative importance of boiling and mixing--It is virtually impossible for a fluid t o boil a f t e r i t s was pointed o u t a t t h e beginning of this section t h a t t e m p e r a t u r e has been lowered by mixing. both boiling and mixing a r e important processes leadThe widespread and intense sericitic a l t e r a t i o n ing t o t h e deposition of base- and precious-metal ores t h a t caps t h e C r e e d e orebodies, t h e widespread in epithermal systems. Drummond and Ohmoto (19851, occurrence of vein adularia in t h e system, and t h e Henley (1985, this volume) and Reed and Spycher deposition of massive amounts of Mn-Fe carbonates in (1985, this volume) have all emphasized t h e efficiency t h e deep parts of t h e southern portions of t h e vein of boiling in t h e precipitation of both ore and gangue systems all argue t h a t very substantial amounts of minerals in epithermal environments. Barton e t al. boiling took place during vein filling. None of these (1977) specifically related t h e precipitation of t h e f e a t u r e s can, however, b e correlated with t h e major C r e e d e o r e s t o cooling and pH changes occasioned by periods of base- and precious-metal deposition. We boiling. The evidence from t h e recent studies of conclude, therefore, t h a t although boiling can b e well C r e e d e c i t e d above, however, would imply t h a t mixing, documented a t C r e e d e by several criteria, t h e r e is no not boiling, was t h e immediate cause of most, if not

evidence t h a t boiling ( a t least in t h e ore zone) was related t o base- o r precious-metal deposition. The evidence f o r mixing of t h e deep, saline, o r e fluid with overlying fresh ground water during sulfide deposition, on t h e o t h e r hand, is overwhelming. As pointed o u t in t h e first section of this chapter, Creede is not unique among the Adularia-Sericite-type deposits in exhibiting a lack of evidence for boiling tied t o m e t a l deposition. Only a t Colqui is t h e evidence compelling t h a t precious-metal deposition resulted from boiling of t h e ore fluid (Kamilli a n d Ohmoto, 1977). Summary of C r e e d e mineralization On both geologic and geochemical grounds, i t has been proposed t h a t t h e C r e e d e o r e s were deposited along t h e t o p of a deeply circulating hydrothermal system a t t h e i n t e r f a c e of t h a t system with t h e overlying ground water. The concept is illustrated in Figure 7.11. The h e a t source driving t h e hydrothermal circulation is speculated t o have been a n intrusion underlying t h e district a t a depth of 3 t o 5 kilometers. This intrusion may have been related t o t h e quartz-latite, ring-fracture volcanism of t h e C r e e d e caldera cycle, or may have been a q u a r t z porphyry related t o t h e later bimodal basalt-rhyolite volcanism. The o r e fluids were dominated by m e t e o r i c waters, whose isotopic composition and salinity evolved by processes of evaporation and diagenesis in t h e playa lake in t h e m o a t of t h e C r e e d e caldera. The episodic introduction of magmatic fluids into t h e circulating system is suggested by t h e isotopic composition of rhodochrosite and siderite in t h e veins. The bulk of t h e lead in t h e C r e e d e ores, and therefore, presumably, most of t h e other metals, appears t o have been leached from Precambrian basement rocks a t depth. The sulfur in t h e C r e e d e o r e s m a y have come from several sources and i t s isotopic composition documents a complex mixing and exchange history n o t now satisfactorily understood. Most of t h e sulfate in t h e C r e e d e system appears t o have undergone considerable isotopic evolution in t h e playa lake sediments, but t h e sulfide sulfur appears t o have been buffered by, and perhaps derived from, a large reservoir of magmatic sulfur a t depth. Precipitation of t h e ores along t h e top of t h e system appears t o have resulted from the dual processes of boiling and mixing, but mixing appears t o have been quantitatively t h e more important mechanism of sulfide deposition. The intense, mixed-layer illite/smectite alteration c a p is interpreted t o have been generated by condensation along t h e top of t h e system of acid volatiles distilled off t h e deeply circulating o r e fluids. The model includes substantial amounts of fluidrock interaction and chemical and isotopic exchange in t h e deeper, higher t e m p e r a t u r e p a r t s of t h e system. In t h e ore zone, however, isotopic exchange between oxidized and reduced aqueous sulfur species was minimal a n d t h e aqueous sulfate apparently did not exchange oxygen with the ore fluid nor was t h e r e significant oxygen isotopic exchange between t h e o r e fluid and t h e unaltered wallrocks. The pH, t h e activity of S2 gas and redox s t a t e of t h e o r e fluids were buffered in t h e o r e zone by reaction with a vein-filling

assemblage comprising: Fe-chlorite + h e m a t i t e + pyrite + K-feldspar + sericite (or mixed-layer illites m e c t i t e clays). Throughout most of t h e history of vein filling, t h e redox s t a t e of the o r e fluid was t h a t of t h e triple point: Fe-chlorite + pyrite + hematite. During the early part of sulfide deposition, however, numerous excursions t o substantially lower oxidation (and sulfidation) s t a t e s occurred. The excursions have been interpreted t o h a v e resulted from episodic introduction of magmatic emanations into t h e circulating system or t o reaction with ferrous silicates episodically exposed t o t h e circulating fluids deep in the system through t e c t o n i c adjustments. Alternatively, t h e relatively oxidized s t a t e of t h e o r e fluid during most of t h e ore deposition cycle could b e due t o mixing of a deep, reduced fluid with surrounding and overlying oxidized groundwater prior t o entering t h e o r e zone, t h e excursions resulting from lesser amounts of mixing. Present evidence does n o t allow us t o choose between these t w o alternatives. THE ACID-SULFATE ENVIRONMENT: SUMMITVILLE AS AN EXAMPLE The recent thesis by Stoffregen (1985) combined with t h e earlier work by Steven and R a t t 6 (1960a) and several other studies (Patton, 1917; Mehnert et al., 1973; Lipman, 1975; Perkins and Nieman, 1983) make t h e Summitville mining district, Colorado, t h e best documented Acid-Sulfate-type deposit and t h e logical choice for illustrating t h e characteristics of AcidSulfate-type epithermal systems. Many of t h e interpretations on Acid-Sulfate-type deposits have been m a d e in light of t h e important insights gained from t h e studies done a t Goldfield, Nevada (Ashley, 1974; Ransome, 1909). Even so, i t should b e noted t h a t t h e observational base for this t y p e of deposit is still much smaller t h a n t h a t for t h e Adularia-Sericite deposits. Since our experience in t h e Summitville district is limited, most of t h e geologic and mineralogic characteristics discussed below, e x c e p t where noted, a r e based on t h e work done by Steven and R a t t 6 (1960a) and by Stoffregen (1985). As was t h e c a s e in using C r e e d e a s t h e exemplar f o r Adularia-Sericite-type deposits, Summitville has s o m e characteristics which a r e not representative of all of t h e Acid-Sulfate-type deposits. In particular, Julcani, Lake C i t y 11, and R e d Mountain a r e silverrich, contain relatively more base m e t a l s (particularly lead and zinc) and appear t o have formed a t greater paleodepths. Although these and o t h e r differences exist, the similarities in mineralogy, alteration, t e c t o n i c setting, and timing of o r e deposition relative t o host emplacement among Acid-Sulfate-type deposits a r e striking, and indicate mineralization in a distinct geothermal environment. Ashley (1982) has summarized t h e observations on a number of deposits of this type into a particularly useful occurrence model which provides further docurnentation on t h e characterisitics of Acid-Sulfate-type deposits. Geologic and Mineralogic Characteristics Volcanic

history--The

Summitville

mining

CHAPTER 7

Figure 7.14. Generalized geology of the Platoro and Summitville calderas, modified from Steven and Lipman (1976). Location of Summitville mining district is shown by pick-and-hammer. Control is moderate to good where boundaries are shown by solid symbols. A-A' marks location of crosssection shown in Figure 7.15.

2z =3

C .-z>/-----------South Mountain volcanic dome

13 I)

Talus-breccia \

district is located a t a n elevation of about 12,000 f e e t on t h e northwest e d g e of t h e Platoro caldera and t h e younger nested Summitville caldera in t h e eastern San Juan Caldera complex, Colorado (Liprnan, 19751 (Fig. 7.1). The quartz l a t i t e porphyry a t South Mountain, which hosts t h e deposit, is a lava dome emplaced 22.8 rn.y a g o (Mehnert e t al., 1973) along t h e western margin of t h e Summitville caldera a t i t s intersection with t h e Pass Creek-Elwood Creek-Platoro fault zone, a major s t r u c t u r e c u t t i n g across t h e c e n t e r of t h e Platoro-Summitville calderas (Fig. 7.14). Drilling has confirmed Steven and Rattg's (1960a) suggestion t h a t t h e dome is funnel-shaped in cross section with a narrow intrusive pipe a t depth which flares out near the surface (Fig. 7.15). The quartz l a t i t e is characterized by 2 t o 8 c m K-feldspar phenocrysts in an aphanitic, greenish groundmass; plagioclase phenocrysts a r e common but a r e finer grained than t h e K-feldspar. Also typical a r e locally resorbed quartz e y e s (0.2 t o 2 cm), euhedral biotite books (1 t o 2 cm), and a p a t i t e up t o 0.5 crn long (Stoffregen, 1985). The q u a r t z l a t i t e a t South Mountain is bordered on t h e north, e a s t , and south by t h e approximately 29 m.y. old Summitville Andesite which filled t h e Platoro caldera a f t e r collapse and on t h e west by t h e rhyodacite of Park C r e e k (about 28 t o 26 m.y.), which comprises lava-dome complexes erupted around t h e northwest margin of t h e Surnmitville caldera a f t e r i t s final collapse (Liprnan, 1975). The rhyolite of Cropsy Mountain, a coarsely porphyritic lava flow, locally overlies t h e q u a r t z l a t i t e south of t h e district. In

,----. ?:

ml

A4

1

--. '.-

7

\

I.---.

E X P L A N A T I O N Cropsy Mountain Rhyolite 20.2 m.y

0Coarse porphyry

j,",r:l South Mountain

Quartz Latite 22.9 m.y.

Quartz-alunite 22.3 m.y.

Summitville Andesite -29 m.y.

0Altered area

Figure 7.15. Geologic cross-section of the restored South Mountain volcanic dome, modified from Steven and ~ a t t b(1960a). Fault is shown with heavy line; contacts shown with thin line; both are dashed where approximate.

D. 0 . HAYBA, P. M.

BETHKE,

c o n t r a s t t o t h e older units, t h e rhyolite of Cropsy Mountain is everywhere unaltered. I t has been d a t e d a t 20.2 i0.8 m.y. (Mehnert e t al., 1973) and a t 18.5 i1.2 m.y. (Perkins and Nieman, 1983). The a g e of mineralization is t h e r e f o r e b r a c k e t e d between 22.8 and 20.2 m.y. in a g r e e m e n t with Mehnert e t al.'s (1973) d a t e of 22.3 m.y. on hydrothermal alunite from Summitville. O r e deposits--The o r e bodies a r e localized along t h e southwestern margin of t h e coarsely porphyritic c o r e of t h e dome, just north of t h e northwest-trending fault zone a t South Mountain (Steven and R a t t 6 , 1960a). The o r e occurs in a series of irregular pipes a n d vein-like masses of q u a r t z a n d alunite t h a t formed largely by r e p l a c e m e n t of t h e original q u a r t z latite. Significant mineralization is confined t o a vertical interval of approximately 400 meters, and t h e surface outcrop of t h e mineralized zone can b e circumscribed by an elipse with a x e s of 1.5 and 1.0 kilometers ( R a t t b and Steven, 1960a). Production through 1947 was 113,000 o z of Au, 240,OO o z of Ag, and 433,000 lbs of C u (Vanderwilt, 1947). Productions since 1947 has been insignificant. Wallrock alteration--The alteration in t h e upper p a r t of t h e deposit shows a well-defined p a t t e r n (Fig. 7.16) from a n intense zone of acid leaching, r e f e r r e d t o a s "vuggy silica alteration", surrounded by quartzalunite alteration grading outward into quartzkaolinite. F u r t h e r out, t h e r e is generally a n abrupt change into a n illitic* zone which grades o u t t o a montmorillonite zone. T h e width of e a c h of t h e zones is highly variable.

P. HEALD, & N. K. FOLEY Unaltered quartz latite

+

Montrnorillonite-chlorite

v

chloride-rich rock

zone A

~ontmdrillonite rich rock

153

QuartzIllite-kaolinite alunite zone zone Y

A

h

Illitid rock Mineralized / vuggy quartz rock

Figure 7.16. Diagram showing hydrothermal alteration pattern in the Summitville district adapted from Steven and ~ a t t s (1960a).

*Stoffregen (1985) has used t h e t e r m illite t o r e f e r t o a fine-grained ( < 2 micron) phyllosilicate with a 10angstrom basal spacing, which does not expand on glycolation, a n d contains less t h a n 5% s m e c t i t e layers. The vuggy silica a l t e r a t i o n is interpreted t o be t h e result of t h e acid dissolution of all t h e primary rock-forming minerals e x c e p t quartz. Most of t h e o r e occurs in this very permeable zone, which is characterized by l a r g e voids due t o t h e removal of t h e K-feldspar phenocrysts. A t t h e surf a c e t h e vuggy silica is quite extensive, but below a depth of about 1000 f e e t , i t becomes much m o r e restricted (Fig. 7.17). In t h e upper p a r t of t h e deposit, t h e surrounding .quartz-alunite zone is up t o 5 0 f e e t wide. Alunite occurs a s pseudomorphous r e p l a c e m e n t s of K-feldspar phenocrysts a n d a s m a t t e d a g g r e g a t e s replacing t h e groundmass. With decreasing elevation (between 11,500 and 11,000 feet), alunite becomes insignificant and quartz-kaolinite r a t h e r than quartz-alunite surrounds t h e vuggy silica "vein". T h e t e x t u r e of this deeper kaolinite indicates t h a t i t has completely replaced pre-existing alunite. Except f o r this d e e p zone, t h e quartz-kaolinite zone is generally thinner (and is locally absent) t h a n t h e quartz-alunite zone which i t surrounds (Stoffregen, 1985). Outside of t h e quartz-kaolinite zone, t h e t e x t u r e of t h e alteration shows a marked change from h a r d and brittle rock, due t o a silicified matrix, t o a soft, incompetent rock, due t o a rnatrix of clay minerals.

Figure 7.17. Schematic cross section of the alteration patterns and mineral zonation of the Summitville deposit, modified from Stoffregen (1985). The clay alteration zones refer to zones 3-6 in Figure 7.16. CV - covellite, luzonite, enargite, pyrite, marcasite, chalcopyrite, trace sphalerite, sulfur, and gold assemblage; TN - chalcopyrite, tennantite, pyrite, plus minor sphalerite and trace galena assemblage. The clay envelope around t h e vuggy silica and quartzalunite-kaolinite a l t e r a t i o n s is generally a t least 100 f e e t wide. According t o Stoffregen (1985), illite predominates n e a r t h e quartz-kaolinite zone, but montmorillonite becomes predominant further away from t h e vein. Kaolinite is present throughout t h e clay zone, but i t d e c r e a s e s in both abundance and

crystallinity away from t h e quartz-kaolinite zone. The illite becomes coarser grained with depth. [Mixedlayered illite/smectite clays (having g r e a t e r than 5% s m e c t i t e ) a r e generally n o t present, occurring only locally in t h e central portion of the deposit where t h e y appear t o b e related t o a post-mineral intrusion. (1985) describes t h e Mineralogy--Stoffregen mineralogy of Summitville in t e r m s of t h r e e main stages: main, late, and supergene. Mineralization from t h e s e t h r e e stages was later than much of t h e acid-sulfate alteration, a s evidenced by t h e presence of o r e minerals in cross-cutting f r a c t u r e s and voids in vuggy silica. In addition, a very minor, early s t a g e of mineralization is also present a s 10 t o 40 micron-sized inclusions in pyrite grains of either pyrrhotite or chalcopyrite + bornite + digenite. The significance of this minor, early s t a g e is not clear. T h e c h a r a c t e r of t h e main-stage mineralization changes with depth in t h e deposit (Fig. 7.17). In t h e deeper p a r t s of t h e deposit a chalcopyrite-tennantite assemblage predominates. Tennantite usually contains only minor antimony and t r a c e amounts of silver, although rarely i t may contain up t o 2.0 wt.-% silver. Pyrite plus minor, low-iron (usually 5 1.0 wt.-%) sphalerite and a t r a c e of galena a r e part of this d e e p assemblage. In t h e upper p a r t of t h e deposit, which contains t h e economically significant precious-metal mineralization, t h e main-stage mineral assemblage consists of covellite, luzonite, and enargite with pyrite, marcasite, locally sulfur, t r a c e amounts of sphalcrite, and gold. Chalcopyrite is also present, but i t decreases in abundance and is more extensively rimmed by covellite with increasing elevation. Gold is found in t h e native state. Silver occurs primarily in argentif erous covellite but lesser amounts a r e also found in either a r g e n t i t e or acanthite (not y e t determined), matildite, stromeyerite, and electrum. Late-stage mineralization a t Summitville is characterized by a barite, jarosite, goethite, and gold assemblage found in t h e uppermost levels of t h e deposit. Sulfides associated with this assemblage a r e extremely rare. A yellow phase included in a fluid inclusion in barite (Stoffregen, 1985) suggests t h e possibility t h a t native sulfur may also be p a r t of this assemblage. This late-stage assemblage, although volumetrically minor, locally contains high grades of gold. Supergene oxidation extends t o a depth of 100 t o 200 f e e t below t h e surface. Copper is essentially completely removed from t h e oxidized zone, immediately below which digenite and lesser chalcocite c o a t and replace other sulfides. Geochemical Environment Defining t h e geochemical environment of t h e Summitville deposits is limited by t h e c u r r e n t lack of fluid-inclusion and isotope data. P a r a m e t e r s such a s t e m p e r a t u r e of mineral deposition, chemical composition of t h e fluids, paleodepth, and origins of t h e fluids and dissolved constituents a r e f a r less well constrained than a t Creede. However, Stoffregen (1985) has done a n excellent job of limiting t h e conditions of alteration and mineralization using equilibrium thermodynamics. We will t a k e t h e s a m e approach, but we will also present s o m e alternative interpretations.

-

Temperature salinity ranges -- According t o Stof f regen (19851, t h e lack of fluid-inclusion d a t a is fluid inclusions. He found due t o a paucity. of primarv . only t h r e e samples with measurable primary inclusions in t h e small euhedral q u a r t z crystals intergrown with sulfides, lining voids in t h e vuggy silica a n d none were found in sphalerite o r alunite. His preliminary data, based on 19 measurements, show homogenization t e m e r a t u r e s in these samples ranged from 230' t o 320 C, with most of t h e values between 230' and 270'~. Two salinity measurements were 4 and 6 wt.-% NaCl equivalent. Perkins and Nieman (1983) also report t e m p e r a t u r e s f o r Summitville of 250' t o 280°c, presumably from fluid-inclusion measurements. In contrast t o t h e lack of good primary inclusions, quartz phenocrysts contain abundant secondary inclusions. The lack of secondary inclusions in quartz phenocrysts outside of t h e vuggy silica and quartz-alunite alteration zones (Stoffregen, 1985) is consistent with Bruha and Noble's (1983) suggestion t h a t these inclusions represent t h e fluids responsible for t h e intense quartz + alunite + pyrite alteration. Bruha and Noble (1983) measured homogenization temperatures of 231' t o 276OC, and salinities of 7 t o 21 wt.-% NaCl equivalent (averaging 10 wt.-%) in these secondary inclusions from o n e sample. They also report a s many a s 5 (unidentified) daughter (or trapped) minerals present in s o m e inclusions. Limited d a t a collected by G. H. Symmes i n our laboratory a r e consistent with those of Bruha and Noble. Based on paragenetic relations and on differences in salinity between these secondary inclusions and t h e primary inclusions in t h e vuggy silica zone, Stoffregen (1985) suggests t h a t t h e high-salinity fluids found in t h e f r a c t u r e s in quartz phenocrysts represent the alteration fluids, but not t h e l a t e r ore-forming fluids. These very limited d a t a may indicate a n evolution from high-salinity t o low-salinity w a t e r s in t h e Summitville hydrothermal system, but many more measurements a r e needed t o substantiate any such conclusion. The high salinity of t h e fluids which preceded o r e deposition is in agreement with observations made by Reynolds (Fluid, Inc., personal communication; also s e e Bodnar e t al., 1985, this volume). H e observed a few isolated, healed microfractures in some early quartz, defined by either vapor-rich H 2 0 + C 0 2 inclusions and/or halite-bearing inclusions with small vapor bubbles. While these inclusions a r e thought by Reynolds t o represent early m a g m a t i c fluids, their genesis is still uncertain due t o t h e limited information. Barite, which is an unreliable host f o r fluidinclusion d a t a (Ulrich and Bodnar, 19841, is t h e only hydrothermal mineral with relatively abundant fluid inclusions. Cunningham (1985) reports preliminary temperatures of roughly 1 0 0 f~o r ~ inclusions in this late-stage barite associated with t h e famous gold "boulder" found in talus slopes of t h e South Mountain dome in 1975. No salinity d a t a were reported. Paleodepth--Based on geologic reconstruction, Steven and R a t t e (1960a) e s t i m a t e t h e paleodepth t o t h e top of t h e o r e between 150 and 400 meters. Using a t e m p e r a t u r e of 250°C, a salinity of 10 wt.-%, and t h e lack of evidence for boiling from t h e limited d a t a

8

D. 0 . HAYBA, P. M. BETHKE, P. HEALD, & N. K. FOLEY on both secondary and primary fluid inclusions, a minimum depth of deposition of 400 m e t e r s is estimated from t h e tables of Haas (1971). Sources of constituents--No hydrogen-isotope d a t a a r e available for anv of t h e Acid-Sulfate-tv~e deposits. Goldfield is t h e o h y o n e with oxygen-isotope data, and therefore i t is used t o help docurnent t h e source of fluids in this t pe of deposit. According t o ' d a t a from Goldfield a r e Taylor (19731, t h e A1O compatible with ore-bearing fluids of essentially 100% m e t e o r i c rigin. Taylor does note, however, t h a t t h e overall depletion a t Goldfield is appreciably smaller than a t nearby Tonopah (an Adularia-Sericite deposit), and suggests t h a t a t Goldfield i t is likely t h a t either t h e alteration occurred a t lower t e m e r a t u r e s (125-20ooc), or from fluids with a higher 6'$ value. While Taylor favored the former, recent fluidinclusion d a t a from Bruha and Noble (1983) indicate temperatures of about 230' t o 270°C a t Goldfield, thus suggesting t h a t t h e fluids may have been richer in 1 8 0 implying a significant m a g m a t i c component. In t h e absence of deuterium data, i t is impossible t o do more than speculate on t h e source of t h e Goldfield o r e fluids. Well-constrained light-stable-isotope studies a r e badly needed on deposits of t h e Acid-Sulfate type. Whitney (1984a,b) has calculated t h e sulfur speciation in quenched magmatic gases evolved from rnagmas of various oxidation states. Relatively oxidized magmas such a s those a t Julcani (Drexler, 19821, and t h a t giving rise t o t h e Fish Canyon ash-flow tuff in t h e C e n t r a l San Juan Mountains (Whitney and Stormer, 1983) produce gases rich in SO2. SO2 gas is unstable a t temperatures below 400°C in t h e presence of water (Iwasaki and Ozawa, 1960; Holland 1965,1967; Sakai and Matsubaya, 1977; Burnham, 1979) and disproportionates into sulfuric acid and H2S gas. In our opinion, t h e intense acid-sulfate alteration characteristic of this deposit t y p e results from t h e t h e a t t a c k on t h e wallrock by t h e H2S04 produced by t h e Although Whitneyls disproportionation of SO2., calculations a r e consistent wlth a m a g m a t i c source for t h e sulfur, they do not demonstrate t h a t such a source was necessary. Stoffregen (19851, taking a different approach, has shown t h a t t h e intense acid-sulfate alteration, which preceded ore deposition, c a n be produced by a fluid whose chemistry is dominated by m a g m a t i c gases. He modeled the interactions between a n ascending sulfur rich "maginatic" gas (idealized a s H20-C02-SO -H2S-HCI) and liquid water using equilibrium ti2+

--,QH+

--.- - - - -,,~rr-::z--I:ZZZ~~~ -,

---'02*

Boil with Frac. Component Species

:

1 " " " " ' 1 " " " " ' ~ " " " " ' ~

B o i l w i t h Frac.

-

@: 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 ( 1 " " " " ' l " " " ' " I " " " " ' .

: B o i l w i t h F r a c . Metal c o m p l e x e ~ ~ :+

-

-

-- B o i l

1 " " " " ' I " " " " ' I ' " " " " .

with Frac.

Acidic

----.--*-------.-------------- HCO~

a=

-

HS

I

F@------__ ---.------_._-_ -----.--_-----------------_------*--=-c----i~+ -'-;.rrc=---c-----

soi2-

- 12

-

100

-

co32-

%--

a ~ +

4 -

--'--

1.

~ ~ 0 4 1 . . . . . , # , , I . . . . . . . . . l . , . . . . . . .

150 200 250 Temperature CDe8.C.)

300

Boi l w i th0u.t F r a c . Minerals

-.---------- -------_____ qz ---.

!

0

4

-6

OD 0

A

I*,;;;,;; @I -M

-

-.-----T

,, , ,,,

0 4

-8 -10

100

,,, ,,,,, ,, -7----

-

, , , ,S T ? ,= , ,= $

Boi l without Frac. Component Species

150 200 250 Temperature CDeg.C.)

300

Figure 11.4. Calculated isoenthalpic boiling of a Broadlandslike water without fractionation of minerals. The gas phase composition is essentially the same as in Figure 11.3a. a). Moles of minerals present in equilibrium with the gas and aqueous phases during boiling. See caption to Figure 11.3 for mineral abbreviations. b). Total molality of a portion of the component species. Activity of hydrogen ion is also shown, for which the ordinate refers to log activity. c). Total molality of o e-metal mmpcnent species. Felt refers to total aqueous Fe. d). Individual species molalities of principal complexes of ore metals.

Figure 11.5. Cooling of a Broadlands-like water without boiling and without mineral fractionation. (a) Moles of minerals present in equilibrium with the aqueous phase. See caption for Figure 11.3 for mineral abbreviations. (b) Total molality of a portion of the component species. Activity of hydrogen ion is also shown, for which the ordinate refers to log activity. (c) Total molality of ore-metal component species. (d) Individual species molalities of principal complexes of ore metals.

l l l l l ' ' l " l l l l l l l l l ' l l l ' ' l r l ' '

@

------------ -- - - - _ Cool ing Only* M i n e r a l s

W

h

l l ' l l ' ' ' l l l l l l l ' ' l ' ' l ' l ' ' T r ' ' l

CI-

~ a +

-- -Si02

7-7-T K+

7 = =

ca2+

-

- - - - - - - - - -w- 2+ - - - - ------.-__-___ s o ~2---.- -. - -. -~h~:---------------,.-.~;-------- - - _ _. --~ a .~ ~ - \ -

'--.Qz-

Component Spec i es 1

1

-

'

Cooling Only

~

1

1

~

1

1

~

1

1

1

"

1

1

"

1

1

'

1

1

l

1

1

"

1

~

1

1

~

1

1

~

1

1

~

1

~

Component Species

1

1

"

1

0;

F.~+

-

-

+=,"

--

/

-

----

-*+-

/.,,,l.,,-rrT;.......,.I.,,,,,,.,: L " " " " ' ~ ' " " " " ~ " " " " ' ~ " " " ~ "

-

--_______ ----- __--------CuIHS12-

.

AgHS

-

-

Cooling Only Metal Complexes 1

150

1

200

1

1

1

1

1

1

1

1

1

1

1

250

Temperature CDeg.C.1

1

--

1

1

3

.

1

1

-

300

"

~

~

~

"

~

'

This reaction also accounts for the replacement of mica by feldspar with decreasing temperature (and increasing pH) shown in Figures 11.3b and 11.4a. Boiling Without Fractionation and Cooling Only Temperature CDeg.C.)

F i g u r e 11.6. Comparison of pH i n c a l c u l a t i o n s of b o i l i n g (lower curve, b t h b o i l i n g calc u l a t i o n s ) and c o o l i n g o n l y (upper curve). See t e x t f o r discussion of differences. by Reed (1985). The high-sulfide concentration also holds Ag' in the stable A HS complex (Fig. 11.3h) throughout cooling from 2788 to 115O, where acanthite precipitates (Figs. 11.3b and 11.3d). In contrast to silver, gold and mercury never precipitate because their sulfide and bisulfide complexes (Fig. 11.3h) are quite stable over the entire temperature and pH range of boiling. Gold and mercury behavior are more thoroughly discussed below. Precipitation of Silicates In both boiling calculations, solid solutions of chlorite and muscovite precipitate as well as quartz and K-spar (Figs. 11.3b and 11.4a). Solid solution compositions (calculated using ideal and ideal multisite mixing) are not plotted separately, but compositions can be read from the figures by comparing the moles of solution end members. The calculated chlorite is iron rich, reflecting the high Fe/Mg ratio of the starting solution (Table 11.11, and the rapid dumping of Fe from solution during cooling as Fe(OHI4- dissociates. In contrast to the sulfides and carbonates, the very presence of silicates is not a consequence of the pH increase upon boiling. The effects of temperature decrease control silicate precipitation, as is obvious for quartz but perhaps not so obvious for the others. Muscovite and K-spar precipitate because A1 is liberated from AI(OHl4- as its stability decreases with decreasing temperature (Reed and Spycher, 1984). This displaces the following example reaction to the right

As discussed above, most of the chemical processes in the closed-system boiling calculation are the same as in the boiling with mineral fractionation. The only important differences are: (a) the lowtemperature sulfide assemblage is dominated by pyrite (Fig. 11.4b) which is entirely absent in the fractionation case (Fig. 11.3d) and (b) talc appears as a silicate phase in the closed-system case (Fig. 11.4a). Both of these differences are a consequence of the redissolution of the chlorite a t low temperature in the closed-system case whereas chlorite is fractionated out of the system (taking Fe and Mg with it) in the fractionation case. Back-reaction of the iron-rich chlorite at 200°C (Fig. 11.4a) provides iron for pyrite, magnesium for talc, and extra aluminum for K-spar. The purpose in executing the cooling-only calculation (Fig. 11.5) was to isolate the effects of temperature change itself from the effects of boiling. The cooling-only calculation results are valuable for this purpose, but are not necessarily valuable as a model of reality. The single most significant difference between the cooling and boiling calculations is that dissociation of weak acids upon cooling causes pH to decrease whereas pH increases in the boiling calculation (Fig. 11.6). This is why very little sulfide precipitates and phyllosilicates dominate instead of feldspar and carbonate (Fig. 11.5a). The decrease in pH with cooling displaces sulfide into H,S (reaction 6) accounting for the decrease in mHS- i n Figure 11.5b. Despite this decrease, metal-bisulfide complexes (Fig. 11.5d) do not dissociate sufficiently to overcome the effect of the acid conditions in enhancing sulfide mineral solubility. Another conspicuous difference in the cooling calculation is that graphite precipitates. The initial water is enriched in aqueous methane which escapes into the gas upon boiling. If boiling is not allowed, the hydrocarbon is forced to go into graphite upon cooling. (Graphite is the only reduced carbon phase, except methane, available in the numerical model.) SUPER- AND SUB-ISOENTHALPIC BOILING

Chlorite precipitates in response to this A1 supply combined with Fe supplied by dissociating Fe(0H);. The validity of these conclusions is indicated by the nearly quantitative removal of aqueous A1 and Fe in both the boiling reactions and in the cooling-only reaction (Fig. Il.5a), where pH decreases (Figs. 11.5b and 11.6). The high pH of the boiling waters fixes feldspar rather than mica in accordance with the following reaction

In addition to the nearly isoenthalpic boiling and cooling-only cases discussed above, we explored numerically the closed-system geochemical space corresponding to fluid ascent with partial conductive cooling and fluid ascent with excess enthalpy boiling. These results are shown in Figure 11.7 in diagrams of gas weight percent. The figures are temperature contoured with mineral stability boundaries, pH, and mole percent C 0 2 gas in the gas phase. The diagrams show the isoenthalpic boiling path (A-D, Fig. 11.7a)

E.

Minerals

Gas Fraction (W. 96)

Gas Fraction (wt. %)

co2 Mole Percent

0

I

I

I

10

20

30

Gas Fraction (wt. YO)

0

Figure 11.7. Calculated results of closedsystem boiling of a Broadlands-like water under a range of conditions corresponding to super-isoenthalpic to conduct ively cooled. The abcissa expresses the gas fraction (including H20 and the dry gases) of the total (liquid plus gas) in weight percent. At any given temperature on the ordinate, only one pressure yields a gas fraction corresponding to isoenthalpic boiling conditions, indicated by curve A-D, in part a). Other pressures yield more or less gas, corresponding to super- and subisoenthalpic boiling conditions. See text. a). Mineral phase diagram showing fields for the presence or absence of the indicated minerals. The shaded area at lower left is the stability field for kaolinite; to the right of the shaded area, moscovite is stable rather than kaolinite. See Figure 11.3 for mineral abbreviations. b). pH of boiling waters. Dotted line shows isoenthalpic boiling curve, A-D from part a). c). Mole-percent C02 in the gas phase in equilibrium with the boiling waters. Contours for 70 and 90% CO are not labeled but are shown to the legt of the 50% contour at low temperature and low-gas fraction. The dotted line shows the isoenthalpic boiling curve, A-D from part a). The other small circles show the points for which calculations were done to produce Figure 11.7.

260

CHAPTER 11

used in t h e previously discussed calculations depicted in Figure 11.4 and t h e cooling-only path (A-B, Fig. 11.7a), representing complete conductive cooling a s depicted in Figure 11.5. The isoenthalpic curve (Fig. 11.7) represents t h e temperature-steam fraction trajectory of a fluid which exchanges no h e a t with i t s wall rocks (Elder, 1965; Grant e t al., 1984). All h e a t used t o vaporize t h e water is supplied from t h e w a t e r itself. As t h e h e a t is removed, t h e w a t e r cools. An isoenthalpic path would b e expected under steady-state flow conditions in a hydrothermal system such t h a t t h e thermal gradient in t h e wall rock has been previously fixed by t h e boiling fluid (Grant et al., 1984). All trajectories between A-D and A-B (Fig. 11.7) represent fluids t h a t boil during ascent, but which simultaneously lose h e a t t o the wall rock. Such trajectories apply when pressure increases over t i m e due t o self-sealing, such t h a t t h e depth of first boiling ascends (from H toward L, Fig. 11.1). As t h e depth of boiling becomes shallower (because boiling is prohibited by increased pressure due t o self-sealing) ascending hot waters encounter rock t h a t is cooler than boiling along t h e newly established pressure-depth regime allows; thus, the rocks e x t r a c t h e a t from t h e water. Depending on t h e t i m e r a t e of self-sealing (and consequent r a t e of pressure change with time) t h e e x t e n t of such heat extraction by t h e rocks could result in cooling trajectories spanning much of t h e Superrange between A-D and A-B (Fig. 11.7a). isoenthalpic trajectories, lying t o the right of A-D (e.g., A-E, Fig. 11.7), represent boiling fluids t h a t e x t r a c t excess h e a t from t h e wall rock a s they ascend. This produces steam quantities in excess of t h a t which t h e ascending water alone could produce. Super-isoenthalpic conditions would prevail, for example, when new h e a t is introduced a t shallow levels by magma intrusion. A more common cause would be breaking of a self-sealed system, resulting in a reversion t o hydrostatic pressure from superhydrostatic conditions. This would cause t h e depth of first boiling t o deepen (downward toward H, and below, Fig. 11.1), exposing hot wall rock t o newly lowered pressure conditions. These s a m e thermal e f f e c t s a r e discussed by Goguel (1982) and Truesdell (1979) in t h e context of exploitation of geothermal energy reservoirs. The common occurrence of self-sealing and rebreaking in active geothermal systems is well established (e.g., White e t al., 1971; Muffler e t al., 1971; Keith et al., 1978; F a c c a and Tonani, 1967; Henley and Ellis, 1983) a s is t h e common occurrence of hydrothermal breccias in epithermal o r e deposits (e.g., Berger and Eimon, 1983; Nelson and Giles, 1985; Hedenquist and Henley, 1985). Thus, i t is m o r e probable t h a t many of the sub- and super-isoenthalpic trajectories represented in Figure 11.7 a r e visited during t h e a c t i v e lifetime of epithermal ore-forming systems. Chemical implications of this a r e discussed further below. The results shown in Figure 11.7 a r e based on calculations of overall heterogeneous equilibrium f o r e a c h of t h e individual points shown in Figure 1 1 . 7 ~ . Most of t h e s e points were calculated using isothermal pressure-drop traverses a t various temperatures.

Although e a c h such traverse is interesting in i t s own right, i t is most useful t o combine a l l results on a f e w diagrams, then explore t h e various ascent trajectories discussed above. The mineral phase boundaries shown in Figure 11.7a a r e best understood a s a consequence of t h e pH variations shown in Figure 11.7b, which a r e primarily a consequence of t h e C 0 2 degassing represented in Figure 1 1 . 7 ~ . The rapid pH increase with early boiling (corresponding t o small gas fractions, Fig. 11.7b), also apparent in Figures 1 1 . 3 ~ a n d 11.4b, a r e a result of t h e early rapid escape of most C O into t h e gas phase a s indicated in Figure 1 1 . 7 ~a n 2 plotted in moles in Figure I l.3a. Figure 11.7b shows t h a t essentially all boiling paths e m a n a t i n g from point A result in pH increases e x c e p t conductive cooling paths near path A-B. The monotonic d e c r e a s e in pH along path A-B is shown in Figure 11.5b. The pH control on gangue silicate mineralogy, for example, is apparent from t h e approximate parallelism of t h e muscovite-K-spar boundary in Figure 11.7a and t h e pH contours of Figure 11.7b. K-spar forms on t h e high-pH side of the boundary in accordance with reaction 8. According t o t h e phase diagram, a l l boiling trajectories t h a t result in more ~ produce d e e p t h a n 10 percent gas a t 1 0 0 ~will muscovite and shallower K-spar. The particular depth where t h e transition from muscovite t o K-spar occurs is fixed by the point of intersection of t h e ascent trajectory (e.g., A-C) with t h e phase boundary in Figure 11.7a. The cinnabar field in Figure 11.7 is particularly instructive because i t demonstrates t h e e f f e c t s of competing reactions on t h e solubility of a sulfide mineral. According t o Figure 11.7a, isoenthalpic boiling does not precipitate cinnabar from this water, but neither does conductive cooling. Consider a n isothermal P-drop traverse (left t o right, Fig. 11.7) a t ~ ~ 170°C; cinnabar a t e m p e r a t u r e between 1 0 0 and remains soluble in t h e absence of boiling (zero gas 11.7) because t h e large H S fraction, Fig. concentration in t h e aqueous phase displaces t z e following reaction t o t h e l e f t HgS(H2S)2 = HgS + (aq) (cinnabar)

W2S (aq)

(9)

The protonated mercury complex and H2S a r e strongly favored by t h e low pH (Fig. 11.7). Upon boiling, reaction 9 is displaced t o t h e right a s H S (gas) escapes, resulting in cinnabar precipitation. dowever, a s boiling proceeds, t h e consequent increase in pH (Fig. 11.7b) shifts sulfide equilibria such t h a t t h e HgST, increases in mercury-sulfide complex, concentration, causing cinnabar to redissolve (reaction 10) despite t h e further loss of H2S gas HgS + (cinnabar)

HS--+ HgS= + H+ (aq)

(aq?

(10)

(aq)

In t h e c o n t e x t of possible fluid ascent trajectories, this behavior of cinnabar means t h a t only paths with partial conductive cooling of t h e fluid will produce cinnabar.

The cinnabar and muscovite-K-spar examples illustrate t h e profound e f f e c t of t h e boiling h e a t budget on t h e expected epithermal vein mineralogy. T o t h e e x t e n t t h a t t h e h e a t budget of boiling fluid "parcels" varies through time, fluid trajectories on Figure 11.7a vary and vein mineralogy must vary. This is o n e probable source of mineral banding in epithermal veins. As a system experiences self-sealing and re-breaking, trajectories shift from isoenthalpic t o sub-isoenthalpic t o super-isoenthalpic on Figure 11.7a. Such a fluid, a t i t s 230°C isotherm (which itself shifts in space) would first precipitate quartz, calcite, K-spar, Fe-chlorite, chalcopyrite, sphalerite, and galena (Fig. 11.4a), then (depending on where t h e subisoenthalpic trajectory falls) perhaps quartz, muscovite, pyrite, and sphalerite (but not galena, chalcopyrite, chlorite, K-spar, and calcite), a s indicated by Figure 11.7a a t 230' and 5-percent gas fraction. A t this s t a g e t h e much lower pH of t h e fluid (Fig. 11.7b) might also e t c h earlier formed minerals and cause muscovite t o replace earlier formed K-spar (e.g., Hedenquist and Henley, in preparation, 1985). Upon re-breaking and a swing t o super-isoenthalpic conditions, the fluid again would precipitate carbonate, feldspar, etc., a t i t s 230' isotherm. Instead of following a fluid parcel of given t e m p e r a t u r e through t h e history outlined above, we may consider how physical and chemical conditions vary a t a point fixed in ,space (in a vein cavity, for example). A gradual pressure increase due t o sealing would correspond to traversing toward high t e m p e r a t u r e in Figures 11.3 and 11.4 for any given position in a vein. Thus, a point initially a t 250°C (Fig. 11.3b) where quartz, K-spar, chlorite, sphalerite, and galena a r e precipitating could experience a shift t o 2 7 8 ' ~ a s pressure increases, resulting in etching of the previously formed minerals, overlapping precipitation of muscovite, and replacement of K-spar by muscovite. Subsequent pressure drop would result in a return t o precipitation of the lower t e m p e r a t u r e assemblage. Other more complicated changes could occur, depending on t h e r a t e of pressure change, resulting in sub- and super-isoenthalpic boiling (Fig. 11.7a) a t t h e given point.

BOILING AND GOLD PRECIPITATION The large sulfide content of the Broadlands-like water used here stabilizes t h e Au(HS)Z complex (Fig. 11.3h) precluding gold precipitation throughout . the t h e boiling range from 2 7 8 ' ~ t o 1 0 0 ~ ~ In calculation of boiling with mineral fractionation (Fig. 11.31, gold is undersaturated by more than t w o orders of magnitude a t 2 7 8 ' ~ (log Q/K = -2.44), but this decreases t o undersaturation by half a n order of ~ ~ Q/K = -0.57). magnitude a f t e r boiling t o 1 0 0 (log Thus, if t h e solution had been saturated with gold a t t h e point of incipient boiling (as in Drummond and Ohmoto, 1985, who used a large gold concentration) or if we had used t h e new value for gold concentration from Henley and Brown (1985, this volume) which is 2 3 times g r e a t e r t h a n t h e 0.1 ppb t h a t we assumed (Table 11.1), gold would have precipitated due t o boiling. The

c r i t i c a l reaction for gold precipitation in t h e acid t o neutral pH range is t h e following 8Au (HS) 5 + 6H+ + 4 ~ 2 0 (aq) (aq) (aq)

~AUO

(gold)

+ S q + I 5H2S (aq)

(g or aq) (11)

This reaction entails reduction of Au+ t o AuO by sulfide which is thereby oxidized t o sulfate. Reaction 11 makes i t clear t h a t pH increase induced by boiling competes with loss of H2S gas in determining whether gold precipitates or not (see also, Drummond and Ohmoto, 1985). The thermochemical d a t a used here for gold-bisulfide complexes (Seward, 1973) indicate t h a t a s pH increases i n t h e range calculated h e r e (6.4 t o 7.81, gold solubility increases (reaction 11) a s H2S dissociates t o HS-. Thus, the calculated approach toward gold saturation upon boiling is principally a consequence of loss of H2S t o t h e gas phase, displacing reaction I I toward gold saturation (see also, Hedenquist and Henley, 1985). The precipitation of sulfide minerals constitutes another sink for sulfide which promotes gold precipitation. However, in t h e c a s e calculated, t h e e s c a p e of H2S t o t h e gas phase is f a r more significant than sulfide mineral formation a s a sink for aqueous sulfide (compare Figs. 11.3a and 11.4a), a s discussed in a preceding section. In any case, gold did not precipitate because of t h e excess sulfide content of original water. Instead, i t stayed in t h e aqueous phase and, in t h e geologic context, i t would b e transported t o t h e surface hot-spring environment (see below). In order f o r gold t o precipitate with sulfides in t h e d e e p vein environment a s i t did, for example, a t Comstock (Bastin, 19221, i t is necessary t o c r e a t e a sulfide-deficient water by removing essentially all of t h e aqueous sulfide t o minerals and gas. This is best accomplished by boiling waters which have a large r a t i o of m e t a l s (Fe, Cu, Pb, Zn) t o sulfide so t h a t precipitation of pyrite, chalcopyrite, galena, and sphalerite along with H2S gas removes a l l sulfide from t h e aqueous phase. The low salinity of t h e Broadlandslike water precludes high-metal contents because chloride is necessary to complex significant concentrations of base metals in t h e presence of highsulfide concentrations. Other calculations (Reed, 1985) on more saline waters (e.g., .5 m NaCI) show t h a t chloride complexes of Pb, Zn, and C u dominate, even in high-sulfide waters. A large metal/sulfide r a t i o is necessary t o assure sulfide depletion with consequent gold precipitation, and large salinities a r e necessary t o large base-metal concentrations. Thus, t h e boiling of saline waters should produce gold precipitation with sulfides in accordance with t h e following example reaction in which galena and gold precipitation a r e coupled

4

8 ~ +~ so4 ' + 30 C I - + 15PbS + 24H' (aq) (galena) (aq) (gold) (aq)

(12)

As boiling proceeds, COZ escapes, driving up pH (reaction 1); this decrease In H+ activity drives sulfide

precipitation (reaction 31, which in sulfide-deficient waters is coupled t o gold precipitation a s in reaction (12). Thus, t h e formation of a hot-spring gold deposit (see below) a s opposed to a base-metal-vein gold deposit may result primarily from t h e distinction between sulfide-deficient a n d sulfide-excess boiling waters, which itself is t i e d t o t h e distinction of lowhigh-salinity waters. Fluid-inclusion d a t a salinity support this distinction.

=

different than t h e boiling calculation. There is a gas phase in equilibrium with a liquid phase. As t e m p e r a t u r e decreases at constant pressure, gas condenses t o liquid and chemical potentials of all species a r e equalized between t h e two phases. The resultant liquid composition at 95OC is given in Table 11.2 and pH and gas composition a r e shown in Figure 11.9. Gas and liquid compositions a r e also shown a t t h e e x t r e m e left-hand ends of Figures 11.10a and 11.10b.

THE HOT-SPRING ENVIRONMENT In t h e hot-spring environment a t and just below t h e surface over a boiling hydrothermal system, t h e e f f e c t s of cool temperatures, atmospheric oxidation, and t h e influx of m e t e o r i c ground w a t e r s determine patterns of rock alteration, water chemistry, and o r e formation. Figures 11.1 and 11.2 show t h e nearsurface geochemical processes t h a t we chose t o simulate numerically for t h e purpose of understanding t h e origins of hot-spring gold deposits (Berger and Eimon, 1983; Henley and Ellis, 1983). In t h e process of unravelling t h e chemistry of gold deposition, we generated numerical models of t h e formation of acid condensates, acid-sulfate waters, and neutral carbonate-sulfate waters. All results were produced using t h e s a m e fundamental approach for multi-component heterogeneous equilibrium calculations (Reed, 1982) t h a t we applied t o boiling. The sequence of operations on t h e gas and liquid phases l e f t over from boiling a r e a s follows (refer t o Figs. 11.1 and 11.2 for points designated below by letters): 1. c o n d e n s e t h e gas phase in open space by cooling from IOOOC t o 93OC (B). 2. Oxidize t h e gas with atmospheric 0 2 (C). 3. Condense t h e gas in cold, oxygenated ground water and r e a c t a t 99OC (D). 4. T i t r a t e (D) t h e acid-sulfate w a t e r from (C) i n t o t h e boiled aqueous phase from (A). 5. T i t r a t e (F)t h e cold, oxygenated ground water into t h e boiled aqueous phase from (A). The results of e a c h of t h e s e calculations a r e presented below in t h e order of t h e above list.

Atmospheric 0 2

T°C Condensate

100 (H20,CO2,H2S)

*

0 2 iqT

H 2 0 C02

Liquid Condensate

95OC B

Condensation of t h e Boiled Gas When t h e gas phase from a boiling hydrothermal system escapes into cool, f r a c t u r e d rock above t h e boiling water table, t h e H 2 0 fraction largely condenses (White, 1957; White e t al., 1971) and carries with i t small amounts of dissolved C 0 2 and H2S. Subsequent, direct atmospheric oxidation (Fig. 11.8) produces acid-sulfate waters. We calculated t h e chemical characteristics of the condensate in order t o provide a starting composition f o r oxidation, but also t o understand b e t t e r how t h e condensate may a c t alone in altering wall rocks and re-dissolving vein minerals. We separated t h e gas phase from t h e li uid phase l e f t over from t h e boiling calculation a t 10173C (Table 11.1, Fig. 11.3a) and numerically condensed i t by cooling t h e gas alone t o 93OC a t a pressure of 1.01 bar. Computationally, such a condensation is no

Figure 11.8. a). Schematic diagram showing the boiled gas phase passing through a temperature gradient in fractured rock above the water table. The H20 condenses, carrying some C02 and H2S with it. Atmospheric oxygen enters from above, oxidizes HZS to sulfuric acid which dissolves in the liquid condensate. b). Diagram of a semi-closed system approximation to the scheme in part a), as set up for numerical oxidation at 9 5 O ~(see text and Fig. 11.10). Oxygen gas is titrated into a two-phase mixture of condensate plus gas, but the original gas phase remains in contact with the liquid throughout the calculation. Thus, the H2S is gradually consumed by oxidation, rather than being continuously re-supplied from belm.

M. H.REED & N. F. SPYCHER Table 11.2--Water

compositions Acid-sulfate water Gas + ground water (0.03 moles of added 02) (72 kg of added water)

Condensed gas (95'~)

Cold ground water

PH

7 (25'~) (molality) (ppm)

C1-

.733e-04

2.6

F-

.174e-04

0.33

C O ~

.202e-02

121

HSCH4 aq.

---

---

---

--3.1

O2 aq.

.323e-04 .250e-03

Si02

.105e-02

63

~1+++

.l lle-06

0.003

~a++

.724e-04

2.9

M~++

.535e-04

1.3

~e++

.179e-06

0.01

K+

.358e-04

1.4

~ a +

.457e-04

10.5

S O ~

5

0

40

:

:

.

.

8

1

~

~

.

~

.

-

-

1

~

-

30 7

:

20

1

:

263

t

Condensation 1 bar, Oar Phars .

L

.

1

.

.

3

I

. . .

1 . . . 1

Condensate Iliquid)

4.0% 92

94

96

98

18%

Temperature