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12 R.G.C. BATHURST CARBONATE SEDIMENTS AND THEIR DIAGENESIS 16 H. H. RIEKE 111 and G. V. CHlLlNGA RIA N COMPACTION OF ARGILLACEOUS SEDIMENTS 18 G. V. CHlLlNGARlAN and K.H. WOLF, Editors COMPACTION OF COARSE-GRAINED SEDIMENTS 19 W. SCHWARZACHER SEDIMENTATION MODELS AND QUANTITATIVE STRATIGRAPHY 20 M.R. WALTER, Editor STROM ATOLITES 25B G. LARSEN and G. V. CHILINGAR, Editors DIAGENESIS IN SEDIMENTS AND SEDIMENTARY ROCKS 26 T. SUDO and S. SHIMODA, Editors CLAYS AND CLAY MINERALS OF JAPAN 29 P. TURNER CONTINENTAL RED BEDS 30 J.R.L. ALLEN SEDIMENTARY STRUCTURES 33 G.N. BATURIN PHOSPHORITES ON THE SEA FLOOR 34 J.J. FRIPIAT, Editor ADVANCED TECHNIQUES FOR CLAY MINERAL ANALYSIS 35 H. VAN OLPHEN and F. VENIALE, Editors INTERNATIONAL CLAY CONFERENCE 198 1 36 A. IIJIMA, J.R. HElNand R. SIEVER, Editors SILICEOUS DEPOSITS IN THE PACIFIC REGION 37 A. SlNGERandE. GALAN, Editors PALYGORSKITE-SEPIOLITE: OCCURRENCES, GENESIS AND USES 38 M.E. BROOKFIELD and T.S. AHLBRANDT, Editors EOLIAN SEDIMENTS AND PROCESSES 40 8. VELDE CLAY MINERALS - A PHYSICO-CHEMICAL EXPLANATION OF THEIR OCCURRENCE 4 1 G. V. CHlLlNGA RIA N and K.H. WOLF, Editors DIAGENESIS, I 42 L.J. DOYLE and H. H. ROBERTS, Editors CARBONATE-CLASTIC TRANSITIONS 43 G. V. CHlLlNGARlAN and K.H. WOLF, Editors DIAGENESIS, II 44 C.E. WEAVER CLAYS, MUDS, AND SHALES 45 G.S. ODIN, Editor GREEN MARINE CLAYS 46 C.H. MOORE CARBONATE DIAGENESIS AND POROSITY 47 K.H. WOLF and G. V. CHILINGARIA N , Editors DIAGENESIS 111 48 J. W. MORSE and F. F. MACKENZIE GEOCHEMISTRY OF SEDIMENTARY CARBONATES 49 K. BRODZIKOWSKI and A.J. VAN LOON GLACIGENIC SEDIMENTS

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Evaporites, Petroleum and Mineral Resources

Edited by

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Data Center ,09126599985,[email protected], For Educational Uses AU‘I’HOK LIS’I‘ V

Donald E. Anders ‘I’imothy K. Lowenstein (Geochemist) State University of New York @ United States Geological Survey Binghamton Mailstop 939, Box 25046 Department of Geological Sciences Lakewood, Colorado 80225 Binghamton, New York 13901

C. Robertson Handford Haq H. Posey ARC0 Oil and Gas Company (Consulting Geologist) 2300 W. Plano Parkway 2020 Routt Street Plano, lexas 75075 Lakewood, Colorado 80215

Robert J. Hite Joseph P. Smoot (Geologist Emeritus) (Geologist) United States Geological Survey United States Geological Survey 10190 W. 78th Avenue Mailstop 939, Box 25046 Arvada, Colorado 80005 Lakewood, Colorado 80225

J. Richard Kyle John K. Warren Department of Geosciences (Principal Petroleum Geologist) University of Texas @ Austin National Centre for Geology P.O. Box 7905, and Geophysics Austin, ‘Texas 78713 GPO Box 498 Adelaide, South Australia 5001

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Data Center ,09126599985,[email protected], For Educational Uses PREFACE

A thorough knowledge of evaporites is essential to professional researchers and geologists in our efforts to determine: the sedimentary histories, depositional environments, eustatics, chemical, climatic, environmental influences, and other factors involved in the deposition of both modern and ancient sedimentary sequences in basins worldwide. Evaporites are found in locations as diverse as Antarctica to the equatorial latitudes, and in depositional settings ranging from intracontinental to marginal marine. Applications of modern scientific techniques and analyses have lead to alternative hypotheses which may challenge conventional depositional models. This volume illustrates the expanding knowledge of evaporites as important reservoir seals, fluid aquitards, ore-hosting sediments, and economicallyviable sediments in their own right. Researchers, oil and gas professionals, minerals resource professionals, environmental specialists and others within geology and the other sciences shall utilize the information within this book in their understanding of the many recent discoveries and concepts developed in the field of evaporite sedimentology. The first three chapters discuss facies, fabrics, textures, crystallography, mineralogy, hydrology, diagenesis, and depositional settings of marginal marine and non-marine evaporites. Woven throughout the first three chapters are discussions of hydrocarbons and mineral resources as they relate to evaporitic basins and deposits. The last three chapters discuss the geochemical, mineralogical, depositional, diagenetic history and economical importance of evaporites and their associated Occurrences with hydrocarbons, ores and bittern salt deposits. Salt dome formation and emplacement are limited to a brief overview with the reader referred to many excellent studies of diapirism and halokinesis. This volume is specifically about evaporite sedimentology, hydrocarbons and mineral resources. Chapter 1 examines controls and crystallography of halite precipitates and the deposition and diagenesis of marine halite as subaqueous cumulates, bottom and intrasediment precipitates. Also included are discussions of syndepositional and early postdepositional diagenesis associated with evaporites in both modern and ancient sabkha and salina depositional settings. Case studies include locations from Africa, Australia, Mexico and the U.S.A. Chapter 2 discusses CaSO, dominated sea-marginal and platform evaporative depositional environments and associated dolomitization. Case studies include the United Arab Emirates, Gulf of Elat, Nile Delta Coast of Egypt, Mediterranean Coast of the northern Sinai, the southern Sinai, Australia and the U.S.A. The importance of displacive and replacive evaporite growth, syndepositional and postdepositional hydrology, and diagenesis is covered. Chapter 3 reviews depositional environments of modern and ancient non-marine

Data Center ,09126599985,[email protected], For Educational Uses VIII PREFACE

evaporites. Lacustrine, fluvial and eolian subfacies are reviewed along with the diagenesis, mineralogy, facies, hydrology and tectonic settings of evaporite formation. Saline soils and their classifications are discussed. Examples from around the globe help illustrate the facies discussed. Chapter 4 investigates the relationship between organic matter, petroleum and evaporites. Evaporites as source rocks for petroleum and evaporite associated oils from marine carbonate, anhydrite, halite-potash and lacustrine carbonate settings are covered. Petroleum resources resulting from halokinesis and diagenesis of evaporites are also discussed. Examples of the relationship between petroleum and evaporites in marine and lacustrine settings include the U.S.A., Gulf of Elat, Turkey, Egypt and the Israel- Jordan border. The authors have included their recently completed unpublished chromatographic data of hydrocarbon analyses from various locations. Chapter 5 discusses halokinesis and cap rock developments with emphasis placed upon the importance of trapping hydrocarbons, sulfur, and metalliferous ores in and around salt domes. Textures, fabrics and mineralogy in the true calcite cap rock, the marine false cap rock, the stock and other associated salt dome areas are discussed. Examples from the U.S.A. and Tunisia are employed. Chapter 6 surveys the economics of evaporites, evaporitic processes and mineral resources. Resources discussed include gypsum, anhydrite, halite and potash, borates, sodium carbonates and sulfates, bromine, iodine, lithium, nitrogen, zeolites, sulfur, barite, celestite, and metalliferous ore deposits containing zinc, lead, silver and copper. Evaporite solution in the creation of ore-hosting porosity and evaporites' involvement with igneous-hosted ore deposits are covered. Evaporites in Mississippi Valley Type (h4VT) and sedimentary exhalative type (Sedex) processes and deposits are covered. The field of evaporite sedimentology has long been considered a "stepchild to carbonate and clastic sedimentology. This book is intended to help bring the latest discoveries and concepts of evaporite sedimentology to the reader and to help bring evaporite sedimentology to its rightful place as an important and significant element in any complete understanding of sedimentology. The editor received manuscripts from the authors in a variety of formats. All manuscripts were prepared by the editor in camera-ready format on an IBM PC. The editor wishes to acknowledge and express gratitude to Joseph V. Rochefort for his contribution to this volume in maintaining references and labeling figures, acting as intermediary to the authors, and for general editorial support. He provided constant encouragement and support to get this project completed.

January 1, 1991 Judith L. Melvin

Data Center ,09126599985,[email protected], For Educational Uses TABLE OF CONTENTS

AUTHORLIST ...... V

PREFACE ...... VII

CHAPTER 1. MARGINAL MARLNE HALITE: SABKHAS AND SALINAS by C . Robertson Handf ord Introduction ...... i Fundamental Controls on Marine Halite Precipitation ...... 2 Halite Crystallography and Crystal Growth ...... 6 Deposition and Early Diagenesis of Marginal Marine Halite ...... 10 Subaqueous Cumulates ...... 10 Subaqueous Bottom Precipitates ...... 13 Intrasediment Precipitates ...... 17 Clastic Halite Particles and Hydrodynamic Structures ...... 22 Syndepositional and Early Postdepositional Diagenesis ...... 23 Dissolution and karst formation ...... 25 Pressure ridges and polygons ...... 29 Cementation ...... 32 Marginal-Marine Halite Depositional Settings .Salinas: Modern and Ancient ... 34 Salina Subenvironments ...... 37 Dry mudflats ...... 39 Saline mud flats ...... 40 Ephemeral salt pans or perennial saline lakes ...... 41 Danakil Depression of the Afar Region, Africa ...... 43 Lake MacLeod, Western Australia ...... 46 Hutt Lagoon, Western Australia ...... 47 Permian Basin, Texas, U.S.A...... 49 Marginal-Marine Halite Depositional Settings - Sabkhas: Modern and Ancient . . 52 Supratidal Mudflats, Northwest Gulf of California ...... 57 Ordovician Joachim Dolomite, Arkansas, U.S.A...... 59 Conclusions ...... 59 References ...... 61

CHAPTER 2 . SUL.FATE DOMINATED SEA-MARGINAL AND PLATFORM EVAPORATIVE SETTINGS: Sabkhas and salinas. mudflats and salterns by John K . Warren Introduction ...... 69 Economic Significance ...... 69 History of Evaporite Research ...... 70 Sea-Marginal Sabkha Model: Arabian (Persian) Gulf Example ...... 74 Subtidal Facies ...... 79

Data Center ,09126599985,[email protected], For Educational Uses X TABLE OF CONTENTS

Intertidal Facies (lagoonal) ...... 83 Supratidal Facies ...... 83 Lower Supratidal ...... 83 Middle Supratidal ...... 84 Upper Supratidal ...... 84 Limitations Inherent to the Sea-Marginal Sabkha Model ...... 87 Sabkha Dolomitization ...... 90 Other Sea-Marginal Sabkhas ...... 93 United Arab Emirates ...... 93 The barrier island - lagoon sabkha ...... 93 Mainland beach - dune sabkha ...... 93 Tidal estuarine sabkha - delta complex ...... 94 Fan-deltasabkha ...... 94 Continental interdunal sabkha ...... 94 Gulf of Suez ...... 94 Nile Delta Coast, Mediterranean Egypt ...... 96 Mediterranean Coast of Northern Sinai ...... 98 Reservoir Occurrence using a sabkha model ...... 99 Modern Sea-Marginal Salina Models ...... 100 Hydrology of Modern Coastal Basins ...... 101 Southern Australian Salinas ...... 103 Lake MacLeod, Western Australia ...... 107 Holocene Salina Evaporites ...... 110 Solar Lake, Gulf of Elat ...... 110 Ras Muhammad Pool, Southern Sinai ...... 112 Hutt and Leeman Lagoons, Western Australia ...... 114 Salina Comparison ...... 116 Modern Salina Dolomite Model - Cooroong, South Australia ...... 118 Lake Stratigraphy ...... 120 Dolomite Occurrence ...... 122 Type 1 lakes ...... 124 Type 2lakes ...... 124 Sedimentary Structures in the Massive Unit ...... 125 Geological Implications ...... 125 Brine Reflux Dolomitization ...... 129 Ancient Platform Evaporites ...... 132 Ancient Evaporite Basins - Depositional Models ...... 133 Platform Evaporites - Depositional Setting ...... 135 Sabkhas and Evaporitic Mudflats ...... 138 Ordovician Red River Formation, Williston Basin, U.S.A...... 138 Upper Minnelusa Formation, Wyoming, U.S.A ...... 142 Basal Seven Rivers Formation, Yates Field, West Texas, U.S.A...... 144 Ancient Salterns ...... 150

Data Center ,09126599985,[email protected], For Educational Uses TABLE OF CONTENTS XI

Ferry Lake Anhydrite .Fairway Field. East Texas. U.S.A...... 152 San Andres Formation .Northwest [email protected] Texas. U.S.A...... 154 Comparisons of Sabkhas and Salinas ...... 164 Diagenesis of Sulfate-Dominated Evaporites ...... 167 Diagenetic CaSO, .Displacive vs . Replacice Growth ...... 167 Evaporite Diagenesis .the Importance of Hydrology ...... 169 Syndepositional Hydrology of Holocene Basins ...... 169 Burial Hydrology: Gypsum .Anhydrite. and Dewatering ...... 173 Summary of Calcium Sulfate Diagenesis ...... 178 Acknowledgments ...... 178 References ...... 180

CHAPTER 3. DEPOSITIONAL ENVIRONMENTS OF NON-MARINE EVAPORITES by Joseph P . Smoot and Tim K . Lowenstein Introduction ...... 189 Modern Non-Marine Evaporites ...... 190 Hydrologic Setting ...... 191 Chemistry ...... 197 Depositional Subenvironments of Non-Marine Evaporites ...... 202 Lacustrine Deposits ...... 203 Perennial Saline Lake Subenvironment ...... 203 Cumulus crystals ...... 204 Evaporitecrusts ...... 205 Detrital deposits ...... 207 lntrasediment crystal growth and cements ...... 208 Saline Pan Subenvironment ...... 213 Floodingstage ...... 213 Saline lake stage ...... 215 Desiccation stage ...... 216 Saline Mudflat Subenvironment ...... 220 Dry Mudflat Subenvironment ...... 228 Shoreline Subenvironment ...... 235 Deltaic deposits ...... 235 Wave-formed deposits ...... 240 Fandeltas ...... 244 Ancient Non-Marine Evaporites in Lacustrine Deposits ...... 244 Fluvial Deposits ...... 253 Alluvial Fan Subenvironment ...... 254 Ephemeral Stream Subenvironment ...... 257 Perennial Stream Subenvironment ...... 258 Ancient Fluvial Deposits ...... 259 Other Deposits ...... 260

Data Center ,09126599985,[email protected], For Educational Uses XI1 TABLE OF CONTENTS

Eolian Dunefield and Sand Sheet Subenvironment ...... 261 Spring Subenvironment ...... 266 Saline Soil Subenvironment ...... 271 Diagenesis ...... 278 Distribution of Subenvironments ...... 282 Alluvial Fan - Saline Pan (Dry Mudflat) ...... 282 Alluvial Fan - Perennial Stream - Perennial Saline Lake ...... 286 Ephemeral Stream - Saline Pan (Dry Mudflat) - Eolian Dunefield ...... 288 Perennial Stream - Perennial Lake - Eolian Dunefield ...... 292 Springs - Saline Pan (Perennial Lake) ...... 293 Ancient Examples ...... 294 Wilkins Peak Member, (Eocene) Green River Formation, U.S.A. .. 294 Ta.jo Basin (Miocene), Spain ...... 298 Newark Basin (Triassic-Jurassic),U.S.A...... 299 Fundy Basin (Triassic-Jurassic), Canada ...... 302 European Rotliegendes and Zechstein (Permian) ...... 303 Other Examples ...... 303 Recognition of Ancient Non-Marine Deposits ...... 304 Economic Aspects ...... 307 Acknowledgments ...... 308 References ...... 309

CHAPTER 4 . PETROL.EUM AND EVAPORITES by Robert J . Hite and Donald E . Anders Introduction ...... 349 Definition of Terms ...... 350 Production of Organic Matter in Evaporite Environments ...... 352 Marine Settings ...... 352 Alviso Salterns. San Francisco Bay. U.S.A...... 353 Solar Lake. Gulf of Elat ...... 355 Orca Basin. Gulf of Mexico. U.S.A...... 360 Lacustrine Settings ...... 361 Lake Van. eastern Turkey ...... 361 Wadi Natrun. Egypt ...... 362 Great Salt Lake, Utah. U.S.A...... 363 Dead Sea, Israel-Jordan border ...... 365 The Spatiotemporal Relationships of Petroleum and Evaporites ...... 366 Evaporites as Source Rocks of Petroleum ...... 368 Evaporitic Carbonates ...... 368 Anhydrites ...... 373 Halite and Potash ...... 375 Lacustrine Evaporites ...... 382 EvaporiteOils ...... 387

Data Center ,09126599985,[email protected], For Educational Uses TABLE OF CONTENTS XI11

Marine Carbonate Oils ...... 388 Anhydrite Oils ...... 388 Halite .Potash Oils ...... 388 Lacustrine Carbonate Oils ...... 392 Paleoproductivity and Preservation of Organic Matter in the Evaporite Environment ...... 393 Mobile Petroleum Systems ...... 397 Vitrinite Reflectance Suppression ...... 398 Evaporites and Reservoirs ...... 399 Conclusions ...... 402 References ...... 406

CHAPTER 5 . HALOKINESIS. CAP ROCK DEVELOPMENT. AND SALT DOME MINERAL RESOURCES by J . Richard Kyle and Harry H . Posey Introduction ...... 413 Geological Setting of Major Salt Dome Provinces ...... 415 Gulf Coast of Southern North America ...... 415 Maghreb Region of Northwestern Africa ...... 419 The Mechanics of Diapirism ...... 420 General ...... 420 Fluid Infiltration vs . Sediment Loading ...... 424 Timingof Diapirism ...... 426 Fluid Migration Around Salt Diapirs ...... 427 Cap Rock Formation ...... 431 General ...... 431 Fluid Volumes Required for Cap Rock Formation ...... 435 Conditions of Cap Rock Formation ...... 436 Salt Dome Mineral Resources ...... 441 General ...... 441 Energy Resources ...... 441 Sulfur Deposits ...... 443 Metalliferous Deposits ...... 446 General ...... 446 Hockley Dome. Texas. U.S.A...... 449 Fedj el Adoum and Bou Grine. Tunisia ...... 452 Origin of Cap Rock Metalliferous Deposits ...... 455 Sourceof Sulfur ...... 458 Timing of Mineralization ...... 462 Classification of Salt Dome Mineral Deposits ...... 463 General Model for Diapiric Halokinesis. Cap Rock Development. and Mineralization ...... 465 Conclusions ...... 467

Data Center ,09126599985,[email protected], For Educational Uses XIV TABLE OF CONTENTS

Acknowledgments ...... 469 References ...... 470

CHAPTER 6. EVAPORITES. EVAPORITIC PROCESS AND MlNEiRAL RESOURCES J . Richard Kyle Introduction ...... 477 Evaporites as Mineral Resources ...... 478 Introduction to Evaporite Formation ...... 478 Gypsum and Anhydrite ...... 483 Salt and Potash ...... 485 Borates ...... 493 Sodium Carbonate and Sodium Sulfate ...... 497 Bromine, Iodine, Lithium and Nitrogen ...... 503 Evaporites and the Origin of Ore-forming Solutions ...... 505 Metalliferous Saline Formation Waters ...... 505 Geothermal Systems ...... 506 Zeolitization in Alkaline Lakes ...... 509 Evaporite Alteration to Precipitate Valuable Commodities ...... 511 Bioepigenetic Sulfur Deposits ...... 511 Carbonate-hosted Zinc - Lead Deposits ...... 511 Sedimentary Copper - Silver Deposits ...... 515 Sedimentary Exhalative Zinc - Lead - Copper - Silver Deposits ...... 516 Superior-type Iron Formations ...... 519 Evaporite-associated Barite and Celestite Deposits ...... 521 Evaporites and Igneous-hosted Ore Deposits ...... 523 Evaporite Solution to Create Ore-hosting Porosity ...... 524 Evaporation as a Hydrologic Agent in Ore Formation ...... 525 Conclusions ...... 526 Acknowledgments ...... 527 References ...... 527

INDEX ...... 535

Data Center ,09126599985,[email protected], For Educational Uses Chapter I

MARGINAL MARINE HALITE: SABKHAS AND SALINAS

C. Robertson Handf ord

INTRODUCTION

Evaporite sedimentology is the "late-bloomingstepchild" of sedimentary geology. From the late 1950's through the 1960's carbonate and clastic sedimentology blossomed when classic studies of various depositional systems, such as deltas, barrier islands, reefs, and carbonate tidal flats were conducted. These studies resulted in the development of numerous clastic and carbonate facies models that gave petroleum geologists new insight to sedimentary rocks, and in particular to the exploration for hydrocarbons in reservoir-prone rocks. Sedimentological studies emphasized reservoir rocks but rarely investigated the rocks which commonly interfinger with and overlie reservoirs to form stratigraphic seals. Evaporites,for example, may be the most important rock type to form stratigraphic seals to carbonate reservoirs. As a result, geologists learned how to recognize stratigraphic sequences formed by, for example, prograding deltas (Scruton, 1960) many years before details of the saline pan succession were eclucidated by Lowenstein and Hardie (1985). Furthermore, today, nearly 30 years after Folk (1959) and Dunham (1962) proposed their carbonate rock classifications, no one has yet devised a widely accepted petrologic classification of, for example, sedimentary halite. Numerous factors are responsible for the late development of evaporite sedimentology. Halite is not a cooperative subject for physical examination. Its extreme susceptibility to dissolution, recrystallization, and flowage discouraged sedimentologists from attempting to identify and distinguish primary and secondary halite, two steps that are necessary to building a viable classification. In addition, much of what we know about ancient sedimentary facies and their depositional environments was derived from the study of modern sedimentary environments. Many of the classic carbonate and siliciclastic environments are fairly accessible and, in some cases, lie in paradise-like settings. However, this is not the case with modern evaporite environments. Most evaporitic environments are remote and isolated in hot, dry climates where virtually no one would dream of having a vacation. They are logistically difficult to reach and maneuver around in, and some are located in politically unstable regions of the world. To compound the problem, geologists have extremely limited opportunities to visually examine ancient evaporites (Kendall, 1979). There are

Data Center ,09126599985,[email protected], For Educational Uses 2 MARGINAL MARINE HALITE virtually no unaltered surface exposures of ancient evaporites and there is a shortage of evaporite cores from the subsurface, relative to other rock types. The beginnings of a practical sedimentary viewpoint toward halite can probably be traced back to Sorby’s (1 858) crystal growth experiments, Walther‘s (191 2) description of rippled modern halite deposits in Mexico, and the pioneering sedimentological work of Dellwig (1953, 1955) and his co-workers (Dellwig and Brigs, 1952; Dellwig and Evans, 1969). Petrographic studies of Permian halite by Schaller and Henderson (1932), Stewart (1949, 1951a,b) and Jones (1965) were marked by detailed descriptions and exceptional illustrations, from which these researchers pioneered the recognition of many primary and diagenetic textures in halite and developed paragenetic sequences. Truly insightful sedimentological approaches were presented by Hardie (1968), Wardlaw and Schwerdtner (1966), and by, perhaps two of the most important contributions yet, Shearman’s (1970) cogent and penetrative description of halite rock from a modern salt pan, and the experimental work of Arthurton (1973). All of the above laid the groundwork for subsequent work on both modern and ancient halite-bearing sediments (Arakel, 1979, 1980; Brodylo and Spencer, 1987; Casas and Lowenstein, 1989; Castens-Seidell, 1984; Fracasso and Hovorka, 1986; Gornitz and Schreiber, 1981; Handford, 1981,1982,1987,1990;Hovorka, 1987; Logan, 1987; Lowenstein, 1982,1987, 1988; Lowenstein and Hardie, 1985; Lowenstein and Spencer, 1990; Lowenstein et al., 1989; Moretto, 1987; Orti Cab0 et al., 1984; Presley, 1987). This chapter presents a summary of marginal marine halite in terms of (1) the fundamental controls on precipitation, (2) halite crystallography and crystal growth, (3) halite deposition and syndepositional to early postdepositional alteration, (4) marginal-marine salina and sabkha depositional settings containing halite, and (5)their vertical sequences.

FUNDAMENTAL CONTROLS ON MARINE HALITE PRECIPITATION

Halite is commonly precipitated as a chemical sediment in sedimentary basins from marine, continental, or mixed marine-continental brines that are supersaturated with respect to NaCI. Supersaturation and precipitation will occur only in sedimentary environments where evaporation exceeds the amount of input supplied by rainfall, rivers, and the oceans (Fig. 1.1). In order for halite to precipitate from a standing body of brine (marine origin), irregardless of its size, there must be partial or complete isolation from open sea so that free circulation cannot take place (Stewart, 1963). Both Lucia (1972) and Kendall (1988, 1989) argued convincingly for complete surface disconnection of evaporite basins from the sea.

Data Center ,09126599985,[email protected], For Educational Uses CONTROLS ON MARINE HALITE PRECIPITATION 3

SEEPAGE REFLUX

Fig. 1.1. Inflow of meteoric, marine, and hydrothermal waters into evaporite basins is offset by excessive evaporation, reflux, and possibly outflow.

There is, however, continuing debate if some ancient evaporite basins were filled with salts precipitated in widespread pans and sabkhas during desiccation (Hsii, 1972) or from deep-water saturated brines (Schmalz, 3 969). The challenges have been flowing from both sides of the issue ever since the theories were presented. Studies promoting deposition in deep water (Anderson et al., 1972; Hite, 1970) have been challenged (Kendall, 1987a,b, 1988; Kendall and Hanvood, 1989), and evaporites long-claimed to be of shallow-water to sabkha origin (Hsii, 1972; Ryan and Hsii, 1973) are now being questioned (Dietz and Woodhouse, 1988, 1989; Schmalz, 1989). Perhaps the truth lies somewhere in between. Seawater must be greatly concentrated by evaporation before marine halite can be precipitated (Herrmann et al., 1973; Usiglio, 1849). For example, McCaffrey et al. (1987) observed that seawater brines from a solar salt production facility in the Bahamas precipitate CaCO, at a brine concentration of about 1.8 times that of seawater. Gypsum begins to precipitate at a brine concentration of 3.8 times seawater, and halite at a concentration of 10.6 times seawater. Magnesium sulfate first precipitated at brine concentrations of about 70 times seawater and potassium-bearing phases began to precipitate at 90 times seawater concentration. In order to reach the concentration values given above to precipitate evaporite minerals in a natural system, a large volume of seawater must be evaporated. Theoretical calculations indicate that a column of seawater 300 m deep, which is completely evaporated, would precipitate -4.8 m of precipitated solids, and of that modest amount, halite would account for about 3.7 m (Fig. 1.2). Similar calculations would suggest, for example, that a single Paleozoic halite unit from the North American craton with a thickness of 50 m would have required, under this simple evaporation mechanism, a seawater column -4OOO m deep. Clearly, cratonic evaporative basins were never that deep, and simple evaporation of deep column of

Data Center ,09126599985,[email protected], For Educational Uses 4 MARGINAL MARINE HALITE

ORDER OF PRECIPITATION SEAWATER AND PROPORTIONS-

I 085 M K 8 Mg SALTS I -

300 M 3 7 M HALITE

173 M GYPSUM 0192 M CARBONATES -.---4 8 M SOLIDS

Fig. 1.2. EvaporationL of a column of seawater 300 m deep theoretically should produce about 4.8 m of solids, of which halite makes up about 78% or 3.7 m. Modified from Holser (1979).

seawater is an unviable mechanism to account for thick evaporites in shallow depositional basins. Ochsenius (1 877) proposed a barred basin model that accommodates the precipitation of thick evaporites in shallow-water basins by (1) importing large volumes of seawater across a barrier, (2) evaporating seawater at a greater rate than is being supplied, and (3) limiting the outflow so that the basin becomes saltier (Fig. 1.1). Isolation of a basin is achieved by a topographic barrier, such as a continuous or nearly continuous reef, barrier island, a dense network of islands separated by small channels, or a tectonic sill. If significant amounts of marginal marine evaporites are to accumulate in the isolated basin, a barrier must allow periodic or continuous inflow of seawater. Inflow can take place through inlets, or channels cutting across a barrier, or by seepage through a continuous barrier. Size of an inlet is critically important to initiation and maintenance of evaporite deposition. Lucia (1972) measured the salinity reached in restricted basins of arid regions by calculating the ratio between the surface area of an evaporite basin (&) and the cross-sectional area of the inlet (At). The ratio reflects the balance between the rates of water loss and gain in evaporite basins. For halite precipitation, the cross-sectional area of an inlet must be at least eight orders of magnitude smaller than the surface area of the basin (Kendall, 1988, 1989; Lucia, 1972) (Fig. 1.3). Lucia (1972) concluded that while bedded gypsum can be produced in standing bodies of water which have a connecting channel to the ocean, the channel would have to be so small as to be

Data Center ,09126599985,[email protected], For Educational Uses CONTROLS ON MARINE HALITE PRECIPITATION 5

_.BODY OF WATER A,/A,

RED SEA 0 3 x 104 PERSIAN GULF x 105 SHARK BAY 2 x lo4 LAGUNA MADRE 9 x lo5 KARABOGAS GOL 2~ 10' DANAKIL DEPRESSION I- I- GYPSUM FIELD FIELD LAKE MACLEOD NO INLET - HUTT LAGOON I 101 , , , , 0 1,50 100j 150 200 250 300 350 SALINITY %o

Fig. 1.3. Relationship between the ratios of the cross-sectional areas of an inlet and the surf ace area of an evaporite basin and salinities, with examples from several restricted marine basins. Modified after Lucia (1972).

geographically insignificant, and that salt deposition must certainly indicate complete surface disconnection from the ocean. This is certainly true today; all modern basins, in which the evaporating waters consistently attain gypsum, halite, or bittern-field concentrations, are disconnected from the sea by some topographic barrier (Logan, 1987). Where inflow of seawater occurs by way of groundwater seepage through a barrier, there must be a hydrodynamic drive, such as an elevation- or density-driven head difference. Lake MacLeod is a classic example of the former. This 200-km2salina lies 3-4 m below sea level and 15 km inland from the west coast of Australia. Despite the slight elevation head difference (3-4 m) between the salina and the Indian Ocean and the width of the barrier, there is sufficient hydraulic head to drive about 404 x 106 m3 of sea water per year through the barrier, at an overall velocity of about 16.09 x lo" cm/sec (Logan, 1987). A barrier can be highly permeable to inflowing seawater and still not exceed evaporative losses. Logan showed that most of the seawater seeping into Lake MacLeod flows through vuggy and cavernous Tertiary limestone aquifers. Kendall (1989), Logan (1987), and Maiklem (1971) claimed that even if a highly permeable barrier is present between an evaporite basin and an adjacent sea, seepage through the barrier cannot offset losses in the evaporite basin due to evaporation. Citing the Middle Devonian evaporite-bearing Elk Point Basin and the flanking Presqu'ile barrier as an

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example, Maiklem (1971) calculated that seepage rates through the Presqu’ile barrier, with a permeability of 73 darcies, would not have been high enough to offset evaporative losses in the basin. Kendall (1989), who used more realistic permeability values of 0.9-2.7 darcies for the barrier, calculated that evaporative losses in the basin would have been 200-800 times greater than the inflow. Large volumes of inflowing seawater must evaporate in order to precipitate small volumes of evaporites. Evaporation is a function of many variables, including temperature, temperature contrast of air and water, wind speed and local turbulence, and relative humidity of the atmosphere. It is widely known, though, that as seawater evaporates and its ionic concentration rises, it becomes increasingly difficult to maintain continued evaporation (Harbeck, 1955; Kinsman, 1976; Schreiber and Helman, 1989). Increased dissolved salt concentration decreases the ionic activity of H20in a solution and, as a result, lowers the equilibrium water vapor pressure exerted by the solution (Kinsman, 1976). This means that evaporation will virtually cease at about the point of halite saturation unless the relative humidity is very low. Kinsman (1976) calculated that halite can be precipitated only where the mean relative humidity of the atmosphere is less than 76%. Schreiber and Helman (1989) suggest an even lower relative humidity (65%). Relative humidity distribution in low latitude, sea level atmospheres is related to proximity to the ocean (Kinsman, 1976). The air masses above oceans have mean relative humidity values close to 100%. Mean relative humidity drops to 70-80% along coastal areas, and intracontinental values may be as low as 60%. To complete the requirements for evaporite deposition to occur, brines must not escape from the basin at a rate greater than that of the inflow. If the evaporite basin is not totally disconnected from the sea by a sill, escape may take place by density underflow if the brine level is above the sill threshold (Scruton, 1953). Owing to their high density, brines may also escape by seeping through the underlying sediments (Adams and Rhodes, 1960; Logan, 1987). Loss of brines translates to a removal of salts that could have been precipitated in the evaporite basin. If the rate of escape exceeds inflow, the evaporite system terminates (Logan, 1987).

HALITE CRYSTALLOGRAPHY AND CRYSTAL GROWTH

Halite is a soft (H = 2.5) rock-forming mineral which belongs to the cubic crystal system (Fig. 1.4). Although colorless or white when pure, halite can typically range from orange or red to yellow, brown, gray, or even blue, depending on the type and amount of impurities, such as iron oxides, clay particles, and organic matter. In the case of blue

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N a+

A 0

Fig. 1.4. The structure o,f halite (A) shows that each Naf cation is surrounded by and related to 6 neighboring Ct anions. The weak electrostatic charges are spread over the entire surface of the nearly spherical ions, such that the bonding is perfectly ionic (Hurlburt and Klein, 1985). The resulting cubic crystal form and cubic cleavage with three planes intersecting at right angles are shown in (B) and (C). halite, its coloration may be due to a permanent disturbance in the crystal structure (Przibram, 19.56) caused by one or more of the following: (1) radiation from the K-40 radioisotope contained within surrounding potassium minerals, (2) heat, and (3) deformation (Shlicta, 1968). Thorough discussions are presented by Sonnenfeld (1984). There are three planes of cleavage in halite with the { 100) being perfect (Fig. 1.4). Its index of refraction is 1S44 (Deer, Howie, and Zussman, 1962) and it has zero birefringence. It is petrographically isotropic and consequently black under crossed nicols. The crystal structure of halite was the first to be analyzed by x-ray diffraction. Bragg (1914) demonstrated that sodium and chlorine ions are arranged alternately along rows parallel to the edges of a facecentered cubic crystal lattice (Fig. 1.4). Thus, each Na' ion is surrounded by six Cl- ions and each Cl ion by six Na' ions in octahedral coordination. This arrangement shows that there is no grouping of atoms in the NaCl structure. Halite is a perfect example of ionic bonding as weak electrostatic charges are spread over the entire surface of the nearly spherical ions (Hurlburt and Klein, 1985). Halite cell dimension is 5.639 A. Sedimentary Occurrences of halite in both natural and man-made (artificial salinas or evaporating pans) settings are due to precipitation from aqueous solutions supersaturated with respect to halite. Crystallization can take place in any of the following sites (Fig. 1.S): (1) at the brine/air interface, (2) on the floor of a brine pool, (3) in brine-soaked sediments as displacive/incorporative crystals, (4) in brine-filled vugs

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3 INTRASEDIMENT 6 CAPILLARY FRINGE 1 BRINE/AIR INTERFACE I 2 BRINE POOL FLOOR I \ 1

\ 5 FRACTURE --I----- 4 CAVITY

Fig. 1.5. Common sites where halite precipitate within evaporite basins include (I) the brine/air interface, (2) the floor of a brinepool, (.?) in a sedimentary matrix, (4)in vugs or dissolution cavities, (5) in fractures as vein-filling fibrous cement, and (6) as efflorescent crusts at the surface or in the capillary fringe of the vadose zone. or dissolution cavities, (5) in fractures as fibrous cement, and (6) as efflorescent crusts in the capillary fringe of the vadose zone. Halite crystals may grow as single, isolated crystals, or be grouped together in clusters as beds or crusts and veins. The growth of a single crystal of halite is generally preceded by the random formation of numerous potential nuclei; however, most of the nuclei do not survive, but instead tend to redissolve in the saturated solution (Hurlburt and Klein, 1985). Dissolution of these tiny nuclei occurs because they present a large surface area to the solution and, thus, have a high free energy. If nuclei grow rapidly enough, thereby reducing surface energy, they may reach the critical size required for survival and can become permanent nucleation sites (Hurlburt and Klein, 1985). For ionically bonded crystals such as halite, the energy of attachment is greatest at crystal corners, intermediate at edges, and least in the middle portions of crystal faces. As explained by Gornitz and Schreiber (1981), the morphology of growing halite crystals is dependent, in part, upon saturation levels and rates of ionic diffusion. Under normal conditions of crystal growth in saturated solutions, cube faces develop parallel to themselves, despite the fact that ions diffuse preferentially toward the crystal corners and edges where the concentration gradients are greater. Gornitz and Schreiber (1981) explain that this occurs because the most rapidly growing faces ((111) or (110)) disappear while the slower growing cube faces survive. However, when supersaturation reaches a critically high point, diffusion takes on greater importance with the result being more rapid deposition on corners and edges to form skeletal or dendritic crystals (Fig. 1.6).

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Fig. 1.6. Depressed cubic faces of a skeletal, hopper-shaped halite cube. Modified after Hurlburt and Klein (1985).

Growth bands frequently develop in halite crystals grown from an aqueous brine and reflect changes in temperature, growth rate and brine composition (Shlicta, 1968). They are most commonly present as fluid inclusions and sediment inclusions. Growth bands may also be induced in the laboratory by irradiating the crystals (Przibram, 1956). Apparently, the presence of polyvalent impurity ions and ionic vacancies provides traps for electrons and forms holes such that the defects are highlighted by radiation (Shlicta, 1968). Fluid inclusions are tiny brine-filled cavities that form during crystal growth (Figs. 1.7 A,B); for primary halite crystals which have not undergone recrystallization, the cavities contain actual samples of the fluid from which the crystals grew. These are

Fig. 1.7. (A) Chevron arrangement of fluid-inclusion-richgrowth bandsin surf ace halite crust from Baja California, Mexico. (B) Chevron arrangement of fluid- inclusion-rich growth bands in Permian-age San Andres Formation of Texas. Close examination reveals the presence of vapor bubbles in several of the fluid inclusions. Photographs by S. Hovorka.

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referred to as primary fluid inclusions, and they are generally arranged and concentrated in zones parallel to the cube face. Inclusion-rich zones form laminae approximately 0.1 to 0.3 mm thick and they alternate with clearer, inclusion-free or inclusion-poor zones (Holser, 1979). Where abundant, the inclusion-rich zones impart a white or cloudy appearance to the halite crystals and are believed to signify episodes of rapid crystal growth (Shearman, 1978). In contrast, zones that are relatively poor in inclusions probably grew at slower rates. Alternating fluid inclusion-rich and inclusion-poor zones are common in halite crystals that have grown in both natural and man-made brine pans. Presumably, this pattern is what would be expected from a daily alternation of rapid crystallization in the daytime (high evaporation rates) and slow crystallization at night (lower evaporation rates) (Holser, 1979).

DEPOSITION AND EARLY DIAGENESIS OF MARGINAL MARINE HALITE

Sedimentary accumulations of marginal marine halite fall generally under four categories or modes of occurrence: (1) subaqueous cumulates, (2) subaqueous bottom precipitates, (3) intrasediment precipitates, and (4) clastic particles. Of these, the first three account for most sedimentary halite recognized in both natural and man-made settings. Although clastic halite is a minor component, its presence certainly is an indicator of high physical energy, most likely in association with a shoreline. After deposition of a halite crystal, it then becomes subject to a range of early postdepositional processes that may alter the original fabric slighlty or severely and even completely remove it by dissolution.

Subaqueous Cumulates

Cumulate is a term borrowed from igneous petrology where it is defined as an igneous rock formed by the accumulation of crystals that settle out from a magma by the action of gravity (AGI Dictionary of Geological Terms, 1984). The term is also ideally suited to define halite crystals that have at first precipitated from a brine, most commonly at the brine/air interface, only later to settle to the bottom of the brine pool under the effects of gravity (Fig. 1.8). Lowenstein and Hardie (1985)first used the term in the sedimentary context of evaporite deposition. Crystals that initially precipitate from brine pools usually form cumulates. As a brine becomes concentrated by evaporation to the point of halite precipitation, halite will usually begin to crystallize at the brine/air interface (Figs. 1.8, 1.9A) as floating, millimeter-size crystals, and then, later, these expand into centimeter-size crystals

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HOPPER RAFT BRINE

OVERGROWTHS

Fig. 1.8. The formation o,f halite cumulates begins with crystal nucleation at the brine/air interface. Individual crystals grow as pyramidal hoppers and plates, coalesce with other floating crystals to form rafts and eventually f ounder and sink to the bottom. There they make up loosely packed layers of cumulates and add on halite cement as overgrowths.

(Arthurton, 1973; Dellwig, 1955; Lowenstein and Hardie, 1985; Shearman, 1978). Arthurton's (1 973) experimentally produced halite, which precipitated at the brine-air interface, consisted of four different varieties of crystals: (1) rectangular plates suspended by surface tension; (2) cuboids with cube corners pointing downwards and each crystal suspended either from a corner or, in the case of larger crystals, from the perimeter of a horizontal truncation face; (3) hollow, inverted,four-sided hopper crystals which were in part suspended by surface tension and also buoyed up by their own displacement of brine (Figs. 1.8,1.9A); and (4) hollow, inverted, six-sided hoppers (Fig. 1.6). The first variety formed at air temperatures in the range of 5-25"C, while the remaining three types formed at air temperatures between 525°C. Four-sided, inverted pyramidal hoppers (Figs. 1.9A, B) are probably the most common type of crystal that forms at the brine-air interface during the early stages of halite precipitation. These crystals precipitate at the brine surface where evaporation is occurring and the concentration of dissolved NaCl is high, and grow laterally outward along cube edges. Single crystals occasionally sink to the bottom, but only if the buoyancy forces holding up these boatlike crystals are destroyed or if surface tension effects, which cause the crystals to adhere to the brine surface, are disturbed, such as

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Fig. 1.9. (A) Drawing of a four-sided, inverted pyramidal hopper crystal with a stepped cube face. (B)Floating, pyramidal hopper crystals in shallow brine pools. Knife is 90 rnrn long. (Lowenstein and Hardie, 1985). Photograph by R. W. Mitchell. when waves are whipped up by the wind. Othenvise, drifting halite crystals will coalesce to form weakly attached rafts which may continue to grow, both by lateral crystal growth and continued coalescence with other drifting crystals and rafts (Fig. 1.10A). Arthurton (1973) showed that growth by coalescence commences with the formation of chains, then progresses to the linking of chains into nets, and finally the formation of continuous mats, or rafts, covering several cm’. These rafts may drift with the wind and eventually wash against the shore of a shallow brine pool where they accumulate like driftwood on a beach (Fig. 1.10B). In some cases, however, the rafts may sink before reaching a shore. This commonly occurs after several days of growth by crystallization at the brine-air interface and after more drifting crystals are taken aboard,

Fig. 1.10. (A) Floating halite rafts at Salina Omotepec, Mexico. Note that the water depth is only a few cm and that the bottom is covered with a halite-encrusted cyanobacterial mat. (B) Cumulate halite. Stranded rafts along the shore of a halite crystallizing pond in Bonaire, Netherlands Antilles. Photographs by C. R. Handford.

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as if they were boats adrift at sea. As a result the rafts become heavy and waterlogged. This, in combination with sudden gusts of wind that disturb the placid brine surface, will cause the rafts to founder like an ocean vessel in distress and quickly sink to the bottom. If the waves are not too energetic, the delicate rafts may remain intact as they drift downward to the bottom. If the cycle of pyramidal hopper growth, formation of rafts and sinking continues during brine pool evaporation, the bottom will become the resting place of layer after layer of sunken rafts, pyramidal hoppers and platy crystals (Fig. 1.1 1) to form a loose aggregate of what can be called an "immature" halite (Lowenstein and Hardie, 1985).

Subaqueous Bottom Precipitates

Although halite precipitation at the brine/air interface can be the dominant mode of crystal growth in shallow brine pools, especially during the initial stages of halite deposition, nucleation and growth of halite on the bottom of brine pools are also very important (Fig. 1.12). Cornet-shaped crystals, cubes, and upward-growing chevrons precipitate on the floor both as bottom-nucleated crystals and as syntaxial overgrowth cements on sunken rafts and other cumulates that have settled to the floor (Fig. 1.12). The formation of bottom-nucleated crystals and syntaxial overgrowths as cornets, chevrons, and cubes marks an important stage in halite precipitation, for it not only is responsible for upward growth and thickening of a halite layer, but it is also responsible for cementation and, hence, the development of a more "mature", and harder halite rock from what was once a loose aggregate of crystals and rafts. In Arthurton's (1973) experiments, bottom nucleation occurred mostly during cooler nighttime hours; new crystals formed on previously unoccupied parts of the brine-pool floor each night. Tiny, transparent rectangular plates formed first, but within a few hours, the plates developed into cuboids. Crystal size seemed to be determined, in part, by brine depth, because brine pools 10 cm deep produced fewer but larger crystals than brine pools 2 cm deep. Bottom-nucleated crystals and cumulates are modified on the brine-pool floor by the addition of overgrowth cements. Where uncrowded conditions prevail, equal growth of exposed crystal faces may occur, thus preserving a cuboid form in the growing crystals, but if neighboring crystals crowd one another and compete for space, the crystals are forced to grow vertically upward (Wardlaw and Schwerdtner, 1966; Shearman, 1970, 1978; Lowenstein and Hardie, 1985). Although vertically upward growth of crystals is dependent upon their competition for space, their morphology is dependent upon the attitude of the parent crystals (Ar-

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Fig. I. I I. Layered cumulates. (A) Modern layered halite curnulater (white)from Salina Omotepec, Mexico. This modern halite-gypsum cemented sediment, held upside down, consists of thin layers of halite rafts (white) overlying pink-tan gypsum. (B) Mold oj founderedhaliteraf t with latarallydirected chevron-overgrowths. Silurian, Indiana. After Schreiber (1988). Photograph by C. F Kahle. (C) Photomicrograph of modern halite crust in (A)showing clear, cubic overgrowth cement on foundered rafts. (D) Crystals less than 1.0 mm in diameter make up foundered rufts prererved in San Andres Formation (Permian) from Texas. C and Dphotographed by S. Hovorka. thurton, 1973; Lowenstein and Hardie, 1985). Where a halite cube or hopper is oriented with a corner or edge pointing upward, a template is established such that syntaxial overgrowth will preferentially develop with chevrons also pointing upward (Fig. 1.13). Crystals oriented with a cube face parallel to the floor, or facing upward, will most commonly develop cornet-shaped overgrowths (Figs. 1.14A,B), which are either flat- topped or have a depressed cube face and widen upward (Arthurton, 1973; Lowenstein and Hardie, 1985). Cuboids form only where crystals are widely spaced and not in competition with each other. In the long term, however, and where cubes grow uniformly in competition with each other, cube corners and edges will advance more rapidly than faces. Thus, in a bed made up of randomly oriented cubes, which are competing for space so that they are forced to grow upwards from the floor of a brine pool, the cubes whose edges and

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- _11 CHEVRONS

Fig. 1.12. Bottom precipitated halite includes cuboids, chevrons ({I I I) facing up), and cornets. Cuboids would be prevalent only where crystals are uncrowded and not competing for space with other crystals. In most cases, crowding and competition do occur and result in the development of vertically oriented fabrics that reflect the preferred survival of chevrons and cornets.

corners (111) are directed upwards will overtake those with cube faces directed upwards. The final product will then consist largely of vertically-oriented elongate crystals with chevron-shaped growth patterns. As the rate of crystallization varies, perhaps due to temperature changes associated with day-night cycles, vertically oriented crystal-growthfabrics are highlighted by alternating bands of cloudy, fluid-inclusion-rich

Fig. I. 13. Bottom-precipitated chevron halite. (A) Experimentallygrown halite. Chevrons with upward pointing fluid-inclusion zones grew on bottom of a 3 cm deep pool over previously f oundered halite hoppers. Chevrons are overlain by newly f oundered cumulates. (B) Ancient example of chevron halite. Vertically upward-grown chevron halite with compromise boundaries between individual crystals. Note the alternating bands of clear halite and f luid inclusion-rich halite, San Andres Formation, Texas. Photographs by S. Hovorka.

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Fig. 1.14. (A) Drawing of cornet crystal illustrates centrally depressed cube-face. (B) Bottom-precipitated cornets f rom Salina Omotepec, Mexico. Photograph by S. Hovorka.

halite and clear halite with few or no fluid inclusions (Fig. 1.7B). Alternations develop only where the brine experiences rapid variation in halite concentration, or temperature, circumstances possible only in shallow brine depths (Kendall, 1979). At this more "mature" stage of halite growth in a brine pool, a wide variety of halite crystal morphologies and fabrics might be expected to result, and to a certain extent, various types do form. However, when given sufficient time and consistently evaporative conditions, a shallow brine pool, which is allowed to completely evaporate, will precipitate a layer of halite that is most commonly made up of a basal zone of cumulates overlain by syntaxially grown, vertically oriented chevrons and cornets (Lowenstein and Hardie, 1985). The relative proportion of cumulates to bottom precipitates may be dependent upon the rate of evaporation of the brine pool (Lowenstein and Hardie, 1985), water depth, and energy (waves and currents). Extraordinarily delicate and complexly branching, dendritic halite crystals have been observed growing on the floors of shallow brine sheets in artificial salinas and documented from Cambrian pseudomorphs of chert after halite (Fig. 1.15). Southgate (1982) described a suite of highly unusual, sheetlike dendritic halite pseudomorphs that are oriented parallel to bedding in Cambrian cherts, suggesting nucleation on the bottom of a shallow brine sheet. The crystal morphologies, however, were so unusual that until laboratory experiments of crystal growth from shallow brines were conducted, they almost defied explanation. Southgate's (1982) experiments and description of Cambrian pseudomorphs led to the documentation of several crystal morphologies

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~~ Fig. 1.15. Excrescent dendritic halite crystals pseudomorphed by chert from Cambrian rocks of Australia. Reticulate ridge halite in (A) has right angle intersection of ridges. (B) Pagoda crystals are present along the ridges. Photographs by P. Southgate.

(Figs. 1.16, 1.17): (1) reticulate (dendritic) halite, (2) horizontal chevron halite, (3) reticulate (dendritic) ridge halite, and (4) pagoda halite. Experiments have shown that if a brine is allowed to reach high supersaturation or if certain organic/inorganic substances are added to the brine, the external crystal form of rapidly precipitating halite crystals can be drastically modified (Cooke, 1966). Extremely dendritic or excrescent crystals, for example, precipitate in either case. Supersaturated layers form at the brine/air interface, but they are extremely thin and difficult to maintain because of excessive halite nucleation at the brine/air interface in the form of halite rafts. In order for dendritic crystals to form on the bottom of a pool, the supersaturated layer must come in contact with the bottom. This can occur when (1) the supersaturated layer sinks to the bottom during overturn of a density stratified body of water, or (2) if evaporation has shallowed the parent brine so much that only a thin brine sheet remains and it is equal to the thickness of the supersaturated layer (Southgate, 1982). With respect to the addition of small quantities of certain organic and inorganic substances, such as ferricyanide or humic acid, they promote dendritic crystal growth by lowering the saturation levels necessary for dendritic growth to proceed.

Intrasediment Precipitates

The third major mode of halite crystal growth is by intrasedimentary precipitation (Fig. 1.18),that is, halite which has grown within soft, brine-soaked sediment (siliciclastic clay, silt, sand, or even carbonate mud) as isolated crystals or as cement. Where sediment has been trapped as inclusions in the growing crystals, the crystals may be

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A PAGODAREGROWTH/+

VARIABLE SCALE

0 UP TO CM

B EVAPORATION DISH

5CM- 1 k------4 RETICULATE HALITE (DENDRITIC) 4 C PAGODA HALITE

2 3 HORIZONTAL RETICULATE RIDGE CHEVRON HALITE HALITE (DENDRITIC)

APPROXIMATE CRYSTAL SCALE

DlRECTlON 0- cI1 '

Fig. 1.16. Diagrams depicting in (A) the pattern of ridge development and pagoda regrowth and in (B) the relationship between skeletal crystal morphology and sites of experimental crystal growth within an evaporation dish. After Southgate (1982). described as sediment-incorporative(Lowenstein, 1982). Crystals that push aside mud are displacive and where displacive and/or incorporative crystals are crowded and competing for space, a chaotic mud/crystal fabric develops (Handford, 1981). According to Gornitz and Schreiber (1981), sediment is incorporated into the crystals most readily when crystal growth is rapid; sediment is pushed aside or displaced

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Fig. I. 17. Experimentally grown halite with excrescent dendritic morphologies. (A) Reticulate ridge and dendritic halite crystals. (B) Horizontal chevron halite. This crystal nucleated on the curved side of the evaporating dish. The central reticulate-ridgef orm passes into dendrites on the updip side of the crystal. Chevron halite crystals have grown in the deeper parts of the brine pool. Photographs by P. Southgate. when growth rates are relatively slow. Thus, crystals made up of clear halite alternating with inclusions of mud along cube faces (Fig. 1.19), have grown in a sediment matrix, perhaps at different rates, by displacing and incorporating sediment. As in the case of hopper growth at the brine/air interface, hopper-shaped crystals that precipitate within the sediment grow preferentially along cube corners and edges from supersaturated brines. At high levels of supersaturation, diffusion of ions to crystal edges and corners promotes dendritic or hopper-type growth, and if the edges and corners contain crystal defects, rapid growth of hopper-type crystals will be promoted (Gornitz and Schreiber,

DlSPL ACI VE

INCO R POR AT IV E I CHAOTlC

Fig. I.18. Intrasediment precipitation of halite creates displacive crystals (push aside sediment) and incorporative crystals (contains sediment inclusions) that are supported by the sediment matrix. If the nucleating crystals are closely spaced or touching so as tof orm some enf acial boundaries, a chaotic f abric of anhedral to subhedral crystals separated by pockets of sedimentary matrix may result.

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Fig. 1.19. (A-B) Ordovician dolomite with pseudomorphed halite-hopper crystals that grew displacively (sediment-f ree zones) and incorporated carbonate mud (sediment-rich zones) duringgrowth. (B) is a view of a slab cut parallel to bedding. Photographs by C. R. Handford.

1981). Deposition along face centers, at this point, may be further hindered by the projecting crystal edges. They could restrict diffusion of ions to the face centers and consequently inhibit deposition there. There are two exceptional modern examples of hopper-shaped cubes of halite. Hoppers rangingfrom 5 to 10 cm in diameter (Fig. 1.20) are present less than 2 m below the surface in a matrix of carbonate mud along the western shore of the Dead Sea (Gornitz and Schreiber, 1981). An origin by upward diffusion is favored by Gornitz and Schreiber (1981), especially where the sediments were subaerially exposed and only intermittently inundated. Continuation of crystal growth would have been promoted as brines in the capillary zone were constantly renewed by lateral recharge from the south basin of the Dead Sea. However, where the sediment has been continually submerged by several meters of water, an origin by downward diffusion would be favored. In this case crystal growth would have been enhanced if the sediment was inundated for a sub-

Fig. 1.20. Recent, giant displacive halite cubes from the Dead Sea (largest crystal about 10 cm in diameter). Photograph by B. C. Schreiber.

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stantial part of the year. Abundant hopper-shaped crystals of halite, larger even than the Dead Sea crystals, are present at Bristol Dry Lake, California. Crystals are commonly more than 10 cm in diameter and, unlike the Dead Sea crystals, which were dredged up, those from Bristol Dry Lake are found in place (Handford, 1982). Blocky cubes, which are slightly depressed along face centers, and highly skeletal crystals are exposed in walls of pits and trenches that were excavated on the playa floor by salt companies. Crystals are both sediment incorporative and displacive. Because dendritic and skeletal crystals are products of rapid crystal growth, they are believed to have grown more rapidly than the cubes with slightly depressed face centers. For all varieties and ages of intrasediment halite, the relative amount of matrix versus displacive and incorporative halite can be extremely variable (Fig. 1.21). There are mudstones with scattered, isolated crystals making up less than 1% of the rock volume, rocks which are made up of approximately equal amounts of matrix and halite, and at the other end of the spectrum, rocks with mostly halite and relatively little matrix yielding a chaotic fabric (Handford, 1981, 1982). In the last case, the final product may have been produced by dissolution of primary subaqueous halite followed by reprecipitation of the NaCl beneath the surface as displacive and incorporative halite. This mechanism presumably could occur repeatedly as shallow salinas and playas are sub-

Fig. 1.21. Core slabs from Permian evaporites of Texasshow a variation in the amount of matrix versus halite. (A) Large, euhedral halite cubes in mudstone, San Andres Formation. Scale in cm. (B) Typical chaotic halite-mudstone fabric with subhedral to anhedral crystal boundaries. Halite and mudstone are roughly equal in volume. Seven Rivers Formation. Scale in cm. Photographs by S. Hovorka.

Data Center ,09126599985,[email protected], For Educational Uses 22 MARGINAL MARINE HALITE jected to multiple cycles of flooding, evaporative concentration, and desiccation. Displacive and incorporative halitecrystal growth is dependent upon diffusion. However, does intrasediment precipitation prefer to take place in sediments that are subaerially exposed or in those that are covered by brine? It may be possible for intrasediment precipitation to occur in both situations, but the direction of brine diffusion may differ in each of those two possible scenarios (Gornitz and Schreiber, 1981). If several meters of nearly saturated brine, which overlie the sediment, are seasonally heated, downward diffusion may be promoted and halite precipitation could occur 1-2 m below the surface. Where brine-soaked sediment is subaerially exposed, upward diffusion probably takes place by capillary action. Summer or daily heating of the surface would promote evaporation and brines from below would diffuse upward to take the place of that which has evaporated. As the brine rises, precipitation would occur at the level where halite supersaturation is reached. The thickness of the zone of halite crystallization will be determined by the thickness of the capillary zone. Thus, relatively thick successions of displacive and incorporative halite may imply the presence of an equally thick capillary zone. And since sediment grain size largely determines the thickness of the capillary zone (fine-grained sediment = thick capillary zone, coarse- grained sediment = thin capillary zone), the development of thick beds of displacive and incorporative halite, which owe their origin to upward diffusion in the capillary zone, is favored in a mud matrix rather than in a sandy one.

Clastic Halite Particles and Hydrodynamic Structures

Rapid cementation and the growth of interlocking halite crystals on the floors of shallow brine pools hinder the formation of halite particles that can remain loose long enough to be physically transported and deposited by waves/currents into primary bedforms. Nevertheless, there are occasions in which foundering cumulates and small bottom precipitates, which have just started nucleating around foreign particles such as peloids or windblown silt, may be abundant enough and uncemented long enough to be concentrated by flows into hydrodynamic structures. Karcz and Zak (1987) described large and small-scale ripples made up of poorly sorted, sand-size halite crystals (single and polycrystalline clusters and aggregates) along a conveyance canal in the Dead Sea salt works. Bedforms were created by unidirection- ally flowing currents of brine (density = 1.29 g/cm’) with velocities of about 30 cm/sec and a flow depth of about 1 m. The authors concluded that the hydraulic behavior of halite/brine is similar to that of quartz/water systems. Ripples have also been noted from the shallow floors of halite crystallizing ponds in Bonaire, Netherlands Antilles.

Data Center ,09126599985,[email protected], For Educational Uses CLASTTC HALITE AND HYDRODYNAMIC STRUCTURES 23

Rippled and crossbedded halite is not restricted to modern halite environments. Ripple forms are preserved along partings in Silurian halite beds from Michigan and New York. Ripple lengths range from 7.5 to 23 cm and heights from 0.6 to 2 cm (Kaufmann and Slawson, 1950). Large-scale crossbeds with sets 30-60 cm thick and foreset slopes of 200 were described by Dellwig and Evans (1969). Additional features associated with mechanical movement and deposition of halite include halite ooids and pisoids. Weiler et al. (1974) discovered halite ooids and pisoids accumulating in a low terrace just above a salt pan halite-encrusted beach in the Dead Sea’s southern basin. The grains are well rounded, polished, spherical, and have diameters between 0.8-4.0 mm. Internally the ooids and pisoids may be composed of single halite crystals or radially oriented halite crystals that have precipitated around nuclei consisting of halite, allochthonous dolomite, limestone, or shale. These coated grains probably form in a similar manner to that of modern carbonate ooids. In moderately agitated water, halite crystals which precipitate at the brine surface sink to the floor, but rather than accumulating as a crust, they are kept constantly in motion by wave agitation. If the energy level is high enough to keep the grains in motion, halite may be precipitated on grain surfaces and subsequently polished as the grains roll around prior to coming to rest on the beach or in deeper, less agitated areas of the salt pan. Modern halite pisoids (1 cm diameter) are commonly present along the windward beaches of artificial salt ponds at the International Salt Company’s Bonaire solar salt works (Handford, 1987,1990) (Figs. 1.22 A,B,C). Like the Dead Sea examples, Bonaire halite pisoids grow when halite is precipitated around wave agitated grains, or nuclei. Precipitation leads to the formation of a radialconcentric structure. Mobile pisoids have smooth surfaces, but when the grains are stationary, tiny halite crystals nucleate on the surface and roughen it so that the pisoids resemble icy snowballs. Because large pisoids require high energy conditions to stay mobile, any cessation in waves, or a significant decrease in wave energy, is quickly translated into rapid cementation of the pisoids by halite cement to form a hard crust along the shores of the artificial ponds. Beach pisoids, which are washed up by wind-generated waves and stranded above the brine pool, are subject to freshwater corrosion and may be destroyed.

Syndepositional and Early Postdepositional Diagenesis

Prior studies (Handford, 1982; Lowenstein and Hardie, 1985; Shearman, 1970) have clearly shown that syndepositional diagenetic modification of halite is both common and extensive. Halite-encrusted surfaces of many marginal marine sabkhas and

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Fig. 1.22. (A) Halitepisoids from Bonaire, Netherlands Antilles are forming in the wave swash zone of an artificial halite crystallizing pond. (B) Loose and cemented pisoids up to I cm in diameter are present in the swash zone of the crystallizingponds. (C) Pisoids have a radial-concentric structure similar to that of many carbonate ooids and pisoids. Photographs by C. R. Handf ord. salinas are altered by both physical and chemical processes, brought about by solar heating and hydrological variables such as frequency of flooding by surface waters, depth to water table, amount of groundwater discharge, composition of the waters, and total salinity of both the surface and groundwaters being evaporated (Hunt et al., 1966). Desiccation, dissolution/reprecipitation, and thermal contraction/expansion are important syndepositionalprocesses that alter primaIy halite at the surface. Dissolution and reprecipitation are chief among processes in layers of halite at and just below the surfaces of modern salt pans, while cementation is dominant at greater but still shallow depths. Alteration of halite in the subsurface environment has received only scant attention. However, new evidence has been found (Casas and Lowenstein, 1989) to suggest that diagenetic alteration continues into the subsurface mainly in the form of halite cementation. In comparison to other sedimentary rocks, cementation is

Data Center ,09126599985,[email protected], For Educational Uses SYNDEPOSITIONAL AND POSTDEPOSITIONAL DIACENESIS 25 completed at very shallow depths. Dissolution and karst f ormation. Halite precipitation is unlikely to proceed continuously in any natural setting; it is more likely that, given the high solubility of halite, periods of evaporation and precipitation from saturated brines are interrupted by perhaps brief but significant episodes of brine dilution and halite dissolution. The intensity and frequency of dissolution would be largely determined in subaqueous settings by brine depth, input of dilute waters, and the permanency of the brine body. Both modern and ancient deposits of halite show abundant evidence of periodic dissolution events interrupting deposition. The presence of dissolution surfaces and cavities (Figs. 1.23, 1.24), which are either empty or later filled with sediment and/or halite cement prior to deposition of the overlying halite layers, are testimony to the contemporaneity of halite deposition and dissolution in salt pans. Lowenstein and Hardie (1 985) stressed that the unequivocal signature of salt pan deposition is scribed in halite layers by dissolution events that occur during subaerial exposure and periodic flooding of halite environments with fresh to brackish meteoric water. The record of dissolution events in this case is frequently repeated throughout the stratigraphic section. Dissolution, however, is not likely to occur frequently in evaporite basins perennially filled with deep brines because the undersaturated waters, having a lower specific gravity, would float or flow over the top of the brine body and, thus, could not descend to the bottom to dissolve the halite.

Fig. 1.23. (A) Photomicrograph of dissolution-createdplanarsurf ace of truncationf rom San Andres Formation, Texas. Dissolution may have been caused by the influx of seawater, undersaturated with respect to halite, or, if the surf ace wassubaerially exposed, to dissolution by rainfall or surface runoff. Anhydrite above the dissolution surface is pseudomorphous after small bottom-nucleated crystals of gypsum, which attest to flooding of the surf ace and resumption of evaporiteprecipitation after the chevron halite was truncated by dissolution. (B) Displacive halite crystals have been partially truncated by dissolution prior to deposition of the overlying siliciclastic sediment. Glorieta Formation (Permian), Texas. Photographs by S. Hovorka.

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Dissolution may be visualized in basins perennially filled with shallow, unstratified brines. In this case, shallow-brine bodies are liable to be diluted periodically by, for example, sudden downpours that lead to the drainage of floodwaters from a landmass into a shallow evaporite basin, or the increased inflow of normal marine water (undersaturated with respect to halite) across or through a physical barrier so that evaporation can no longer maintain high levels of halite saturation. Repeated dilution of shallow brine bodies and concomitant dissolution may result from short-term climatic changes (increased humidity, change in wind direction and efficiency of evaporation) or the destruction of a sill or barrier separating the evaporitic basin from normal marine water by, for example, a slight rise in sea level, or increased wave and current erosion. In any event, if halite is either subaerially exposed or flooded by undersaturated waters, dissolution will take place. When undersaturated waters first come in contact with the top of the halite layer, the upper portion is stripped off by dissolution and is marked by a smooth dissolution surface (Fig. 1.23). As a result, the tops of vertically oriented chevron and cornet crystals are truncated (Arthurton, 1973; Lowenstein and Hardie, 1985; Shearman, 1970, 1978). For both modern and ancient layered halite deposits, dissolution surfaces may occur throughout and at intervals ranging from approximately 0.5 to several centimeters. The dissolution surface is marked below by truncated primary halite crystals and above by scattered or concentrated primary gypsum

Fig. 1.24. (A) Slab of modern chevron halite from Salina Omotepec, Mexico is well cemented but also riddled with dissolution voids that appear as vertically oriented or pipeshaped f eatures in (B)photomicrograph. As suggested by Shearman (1970), the voids were probably f ormed by dissolution along halitegrowth planes by descending undersaturated waters. These waters may eventually reach supersaturation with respect to halite and precipitate clear halite cement in the voids. In fact, this thin section consists of patches of clear halite that appear to truncate primary chevron halite and may represent a cement that precipitated in earlier-f ormed dissolution voids. Photographs by C. R. Handf ord and S. Hovorka, respectively.

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crystals, a thin clastic-gypsum layer and/or siliciclastic mud resting directly upon the truncated halite. When dissolution surfaces are created, undersaturated waters can percolate downward into halite layers, dilute the existing pore waters and dissolve halite to form cavities. If, at the time of flooding, the water table lies below the affected halite crust, dilute floodwaters seep downward through the vadose zone and dissolve out vertically elongate cavities (Fig. 1.24), such as pipe- or tube-shaped voids and pits. Formation of vertical cavities is probably enhanced by the underlying vertical grain or orientation of competetively grown halite crystals, such as chevrons, that have vertical compromise boundaries. These boundaries then function as templates along which dilute waters percolate and excavate cavities. A case in point is the documentation by Lowenstein and Hardie (1985) of preferential dissolution along original crystal growth bands such that the surviving halite remnants look like spiny fish skeletons. Where thick halite crusts are exposed to surface weathering, especially dissolution by rainwater and flooding, they erode into jagged pinnacles and serrate ridges (Stoertz and Ericksen, 1974). These solution-formed features impede any sort of ground transportation for their sharp and pointed crests are as much as 0.5 m high in Badwater basin of Death Valley (Hunt et al., 1966). Pinnacles result from corrosion by rain, and where windblown rain falls, the pinnacles and ridges are marked by solutionformed rills and sharp points oriented parallel to the direction of the wind (Stoertz and Ericksen, 1974). The most jagged surfaces are best developed in thick halite crusts where flooding is rare and erosion is largely caused by rainwater. Thick halite crusts may also contain vertical solution tubes which drain rainwater (Stoertz and Ericksen, 1974). These tubes form in the lowest parts of small depressions and lead down from the surface to the the water table, which may be just a few cm to several meters below the surface halite crust. In both Death Valley and the salars, or salt pans, of Chile, springs of fresher water discharge in the salt pans and dissolve the overlying halite (Fig. 1.25). This often results in the formation of circular collapse structures or sinkholes filled with brine. Perennial circular pools between 1-10 m in diameter are present in the Chilean salars, and those which are believed to occur along faults are 10 m deep with vertical to overhanging walls of white halite (Stoertz and Ericksen, 1974). Dissolution features formed by subaerial exposure have been described from Permian marine halite (Hovorka, 1987). Karstlike pits, up to 10 cm wide and 2 m long are present in San Andres halite (Fig. 1.26). They are partially to completely filled with geopetal mud and anhydrite, and most of the dissolution pits are shaped like irregular tubes, which are circular in plan view, and may widen or narrow in a downward

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Fig. I.25. Meteoric water spring and associated solution-pitted halite crust at Devils Golf Course, Death Valley, California. Note the floating rafts of halite. Pool is about I m in diameter. Photograph by B. Lock. direction. The strongly vertical orientation of the pits and the depth (several meters) at which they formed and extend below the paleosurface suggest that these features formed as karst in the vadose zone above the brine-water table by downward percolating water.

Fig. 1.26. (A) Vertical pipes (white) in San Andres halite are marked by the truncation of primary halite by irregularly to vertically oriented clear halite that represents either recrystallized primary halite or cavity-f illing halite cement. (B) Microkarst pits in San Andres halite are filled with dark, anhydritic mudstone that filtered down from durk mudstone interbed near top of slab. The mudstone interbed was deposited during or immediately after a period of subaerial exposure and karstif ication. Photographs by S. Hovorka.

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Powers and Hassinger (1985) described synsedimentary dissolution pits in the Salado Formation that could owe their origin to dissolution in a subaqueous setting rather than a subaerial one. These pits average 25 cm across, 25 cm deep, and are U- shaped in cross section. The pits truncate underlying halite layers and are filled with coarsely crystalline halite with primary growth structures (chevron crystals). The authors suggest they formed by dissolution as the Salado shallow salina was freshened by either fresh sea water or meteoric runoff. Later, as the ambient water became hypersaline again and halite precipitation resumed, the pits were filled with primary halite. Pressure ridges and polygons. Desiccation and subaerial exposure of halite crusts can lead to a host of surface deformation features formed by thermal contraction/expansion and expansion by the force of crystallization. Resulting deformation leads to an orderly arrangement of surface features such as polygons, circles, nets, steps, and stripes. Polygonal fractures in evaporitic mud-flat environments are desiccation features where water loss from sediment causes a decrease in volume of the sediment so that it shrinks and is cracked at the surface (Tucker, 1981). Since the desiccation cracks can only penetrate to a depth where an overall water loss occurs, escaping brines commonly leak into the cracks, where they evaporate and precipitate halite. Thus, it is not unusual to see polygonally mudcracked surfaces in subaerial mud flats in which the cracks are filled with halite. Precipitation of halite in mudcracks may continue to the extent of creating ridges or small ramparts which are able to pond water within the polygons (Hunt et al., 1966). As the ponded water evaporates, it will precipitate thin halite crusts. Patterned-ground features also include salt blisters, pressure ridges, salt polygons, and salt saucers (Fig. 1.27). Although those features form in intensely evaporative environments, often in proximity to polygonally cracked mudflats, desiccation alone cannot account for their origin. They, in fact, show physical evidence supporting an origin by expansion due to halite crystallization, and thermal expansion/contraction. Where a halite crust covers the surface of a desiccated salina or a once-flooded sabkha, the brine water-table is usually at or slightly below the surface. If previously established polygonal patterns formed by desiccation of the underlying mud are present, they can continue to develop. In addition entirely new polygons may develop, but not by desiccation. Tucker (1981) stated that it would be impossible to form open cracks in halite purely by desiccation because additional water loss would lead to the precipitation of yet more halite such that there would be a net expansion of the bed, rather than contraction. In hot deserts, such as Death Valley, California, the diurnal temperature ranges are extremely high. Although daily air temperatures vary by 54"C, even greater temp-

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Fig. 1.27. Polygonal pressure ridges and fractures. (A) Slightly buckled halite crust at Bahia Adair, Sonora, Mexico. Scale = 30 cm. (B) Spectacularlydef ormedpressure ridges and polygonal salt saucers at Devil's Golf Course, Death Valley, California. Large tepee ridge in center is less than 0.5 m high. Photographs by B. Lock. erature variations characterize the surface (60-70°C) (Tucker, 1981). These extreme variations in temperatures affect the physical integrity of the surface crust by forcing the halite to expand and contract daily. For every 25°C change in temperature, there is a 0.10% change in length of a 2-dimensional bar of halite (Tucker, 1981). This means that to produce a 2-cm wide crack around a 14-m polygon, a temperature variation of 36°C is required, which is well within the observed values for Death Valley. Thus, according to Tucker's thermal contraction hypothesis, polygons that form in halite crusts are not initiated by desiccation, but rather by both thermal expansion and contraction of the crust as driven by the diurnal/seasonal temperature changes. After a thin bed of halite is precipitated on a subaerially exposed surface, thermal contraction cracks form and penetrate downward through the crust and into the underlying sediments. Brines are drawn upward and through the cracks by capillary action into the halite crust where they deposit vein-filling halite and the characteristic ridge of halite around the margins of each polygon. Continued addition of halite will result in a net expansion of the crust and cause it to buckle into upward thrusted pressure ridges. The ridges may range initially from afew centimeters to, in a later stage of development, several 10s of cm high, and the distance between polygonal fractures can range from a few meters to over 100 m. Where, as in the latter case, giant salt polygons are present, the affected halite crust must be perennial, hard and brittle, and thick enough to transmit stress for long distances (Stoertz and Ericksen, 1974). With sufficient time, as thrusting and the addition of vein halite proceed, the polygons take on extreme characteristics and spectacular development. Hunt et al. (1966) documented huge polygonal slabs with grotesque forms that are locally referred

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to as salt saucers (Fig. 1.27B). They are as much as 10 m in diameter and about 30 cm thick. The edges of the saucers or polygons are upturned as much as 30" and are as high as 0.5 m. In cross section the saucers are arranged like shingles so that they overlap one another as thrust sheets. In fact, Hunt et al. (1966) showed that some are consistently thrust northward, which suggests that the direction of solar exposure has affected their movement. Stalactites of halite often form on the undersides of overhanging saucers or polygon slabs. Those in the Chilean salars are about 1 cm wide and 13 cm long, and they form by reprecipitation of halite which was carried downward in solution by rainwaters migrating through pores and cracks in the broken crust. Polygonal fractures in halite crusts are often filled with veins of salt which are deposited by brines that bleed into the fractures from the groundwater below. These make up the small ridges and ramparts of salt polygons and pond water within the polygons. Many of the veins are banded parallel to the sides of the fracture walls (Stoertz and Ericksen, 1974), an observation which suggests halite was precipitated episodically to form multigenerationsof vein-filling cement. Halite crusts in the Chilean salars are cut by numerous, widespread veins up to 5 cm wide. Vertical to inclined fractures in shales, filled with red halite, were documented from Permian evaporites in Kansas, Texas, and New Mexico (Schaller and Henderson, 1932; Dellwig, 1962). Schaller and Henderson (1932) briefly noted that they are especially associated with clay, and that the halite has a fibrous character. Dellwig (1962) described polygonal fractures in shale beds which are filled with banded red halite whose acicular crystals are oriented normal to the vein walls (Fig. 1.28). He further reported that the insoluble residue in the halite consisted of 23% ferric iron and contained bacteria and related organic forms. The precipitation of iron was attributed to bacterial action and the polygonal veins were interpreted to represent mud cracks. A section of roof in a Kansas salt mine contained a well-developed polygonal pattern with individual polygons ranging between 1 to 3 m in diameter. The polygons are outlined by borders of anhydrite and veins of red halite are present within the polygons (Fig. 1.28). In a later paper (1968), Dellwig attributed the formation of these polygons to thermal contraction and expansion. Red to orange, fibrous halite also occurs in "late-stage" fractures in Permian evaporites of Texas. Fractures of various orientations cut through mudstones and chaotic halite-mudstone and and show no apparent relationship to polygonal-type patterns of fractures. Their origins have not been thoroughly addressed, but the fractures may have more to do with rock volume reductions coincident with mineralogical changes (i.e., gypsum to anhydrite) than to any tectonic deformation. Polygonal patterns in Triassic halite from England were described and interpreted

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Fig. I. 28. Kansas mine roof exposespolygonalfractures in Permian halite that are filled with fibrous orange halite. Tape scale on left is 30.5 cm long. Photograph by L. Dellwig. by Tucker (1981). Polygons range from 6 to 14 m across and the veins or V-shaped fissures are as much as 6 m deep. The fissures cut through banded, pure halite and the fissures themselves are filled with colorless to pink halite, anhydrite, detrital quartz, feldspar, and clay. The vein-fills are banded parallel to the sides of the fissures and consist of lcm thick units of halite and clastic-rich halite. Tucker (1981) called upon a thermal contraction mechanism, which was previously described above, to account for the formation of the polygons. Cementation. Like continental salt pans, many marginal marine salinas are subject to alternating cycles of flooding, evaporative concentration, and desiccation (Lowenstein and Hardie, 1985). Halite is often reprecipitated as a cement in both primary and previously formed dissolution pores within surface layers of halite as the depositional basin passes through the desiccation stage. Pore waters reach halite saturation during evaporative concentration but, as Shearman (1 970) demonstrated, precipitation in the pores and cavities at the desiccation stage is dominated by crystallization of clear halite cement (Figs. 1.29 A,B). This cement can partially to completely fill dissolution cavities. Its clear, sparry nature contrasts with the cloudy, inclusion-rich halite precipitated earlier as primary bottom precipitates or overgrowths. Halite, which grows first as cumulates, later develops syntaxial overgrowths, and then is subjected to dissolution and cementation. Lowenstein and Hardie (1985) stated that it is these layers of older buried halite, modified by numerous cycles of dissolution and recementation so that only patchy remnants of chevrons and cornets survive, that characterize "mature" salt-pan halite and resemble ancient halite deposits. Originally porous (commonly >50%) at the surface, owing to the numerous dissolution events the layers are subjected to, halite loses most of its porosity when buried. In fact, once buried below a few meters, halite is no longer susceptible to

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Fig. 1.29. (A) Clear halite cement is prevalent in Quaternary halite from Bristol Dry Lake, California. Note that the chevron crystal is embayed and encased by clear, secondary halite cement. This is a good example of Lowenstein's and Hardie's (1985) "mature"halite, which has undergone numerous cycles of dissolution and recementation so that only patchy remnants of primary chevron halite still survive. Despite the intense diagenesis, this halite rock is Quaternary and lies only I m below the surf ace of Bristol Dry Lake! (B) Chevron in center of a halite crystal is rimmed by clear halite, possibly a cement similar to that in A. Alternatively, the clear halite may be a neomorphicproduct from the earlyrecrystallization of chevron halite. Tansill Formation (Permian),Texas. Photographs by S. Hovorka. dissolution from floodwaters, but the remaining cavities will be cemented by clear halite (Casas and Lowenstein, 1989). Quaternary halite layers at about 10 m have porosities of

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30 -- MODERN SALT-PAN HALITE DEPTH MAN-MADE BRINE PONDS (GARRETT, 1970) 0 SALINE VALLEY, CA

A QARHAN SALT PAN, CHINA

50 SEARLES LAKE, CA

0 10 20 30 40 50 I POROSITY (%)

Fig. 1.30. Plot of porosity versus burial depth (based on estimated visibleporosity in about 100 thin sections). Modified from Casas and Lowenstein (1989). halite-saturated brines (Lerman, 1970; Raup, 1970). Although evaporative concentration of groundwater brines has been shown from modern salt pans (Valyashko, 1972) and demonstrated in laboratory experiments (Hsu and Siegenthaler, 1969; Casas and Lowenstein, 1989) question if this mechanism can produce a groundwater brine, supersaturated with respect to halite, which will precipitate halite in the shallow subsurface. In their studies of salt pans, Casas and Lowenstein (1 989) reported that surface brines are heated by solar radiation to 70°C in the summer and that groundwater brines range from 10-25°C in the winter. The authors suggest that a drop in brine temperature from 70°C to 20°C in the system NaCI-H,O could, when repeated daily or seasonally, lead to the precipitation of large volumes of diagenetic halite in the shallow subsurface.

MARGINAL-MARINE HALTTE DEPOSITIONAL SETTINGS - SALINAS: MODERN AND ANCIENT

Marginal marine halite is deposited in subaerial and subaqueous hypersaline environments that lie in (1) peritidal settings or in (2) settings never affected by the tides or storm-washovers but nonetheless flooded by marine waters that are periodically or continuously seeping through a physical barrier separating the evaporite basin from the

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sea (Fig. 1.31). Meteoric waters may enter these environments, too, as rainfall, surface runoff, or as groundwater. With regard to subaqueous environments, only shallow standing bodies of water, which will be referred to here as salinas, are considered, and subaerial environments are grouped under the general heading of sabkha. Following are chief characteristics of marginal marine sabkha and salina environments and the halite that forms in them. Marginal marine salinas are evaporite systems that are partially to completely isolated from the open Ocean and they are covered by water for brief to long periods of time. If covered for long periods, the term perennial salina is preferred, but those which undergo alternating periods of flooding, evaporative concentration, and desiccation (Lowenstein and Hardie, 1985) will be called ephemeral salinas. In both cases, however, the bulk of evaporite deposition is by precipitation from a standing body of water. These environments probably precipitate and preserve far larger volumes of halite than sabkhas (Tables 1.1 and 1.2). Although perennial salinas may undergo water-level fluctuations and thus shoreline transgressions and regressions, their shorelines do not generally migrate as wildly as ephemeral salinas. An ephemeral salina may be totally filled with water after a flooding event only to be totally desiccated a few months later (Fig. 1.32). Both ephemeral and perennial salinas are widely variable in areal extent and water depth; they range from less than a few hundred m2 to thousands of km2 and water depths can vary from 0 to perhaps several meters. Siliciclastic sediments, derived from adjacent marine or continental sources, are deposited in salinas by fluvial, eolian and storm-generated marine processes, and carbonates are precipitated from salina waters or transported from a marine environment into salinas by storm washover.

A PER IT1DAL

6 TIDAL RANGE SALINA SABKHA -* cw SEEPAGE \--’

Fig. I.31. Marginal marine evaporites are deposited in (A)peritidal and (B) other settings at or below sea level but not directly affected by the tides.

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Table 1. I. Modern marginal marine salinas with halite.*b

NAME/LOCATION REFERENCE

Afar region, Africa (Danakil Tiercelin and Faure, 1978; Depression and Lake Asal) Valette, 1975; Hutchinson and Engels, 1970

Sabkha el Melah. Tunisia Busson and Perthuisot, 1977

Lake Macktd, Western Australia, Lxgan, 1987

Hutt and Lecrnan Lagoons, Arakel, 1979,1980 Western Australia

Artificial salt ponds, Bonaire, Handford, 19W7,lO Netherland5 Antilles

Artificial salt ponds, Great McCaffrey et al., 19x7 Island, Bahamas

Ras Muhammad, Sinai Peninsula Friedman, 19x0

Carmen Island, Baja California, Kirkland et al.. 1066 Mexico

Ojo dc Liebre, Baja California. Phleger, 1W19 Mexico

Layna Mormona, Baja California, Vonder Haar and Gordine, 1977 Mexico

Salina Grande, Sonora, Mexico Lock et al.. 19x9

Bocana de Virrila, Peru Moms and Dickey, 1957; Brantley et al., 1W

a Not intended to he a cornpletc list. Includes ephemeral and permanent accumulations of halite

Marginal marine salinas that occur above sea level are present in supratidal coastal sabkhas, distal margins of fan deltas and ephemeral stream deltas, and interdune depressions (Fig. 1.33). High spring tides, storm tides, and/or wind tides force marine waters into depressions that lie on top of the elevated surfaces. Meteoric water flows into the depressions as surface runoff and spring discharge from adjacent higher ground. Marine water adds new soluble salts from the sea and, like meteoric water, may dissolve

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Table 1.2. Modern marginal marine sabkhas and tidal flats with halite."b

NAME/LOCATION REFERENCE

Abu Dhabi, and nearby sabkhas Kinsman, 1969; in United Arab Emirates Warren and Kendall. 198.5

Mediterranean coast of Egypt West et al., 1979

Ranns of Kutch, India' Glennie and Evans, 1976

Laguna Madre, Texas, U.S.A.' Miller, 1975

Sabkha Faishakh, Qatar flling et al., 1965

Northwest Gulf of California, Castens-Seidell, 1084, Mexico' Shearman, 1970; Thompson, 1968

* Not intended to bc a complete list. No permanently preserved beds, only ephemeral crusts. ' Subaqueous prccipitation in a supratidal sctting. previously deposited evaporites and reprecipitate them once evaporation is well under way. Marginal marine salinas present at or below sea level occur in tectonic depressions or basins, and behind coastal barriers, such as beaches or storm rubble ridges (Fig. 1.34). Though perhaps not as large as the giant salinas that existed in the past (Fig. 1.35), some modern coastal areas are flooded extensively each year. The Ranns of Kutch, India, for example, is largely a supratidal area, but approximately 30,000 km' of it are subject to annual flooding (Glennie and Evans, 1976), and the modern salina of Lake MacLeod, Western Australia covers an area of approximately 2000 km2 (Logan, 1987).

Salina Subenvironments

Marginal marine salinas are similar to continental saline lakes. Both consist of subaqueous and subaerial environments within which evaporites form, and the depositional processes and products of both are similar. Therefore, Hardie's et a]. (1978) concepts and terminology for saline lake subenvironments, depositional processes and sedimentary products are employed here. A marginal marine salina is likely to occur in a coastal plain with either high or low topographic relief. Where relief is high, alluvial fans may flank the landward sides,

Data Center ,09126599985,[email protected], For Educational Uses 38 MARGINAL MARINE HALITE

Fig. 1.32. Aerial view of flooded Salina Omotepec on the supratidal mud flats in the northwest Gulf of California (Handf ord, 1988). Fresh to brackish water covered most of the sabkha and salina for several months and dissolved evaporite deposits. Later, this flooded area desiccated and left behind a salt-encrusted pan. Photograph by C.R. H andf ord.

and if relief is low, ephemeral streams or wadis commonly empty into a salina. Other flanking subenvironments around landward margins include eolian sand flats or sand seas, springs, and shoreline features. Topographic barriers separating the salina from the sea can be one of several features, including a barrier beach ridge or spit, coastal

EPHEMERAL C STREAM DELTA D COASTAL INTERDUNES

Fig. 1.33. Marginal marine salinas that form above sea level are present (A) on top of coastal sabkhas, (B) at the distal margins off an deltas and (C) ephemeral stream deltas, and (D) in interdune depressions.

Data Center ,09126599985,[email protected], For Educational Uses SALINA SUBENVIRONMENTS 39

eolian dunes, a reef or reef debris thrown up by major storms, lava flows, and perhaps a tectonic barrier such as a horst block (Figs. 1.34 and 1.36). A salina is usually marked by the presence, too, of a landward dry mud flat that passes into a saline mud flat and eventually into a standing body of water, which may be present for short to long periods of time. Perennially flooded areas that are small may be called pools and ponds, and larger bodies have been called lakes and lagoons. Those that are ephemeral and tend to desiccate are called pans. Dry mud f fats. These are subaerially exposed plains that are generally made up of fine-grained sediments with abundant mudcracks and thin ephemeral salt crusts (Fig. 1.37). Patches of windblown sand and mud-chips (intraclasts) may be present across the surface, too. Sediments are deposited by (1) ephemeral stream or wadi floodwaters that occasionally course their way across the mud flats after substantial rainfalls, and (2) by lake waters that periodically encroach onto the mud flats. A rise in lake level may be due to an increase of freshwater runoff or an increase of seawater inflow across the barrier separating the salina from the sea. Suspension deposition of silt and clay lead to the formation of laminae, but they are commonly disrupted by mud cracks, burrows and evaporite crystals that have precipitated within the sediment (see section on Intrasediment Precipitates). Although halite is not likely to be preserved in these sediments, the presence of hopper-shaped

A REEF RUBBLE

B BEACH-DUNE RIDGES

C HORST BLOCK

Fig. 1.34. Salinas that occur below sea level may be separated from the sea by (A) a reef or reef debris, (B) coastal beach and dune ridges, (C) or by a tectonic barrier.

Data Center ,09126599985,[email protected], For Educational Uses 40 MARGINAL MARINE HALITE

KEG-PRESQUILE BARRIER REEF

BRlTlSH COLUMBl .~ if. ,, ,. ,. h

MIDLAND BASIN

100 km

Fig. 1.35. Lithofacies and paleogeography of several saline giants. (A) The Texas Panhandle during early Permian time (Leonard);lower Clear Fork Formation. Modified from Handford (1981). (B) Elk Point Salt Basin in Devonian time. Modifiedfrom Fuller and Porter (I969). (C)Amazon Basin of South America in Permian time. Modified f rom Szatmari et al. (I979). cubic molds up to several cm wide records precipitation in brine-soaked sediment after one or more cycles of flooding and desiccation. Saline mud f lats. These are transitional subenvironments between dry mudflats and the subaqueous salt pan or lake. They are better developed where lakes expand and shrink, and thus can record both subaqueous deposition of muds during the expansion phase followed by desiccation (lake shrinkage) and widespread intrasediment precipitation of halite crystals from groundwater brines. Although the term saline mud

Data Center ,09126599985,[email protected], For Educational Uses SALINA SUBENVIRONMENTS 41

Fig. 1.36. Depositional topographic barriers .separating a salina f rom the ocean include (A) beach and dune ridges associated with a barrier island or accretingspit, (B) backshore dune fields, (C) carbonate reef and reef debris, and (D) volcanic flows. flat is used here, sabkha is an equally valid term, too, because the record of evaporite deposition is largely beneath the surface of a subaerially exposed surface. Primary laminae are produced by suspension deposition of mud from floodwaters, but they are subsequently destroyed by intrdsediment growth of halite. This commonly results in a chaotic mixture of halite crystals floating in a mud matrix (Figs. 1.18-1.21). Ephemeral salt pans or perennial saline lakes. Subaqueous subenvironments or shallow bodies of water, whose salinities may be perennially high or periodically low and high (dependent on frequency of flooding and desiccation) are called ephemeral salt pans and perennial saline lakes (Fig. 1.38). For purposes of this discussion, water depths are usually less than 5 m. Water in these subenvironments may come from marine and/or continental sources (Fig. 1.1). Meteoric water arrives as runoff, direct rainfall, and meteoric groundwater discharge. Marine water is derived from (1) surface inflow across a barrier, (2) seepage horizontally through a barrier and discharging along the shoreline at sea level, or (3) in the case of a below-sea-level salina recharged by potentiometric-driven groundwater flow, seepage occurs through a barrier and seawater discharges from springs along the shore or on the lake bottom. Evaporite deposition in the lakes and salt pans is through evaporative concentration of the water to a brine, followed by precipitation of crystals at the brine surface and on the floor. The salt pan cycle of flooding, evaporative concentration, and desiccation

Data Center ,09126599985,[email protected], For Educational Uses 42 MARGINAL MARINE HALITE

Fig. 1.37. View across (A) supratidal flats with ephemeral halite crust at Salina Grande, Sonora, Mexico. Note the eolian dunes in background. Photograph by B. Lock. (B) Mudcracks on supratidal mudflats of northwestern Gulf of California. Knife = 9 cm long. Photograph by C. R. Handf ord.

has been thoroughly discussed by Lowenstein and Hardie (1985). Perhaps the greatest insight geologists have gained in understanding the sedimentology of evaporites has come from the study of modern evaporite settings. A study of a modern evaporite environment gives us first-hand opportunity to (1) view the day-to-day processes affecting or controlling evaporite sedimentation and (2) examine the precipitational crystal products by evaporation over a range of evaporite subenvironments. In some cases, cores and trenches have laid open the stratigraphy for viewing and have allowed us to determine the depositional and diagenetic history of the evaporites. Although these studies are rare, they have been used with great success to infer the depositional and diagenetic origins of ancient evaporites. This section will examine various modern salinas (Table 1.1) and inferred ancient salina halite deposits and their vertical successions.

Fig. 1.38. Lake MacLeod, Western Australia from (A) an aerial view shows a dry, evaporite-encrusted surf ace of the ephemeral pan and (B) the ground shows a brine sheet just a ,few cm thick covering the pan surf ace. Photographs by C.R. Handf ord.

Data Center ,09126599985,[email protected], For Educational Uses DANAKIL DEPRESSION - AFAR REGION, AFRICA 43

The use of vertical successions to infer the lateral facies distribution is an important and useful concept (Walther's Law) in sedimentology. It has been successfully applied to many carbonate and siliciclastic strata deposited in all major depositional environments. However, Kendall(l988) has recently shown that the genetic relationship implied by Walther's Law between vertical and lateral facies relationships is commonly not applicable to evaporite sequences. Dogmatic and careless utilization of Walther's Law can result in false and misleading interpretations of the stratigraphic relationships between different lithofacies deposited in evaporite basins. It has previously been pointed out (Middleton, 1973) that Walther's Law is valid only in situations where there are no unconformities separating strata and where the change in facies over a given location is a simple consequence of facies migration. Kendall (1988) explained that evaporite systems are delicate balances between rates of inflow, outflow, and evaporation and that slight changes in external conditions (sea-level, climate, integrity of a barrier, etc.) can upset the balance and result in major changes in the depositional environment. Thus, environments present at one stage in a basin's history may be destroyed and replaced by another that did not exist prior to the changes brought on by external factors. For example, Kendall (1988) cited a case in which redbeds deposited in mud flats were overlain by bedded halite of an ephemeral salt pan. An application of Walter's Law to this succession would likely result in an interpretation of mud flats flanking a salina. However, careful stratigraphic analysis showed that each lithofacies apparently occupied almost the entire basin during their deposition. That is, redbeds were deposited in mud flats that occupied most of the basin and this event was followed by the formation of a halite salina that supplanted the mud flats and occupied most of the basin. As in the example cited by Kendall, most evaporites occur in stratigraphic cycles and each cycle component represents a stage in the evaporative concentration of seawater (Krumbein and Sloss, 1963). Thus, although Walter's Law may not apply in every case, the sequences at the very least record the history of a basin in any one location, and we still profit by examiningvertical sequences of evaporite cycles deposited in salinas.

Danakii Depression of the Afar Region, Africa

When a continent separates along pre-existing lines of weakness or over hot spots by extensional faulting to form major boundary faults during the initial stages of rifting, linear basins made up of grabens, half-grabens and tilted fault blocks form rift valleys over uplifted arches (Veevers, 1977). During these early stages, the graben floors are

Data Center ,09126599985,[email protected], For Educational Uses 44 MARGINAL MARINE HALITE filled with nonmarine clastics, volcanics, and nonmarine or marine evaporites. High rates of subsidence occur such that sediments can accumulate up to 5 km thick in the grabens. Tiercelin and Faure (1978) reported depositional rates of 10-35 m/1000yrsfor halite in the Ethiopian Rift and Danakil Depression of the Afar region, Africa. Less rapidly subsiding basins typically form outside rifted arches and between rifted arches. The Danakil Depression of the Afar Region of Africa (Fig.1.39) is an active rift basin filled with Pleistocene and younger evaporites. The depression is a salt plain, 40 km by 10 km, which lies 125 m below sea level between the Ethiopian Plateau to the west and the Danakil Alps to the east. The Danakil Alps separate the Danakil Depression from the southern Red Sea. A drillhole near the center of the depression encountered halite to its TD of 975 m (Holwerda and Hutchinson, 1968), and seismic reflection profiling indicates that halite is up to 2.2 km thick (Behle et al., 1975). According to Hutchinson and Engels (1970), rifting in the Afar region apparently began during Miocene time and resulted initially in the deposition of more than 3600 m of Miocene halite in a Red Sea basin bounded to the west by the Ethiopian Escarpment. Later episodes of faulting and asymmetrical subsidence resulted in the formation of the Danakil Alps and Danakil Depression (Fig. 1.39). Seawater flowed over, or through, a barrier provided by the emerging Danakil Alps, flooded the depression and subsequently precipitated extensive halite beds. This would help explain the westward facies changesfrom red beds and intercalated basalt flows in the east, near the Danakil Alps barrier, through succeeding gypsum beds and basalt, and into thick halite beds and bittern evaporites to the west (Hutchinson and Engels, 1970). However, since the depression floor is periodically flooded by water draining the western escarpment through alluvial fans, and later totally desiccated, there may also be some input of salts from the west. Evaporite deposition is currently taking place south of the Danakil Depression in Lake Asal (Fig. 1.39). This shallow to relatively deep (average depth = 4 m, maximum depth = 40 m), hypersaline lake lies about 155 m below sea level (Langguth and Pouchan, 1975) where it occupies a large graben system. It is separated by a topographic barrier from Ghoubet el Kharab ("Bottomless Bay"), a restricted bay or lagoon at the western end of the Gulf of Tadjoura. Marine water seeps across the barrier and discharges into the lake through numerous hot and warm springs located along its eastern and southeastern shores. High rates of evaporation have led to the precipitation of gypsum and halite in the saline flats that flank the lake's western shore (Valette, 1975). In addition, shallow gypsum banks have apparently formed caps to the horst blocks that rise above the more deeply flooded grabens of Lake Asal (Langguth and Pouchan, 1975; Valette, 1975).

Data Center ,09126599985,[email protected], For Educational Uses DANAKIL DEPRESSION - AFAR REGION, AFRICA 45

PRECAMBRIAN MESOZO,C VOLCANIC5 EVAPORITES BASEMENT INTRUSIONS 100 KM

Fig. 1.39. Danakil Depression of the Afar region is present in grabens associated with rifting of the east African continent from Arabia. Modified from Hutchinson and Engels (1970).

Although shallow brine pools cover the floor of the Danakil Depression today, the depression may have been filled with deep-water in the recent past. Tazieff (1970) hypothesized that the northern half of the Afar Region, including the Danakil Depression, was covered by sea water as recently as 10,OOO years ago. Citing evidence such as stone axe artifacts (200,000 yrs old) encrusted with "sea shells", the presence of Quaternary coral reefs in subaerially exposed lava fields, flat-topped volcanoes, or guyots, and volcanic deposits that form only in submarine environments, Tazieff suggested that the location of these features indicate that the depth of flooding could have been nearly 500 m. He further suggested that the region was later uplifted and

Data Center ,09126599985,[email protected], For Educational Uses 46 MARGINAL MARINE HALITE drained of seawater. If Tdzieffsflooding hypothesis is correct, those areas that were still below sea level after uplift could not have been drained like the areas lifted above sea level were. However, is it possible that these areas could have desiccated, leaving behind an evaporite deposit that resulted from evaporation of a once deep body of seawater?

Lake MacLeod, Western Australia

The western continental margin of Australia has been a passive margin since Late Carboniferous time when the breakup of Pangea began. According to Veevers and Cotterill (1978), much of Australia’s western continental margin evolved by plate divergence from rift valley systems. Even today, a series of rift valleys and grabens are present there, and it is in one of these, the MacLeod graben, that the spectacular Lake MacLeod evaporite basin (Logan, 1987) has formed (Fig. 1.40). The MacLeod graben,

INDIAN

WESTERN AUSTRALIA

S N BEJALING BARRIER

IBIS GYPSITE 2 0 -2 -4 -6 -8 rn -10 -12 m CYGNET CARBONATE TEXADA HALITE

Fig. 1.40. Lake MacLeod of Western Australia and a north-south crosssection. Modified from Logan (1987).

Data Center ,09126599985,[email protected], For Educational Uses HUTT LAGOON, WESTERN AUSTRALIA 47

which is part of the Carnarvon Basin, was initiated by mild tectonism in late Miocene time. The graben was flanked to the west by the Quobba ridge horst, which formed a barrier between the graben and the Indian Ocean. Pleistocene eolian dunes and cavernous Tertiary limestones comprise the ridge, which is 3-15 km wide, 180 km long, and as much as 90 m high. Logan (1987) has shown that the southern end of the MacLeod graben was connected to the Indian Ocean by a narrow silled passage around 6500 yrs ago. Over the ensuing 1500 yrs, the passage was closed by beach and dune ridges that accreted northwestward by longshore currents and thereafter, the graben received marine water only by seepage across the barriers separating it from the Indian Ocean. In less than 300 years Lake MacLeod was transformed from marine-basin environments of carbonate deposition into an isolated evaporite basin (Figs. 1.38 and 1.40), whose shallow floor laid about 10.5 m below sea level. Marine water seeping across the barriers accumulated in shallow but vast brine ponds that occupied the basin floor. With the limited outflow of brine and the high rates of evaporation, the brines quickly reached halite saturation levels and a 5-6 m thick succession of Holocene Texada halite (Fig. 1.41) was deposited above a 1.5 m thick succession of carbonate, and followed by 2 to 6 m of gypsum. This halite unit was deposited chiefly as bottom precipitates (cornets and chevrons) in the MacLeod salina at rates ranging from 3 to 20 mm/yr from about 5100 to 3600 yrs B.P. Layers ranging from 4 to 30 cm thick are prevalent.

Hutt Lagoon, Western Australia

Evaporite sedimentation was studied by Arakel (1979, 1980) at Hutt Lagoon, Western Australia (Fig. 1.41 and Table 1.1). This lagoon, or marginal-marine salina, is an elongate depression about 70 km2 in area and most of it lies a few meters below sea level. It is separated from the Indian Ocean by a beach barrier ridge and barrier dune system. Similar to Lake MacLeod, Hutt Lagoon is fed by marine waters through the barrier ridge and by meteoric waters through springs. Due to the salina’s below sea-level position and consequently the hydrostatic head of the nearby Indian Ocean, seepage of seawater into the salina is continuous year round. During the winter wet season, the amount of water coming into the salina is substantially increased by the influx of meteoric groundwater. The climate at Hutt Lagoon is Mediterranean; high evaporation rates (2150-2400 mm) are characteristic of the summer and there is moderate rainfall in the winter. These factors combine to form a setting within which sedimentary halite is deposited seasonally and the rates and style of precipitation follow a balance between influx of water and its removal by evaporation and/or reflux.

Data Center ,09126599985,[email protected], For Educational Uses 48 MARGINAL MARINE HALITE

MODERN MARGINAL MARINE SALINA SEQUENCES

BRISTOL DRY LAKE, CALIF. LAKE MACLEOD, W. AUSTRALIA M 0 DRY MUD FLAT GYPSUM SALINA ,815 C."PI,lt

1 SALINE MUD FLAT 5 HALITE SALINA rk "&"A ,,A, ,I/ 2 - SALT PAN SUBTIDAL EMBAYMENT CICNi 1 8

HUTT LAGOON, W. AUSTRALIA NEW LAKE, S. AUSTRALIA M M

EOLIAN GYPSITE -

SALINA

iiltNlll D"Mt5 - EOLIAN DUNES

I LITHOLOGY

Fig. 1.41. Vertical successions in modern salinas. Bristol Dry Lake (nonrnarine) from Handford (1982);Lake MacLeodfrom Logan (1987);Hutt Lagoon from Arakel(1979); New Lake from Warren and Kendall(198.5).

Data Center ,09126599985,[email protected], For Educational Uses PERMIAN BASIN, TEXAS, U.S.A. 49

Two types of halite are present; (1) a thin layer of cumulates overlain by bottom precipitates of chevron and clear halite, and (2) clastic halite made up of broken, abraded, and frosted halite crystals that comprise halite layers. The bottom precipitates form after the winter flooding, when the salina is transformed into an ephemeral lake with waters up to 65 cm deep. These flood waters initially dissolve preexisting halite but after a few days, halite precipitation is renewed and the pond slowly shrinks by desiccation to its summer playa or salt-flat phase. During the summer, when about 95% of the salina surface is a dry salt flat, an area of about 8 x 2 km is underlain by bedded halite up to 50 cm thick. Hot summer winds rework some of the halite into layers of clastic halite particles.

Permian Basin, Texas, U.S. A.

The Permian Basin, which includes the Delaware and Midland Basins of Texas and New Mexico, and companion Palo Duro Basin of the Texas Panhandle (Fig. 1.42) were all relatively deepwater basins flanked by mixed carbonate/clastic platforms with giant salinas (Figure 1.35A) in early Permian time. However, in a span of about 30 m.y., shelf edges prograded across the Palo Duro and Midland Basins and previously deep marine environments, covering approximately 52,000 km', were transformed into platform-wide salinas and desert wadi/eolian plains. Thus, by late Permian time, only the Delaware Basin remained deep, but not for long, because it quickly became evaporitic and filled with 600 m of Castile evaporites in about 200,000 yrs (Anderson et al., 1972; Dean and Anderson, 1982). Evaporites that were deposited in the shallow platform salinas north of the Midland Basin are about 1200 m thick (Presley, 1987) (Fig. 1.42B). They include interbedded evaporites (anhydrite and halite), dolomite, and siliciclastics. Interpreted deep-water evaporites (Castile/Salado anhydrite and halite) are approximately 900 m thick in the Delaware Basin (Adams, 1969). Leonard and Guadalupian-age evaporites of the Palo Duro Basin have been studied almost continuously since 1977 by the Texas Bureau of Economic Geology (University of Texas at Austin). Several long cores and numerous subsurface logs from exploration and production wells make up a data base from which Bureau researchers have constructed numerous regional cross sections, subdivided formations into genetic units, and mapped their distribution, and developed depositional models (Fracasso and Hovorka, 1986; Handford, 1981; Hovorka, 1987; Presley and McGillis, 1982; Presley, 1987). Numerous, regionally correlative cyclic sequences make up the evaporite intervals

Data Center ,09126599985,[email protected], For Educational Uses 50 MARGINAL MARINE HALITE

I*,*DYi/ *EL"

. ~ ~, NORTH SOUTH B

POSTSANANDRES

SAN ANDRES

U CLEAR FORK-GLORIETA

L CLEAR FORK-TUBB Chaotc Mudstm-Hallte 1

WICHITA-RED CAVE Banded Halite } a AnhydriIe [ 500 m

Fig. 1.42. (A) Regional mapof the Permian Basin and adjoining Palo Duro Basin ofwpst Texas and the panhandle region. Modified from King (1948) and Nicholson (1960). (B) North-south cross section of Permian evaporites in the Texas Panhandle. Major transgressive-regressive sequences are identified by formation names and solid dark arrows. Modified from Presley (I987) and Handf ord and Fredericks (I980).

(Figs. 1.42B-1.43)and the vertical succession of lithofacies in these successions changes with the geographic position, and hence paleoenvironment, of the reference point. Sequences that lay nearer to open marine environments typically consist of a thin, dark siliciclastic-rich mudstone at the base that pass upward into inner shelf dolomite with locally present supratidal features, followed by nodular anhydrite, laminated anhydrite

Data Center ,09126599985,[email protected], For Educational Uses PERMIAN BASIN, TEXAS, U.S.A. 51

YI 0 CROSS SECTION MIDDLE SAN ANDRES GENETIC CYCLES '' A

CICLES IN 'CROSS SECTION

Fig. 1.43. Stratigraphic divisions of the San Andres Formation with a reference well and log from the central Palo Duro Basin and cross section A-A' (see Fig. 1.42A for location) of middle San Andres Formation with numerous genetic cycles. See Fig. 1.41 for key to lithology, structures, and grain types. Modified from Fracasso and Hovorka (1986). and massive to rhythmically layered or banded halite (Fig. 1.44). In some cases, siliciclastic mudstone and siltstone may cap the sequence. Most, if not all, of the evaporite strata were deposited in shallow subaqueous environments,with water depths ranging from perhaps a few cm to several meters.

Fig. 1.44. Two core slabs (A-B) of Permian halite showing well developed layers or bands of translucent to cloudy, or f luid-inclusion-rich, halite with vertically oriented chevron crystals. Photographs by C.R. Handford (A) and S. Hovorka (B).

Data Center ,09126599985,[email protected], For Educational Uses 52 MARGINAL MARINE HALITE

Cyclic sequences that lay in more landward positions contain progressively less dolomite and anhydrite. Banded halite is less abundant, too, and is supplanted by a mixture of halite and siliclastic mudstone, referred to as chaotic mudstone/halite (Fig. 1.21). Sequences commonly include banded halite passing upward into chaotic- mudstone/halite and a capping unit of siliciclastic mudstone and siltstone (Fig. 1.45). Deposition of the banded halite was subaqueous, with intervals of subaerial exposure, while the chaotic mudstone/halite and siliciclastic mudstone and siltstone represent saline mud flats to dry mud flats and distal alluvial and eolian plain environments.

MARGINAL-MARINE HAL.ITE DEPOSITIONAL SE'ITINGS-SABKHAS: MODERN AND ANCIENT

Marginal marine sabkhas are wind-deflation surfaces that lie either above or below sea level and are made up of siliciclastic and/or carbonate sediment, partially cemented by evaporites. Though deposited periodically by floodwaters, the matrix sediments are subaerially exposed, either above or below sea level, for much of the year. Evaporites are precipitated (1) at the surface by evaporation of ponded floodwaters and (2) beneath the sediment surface by capillary brines. Surface accumulations are commonly reworked by wind and reworked and/or dissolved by floodwaters. However, some of the dissolved evaporites may later precipitate beneath the surface from evaporatively concentrated floodwaters that seeped downward following a flood event. Their reprecipitation as cements sugests that many sabkha evaporites are diagenetic in origin (Castens-Seidell, 1984). Marginal marine sabkhas (Table 1.2) are present in (Fig. 1.46): (1) narrow to broad, supratidal flats along a coastline, but they may grade almost imperceptibly into inland or continental sabkhas (Kinsman, 1969); (2) coastal interdune depressions, and (3) in depressions that lie below sea level and are separated from the sea by a physical barrier. The following discussions will only address the first two varieties. Depressions that lie below sea level are usually flooded perennially or periodically and, thus, salinas rather than sabkhas are more likely to dominate there. Prograding coastal sabkhas leave behind a shallowing upward record of, from bottom to top, subtidal, intertidal and supratidal facies, but numerous and important variations in vertical sequences do exist and should be expected in the ancient rock record (Fig. 1.47). As shown by Warren and Kendall (1985) and numerous others, gypsum and anhydrite are common evaporite minerals in coastal sabkhas. Halite is rarely preserved, and then only as molds and pseudomorphs. The only likely evaporite to be preserved would be gypsum/anhydrite deposited as displacive nodules or as dia-

Data Center ,09126599985,[email protected], For Educational Uses SABKHAS: MODERN AND ANCIENT 53

ANCIENT MARGINAL MARINE SALINA SEQUENCES

TELEGRAPH SALTS (DEVONIAN) CLEAR FORK FM. (PERMIAN) SAN ANDRES FM. (PERMIAN) CANADA TEXAS TEXAS

IIDAL?

SALAD0 FM. (PERMIAN) FERRY LAKE ANHY. (CRETACEOUS) NEW MEXICO TEXAS 0 - 0 \SA0KHA

SALINE MUD FLAT 1 INTERTIDAL - 2 SALT PAN 5 - 3 GYPSUM SALINA

GYPSUM SALINA 1 - SHALLOW LAGOON 10 5 M

LITHOLOGY

Fig. 1.45. Vertical successions of depositional cycles present in some major evaporite units of probable marginal-marine salina origin. Telegraph salts from Brodylo and Spencer (1987); Clear Fork from Handford (1981); San Andres from Fracasso and Hovorka (I986) but reinterpreted here; Salado Formation from Lowenstein (1982); Loucks and Longman (I982). genetically altered gypsum pan layers (Castens-Seidell, 1984; Hardie, 1986). Modern coastal sabkhas are created by depositional offlap of marine sediments of subtidal, intertidal, and supratidal environments. The shoreline progrades seaward

Data Center ,09126599985,[email protected], For Educational Uses 54 MARGINAL MARINE HALITE over previously deposited subtidal sediments and results in the formation of a shoaling-upward,peritidal sequence in which the supratidal unit is usually no more than 1 m thick (Warren and Kendall, 1985) (Fig. 1.47). Modern coastal-sabkha sediments consist of any mixture of siliciclastic or carbonate muds and sands forming the matrix, and gypsum/anhydrite and halite making up the evaporites. Marine sediments are transported from offshore and deposited on sabkhas by spring tides, storm tides and wind tides while continental sediments are transported from the hinterland and usually deposited in the updip sabkha areas by fluvial processes. Modern coastal sabkhas that receive sediment from marine sources include the carbonate mud flats along the southern coast of the Arabian (Persian) Gulf, such as the Abu Dhabi sabkhas, and the siliciclastic mud flats in the northwest Gulf of California (Castens-Seidell, 1984; Thompson, 1968). Input of siliciclastics from updip fluvially drained sources characterizes the narrow sabkhas present in the distal fringes of fan deltas and wadis of the gulfs of Suez and Elat (Sneh and Friedman, 1984), the most updip portions of the supratidal flats in the northwest Gulf of California (Castens-Seidell, 1984; Thompson, 1968), and fluvial-dominated sabkha of Gladstone Embayment of Shark Bay, Western Australia (Davies, 1970; Handford, 1988). Wind may rework previously deposited marine and continental sediment on sabkhas to form eolian deposits. Other than that which occurs in supratidal pans, very little evaporite deposition takes place at the sabkha surface; most of it occurs below the surface in the

SUPRATIDAL SABKHA INTERDUNE SABKHA

BELOW-SEA-LEVEL SABKHA

Fig. 1.46. Diagram illustrates marginal marine sabkhas form in (A) supratidal flats, (B) coastal interdune depressions, and (C) below sea-level in depressions.

Data Center ,09126599985,[email protected], For Educational Uses SABKHAS: MODERN AND ANCIENT 55

MODERN-ANCIENT MARGINAL MARINE SABKHA SEQUENCES

ARABIAN (PERSIAN) GULF NORTHERN EGYPT NW. GULF OF CALIFORNIA

M M M

SUPRATIDAL SABKHA SUPRATIDAL SABKHA SABKHA -

INTERTIDAL - INTERTIDAL - LAGOON SUBTIDAL

BUCKNER FM. (JURASSIC) RED CAVE FM. (PERMIAN) JOACHIM DOLOMITE (ORDOVICIAN) M ALABAMA TEXAS ARKANSAS 0 0

SUPRATIDAL GlPZUU PAN "OM SUPRATIDAL SABKHA 1 SABKHA - INTERTIDAL 05 LIPS",, DAN - - -INTERTIDAL 2 SUBTIDAL SUBTIDAL RESTRICTED LAGOON - YNTERTIDAL 3 10

I GRAINS

Fig. 1.47. Vertical successions through modern and ancient sabkha sequences. Arabian (Persian) Gulf from Shinn (1983); northern Egypt from West et al. (1979); northwest Gulf of California modified from Castens-Seidell (I984); Buckner Formation from Lowenstein (1987);Red Cave Formation from Handf ord and Fredericks (1980);Joachim Dolomite from author's unpublished work.

Data Center ,09126599985,[email protected], For Educational Uses 56 MARGINAL MARINE HALITE sediment and just above the water table (Warren and Kendall, 1985). Subaerially exposed mudflats in marginal marine sabkhas are often covered with thin crusts of efflorescent halite (Fig. 1.37A). These crusts are precipitated from pore waters drawn to the surface by capillary action from the phreatic zone. Surface evaporation causes the brines to become highly concentrated and, subsequently, to precipitate halite crusts at the surface. At night, the efflorescent crusts may deliquesce in the damper air or dissolve in dew if atmospheric temperatures drop to the temperature at which water vapor condenses (dew point). The crusts almost certainly will be dissolved if rain water falls on the surface or if the surface is flooded by (1) sheetwash from adjacent continental areas or (2) marine waters driven up onto the mud flat by onshore winds and spring tides. Furthermore, as strong winds sweep across the mud flats and deflate the surface, halite is picked up by the wind and carried away as an aerosol or dust. Although efflorescent crusts form quickly and are very common on sabkha surfaces (Figure 1.37A), they are rarely, if ever, preserved in the geologic record. Coastal interdune sabkhas within eolian dune complexes frequently are present along arid coastlines (Fig. 1.48). In some cases, dunes are of local extent, with the sand being supplied solely by the foreshore zone of a beach, such as that occurring in the backbarrier sand dunes of Padre Island (Weise and White, 1980). Otherwise, sand seas may encroach coastal environments from the hinterland. In either case, low-lying coastal interdune depressions may develop into sabkhas. Although marine carbonates may be locally present, interdune sabkhas are more commonly siliciclastic-rich,with the bulk of the sediment left behind as wind-deflation lag. Gypsum and halite are precipitated from brines that seep into the interdune depressions from marine or meteoric sources. Evaporitic interdunes consist of layered to disrupted evaporites and evaporite-cemented siliciclastic sediments. Fryberger et al. (1983) documented poikilitic halite-ce-

Fig. 1.48. Aerial view of coastal dunes and interdune sabkhas in Qatar along the .southern coast of the Persian (Arabian) Gulf. Photograph by P. Scholle.

Data Center ,09126599985,[email protected], For Educational Uses SUPRATIDAL MUDFLATS, GULF OF CALIFORNIA 57

mented sands, polygons (1-2 m across) of halite pressure ridges, or tepees, and layered halite up to 1 m thick incorporating windblown sand in interdune sabkhas. Preservation of primary halite is not likely, however; it would probably be dissolved by fluids moving through the highly porous and permeable dune sands during and after halite deposition. This does not preclude cementation of dune sediments by halite during later subsurface diagenesis.

Supratidal Mudf lats, Northwest Gulf of California

A siliciclastic-rich supratidal sabkha has formed near the mouth of the Colorado River in the northwest Gulf of California (Fig. 1.49). These 20 km-wide supratidal flats, which are flanked to the west by alluvial fans emanating from the mountainous spine of the Baja California peninsula, lie about 3 to 5 m above mean sea level but, nonetheless, are occasionally flooded by runoff from the adjacent alluvial fans and infrequently by abnormally high spring tides (Fig. 1.29A). Floodwaters drain into slight topographic depressions on the sabkha surface to form shallow ephemeral salinas. One of these is known as Salina Omotepec (Fig. 1.32) and it was from this location that Shearman (1 970) documented the saline pan cycle of halite deposition/dissolution and resulting halite crystal fabrics. When the supratidal surface and Salina Omotepec are initially flooded, the previously deposited halite crust covering the surface of the salina is rapidly dissolved by the brackish marine and/or meteoric water. However, in a short time the salina floodwaters are evaporated to a brine, and gypsum, followed by halite, will precipitate subaqueously. As documented by Shearman (1970), small pyramidal hopper crystals nucleate at the surface of the brine, where they float and may coalesce with other crystals to form rafts (Fig. 1.1OA). Subsequently, many of these will sink to the bottom, or drift into shore, and are deposited as halite cumulates above a previously deposited layer of gypsum. The layer of cumulates may reach 5 cm thick. Many of the foundered rafts enlarge into halite cubes by the precipitation of subaqueous overgrowths. Many of the overgrowths are oriented with cube corners and edges facing up so that a chevron arrangement of fluid inclusions is formed. Shearman (1970) described the surface crust of halite and then compared it and some ancient halite deposits to a 25 cm-thick bed of Recent but very lithified halite that was present just beneath the surface of the salina (Fig. 1.24A). The bed consisted halite layers, ranging from 1 to 8 cm in thickness, alternating with thin layers of gypsum sand. Although the bed was rock-hard, it was also vuggy; small vertically-oriented dissolution pipes riddled the rock. In thin section, Shearman (1970) noted that the rock was made

Data Center ,09126599985,[email protected], For Educational Uses 58 MARGINAL MARINE HALITE

up of zoned or inclusion-rich chevron halite and patches of clear halite, which crosscut the chevron growth planes. He recognized that the clear halite patches were similar in size and shape to the dissolution pipes and suggested that the areas of clear halite record former dissolution voids. The arrangement of dissolution cavities, patches of clear halite, and areas of chevron halite within the rock layers testifies to a history of repeated episodes of dissolution and precipitation. Shearman proposed that the vertically-oriented dissolution cavities were created during flooding events by the downward percolation of dilute waters into the halite bed and that the evaporative reconcentration of these waters, which followed, precipitated the diageneticclear halite cement in the brine-filled dissolution cavities. Thus, the halite rock layers, similar to many ancient examples, were the product of a long period of reworking of salt pan halite and groundwater brine diage nesis.

I ALLUVIAL SAL!NA

Fig. 1.49. Northwest Gulf of California supratidal mud flats and a schematic cross section through the alluvial fan and mud-flat systems (modified from Thompson, 1968 and Walker, 1967).

Data Center ,09126599985,[email protected], For Educational Uses CONCLUSIONS 59

The Recent halite rock disappeared from the sediment record in early 1983. Itr disappearance was due to dissolution by brackish waters that flooded much of the supratidal flat when unusually high amounts of rain fell in Baja California (Fig. 1.32) This event substantiates the view that halite is an ephemeral component in sabkhas anc is unlikely to be preserved in ancient sabkha deposits. Although a hard halite-rock layer made up of several halite depositional units (saline pan cycles) was present for man] years on the high flats beneath a protective cover of supratidal muds and additiona saline deposits, the flooding events of 1983 were too great and eventually led to it! dissolution.

Ordovician Joachim Dolomite, Arkansas, U.S. A.

In spite of its rare preservation in sabkhas, the former presence of halite may be indicated by crystal molds and casts, or pseudomorphs. Halite molds are present in supratidal mudflats of the northwest Gulf of California (Castens-Seidell, 1984), and in ancient supratidal rocks, too. Abundant and delicately preserved pseudomorphs of calcite/dolomite after halite (Figs. 1.47and 1.19A-B)were documented from the Middle Ordovician Joachim Dolomite in Arkansas by Handford and Moore (1976). The Ordo- vician varieties show clear evidence of having grown displacively in brine-soaked, laminated dolomitic mudstones whose mudcracks, fenestral fabrics, and rip-up clasts indicate an intertidal to supratidal setting. The halite crystals are believed to have precipitated in low-lying, hypersaline ponds present on supratidal mud-flats. Many of the pseudomorph crystals are asymmetrical and indicate preferred growth in a downward direction at corners and edges. Such a style of crystal growth would imply an upward movement of pore fluids toward the surface, as if evaporative pumping (Hsu and Siegenthaler, 1969) were the driving mechanism. Evaporative pumping would not only explain the downward orientation of the displacively grown halite crystals, but could also account for the early dolomitization of the supratidal mud matrix.

CONCLUSIONS

Precipitation of halite in marginal marine settings leads to the formation of a variety of primary and early postdepositional fabrics in subaqueousand intrasedimentary sites. Distinguishing between these two types of fabrics (Hardie et al., 1985) becomes an important, prerequisite task to interpreting depositional settings. Most of the criteria for distinguishing primary depositional fabrics and early postdepositional diagenetic fabrics were developed from experimental studies and petrographic examination of salt

Data Center ,09126599985,[email protected], For Educational Uses 60 MARGINAL MARINE HALITE pan halite (Arthurton, 1973; Casas and Lowenstein, 1989; Handford, 1990; Lowenstein and Hardie, 1985; Sherman, 1970). However, few criteria exist for determining paleowater depths. For example, bottom precipitates such as cornets and chevrons are known to form in water depths of a cm or less, and there is no reason to suppose that they cannot form in deep-water settings, given a brine solution with sufficient halite concentration levels. Disagreement continues, though, about how deep a body of concentrated brine is possible, or has ever been maintained in a marine basin, to yield a halite deposit. Furthermore, there is the perennial question - "How deep must the brine be before it can be called a deep-water evaporite'?" Despite the relatively few number of researchers in evaporite sedimentology, the field has developed quickly in the past 20 years. Recently, researchers have proposed new ideas and observations about paleowater depths, depositional models, and fluctuating groundwater tables and/or eustasy. Clues about relative energy levels and water depths of halite deposition have recently come from descriptions of the depositional fabrics of both modern (Handford, 1990) and ancient (Lowenstein et al., 1989) examples. Alternative hypotheses are challenging standard depositional models of some classic evaporites as they undergo a fresh round of reexamination (Dietz and Woodhouse, 1988, 1989; Kendall, 1987a). The applicability of "Waltherian" facies models to evaporites has been recently questioned as new viewpoints are issued on the stratigraphic relationships between evaporites and associated strata (Kendall, 1988). Recognition of karst features in halite (Fracasso and Hovorka, 1986) offers hope for inferring paleowater tables and determining eustatic drops in sea level. Evaporites are "the premier recorders of the chemisty of ancient sea waters, lake waters, and other surface waters" (Hardie et al., 1985), and, thus, may hold clues to paleoclimate and geochemical cycles on a global scale. Advances and newly emerging concepts outlined above indicate that halite sedimentology offers promise for future research and that it may help answer fundamental questions about earth history on both the local and global scale.

Data Center ,09126599985,[email protected], For Educational Uses REFERENCES 61

REFERENC‘ES

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Data Center ,09126599985,[email protected], For Educational Uses MARGINAL MARINE HALITE

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Holwerda, J.G. and Hutchinson, R.W., 1968. Potash-hearing evaporites in the Danakil area, Ethiopia. Eon.(;eel., 63: 124-1.v). Hovorka, S., 1987. Depositional environments tf marine-dominated hedded halite, Permian San Andres Formation, Texas. Sedimentology, 34: 1~29-1054. Hsfi, K.J., 1972. Origin ti' saline giants: A critical review after the discovery of the Mediterranean evaporite. Earth Sci. Rev., 8: 371 -396. Hsii, K.J. and Siegenthaler, C., 1969. Preliminary experiments on hydrodynamic movement induced by evaporation and their hearing on the dolomite problem. Sedimentology, 12 11-25. Hunt, C.B., Robinson, T.W., Bnwles, WA. and Washburn, A.L., 1966. Hydrologic Basin, Death Valley, Calfornia. LJ. S. Geol. SUIV.,Prof. Pap., 494-B: 18pp. Hurlhurt, C.S.,Jr. and Klein, C., 1985. Manual d Mineralogy (after J.D. Dana). John Wiley and Sons, New York, 596 pp. Hutchinson, R.W. and Engels, G.G., 1970. Tectonic significance d regional geology and evaporite lithdacies in northeastern Ethiopia. Phil. Trans. Roy. Soc. London, A. Math. and Phys. Sciences, 267: 313-329. nling, L.V., Wells, A J. and Taylor, J.C.M., 1965. Penecontemporary dolomite in the Persian Gulf. In: L.C. Pray and R.C. Murray (Editors), Dolomiti7ation and Limestone Diagenesis. Soc. Eon.Paleon. Mineral., Spec. Publ., 13: 89.111. Jones, C.L., 1965. Petrography d Evaporitesfrom the Wellington Formation near Hutchinson, Kansas. U. S. Geol. Surv. Bull., IZO-A: 70 pp. Karci, I. and Zak, I., 19x7. Bedforms in salt deposits of the Dead Sea brines. J. Sed. Petrol., 57: 723-735. Kaufmann, D.W. and Slawson, C.B., 1950. Ripple mark in rock salt d the Salina formation. J. Geol., 58: 24-29. Kendall, A.C., 1979. Facies models - 13, continental and supratidal (sahkha) evaporites. In: R.G. Walker (Editor), Facies Models. Geosci. Canada, Reprint Ser., 1: 145-157. Kendall, A.C., 1987a. Depositional mtdel for carbonate-evaporite cyclicity: Middle Pennsylvanianof Paradox Basin. Bull. Am. Assoc. Pet. Geol., 71: 576 (abstracts). Kendall, A.C., l9X7h. Early salt dissolution: Pennsylvanian of Paradox Basin, Colorado and Litah. Soc. Eon. Paleon. Mineral., Annu. Midyear Mtg., 4 41 (abstracts). Kendall, A.C., 19%. Aspects ti' evaporite hasin stratigraphy. In: B.C. Schreiber (Editor). Evaporites and Hydrocarhons. New York, C'olumhia LJniv. Press, pp. 11-65. Kendall, A.C., 1989. Brine mixing in the Middle Devonian d western Canada and its possihle significance to regional dolomitii.atitin. Sed. Geol., 64: 271-2235, Kendall, A.C. and G.M. Harwood, 1989. Shallow-water gypsum in Castile Formation - Significanceand implications. Bull. Am. Asstr. Pet. Geol., 73: 371 (abstracts). King, P.B., 1948. Geolog ci' the southern Guadalupe Mountains, Texas. LJ. S. Geol. SUIV.,Prof. Pap., 215 183 pp. Kinsman, DJJ., 1969. Modes of formation, sedimentary associations, and diagnostic features of shallow-water and supratidal evaporites. Btdl. Am. ASROC.Pet. Geol., 53: 830-840. Kinsman, DJJ., 1976. Evaporites: Relative humidity control d primary mineral facies. J. Sed. Petrol., 46 273-279. Kirkland, D.W., Bradbury, J.P. and Dean, W.E., Jr., 1%. Origin d Carmen Island salt deposit Baja California, Mexico. J. Geol., 74 932-93. Krumhein, W.C. and Sloss, L.L., 1963. Stratigraphy and Sedimentation. Second Edition. W. H. Freeman Co., San Francisco, 6ho PP. Languth, H.R. and Pouchan, P., 1975. Caractkres physiques et conditions de stabiliti. du Lac Assal (T.FA A. Riisler (Editors), Afar Depression of Ethiopia. 1. Inter-Union on Geodynamics Scientific Rept. 14, E. Schwcixrhart'sch Verlagsbuchhandlung, Stuttgart, pp. 250-258. Lerman, A,, 1970. Chemical equilibria and evolution of chloride brines. Min. Sw. Am., Spec. Pap., 3: 291-30(>. I,crk, BE., Sinitiere, S.M. and Williams, L.J., 1989. Bahia Adair and vicinity, Stonora: Modern siliciclastic-dt>minatedarid macro-tidal coastline. Bull. Am. ASS(F. Pet. Geol., 73: 381 (abstracts). Logan, B.W., 198'7. The Mack4Evaporite Basin, Western Australia - Holtrene Environments, Sediments and Geological Evolution. Mem. Am. Asstr. Pet. Geol., 44: 140 pp. Loucks, R.G. and Lcmgman, M.W., 1982. Lower Cretaceous Ferry Lake Anhydrite, Fairway Field, East Texas: Prduct of shallow-suhtidal deposition, In: C.R. Handford, R.G. Loucks and G.R. Davies (Fditors), Depositional and Diagenetic Spectra of Evaporites - A Core Workshop, Soc. Eon. Paleon. Mineral., Core Workshop No. 3: 130-173. Lnwenstein, T.K., 1982. Primary features in a potash evaporite deposit, the Permian Salado Formation d' west Texas and New

Data Center ,09126599985,[email protected], For Educational Uses 64 MARGINAL MARINE HALITE

Mexico. In: C.R. Handford, R.G. Imucksand G.R. Davies (Editors), Depositional and Diagenetic Spectra of Evaptnites -A Core Workshop. Soc. &on. Paleon. Mineral., Core Workshop No. 3: 276-304. L,owenstein,T.K., 1987. Evaporitedepositionalfahricsin thedeeplyhuriedJurassic Buckner Formation, Alabama. J. Sul. Petrol.. 57: 108-116. Lowenstein, T.K., 1988. Origin d depositional cycles in a Permian "saline giant": The Salado (McNutt 7one) evaporitcs tfNew Mexico and Texas. Geol. Scr. Am. Bull., 100: 592-608. Lowenstein, T.K. and Hardie, L.A., 1985 Criteria for the recognition d salt-pan evaporites. Sedimentology, 32: 627-644. Lowenstein, T.K., Hardie, LA., Casas, E. and Schreiber, B.C., 1989. Suh-Mediterranean giant salt: A perennial suhaqueous evaporite. Ceol. Stx. Am., Annu. Mtg., 21: AM(Abstracts with Programs). henstein, T.K. and Spencer, RJ., 1993. Syndepositional origin of potash evaporites: Petrographic and fluid inclusion evidence. Am. 1. Science, 290: 142. Lucia, FJ., 1972, Recognition of CVdpOritC-CarhOnate shoreline scdimentatitm In: J.K. Rigby and W.K. Haniblin (Editors), Recogmition of Ancient Sedimentaly Environments. Sw. &on. Paleon. Mineral., Spec.Publ., 16: lho-191. Maiklem, W. R., 1971. Evaporative drawdown - a mechanism for water-level lowering and diagenesis in the Elk Point Basin. Bull. Can. Pet. Getd., 19487-503. McCaffrey, M.A., LaTar, B. and Holland, H.D., 1987. The evaporation path (f =water and the coprecipitation of Br and K' with halite. J. Sed. Petrol., 57: 928-937, Middlelon, G.V., 1973. Johannes Walther's Law cf the correlation of facies. Geol. Soc. Am. Bull., x4: 970-988. Miller, J.A., 1975. Facies characteristics of Laguna Madre wind-tidal flats. In: R.N. Ginsburg (Editor), Tidal Deposit\: A Casehook of Recent Examples and Counterparts. Springer-Verlag, New York, pp. 67-73. Moretto, R., 19111. Etude Sedimentologique et Getxhimique des Depots de la Sene Salifere Paleogene du Rassin de Bourg-en-Bress (France). Mem. Sci. Terre, 50: 252 pp. Moms, R.C. and Dickey, PA., 1957. Modern evaporite deposition in Peru. Bull. Am. Assoc. Pet. Geol., 41: 2467-2474. Nicholson, J.H., 1960. Geology of the Texas Panhandle. In: A~pectsof the Geolcgy d Texas, A Symposium. LJniv. (fTexas, Bur. Eon. Geol., Puhl., 6017: 51-64. Ochsenius, C., 1877. Die Bildung der Steinsablager und ihrer Mutterlaugensal7~unter spe7ieller Beriicksichtigung der Flii~e von Douglashall in der Egeln'schen Mulde. Halle, C.E.M. Pfeffer Verlag, 172 pp. Orti Caho, R., Pueyo Mur, JJ., CXsler-Cussey, D. and Dulau, N., 1W.Evaporitic dimentation in the coastal salinas of Santa Pola (Alicante, Spain). In: F. Orti Cab0 and G. Busson, (Editors), Introduction a la Sedimentol(@e des Salines Mantimes de Santa Pola (Alicante, Espagiia). Rev. Dlnvest. Geol., [Jniv. Barcelona, 38/39 169-220. Phleger, F. B., 1969. A modern evaporite deposit in Mexico. Bull. Am. Asstr. Pet. Geol., 53: 824-829. Powers, D.W. and Hassinger, B.S., 1985. Synsedimentaly dissolution pits in halite ofthe Permian SaladoFormation, southeastern New Mexico. J. Sed. Petrol., 55: 769-773. Presley, M.W., 1987. Evolution d Permian evaporite basin in Texas Panhandle. Bull. Am. Assnc. Pet. Geol., 71: 167-190. Presley, M.W. and McGillis, KA., 19a. Coastal Evaporite and Tidal-flat Sediments of the upper Clear Fork and Glcnieta Formations, Texas Panhandle. 1Jniv. of Texas, Bur. &on. Geol., Rept. Invest., 115: 50 pp. PrAbram, K., 19%. Irradiation Colours and Luminescence. Pergamon Press, Inc., London, 332 pp. Raup, O., 1970. Brine mixing: An additional mechanism for formation of basin evaporites. Bull. Am. Assoc. Pet. Geol., 54: 2245-2259. Ryan, W.B.F. and Hsii, KJ., 1973. Initial reports cf the Deep Sea Drilling Projects, Leg 13. US. Gtw't Printing Office, Washington, D. C., 1447 pp. Schaller, W.T. and Henderson, E.P., 1932. Mineralogy of Drill Cores from the Potash Field of New Mexico and Texas. 17. S. Geol. Surv. Bull., 833 I24 pp. Schmalz, R.F., 1969. Deepwater evaporite deposition - a genetic model. Bull. Am. Assoc. Pet. Geol., 53: 798-823. Schmalz, R.F., 19%). Curiouser and curiouser. Ceol. Soc. Am., Annu. Mtg., 21: A363 (Abstracts with Programs). Schreiber, B. C., 1988. Subaqueousevaporite deposition. In: B. C. Schreiber (Editdr), Evaporitesand Hydrocarbons. New York, Columbia Univ. Press, pp. 198-255. Schreiber, B.C. and Helman, M.L. 1989. What are the problems in forming deepwater Messinian evaporites? Geol. Soc. Am., Annu. Mtg., 21: A363 (Abstracts with Programs). Scruton, P. C., 1953. Deposition of evaporites. Bull. Am. Assoc. Pet. Geol., 372498-2512.

Data Center ,09126599985,[email protected], For Educational Uses REFERENCES 65

Scruton, P. C., 1W.Delta huilding and the delta sequence. In: F. P. Sheprd and T. H.von Andel (Frlitors), Recent Sediments, Northwest Gulf tf Mexico, 1951-1958. Am. Assoc. Pet. Geol., pp. 82-1(12. Shearman, DJ., 1970. Recent halite rock, Fhja California, Mexico. Tpans. Tnst. Min. and Metall., R., 79 B155-162. Shearman, DJ., 1978. Halite in sahkha environments. In: W.E. Dean and R.C. Schreikr (Editors), Marine Evaporites. Soc. Frnn. Paleon. Mineral., Short Course, 4 3042. Shinn, E.A., 1983. Tidal flat envirnnment. In: PA. Schnlle, D.G. Behut and C.H. Moore (Editor), Carhnnate Depositiclnal Environments. Mem. Am. Assnc. Pet. Chd., 33: 171-210. Shlicta, P.J., 1%. Grtnvth, deformation, and defect structure (f salt crystals. In: R.B. Mattox (Fditor), Saline Deposits. C;eol. Soc. Am., Spec. Pap., 88: 597-617. Sneh, A. and Friedman, G.M., 19x4. Spit complexes alnng the eastern coast id' the Gulf nf Sue/. Sed. Genl.. 39 21 1-226. Snnnenfeld, P. 1984. Brines and Evaporites. Academic Press, Toronto, 613 pp. Sorhy. H.C., 18-53. On the microscopical structure d crystals, indicating the origin td' minerals and rocks. Geol. Str. London, Chart. J., 14453-.500. Southgale, P.N.. 19x2. Cambrian skeletal halite crystals and experimental analoyes. Sedimentolngy, 20: 391407. Stewart, F.H., 1949. The petrolngy d the evaporites tf the Eskdale no. 2 horing, east Yorkshire; pt. 1 - The Lower Evaporite Bed. Mineral. Mag., 2% 621-675. Stewart. F.H., 1951a.The petrnlngy of the evaporites d the Eskdalc no. 2 horing. east Ynrkshire; pt. 2 - The Middle Evaporite Bed. Mineral. Mag., 29 445-47.5. Stewart, F.H., 1951h. The petrology tf the evaporites nf the Eskdale no. 2 horing. east Yorkshire; pt. 3 - The LIpper Evaporite Bed. Mineral. Mag., 29 557-572. Stewart, F.H.. 1963. Marine evaptirites. In: M. Fleischer (Editor). Data d Getrhemistry. Sixth Edition. Chapter Y, (1. S. Geol. Surv., Prnf. Pap., 440-Y 52 pp. Stocrli. G.E. and Ericksen, GE., 1974. (;enlogy of Salars in Northern Chile. LJ. S. Geol. Siirv., Prnf. Pap., 811: 65 pp. S~atmari,P., Carvalho, R.S. and Simocs, LA., 1Y79. A comparison d evaporite facies in the Late Palewoic Amamn and the Middle Cretaceous South Atlantic Salt Basins. Fain. Genl.>74 432- 447. Tafieff, H., 1970. The Afar triangle. Sci. Am., 1-222: 12-40, Thompson, R.W., 1%. Tidal Flat Sedimentation \>nthr Colorado River Delta. Northwestern Gulf of California. Geol. Snc. Am. Bull., 107: 133 pp. Tiercelin. J.J. and Fame, H., 1978. Rates d rcdiment;ition and vertical siihsidcnrc in neorifts and paleorilts. In: 1.B. Ramkrg and E.R. Neumann (Fditors). Tectonics and Geophysics nf Continental Rifts. D. Reidel Pub.Co., Dordrwht, Hdland, pp. 41-47. Tucker, R.M., 1981. Giant pnlyg,ms in the Triassic salt (d' Cheshire, England: A thermal cnntraction mdel for their origin. J. Sed. Pctrtd., 51: 779-726 Llsiglio, M.J., 1x49. hides sur la compnsition de l'eao de la Mediterrank et sur l'exploitatinii dcs sel qu'elle conteint. Annu. C7lim. Phy\., 27: 172-191. Valette. J.N., 1975. Gctrhemical study of Lake 4-1 and Ghouhet el Kharah (T.F.A.I.). In: A. Pilger and A. Riisler (Fditors), Afar Depression d Ethiopia. 1. Inter-linion ('ommission on (ietdynamicsScientaic Rpt. No. 14, E. Schwei/erhart'sche Verlagshuchhandlung, Stuttgart, pp. 239-2SlcO. Valyashko, M.G., 1412. Playa lakes - a necessary \(age in the development (f salt hearing hasins. In: G. Richter-Bernkry (Editor), Geolcgy of Saline Dcpsits. LJNESCO - Prtr. Hannvcr Symp., pp. 41-51. Veevers, J.J., lY77. Rifted arch hasins and post-hrcakup rim hasins trn passive ctmtinental margins. Tectonophysics, 41: T1-T5. Veevers, J.J. and Cotterill, D., 1978. Western margin td' Australia: Evolution of a riftcd arch system. Geol. Snc. Am. Bd., 89 337-355. Vonder Haar, S.P. and Gorsline, D.S., 1977. Hyperwline lagmn depnsits and prncesxs in Baja California, Mexico. In: HJ. Walker (Frlitor), Recearch Techniques in Coastal Envirnnments. Geosci. and Man, 1X: 165-177. Walker, T.R., 1967. Formation of red kds in ancient and mdern deserts. Geol. Snc. Am. Bull., 78: 353-W. Walther, J., 1912. Das (;eset/ der Wustenhildung. Sectnd Fxlition. D. Reimer, Berlin. pp. 24-242. Wardlaw, N.C. and Schwerdtncr, W.M., 1966.Halite-anhydrite ccasonal layersin Middle Devonian Prairie EvaporiteFormatim, Saskatchewan. Canada. Geol. Str. Am. Bull., 77: 331-342. Warren, J.K. and Kendall. C'G, St. (',, 1985. Ctimparison nf sequences formed in marine sahkha (suhaerial) and salina

Data Center ,09126599985,[email protected], For Educational Uses 66 MARGINAL MARINE HALITE

(suhaqueous) settings - mtdem and ancient. Bull. Am. Assoc. Pet. Geol., hY 1013-1023. Weiler. Y., Sass, E. and Zak, I., 1Y4. Halite oolites and ripples in the Dead Sea, Israel. Sdimentolnpy, 21: 623432. Weise, B.R. and White, W. A,. 1980. Padre Island Seashore - A Guide to the C~olngy,Natural Environments and History nf' a Texas Banier Island. Univ. nf' Texas, Bur. &on. Geol., Guidehook, 1794 pp. West, I.M.,Ali, YA.and Hilmy, M.E., 1979. Primary gypsum nodules in a modern sabkha on the Mediterranean coast td Egypt. Ged., 7 3.543%.

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Data Center ,09126599985,[email protected], For Educational Uses Chapter 2

SULFATE DOMINATED SEA-MARGINAL AND PLATFORM EVAPORATIVE SE'ITINGS: Sabkhas and salinas, mudflats and salterns

John K. Warren

INTRODUCTION

Economic Significance

Many explorationists consider the study of evaporites as an erudite pastime, yet these same explorationists worry about seal continuity and capacity, timing of trap formation, and source potential. In a basin containing evaporites the answer to some of these problems lies in a study of the evaporites themselves; why they occur where they do, how dolomitization is repeatedly associated with evaporites, and how salt can flow to generate giant traps in subjacent and overlying sequences. Evaporites overlie carbonates that enclose an estimated one-half of the worlds known petroleum reserves (Kirkland and Evans, 1981). About 70% of the world's giant oil fields in carbonate rock.. have a relationship to evaporites (Zhang Yi-gang, 1981) and all of the giant gas fields in thrust belts possess an evaporite seal (Downey, 1984). The association among evaporites, carbonates, and hydrocarbons is more than accidental as evaporites constitute less than 2% of the world's platform sediments (Ronov et al., 1980). The partnership persists for various reasons. Bedded ancient evaporites overlying or updip from porous sediments set up effective seals to underlying potential reservoirs. Evaporite seals typically embody capillary entry pressures in excess of lo00 - 1500 psi causing thin evaporite units to be able to support tall oil columns. The contact between the underlying reservoir and many bedded evaporites is often knife sharp. Abrupt contacts are due to the very different conditions that were required to deposit thick subaqueous evaporite sequences (not sabkhas) compared to the underlying, often marine reservoir rock (e.g. a reef). Because the units are separated in time and location by Walther's Law, the contact is typically sharp. This contrasts to many shale and micrite seals where the seal facies is a lateral time equivalent to the reservoir. This lateral continuity and time-equivalence of non-evaporite environments often leads to an uneconomic transitional zone of silty sand that contains oil and gas and separates the sand reservoir from the shale seal (e.g. floodplain shales sealing barrier island and

Data Center ,09126599985,[email protected], For Educational Uses 70 SEA-MARGINAL AND PLATFORM CALCIUM SULFA’I’l: deltaic reservoirs). Thick, buried halite units are often remobilized into salt domes, which in turn create potential reservoirs in the surrounding sediments (see Halbouty, 1979; Jenyon, 1986; Warren, 1990a for summaries). The subsurface movement of pore waters can leach evaporites in carbonate and siliciclastic matrices and so create areas of secondary porosity. Evaporite diagenesis can release large quantities of magnesium-rich brine to form sucrosic dolomite, an excellent potential reservoir, in adjacent limestones. An association may even exist between evaporite depositional settings and source rock formation (Evans and Kirkland, 1988; Kirkland and Evans, 1981; Warren, 199Oc). Evaporites and evaporitic carbonates are deposited in arid environments lacking fresh water, hence sediments do not contain a high proportion of organic matter derived from higher plant debris. Rather, the elevated salinities and general environmental stress means the organics are often derived from algae and zooplankton. Most modern evaporites are deposited adjacent to, or within, areas of high algal productivity and accumulation. Thus evaporitic source rocks, frequently laminated carbonates, are more likely to be oil producing than gas producing. Most modern evaporites are deposited adjacent to, or within, areas of high organic productivity and accumulation. Similar ancient settings, such as rifts and collision basins, are conducive to the preservation of this organic matter until it reaches the zone of catagenesis (Evans and Kirkland, 1988; Warren, 1986b).

History of Evaporite Research

The first serious study of the depositional origin of evaporites began with the work of the ltalian chemist Usiglio (1849). He was the first to establish the order of crystallization of salts from concentrating seawater, and for the less soluble salts such as gypsum and halite, his obsewations have yet to be improved. The idea that thick evaporites were formed by the evaporation of ponded seawater in a barred-basin located behind a restricting sill was advanced by Bischof (1 854) and further expounded by Ochsenius (1877). This barred-basin model held sway well into the 1960s when studies of modern sabkha (salt flat) sediments convinced many geologists that tidal flats were the key in interpreting many ancient evaporites (Illing et al., 1965; Shearman, 1963, 1966; Butler, 1969). Our improved understanding and modeling of evaporite deposition and diagenesis in the last three decades is based on improved understanding of Quaternary depositional settings. It began in the 1960s with detailed studies of sabkhas in the Arabian (Persian) Gulf and bedded salts in the Dead Sea (Illing et al., 1965; Shearman 1963, 1966; Neev

Data Center ,09126599985,[email protected], For Educational Uses HISTORY OF EVAPORITE RESEARCH 71

and Emery, 1967). Next platform gypsum was described in the Permian of Canada and the Miocene of the Mediterranean (Davies and Nassichuk, 1975). Then texturally equivalent subaqueous gypsum was described from Holocene coastal lakes in West and South Australia (Ardkel, 1980; Logan, 1981; Warren, 1982a,b) and in shallow-water gypsum pseudomorphs in cratonic rocks 3.45 billion years old in Western Australia (Fig. 2.1; Dunlop, 1978; Lowe 1983). In the 1970s and 80s, as more and more environmental complexities were recognized in ancient evaporites, geological models moved away from pure sabkha or salina interpretations into analyses that encompassed a spectrum of depositional settings. One of the first spectral reconstructionsof evaporite depositional settings was that of the Messinian evaporites in the Mediterranean region. In the late sixties and early seventies, the Deep Sea Drilling Program (DSDP) discovered a 2 km thick sequence of Late Miocene shallow-water evaporites beneath the Mediterranean Sea. The unit is sandwiched between deep-water carbonates, and extends beneath large areas of the Mediterranean Sea and surroundings. Textural analyses of DSDP cores indicated a mixture of shallow subaqueous and sabkha evaporites with rare deeper water intervals. This led to the novel idea of a desiccated deep-basin with a basin floor covered by brine only a few meters deep, yet situated thousands of meters below sea level (Hsii et al., 1977 for summary of relevant literature). This innovative hypothesis was not widely accepted by land-based geologists who preferred a deep-water deep-basin origin of the evaporite unit. Subsequent onshore work in the Late Miocene evaporites of the Sicilian and Italian basins confirmed a shallow-water origin (Schreiber and Decima, 1976,1978). Today we realize that there is no all-encompassing depositional model for ancient platform and basinal evaporites, and that models based solely on the scale of Holocene counterparts are limiting and sometimes misleading in the interpretation of ancient

Fig. 2.1. Evaporitepseuaornorpvls Dottom-nucleateagypsurnfrorn tne 3.45 billion year old Warrawoona Group, North Pole, WesternAustralia. Thepseudomorphsarenow composed of barite. (Sample courtesy of John Dunlop).

Data Center ,09126599985,[email protected], For Educational Uses 72 SEA-MARGINAL AND PLATFORM CALCIUM SULFATE evaporite sequences. Contemporary marine evaporites are confined to sabkha and sea-marginal subaqueous lakes with thicknesses measured in tens of meters and areas measured in hundreds to thousands of square kilometers. At various times in the past, evaporites were deposited in depositional settings ranging from continental to platform to basinal with thicknesses in hundreds of meters and aerial extents measured in thousands to millions of square kilometers (Fig. 2.2). This chapter is restricted to platform evaporites which are defined as stratiform units (<50 m thick, often < 5 to 10 m thick) of mixed evaporitic mudflat and saltern evaporites, with rare deeper water intervals which are all deposited on a continental platform and often intercalated with other shoal-water marine-influenced platform sediments. Platform evaporites not deposited as part of a basin-wide setting, passed seaward into time-equivalent open-marine deep-water sediments (Warren, 1989). Platform evaporites are commonplace in the rock record and in the last few years they have been interpreted using textures documented in modern sea-marginal settings. Modern sea-marginal evaporites form in one of two settings, subaerial or subaqueous (Table 2.1). Modern subaerial dominated settings are coastal sabkhas while subaqueous dominated settings are coastal salinas. The two settings are by no means mutually exclusive, modern salinas and sabkhas often grade into one another both laterally and vertically. Large parts of many ancient arid platforms were evaporitic mudflats; depositional mosaics of salinas and sabkhas. Other arid zone platforms were covered by salterns; huge evaporite-depositing seaways dominated by laterally extensive subaqueous evaporite units (Warren, 1989).

CONTl NENTAL PLATFORM SLOPE & RISE BASIN ___~____~ ~ sabkhas, salt lakes ramp, shelf slumps, turbidites laminates

EVAPORlTlC MUDFLAT SALTERN laterally extensive laterally extensive evaporite mosaic of seaway with evaporitic sabkhas and salinas mudflats in shoal areas

Fig. 2.2. Environmental spectrum seen in ancient evaporites of whichplatform sequences are one of the major subsettings. Platforms are,further divided into salterns and evaporitic mudf lats.

Data Center ,09126599985,[email protected], For Educational Uses HISTORY OF EVAPORITE RESEARCH 73

Table 2. I. Sedimentological characteristics of typical sabkha and salinas.

SABKHfi SALINA-SUBAQUEOUS

Sulfate-dominated facies Sulfate-dominated facies

Evaporite units are supratidal, matrix dominated, Evaporite units are subaqueous, relatively pure, usually < 60% sulfate often > 70% sulfate

Each evaporite (supratidal) depositional unit is Each evaporite (subaqueous) depositional unit is thin, typically <1-2 m thick, 1-20 m

Displacive and replacive nodular and enterolithic Bottom-nucleated evaporite crystal textures; often textures laminated, laminae can be laterally continuous but do not extend across the whole basin

Evaporite crystals are diagenetic and grow in a Deposition can be mechanical; evaporites contain matrix of mechanicallydepositedcarbonate or clastic textures; gypsolites, wave and current siliciclastic sediments ripples, cross-beds, rip-up breccia, reverse and normal graded beds

Associated with sabkha tepees, flat-laminated and Associated with groundwater tepees, laminar polygonal algal mats algal mats and domal subaqueous stromatolites

Facies units tend to be laterally extensive, Facies often symmetric or asymmetric bull's-eye parallel to shoreline; deposited as "pentidal" patterns; sometimes as vertical aggrading facies trilogy

Carbonate matrix washed in from lagoon during Carbonate facies outline areas of less saline water storms; sands can be blown in from adjacent sand in basin; vadose pisolites can be common in seas (ergs) such areas

Hydrology dominated by storm recharge and Hydrology dominated by evaporative drawdown shallow brine reflux and deeper brine reflux

Halite-dominated facies Halite-dominated facies

Tend to be matrix-rich units usually less than 2 Units are rclatively pure and thick; composed of m thick. Chaotic halite muds and karstic superimposed halite crusts separated by surfaces are commonplace terrigenous or carbonate-sulfate laminae. Each depositional unit was as thick as the brine pond was deep, 2-10 m

Crystals tend to be displacive or replacive with Bottom-nucleated chevrons, rippled and cross common inclusions of matrix in the halite bedded accumulations, and cumulates; chevrons crystals; "pagoda" or skeletal halites are good in crusts deposited in ponds with fluctuating indicators of mud flat deposition salinities or on flats which may later exposed. Chevron halite oftcn composed of brine- inclusion-rich, upward aligned crystals. In ephemeral water environments these crystals can be crosscut with clear halite filling karst cavities

This chapter was written to aid the exploration and reservoir geologist in interpreting shoal-water and subaerial facies in ancient sulfate-dominated platform settings. Interpretation of platform evaporites almost always includes a study of dolomite

Data Center ,09126599985,[email protected], For Educational Uses 74 SEA-MARGINAL AND PLATFORM CALCIUM SULFATE

and evaporite diagenesis. Included are some of the more relevant aspects of dolomitization at the end of the appropriate subsections and a short summary of pertinent aspects of evaporite diagenesis at the end of the chapter. For more comprehensive summaries of these complex topics the interested reader is referred to Hardie (1987), Holser (1979), Land (1985), Machel and Mountjoy (1986), Schreiber (1988) and Warren (1990ac).

SEA-MARGINAL SABKHA MODEL: Arabian (Persian) Gulf Example

The word "sabkha" (plural is sabkhat) comes from the Arabic word sabaka meaning salt marsh or salt swamp. Geologically it describes a salt flat with an equilibrium geomorphic surface whose level is dictated by the local water table (Shearman, 1966). In a sea-marginal setting the sabkha surface is usually no more than a meter above mean sea level. Sabkhas are occasionally covered by ephemeral shallow water. They were first documented along the carbonate dominated coast of the southern Arabian (Persian) Gulf by Curtis et al. (1963) but were subsequently described on the siliciclastic coast of Baja California, Mexico (Phleger, 1969; Shearman, 1970; Castens-- Seidell, 1984), the coast of Sinai (Gavish, 1974; Levy, 1977), and along the edges of the Mediterranean including the coasts of Libya, Tunisia and Egypt (West et al., 1979). They are also found at the northern ends of Gulf St. Vincent and Spencer Gulf in South Australia and about the edges of Shark Bay in western Australia and portions of the northwestern Australian coastline. If the word sabkha is used objectively to imply "salt flat" then there are both coastal and continental sabkhas. In fact, some coastal sabkhas, such as those along the Trucial Coast, pass laterally into continental and eolian sabkhas without any noticeable change in surface morphology on the sabkha (Kinsman, 1969). The marine portion of the sabkha is characterized by a matrix of marine sediments soaking in largely marine-derived groundwaters, and the continental portion by non-marine sediments and groundwaters. Kendall (1979) uses continental playa as an equivalent to continental sabkha. Other workers use the term sabkha in a more narrow sense to describe marine sabkhas as exemplified by the Arabian Gulf sabkhas. This chapter uses the term sabkha to describe marine and continental salt flats where displacement and replacement evaporites areforming in the capillary zone above a shallow saline water table. This definition has the advantage that it is a workable subsurface definition when information on the horizontal extent of a particular ancient evaporite is often sparse or lacking.

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As more modern salt flats are described in settings as varied as the Andean Altiplano in South America and the interdunal flats around Lake Eyre in Australia, it becomes more difficult to retain narrower usage of the word sabkha exemplified by the Arabian Gulf coastal sabkha (Handford, 1981a). Just as major siliciclastic depositional systems exhibit a spectrum of morphologic types (e.g. river verses tide verses wave-dominated deltas), so there are significant variations among sabkhas. Physical processes thought to be most important (besides evaporation) include those operative under: a) marine-coastal, b) fluvio-lacustrine, and c) eolian-dominated conditions. Dominance of any one of these processes gives rise to sea-marginal sabkhas, continental sabkhas, or interdune sabkhas. An interdune sabkha can be a subfacies of either a marine coastal or a continental sabkha. Sedimentary responses to the dominant physical processes lead to the development of sabkhas consisting of combinations of; a) terrigenous clastics, b) carbonate-sulfate (anhydrite-gypsum minerals), and c) soluble salts (halite, sylvite, polyhalite, etc.) also known as the bittern salts which do not occur in modern sea-marginal settings. The Arabian (Persian) Gulf sabkha should not be used to generate an all encompassing model of evaporite deposition. A sabkha classification must consider the sabkha system ;is an environmental and lithological continuum. The only single criteria to describe evaporite sediments, including sabkhas of the United Arab Emirates, is their variability. Many ancient sabkhas probably evolved from sea-marginal to fluvio-lacustrine dominated systems during a single progradational episode of the coastal plain. However, this chapter focuses on sea-marginal and platform settings, it does not present models for continental sabkhas as they are discussed elsewhere in this volume, and by Hardie, 1984; Kendall, 1984; Lowenstein 1982; and Warren, 1989. Coastal sabkhas are to be found along the western and southern coasts of the Arabian Gulf, the coasts of northwestern Australia, Northern Africa, and Israel. The best studied marine sabkhas in the world today occur along the coast of the United Arab Emirates between the Qatar Peninsula and Ras Ghanadah in the southeast Ara- bian Gulf, especially well studied are the sabkhas southeast of Abu Dhabi City (Fig. 2.3A). These salt flats are the best documented Holocene counterparts of ancient salt flat sequences deposited across carbonate platforms with eolian and "wadi" (ephemeral stream in Arabic) input from the landward side. The Arabian Gulf sabkha is approximately 200 miles long and up to 15 miles wide; it has prograded 1 mile every 1000years (Kinsman, 1969). The salt flat occurs landward of coralgal reefs, sponge banks, oolite shoals, skeletal grainstone banks and shallow lagoons, all deposited atop drowned Pleistocene dunes and sand sheets (Fig. 2.3). The Abu Dhabi sabkha is unique in the large volumes of nodular anhydrite and

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SCALE Km

ARABIAN €2 GULF

Fig. 2.3. Surface lithofacies in the varioussabkhas of the United Arab Emirates A) Abu Dhabi area, barrier island - lagoon sabkha. Most of the detailed work was done in the 1960s in the supratidal flats behind Khor Qirqishan; B) Ras Ghanadah to Dubai, mainland tidal - beach-dune sabkha; C) Umrn a1 Qaywayn, estuarine sabkha; D) Ras a1 Khaymah, f an-delta sabkha. Base map sho wing geology courtesy of Ministry of Petroleum and Mineral Resources, Abu Dhabi.

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1 SCALE: Km 15 OA

ARAB IAN

GULF

EXPLANATION

Reef - Holocene coralgal boundstones and related deposits (Qc).

Tidal Delta or Shoals - Mainly oolitic sands wlth lesser water logged marshes and mangroves (Qsh).

Beach-Dune Grainstones - Light coloured calcareous, quartzose. and oolitic sands (Qbs)

Coastal Sabma Calcarco-s s ds and mbaoy SanaS yi In a sp ac ve s,.lale nodr es an0 salt crusts F ooaed 0, stom an0 1spr ng Iaes ana octas onal uaa rJnall lacs Continental Inland Sabkha - Silt and muddy sands with displacive @# and sulfate and halite evaporites. Flooded by wadi or rising groundwaters (Qis)

Calcareous sandstones and grainstones - Porous cross-bedded sands (Miliolite) commonly with well-rounded grains (Qm).

Gravel Plain and Desert Dune - Gravel covered desert floor, locally with sand veneer and sulface evaporites (Qesl,2) a Alluvial Fan - Boulders, gravel and sand deposited by ephemeral runoff from the Oman Mountains (Qg)

Mountains - Upllfted Tertiary and Cretaceous sediments

Holocene dolomite currently forming in the supratidal zone. This uniqueness first attracted the attention of the geological community in the late 1950s and early 1960s. The initial works by Curtis and others (1963), Evans and others (1969), Kinsman (1963, 1965, 1969), Shearman (1963), Butler (1966, 1969), and Kendall and Skipwith (1969), provided the observational data for current sabkha models. Subsequent work in the Arabian Gulf sabkha usually considered only a specific problem within the sabkha

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setting such as; the hydrological setting (Patterson and Kinsman, 1981; McKenzie and others, 1980), dolomitization (Patterson and Kinsman, 1982; McKenzie, 1981), or it has reinforced the models of earlier workers (Butler and others, 1982). To understand sabkha deposition requires an understanding of the climatic and hydrological factors which create the conditions suitable for sabkha progradation. The Arabian Gulf is probably the warmest sea in the world; in the open gulf the surface temperatures range from 2OT in February to 34°C in August. Water temperatures on the open shelf in the vicinity of Abu Dhabi Island range from 23-34"C, and in the inner lagoon areas from 1540°C with an annual average temperature of 29°C. The shallow, inner lagoon waters of the Khor al Bazani have diurnal ranges of up to 10°C. Air temperatures on the sabkha can be as low as 5°C and as high as 50°C with average temperatures ranging from 23-33°C. Sediment surface temperatures on the sabkha are even more variable, with reported values up to 60-80°C on the surface of the algal mats (Kinsman, 1965). Shallow subsurface temperatures range from 22-40°C. In the summer the sabkhacan be extremely humid, especially at night when the humidity is often above 95%, yet at mid-afternoon the humidity can fall to around 30% (Bush, 1973). Salinities in the open Gulf are 3-4%0 above normal open marine salinities. In the vicinity of Abu Dhabi where the coastal islands restrict exchange with the open Gulf the lagoon salinities range between 54-67/00.To date, no free-standing gypsum-precipitating brine bodies have been found in the inner lagoon. The major sources of lower salinity waters moving into the Gulf are the normal marine waters flowing through the Straits of Hormuz and the fluvial inflow of the Shatt el Arab delta. The Shatt el Arab delta is the confluence of the Tigris, Euphrates and Karoon rivers. It forms an enormous deltaic wedge at the northwestern end of the Gulf and is gradually spreading out into the dominantly carbonate province of the open gulf. There are small sabkhas dominated by siliciclastic sediments on the Kuwaiti portion of this substantial wedge of deltaic sediments (Gunatilaka et al., 1987). Surface runoff also adds small volumes of siliciclastics to the carbonate sediment of the northern Gulf, especially along its northeastern shore where ephemeral streams sourced in the Zagros Mountains are feeding alluvial fans along the Iranian coast. To the south along the Emirates sabkhat, evaporation is aided by the prevailing onshore wind (the Shamal) which is most intense in the winter months. Shamal winds supply predominantly carbonate sediments to the sabkha as it drives sediment-laden lagoonal water onto the sabkha surface. At other times, offshore winds blow continental dune sand across the back side of the sabkha. Rain in the region is infrequent but torrential, especially near the Oman Mountains. Flooded wadis can carry continental quartzose sediments for several kilometers seaward across the continental side of the

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sabkha. As in any desert the annual rainfall is highly variable, in the six year period from 1958-1964 the highest annual reading on the sabkha was 6.73 cm, the lowest 0.33 cm. Rain floodwaters can sometimes cover portions of the sabkha for 2-3 weeks dissolving any surface and near-surface halite crusts. Physiographically the sabkha is characterized by its billiard table flatness giving a seaward slope of between 1:2000 and 1:3000 (Butler, 1970). In many areas one can stand on the sabkha and see nothing but the perfectly smooth spread of the sabkha surface out to the horizon in all directions (Fig. 2.4A). Locally this smooth surface is interrupted by small 'Sebels" (hill or mountain in Arabic) of older Tertiary rocks, by platforms of Pleistocene rocks (including miliolites), by tidal channel depressions and by low beach ridges. Wind deflation, infill of depressions by eolian and water-carried sediments, and subsequent equilibration to the capillary zone of the local groundwater table account for the flatness of the sabkha. The sabkha near Abu Dhabi extends some 15-20 km onshore from the shallow waters of the lagoon, across the algal mats and mangroves of the intertidal zone into low beach ridges and up onto the extensive flat surface of the sabkha supratidal (Fig. 2.3). The present sabkha sediments reflect 7,000 years of depositional history (Evans and others, 1969; Kinsman, 1969). Holocene deposition began some 7,000 years b.p. as a rapid transgression covering Pleistocene eolian sediments with lagoons and mud flats. About 4,000 years b.p. the Arabian Gulf shoreline reached its maximum Holocene strandline. This was followed by progradation and accretion of the sabkha to its present configuration. The modern sabkha depositional system is divided into three areas, the subtidal, the intertidal, and the supratidal (Figs. 2.4,2.5). Sediment successions deposited in these areas are assigned to subtidal, intertidal and supratidal facies. In a prograding sequence the three facies are stacked one atop the other to form the well known shallowing-upward sabkha cycle termed "the Trilogy" by Warren and Kendall (1985). There are numerous subfacies which can be identified within this tripartite division.

Subtidal Facies

Open marine sedimentation occurs on the seaward side of the northeast trending barrier island system (Fig 2.3A). The islands are composed of Holocene and cemented Pleistocene oolite shoals separated by channels several kilometers wide and 7-10 meters deep (Purser and Evans, 1973; Loreau and Purser, 1973). Channels are flanked by shoal-water deltas of oolitic sand where grains are agitated by opposing onshore waves and offshore ebb-tidal currents. In the shoals the mid grainsize and cortex thickness

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Fig. 2.4. A) Smooth. flat sabkhasurface covered in pressure ridges and eolian sifts of the ephemeral halite crust. A1 Zihaya peninsula, United Arab EmiratPs. B) Lagoonal sediments from the Khor al Bazam Lagoon, south o,f Abu Af Abvad I.sland, United Arab Emirates. Much of' the brown algae follows breaks in the .subtidal hardground extending across the lagoon floor (photo courtesy of Chris Kendall). C) Sabkha tepee atop the intertidal sediments of Khor Al Bazam south of Abu a( Ahyad Island. D) Algal mat sediments on southern edge of Khor a1 Bazam Lagoon, the mat is composed of laminated mat in the meandering tidal channels andpuffy mat in the intervening areas. E) Core with ulgal laminations cross cut by displacive gypsum prisms. Core taken f rom southeast of ilbn Dhabi city (see Butler et al 1982). F) Benchridge on the Al Zibaya .rabkha, Abu Dhabi. Rusted f orty-f our gallon drum .for scale.

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BEACH RIDGE

ALGAL MAT ALGAL MAT 13)

INTERTIDAL PEAT 12)

LAGOONAL SANDS/ ANDMUDS WITH ERTIARY ROCK HARDGROUNDS 1 I1 WITH OUTWASH FANS

ANHYDRITE TOGYPSUM

JUST ABOVE HWM MIDDLE SABKHA HALITE CRUST (91 HYDRITE POLYGONS (71 ANHYDRITE LAYER (81 GYPSUM-CARBONATE MUSH I61

LARGE GYPSUM T CRYSTALS I41 Pm Fig. 2.5. Composite idealized model showing vertical and horizontal distribution of sabkha sediments based on the barrier island - lagoon sabkhas of the Abu Dhabi region (after Warren and Kendall, 1985). All the peripheral diagrams are keyed to the central block diagram. HWM = High Water Mark. (I) Lagoonal carbonate sands and/or muds with carbonate hardgrounds, (2)Vaguely laminated lower tidal-f lat algalpeat, (3)Upper tidal-f lat algal mat; commonly laminated and cross-cut bydesiccationpolygons, (4) Large gypsum crystals (prismatic or lenticular), (5) Cemented carbonate layer, (6) Hightidal- flat to supratidal mush of gypsarenite and carbonate, (7) Supratidal anhydrite polygons with windblown carbonate and quartz, (8) Anhydrite layer replacing gypsum mush and forming diapiric structures, (9) Halite crust formed into compressional polygons, (I0) Deflated beach-ridge of cerithid coquina and carbonate sand.

Data Center ,09126599985,[email protected], For Educational Uses 82 SEA-MARGINAL AND PLATFORM CALCIUM SULFATE decreases away from the channel levees. At about 2 m water depth there is an abrupt change into a mixed ooid-bioclastic grainstone. Patch reefs occur immediately seaward of the barrier islands in the areas between the ooid shoals. When compared to the reefs of the open Indian Ocean they have a sparse restricted coral fauna dominated by genera of Porites and Acropora with Platygyra, Cyphastrea, and Styfophora. The depauperate reef biota is a response to the warmer and more saline waters of the Gulf. Many of the corals are dead or appear unhealthy and are coated by a flourishing coat of calcareous algae (Lithothamnion, Lit h ophy 11 u m and Gon i o 1it h on). Seaward of the Pleistocene barrier islands and their reefs is a broad extensive platform ramp dominated by skeletal sands. Offshore it passes into progressively deeper water floored by the finer grained carbonate silts and muds of the Gulf center. Even here in the open gulf, coarse sediments can be found as haloes around sea floor shoals and islands that are located above salt structures sourced in the Hormuz Formation (Purser, 1973; Warren, 1989). Restricted marine sedimentation occurs along the landward side of the barrier islands. This region includes the lagoon and the most seaward sediments of the sabkha. Brown algae thrives in the breaks in the subtidal hardground extending across the lagoon floor (Figure 2.4B). These extensive shallow water lagoons are filled with bioturbated packstones and wackestones that contain a diverse biota of many benthic species but relatively few phyla (Kendall and Skipwith, 1969). In the more open areas of the lagoon there are molluscan sands and Peneropfid forams, while in the more protected reaches of the lagoon there are pelleted carbonate mudstones. Major bioclastic sediment components such as molluscs, benthic forams, calcareous algae and pellets are present in similar proportions to those found in lagoonal settings of the Bahamas and Florida. In addition, in the Arabian Gulf lagoons, there are detrital eolian grains as well as gypsum and anhydrite in the more saline areas. Tepee structures are common on the subtidal floor of the lagoon as well as up-slope in the intertidal/supratidal zone (Fig. 2.4C; Kendall and Warren, 1987). Both geographical areas where tepees can be found have their own characteristic signature; lagoonal tepees contain an open marine fauna while intertidal/supratidaI tepees are dominated by cerithid gastropods, pelleted mudstones and other indicators of more restricted conditions of formation. Tepees in the lagoon grow by aragonite crystallization within the crust; a probable result of supersaturated bottom waters immediately above the crust (Shinn, 1968). In contrast, the intertidal/supratidal sabkha tepees form under repeated cycles of thermal cracking, mud infill, and cementation (Kendall and Warren, 1987).

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Intertidal Facies (lagoonal)

The intertidal is divided into an upper and lower intertidal facies, though the lower intertidal facies of some workers is equivalent to the most restricted lagoonal facies of others. The lower intertidal facies is an algal peat composed of the bioturbated remnants of an algal mat. The upper intertidal facies is a laminated or pustular algal mat facies spreading across an extensive belt of algal flats up to 1-2 km wide (Fig. 2.4D). Mats accumulate in areas too hostile for the cerithid gastropods and grazing animals living in the lagoon. Laminated mats grow in channels while the intervening areas are composed of pustular mats. Cores through the pustular zones are composed of a poorly layered algal mush up to 60 cm thick. Cores through the laminated mats are composed of laminated mudstones up to 50 cm thick composed of cyanobacteria interlayered with aragonite muds. Mat sediment is crosscut frequently by desiccation cracks and contains lenticular (diagenetic) gypsum crystals up to 2 mm long (Fig. 2.4E). Aragonite, magnesite and dolomite locally cement the surface sediments of the algal flats (Butler and others, 1982). When these intertidal muds are covered by prograding supratidal sediments they can be intensely dolomitized in the shallow subsurface (McKenzie, 1981). Arabian Gulf sabkhas, along with tidal hammocks in the Bahamas and the near coastal lakes of the Coorong, South Australia are some of the few areas on the world's surface of marine-associated Holocene dolomites. Along many parts of the active shoreline of the sabkha are low beach ridges of skeletal grainstones that separate the intertidal and supratidal zones (Figs. 2.4F, 2.5). Buried equivalents of these grainstone ridges, along with ridges composed of reworked eolian sands, have been observed on the mid-sabkha west of Abu Dhabi island (Warren and Kendall, 1985). Better documented Holocene equivalents of buried beach ridges within sabkhas are described from Kuwait by Picha (1978) and from NE Qatar by Shinn (1973). Holocene beach ridges in the Abu Dhabi sabkha are not to be confused with buried Pleistocene eolianites that form the basement to much of the sabkha Holocene sediments (Fig. 2.5).

Supratidal Facies

The supratidal of the sabkha is subdivided into three facies; lower, middle and upper supratidal (Fig. 2.5). Lower supratidal. The lower supratidal forms a belt up to 2.5 km wide just above the high water mark (HWM), and is typically a gypsum mush up to 30 cm thick.

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This zone is flooded once or twice a month during high spring tides. Merging with, and adjacent to, the mush is the crinkly to pustular algal mat of the intertidal algal flat. In a complete section, the gypsum mush consists of an upper and a lower part. The lower part is composed of aragonite mud containing algal filaments and isolated gypsum crystals. The upper part is a thixotropic mush of lenticular gypsum crystals mostly of sand size with occasional larger crystals up to 1-1.5 cm long. The mush precipitates within and at the top of the more marine portions of the supratidal capillary zone to create a bed of lenticular gypsarenite. Surficial gypsum mush can be altered to bassanite (CaSO, %H,O) or anhydrite during hot, dry periods and then back to gypsum during cooler, humid periods (pers. obs.). Middle supratidal. The middle supratidal forms a belt up to 1.6 km wide and is flooded by lagoonal water on less than monthly intervals. It is an area characterized by a surface crust of ephemeral halite and silty muds that is underlain by sediments containing diagenetic nodular anhydrite (Fig. 2.4). Pockets of gypsum crystals in the halite crust, especially in the zone of overthrust salt ridges, are diagenetically altered to anhydrite, while in the subsurface of this zone the partial replacement of the gypsum mush by anhydrite has begun. Subsurface and surface anhydrite nodules increase in size and number in a landward direction, until in some areas they eventually form interlocking polygons in anhydritic beds 60 to 70 cm thick. Still further landward the surface anhydrite is periodically covered by quartzose sediments so that the new matrix allows further CaSO, growth such that the buried anhydrite polygons are festooned and erosionally truncated. At the same time the anhydrite-gypsum is growing in the capillary zone, there is widespread precipitation of displacive and pore-filling gypsum in those intertidal sediments that lie beneath the gypsum-anhydrite unit and the sabkha water table. This displacive gypsum grows thr.ough the algal' mats disturbing and destroying the now buried laminae (Fig. 2.4E). The middle supratidal is also a zone where shallow flushing of marine-derived brines generates dolomite in lagoonal and algal mudstones. Upper supratidal. In the vicinity of Abu Dhabi, the upper supratidal zone is up to 4.8 km wide and is only flooded once every 4-5 years. According to much of the literature the gypsum mush has been completely replaced by coalesced nodules of "chicken-wire" anhydrite in units up to 30 cm thick (Fig. 2.6; Kinsman, 1969). The ongoing growth of anhydrite in layers creates thrust folds, ptygmatic and enterolithic folds, and diapir-like structures. Anhydrite cement infills the internal molds of gastropods. In the subsurface, the algal lamination in the intertidal unit has been partially or completely obliterated by the continuing growth of CaSO, nodules and large gypsum crystals (6-25 cm diameter).

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Fig. 2.6. Nodular and enterolithic anhydrite from the upper sabkha south of Khor a1 Bazam Lagoon, shows truncation of enterolith and subsequent accretion with the enterolithic anhydrite zone. Coin for scale (photo courtesy of Mohammed a1 Dubal).

Contrary to much of the literature, the supratidal of the Arabian Gulf sabkha is typically gypsum with lesser to minor anhydrite. In the mid- and upper-sabkha near the Al Zibaya Peninsula, some 40 km west of Abu Dhabi Island (Fig. 2.3A), there is an extensive supratidal unit (20-30 cm thick) composed predominantly of subhedral prisms and discs of gypsarenite. This layer develops where, according to commonly held perceptions of an Arabian Gulf model, the supratidal anhydrite should occur. Occasionally the unit contains nodules of anhydrite a few centimeters across; some nodules have a core of anhydrite surrounded by radial gypsum blades, others have a core of gypsum surrounded by anhydrite. In isolated tracts in the region, a few tens of meters across and tens of centimeters thick, the gypsum layer is completely converted to chaotic layers of nodular and enterolithic anhydrite. The two sulfate phases transform back and forth as they grow displacively in the supratidal sediments. In other words the supratidal of an Arabian Gulf sabkha is not always or everywhere dominated by nodular anhydrite nor is aJ the anhydrite primary. Observations by the author on the Al Zibaya sabkha agree with the work of Bush (1973) who argued much of the anhydrite in the Abu Dhabi sabkha is after a gypsum precursor. Whatever the mineralogy, the growth of displacive and replacive sulfates in the capillary zone above the water table jacks up the sabkha surface until the near surface sediments leave the capillary zone and enter the vadose zone (Fig. 2.7). There they dry out and are blown away by the wind leaving behind the characteristic erosion surface which caps a sabkha sequence. This process can happen a number of times where an early erosion surface is preserved within the nodular anhydrite zone (Fig. 2.6). Erosion surfaces within and capping nodular sulfate horizons are the most important features of a sabkha sequence. They show that sulfate growth was syndepositional and controlled

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raised sahkha surface

sabkha surface ----_-

displacive growth of anhydrite matrix of capillary and gypsum carbonates raises sabkha andh surface siliciclastics I I. ;.I t--sabkha'water table --* algal ma1 I phreatic TIME 1 zone TIME 2

~ - deflationvadose sediment of leads to truncation of uppermost sabkhasequence

TIME 3

Fig. 2.7. Sabkha hydrology showing the effect of displacive evaporite growth (black) within the sediment column. The increased volume of evaporites in the column lifts the near surf ace sediments into the vadose zone where they dry and are deflated to create an erosion surf ace. by sabkha hydrology. On the most landward side of the upper supratidal the anhydrite is reconverted to gypsum by the slow seepage of continental groundwater into the sabkha. Displacive halite crystals (1-2 cm across) can be found in the upper few centimeters of sediment in the upper supratidal attesting to the infrequency of flooding in this region. These halite containing sediments are covered by an ephemeral crust of halite that is crosscut by pressure ridges. Displacive halites are important indicators of supersaturated salt flat pore fluids and contradict the notion that no halite is presewed in the Arabian Gulf sabkha.

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Landward of the sabkha is an escarpment of Tertiary limestones lined and capped by Pleistocene sediments (Fig. 2.3). The Pleistocene sediments are made up of cross-bedded windblown carbonate and quartz sands. The carbonate portion of these sediments is known locally as "miliolite",so called because of the abundance of forams. The quartzose portion is the outcropping equivalent of the Pleistocene eolian dunes which underlie the Holocene sabkha. The dunes were deposited during the last glacial eustatic low in the gulf when an erg was active over much of the area between the Oman Mountains and the position of the present coast. An erg is an extensive area covered by shifting sands (sand sea) with a complex suite of dune types. The "miliolites" were probably deposited during sea level highstands of the Pleistocene, but age relationships between the quartzarenite dunes and the "miliolites" are not well understood. Both can be easily differentiated from the Holocene beach ridges which subcrop in the mid-sabkha beneath the halite crust. The beach ridges are shoestring sands dominated by cerithid grainstones, the "miliolites" are shoestrings dominated by miliolid foram grainstones, and the ergs are sheet-sands dominated by quartzarenites.

Limitations Inherent to the Sea-Margin Sabkha Model

An Arabian Gulf sabkha is a prograding sequence composed of subtidal, intertidal, and supratidal sediments. In Abu Dhabi the Holocene is up to 2.7 m thick and is composed predominantly of lagoonal carbonate sediments. The Quaternary sediments beneath the Abu Dhabi sabkha are up to 12 m thick - they are predominantly Pleistocene dune sands and directly overlie the Miocene carbonate basement (Fig. 2.5; Butler, 1970). The subtidal contains a restricted marine biota, the intertidal sediments are often laminated mudstones, the supratidal unit has a matrix of clastics (carbonates and siliciclastics) that typically incorporates penecontemporaneous gypsum and/or anhydrite, and the whole succession is capped by, and/or contains, erosion surfaces (Figs. 2.5,2.8). Recognition of similar shallowing-upward carbonate trilogies is used to interpret ancient evaporites as sabkhas. In the 20 years since the first documentation, the sabkha has come to dominate the thinking of many geologists working with ancient marine sequences containing nodular anhydrite. As with any model based on a modern depositional system, there are modifications and simplifications made to produce applicability for ancient sequences. There are still many gaps in our understanding of modern sabkhas. At times the Abu Dhabi sabkha analogy has been applied to ancient successions without a full understanding of the implications of the Arabian Gulf sabkha hydrology and depositional setting (Warren and Kendall, 1985). For example, many thick, ancient

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Crossbedded carbonate or quartz EOLIAN-WADIE sand-silt with mot marks. Deposited QQuartz sand-silts as desert dunes, beach ridges, wadi variable thickness sheetflood, and tidal channel till * Unconfonnity (erosion surface) Anhydrite or gypsum nodules and sands often show "chicken-wire" and SUPRATIDAL "enterolithic" shucture. Thickness less than 1 meter Evaporites form in a matrix of (Height of carbonate or terrigenous clastics. May Capillary Zone) be dolomitized.

> > Gradational contact Organic-rich algal-laminated mud with fenestral voids and mudcracks. Lamination decreases downward and passes into burrowed pelleted lime-mud or dolomitized mud. ** Bored grainstone crust with tepees 0 Cross-bedded carbonate grainstone (tidal bar or beach ridge). May be oolitic or siliciclastic. Often underlain SUBTIDAL by coral or other open marinefacies. Variable thickness (Thickness dependant @ Lagoonal facies. Gray bunow-mottled on laaoon water lime mudthat mav be dolomitized

Fig. 2.8. Complete Holocene sequence from a coastal sabkha based on the Arabian Gulf subkhat (not to scale). The subtidal is divided into a high-energy shoal and a low-energy lagoonal sequence.

sequences of pure nodular anhydrite have been explained as stacked sabkhas deposited under conditions of slowly rising sea level. Yet for ancient thick and pure platform evaporite sequences this interpretation is opened to question by observations of the thickness and lateral extent of capillary zones in modern sabkhas. Within the capillary zone, in a matrix of biogenic and siliciclastic sediment, sabkha evaporites grow displacively as nodules and crystals. The thickness of the anhydritecon- taining portion of the intertidal/supratidaI is limited by the thickness of the sabkha capillary zone (Fig. 2.7). As the laws governing capillary pressures have not changed through time, the evaporite dominated portion of any single sabkha cycle can be no more than a meter or two thick (Warren and Kendall, 1985). In addition to thickness, it is also difficult to form a pure evaporite sequence within a capillary zone. Matrix must exist before the evaporite nodules can grow. Hence the difficulty of explaining meters-thick pure evaporites as stacked sabkhas. The only exception was to be found in ancient platform evaporitic mudflats that were located many kilometers from any detrital source other than reworked evaporite. Then the

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matrix could be sulfate, originally from a nearby gypsum-depositing lagoon or salina, that was reworked into the salt flat by storm or eolian transport. When such a succession was preserved, the result was a sulfate-dominated sabkha unit interlayered with subaqueous sulfate units. Once buried and converted to nodular anhydrite, such sequences are difficult to interpret in terms of depositional setting. The delicate balance between deflation of the sabkha surface and the position of the water table means the supratidal zone of a sabkha sequence has a maximum width and thickness dependant on the hydrology. The presetved sabkha may represent only a portion of what was a much wider and thicker supratidal succession. As the modern Arabian Gulf sabkha progrades, the water table on the landward side falls, allowing the sediment above the capillary zone to erode and deflate (Patterson and Kinsman, 1981). Deflation of the modern sabkha slows or stops at the cemented layer atop the subtidal and intertidal sediments (Sabkha Mutti, Arabian Gulf). Thus, in some laterally extensive ancient sabkhas, the supratidal sediments on the backside of the sabkha could have been completely removed by deflation. Instead, the subtidal/intertidal sediments were covered by a deflation lag or by eolian sediments. In this type of setting the subtidal and intertidal gypsum which formed syndepositionally below the crust would be converted to nodular anhydrite with burial. Intertidal/subtidal carbonate cycles in the Permian Tansill Formation that contain early dolomites and nodular anhydrite after lenticular gypsum, are probably a result of this process (adjacent to the Capitan Formation, Guadalupe Mountains, West Texas; Warren and Kendall, 1985). In the modern sabkha, the evaporites of the sabkha only overlie restricted lagoonal sediments,they never overlie shelf carbonates (Fig. 2.8). Yet in ancient sabkhas the. evaporites of the sabkha often seal open-marine shelf carbonates. Perhaps insufficient time has elapsed in the Holocene to form a geologically realistic sabkha analog, or some ancient nodular anhydrites have been mistaken for sabkha deposits. In a few millennia, and only if the Trucial Coast undergoes suitable subsidence, sabkha evaporites may prograde over the Holocene carbonates of the oolite shoals and the open shelf. And yet, the modern sabkha progrades across the inshore lagoon in response to the protecting action of the offshore oolite-reef shoal. The barrier islands screen the soft sabkha muds and evaporites from the erosive forces of the open gulf waves and the Shamal storms. Once the sabkha has infilled the lagoon behind this buffer, it is unlikely the sabkha will prograde much further seaward without the deposition of another sand shoal or reef, a little further offshore. Yet the 'Trilogy" is supposedly composed of three laterally continuous facies belts with subtidal overlain by intertidal overlain by supratidal sediments. In reality, sabkha progradation, using an Arabian Gulf analog over geologically significant time scales, is likely to be a punctuated process. It would not be

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a simple continuous seaward movement of the ‘Trilogy”,never broken up by coquinas, barrier islands, ooid shoals, beach ridges or reefs (Warren and Kendall, 1985). Geologically realistic models of ancient sabkhas should contain shoestring sands or barrier islands separated and sealed by sabkha sequences. Carbonate shoals (reservoirs) of the shoreline were frequently composed of off-lapping shoestrings that were separated by sabkha/salina filled depressions; e.g. oolite shoals in the Smackover Formation of the Gulf Coast and Mississippian formations composed of oolite-coquina shoals within the Williston Basin (Moore, 1984). The presence of beach ridges in the middle of the modern sabkha flat and along portions of the present shorelines of both the Abu Dhabi and the Qatar sabkhat, suggest that even the Holocene progradation of the sabkha has not been as simple as some sabkha proponents have suggested. Models used to interpret ancient sabkhas typically regard the Arabian Gulf as a carbonate province with the siliciclastic components ignored, or mentioned only in passing. This should not be the case. In almost all coastal areas along the southern Arabian Gulf the most landward portion of the sabkha is a sediment belt 2 to 6 kilometers wide with evaporites in a siliciclastic matrix. This passes seaward into a predominantly carbonate matrix. In ancient counterparts, such as parts of Seven Rivers Formation in West Texas (Spencer, 1987; Spencer and Warren, 1986,1987) much of the sabkha was preserved with a high quartz sand content.

Sabkha Dolomitization

Sabkha dolomite domains are erratic; dolomite is commonplace in intertidal/supratidal sabkha sediments but is also found in subtidal sediments (llling et al., 1965; Gunatilaka et al., 1984; Shinn, 1983). Most dolomite in the Abu Dhabi sabkha is diagenetic, not primq nor detrital; penecontemporaneous dolomite has never been found in the present lagoon or in actively growing algal mats. Sabkha dolomite occurs below the surface of the prograding sabkha, usually as partially replaced aragonite muds in units less than a meter thick. On sabkhas bordering the Qatar Peninsula the dolomite is diagenetic and most abundant beneath the high intertidal zone where it forms 1-5 pm euhedral crystals in aragonitic muds (Fig. 2.9A). There is an inverse relationship between aragonite and dolomite in the Abu Dhabi sabkha; as the proportion of dolomite increases, the proportion of aragonite decreases. In other words, the dolomite is replacing an earlier aragonite phase. This penecontemporaneous “secondary”dolomite forms in isotopic equilibrium with brines at temperatures between 34-49°C (McKenzie, 1981). Aragonite, and perhaps high Mg-calcite, serve as intermediates in the formation of dolomite which proceeds via a

Data Center ,09126599985,[email protected], For Educational Uses SABKHA DOLOMITIZATION 91

la1

Fig. 2.9. Dolomite in sabkhat A) Cross section of the peritidal zone on the Qatar Peninsula sabkha, Arabian Gulf with the percentage of dolomite in the fine sediment fraction (after Bush, 1973). B) Zone of dolomite formation on the Abu Dhabi sabkha, Arabian Gulf. It is defined by the area of maximum seawater flushing of the sabkha supratidal sediments (after Patterson and Kinsman, 1982). dissolution-reprecipitation process. With time, the dolomite appears to "age"; there is an increased ordering in the crystal lattice, an increase in the crystal size, and continued isotopic equilibration with brines at the lower temperatures found deeper under the sabkha surface. A process of diagenesis and isotopic re-equilibration also occurs in the co-existing CaCO, phases. Sabkha dolomite is forming in areas of maximum flux of seawater through the Abu Dhabi and Qatar sabkhat (Bush, 1973; Patterson and Kinsman, 1981). Areas of dolomite growth follow the strandline, heading inland in the vicinity of tidal channels and moving seaward near tombolo islands (Fig. 2.9B). Storms and high spring tides drive seawater from the lagoon up onto the sabkha flat. The seawater sinks into the sabkha and evaporates as some of the pore water returns to the surface via the capillary zone. As it does, the remaining water concentrates to the point where it deposits aragonite

Data Center ,09126599985,[email protected], For Educational Uses 92 SEA-MARGINAL AND PLATFORM CALCIUM SULFATE and then a CaSO, phase. Precipitation of these diagenetic minerals removes calcium from the solution, raises the proportion of magnesium, and so raises the Mg/Ca ratio. The dense Mg-bearing hypersaline pore fluids sink downward and flow seaward through the sabkha sediments, dolomitizing the aragonite as they go. Dolomite occurs in areas landward of the algal mat to depths of 2-3 meters with the most intense dolomitization in buried algally-bound and/or lagoonal muds (Fig. 2.9B; Patterson, 1972; Bush, 1973). The amount of new dolomite falls off in more landward positions as these areas are infrequently flooded and less subject to flushing by Mg-rich waters. Hsii and Schneider (1973) proposed an alternate hydrological model to that of surface flooding, they called it evaporative pumping. It involved continuous subsurface flow of lagoonal seawater into the sabkha to replace groundwater lost by evaporation at or near the sabkha surface. With evaporative pumping, seawater flows landward in the subsurface of the sabkha flat; the reverse to the surface flooding model. Field evidence for evaporative pumping is contradictory and many authors feel it is not a viable mechanism to explain the distribution of dolomite in the sabkha (Morrow, 1982). If landward seepage was the sole supply mechanism for sabkha brines, marine evaporites would quickly precipitate in the sabkha matrix. This would reduce the permeability of the sediments, cancel the head difference, and so choke off landward flow before sufficient volumes of magnesium had flushed through the system (Land, 1983). Perhaps the mechanisms supplying seawater to the Abu Dhabi sabkha are not as simple as the preceding either/or presentation implies. Work by McKenzie et al. (1980) has shown that the sabkha hydrology may be characterized by three sequential stages; flood recharge, capillary evaporation and evaporative pumping. All three stages occur as the dolomitizing brine evolves. They also demonstrated that the sabkha hydrology is not operating within a simple homogeneous aquifer, rather the sabkha plumbing is composed of three aquifers separated by two cemented crusts which perform as aquitards. In Al-Khiran lagoon in Kuwait there is dolomite, it is diagenetic but is unlike the dolomite of the Trucial Coast. The Kuwaiti dolomite is found in the upper 10 cm of the lagoon sediments and is not associated with gypsum or other evaporites, nor is it detrital. Gunatilaka et at., (1984) proposed that sulfate reduction also played a role in the formation of the Kuwaiti dolomites. Nearby, in the sabkha behind the Al-Khiran lagoon is Mg-rich dolomite in thin layers that alternate with layers containing the more commonplace Ca-rich dolomites. This Mg-rich variety occurs in parts of the sabkha that today are under the influence of continental groundwaters and rarely covered by seawater (Gunatilaka et al., 1987). In this respect, the Kuwaiti dolomites are more akin to Coorong dolomites than to the seawater derived dolomites of the Abu Dhabi and

Data Center ,09126599985,[email protected], For Educational Uses UNITED ARAB EMIRATES 93

Qatar sabkhat. Not all the dolomite in the Arabian Gulf sabkha environment is diagenetic. Shinn (1 983) documents detrital dolomite with quartz-silt and various clay minerals as detrital components. This detrital dolomite occurs not only in association with the diagenetic dolomite in the intertidal and supratidal, but also occurs in bottom sediments of the open Gulf. Pilkey (1966) documented 10% detrital dolomite in fine-grained sediment from the central axis of the Gulf. Wind-carried silt-sized detrital dolomite must add a component to the volume of dolomite found in the sabkha sediments. However, the contribution is probably minor. Whether or not it is necessary for the nucleation of the calcian-rich diagenetic dolomite that forms the bulk of the sabkha dolomite has yet to be addressed. No matter what the origin of the sabkha dolomite, in terms of ancient counterparts, sabkha dolomites are penecontemporaneous, thin (<1-2 m thick), stratiform, and confined to restricted marine to supratidal facies. Sabkha dolomitization cannot explain pervasive dolomitization of marine limestones in carbonate-evaporite platforms (Warren, 1989).

OTHER SEA-MARGINAL SABKHAS

United Arab Emirates

There are at least five different types of sabkha in the United Arab Emirates region (Fig. 2.3) which are: 1) barrier island - lagoon sabkha, 2) mainland beach-dune sabkha, 3) tidal - estuarine sabkha, 4) fan-delta sabkha and 5) continental interdunal sabkha. Only one of the five, the barrier island - lagoonal sabkha, has been well studied. My co-workers and I are currently documenting the other types of sabkhas found along the Emirates coastline. Barrier island - lagoon sabkha. It typifies the area to the southwest of Ras Ghanada as the most aerially extensive sabkha on the coast of the Emirates. The sabkha is prograding into a barrier lagoon which formed behind a series of Pleistocene barrier islands (Fig. 2.3A). It is the most well studied region and is the type area for the Abu Dhabi sabkha model. Yet even this sabkha has complexities, such as buried beach ridges, abundant siliciclastic matrix, and a facies mosaic distribution of anhydrite-gypsum-dolomite-magnesite, which are not often considered in sabkha models. Mainland beach-dune sabkha. Located to the northeast of Ras Ghanada, it infills a depression behind an extensive Holocene mainland beach-dune that extends all the way from Ras Ghanada to Dubai (Fig. 2.3B). This type of sabkha lacks low-energy

Data Center ,09126599985,[email protected], For Educational Uses 94 SEA-MARGINAL AND PLATFORM CALCIUM SULFATE lagoonal sediments and its evaporites formed in the mixing zone between marine- and continental-influenced groundwaters. Tidal estuarine sabkha-delta complex. Located further to the northwest, the coastal section from Dubai city to Ras a1 Khaymah encompasses many tidal estuaries in the area characterized by wide tidal channels which cut back as much as 20 km into the coastal plain (Fig. 2.3C). Tidal flats off the sides of these channels are inundated by tidal waters and extensive sabkha flats line the channels to form tidal-estuarine sabkhas. The largest such complex can be seen at Umm al Qaywayn where a tidal estuarine sabkha-delta complex measures 12 by 40 kilometers (Fig. 2.3C). The sabkha in this area is subject to the effects of regular marine inundation and is characterized by abundant tidal channel sediments. Fan-delta sabkha. Ras al Khaymah and Rams are both areas dominated by fan deltas with sabkhas on their most seaward regions (Fig. 2.3D). The alluvial fans are fed by ephemeral streams from the Oman mountains and abut directly into the waters of the Gulf. In sabkhas on the distal portions of these fan deltas, nodular and bedded gypsum plus displacive halite grow in a matrix of siliciclastic fan delta sediments. Continental interdunal sabkha. Inland of the main coastal sabkhas are many interdunal sabkhas. In the late Pleistocene most of the Emirates were covered by an active sand sea (erg) covering the northeastern limit of the Rub al Khali. In many of the interdunal corridors there are today elongate interdunal sabkhas up to 60 km long and 10 km wide (Figs. 2.3A,C). These continental interdunal sabkhas are filled with late Pleistocene/Holocene evaporitic sediments. The evaporites are displacive sulfates (gypsum and anhydrite) mixed in with displacive and bedded halite in a siliciclastic matrix. There appears to be a much higher proportion of halite preserved in these continental sabkhas compared to the coastal sabkha. The occurrence of five different sabkha types on the coastal plain of the United Arab Emirates implies we can expect similar facies complexities in ancient counterparts. A sabkha in the Arabian Gulf is not a simple prograding belt of parallel facies tracts. Reservoirs in ancient sabkha associated sequences will show a variety of geometries both parallel and perpendicular to the paleoshoreline, dependant on the type of original sabkha (Fig. 2.3).

Gulf of Suez

After the Arabian Gulf, the next best understood area of sabkha deposition is probably the western margin of the Gulf of Suez along the coastal strip from Abu Rudeis to Ras Muhammad (Fig. 2.10A; Gavish, 1974,1980). This area, named El-Qaa,

Data Center ,09126599985,[email protected], For Educational Uses GULF OF SUEZ 95

GULF OF SUEZ NE sw SUPRATIDAL INTERTIDAL SUBTIDAL MEDITERRANEAN SEA I I H Beachrock .: :: --- -;--- Gypsum + .?4?: ~?;;~;;.'y..-:.a.w.. -I ...... ' 0.. e .., ., ,. ;: W Anhydrite Carbonate mud -1 . Groundwater

Halite Anhydrite + Gypsum Calcite Mg-Calcite Aragonite D A IJ Dolomite-- 12345670 DISTANCE IN m x 10

Fig. 2.10. A) General map of the Sinaipeninsula showing various sea-marginal evaporite settings mentioned in text (after Gavish, 1980). B) Generalized cross section through the coastal zone, Gulf of Suez (after Gavish, 1980) showing typical sulfate distribution in relation to the water table and distance from shoreline, as well aspore water salinities. There is no distinction made between authigenic and clastic components. is a flat alluvial plain composed of alluvial silicate minerals derived from Precambrian and granitic massifs to the east. El-Qda includes some of the hottest and driest conditions in the Sinai. Carbonate components in these sabkhas are less common than in Arabian Gulf sabkhas. Much of the matrix is siliciclastic with evaporites scattered as interstitial cements and only rare displacive textures. Evaporites are not as extensive as in the Arabian Gulf sabkha and have not undergone as much diagenesis (Gavish, 1974). The sabkha environment extends from the shoreline to the base of the alluvial fans (Fig. 2.10B). Subtidal sediments are mixtures of clastic fine sands derived from the alluvial input, and carbonate sands composed of skeletal fragments, pellets, ooids, and sporadic intraclasts. Intertidal sediments vary from abundant fine-grained marine carbonates on the seaward side to coarse-grained clastics on the landward side. In many places there is an elevated sandy berm or a strip of pebbly beach rock which clearly

Data Center ,09126599985,[email protected], For Educational Uses 96 SEA-MARGINAL AND PLATFORM CALCIUM SULFATE defines the upper limit of the intertidal zone. The supratidal zone generally extends a few hundred meters inland with the distance dependant on the regional slope. The most landward extent of the sabkha is generally defined as the maximum extent of the intrasediment evaporite minerals. In places where the beach-berm is non-existent, storm water occasionally covers large portions of the supratidal area. In such places a thin upper layer of halite may form over the top of the sabkha, but it generally dissolves after a few weeks. A generalized view of the mineral distribution in a typical Gulf of Suez sabkha is very similar to the mineral suite in an Arabian Gulf sabkha. Like the Arabian Gulf, the supratidal zone is where most of the evaporites are precipitated. This creates a matrix-rich evaporite unit no more then 60-80 cm thick. In the Gulf of Suez, gypsum is the dominant mineral in the recent sabkha and anhydrite is found in the more elevated older portions of the sabkha where it recrystallized from gypsum, thus in the Gulf of Suez the anhydrite is a replacement mineral. Gavish (1974, 1980) attributes most of the ion supply to the sabkha to be via subsurface seepage of seawater after the evaporative pumping model of Hsii and Schneider (1973).

Nile Delta Coast, Mediterranean Egypt

A modern sabkha is forming on the coastal edge of the Nile Delta along the Mediterranean coast near El Hammam; about halfway between Alexandria and El Alamein (West et al., 1979). There the coastal plain consists of ridges of oolitic and skeletal beach and dune grainstones. All the ridges, except the most recent coastal ridge, are cemented into hard limestone (Fig. 2.1 1A). Between the ridges are elongate interdunal depressions floored by sabkhas. The first sabkha depression immediately behind the coastal ridge of oolitic sand is more than 8 km long and a few hundred meters wide. Displacive nodular gypsum is growing in the carbonate-rich clastic sabkha of this most seaward depression (Fig. 2.1 1B). Landward, and to the south beyond the Abu Sir limestone (Pleistocene), is another larger interdunal depression typified either by a mosaic of halite-encrusted, algal-mat fringed lakes, or by sabkha flats characterized by small gypsum nodules and lenticular crystals. Until the 12th century a branch of a Nile delta-lagoon (Lake Maneotis) occupied much of the second depression. Sabkha evaporites presumably postdate the fall in water table that was tied to the drying of the basin. Much of the matrix is brown, laminated, fine-grained quartz and carbonate sandy silt carried into the depression by floods or wind. It was originally deposited during floods of the lagoon by Nile waters. Intrastratal gypsum nodules push aside sediment laminae as they grow in

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1.5 B

Fig. 2.11. Coastal Sabkha, Mediterranean coast, Egypt (after West et al., 1979) A) Surficial geology showing the location of the sabkhat in interdunal depressions. B) Vertical section through first sabkha depression showing groundwater flow and the resultant evaporite mineralogies. this matrix. Nodules are typically spherical or elliptical with diameters between 1 and 4 cm. Sometimes they form upward-bulging almost diapiric structures that are incipient enterolithic structures. Most gypsum nodules form in the capillary zone to give a nodular

Data Center ,09126599985,[email protected], For Educational Uses 98 SEA-MARGINAL AND PLATFORM CALCIUM SULFATE evaporite unit up to 60 cm thick. The Nile Delta sabkhas are interesting for two reasons. First, they are examples of matrix-dominated siliciclastic sabkhas forming in a lagoon system on the coastal edge of a huge deltaic complex that is locally dominated by carbonates, including oolite grainstones. Second, the nodular structures in the sabkha facies are the result of the growth of displacive gypsum, not anhydrite.

Mediterranean Coast of Northern Sinai

The northern Sinai coast is bordered for about two-thirds of its length by the Bardawil Lagoon. Along the lagoon’s mainland shore is an extensive mosaic of sabkhas and dry sand flats (Fig. 2.10A; Levy, 1977, 1980). Sabkhas are of two types; 1) coastal sabkhas (also called lagoon-beach-sabkhas) at present connected to the lagoon and flooded occasionally by seawater, and 2) inland sabkhas cut off from the lagoon by sand dunes and forming under the influence of groundwater discharge with occasional flooding by rainwater (Fig. 2.12A). Both types of sabkha are typified by siliciclastic matrices. Displacive evaporitic salts and carbonate crystals grow in the capillary zones to form supratidal evaporitic units up to 60 cm thick. The largest documented coastal sabkha, Sabkhat Nathal Yam, is about 7 km long and extends about 1 km inland. In the coastal sabkhas, the sulfate mineral is always gypsum and no anhydrite has been found in the sabkha. Magnesium calcite, dolomite and halite were identified as minor components with scant halite crusts only in the most restricted parts of the sabkha. To the south of the lagoon-beach sabkhas, sand dunes extend inland to a distance of about 4 kilometers. The inland sabkhat form as small interdunal areas, up to 1 km across, along the coastal side of this large dune field. The inland sabkhas were once connected to the lagoon but are now isolated from any surface connection, most of the present water supply is via groundwater discharge into the sabkha depressions, especially after rainy periods. The age of the inland sabkha sediment and the origin of the waters (continental, marine or mixed) is not well known. Some of the inland sabkhat may be late Pleistocene. There is a soft white unit of calcian-rich dolomite (20-30 cm thick) in some of the inland sabkhas, and beneath this are well preserved aragonitic gastropods embedded in a muddy calcitic groundmass (Fig. 2.12B). Radiocarbon dating shows the dolomite is older than 35,000 years and so may not be marine in origin. Sea level was always tens of meters below its present level for the period 10,000 to 100,000 years b.p. Waters seeping into the interdunal depression in this period were unequivocally continental.

Data Center ,09126599985,[email protected], For Educational Uses RESERVOIR OCCURRENCE IN A SABKHA MODEL 99

SCHEMATIC CROSS SECTION THROUGH A LAGOON BEACH SABKHA. Sabkhat Hawash ..N Lagoon Water body c J u J,.:..~~:::::;.:~.i.i:::.:... , ...:'.: :,::.. - . ,. . . ,. .._-2A- :. .:. . , .;..., .- ,~ -.y ,_... 7.~. . :,.:...... : (....., . '. . .::. . ..

SCHEMATIC CROSS SECTION THROUGH ANCIENT SABKHAS NOT CONNECTED TO THE LAGOON Lagoon

A Halite 1",","1 Gypsum Water body

COLUMNAR LITHOLOGIC SECTION IN A LAGOON BEACH SABKHA (Nahal Yam) Depth cm

20 Detrital sand 30

60 Detrital sand

B 70

Fig. Sabkhas of the Bardawil lagoon, Northern Sinai (fig. OA for .xa..ty) A) Schematic cross sections through lagoon beach and inland dune sabkha showing the differences inscale andgeometriesof the evaporitic unit (after Levy, 1977). B) Composite vertical section throLgh a lagoon beach sabkha with a siliciclastic matrix (Nahal Yam) based on cores taken at.20,500 and 900 meters from beach line (after Levy, 1977).

Reservoir Occurrence Using a Sabkha Model

Potential reservoirs in a prograding coastal sabkha lie not in the evaporite dominated section of the sabkha supratidal, but in the associated subtidal/intertidal grainstones and dolomitized mudstones (Fig. 2.13). The evaporite-plugged supratidal sediment acts as a seal to these underlying porous units. Grainstones form in a number of settings; as a beach-barrier system separating the open marine from the restricted marine of the lagoon, as beach ridges which become encased in the evaporites of the mid-sabkha, as tidal deltas which cut the barrier islands, and as sandy infills of tidal channels that cut across the mudflat. Barrier islands and

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tens of kms-

CONTINENTAL FACIES 500

IDA1

Tidal sand shoal

Fig. 2.13. Block diagram off acies relationships and reservoir occurrence in prograding sa b k ha. beach ridges form shoestring sands that typically parallel the paleoshoreline and depositional strike. Tidal deltas form sand lenses perpendicular to the depositional strike, while the tidal channel sands infill furrows that meander down the paleoslope. Dolomitized peritidal mudstones, created by fluxing sabkha brines, follow mud-rich trends of the lagoon and algal flats, and so parallel the shoreline. Most modern dolomitized mudstones were deposited as sabkha lagoonal and intertidal muds and were subsequently dolomitized beneath supratidal evaporites (Patterson and Kinsman, 1982). Thus in ancient counterparts, the best diagenetic permeability was created along the interface between the sabkha supratidal and the marine. In ancient sequences it is important to distinguish between sabkha dolomites and brine reflux dolomites. Sabkha dolomites are thin, finely crystalline, penecontemporaneous units usually no more than a few meters thick that are confined to shallow subtidal and intertidal muds. Reflux dolomites are thick, medium- to coarsely-crystalline sucrosic textured units often located beneath thick evaporite beds; they are pervasive dolomites often tens to even hundreds of meters thick and typically form from a precursor of restricted-marine and open-platform limestones (Warren, 1989).

MODERN SEA-MARGINAL SALINA MODELS

In the AGI glossary, a salina is defined as a body of saline water such as a salt pond, lake, well or spring or a playa lake. The term salina is best described as small-scale bodies of hypersaline water or their associated subaqueous evaporites. By this definition

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salina evaporites are subaqueous units, less than 10 - 15 meters thick with lateral extents less than tens of kilometers. Like Holocene sabkhas, Holocene salinas can be either coastal-marine or continental depending on the predominant water supply to the basin. The most important characteristic of salina deposits is that they were precipitated from a standing body of brine. Until the 197Os, most subaqueous textures had been documented in ancient strata such as the Silurian salts of the Michigan Basin (Dellwig, 1955) and the late Miocene evaporites of Sicily (Schreiber and Decima, 1976). The disadvantage of such evaporite studies was that water depths had to be inferred and could not be measured directly. Many ancient evaporite basins had water depths that were subject to rapid and large fluctuations. A more reliable indicator of water depth than inference was required for ancient strata. In the late 1970s and early 80s subaqueous sulfate textures and observed water depths were published based on Quaternary salina evaporites in Baja California, southwest USA, Sicily, Southern and Western Australia, and the Middle East. This led to an improved understanding of the significance of laminar and bedded sulfate evaporites but, like many models based on Holocene sedimentation, suffered from the parochialism of lateral scale that is endemic to much of Holocene studies (Davies and Nassichuk, 3 975; Loucks and Longman, 1982). In the last decade more and more ancient evaporites, previously interpreted as sabkha sequences, are being re-evaluated and have often been found to contain evidence of a number of depositional settings. Most ancient evaporites are not simply sabkha or subaqueous, rather they are combinations of depositional settings, evolving over time and distance from one setting, say sabkha, into another perhaps subaqueous environment (Fig. 2.2; Kendall, 1984; Schreiber 1988; Warren, 1989). With the models available at the present time the distinction between ancient subaqueous and subaerial evaporites is readily made (Warren and Kendall, 1985). However, it is much more difficult to assign water depth in ancient subaqueous evaporites, especially in calcium sulfate settings where gypsum converts to anhydrite with burial.

Hydrology of Modern Coastal Basins

Marine and meteoric waters flow into modern sea-marginal basins as surface inflow via inlets and rivers, and as subsurface inflow via resurging groundwaters and seeps (Fig. 2.14; Warren 1986a). Marine-derived waters in coastal basins are freshened by rainfall and riverine input, or are concentrated by evaporative loss. No matter if the basin is viewed across weekly or millennia1 perspectives, the hydrology of such shallow-

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I 1. CONTINUOUS SURFACE CONNECTION 2. PERIODIC STORM & TIDAL INFLOW i

OCEAN EVAPORITE BASIN OCEAN EVAPORITE BASIN

I 3. SUBSURFACE INFLOW (SPRINGS) I 4. SUBSURFACE & PERIODIC SURFACE INFLOW

OCEAN EVAPORITE BASIN OCEAN EVAPORITE BASIN

Pig. 2.14. Hydrological setting o,f a sea-marginal evaporite basin (after Warren, 1986a). I) Continuous surface inflow; water loss from the basin is replenished by an influx of surf ace waters from the adjacent ocean. 2) Ephemeral surf ace inflow; basin waters are replenished by periodic storm and tidal inflow during times of storm surge and/or high spring tide. 3) Subsurface inflow; basin is connected to the oceans by a series of springs or seeps. 4) Subsurf ace and ephemeral surf ace connection. Stippled pattern represents a relatively impervious basement to groundwater flux; dot pattern represents a more permeable, unconfined or confined aquifer. water evaporite basins changes with time. Short term or seasonal variations in weather patterns affect the mineralogy of salina precipitates. A sudden increase in rainfall or humidity will decrease brine salinity to produce less saline laminae such as carbonates in gypsum, or gypsum in halite. Across weekly to annual time frames, marine-fed salinas are less susceptible to changes in average water level than continental salinas. Marine coastal basins are connected, either at the surface or in the subsurface, to the isochemical reservoir of the ocean. Water input into continental salinas is much more subject to the vagaries of local climate. Due to shifts in relative sea level, coastal salinas are also affected by longer term hydrological changes (100s to loooS of years). A transgression raises the regional water level in the coastal plain and may flood a sea-marginal basin with seawater. This prevents or restricts the deposition of evaporite minerals in the salina. A regression, or progradation, in an arid area will lower the regional water level as the major source of water moves further basinward from a particular locale. This may isolate the salinafrom

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marine waters, so converting it to a coastal sabkha or a continental basin, or due to a lack of mother fluid, it may stop evaporite deposition as earlier deposits are karstified. Falls in the regional water level also reduce the hydrodynamic head supplying seawater to groundwater-fed coastal basins. This may not completely dry out the basin, but it can increase the salinity of waters, changing the basin from a carbonate-gypsum depositing area to a gypsum-halite depositing one. Evaporite textures and diagenesis in modern coastal basins are direct responses to the hydrological setting of the basin. When considering ancient evaporites one should always try to ascertain the most plausible hydrology at the time of deposition. This will greatly enhance the predictive powers of any sedimentation model. Modern shallow-water marine evaporite deposits are only found in coastal lakes or salinas and, as will be shown later, a too strict application of the law of uniformitarianism to these deposits can be misleading in interpreting ancient subaqueous evaporites. However, the textures observed in these modern subaqueous settings are indispensable in interpreting ancient platform evaporites. Coastal salinas are found around the edges of the Mediterranean, the Black Sea, the Red Sea, and the southern and western coast of Australia. Some of the best documented Holocene examples of widespread subaqueous deposition come from southern Australia and Lake MacLeod in western Australia.

Southern Australian Salinas

Individual salinas measure up to 20 x 12 km, with centers filled by as much as 10 meters of laminated, subaqueous gypsum (Warren, 1982a). The salinas formed in the last 6000-7000 years as isolated lacustrine depressions within surrounding Quaternary calcareous coastal dunes. Inflow into the depressions is mainly from resurgence of marine-derived groundwaters which seep from the sea through the dune aquifer into the salina depressions. The driving force supplying the water to the basin is evaporative drawdown; whenever salina waters evaporate to levels below the adjacent ocean, this moves seawater plus some meteoric groundwater through the dunes and into the salina basin edge. This inflow is reflected in the bull's-eye pattern of sedimentation with a carbonate rim surrounding a central zone of laminated gypsum (Fig. 2.15). The carbonate unit of the lake edge is a distinct unit from the more central gypsum with a sharp, often near vertical diagenetic contact (Warren, 1982b). The CaSO, phase in the basin is always gypsum, not anhydrite, as temperatures are not high enough, nor the baGn arid enough to deposit anhydrite.

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GIPSARENITEIIOItULlbN

LIMINATED GIPSARENITEI9I

LCiMlNATEDSELENlTE 181

SELENITE DUMESSI

CARBONATE FRINGE BASIN EVAPORITES

Fig. 2.15. Depositional facies in a modern subaqueous setting, based on the salinas of South Australia (after Warren and Kendall, 198.5). The facies can be subdivided into a bedded subaqueous gypsum f acies surrounded by a carbonate sequence. (I) Boxwork limestone showing remnant algal or evaporite structures, (2) Veneer limestone, a hard indurated crust that is overthrust into tepees in manyplaces, (3) Groundwater tepee, (4) lntraclast breccia, (5) Stromatolites - domal f orms commonly cap tepees, larger laterally linked f orms are more common in deeper waters ( >SO cm depth);stromatolites and tepees both form in megapolygonal patterns, (6) Coarse-grained gypsum (selenite) domes, (7) Poorly layered coarse-grained gypsum, (8) Laminated coarse-grained gypsum, (9) Laminatedgypsarenite - in some solinas this unit dominates, (I0) eoliangypsarenite, (I1) Gypsite (silt-sized gypsum) -typically f orms on vegetated surf aces.

The South Australian coastal zone experiences Mediterranean style weather with cool, wet winters and hot, dry summers. Annual rainfall is 300-600 mm/year and annual evaporation is 1500-2250mm/yr. Winter temperatures are 10-20°C and rarelyfall below freezing, summer temperatures are often in the thirties but can climb into the forties.

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Much of the salina gypsum is coarse-grained and bottom-nucleated; pore infilling gypsum is rare (Warren, 1982a). The gypsum is usually more than 90% pure and is laid down as part of a shallowing upward evaporite unit (Fig. 2.15). At the base of a typical succession are massive, poorly layered domes of coarse-grained gypsum (selenite). Elongate gypsum crystals in this zone show little or no preferred orientation and carbonate impurities are distributed randomly through the gypsum (Fig. 2.16A). Higher in the section the degree of lamination increases as the dome amplitude decreases. Domes pass into horizontally laminated selenite where individual laminae appear to crosscut large upwardly-aligned 40-60 cm tall gypsum crystals (Fig. 2.16B). These large ("giant") crystals at first appear secondary and their Miocene equivalents in the Sicilian Basin were once interpreted as secondary gypsum after anhydrite (Ogniben, 1957). Studies of identical crystals in the South Australian salinas have confirmed the interpretation of Schreiber and Decima (1976) that large, aligned gypsum crystals (30-50 cm long) are primary, not secondary, crystals that grew in waters a few meters deep. They grew on the salina floor with their long axes perpendicular to bedding, an effect of crystal impingement and growth alignment (Warren, 1982a). Large subaqueous gypsum crystals on the floor of the South Australian salinas actually grew upward each season by millimeter scale increments, each step was in crystallographic continuity with its underlying parent crystal. The carbonate laminae in the gypsum units formed by the precipitation and accumulation of aragonite peloids during the spring and early summer. The peloids are ostrocod and brine shrimp pellets, mixed with the micritized remnants of algal tubules. At first the peloids mantled the underlying gypsum and were often captured by a cyanobacterial mat that covered the gypsum crystals. During the ensuing summer and fall, peloids were encased by the upward poikilitic growth of the large gypsum crystals. The end result of ongoing poikilitic enclosure was a bed of coarsely crystalline, upwardly aligned crystals encasing laminae that appear to penetrate the crystals. In a complete cycle these coarse-grained gypsum beds are overlain by laminated, sand-sized gypsarenites (Fig. 2.15). The upper portion of the gypsarenite bed could be reworked to form a wave-rippled sand. Laminated and rippled gypsum is in turn overlain by a thin, massive, poorly-bedded gypsarenite unit deposited under seasonally vadose or subaerial conditions. Capping the whole succession is a unit of cross-stratified eolian gypsum and, in areas stabilized by vegetation, a pedogenic crust of gypsite (silt-sized) caps both lacustrine and eolian sediments (Fig. 2.16C). The carbonate unit that develops about the edges of many South Australian salinas, is composed of a boxwork limestone. It is 4-5 m thick and overlain by a fenestral limestone sheet that is less than 1 m thick (Warren 1982b). Boxwork limestone is a

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Fig. 2.16. Salina sediment, South Australia. A) Coarse-grained gypsum in the domal core at the base of the gypsum sequence, New Lake, Kangaroo Island. Pen for scale. B) Primary coarse-grained gypsum which haspoikilitically enclosed aragonite laminae, New Lake Kangaroo Island. Note aragonite laminae appear cross-cut by largegypsum crystals. C) Eolian gypsum sands capped by gypsite in the soil zone from Cooke Plains playa, South Australia. Spatula is 30 cm long and handle is at contact. D) Boxwork limestone showing indistinct layering and high porosity, Marion Lake, South Australia. E) The two halves of a mm-laminated stromatolite collected from the crest of a tepee, Marion Lake, South Australia. Scale is in cm. F) Tepee structure, Deep Lake, South Australia. The tepee suture shows multiple episodes of cementation and is surrounded by a surface covered with intraclast breccias.

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diagenetic unit created by the cannibalization of earlier lacustrine carbonate by inflowing ‘Tresher”,but still saline, marine-derived groundwaters. The “boxwork”contains algal structures, evaporite pseudomorphs and other highly altered bedded fabrics dependant on the pre-existing sediment types (Fig. 2.16D). Some areas of the limestone sheet above the boxwork limestone consist of subaerial algal tufas or subaqueous mm-laminated stromatolites and algal mats (Fig. 2.16E), other areas consist of fenestral capillary crusts crosscut by tepees (Fig. 2.16F). The stromatolites grew as current-aligned domes up to 40 cm tall in the continuously subaqueous parts of the lake edge or as mats in areas of very shallow and sometimes ephemeral surface water along the lake shoreline (Von der Borch et al., 1977; Warren, 1982b). In some areas a little further out in the lake the stromatolite domes cap tepee structures. Like stromatolites, the tepee structures are confined to the capping carbonate unit and form an extensive fenestral limestone sheet up to 60 cm thick. Tepees are typically found in areas of resurging groundwater where seasonal changes in the groundwater head encourage cementation and tepee growth (Warren, 1982b). Lime- stone sheets caught up in the tepee structures are mostly capillary crusts composed of fenestral lime mudstones containing stromatactic-like voids, pisolites, laminar cements and internal sediment. Marion Lake and nearby lakes deposited gypsum and stromatolites only when the narrow marine entrances to the embayments were blocked by a build up of sand spits and sea-grass banks (Olliver and Warren, 1976; Von der Borch et al., 1977). In contrast to most of the gypsum-filled salinas along the southern Australian coast, Marion Lake and some nearby salinas were not always enclosed depressions fed by marine springs. In the early Holocene, Marion Lake was part of a marine embayment depositing open-marine skeletal grainstones and wackestones on the lagoonal floor. Today these sea-grass bound sediments lie beneath the stromatolites and subaqueous gypsum of the ensuing lake fill. A similar scenario can be envisioned for the future of Shark Bay in Western Australia. That is, Shark Bay will evolve into a gypsum-filled salina as the accretion and progradation of the seagrass banks and sand shoals completely isolates Hamelin Pool and Depuch Loop from the Indian Ocean.

Lake MacLeod, Western Australia

Lake MacLeod is an elongate halite-gypsum salina, 140 km long and 50 km wide, located in the MacLeod graben at the northern end of Shark Bay (Fig. 2.17; Logan 1981, 1982,1987; Logan and Brown, 1986; Handford et al., 1984). The climate is arid with an

Data Center ,09126599985,[email protected], For Educational Uses 108 SEA-MARGINAL AND PLATFORM CALCIUM SULFATE annual evaporation rate of approximately 2600 mm and an annual rainfall of 230 mm. Lake MacLeod was linked to Shark Bay during Pleistocene marine phases, and again early in the Holocene. Around 5000 years ago the lake was isolated from Shark Bay and filled with a 12 m thick unit of halite, gypsum and evaporitic carbonates. The present surface of the lake is up to 3-4 meters below sea level and the salina is fed by evaporative drawdown from the nearby Indian Ocean. Some seepage areas are marked by perennial brine pools - Cygnet and Ibis ponds with surfaces 2.6 and 2.9 m below sea level. To the north and west of the lake, and at the surface, the lake is filled by a massive gypsum layer up to 5 meters thick - the Ibis Gypsite (Fig. 2.17). Ibis Gypsite, as used here, is a predominantly clastic evaporite composed of sand-sized gypsum crystals with less than 10% carbonate; locally it contains aligned subaqueous gypsum prisms. Typically the unit is strongly layered and contains current bedding, ripple marks, and scour and fill structures. Other definitions of gypsite used by others is as an efflorescent crust atop a mother gypsum bed (AGI Glossary) and any unit which is more than 50% gypsum (Logan and Brown, 1986). To the south, beneath a much thinner gypsum unit, there is a massive halite unit up to 7 meters thick - the Texada Halite. Texada Halite is a bottom nucleated unit composed of bedded halite with thin gypsumcarbonate stringers. It was deposited subaqueously as pyramidal and columnar (chevron) halite in water depths of 0.5-2 meters. Halite deposited in waters deeper than 2 m was a halite silt. Both the Ibis Gypsite and the Texada Halite overlie a thin Holocene unit of marine-lagoon carbonate - the Cygnet Carbonate (up to 1 m thick). The Cygnet Carbonate is an aragonite mudstone containing some coated grains and ooids at the northern end of the basin. Its upper part is intensely cemented by gypsum. This created a hydroseal, which in turn allowed brine to pond and salinities to rise into the halite stability field, hence the Texada Halite. The Cygnet carbonate is underlain by another marine unit -the Coolan Carbonate Member (up to 0.5 m thick). The Coolan Carbonate Member is made up of skeletal sands and coquina, mixed with lithoclasts and detrital quartz that were reworked from the underlying Pleistocene. ln areas of groundwater seeps about the salina margin, especially on the northwestern seaward side, a carbonate is precipitating - the Egret Carbonate (up to 1 m thick). The unit is coeval with the Ibis Gypsite. It lies on an irregular unconformity and tends to thicken in local depressions to 5 meters and thin to zero over Pleistocene highs. This unit contains pisolites and tepees - groundwater resurgence indicators - that were described by Handford et al., (1984).

Data Center ,09126599985,[email protected], For Educational Uses LAKE MACLEOD, WESTERN AUSTRALIA 109

S N 6 BEJALING BARRIER (lorn)

Beialina Sands A ttiRET MBR. SEA LEVEL Boolathanna Fm

HALITE --lo I Pleistocene 8 Tertiary Fm b’”’yc’ -CARBONATE 0 10 20 Km -12m

E PLATFORM --6 --a --I0 Pleistocene 8 Tertiary Fm

Fig. 2-17. Lake MacLeod, Western Australia. Stratigraphic cross sections show thickness of various members and their relation to basin morphology, note the Egret Carbonate Member is limited to marginal areas in thc northern parts of the basin (after Logan and Brown, 19x6).

Deposition of the various units in Lake MacLeod reflects the evolution of the basin hydrology. Seawater today flows into the lake as a series of carbonate precipitating springs fed from fractures and caves in underlying calcareous coastal dunes. Water entering Lake MacLeod.has normal marine salinity which increases as the waters move out over the lake floor and evaporate. Similar spring-fed normal-marine waters are thought to have fed the lake some 11,OOO to 7900 years ago. Then, 7900 years ago the transgressing Australasian Sea rose to within 5 meters of its present level and established a surface connection to the Indian Ocean across the Texada sill. Six thousand years ago the MacLeod Basin was joined to the Ocean by a 5 meter deep channel. The time of open marine connection is characterized by deposition of the Cygnet Carbonate and the Coolan Carbonate Member. From 5800 to 5100 years b.p., surface inflow was reduced to a few narrow shallow channels. Free ingress of the Indian Ocean across the Texada sill ceased at 5100 years b.p. as beach dune accretion completely blocked the surface passage. Lake MacLeod became isolated and marine brines, supplied by evaporative drdwdown, cemented the upper portion of the Cygnet carbonate. Isolation marked the

Data Center ,09126599985,[email protected], For Educational Uses 110 SEA-MARGINAL AND PLATFORM CALCIUM SULFATE onset of Texada Halite deposition which took place from 5300 to 3800 years b.p. As the salina filled with evaporites, the groundwater head drawing water into the lake was reduced, and the perennial water body finally became ephemeral. Meteoric input became more important at this late stage in the depositional fill, in the same way as in the South Australian salinas (Warren, 1982a). The salinity of the surface waters was once again reduced to where gypsum was the stable mineral phase and the Ibis Gypsite unit was deposited across the lake (Fig. 2.17). Local springs fed the lake throughout its Holocene history, Egret Carbonate was deposited adjacent to the springs and gypsum and halite filled the salina in an asymmetrical fashion away from these point water sources (Handford et a1.,1984).

HOLOCENE SALINA EVAPORITES

Solar Lake, Gulf of Elat

The Northern Sinai coast of the Gulf of Elat (Fig. 2.10A) is characterized by a steeply inclined topography and a clearly defined fringing reef (Aharon et al., 1977; Gavish, 1980). In most areas the mountainous granitic and metamorphic massif meet gulf waters to yield a narrow steep shore zone composed of slump blocks, coarse beach-rock plates and narrow reef flats. Locally, wadis extend as fan deltas into the gulf to create a wider less steep coastal plain. Small brine ponds or salinas have developed in depressions on this narrow coastal plain. Probably the best known of these is Solar Lake, located immediately north of a large fan delta (Wadi Murah) and 18 km south of Elat. Solar Lake is a meromictic brine pond about 140 m long and 65 m wide with a maximum depth of 5 m (Fig. 2.18). The Solar Lake pond is unusual in that its waters are heliothermal with bottom water temperatures reaching 57°C in the winter and spring. Year round surface water temperatures range from 20-28°C. The pond is separated from the open gulf waters by a beach-barrier ridge, some 60 m wide and up to 3 m high, that is composed of coarse-grained molluscan sand and pebbles of igneous and metamorphic rocks. Evaporative drawdown induces seepage of Gulf waters into the pool during high tide; this is the main supply mechanism for water into the pond. Minor amounts of seawater may also enter the pond by splashing over the barrier during strong winter storms. In summer, evaporation exceeds inflow, raising the total dissolved solids content of the pan brine to 165 gm/l. In winter, the inflowing seawater and minor amounts of rainwater form an upper water layer of relatively low salinity. This results in stratifica-

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B DOLOMlTlZATlON IN THE GULF OF ELAT SL-2 SL-4 SL-4 SL-5 E m 4 2 0 -2 -4 -6

25 50 75 100 125 150 175m (vertical exaggeration x 3) SL-2 SL-1 SL-1 SL-5 LEGEND

0- Conglomerate and - 10 loose gravels 20 - Algal mats 30 - - Bedded gypsum 40 with carbonate 50 - laminae 60 - Carbonate mud 70 - Li thoclastic 80 - sediment 90 - Hand cores 100 - Drill cores 110 - 120 - 130 - 140 -

Fig. 2.18. Solar Lake, Gulf of Elat. A) algal mats (dark brown rim) about the lake margin and the proximity of the lake to the Precambrian Massif (courtesy W. B. Lyons). Gulf of Elat to right, see f ig. 2. I OA f or locality. Vehicle at right of photograph f or scale. B) West - east topographic and lithologic traverse across Solar Lake showing the core locations and interpreted geology of Aharon et al., 1977.

Data Center ,09126599985,[email protected], For Educational Uses 112 SEA-MARGINAL AND PLATFORM CALCIUM SULFATE tion of the water body across a well developed halocline and corresponding thermocline. During times of stratification, the lower layer absorbs heat, convection is inhibited, and the bottom layer becomes anoxic as H2Sconcentrations climb in the hot bottom waters. By summer, the evaporation of the upper water body causes a weakening of the density stratification and finally overturn. At that time the water level in the pond is as much as 0.5 m below sea level. As the water body becomes homogeneous and aerobic, it cools down by convective heat loss to a consistent 28°C. The sediments of the pond reflect the mode of water supply and the associated density stratification of pond waters (Fig. 2.18B). The margins of the pond are covered by well developed algal mats, while its slopes are composed of a white crust of large bottom-nucleated gypsum crystals and aragonite. The deeper bottom is covered by black mud and a loose gypsum mush (Lyons et al., 1984). The algal mats are crosscut by well-developed polygonal cracks giving the mat a "lily-pad"appearance. The cracks form when the mat surface is subaerially exposed during the late summer and autumn. Beneath the algal and gypsum layers of the pond is a carbonate layer with Cerithium gastropods indicating deposition during the initial restriction of the pond. Carbonate laminae in the algal mats and in the gypsum crust are composed of authigenic aragonite, Mg-calcite, dolomite and calcite with the dolomite forming as a replacement of the aragonite (Friedman et al., 1973; Aharon et al., 1977). The bulk of the authigenic dolomite occurs in the gypsiferous sediment, not in the algal mats, and has an aragonite precursor. The dolomite is a form of brine-reflux dolomite, a result of the throughflow of hot hypersaline brines with a Mg/Ca ratio increased by the precipitation of gypsum.

Ras Muhammad Pool, Southern Sinai

At the most southern tip of the Sinai there is a small peninsula, Ras Muhammad, that is composed mostly of uplifted Pleistocene reefs (Figs 2.10A, 2.19; Gavish, 1980; Kushnir, 1981). At the termination of a 1 km long tidal inlet on the tip of the peninsula is a small brine pool called Ras Muhammad Pool; an oval feature about 100-170 m in diameter containing waters not more than 50 cm deep (Fig. 2.19A). It is separated from the tidal inlet by a porous barrier some 250 m wide composed of reef material and carbonate sand. Water enters the pool as seawater springs about the edge of the pool so that even in this very arid area the pool never dries up, although the extent and depth of water in the pool varies with the seasons. During summer and fall, the water volume in the pond is greatly reduced by high evaporation and salinities range up to 300%0. In winter and spring the pond expands and salinities fall to 1W!oo. Sediments in the pool are composed of alternating layers of gypsum mush and

Data Center ,09126599985,[email protected], For Educational Uses RAS MUHAMMAD POOL. SOUTHERN SINAI 113

B RAS MUHAMMAD N sr Seawater Halite crust POOL WATER

W w -1 100m Carbonate mud U

Fig. 2.19. Ras Muhammed pool, Southern Sinai peninsula. A) Aerial view showing proximity to the adjacent tidal lagoon and the Red Sea (courtesy G.M. Friedman). North is to lower right of photo B) Generalized geological cross section showing the asymmetric sediment distribution related to the position of the seawater springs. In the marginal flats gypsum occurs below the water table and anhydrite above. algal mats, there is no primary or diagenetic formation of anhydrite in this pool (Fig. 2.20A,B; Kushnir, 1981). Thickness of the gypsum layers ranges between 0.3 and 6 cm, averaging 1.8 cm. Porosity in the gypsum layers ranges between 64-69 percent. The gypsum in the layers is mainly prismatic gypsum (up to 6 mm long) associated with minor amounts of aragonite, calcite, quartz and feldspar. In the mid-portions of the cores through the algal sediments (about 60 cm deep) some of the algal layers contain gypsum crystals which appear to have grown displacively within the algal layer (Fig. 2.20B). The continued growth of gypsum in the algal layer eventually destroys the lamination. After the decomposition of the algal material the evidence of aformer algal origin is ultimately destroyed. Only the carbonate grains captured by the algal mat remain to indicate its former presence. The final fabric, as revealed in the lowermost layers of the cores (about 80 cm deep) taken by Kushnir, is a laminated gypsum mush with thin interlaminations of carbonate crusts and ooids with very little identifiable algal

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AUSTRALIA ,

Huff Lagoon

PERTH

1200 130-

Calcrefized eolianite terrain Alluvial sands f3 and calcrete sods Vegetated sand dunes

Mobile sand dunes -Salt Mrsh flats I I Barren zone U River Fan depostt B

Mobile dunes Ephemeral halite unit Pisolitic loose soils aGypsite unit mm] intraclast veneer unit Gypsum-mud unit

Ic D WRiver fan system m =Sand sheet unit >SEQUENCE Basal sheet unit Tamala eolianite m(calcrete capped) IA Hutt B

HUTT LAGOON

Fig. 2.21. A) Map of Hutt Lagoon showing the principal physiography and surficial geology. B) Stratigraphic cross sections of Hutt Lagoon (A& B after Arakel, 1980).

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marine lagoonal stage; accretion of coastal barrier-beach and dune complexes progressively restricted open marine circulation resulting in an upward fining marine sediment sequence. In Hutt Lagoon there was a simultaneous progradation of alluvial units from the Hutt River, 3) evaporitic pond stage; following the complete surface isolation of the lagoon from the sea, a sequence of gypsum was deposited in the lagoonal brine pond layered and laminated by mud. Most of the gypsum in this facies is composed either of bottom-nucleated gypsum crystals or reworked gypsarenites and less common gypsru- dites. The finer-grained muds form interlayers in the gypsum; they are composed of micrite pellets in Leeman Lagoon and calcareous clays in Hutt Lagoon. Gypsum layers are white and about 5 mm thick, while the muds are dark gray laminae 4-15 mm thick. Localized subaerial exposure resulted in the reworking of some evaporites toward the lagoon center. A pellet-intraclast grainstone, found about the edges of Hutt Lagoon, covers the gypsum-rich units of Leeman Lagoon. Arakel (1980) also mentions the occurrence of hemipyramidal anhydrite and nodular gypsum in this veneer. This is a rather unusual Occurrence of anhydrite given the semi-arid climatic setting of this coastal region of Western Australia. Both the nodular gypsum and the anhydrite are diagenetic phases forming under vadose, probably capillary zone conditions and 4) pond-playa stage; is the current stage. Evaporites precipitate both as ephemeral halite crusts from ephemeral pond waters and as vadose gypsum cements within subaerially exposed lagoonal sediments (Arakel, 1980).

Salina Comparison

Preservation and dominance of a thick halite unit in Lake MacLeod in Western Australia is unique amongst Holocene coastal salina evaporites. It is due to the arid setting of this salina compared to the semi-arid climate of the others which are the gypsum-filled lakes. In the southern Australian examples, meteoric water played a more important role throughout the depositional history and became increasingly important as the lakes filled with sediment (Warren, 1982a). Rainfall and resurging groundwaters periodically diluted salina waters keeping the surface waters below halite saturation levels, but in the gypsum precipitation field. In Lake MacLeod, more arid conditions meant seawater was always the dominant inflow. Once the surface connection to the Ocean was lost, the area’s greater aridity kept the surface waters at higher salinities than their southern Australian counterparts. This allowed gypsum cements to clog salina carbonates immediately prior to the onset of halite deposition. The resulting hydroseal allowed brines to pond and concentrate to where halite was a stable precipitate (Logan and

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Brown, 1986). The lack of preseived halite beds in Holocene salinas of the Middle East is at first puzzling as they are located in climatic zones that are even more arid than those of Lake MacLeod. However, water bodies in the region are much smaller than Australian salinas and are underlain by highly porous Holocene sediments. They are also located less than a few hundred meters from the open ocean. The associated high humidities and active flushing tend to prevent the salina bottom waters from precipitating a hydroseal. A seal that would otherwise allow ongoing brine ponding and storage at halite saturation levels. In general, independent of the degree of aridity, the inflow zones in salinas are concentrated on the seaward side of the salina margin and are marked by carbonate precipitation and algal mats. The dominance of a subsurface boxwork limestone around the edge of many South Australian salinas and its lesser importance or absence in salinas in more arid settings is probably due to relative differences in the influence of inflowing meteoric mixing in the subsurface with marine waters. If so, the presence of a thick boxwork unit surrounding an evaporite-filled lake may well be an indicator of a semi-arid verses arid setting (Warren, 19%). Climate controls the hydrology which in turn controls the evaporite mineralogy; halite is precipitated in almost all salinas, but is preserved in a handful. In salinas formed in arid settings, halite, a more saline mineral, can dominate the subaqueous parts of the pond. In semi-arid settings, and in the more marine-influenced parts of arid salinas, the preserved subaqueous assemblage is gypsum. The presence of beds of coarse-grained gypsum (selenite) makes the southern Australian salinas unique amongst Holocene salina gypsum deposits. Worldwide, the more usual form of gypsum is as layered to laminated gypsarenite-gypsite units that are composed of discs, needles, or cleavage fragments. Such gypsum occurs in the majority of modern coastal salinas, including Streaky Bay and Snow Lakes in South Australia, Lake MacLeod and Hutt Lagoon in Western Australia, in lagoons about the Red Sea, and in continental playas. Typically this clastic gypsum is deposited as laminated to cross-laminated and algally-bound sediments. Basal scour surfaces and intraclast breccias show that some of these clastic evaporites were subjected to further reworking during storms. Coarse-grained bottom-nucleated gypsum only forms when the bottom conditions are stable enough to allow sluggish nucleation of new crystallites (Warren, 1982a). The semiarid conditions of coastal South Australia, with annual evaporation rates between 1500-2250 mm and annual precipitation rates between 300-500 mm, favors the relative stability of the bottom brines in many of the density-stratified coastal salinas. More arid conditions of other coastal salinas around the world, means the salinities on the salina

Data Center ,09126599985,[email protected], For Educational Uses 118 SEA-MARGINAL AND PLATFORM CALCIUM SULFATE floor change more rapidly. Then multiple nucleation of crystallites occurs and gypsarenites or gypsites are precipitated. Controls on the relative rarity of bottom-nucleated giant gypsum in modern salinas also explains why giant gypsum crystals were commonplace in many ancient bedded sulfates. Unlike the small sea-marginal brine pans of today, many ancient sulfates were deposited on the floor of permanent shoal-water seaways that coveredvast areas of the continental platform. Bottom conditions on the floors of these ancient seas were relatively stable and rates of change of salinity were slow enough that gypsum could grow as large aligned crystals that either pushed aside impurities or poikilotopical- ly enclosed them.

MODERN SALINA DOLOMITE MODEL - Coorong, South Australia

Holocene dolomite is a relatively minor component in all of the subaqueous sea-marginal settings so far described. Where it does form, it is a replacement after an aragonite precursor. There is only one coastal area in the world where thick units of Holocene dolomite have accumulated as "primary"precipitates in association with other sea-marginal sediments-- the salinas of the Coorong Region in southeastern South Australia. Dolomite-forming lakes of the Coorong are situated on the prograding coastal dune plain of southeastern South Australia (Fig. 2.22). The Coorong Region of South Australia was first documented as an area of "primary"dolomite by Mawson (1929), but was not studied in any detail until the work of Alderman and Skinner (1957), Skinner (1963), Von der Borch (1965,1976), Von der Borch and Lock (1979), Rosen et al. (1987, 1989), and Warren (199Od). The climate in the Coorong region is Mediterranean-like, with hot, dry summers (Dec.-Mar.), and cool wet winters (June - Sept.). Prevailing winds in the area are westerlies and southerlies and afternoon sea breezes dominate throughout the study area. Maximum rainfall is 800 mm in the far southeast of South Australia, 600 mm around Kingston, and 400 mm in the area about Salt Creek (Fig. 2.22). In the Salt Creek region, evaporation rates exceed precipitation except for a few weeks in June and July (South Australian Bureau of Meteorology, public data). To the southeast, near Kingston, precipitation exceeds evaporation for 3-4 months each year (May-August). Air temperatures range from -1°C to 38°C with a mean annual air temperature of 13.5"C. Water temperatures vary between 10-28°C and the pH from 8-10. During 1982-85 the summer temperatures within subaerially exposed lake sediments were as high as 5VC, winter temperatures were as low as 5°C.

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Fig. 2.22. Regional setting and surf icialgeology of Coorong Lakes in the Salt Creek area. Shows the locality and major mineralogy of surf icial sediments as well as inset enlarge- ment of the Salt Creek lake chain [North Stromatolite Lake (13), Pellet Lake (2), Dolomite Lake (3), Halite Lake (12) and Milne Lake (I)]. Mineralogies of surficial sediments f rom Von der Borch (I965) and Warren, 1988.

Von der Borch (1976) observed that dolomite in the Coorong region, between Kingston and Salt Creek, formed in areas of the coastal plain where rainfall was less than 700 mm (Fig. 2.22). He suggested that south of the 700 mm isohyet, the reduced evaporation rates and higher rainfall prevented the concentration of lake waters to the range of salinities where dolomite could precipitate. Reconnaissance of the area confirms this observation and goes further by showing that the more magnesian-rich dolomites tend to be found at the more arid northwestern portion of the coastal

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Coorong plain near Salt Creek (Warren, 1990d). Detailed analyses of some Coorong lake waters can be found in Von der Borch (1963, and the hydrogeochemistry of the regional groundwaters in O'Driscoll (1960). Mg/Ca ratios in different lakes varies from 1 to 20. Sulfate concentrations measured by Skinner (1963) in several lakes in the coastal plain range from 6.1 to 8.2% and are close to the expected value for seawater sulfate (7.68%).

Lake Stratigraphy

In terms of texture, similar vertical sequences occur in all Coorong lakes no matter what the mineralogy of the Holocene sediment fill; the dominant facies is always a mm-laminated unit. Facies changes up-section and across the various lakes are controlled by two factors; 1) did the lake ever possess a Holocene connection to the marine-estuarine waters of the Coorong lagoon? and 2) what were the energy levels and degree of bioturbation at the depositional site? Lf the lake possessed an early Holocene connection to the marine waters of the lagoon then it is a type 2 lake and the basal unit is dominated by marine/estuarine skeletal grainstones-packstones (Fig. 2.23A,B). If there was no Holocene marine connection to the lake, it is a type 1 lake and the typical basal unit is a quartzose packstone to wackestone. Basal facies in either lake type contain small percentages of basal dolomite. Above the basal facies in some lakes is a massive to faintly laminated organic-rich unit. In lakes with a marine connection that was lost early in the Holocene (type 2a lakes), the levels of total organic content (T.O.C.) in this unit can be as high as 12%. The organic matter is repeatedly an oil-prone proto-kerogen. Overlying this unit is a mm-laminated unit of pelletal packstone to mudstone deposited once the estuarine connection to the open Coorong Lagoon was completely cut off by beach-ridge accretion. The proportion of grains in the laminated unit is a direct reflection of energy level and organic binding when the lake sediments were deposited. In areas of higher wave energy/bottom currents and little or no algal binding, the sediment is a packstone, such areas were more common about the lake edge. In areas of lower wave energy/bottom currents in relatively deeper water and algal binding, the sediment is a wackestone or at times a mudstone. These muddier sediments are more common in the central parts of the lakes. In a few lakes such as Milne Lake this laminated unit composed of evaporative dolomite up to'5 meters thick, more often is composed of varying proportions of aragonite, hydromagnesite, magnesian calcite and rarely of laminated gypsum/aragonite couplets. Capping the lake sediments is a "massive" unit of poorly layered packstone-mud

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I TYPE ZA FACIES ASSOCIATION ,I TYPE I FACIES ASSOCIATION I

Fig. 2.23. A) Schematic vertical section of a Coorong Lake showing f acies associations. B) Schematic cross section of a Coorong Lake showing f acies stratigraphy in Type 1 and 2 lakes; see text f or definition of lake types (after Warren, I988). stone usually less then 60 cm thick and showing a varying degree of induration. This unit usually holds the bulk of the evaporative dolomite found in the Coorong but can also be composed of margin dolomite, aragonite, hydromagnesite, magnesian calcite and magnesite. The massive unit contains a characteristic suite of sedimentary structures including: bioturbation structures, stromatolites, mud cracks, extrusion structures, tepees, and breccia fragments; all features indicative of at least occasional desiccation. The presence of this capping unit in conjunction with underlying laminated sediments is the most reliable diagnostic for an ancient Coorong counterpart.

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Dolomite Occurrence

Dolomites in the various lakes of the Coorong coastal plain in South Australia occur in three mineralogical associations (Warren, 199Od); 1) dolomite * Mg-calcite (widespread), 2) dolomite ~t Magnesite (common in northwest) and 3) dolomite f aragonite f hydro-magnesite (rare and only in the northwest). Most Holocene dolomite Occurrences in the near coastal lakes of the Coorong region are associated with Mg-calcite. Only in the more arid settings found at the northwestern end of the Coorong region, near Salt Creek, is dolomite commonly found in association with magnesite and occasionally with hydromagnesite and aragonite (Fig. 2.22). Here too the thickest and richest dolomite unit is found. Dolomite in the near surface sediments of lakes near Salt Creek, such as Pellet Lake and Milne Lake, is often associated with magnesite in the lake center and Mg-calcite about the lake margin. The dolomites associated with magnesite tend to be more magnesian-rich than dolomites associated with Mgcalcite, a direct reflection of the chemistry of the mother waters. Waters which precipitate magnesite are typically more concentrated and more magnesium-rich than waters precipitating Mg-calcite, hence co-precipitated dolomites will also be more magnesium-rich. Coorong dolomites are true "primary dolomites"; dolomite is precipitating as dolomite and not replacing an earlier carbonate mineral. Coorong dolomite occurs as two mineralogically distinct forms - type A and type B. They are deposited as three stratigraphically distinct forms - evaporative, margin, and basal dolomite (Fig. 2.24A,B; Warren, 1988). Type A dolomites tend to be Mg-rich with a crystalline but heterogeneous microstucture, the unit cell is contracted in the a, dimension; it has a heavier carbon isotope value and a well-clustered oxygen signature compared to type B (Rosen et al., 1987). Type B dolomites tend to be calcian-rich to near-stoichiometric, with expanded unit cell dimensions (ac) and co), they are more crystalline with a more homogeneous microstructure compared to type A, and show isotopically lighter carbon values, as well as a larger variation in oxygen isotope values. Stratigraphically the most voluminous type of dolomite is an evaporative dolomite, deposited as the last episode of sedimentation in almost all Coorong Lakes which contain dolomite. Typically, it is a capstone unit no more than tens of centimeters thick, although in Milne Lake near Salt Creek it has infilled the lake to form a dolomitic unit up to 4-5 meters thick. Evaporative dolomite is the most common dolomite type in most Coorong Lakes, and in the vicinity of Salt Creek it is dominated by "type A" dolomite. Further south near Kingston, the evaporative dolomite unit is "type B" dolomite (Fig. 2.22 for locality). Evaporative dolomite is the dolomite unit sampled by studies

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TYPE 21 FIClES ASSOCIATION TYPE I FACIES ASSOCIATION I

Fig. 2.24. Schematic crosssection oj lake dolomitessuperimposed onfacies distribution: A) stratigraphic dolomites; horizontal stipple represents evaporative dolomite, random stipple represents margin dolomite. dots represent basal dolomite B) geochemical dolomites showing distribution of type A and type B dolomite (after Warren, 1988).

published prior to 1987. In some lakes near Salt Creek that contain type A dolomite in the lake center, there is mineralogically distinct surficial dolomite that is found about the lake edge which is a more calcian-rich margin dolomite named a type B dolomite (Rosen et al., 1987, 1989; Warren 1988). Like the evaporative dolomite, this type B margin dolomite is not intergrown with other carbonate phases and defines areas where resurging continental groundwaters first enter the lake margin. There is a third type of dolomite in many Coorong lakes. This is a basal dolomite, a type B dolomite, which is more crystalline than the other two forms of dolomite. It formed some 6OOO years ago when the rising Pleistocene water table (driven by a transgressing sea) first caused continental groundwaters to crop out and evaporate at the surface. The various Coorong dolomites form by the evaporation of continental

Data Center ,09126599985,[email protected], For Educational Uses 124 SEA-MARGINAL AND PLATFORM CALCIUM SULFATE meteoric waters, and this controls the Occurrence of dolomite versus other lacustrine carbonates. Type 1 lakes. The thickest evaporative dolomites form in type 1 lakes where even the laminated sediments are dominated by dolomite (Fig. 2.23A). Type 1 lakes that are now fed by resurging continental groundwaters never possessed a surface connection with the marine waters of the Coorong Lagoon. Milne Lake in the Coorong National Park is a dolomite dominated type 1 lake as is Lake 5 (Fig. 2.22). Milne Lake contains a laminated magnesian-rich dolomite-magnesite aragonite-hydromagnesite unit up to 4 meters thick, capped by a massive, bioturbated dolomite-magnesite unit, half a meter thick. Not all type 1 lakes contain dolomite. Type 2 lakes. The majority of all those Coorong lakes which contain dolomite are either type 2a or 2b lakes (Warren, 1988). Sediments from type 2 lakes have basal unit sediments which are dominated by a marine biota. Unlike dolomite-filled type 1 lakes, the evaporative dolomite in type 2 lakes does not dominate the lake succession, rather it forms a massive capping unit up to 80 cm thick above other carbonates and magnesian carbonates (Figs. 2.23A, 2.24). Type 2a dolomite lakes possessed a surface connection with the marine realm that was lost early in the Holocene. Pellet Lake, Halite Lake, North and South Stromatolite Lakes, Dolomite Lake and numerous other unnamed lakes sitting in interdunal areas behind the landward margin of the Coorong Lagoon are type 2a lakes (Fig. 2.22). In type 2a lake sequences the deposited sediment was relatively fine-grained and the majority of the section composed of laminated mudstone-wackestone. Type 2a sequences often have up to 12% T.O.C. due to the finer grained nature of the organic-rich unit (Warren, 1988, Hayball et al., 1991). Lakes dominated by type 2a sequences often contain a basal dolomite within basal sands which were deposited immediately prior and during the Holocene incursion of estuarine/marine waters into these lakes. The calcian-rich marginward dolomite is also common in the lakes dominated by type 2a sequences. Areas of marginward dolomite are often found beneath capillaly crusts about the lake margin. Type 2b lakes possessed marine connections until much later in their Holocene history. Most type 2b lakes were arms of the Coorong Lagoon that were cut off by spit accretion and coastal-dune migration. Hence type 2b sequences are common in the southern reaches of the Coorong Lagoon (Warren, 1988). Compared to type 2a sequences, type 2b sequences typically contain coarser grained sediments with the bulk of the laminated unit composed of skeletal packstones. Type 2b sequences pass directly up from bedded to laminated packstone-grainstones of the marine/estuarine Coorong lagoon into a "massive" unit which in some lakes

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contains afew tensof centimeters of evaporative dolomite. Evaporative and marginward dolomite formed when the type 2b lake depression was isolated from the open waters of the Coorong Lagoon. Type 2b lakes that contain evaporative and marginward dolomite may not possess a basal dolomite as the early Holocene sediments were deposited under conditions that were too marine and too energetic for dolomite to precipitate.

Sedimentary Structures in the Massive Unit

The bulk of the sediments in the Coorong lakes are characterized by monotonous and nondescript lamination in the subaqueous mudstones (Warren, 1988). Far more diagnostic of the Coorong signature are the sedimentary structures to be found in the overlying massive unit (Fig. 2.25). They include bioturbation structures, mud cracks, extrusion tepees, breccia fragments, and algal biscuits, stromatolites - all features indicative of at least occasional desiccation. In winter, when the lakesfill with water, the more basinward dolomitic mudstones possess the consistency of yogurt and are partially bound by algae and "duckweed" rootlets. In late summer and early fall, when dried, the same evaporative dolomite mud often has a leathery consistency and is cross cut by desiccation cracks (Fig. 2.2SA). About the edges of most of the lakes, the same mud forms a hard indurated crust, where almost year-round evaporation has converted the mud to a cemented capillary crust (Fig. 2.25B). The crusts contain extrusion 'tepees and intraclast breccias (Figs. 2,2SB,C,F;Kendall and Warren, 1987). Marginward dolomites are often associated with this crust along with local cherty cements. Algal biscuits and domal stromatolites are also present in some lakes (Figs. 2.2SE,D).

Geological Implications

A model of Coorong-styledolomites must include the followingobse~vations(Fig. 2.26): 1) there is a typical vertical sequence dominated by the laminated unit and overlain by the massive unit, 2) dolomite is a primay precipitate in the Coorong lakes. There are three distinct dolomite assemblages precipitating in the Coorong lakes; a) dolomite associated with Mg-calcite occurs throughout the Coorong coastal plain, b) dolomite associated with magnesite or more rarely, c) with hydromagnesite. Assemblag- es b) and c) are found only in the more arid northern end of the Coorong near Salt Creek and not in the more humid region near Kingston, 3) type A dolomites precipitate in association with hydromagnesite and/or magnesite, Type B dolomites in association

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,Extrusion tewe

Indurated mud crust (layered and laminated)

Fig. 2.25. Sedimentary structures in massive unit A) Desiccation cracks forming in Coorong dolomites, partially filled with yogurt dolomite that is blown about the lake in thin sheets of surface water. Pellet Lake, South Australia. B) Capillary crust in the Coorong dolomite, cross-cut by extrusion structures. Lake 5, Coorong, South Australia. See Fig. 2.22 for location. C)Schematic of extrusion structure in capillary crust, .showing succP.ssivr episodes of fill and cement. Coorong. South Australia. D) Algal bioherm in North Stromatolite Lake, Coorong, South Australia. These delicate algal forms were normnlly water covered. Some of the hioherms have erosional edgrs, others have been covered by a growing algal cover forming low domal morphologies. Spatula handle is 12 cm long. E) Algal biscuits in South Strornatolite Lake, Coorong, South Australia. F) Magnesite breccia in a dolomitic matrix, Milne Lake, Coorong National Park.

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Fig. 2.26. Three dimensional schematic of ancient Coorong counterpart showing the f acies mosaic nature of lacustrine deposition within an extensive paleoaquif er. Groundwater flow lines show a dominantly seaward flow of brackish to saline meteoric groundwaters. Meteoric waters crop out and evaporate best in near-coastal depressions due to presence of the subjacent seawater wedge (after Warren, 1988). with Mg-cdkite and 4) dolomite forms in lakes and ponds atop a much more extensive Pleistocene aquifer that supplies ions to the evaporating waters of the lakes. In the Coorong region the aquifer is a Quaternary coastal dune system. On its passage to the coast, groundwater dissolves magnesium from the calcareous matrix of calcareous dunes that surround and underlie the Coorong lakes (Schwebel, 1984). These Quaternary bioclastic dune grainstone units are covered by cdlcrete and are extensively karstified. Most meteoric water is quickly cycled back to the atmosphere through the soil moisture zone, or channeled down to add to the volume of phreatic water. There is very little rainwater runoff in the dune areas. As groundwater seeps to the surface in the near-coastal region it mixes with any surface water in the lake and evaporates to form dolomite. Areas subjacent to a lake strandline are the supply conduits of magnesian-rich groundwaters to the lake floor. In the more arid northern end of the study area, cdlcian-rich (type B) dolomites precipitate from these waters about the lake margin and more magnesian-rich (type A) dolomite assemblages precipitate in the more evaporitic lake center. When tne present distribution of dolomite containing lakes is compared with the areal extent of the coastal plain formed in the last 120,000 years, then Holocene dolomite-forming areas constitute less than 15% of the total area of the coastal plain.

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The lack of areally extensive, laterally continuous dolomite units across the coastal plain is one of the most important and fundamental observations when developing a model of sediment distribution for the Coorong dolomites (Fig. 2.26). Hydrology of the Coorong lakes is such that the dolomite-forming lakes are always surrounded and underlain by an extensive aquifer. Coorong dolomites cannot form without the surrounding dune sands to create the groundwater head and carry groundwater to the lake depression. The dune aquifer supplies, in solution, the Ca", Mg", and HCO; necessary to precipitate all the Coorong carbonates. Coorong dolomite forms by the evaporation of meteoric continental waters, not seawater, although dolomite in the basal unit may have formed by evaporative concentration of groundwaters locally influenced by seawater (Rosen et al., 1989; Warren, 1988). Vertical sequences typical of the Coorong lakes are independent of the mineral assemblage within a particular lake. Whatever the mineral assemblage, the dominant textural feature in the sediment column is lamination passing up section into a more massive unit. Coorong dolomites are penecontemporaneous carbonates not associated with preserved evaporites (Von der Borch and Lock, 1979). Yet, the best dolomite forming lake in the Coorong region (Milne Lake) is situated less than a kilometer from a previously undocumented lake filled with a laminated gypsum-aragonite sequence and capped by a massive hydromagnesite-aragonite unit (Halite Lake). Milne Lake is fed by continental groundwaters while Halite Lake is fed by marine groundwaters during the gypsum-depositing stage. The close juxtaposition of these mineralogies means that ancient Coorong counterparts may also show a close association of evaporative dolomite with more saline salts such as gypsum and halite. Given the abundance of associated magnesian-carbonates in the lacustrine sediments of the Coorong lakes, the lack of more widespread dolomite in the lake sediments may be in part an artifact of lack of time for full dolomite development in the Coorong region. But it is true to say that in any ancient Coorong-style sequence the penecontemporaneous dolomite phase would be deposited as a facies mosaic in depressions on top of a paleoaquifer. Probably the best ancient counterparts to the Coorong Lakes occur not in sea-marginal settings, but in the lacustrine realm. The carbonate mudflats of the Green River Formation, Utah, contain all the structures and textures present in the Coorong Lakes, as well as many of the same vertical transitions from laminated calcium carbonate sediments to massive dolomites. The Green River Formation was deposited in a large lake complex, part of an Eocene foreland basin that contained no marine waters. Green River carbonates consist of a series of stacked shoaling lacustrine cycles composed of laminated dolomites and limestones (Wolfbauer and Surdam, 1974).

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Individual units shoal upward and contain stromatolites, ripples, flat pebble breccias, oolites, mudcracks and extrusion tepees (personal observation and Ron Surdam personal communication, Nov. 198.5). Coorong-style models of evaporative dolomite cannot explain the bulk of ancient platform dolomites which formed as replacement minerals in thick beds that spread across laterally extensive areas of platform carbonates. Ancient Coorong equivalents in a sea-marginal setting are most likely to be relatively thin, laminated units which show facies mosaic.. in plan view. Dolomite is most likely to be found as a capstone in paleo- topographic lows (Fig. 2.26). A Coorong analog cannot be used to interpret extensive, thick platform dolomites, rather it should be used to explain primary dolomites found as laminated or massive units surrounded by an extensive paleoaquifer or fed by meteoric surface waters.

BRINE REFLUX DOLOMITIZATION

The sabkha and Coorong dolomitization models explain thin (1-5 meters), localized (10s to 100s of square kilometers), supratidal-associated "primary"dolomites that preserve characteristic depositional signatures. They do not explain the widespread dolomites found in close association with evaporitic sediments across vast areas of ancient platforms. Such dolomites formed by a process of reflux dolomitization. The brine reflux model explains thick (10s to 100s of meters), laterally extensive (thousands to millions of square kilometers), ancient "secondary" dolomites that replace restricted to open-marine platform limestones (Fig. 2.27). Reflux dolomites are unfailingly associated with platform evaporites or evaporite-dissolution breccias. For full discussion of this and other dolomitization models see Hardie (1987), Land (198.5), Machel and Mountjoy (1986), and Warren (1989). Brine reflux was first proposed by Adams and Rhodes (1960) to explain extensive lagoonal and reefal dolomites of Guadalupian age in the Permian Basin of West Texas. Reflux dolomites form when, "hypersaline brines eventually become heavy enough to displace the connate waters and slowly seep downward through the slightly permeable carbonates at the lagoon floor". When CaCO, and gypsum precipitated in hypersaline areas on the platform, the solution density was as much as 1.2 gm/cc. The Mg/Ca ratios of the remaining brines had risen, and these dense Mg-enriched solutions sank through the underlying platform sediments. As solutions seeped basinward through platform limestones they displaced subsurface pore fluids which were originally marine-derived waters with densities around 1.03 gm/cc. In this way large volumes of Mg-rich waters passed through previously deposited shelf limestones and converted them to dolomite.

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Adams and Rhodes set up their model to explain dolomites that formed beneath a lagoon that was behind some sort of sill or reef. The latter acted as a physical barrier to the free inflow of unevapordted seawater. On ancient epeiric platforms, such as the ramps on the Northwest Shelf of the Permian Basin, there may have been no actual physical barrier. The large lateral distances across a shelf/ramp that was covered by very shallow water may have sufficiently restricted seawater inflow to where saltern evaporites were precipitated on the seafloor (Bein and Land, 1982; Hovorka, 1987; Warren, 1989). Residual Mg-rich brines sank into the underlying platform carbonates. Minor falls in sea level furthered the reflux process by lowering the marine water table across the platform. This established new drainage pathways for Mg-enriched brines that moved out from evaporitic lagoons to sink through the more accessible platform limestones. In the 196Os, seepage reflux had gained a large number of proponents, and in 1965 Deffeyes, Lucia and Weyl reported modern Mg-rich brines in a marine spring-fed gypsum salina (Pekelmeer Lagoon) on Bonaire in the southern Caribbean. From the water chemistry of the Holocene sediments, they predicted the brines should be sinking and forming dolomite in pre-Holocene sediments beneath the lake. They obselved micritic dolomite in Holocene lagoon sediments beneath evaporite crusts, and on the opposite end of the island, well away from the present lagoon, they located a dolomite that was replacing a Pleistocene limestone. They inferred this dolomite had formed during an earlier episode of brine reflux. Lucia (1968) drilled beneath the Pekelmeer Lagoon and found no widespread dolomites in the underlying Pleistocene carbonates. He also found the sub-lake pore-waters were of normal marine salinity in areas where a brine was predicted by a seepage reflux model. A thin clayey ash layer of generally low permeability (a hydroseal)

basinward landward 4

Reflux of dense brines dolomitizes adjacent carbonates

platform evaporites platform carbonates

I 100 m I100's - 1,000's of km

Fig. 2.27. Hydrological model for brine reflux model.

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was found which separated sediments below with normal salinity from lagoonal sediments above with elevated salinity. Thus reflux through the lake floor was probably not occurring in the volumes postulated by Deffeyes et al., (1965). Seawater springs supplying seawater to the salina were local sites where the ash bed was broken or missing. The springs were the terminations of permeable flow pathways in the underlying Pleistocene (Murray, 1969). For most of the year seawater seeped into the springs by evaporative drawdown. During a short period in the late summer the springs were inactive and the pressure on the landward side was greater than on the seaward side. Murray concluded that return flow or reflux should occur at that time. The claim that Pekelmeer was forming extensive seepage reflux dolomites in the Pleistocene sediments beneath an evaporite lagoon was laid to rest when Sibley (1980) concluded the solutions responsible for extensive Piio/Pleistocene dolomitization on Bonaire, Netherland Antilles, were probably fresh to brackish waters. Thin dolomite layers in the Holocene salina sediments may well form by brine reflux but dolomites in a Plio/Pleistocene matrix do not. There are other islands in the tropics where brine reflux associated with hypersaline conditions is locally forming Holocene dolomites. Miiller and Teitz (1971) documented brine reflux dolomite replacing an earlier cement in skeletal grainstones from the shoreline in Fuerteventura in the Canary Islands. Kocurko (1979) found brine reflux dolomite a few meters above the high tide line, in the spray-zone pools of the shoreline of San Andres, Columbia. As discussed in the previous sections, thin, localized brine reflux dolomites can also be found about the edges of Ras Muhammed Pool and Solar Lake on the Southern Sinai Peninsula. All modern examples of brine reflux dolomites are small in scale and are found in coastal lakes close to a shoreline - they never approach the scale of brine reflux processes which dolomitized shelf carbonates in ancient successions. From a study of modern reflux dolomites, one would conclude brine reflux probably cannot operate on the regional scale for which it was first proposed (Land, 1983). However, as will be discussed in the next section and by Warren (1989), the present is not a good time to model ancient platform evaporites - to state it simply, there are no modern counterparts to ancient platform evaporites, hence there are no modern counterparts to brine reflux dolomites on such platforms. In modern marine settings, areas of evaporite precipitation are localized to coastal zones and the small fluid potentials caused solely by density differences cannot move brine very far through subadjacent sediments. Elevation head is even required to generate a sabkha dolomite which in itself is a small-scaleversion of reflux dolomite. In ancient evaporite settings, the lack of a shoreline-driven groundwater head was

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probably not a problem as large areas of the platform were covered by thick evaporites that supplied dense Mg-rich waters to underlying limestones. Reflux dolomitization is often easy to recognize, even in core, as the intensity of dolomitization decreases the further distance in depth away from the evaporitecarbonate contact (e.g. the lower San Andres cycles in the Levelland-Slaughter trend, Elliott and Warren, 1989).

ANCIENT PLATFORM EVAPORITES

Comparison of modern and ancient evaporite deposits underscores the limitations of a strict application of the Law of Uniformitarianism (Kendall, 1984; Warren, 1989). At different times in geological history, but not in the Quaternary, evaporites, like carbonates and siliciclastics, formed in depositional settings ranging from pedogenic to lacustrine on the continents, and from the supratidal to the deep subaqueous in the marine realm (Fig. 2.2). Process and texture can be studied in modern sea-marginal settings and applied directly to ancient sediments, but models so derived should not limit concepts of the scale or setting of ancient sequences. Ancient evaporite deposits typically possessed thicknesses and horizontal extents two to three orders of magnitude greater than those of modern evaporites (Fig. 2.28). The wider extent, greater depositional diversity, and greater thickness of evaporites deposited on ancient platforms can be explained by warmer climate worldwide, creating wider latitudinal belts suitable for evaporite deposition and preservation. For example, there is a clear diminution of evaporite extent from the Jurassic to the Cretaceous coinciding with a change from an arid to a more humid worldwide climate (Gordon, 1975; Meyerhoff, 1970; Hite and Anders, this volume). Nevertheless the extent of evap-

600 500 400 300 200 100 0 Time (million years B.P.)

Fig. 2.28. Vuriation in percent of evaporites on the world’splatf orm sediments across the Phanerozoic (ufter Ronov et al., 1980).

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orite deposition in the Cretaceous was considerably wider than today with the transition from evaporites to coal at around 55" latitude. Huge shallow epeiric seas, often formed large inland seaways in arid areas, so creating areas of extensive platform and intracratonic evaporites. This type of evaporite deposition was most common during times of long-term tectonic quiescence along with extensive carbonate deposition (Hay, Rosol and Sloan 11, 1988). Sediments were deposited as stacked platform cycles of relatively thin evaporite units (typically < 10 - 20 m) that were interlayered with open marine shelf carbonates and siliciclastics. Shoal-water platform strata of the Permian Basin of West Texas and New Mexico were deposited in this fashion as were the Arab cycles of the Middle East. Set up of the correct tectonic conditions for the formation of extensive basin-wide evaporite deposits with thicknesses exceeding 100 m (300ft) and extending across whole basins were formed during times of tectonic activity that generated restricted seawater inflow during incipient rifting, continental collision, and/or transtensional faulting. The last such saline giant, dominated by shallow subaqueous evaporite deposition, formed on the earth's surface during the late Miocene when the collision of Africa and Eurasia converted the Mediterranean Sea into a complex of shoal-water evaporite lakes and the Zagros Basin into an evaporite-filled foreland depression (Hsu et al., 1977).

ANCIENT EVAPORTTE BASINS - DEPOSITIONAL MODELS

Ancient marine-associated evaporites were deposited in two interrelated settings: 1) platform evaporites as stratiform units ( SO m thick) of deep water/shallow water evaporite deposits containing textural evidence of many different depositional settings including; mudflat, saltern, slope, and basin. The whole basin was evaporitic and showed a distinctive platform-slope-basin profile when evaporite deposition took place (Warren, 1989). Based on the presence or absence of a steeply dipping continental slope, ancient continental platforms can be subdivided into ramps and shelves (Fig. 2.29). Such ancient platforms were typically wider than today (1000s of km versus 10s of km) and shallow epeiric seas extended far into continental interiors (Hallam, 1981). The Jurassic Hith Anhydrite of the Arabian Gulf is such a unit. It is dominated by shoal-water evaporites

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,I No shoal

Shoal (zone Y)

I1 B I1 Shelf-edoe shoal

S H L C I I IL I Mid-shelf shoal

I D

Tvoicallv arealer than 1000 km . r Landward ==b

Fig. 2.29. Evaporite platform depositional styles - ramp and shelf. A) Gently deepening ramp. B) Ramp with mid-platform shoal (zone Y).C) Rimmed-shelf with topographically highest part located at the shelf break. D) Shelf with deep rim, topographically highest part is a shoal or island-crest facies located on the mid shelf. and over a vast area of more than one million square kilometers it forms a regional seal to much of the Middle East's oil (Murris, 1980). Intrashelf basins sometimesformed in epeiricseaways, they were prone to sluggish circulation and were often anoxic in their deeper water portions. In the Middle East this led to the deposition of prolific source rocks such as the Jurassic Diyab and Hanifa Formations (Ayres et al., 1982; Loutfi and El-Bishlawi, 1986). The intensity of oil migration upward from these source rocks, and its final column height in reservoirs in the Jurassic section was strongly influenced by the evaporite facies present in the seals to the overlying Arab cycles and the Hith Anhydrite (pers. obs.).

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PLATFORM EVAPORITES - DEPOSITIONAL SE'ITING

Ancient platform evaporites were deposited either on mudflats or subaqueously as shoal-water saltern deposits (Fig. 2.30). Evaporitic mudflat sediments were deposited as a facies mosaic of both dry and saline mudflats (sabkhas) separated by local brine pan depressions (salinas). These small depressions on the mud flat surface contained salinas, a few to tens of kilometers across, that deposited lenses of shallow-water evaporites as clean and pure evaporite units that were tens of centimeters to a few meters thick. This mosaic formed shoaling upward platform evaporite cycles 1 to 5 m thick. Individual cycles are not easy to correlate across the platform due to this limited continuity of individual sabkha and salina facies (Fig. 2.30A,C). However, the evaporite mudflat unit can extend laterally for hundreds or thousands of kilometers. Surface waters on the mudflat were ephemeral, or extremely shallow, and so sediments often contain evidence of subaerial exposure such as mudcracks, erosional, or karstic surfaces. Displacive evaporites in the forms of nodular, enterolithic, and diskoidal gypsum or anhydrite, and skeletal halite grew beneath the mudflat surface. Evaporite mineralogy varied according to the degree of aridity at the time of deposition, and the chemistry of inflow waters. Modern counterparts to evaporitic mudflats, that is counterparts in terms of size and scale of individual sabkha and salina facies within the mosaic, but not the overall extent of ancient mudflats, can be found in the sabkhas of the Arabian Gulf and the subaqueous coastal lakes of South and West Australia described in the preceding sections of this chapter. Evaporitic mudflat beds were often interlayered with eolian or fluvial sediments that, when sealed by evaporitic mudflats, form typical productive reservoir facies in Permian sediments of the southern United States. Examples include; the Minnelusa Formation in Wyoming (Achauer, 1982), the seal facies in the Basal Seven Rivers Formation in Yates field, West Texas (Spencer and Warren, 1986), the northern regions of the Palo Duro Basin during deposition of the Upper San Andres Formation (Hovorka, 1987), and the Lower Clear Fork Formation of the Texas Panhandle (Handford, 1981b; Presley and McGillis, 1982). Some authors call these evaporitic mudflats tidal, but most transitions from subaqueous to subaerial on the mudflat were not tidal but related to wind and storms with the associated wetting and drying episodes. Generally, the term tidal flat evaporite is ill-suited for evaporitic mudflat sediments in epeiric settings (usage after Hallam, 1981). In many areas water came across the mudflat as thin, predominantly wind-driven sheets of brine. Thus, the term strandline, rather than shoreline or tidal zone, is appro-

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Fig. 2.30. Depositional settings for ancient shoal-water evaporites: A) an rvaporitic mudflat ramp profile - a facies mosaic of sabkhas and salinas, 6) on a saltern ramp - a seaway f loored by subaqueous evaporites. C)an evaporitic mudflat shelf - a f acies mosaic of sabkhas and salinas, D) on a saltern shelf - a seaway f loored by subaqueous evaporites. priate in describing the water or brine-sheet edge. In contrast to the sabkha-dominated surface of an evaporitic mudflat, salterns were areas of widespread subaqueous evaporite deposition. Saltern is a new term to describe extensive shallow-subaqueous evaporite sediments that formed continuous depositional units across hundreds of kilometers (Warren, 1989). Saltern deposition occurred on platforms and in basin-wide settings (Fig. 2.30B,D). There is no known modern counterpart, although shallow-water evaporite textures can be observed in many

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small coastal lakes along the coasts of Australia and the Mediterranean. Saltern sediments were typically cyclic and made up of shoaling units 2 to SO m thick, sometimes capped by thin sabkha sequences. Cycles tended to be widespread and correlatable over distances of tens to hundreds of kilometers. Individual evaporite beds were dominated by subaqueous, often laminated and layered evaporite textures. Horizontal laminae were often crosscut or interrupted by coarse-grained to giant, vertically aligned crystals of calcium sulfate or halite. When such gypsum was buried and converted to anhydrite, it was sometimes preserved as vertically aligned nodular textures. Saltern evaporites in platform settings were deposited in lagoonal or intrashelf basin areas; either on rimmed carbonate shelves or on platform ramps (Fig. 2.29; carbonate terminology of Read, 1985; Warren 1989). Platform cycles often had open-marine or restricted marine shoal water carbonates at their base, passing up into saltern evaporites. Outside of the main sulfate depositing area, shallow water conditions about the edges of many platform evaporite basins were demonstrated by the Occurrence of micritic carbonates that contain domal stromatolites, tepees, mudcracks, bird or dinosaur footprints, fossil brine shrimp and ostrocodes (e.g. Calcare de Base in Sicily, Decima et al, 1988; Seven Rivers Fm. outcropping in Rocky Arroyo, near Carlsbad New Mexico, pers. obs.). In the subsurface and in wireline logs an ability to correlate clean evaporites over large areas and to define the mineralogical purity as more than 80% evaporite (on cross-plots) tends to separate saltern evaporitesfrom mudflat evaporites. Ancient saltern examples include the Cretaceous Ferry Lake Anhydrite in the northern Gulf of Mexico (Pittman, 198S), and the P1 and P2 evaporites of the San Andres Formation on the northwest shelf of the Permian Basin, and the Lower San Andres Formation of the Texas Panhandle (Elliott and Warren, 1989). Mudflats and salterns were often part of the same formation. Mudflats tended to dominate about the edges of the depositional basin or subaerial areas within saltern seaways. The rapid depositional rate of saltern evaporites, compared to slower rates of change for the water level in the basin, meant many subaqueous saltern evaporites started precipitating in deeper water (few to tens of meters) but quickly filled the basin with sediment up to the water surface. Any further deposition of sediment above this level was in a mudflat or sabkha setting. With the infill of the subaqueous basin, the depositional signature passed up section from subaqueous to sabkha. Many ancient mudflats held local brine ponds filled with subaqueous evaporites (Seven Rivers Fm, Spencer and Warren, 1986), and many large scale evaporite basins that filled with saltern salts were fringed, and later covered by sabkhas (San Andres Fm, Elliott and Warren, 1989). Changes in water level created alternations of sabkha and

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subaqueous signatures, a rise in water level changed deposition from sabkha to subaqueous, while a fall in brine level probably "sabkharized' the upper portion of many subaqueous sequences. Slight changes in climate, subsidence rate or relative sea level could rapidly change the platform from a saltern-dominated to a mudflat-dominated area of deposition.

SABKHAS AND EVAPORITIC MUDFLATS

Probably the first ancient evaporitic mudflat sediments to be recognized were sabkhas in the lower Purbeck, uppermost Jurassic of England (Shearman, 1966). Since that time many carbonate-sulfate sequences have been interpreted as ancient sabkhas, examples include the Visean of Eire (West et al., 1968), the Upper Devonian of western Canada (Fuller and Porter, 1969), the Middle Carboniferous of the Canadian Maritime provinces (Schenk, 1969), the Jurassic Arab cycles of the Arabian Gulf (Wood and Wolfe, 1969), some backreef strata of Guadalupian and Leonardian age in the Permian Basin (Kendall, 1969; Ramondetta, 1982) and the Smackover/Buckner Formations and equivalents in the Gulf of Mexico (Budd and Loucks, 1981). All of these cited examples are sabkhas dominated by a carbonate matrix which may pass laterally and up dip into redbeds. In the Lower Clear Fork Formation of the Palo Duro Basin in West Texas, there are Permian sabkha dominated mudflats with siliciclastic matrices and a large quantity of halite both as interbeds and sporadic displacive crystals (Handford, 1981b). The following discussion of ancient sulfate dominated evaporitic mudflats concentrates on those examples which are associated with economic hydrocarbon accumulations.

Ordovician Red River Formation, Williston Basin, U.S.A.

The Williston Basin is located on the southern edge of the Canadian shield, straddling the U.SA.-Canadian border and centered on Saskatchewan and North Dakota. The basin itself is circular, about 800 km across and filled with 4600 m of Phanerozoic sediments (Fig. 2.31). Ordovician carbonates were deposited in a shallow epeiric sea and show evidence of shoaling cycles containing environments ranging from shallow marine to saltern to supratidal (Roehl, 1967). Lithologies include bioclastic and argillaceous limestones, dolomites, and evaporites. Ruzyla and Friedman (1981, 1985) studied the Ordovician Red River Dolomite in the Cabin Creek field on the Cedar Creek anticline in Montana on the southwest side of the basin. They concluded the Red River in that area was deposited in an evaporitic mudflat setting dominated by sabkhas

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I I Cq SASKATCHEWAN I MANITOBA

Fig. 2.31. Location of Cedar Creek .4nticline and Cabin Creek Field in Williston Basin. Contour interval of basin sediment is 1000 feet (from Warren, 1989, based on Ruzyla and Friedman, 198.5).

(Fig. 2.32). The Red River Formation is 90-180 m thick, the lower part is a nonporous marine limestone, the upper part is 40-50 m thick and is composed of 3 cycles, each consisting of a limestone-dolomite couplet deposited in environments ranging from subtidal to supratidal. Subtidal lithologies are composed of mottled, dolomitic, biomicrite wackestones that contain fragments of echinoids, brachiopods, bryozoans, trilobites, molluscs, ostrocods, and gastropods with lesser amounts of solitary corals and stromatoporoids. These units are poorly bedded and heavily bioturbated, they were laid down in a low energy marine setting. The intertidal sediments are also dolomitic, biomicrite wackestones but they were deposited under higher energy conditions. The deposits contain graded beds, crossbeds, ripple marks, imbricated shells and conglomerates. They also contain allochems characteristic of an intertidal setting such as oncolites, ooids, intraclasts and peloids. Supratidal sediments in the Red River Formation include wavy-laminated micritic

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Fig. 2.32. Typical columnar section showing wireline signature, porosity and permeability plus idealized interpreted sequence. Note that the base of the dolomite is transitional down into limestone (,from Warren, 1989, based on Ruzyla and Friedman, 1985).

dolomites and organic-rich mudstones. They contain evidence of restricted and at times desiccated conditions in the form of desiccation cracks, erosion surfaces, intraclast breccias, solution collapse breccias and nodular anhydrite. The supratidal sediments and some intertidal sediments show evidence of supratidal leaching as they contain empty gypsum and anhydrite molds, corroded anhydrite crystals and solution collapse breccias. Notably, it is the micritic dolomites of the supratidal which best retain the original depositional fabrics. To the east of Cabin Creek field, and further into the Williston Basin, the evaporites change from nodular anhydrites with a dolomite matrix (evaporitic mudflats) into thicker bedded evaporites (salterns). Longman (1982) and Longman et al. (1982) studied the distribution of dolomite in the Red River Formation about the margins of the Red River Basin and proposed the bulk of the dolomite in cycles C and D were from brine reflux (Fig. 2.32). Ruzlya and Friedman (1985) and Roehl(l985) conclude that in cycles A-C in Cabin Creek field there is an early sabkha dolomite in the supratidal and a subsequent brine reflux stage related to sinking of brines from this mudflat. Similar conclusions were reached by Clement (1985) in his study of Pennel field. Generally, Red River Dolomites show at least two closely associated episodes of diagenesis; an early penecontemporaneous

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evaporitic mudflat stage which precipitated dolomite in supratidal sediments, and a subsequent, but related, brine reflux stage which dolomitized portions of the underlying intertidal and subtidal sediments. In the first stage, a sabkha dolomite is pervasively dolomitized in the supratidal portions of the sedimentation cycles. Such dolomites are finely crystalline and occur in the wavy-laminated or peloidal mudstones. The second stage, a coarser-grained brine reflux dolomite occurs in the intertidal and subtidal sediments immediately below the supratidal dolomites. The brine reflux dolomites selectively replace skeletal fragments and matrix and are a consequence of the sinking of dense brines which originated in the mudflat. The degree of reflux dolomitization decreases downward from the base of each supratidal unit into the underlying shelf carbonates. The decrease in dolomite downward is due to the decreasing influence of reflux brines; the further the carbonate is beneath the mother salt bed, the less the degree of reflux dolomitization. Because the dolomite- limestone units stacked up in the section, the timing of the reflux dolomitization must have been relatively early and preceded the deposition of the limestones of the next cycle. Derby and Kilpatrick (1985) recognized the above two dolomite types in Killdeer field, in addition they recognized what they interpreted as a mixing zone dolomite in Cycle D. Economic porosity in the Red River Dolomite is restricted to the dolomitized units. Dolomitization is controlled by the depositional setting, specifically the paleo-occurrence of hypersaline and fresher waters within the mudflat (Clement, 1985; Derby and Kilpatrick, 1985; Longman, 1982; Roehl, 1985; Ruzyla and Friedman, 1981, 1985). Some interparticle porosity is preserved between peloids, intraclasts, ooids, and skeletal fragments in the high energy intertidal sediments. Intercrystallineporosity is best developed in dolomitized mudstones and occurs as the spaces between individual dolomite rhombs. In Cabin Creek field, porosities up to 25% (13% av.-dolomite) and permeabilities up to 142 md. (7.9 md. av.-dolomite)occur in areas with good intercrystalline porosity. In addition, vuggy and moldic porosity occur throughout the Red River Dolomite wherever the sequence has been leached by undersaturated waters. Production is from the porous zones in the dolomitized units in the Red River, dolomites controlled by diagenesis in a sabkha-dominated mudflat. In Cabin Creek field the oil is 33" APT, with an average S0=30% and S,= 30%. The average depth to the economic zone for hydrocarbons across the field is 2750 m, with a 15.2 m net pay thickness, a productive area of 30.5 km2 (7620 acres) and an estimated 224 million barrels of oil in place with an ultimate recovery of 75 million barrels of oil.

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Upper Minnelusa Formation, Wyoming, U.S.A.

Achauer (1982) described interdunal sabkhas from the Upper Minnelusa Formation (Permian), in Rozet Field, of the Powder River Basin, Wyoming (Fig. 2.33). He studied the Upper Minnelusa from the Minnelusa-Opeche contact down to the top of the "Mel" marker (Fig. 2.34). This is a widespread marker easily recognized on gamma and sonic logs. When the Upper Minnelusa succession was deposited, the area was a broad marine platform adjacent to an extensive eolian dune field. Fluctuations in

=EOLIAN =DELTAIC

=FLUVIAL =SHELF, MARINE and EOLIAN

=HIGHLANDS LAGOON ..-.. EASTWRD LIMIT OF SANDSTONE UNITS >50' X EOLIAN SANDSTONE REPORTED IN UPPER MINNELUSA

Fig. 2.33. Paleoenvironmental reconstruction f or the Northern Rocky Mountain region during Upper Wolf campian time (afterAchauer, 1982). X shows documented locations where marine marginal and eolian deposits interf inger in the subsurface. This area was part of the Lusk embayment.

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relative sea level meant conditions on the platform alternated between open marine and eolian deposition. The Upper Minnelusa Formation in Rozet Field is composed of a number of shoaling carbonate cycles intercalated with laminated and cross-bedded quartzarenites (Figs. 2.33, 2.34). Quartz sandstone is the reservoir facies wherever it is not plugged by an hydrite. A complete carbonate cycle within the Upper Minnelusa is composed from top to bottom of: 1) supratidal facies of anhydrite with "chicken-wire''texture and/or a capping erosion surface with rip-up breccias. These evaporite units are no more than a meter thick, 2) intertidal facies of algal-laminated or layered dolomicrite which may be desiccated or slightly brecciated. Thicknesses rarely exceed a few tens of centimeters,

LITHCCM;Y I STRUCTURES

LITHOLOGY STRUCTURES 0Sondstone,quortz Inclined bedding Anhydrite "Chtckenuir$ Algal laminoled Dolomite. anhydritic w layered SUB = Subtidal SUP= Supratidal I NT = Intertidal Ss i Sandstone

Fig. 2.34. Gamma ray - sonic log profiles through the Upper Minnelusa Formation, the Opeche Shale and the Minnekahta Dolomite (after Achauer, 1982). Also shows an expanded graphic log and interpretation of the two depositional cycles documented on the wireline log. Logs are from the Arc0 Peter Svalina No. I5 well, NENE, see. 25, T50N - R70W, South Rozet Field Wyoming.

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and 3) subtidal facies of peloidal dolomicrite or dolomitized packstone containing fusulinid or crinoid fragments. Thicknesses range from 1-5 meters. Complete carbonate cycles are typically interrupted by exceptionally clean, well-sorted quartzarenites deposited in a variety of subjacent settings including eolian, beach, and very shallow subtidal (foreshore) environments (Figs. 2.33, 2.34). The vertical succession fits the classic Arabian Gulf sabkha model even in terms of lateral extent of facies and the thickness of the supratidal anhydrite facies which are consistently less than a meter thick. Achauer's detailed mapping of the subtidal dolomite units, especially the lowermost unit directly above the "Mel" marker, further refines the sabkha interpretation (Fig. 2.35). The dolomite forms a series of narrow northwest trending channels separated by quartzarenites. This trend is the erosional remnant of an eolian sabkha trend, with the sabkha dolomitesforming in the lows between a series of parallel crested desert dunes. The depositional setting was an area adjacent to an open marine sabkha. It was similar to a region south of Abu Dhabi where the modern coastal sabkha of the Arabian Gulf passes landward into interdunal continental sabkhas (Fig. 2.3A). The quartzarenite dune-sabkha association of Rozet field is part of a wider Minnelusa erg composed of coastal dune and shoreline sands about the Lusk embayment (Eschner and Kocurek, 1988). Erosional remnants of eolian sabkhas can also be seen in Rourke Gap field and other areas in the Powder River Basin. Preserved topography on the quartzarenites typically supply the closure, and the overlying evaporitic dolomites the seal. However, the preserved topography is not pristine, much of the complexity of the individual coastal dunes within the erg was lost during the transgression of the Lusk seaway. Only sand thickenings related to erg-scale waveforms have been preserved (Eschner and Kocurek, 1986, 1988).

Basal Seven Rivers Formation, Yates Field, West Texas, U.S.A.

Yates field is a shallow reservoir at 300-460 meters (1000-1500 ft) in depth. It producesfrom the San Andres Formation with additional production from the overlying Grayburg, Queen and Seven Rivers Formations. The field is located at the southernmost tip of the Central Basin Platform in Pecos County, Texas and covers approximately 105 km2 (41 sq. miles; Fig. 2.36A). As of 1986, some 1600 wells had been drilled in the field. At the level of the Permian section, the field sits on the structurally highest point of the Central Basin platform and is dominated by a broad asymmetric domal or anticlinal feature. The anticline appears to be formed by the intersection of two separate tectonic structures that create a horseshoe-shaped high through the field (Fig. 2.36B). The major axis of the two intersecting limbs is directed northwest-southeast and forms

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Fig. 2.35. Distribution of the Dolomite 1 member hung from the "Me1 Marker" (after Achauer, 1982). Note the elongate interdunal trends of the sabkha dolomites, arrow points to Arc0 Peter Svalina No. 1.5 well. the southeastern-most fold culminating in a series of folds in Pecos, Crane and Upton Counties (Gester and Hawley, 1929; Hennen and Metcalf, 1929). The minor axis in the southern part of the field is directed east-west and may be a structural continuation of the Fort Stcckton high (Gester and Hawley, 1929) or more likely is related to shelf margin buildups forming on the Sheffield Channel (D.H. Craig, oral communication). Yates has proven to be one of the most prolific reservoirs ever discovered in the United States. The high pressures, supplied by the combined gravity segregation, bottom water and gas cap drives, resulted in unusually high initial production rates. In 1929 the official proration test on the Mid-Kansas Oil & Gas Co. and Transcontinental Oil Co. #30 I. G. Yates "A"gave the well a daily potential production of 204,681 barrels of oil

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0 Irnl 0- lkm .CORES IN STUDY 0 LOGS FOR CROSS SECTIONS * CORE SHOWN IN FIGURES

Fig. 2.36. Yates Field: A) Locality map of Yates Field. B) Structure on top of the Seven Rivers Formation. Map also shows location of cross sections, Yates f ield unit boundary and well 1ocalitie.s. (from Spencer, 1987; Warren, 19906).

(Hennen and Metcalf, 1929), ranking it worldwide as one of the highest daily production rates from a single well. Evaporite deposition in Yates field area during basal Seven Rivers time occurred in two major subenvironments; sabkha and salina. Together they constituted the seal facies of an extensive evaporitic mudflat. Lithofacies analysis of the anhydrite seal (basal Seven Rivers Fm.) over the giant Yates field (4 billion barrels original oil in place) can be used to determine the quality of the underlying reservoir (San Andres Formation). The aerial distribution of the sabkha to salina facies across Yates Field is approximately 1:l. Ties of the seal facies to production statistics shows that in the Yates field unit, as of October 1986, only 85 million barrels of oil had been produced from under the thick salina accumulations. Yet, almost 985 million barrels have been produced from the area under the sabkhas (Spencer and Warren, 1986; Spencer, 1987; Warren, 1990b). Wirelines and core logs through evaporite portions of the seal show a blocky, thicker log character due to stacked cycles of relatively pure subaqueous salina evaporite. Thinner, spikier signatures indicate the thinner and matrix-rich sediments of a sabkha (Fig. 2.37A,C). Subaqueous signatures dominate in what were the tOpOgrdphicdlly lower parts of Yates Field while sabkha facies are found in the topographically higher parts of the basal Seven Rivers (Fig. 2.37). Massive salina anhydrite is found only over the central and western parts of the

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W Sabkhas show an irregular spiking response on the Gamma-ray, neutron and bulk density Massive anhydnte gives a very low gamma- CuNeS. reflectina the thin stacked sabkha ray reading and a low poros~tyresponse on sequences H,ih Gamma responses the neutron and a high bulk denslty A thln correspond to sandy dolomite and gypsum at the base of the masswe anhydrite Sandstones at the base of each cycle gives a false high POTOSI~Ydue to bound water Increasing anhydrite and decreasing sandstone give lower gamma ray readings and lower porosity Anhydrlte rehydratlng to gypsum may give false high porosity

Top cycle Sharp vpper content Supratidal t NodYlOr onhydrlte increoring upward 10 increarinp vprord m 3-5fl 5#111tone ond dolomite (1-15mI mofrl"

Intertidal

Dolomite and silty dolomite Subtidal Base cycle t Sharp cmtm__-

~ANHYDRITE =SILTY DOLOMITE BANHYDRITE LAMINATlONSicyonOboCter~Dl?~ BSILTYDOLOMITE a MARINE FOSSILS

Fig. 2.37. A) Wireline signature and core log of massive anhydrite (white) and dolomite (stippled). See fig. 2.36B for location of w~ll# 10. B) Wireline signature and core log of sabkha anhydrite. See fig. 2.36B for location of well # 5. C) ldealized cycle from the massive anhydrite. D) Idealized cycle from the sabkha anhydrite. (after Warren, 1990b). field. The massive anhydrite thins to zero roughly parallel to the platform margin, and landward of the San Andres platform edge. The isopach map (Fig. 2.38) of the massive anhydrite shows closed contours surrounding pod shaped areas of anhydrite up to 15 m (50 ft) thick. The pod shaped geometry, is the predicted style for massive gypsum deposited in a salina (Table 2.1). Two major areas of salina evaporites were normally present during early Seven

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ConlOYr lnlerrol i 10 fl

aANHYORITE THICKNESS 30fl

sw B NE A I5 16 17 A' Gamma- Densitv Gamma- hnrity _.

SEVEN RIVERS

3K-r-GRAYEURG/SAN ANDRES 100 ft

I 0.5mi

Fig. 2.38. A) lsopach map of massive anhydrite unit in base of Seven Rivers Formation. Map is based on core description and log response. The solid line shows the location of northwest trend and the dashed line the location of the northeast embayment (after Spencer, 1987; Spencer and Warren, unpublished; Warren, 19906). B) Crosssection A-A' showing change in log signature (see f ig. 2.368 for locality).

Rivers time; one on the northern side of the field and one on the southern side. The north and south ponds were flanked by smaller pod shaped thickenings and separated by a non-evaporite deposition trending north-northwest through the central part of the field. This linear feature is referred to as the northwest trend. Massive anhydrite was

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also absent in cores and logs in the northeast corner of the field. This area is oriented perpendicular to the shelf edge and intersects the northwest trend. Geometrically, this area formed an embayment devoid of evaporite and was named for the purposes of this study, the northeast embayment. This term was used without any genetic connotation. Because the M datum in the Seven Rivers approximates a level depositional surface (Spencer and Warren, 1986), the isopach maps of the unit from the M datum to the top of the San Andres approximated the depositional topography on the top of the respective formations. The map showed thick areas in the paleo-lows and thin areas across the ancient highs. The Queen Formation isopach map (Spencer, 1987), illustrates the same relationship of a local thickening of section under the evaporite ponds, and a thinning toward the field margins, as well as across the northwest trend and the northeast embayment. The units from the M datum to San Andres thin parallel to the platform margin, then thicken both into the basin and onto the platform. The Queen Formation thickens generally toward the basin but is thickest where the evaporite is thickest, and thins locally over the northwest trend and just shelfward of the platform margin. In summary, combining log and core data with maps and cross sections, it is possible to illustrate that the evaporite facies distribution in the basal part of the Seven Rivers Formation was controlled by paleotopography. Seven Rivers topography was in turn controlled by the depositional facies in the underlying San Andres Formation (Fig. 2.39; Spencer and Warren, 1987). The San Andres lagoon represented an area both originally topographically low and one that remained low because of dewatering and compaction of shales during early burial. The thick section of massive anhydrite accumulated in low areas situated above the dewatering San Andres lagoon. The lagoon facies was shale-rich and wells through the unit below the massive anhydrite typically show low hydrocarbon production. The massive evaporite that formed over the lagoon facies provided a lateral seal to the much more productive adjacent reservoirs located within karstified San Andres carbonate mounds. The basinward edge of the massive anhydrite is located parallel to the field margins and appears to have thinned onto the topographic high formed by the rigidly cemented San Andres carbonate mounds. San Andres carbonate mounds were karsted and cemented prior to the deposition of the Queen and Seven Rivers Formation (Fig. 2.39A). With burial, the mounds remained rigid bodies that did not compact as much as the adjacent lagoon sediments. Thus the mounds were topographic highs and were covered by sabkha facies during the deposition of the Basal Seven Rivers (Fig. 2.39B). These sabkhas formed an edge facies to the pond deposits of massive anhydrite. As Seven Rivers deposition con-

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Fig. 2.39. Block diagram illustrating interrelationship between evaporite f acies in Seven Rivers time and San Andres time. A) islands of San Andres f acies undergoing karstif ication, B) Salinas in Basal Seven Rivers Formation formed in lows located atop lagoonal sediments of San Andres Formation, C)Seven Rivers Formation subsequent to time shown in B. The area is covered by sabkha.s with only small localized ponds. tinued, the pond topography was infilled and the upper portion of the Seven Rivers was deposited as a sabkha-dominated evaporitic mudflat (Fig. 2.39C). Further discussion of these relationships can be found in Spencer (1987), Spencer and Warren (1986), and Warren, (1990b).

ANCIENT SALTERNS

Compared to sabkhas and evaporitic mudflats, ancient saltern evaporites in platform settings (Fig. 2.30B,D) are not well documented in the geologic literature. This is a consequence of many factors. Until the last decade there were no models of textures likely to be found in laterally extensive, subaqueous evaporites. Subaqueous evaporites tended to be relatively pure massive units (Warren and Kendall, 1985), easily picked on wireline logs but rarely cored in a drilling program (Nurmi, 1978). Subaqueous evaporite seals to stratigraphic traps tended to form best in off-structure positions; locations that are rarely cored (Warren, 1989). Relatively pure subaqueous evaporites, especially gypsum, were subject to early recrystallization, flowage and intense burial diagenesis which left only indistinct relics of the original texture. This was not so important in sabkha sequences where the matrix tended to be more rigid as it was composed of less mobile carbonates and siliciclastics

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(Warren and Kendall, 1985). Even when much of the original depositional texture was obliterated it has become increasingly apparent that saltern evaporite deposition played a major role in many ancient evaporite basins such as; the basin-wide Late Miocene (Messinian) evaporites of the Mediterranean (Cita, 1983; Schreiber and Decima, 1976), the platform evaporites of the Cretaceous Ferry Lake Anhydrite in the Gulf of Mexico (Loucks and Longman, 1982; Pittman 1985), the platform sulfates of the Permian Zechstein Basin (Taylor, 1984), and the ramp evaporites of the Palo Duro Basin (Handford, 1981b; Hovorka, 1987). Water or brine levels in platform evaporite basins were not controlled by the same factors that controlled sea level in many carbonate platforms. Ancient saltern basins were subject to evaporative drawdown but, after the initial drawdown stage, the saltern bed typically showed little internal evidence for long term desiccation. This is an indication of an evaporite bed‘s unique ability to hydroseal the basin. When a saltern basin was established, the water in the basin began to evaporate and the floor may have completely dried up in the early stages of the saltern’s history. This concentration stage deposited a widespread evaporite cement in the underlying carbonate or siliciclastic unit that was responsible for much of the evaporite plugging associated with many platform reservoirs. Once the upper part of the subadjacent unit was cemented, it formed a hydroseal or aquitard layer which greatly slowed the rate of loss of brine through the basin floor. Rather than quickly sinking through the saltern floor, further influxes of marine-derived brines started to pond and concentrate atop this aquitard. As the brines concentrated and precipitated evaporites, the air mass above the brine body quickly approached the dew point (Myers and Bonython, 1958). This further slowed the rate of water loss from the saltern. Similiar retardation effects can be seen today in working solar salt factories where in some cases huge fans are installed to encourage further evaporation during times of high humidity. Thus, ancient salterns trapped humid air over the water body and this meant that large areas of the saltern brine couldn’t concentrate much beyond the gypsum or halite precipitation stage, hence, the relative purity and monomineralic or bimineralic composition observed in many ancient saltern deposits. Kinsman (1976) discusses similar evaporation kinetics in a sabkha setting. Marine saltern units were deposited in brine-equilibrium conditions where ongoing evaporation kept the brine surface below sea level and the water body only tens of centimeters deep. The hydroseal, the high humidities in the overlying air mass, and the frequent resupply from storms and seepage, usually kept the saltern from completely drying up.

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Highly concentrated brines capable of precipitating bittern salts could form only in areas where dry arid winds blowing off nearby deserts had removed the water saturated air mass. Perhaps this occurred about the edges of some marine salterns, but there are no modern sea-marginal potash salinas or sabkhas, and ancient potash deposits seem to be associated with continental saltern and playa deposits. Textures commonly show early diagenetic replacement of less saline minerals (Lowenstein, 1982; Hardie, 1984; Kezao and Bowler, 1985). Saltern evaporites were deposited on a scale and in settings which do not occur today. For example, when the San Andres evaporites were deposited on the Palo Duro ramp in the Texas Panhandle, one could have probably walked seaward, across the platform, in ankle to waist-deep brine for 100 or more kilometers. Or, with enough drinking water and a few camels, one could have walked across the seafloor of the Mediterranean some 5.5 million years ago. At that time, the Mediterranean was a huge desert-saltern complex fed by marine and continental waters (Hsu et al., 1977). There are no modern equivalents in scale or depositional diversity to salterns. This section concentrates on saltern evaporites in platform settings; for a full discussion of salterns in basin-wide and other platform settings the interested reader is referred to Warren (1989).

Ferry Lake Anhydrite - Fairway Field, East Texas, U.S.A.

The Fairway field is located in Henderson and Anderson counties in east Texas. The Ferry Lake Anhydrite is a Lower Cretaceous unit sitting above the James Limestone which forms the major reservoir in the field. Loucks and Longman (1982) documented ancient subaqueous gypsum which occur as vertically or near vertically aligned anhydrite nodules ("gypsum ghosts"). The outline has been preserved as the original gypsum crystal surrounded by an impurity of either carbonate, organic matter or clay (Fig. 2.40). The outline is recognizable because the conversion from gypsum to anhydrite took place relatively shallow and slowly, but not so slowly that the gypsum became overpressured and thixotropic. The Ferry Lake Anhydrite is composed of alternating carbonate and evaporite units that were deposited in a broad, shallow subtidal hypersaline lagoon/saltern. The evaporites give a blocky log signature that can be correlated over distances of more than 150 km (Fig. 2.41; Pittman, 1985). The saltern was up to 260 km wide and separated from the sea by shelf-margin rudist banks and possibly a tectonic sill (Fig. 2.42). Evaporite units contain vertically to randomly aligned (elongate) mosaic nodular anhydrite originally deposited as subaqueous gypsum. Sabkha evaporites are rare in the

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Fig. 2.40. Vertically aligned anhydrite ("gypsumghosts") f rom the Ferry Lake Anhydrite, J. B. Kitchens No. I Well, depth 9332 ft. (courtesy of Bob Loucks).

Ferry Lake implying most of the anhydrite was originally deposited as gypsum on the floor of a subaqueous lagoon/saltern. Individual anhydrite units are up to 7 m thick and contain carbonate mudstone layers with anhydrite rip-up clasts which represent exposure or storm flooding of the saltern lagoon. Carbonate units that are interbedded with the anhydrites are predominantly mudstones and wackestones that were deposited in a restricted lagoonal setting. The fauna was primarily ostrocodes and mollusks. Several units contain Orbitulina and/or echinoids indicating normal marine waters were present at times in the depositional history of the Ferry Lake Anhydrite. Ferry Lake Anhydrite carbonates are not highly productive, but their Occurrence over deeper hydrocarbon reservoirs makes them a potential secondary objective in east Texas. The Ferry Lake Anhydrite produces hydrocarbons in the Fairway Field and in other areas in the northern Gulf Coast Basin. Fox (1963) reported production as of 1963 totalled 185,000 bbl of oil producing from thin oolitic stringers in the anhydrite section, Terriere (1963) reported one well had produced 40,OOO barrels of oil from the Ferry Lake. The stacked nature of the Ferry Lake Anhydrite cycles with each cycle passing up section from marine carbonates to massive saltern evaporites is a common transition in many other saltern settings including the San Andres Formation of West Texas and the Red River Dolomites of the Williston Basin. The upward shoaling, upward salting cycles probably reflect an inbuilt response of shallow water evaporites to episodic rises in

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Fig. 2.41. North-south stratigraphic cross section through the Ferry Lake Anhydrite from Howard County, Arkansas to Sabin? Parish, Louisiana. Datum is top of Ferry Lake Anhydrite - FI1 (af ter Pittman, 1985).

relative sea level. Subaqueous evaporite deposition in shallow brines is very rapid, much more rapid than carbonate deposition (Warren, 1989). Nevertheless, before saltern evaporites can form in lagoonal settings, some sort of barrier or inflow restriction must form in the carbonate zone on the seaward side. Otherwise, shelf/ramp waters are always too fresh for evaporite deposition and only open-marine to restricted-marine carbonates can accumulate. The base of an ideal saltern evaporite cycle consists of a thin transgressive unit that reflects an initial relative rise of sea level. This is overlain by open marine limestones deposited on the shelf/ramp. At this time of open marine limestone deposition one of two events occurs to instigate saltern deposition; either an offshore bar/reef builds up to the new relative sea level so that waters behind it become much more restricted, or a vast restricted seaway forms to limit seawater inflow causing brines to accumulate along the more landward portions of the seaway (zone Z of Shaw, 1964). At first a restricted marine carbonate is laid down, and then, when lagoon/saltern salinities become high enough, a rapid final infill by subaqueous evaporites occurs, perhaps capped by an evaporitic mudflat. The deposition of the saltern evaporites was often accompanied by extensive brine flushing and the precipitation of replacive reflux dolomites in the underlying platform carbonates.

Sun Andres Formation - Northwest Shelf, West Texas, U.S.A.

The San Andres Formation (Leonardian - Guadalupian) extends from Texas to

Data Center ,09126599985,[email protected], For Educational Uses SAN ANDRES FORMATION - U.S.A. 155

GENERALIZED DEPOSITIONAL EPPORITE STRUCTURES

Eolion Sill and Cloy Sabkho

\- EXPL4NATlON IT?~ Subaqueous Polmale Gypsun Intartidol Shoak with YM AQoI Mot 1c Echmmdi L C.biIalina 'Subtido1 May become ruboeriolly exposwed during long periods of eruporaflon

Fig. 2.42. Depositional model for the Ferry Lake Anhydrite in East Texas and a general depositional cycle as interpreted from the J. B. Kitchens No. I well. (after Loucks and Longman, 1982).

as far west as Arizona and Utah. Subsurface studies on the San Andres have focused on two localities: on the Permian Basin of southeast New Mexico and West Texas which includes the North, Northwest Shelves and Central Basin Platform; and on the Palo Duro Basin in the Texas Panhandle (Fig. 2.43). Most studies concentrated on the North and Northwest Shelves and the Central Basin Platform as these areas contain highly productive hydrocarbon trends (Silver and Todd, 1969; Meissner, 1972; Todd, 1976; Longacre, 1980). As of 1982,80% of the oil produced in the Northern Shelf was taken from reservoirs within the lower San Andres Formation (Ramondetta, 1982). The San Andres Formation is also the most productive unit in the whole of the Permian Basin of West Texas. Recently, the San Andres of the Palo Duro Basin was an area of interest as a potential radioactive waste repositoy (Bein and Land, 1982; Hovorka, 1987). Unlike the producing dolomitic intervals to the south, the San Andres evaporites in the Palo Duro

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Fig. 2.43. Majorgeologic featurer of the Permian Barin, west Texus anti New Mexico (in part after Silver and Todd, 1969).

Basin are composed primarily of salt, a relatively impermeable and self annealing mineral. Most of the carbonate intervals in the Palo Duro basin are nonproductive as they are plugged by anhydrite and halite. The more porous San Andres units of both the Central Basin Platform, and the North and Northwest Shelves are cyclic dolomites usually interpreted as brine reflux and sabkha dolomites formed in progradational or regressive limestones. The cyclical nature of the San Andres is typical of Permian deposits around the world and is often explained by the glacial-eustatic sea level changes induced by the Permian Ice Age. Chuber and Pusey (1972) subdivided the San Andres cycles into high energy shelf-edge and low energy back-shelf cycles. Almost all workers, with the exception of Hovorka (1987) and Elliott and Warren (1989), have interpreted the San Andres sediments as a sabkha sequence deposited under the influence of processes very similar to those occurring along the southern shore of the modern Arabian Gulf (Ramondetta, 1982). The Permian San Andres Formation of the Northwest Shelf was deposited on a Late Leonardian to Guadalupian platform, the northwest boundary of the deeper

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Midland and Delaware Basins (Fig. 2.43). It preceded the Goat Seep and the Capitan reefs and was deposited on a ramp with no striking break in slope from the continental shoreline out to the open marine of the Delaware and Midland Basins. The San Andres Formation on the Northwest Shelf is divided into two members each containing several shoaling evaporitic cycles previously interpreted as sabkha cycles (Longacre, 1980; Meissner, 1972; Silver and Todd, 1969; Todd, 1976; Ramondetta, 1982). The lower member can contain two well-defined hydrocarbon reservoirs (P1 & P2), they form the reservoir in the highly productive Levelland-Slaughter trend (Figs. 2.44A-C). The producing style is stratigraphic, with updip pinchouts from porous dolomite into nonporous dolomite and evaporites. Production is difficult to predict, most often it is slow but constant. Cycles in the lower San Andres member can be divided into 5 lithofacies arranged in a shoaling upward succession. Each cycle is 15-45 m (50-150ft) thick (Figs. 2.45A,B). The first stage is activated at the base of each cycle when shaley carbonate mudstones, indicative of low energy deepwater sediments, are deposited after the initial transgression. These are overlain by a biomicrite unit (10 m) containing brachiopods, bryozoa, crinoids, mollusks, cephalopods sometimes forming low relief bioherms (Fig. 2.46A). The unit was deposited on an open marine shelf, below wave base under low energy conditions. The second stage was dolomitization, by far the most variable unit, but the most important as it is the reservoir facies. It is classified into three subtypes (Figs. 2.46B,C): 1) oolitic peloid-foram grainstones forming as localized shoals, 2) wispy laminated to microstylolitized biomicrites, intramicrites and peloidal micrites (abundant unit); it contains a poorly preserved open marine fauna now present as molds sometimes filled with anhydrite, and 3) intertidal to supratidal mudstones with fenestra, pisolites, cryptal- galaminates and possible desiccation cracks (uncommon unit). The third stage was deposition of a massive anhydrite unit, previously interpreted as sabkha. In the western Levelland-Slaughter area it is meters thick and composed of nodular anhydrite interbedded with laminated anhydrite. In some cores there are vertically aligned anhydrite nodules which pass downward into laminated and bedded nodular mosaic anhydrite. The alignment, the lamination, the purity, the thickness of the anhydrite units and the fact it is laterally equivalent to well preserved pseudomorphs after subaqueous gypsum in lateral equivalents to the north implies this sulfate unit was laid down subaqueously in the western Levelland-Slaughter area (Fig. 2.46D; Warren and Kendall, 1985; Elliott, 1985; Elliott and Warren, 1989). The final stage was the capping halite unit which is not present in the Levelland- Slaughter trend, but occurs to the north as subaqueous and recrystallized halite (Fig.

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Fig. 2.44. A) Surfacr to subsurface correlation of the Sarz Andres Formation near Roswell, New Mexico B) North-south cross section through Chaves and Roosevelt Counties, New Mexico. The cross section shows the northward facies change from dolomite to anhydrite in each cycle and the resulting porosity pinchouts (after Gratton and LeMay, 1969). C) Location map showing the position of wells studied within the Levelland-Slaughter trend. Numbered circles indicate well locations; I) Depco #8 Rose Federal, 2) Ralph Nix #I Cherry, from Diablo Field. 3) Stevens #7 Citgo State, Twin Lakes Field. 4) Tenneco #18-4Santa Rita Moonshine, Twin Lakes Field. 5) Tenneco Oil #I Lee Carter, east of Milnesand Field. 6) Dalport Oil #I Tenneco Federal. The rectangle (marked fig.2b) shows outcrop area studied by Elliott and Warren, 1989.

Data Center ,09126599985,[email protected], For Educational Uses SAN ANDRES FORMATION - U.S.A. 159

CITIES SERVICE No I GOVERNMENT K tll I 24 T7S RZSE SAN ANDRES CYCLE

HALITE: Bedded and recrystallized. Bedded units show aligned chevrons. rafts. and hoppers Abundant evidence of penecanlemporaneous karst as halite pits and p~pes.

ANHYDRITE: Nodular mosaic and laminated sequences showing relict bedding and "ghosts" of original subaqueous gypsum crystals.

' FEET DOLOMITE. Helerogeneous UETERS dolomitized shoaling upward sequences composed of 3 vertical sequences. Sequences a) and c) are care, b is common 8) subtidal grainstones subtidal wackc and packstones b) rubtidal restricted mudstones subtidal near-normal marine mudstones c) Inlertidal/supratidaI sediments.

LIMESTONE: Biomicritic wackestaneslpackstanes deposited under open-restricted marine conditions.

MUDSTONE. Transgressive shaly I.s.

Fig. 2.45. A) Type log of the carbonate-dominated San Andres Formation from the Levelland-Slaughter trend, Chaves County, New Mexico (after Gratton and LeMay, 1969). Depth is in f eet. Gamma-ray/Density is an excellent combinationf or diff erentiat- ing major evaporite lithofacies; note the clean blocky log and thickness of the PI and P2 evaporites. Note also the change in the log signature between the lower and upper members; this reflects an up-section change from saltern-dominated evaporites in the lower member to evaporitic mudflat sequences in the upper member. B) Geology of an idealized PI or P2 cycle in the San Andres Formation in the western Levelland-Slaughter area, New Mexico. Shows increasing up-section restriction and finally evaporite deposition.

3.7B, this volume; Hovorka, 1987). In outcrop to the west of Roswell, the evaporites occur as dissolution breccias which can be used as surface mapping datums (Fig. 2.44A7C;Elliott, 1985). In fact it is possible to correlate the 5 cycles of the lower member of the San Andres from outcrop down into the subsurface near Roswell and then in the subsurface all the way into the Palo Duro Basin (Elliott and Warren, 1989). In outcrop, the lateral facies changes can

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D Fig. 2.46. San Andres Lithqfacies. A) Open marine facie.7 of' the basal limestone containing bryozoa. crinoids, brachiopods; Hondo West outcrop. San Andres Formation, Northwestern Shelf. B) Dolomite f acies with pel-oomoldic porosity, photomicrograph from Depco #X Rosa Federal 1021 ft. C) Dolomite-pellet grainstone ,facie.s; shows the consistent size of the pellets. Pelmoldic porosity is typical of this f aries although it is of'tpn occluded by an anhydrite cement. Bar scale is 0.6 mm (Depco #H Rose Federal; depth 10.54 ft); D) Vertically aligned subaqueous gypsum now replaced by halite, Sari Andres Formation, Paln Duro Basin. be seen in the carbonates of each cycle; the carbonate intervals were deposited as a facies mosaic. In the subsurface the carbonates are overlain with a sharp contact by the laterally extensive subaqueous evaporites; sulfate units and the Pi marker are lost by dissolution in the vicinity of the Pecos River. Continuity of evaporites in the subsurface has led Elliott and Warren (1989) to propose a revised model for San Andres deposition; a model which sees sabkhas as a relatively minor facies compared to the salterns that extend northward for hundreds of

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kilometers from the western Levelland-Slaughter trend. Each complete cycle ( P1, P2, etc.) in the San Andres records a lowering of relative sea level (Fig. 2.47). Time I: highest stand of relative sea level stand - initially low energy deposition or nondeposition, then the development of normal marine with open circulation, and the occasional bioherm. This period was typified by "starved basin" or mudstone deposition. Time 2: intermediate stand of relative sea level - skeletal banks and oolite shoals developed where wave base impinged on the sediment water interface. The shoals, and the great width of the platform, along with high evaporation rates, eventually proved sufficient restriction to develop a mosaic of carbonate shoals and islands separated by a pellet grainstone lagoon from the mainland which lay to the north. The high energy shoals occur to the north of the study area. Time 3: lowest stand of relative sea level - accompanied by climatic change to more arid conditions. An extensive gypsum depositing saltern formed, still in partial contact with the more open marine platform to the south. There were few, if any, sabkhas or evaporitic mudflats present during the early stages of saltern deposition. Extremely high evaporation rates and very shallow water sometimes lead to local desiccation about the saltern edge and over topographic highs in the Northwestern Shelf. Highs were often located above buried shelf-margins and other well-cemented shoals. This was a time of extensive brine reflux, resulting in dolomitization and evaporite plugging of the underlying sediments. Karst and local evaporitic mudflats developed in such areas, especially during the latter stages of each saltern stage when the basin filled with evaporites to the ambient brine level. In the northern parts of the Palo Duro Basin this led to the deposition of chevron halite crosscut by syndepositional karst (Hovorka, 1987; Hovorka et al., 1985). San Andres exploration takes place in a very mature province where production is typically enhanced by secondary (water-flooding) and tertiary (CO, injection, polymers) recovery techniques. Hence an understanding of porosity controls and, if possible, predictive models would enhance future exploitation of the basin. San Andres reservoirs are characteristically slow but prolific producers of oil and gas, and are commonly found at shallow depths of less than lo00 meters. The source for the hydrocarbons in the San Andres reservoirs are perhaps the underlying basinal shales of Wolfcampian age (Ramondetta, 1982). Migration probably occurred via vertical fractures upward into the reservoirs of the San Andres. Permeabilities in the reservoirs are typically low and reservoir quality hard to predict due to patchy evaporite plugging. Traps are a result of several factors including; stratigraphy, diagenesis, structure, and hydrodynamics. The most important trapping

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NORTH LANDWARD

TIME I

SOUTH LOtilude of Stu BASINWARD are0 Restricted Lagoon

TIME 2

,Br,"e PO" NORTH LANDWARD

TIME 3

Progroding Sobkho SOUTH BASINWARD

Fig. 2.47. Deposition sequence of a .single cycle during deposition of PI and P2 cycles in the lower San Andres Formation in the west Levelland-Slaughter trend (after Elliott and Warren, 1989). This is a regional reconstruction with the block diagrams oriented north - south. Time I) Highest stand of relative sea levcl when seawater covered the entire platf orm; energy levels in the study area were low. Time 2) Intermediate stand of relative sea level. Time 3) Lowest stand of relative sea level.

mechanisms are usually stratigraphic,although combination traps arecommonplace.The most important porosity generating mechanisms are dolomitization (intercrystalline) by brine reflux and the partial retention of original intergranular porosity. Oil column heights greater than structural closure confirm the importance of stratigraphic traps within many San Andres fields, even in areas dominated by combination traps. The most common trap for San Andres reservoirs is a regional updip porosity

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pinchout as porous dolomite passing up dip into evaporite-plugged, low-porosity dolomites with anhydrite caps (Fig. 2.44B). Such pinchouts are intimately related to the cyclical nature of San Andres sedimentation and the progradational basin architecture during San Andres time. Dolomitized porous reservoir rocks interfinger with tight updip evaporitic seals to create an efficient reservoir trap mechanism. The reservoir units are usually subtidal and intertidal dolomitized limestones with good intercrystalline porosity while the nonporous evaporitic seal is composed of similar carbonates that have been completely plugged by anhydrite and are overlain by massive anhydrite (Fig. 2.45). Previous authors interpreted the anhydrite cap as Arabian Gulf-type sabkhas, and the reservoir facies as paleoshoreline sands deposited in the same manner that porous ooid shoals follow the shoreline trends in Abu Dhabi (Dunlap, 1967; Gratton and LeMay, 1969; Zaaza, 1978; Pitt and Scott, 1982; Ramondetta, 1982). In the past, the Levelland-Slaughtertrend on the Northwestern Shelf has been used as the type example of sabkha-associated traps (Shinn, 1983). The model of Elliott and Warren (1989) offers a different explanation and exploration strategy for the San Andres reservoirs on the Northwestern Shelf of the Permian Basin. Many dolomitic reservoirs in the western Levelland-Slaughtertrend (Fig. 2.44C) did not form by the updip pinchout of marine grainstones and mudstones into stacked sabkhas. Rather, the updip carbonate portion of evaporite-dominated San Andres cycles across the northern Northwestern shelf and in the Palo Duro Basin underwent pervasive pore occlusion during intense episodes of subaqueous evaporite deposition. Brines flushed downward and seaward, plugging and dolomitizing the carbonates along the way. The main reservoir trend lies downdip of this evaporite plugged zone and defines the transition from up dip, evaporite plugged, partially dolomitized saltern sections into downdip compacted and pervasivelycemented open-marine limestones. The reservoir trend in the western Levelland-Slaughter trend formed not as a paleoshoreline but as a diagenetic front (Fig. 2.47). The front defines a transitional dolomitized zone, a subtidal-intertidal platform sequence, immediately seaward of the saltern-dominated zone. It is not a sabkha shoreline in the sense of the classic Abu Dhabi sabkha model. Dolomitized carbonates in the reservoir trend are a heterogeneous mosaic of pelletoidal open-marine to restricted-marine platform facies; they are not dominated by the oolitic grainstones and reefal carbonates predicted by the Arabian Gulf sabkha model. The reservoir trend is a transition zone that experienced brine reflux and evaporite plugging during relatively short periods of lowered relative sea level (Fig. 2.47C). Carbonates were dolomitized but not completely plugged by evaporites. When the relative sea level was higher, the reservoir trend was covered by restricted marine

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waters that were not conducive to brine reflux and evaporite plugging. The reservoir prone zone in turn passed downdip into uneconomic open-marine subtidal limestones lacking economic levels of porosity. Todd (1976), argued that fresh water diagenesis in the Levelland-Slaughter trend effected subaerially exposed highs during low sea level stands and created secondary porosity as vugs and solution channels. An abundance of fresh water is difficult to envisage in an area of widespread evaporite deposition. Bein and Land (1982) argued that identical textures identified by Todd (1976) formed in the San Andres Formation on the Northern Shelf during brine reflux. These vugs and channels formed totally independent of the effects of fresh water. Tertiary age hydrodynamics within the Permian Basin may have influenced trapping configurations. Eastward flowing groundwater has created eastward dipping oil-water contacts which further complicate an already complex group of traps (Ramondetta, 1982). Groundwater flushing has also degraded oil quality in many shallow reservoirs. Preliminary studies, using mercury injection data, imply some of these tilted oil-water contacts in the Northwestern Shelf may not be a result of flowing groundwaters, rather they reflect differences in oil column heights that can be related to the different pore throat geometries of various depositional and diagenetic facies.

COMPARISONS OF SABKHAS AND SALINAS

In modern marine-associated evaporites, the subaerial setting is called a coastal sabkha and the subaqueous setting a salina. Both settings deposit shoaling-upward sequences. In a sabkha, most of the evaporites are deposited as intrastratal displacive and replacive crystals within carbonate and siliciclastic matrices in the uppermost phreatic and capillary zones (Fig. 2.48, Table 2.1). Thus, marine-sabkha sequences are shoaling peritidal cycles with nodular, often displacive evaporites, in their upper part. Each cycle of deposition is capped by an erosion surface, the result of wind erosion after displacive crystal growth has pushed up the surface of the wet mudflat into the dry conditions of the vadose zone. The evaporitic supratidal unit capping a sabkha cycle is usually less than a meter thick and matrix dominated; the thickness of the supratidal unit is inherently limited by the sabkha hydrology. Prograding marine supplied sabkha cycles often fine upward (Abu Dhabi sabkha), whilst coastal sabkhas with matrices supplied from prograding continental sediments from the hinterland tend to coarsen upward (Gulf of Suez sabkha). Grainstone and sandy reservoirs in marine sabkhas are composed of subtidal and intertidal sand bodies or wadi and eolian dune sands; all sealed and cemented by sabkha

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lncreosing salinity __t SULFATE SALT

Suboeriol a&?, ond 000 -ocustrini r__ Continental - Sobkho lisplocive and Halite COASTAL Vodose POND and (0s in Shallow su boaueouc Phreotic

- lncreosing turbulence ---c Shallow zs!zZ=- SgzS Lominoted Cross-lominoted om Chevron-halite beds rippled

Shelf rystollinern&! with Hololites Corbonote

.o~inoecomposed d hopper' rafts

-siS= DeeD - mm Lominoted

Fig. 2.48. Summary of the various evaporite textureS f or calcium sulf ate and halite which are indicators of particular physical environments of deposition (after Schreiber and Decima, 1976, 1978; Kendall, 1984). evaporites. Dolomite is currently forming in the intertidal and lower supratidal sediments of the Abu Dhabi and Kuwaiti sabkhas. At this stage the controls on sabkha dolomite distribution are not well understood. Salinas are small-scale bodies of hypersaline water or their associated subaqueous evaporites. By this definition salina evaporite units are less than 10-15 meters thick and have lateral extents less than tens of kilometers. The textures found in modern salina deposits are used to model ancient subaqueous evaporite textures. In a subaqueous calcium sulfate succession, the stability of the water column and rate of change of salinity are paramount in controlling mineralogy and textures. When brine conditions are stable, large bottom-nucleated crystals are deposited as growth aligned crystal aggregates. Under more changeable conditions of salinity and energy level, mechanically reworked accumulations of sulfate evaporites are deposited. Salina evaporite beds

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deposited in the larger modern coastal lakes are at least 1-2 meters thick, and so are thicker than modern supratidal sabkha beds which are typically less than 0.5-1.0 meter thick. In the nearcoastal continental salinas of the Coorong region, thick Holocene dolomites are forming from the evaporation of continental groundwaters. Intrinsic textural properties of sabkhas and salinas are controlled by the hydrology of their respective depositional settings; this leads to differing levels of purity and hence relatively distinct wireline properties in ancient counterparts. Sabkha evaporites (calcium sulfate and halite) tend to grow as displacive crystals in the capillary zones of mudflats adjacent to a water body. Individual supratidal cycles are thin matrix-dominated beds less than a meter thick; salina cycles are thicker evaporite-dominated beds that can be meters thick. The thin, matrix-rich nature of the sabkha evaporites often gives a high frequency, spiky signature to a wireline log taken in a well that intersected an ancient counterpart. Lithological crossplots from a wireline through an ancient sabkha dominated sequence show high percentages of matrix minerals such as dolomite and silts compared to anhydrite. Aligned giant crystals and their pseudomorphs (1 to 100 cm long), as well as mechanically reworked evaporites, are most often subaqueous deposits. In wells drilled through an ancient salina (saltern) counterpart, the greater purity and thickness of subaqueous evaporites gives the wireline signature of a subaqueous evaporite a lower frequency and blocky appearance compared to the spiky sabkha signature. Lithological crossplots of salinas tend to show higher percentages of evaporite compared to sabkha sequences. Ancient evaporites were deposited in settings ranging from continental to basinal. Ancient platforms were covered occasionally by evaporite-depositing marine derived brines that were capable of laying down sediments either in saltern or evaporitic mudflat settings. Saltern deposits were subaqueous evaporite units up to tens of meters thick that were deposited on the floors of evaporite seaways. They are pure evaporite units with log signatures that can be correlated for lateral distances measured in tens to hundreds of kilometers. Evaporitic mudflats were laterally extensive deposits composed of a mosaic of salina and sabkha evaporites. The mosaic style of deposition means the wireline signatures are spiky and individual spikes are hard to follow across more than a few kilometers. The relative lateral continuity of these two settings allows their definition in the subsurface with only minimal core control. On many saltern covered platforms each cycle is characterized by an initial episode of transgressive starved-basin deposition followed by marine carbonates passing upward into saltern evaporites. Brines sinking through the underlying carbonates often precipitate widespread reflux dolomites and pore plugging evaporites.

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DIAGENESIS OF SULFATE-DOMINATED EVAPORITES

Because textures found in evaporite cores are vicarious indicators of the original depositional conditions, this short subsection to the chapter will explain some of the complexities in interpreting textures in ancient sulfate-dominated evaporites. During late, early, and even syndepositional diagenesis, primary crystal textures of many ancient evaporites were altered by recrystallization, displacement, and replacement growth (Warren, 1989). Pseudomorphs and reaction rims to evaporite crystals are common, as are relict mineral cores encased in the mineral that replaces it. Documented examples include pseudomorphs of anhydrite after subaqueous gypsum, reaction rims of polyhalite around earlier gypsum or anhydrite, relics of carnallite in secondary sylvite and pseudomorphs of sylvite plus kiesel-ite after earlier langbeinite (Hardie, 1984). The transformation of depositional texture is typically a consequence of diagenesis associated with burial. Gypsum precipitated at the suiface is converted to anhydrite with burial and back to gypsum with re-emergence. Other evaporite salts such as halite and the bittern salts often recrystallize or transform to another mineral during dewatering associated with shallow or deep burial. Evaporite textures are also altered during the deformation and recrystallization associated with rheotropic flow in the subsurface. Buried salt can flow and fold into intricate salt structures on scales ranging from millimeters to kilometers. As buried gypsum converts to anhydrite, it can develop enterolithic folds similar to those formed in a sabkha. Some saline minerals flow more readily than others; in bedded evaporites in the Realmonte Mine in Miocene potash salts of Sicily, a layer composed of carnallite and halite is isoclinally folded and yet sits between two relatively undeformed laminated anhydrite beds (pers. obs). The following section concentrates on the calcium sulfate evaporites, for a discussion of the diagenesis of other evaporite salts the interested reader is referred to Hardie, 1984; Holser, 1979; Schreiber, 1988; Warren, 1989.

Diagenetic CaSO, - Displacive vs Replacive Growth

Primary evaporites are those deposits which were precipitated from a standing body of brine with no further alteration to the crystals. Outside a few Neogene evaporites there are no "primary evaporites" on the earth. Syndepositional or early burial diagenesis altered the evaporite rock fabric so that the original depositional texture of the matrix is modified in the weeks, months, or millennia after the crystal was first precipitated. In evaporite studies we often talk about displacive verses replacive growth of

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evaporites. Displacement is a process by which something is moved or displaced. Displacive crystal growth is a term used to describe the situation where a growing crystal pushes aside the surrounding matrix. This is seen by most workers to be the dominant mode of growth of the anhydrite nodules in the Arabian Gulf sabkhas. In metamorphic studies replacement is defined as "the process of practically simultaneous capillary solution and deposition, by which a new mineral of partly or wholly differing chemical composition may grow in the body of an old mineral or aggregate" (AGI glossary). When a crystal grows replacively, often there are small remnants of the original material left within the new crystal. Usually the term replacive growth is used to describe evaporite pseudomorphs in ancient strata such as anhydrite or calcite replacing gypsum crystals. If the original lozenge or prism shape of the gypsum crystal is presewed, it is called a pseudomorph. The processes of replacement and displacement are not easily separated in many evaporite units. Consider the situation of a gypsum crystal growing in a clayey matrix such as one finds in the intertidal and subtidal sediments beneath the Arabian Gulf sabkhas. Kastner (1970) modeled this situation where she found that gypsum growing rapidly in a matrix will poikilitically incorporate some of the matrix. Yet when the same crystal grows a little slower it does not incorporate any matrix but pushes the matrix aside. There is yet another control on evaporite deposition, namely the permeability of the matrix. A gypsum crystal growing in a sand often envelops the sand grains as it grows by passive precipitation in the pore space; this creates a poikilotopic texture. Huge crystals, up to a meter long, grow by this process in strandline sands in many continental playas near Kalgoorlie in southwestern Australia. Yet when giant gypsum crystals of the same size grow in a clay matrix in the same playa, they typically grow by displacing matrix with only an occasional rapid burst of poikilitic growth. How should their ancient counterparts, anhydrite nodules containing varying proportions of matrix, be interpreted? An initial reaction would probably be to call them replacive. Yet they may have grown by combinations of passive precipitation, displacement, and replacement! Consider another common situation in many ancient marine carbonates which contain anhydrite nodules. Such nodules often have very pure anhydrite centers but the edges of the anhydrite nodules have enclosed or dissolved the original carbonate matrix. How do we interpret these nodules - are they passively precipitated in the pore space or replacive or displacive? As a general rule of thumb, it seems gypsum is only rarely replacive, and is usually a passive precipitate in porous media and displacive in fine grained media. It is extremely unusual to find gypsum replacing carbonates and feldspars. On the other hand anhydrite readily replaces many carbonate and siliceous

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phases. This may reflect the truism that most anhydrite is formed in the subsurface at elevated temperatures, pressures and salinities which are the same chemical conditions where the dissolution of the more stable mineral phases (calcite, feldspar, etc.) is favored by the same conditions that are capable of precipitating anhydrite.

EVAPORITE DIAGENESIS - the importance of hydrology

Evaporite deposits are created by chemical changes in a concentrated aqueous solution, either at the surface or in the subsurface - this solution always precipitates and surrounds the evaporite deposit. In the early stages of the history of the evaporite deposit the changes in the solution are brought about by evaporative concentration and meteoric dilution, both at the surface and in the shallow subsurface. Later in the history of the evaporite unit, the changes in the chemistry of the pore fluids and the associated stability fields of the various evaporite minerals are a response to changes in temperature, pressure, and the regional basin hydrology. In every case, an evaporite unit sampled in the subsurface reflects the geohydrological history of the basin. Its textures and its mineralogy are no more than a response to hydrological changes. Thus to understand an evaporite deposit needs an understanding of basin hydrology both during and after deposition.

Syndepositional Hydrology of Holocene Basins

A sulfate-depositing evaporite basin is connected to its main water source, the ocean, in four ways (Fig. 2.14; Warren, 1986a); 1) a direct surface connection to the Ocean (Bocana de Virila - Morris and Dickey, 1957), 2) a surface connection during storm flood or spring tide (sabkhat - Abu Dhabi; Gulf of California - Castens Seidell, 1984),3) a subsurface groundwater connection (Lake MacLeod, Marion Lake Complex, the Coorong Lakes, Solar Lake) or 4) a subsurface connection for much of the year, but also a surface connection during occasional flooding events (Los Roques, Venezuela; Goto Meer and other ponds in the Netherlands Antilles - Sonnenfeld and Hudec, 1980). Only rarely is one of the supplying mechanisms the sole supplier of seawater to the basin, but one of the four usually dominates the hydrological history. Basins with a perennial connection to the sea either as surface channels or as spring fed inflow, tend to contain perennial brine ponds in the depressed areas of the basin (South Australian and West Australian salinas). Sediments in these areas of ponded inflow are algal-laminated carbonates or subaqueous, bottom-nucleated, gypsum. In most basins away from the strandline zone there is little if any early

Data Center ,09126599985,[email protected], For Educational Uses 170 SEA-MARGINAL AND PLATFORM CALCIUM SULFATE syndepositional diagenesis in the evaporites after they were deposited. This is probably due to the one way flow of saline solutions and the near constant salinities of the bottom waters of the perennially brine covered areas of the system. The associated downward reflux of these near constant density brines maintains near constant salinities in the pore brines directly below the sediment surface. Once in the subsurface these brines cannot be further evaporated and so new minerals tend not to form. For example, the basal gypsum of a 10 m thick gypsum sequence in Lake MacDonnell, a large South Australian salina at the head of the Great Australian Bight, still retains its original depositional signature and porosity. There are no reflux dolomites to be found in the Pleistocene carbonates beneath the lake although the occasional pore is filled with gypsum (Warren, 1980 and pers. obs. 1978, during drilling for gypsum reserves). There has been no further intrasediment deposition of gypsum in the 4,000 - 6,000 years after it was deposited. Today, none of the worlds Holocene subaqueous gypsum beds have sufficient thicknesses, nor lateral extents, nor have they subsided deep enough, to fully test a hypothesis of reflux diagenesis. Modern sea-margin basins with ephemeral flood events tend to dry up and refill many times in their history. Strandlines migrate back and forth across large distances in the basin as the water table rises and falls within the sediments. Near-surface sediments pass many times through subaqueous to capillary to vadose environments and depositional signatures reflect thisvacillating hydrology. Periodicsheetflood and deflation are common in such areas (Fig. 2.49A,B). Surface and groundwater flow patterns are often reversed and a whole suite of early diagenetic minerals can form within weeks, but more often months or years, after a flooding event. Because of these hydrological instabilities, evaporative mudflat sediments are characterized by a lack of long-term pore water stability. Sediments undergo penecontemporaneous diagenetic mineral changes such as; formation of penecontemporaneous dolomite, the back and forth transformation of gypsum and anhydrite, the growth of fractionated evaporite accumulations, the replacement of aragonite by gypsum and vice versa, and the formation of pseudomorphs (see appropriate earlier sections in this chapter). The possible exchange pathways of the ions in areas of oscillating strandlines and sabkhas provide some idea of the complexity of possible hydrological/mineral interactions (Fig. 2.50). Of course the two typesof settingsjust described are end members of a hydrologic spectrum. Areas which are usually subaqueous in a sea-margin setting can occasionally dry up, and areas which are subject to periodic wetting or drying can be subject to occasional longer periods of subaqueous or subaerial deposition. As a general rule of thumb, recent gypsum deposited in a continuously subaqueous environment is less

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Fig. 2.49. Sabkha el Melah de Zarzis, Tunisia. A) Pavement of diageneticgypsum crystals formed in the capillary/phreatic zone and then later left as a lag deposit when the surrounding f iner grained matrix was eroded by marine sheetflood and eolian deflation. This is a gypsrudite formed by a feedback between capillary evaporation (a diagenetic process) and sheetf lood and deflation (depositional processes). B) Detail of puvement in A. Both photos courtesy J. P. Perthuisot.

subject to pervasive diagenesis than gypsum laid down in an ephemeral surface water setting. Although not seen in the 6,000 year time frame available in modern salina sediments, diagenetic changes during early burial do effect subaqueous gypsum. This appears to happen once the subaqueous part of a depositional system is covered by waters capable of precipitating a different mineral suite, or when the gypsum unit is buried a little deeper and taken out of this relatively stable salina hydrology. This is evidenced in beds of Permian San Andres and Clear Fork Formations where anhydrite and halite textures mimic the original gypsum, and the underlying carbonates are pervasively altered by brine reflux dolomites (Fig. 2.51A). In these Permian evaporites the transformation of gypsum to anhydrite may have occurred at depths of only 1-2 meters (Hovorka, 1988). The transition from limestone-depositing to gypsum-depositingconditions on the floor of an epeiric seaway may also generate vugular porosity in the limestone immediately below the gypsum. In a study of vuggy porosity in the Mission Canyon Formation along the eastern margin of the Williston Basin, Luther (1988) found that vuggy dissolution occurred in a shallow subaqueous setting. The most pervasive vuggy

Data Center ,09126599985,[email protected], For Educational Uses 172 SEA-MARGINAL AND PLATFORM CALCIUM SULFATE or enlarged fenestral porosity was located near the contact between limestone and penecontemporaneously deposited gypsum. He postulated the deposition of the gypsum (now anhydrite) took place under less oxygenated conditions than the deposition of the limestone. Low oxygen levels are the norm in gypsum-depositing environments (Warren, 1986a for discussion). Microbial processes in the less oxygenated gypsum precipitating areas reduce free sulfate ions to sulfides. These sulfide-enriched brines flow seaward across the platform into the more oxygenated limestone forming areas. The introduction of sulfides into these more oxygenated environments produce an acid corrosive to limestone by the reaction:

HS- + 20, --> SO,"- + H'

Dissolution of the limestone by this acid probably took place near the sediment-water interface down to a shallow depth in the underlying sediments. The vugs formed by this process are morphologically similar to those formed by vadose processes. If the process is a common one, then there is the possibility of extensive vuggy porosity in depositional topographic lows and near many limestone-anhydrite contacts.

seawater flooding, rainfall, runoff eolian + I and Vl + detrital material surface * * water + A biological l and -_- chemical >v -- reactions sediments + eolian + and deflation porewater A

V groundwater groundwater groundwater outflow * pool inflow -

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Fig. 2.51. A) Subaqueous gypsum texture: AltPrnating luminoe of clear halite-f illed peudomorphr aftergypsum arid detrifalgptnm,now anhydrite. D. 0.E./Gruy Federal #I Grabbe (3641.0 ft.). Photo court of Joe Ramage. R) Constant vergence in a gvpsum/anhvdrite luyer in outcropping Cnstilr Formation, west Texas. The constant vergence in thir case ir probably due to the flow of laminated gypsum into a dissolution cavern created by the shallow subrurface dissolution of Castile halite.

Burial Hydrology: Gvpsum - Anhydrite, arid Dewateririg

'The calcium sulfate minerals Bpsum and anhydrite are important, often puzzling, phases that change syndepositionally or in the burial environment according to the prevailing temperature, pressure and salinity (Kinsman, 1974; Holser, 1979). Gypsum formation is favored by lower temperatures, lower pressure3 and relatively lower salinities; anhydrite is favored by higher temperatures, pressures and salinities (Fig. 2.52A). Gypsum is by far the most common calcium sulfate phase at earth surface conditions. Anhydrite can precipitate and grow by capillary evaporation in the unusually high salinities and temperatures of the Arabian Gulf sabkhas, or it can replace gypsum already present in the sabkha. At times after a heavy rainstorm the partially hydrated form of gypsum (bassanite - CaSO, 0 %H20)is as common as anhydrite on the sabkha surface (pers. obs. and Shearman, pers comm.). Anhydrite is less common in many other modern coastal evaporitic areas around the world where temperatures and humidities are modified by the coastal setting; in such settings, gypsum is the more likely sulfate and perennial halite is relatively uncommon (Kinsman, 1976). Penecontemporaneous anhydrite is more common under the extreme temperatures and closed basin hydrologies found in continental deserts.

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Even there, in many near-surface settings, gypsum not only grows but persists at temperatures and brine concentrations where anhydrite should be the stable phase. In this respect anhydrite is an exceptional mineral, it is easier to make it in the laboratory under conditions equivalent to natural surface settings than it is to find it in nature. Certain organic compounds and the high relative humidity of many evaporitic settings probably discourage the surface formation or preservation of penecontemporaneous anhydrite (Cody and Hull, 1980). Since their discovery in Holocene sediments of the Arabian Gulf, enterolithic and "chicken-wire" structures composed of anhydrite nodules have frequently been used as a diagnostic indicator of supratidal deposition in ancient carbonate-sulfate sequences. Shearman and Fuller (1969) noted the presence of anhydrite nodules with subparallel

100

80 anhydrite + brine 6O - Temperature "C brine halite + 40 - anhydriu +brine

20 - halite + gypsum + bnne 0 I I I I I I 0 2 4 6 8 101214 A Concentration Factor (c.f. seawater)

Temperature (DF)

I I I I 0 i' i' 2000- f Transition / I depth 4000 - Depth (ft) j 6000 - i' ELBrine 8000 - i' Gypsum $ loo00 -~ Ihhydrite B

Fig. 2.52. Gypsum-anhydrite transition. A) The stability of the calcium sulfate minerals (after Hardie, 1967). B) Stability of the calcium sulf ate minerals in terms of pressure and depth. Note the importance of pore f hid salinity (after Holser, 1979).

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crystal textures along the edge of the nodules in the Arabian Gulf and suggested this texture to be an indication of penecontemporaneous displacive growth. Early anhydrite crystals were displaced and fractured by later crystals growing in a framework of the earlier ones. However there are other ways anhydrite nodules can form enterolithic and "chicken-wire'' textures. When gypsum is buried and the temperature rises above 60°C it is transformed to anhydrite which often has a nodular and enterolithic texture (Fig. 2.52B). The process is complete at a depth of 1000 meters (3000-4000 ft). When the pore fluids are saline, the transformation may occur at depths as shallow as 1-2 meters (Holser, 1979; Shearman, 1985; Hovorka, 1988). The conversion reaction is:

CaSO, 2H,O --> CaSO, + 2H,O

CaS0,-saturated water is released upon burial as the gypsum is transformed to anhydrite. In a sluggish drainage situation, this adds a water-filled porosity equivalent to 38% of the original volume sulfate unit. As it drains from the mother bed, this water of dehydration interacts with its surrounds. A likely early effect of this dewatering is a decrease in the strength of the anhydrite unit by increased lubrication and perhaps overpressuring. If the water cannot drain freely from the dewatering gypsum, the bed may well become plastic. This explains some of the intense deformation and enterolithic textures observed in many ancient bedded anhydrites that have undergone only minor tectonism (as occurs in the Italian dolomites; figs. 13-16 in Schreiber et al., 1982). Dean et al. (1975) have shown that by themselves nodular anhydrites, including "chicken-wire"and enterolithic textures, are not reliable indicators of ephemeral water conditions. Although the origin of some of the nodules in the Castile Formation in the Delaware Basin is controversial, Dean et a]. (1975) concluded that this laminated deepwater gypsum had been converted to nodular and contorted nodular beds by the recrystallization of gypsum to anhydrite upon burial and not during evaporative drawdown. Others contend these nodules are supratidal. Constant vergence in some enterolith like structures in the outcropping Castile Formation suggests this folding of the lamination is a late-stage event (Fig. 2.51B). Much of it was generated during the rheotropic collapse of the laminar sulfate units into dissolution cavities created by relatively recent dissolution of interbedded halite units in the shallow subsurface. It is totally unrelated to the original depositional setting. Generally, enterolithic and nodular anhydrite is not environmentally diagnostic, it can form in several different ways; 1) in the supratidal zone of an Arabian Gulf sabkha, 2) during burial as gypsum dehydrates and compacts to anhydrite, 3) as a near

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surface, soft sediment slumping into the basin, 4) crumpling as diagenetic anhydrite hydrates back to gypsum - a 64% volume increase and 5) as shallow-burial features where laminar units slump into dissolution voids. Ancient nodular sulfates are always part of a vertical sequence, it is the changes up and down section, along with considerations of purity and thickness, that allow depositional interpretation (Warren, 1989). Enterolithic anhydrite formed tectonically or as slumps in bedded dewatered gypsum tends to show localized constant vergence of the fold axes. This can be used to distinguish some later diagenetic "enterolith-like" structures from early sabkha anhydrite. In the latter case, the orientation of the enterolith crests is almost always random (Fig. 2.6). As the water of crystallization drains out of the sulfate unit it can lubricate nearby growth faults and other relatively shallow slippage features including thrusts nappes. If drainage of the water of dehydration is sluggish, or if the replacement occurs in a zone of shallow brine flux with little overburden pressure, then the gypsum bed can retain enough rigidity to form aligned, nodular anhydrite pseudomorphs or ghosts after the original growth-aligned gypsum crystals (Hovorka, 1988; Shearman, 1985; Warren and Kendall, 1985). In the San Andres of the Palo Duro Basin, the very early replacement of gypsum from shallow refluxing brines formed pristine gypsum ghosts composed of anhydrite and halite at depths of a few meters (Fig. 2.51A; Ramage, 1987; Hovorka, 1987). Similar gypsum "ghost"textures formed smn after deposition in the 800 m thick Paleozoic Carribuddy "B' halite of the Canning Basin, Australia (Warren, 1990b). Water of crystallization escaping in the deeper subsurface can also effect units subjacent to the dewatering gypsum-anhydrite. The escaping water is saturated with respect to CaSO, but undersaturated with respect to other evaporite minerals including any nearby carbonates, halite and the bittern salts (Sass and Ben-Yaakov, 1977). As it flows through the surrounding rocks it can leach minerals to create secondary porosity or bring about mineralogical transformations. During brine flux in the shallow subsurface of the Palo Duro basin the flow of such waters dissolved marine carbonate to form anhydrite filled molds (Bein and Land, 1982). With dewatering a gypsum bed 10 meters thick will convert to a 6.2 meter thick anhydrite bed. This releases 4.9 m' of water per square meter of cross sectional area. At 30°C this volume of CaS0,-saturated water can dissolve 0.8 m3 of halite or 5.4 m3 of carnallite or convert 8.1 m3 of carnallite to sylvite (Blatt, 1982).

KMgCl, 6H,O + 4H,O --> KCI + Mg" + 2C1- + 10H,O carnallite syhite

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If these waters are also relatively hot they can carry and transform immature, shallow hydrocarbons (Warren, 1986a). Dewatering a thick gypsum unit causes a decrease in sediment volume and thickness as the unit compacts with a loss of 38% water and associated porosity. Usually this compaction will occur evenly across the whole evaporite unit. If the unit is widespread and of equal thickness the end result is not all that noticeable. But in certain depositional settings dewatering can enhance hydrocarbon migration pathways by differential compaction. Consider the example of a carbonate buildup capped and surrounded by a CaS0,-rich evaporite unit (Fig. 2.53). Evaporites off the crest of the structure would most likely be thick sequences of subaqueous, relatively pure gypsum and halite, the on-structure cap would most likely be an anhydritic, matrix-dominated sabkha succession (Fig. 2.53A; Warren and Kendall, 1985). When evaporite deposition ceased in the basin the upper surface would have been a near horizontal plane. As the succession was buried, the gypsum-rich off-structure evaporite sequence would dewater

burial compacbI

Original evaporite I thickness

Fig. 2.53. Schematic of effects of gypsum compaction during burial transformation to anhydrite. A) depositional setting. B) Burial compaction of the off -structure gypsum has created a new reservoir due to the greater volume loss in the salina-dominated areas verms the on-structure sabkha dominated areas.

Data Center ,09126599985,[email protected], For Educational Uses 178 SEA-MARGINAL AND PLATFORM CALCIUM SULFATE to around 60% of its original thickness. The on-structure matrix-rich sabkha sequence would also compact but not to the extent of the off-structure succession (Fig. 2.53B). The result would be a differentially compacted sequence draping over the buildup with excellent migration pathways into the region above the buildup. No matter if the underlying carbonates were tight, the differential compaction of an evaporite unit above the buildup could create potential reservoirs above the reef, a bonus if the reef were to be drilled. The reverse process, the conversion of anhydrite back to gypsum during erosion and exposure, occurs as the sulfate re-enters the meteoric realm and is bathed in cooler and fresher pore fluids. This process can form a range of diagenetic features in the newly precipitated gypsum including nodules, gypsum daisies, and pseudo-satinspar and is intimately associated with karstification (Warren et al., 1990). Hence, textures are more reliable depositional indicators of ancient evaporites in samples from the subsurface rather than from outcrop or subcrop.

Summary of Calcium Sulfate Diagenesis

Evaporites undergoing diagenesis experience both chemical and compactional effects. Chemical changes convert one mineral phase to another, often leaving evidence of the former mineralogy as pseudomorphs. Such reactions also release pore fluids which can then further react with the surrounding evaporite minerals. During burial the loss of water of crystallization from many of the early evaporite salts such as gypsum to anhydrite and carnallite to sylvite can lead to compactional draping with enhanced migration pathways in the units laid down above the evaporite unit. As it flows out of the system this water also lubricates the internal part of an evaporite unit to form flow textures (e.g. enteroliths) and may also enhance faulting, thrusting, and the growth of salt diapirs. The diagenetic complexity in an evaporite basin (Fig. 2.54) furnishes a feeling of the diverse feedback that occurs in an evaporite system from the time a unit is buried until the time it is eroded onto the surface or is completely dissolved in the regional groundwater flow system.

ACKNOWLEDGMENTS

I would like to thank the management and geologists of the Abu Dhabi National Oil Company, and the Ministry of Petroleum and Mineral Resources of Abu Dhabi for their support while I was in Abu Dhabi, most especially Ghelal Loutfi and Mohammed a1 Dubal. Thanks also to my present and former students, without their research much

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of the material presented in this chapter would not exist. Thanks also to Prentice-Hall for allowing me to publish portions of this chapter. Many thanks to my wife Jennifer who did the initial drafting of many of the figures on our "Mac" and also proofed the many drafts of this chapter. Finally, thanks to Sherry Proferes for drafting many of the "non-Mac"figures.

SEDIMENT (SHELL) (PELLETS) (PHOTOSYNTHETIC PRECIPITATES)

. (ALLUVIAL FANS) (EPHEMERAL STREAMS) SYNDEPOSITIONAL SHALLOW +- . (EOLIAN SEDIMENT) BURIAL SUBSURFACE t II I DEFLATION

(NODULES, CRYSTALS) DISPLACIVE - REPLACIVE CEMENTS

SHALLOW-INTERMEDIATE BURIAL t 1 BURIAL DIAGENESIS 4 (REPLACIVE - DISPLACIVE) (DISSOLUTION OR CEMENT) \ J 1 + METAMORPHISM

Fig. 2.54. Flowchart of the comp/ex diagenetic interactions an evaporite unit experiences from the time of first deposition until metamorphism. erosion or complete dissolution (after Warren, 1989).

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Data Center ,09126599985,[email protected], For Educational Uses Chapter 3

DEPOSITIONAL ENVIRONMENTS OF NON-MAR INE EVAPORITES

Joseph P. Smoot and Tim K. Lowenstein

INTRODUCTION

Non-marine evaporite deposits are common features of modern arid closed basins, but relatively few have been recognized in the geologic record. They are important because their depositional setting reflects both climatic and tectonic conditions and they may contain economically important mineral resources. The identification of ancient non-marine evaporites is straightfotward in some cases, such as the Eocene Green River Formation, based upon diagnostic saline mineral assemblages, fossils, or the regional depositional setting. Commonly, however, these features are absent or ambiguous, which makes the distinction between marine and non-marine evaporites more difficult to assess. The distinction is further complicated by the possibility of intercalation of non-marine deposits with coastal marine saline deposits and by the mixing of marine and non-marine waters in isolated coastal lakes. This overlap between marine and non-marine evaporites suggests the division should not be treated as absolute. Most modern non-marine evaporites and almost all ancient non-marine evaporites occur in saline lake deposits or their associated facies. The term saline lake includes perennial lake, ephemeral lake, playa-lake, playa, salina, salar, chott, salt pan, inland sabkha, and any other terms used to refer to areas that, at least intermittently, hold standing bodies of water. The major exception to lacustrine-associated non-marine evaporites is saline soils. Saline soils are common in modern arid and semi-arid regions, but are rarely recognized in the rock record. The most comprehensive overviews on the sedimentology of non-marine evaporites are Strakhov (1970), Glennie (1970, 1987), Cooke and Warren (1973), Eugster and Hardie (1978) and Hardie and others (1978). General descriptions of non-marine evaporites are included in summaries on lacustrine deposits (Reeves, 1968; Picard and High, 1972a, 1981; Fouch and Dean, 1982; Dean and Fouch, 1983; Eugster and Kelts, 1983) or evaporites (Kendall, 1979, 1984; Schreiber and Hsii, 1980; Sonnenfeld, 1984; Warren, 1989). This paper summarizes the sedimentary characteristics of non-marine environments in which evaporites occur with emphasis on depositional and diagenetic processes.

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Our approach is modeled after Hardie and others (1978) treatment of saline lake deposits. We review hydrologic and chemical settings of non-marine evaporite accumulation and then describe the sedimentary characteristicsof subenvironments that contain non-marine evaporites or that are associated with them. Map distributions of modern evaporites and stratigraphic sections of ancient deposits will illustrate the lateral and vertical association of the subenvironments. Finally, we discuss criteria for distinguishing between non-marine and marine evaporite deposits.

MODERN NON-MARINE EVAPORITES

‘The general distribution of modern non-marine evaporites is shown in Figure 3.1 and references are presented in Table 3.1. Not all of the lakes included in the Table are presently forming evaporites, although they have the potential. Also, the criteria for defining a saline lake have not been evenly applied. For instance, Hardie and others (1978) chose a minimum concentration of SO00 ppm for a saline lake after Beadle’s (1 974) upper salinity tolerance of freshwater organisms, whereas Australian limnologists use 3000 ppm (Geddes and others, 1981) as the boundary between freshwater and saline faunas.

Fig. 3. I. Distribution of areas without surface drainage and areas with interior basin drainages. This is essentially equivalent to the distribution of all modern 0ccurrence.s of non-marine evaporites other than Antarctica. (Redrawnf rom Cooke and Warren, 1973).

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Hydrologic Setting

Thick accumulations of non-marine evaporites may form where evaporation rate exceeds inflow, where inflow is sufficient to supply solutes, and where the inflow accumulates in a closed basin (Langbein, 1961) or in a basin with restricted outflow. Arid or semi-arid conditions occur in the global high pressure belts of the subtropical horse latitudes and the poles, in the midlatitude intracontinental deserts and steppes of Asia and North America that are isolated from oceanic moisture, and in orographic (rain shadow) deserts that may occur at any latitude (Fig. 3.1 and Fig. 3.2). Solutes are introduced by rivers and streams, springs, groundwater, rainwater, and by wind-blown aerosols and dust. Closed basin conditions can occur in a myriad of ways: 1) tectonic basins, including fault-related intermontane basins and intracratonic structural sags, 2) interdunal depressions in sand seas, 3) wind deflation hollows, 4) blocked or abandoned fluvial channels or glacial valleys, 5) volcanic or meteor impact craters, and 6) combinations of two or more of the above (Fig. 3.3). Orographic deserts in fault-bounded intermontane basins produce the thickest modern non-marine evaporite deposits. Moisture trapped in the surrounding mountains provides a steady source of solutes to the basin, while the structural and climatic setting allows the basin to be a long-lived sink for water and sediment. The volume of water entering saline lakes and leaving by evaporation can be veIy large. For instance, the Jordan River brings about 1.25 km3 of water per year into the Dead Sea, which is about 80% of the annual inflow, but the evaporation rate of 1.6 m/year removes about 1.58 km3 of water from the lake (Hardie and others, 1978). A single flood brought 39.3 km’ of water into the Lake Eyre basin in 1974, but 39.5 km3 of water was removed during the next two years by evaporation (Kotwicki, 1986). Closed basin lakes are unstable features that exist through a delicate balance of inflow and evaporation (Langbein, 1961). For instance, over the period of 1931 to 1976 the net inflow into Great Salt Lake, Utah was essentially equal to the net evaporation. If evaporation exceeds inflow by too large a margin, the lake level will drop, sometimes to complete desiccation. If inflow exceeds evaporation, lake level will rise eventually spilling the lake out of the basin. Examples of modern closed lake fluctuations in response to droughts and periods of high rainfall are common; for instance, historical records of Great Salt Lake show fluctuations in depth of several meters (Arnow, 1980; Spencer and others, 1984) which produced lateral shoreline shifts of over 10 km (Currey, 1980). The man-made diversion of the major river inflow to the perennial Owens Lake resulted in its complete desiccation and present status as a saline pan (Friedman and others, 1976; Alderman, 1985). Many Pleistocene pluvial lakes spilled out of closed

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Table 3. I. Selected references on Modern and Quaternary non-marine evaporites.

Central North America Northern United States: Ahlbrandt, 1973;Ahlbrandtand Fryberger, 1980,1981;Bradley and Raintree, 1956; Condra, 1918; Grossman, 1949,1968; Hicks, 1%1 Southern United States: a) Salt Flat Graben - Boyd and Kreitler, 1986; Friedman, 1965; Hussain and Warren, 1988,1989 b) Rio Grande Rat - Allmendinger, 1971; McKee, 1%; McKee and Douglass, 1971; McKe and Moiola, 1975; Neal and others, I%, Schenk and Fryberger, 1% Simpson and Loc~pe,1985; Talmage, 1932 c) Other - Carlisle and others, 1978; Ostercamp and Wd,1987; Powers, 193% Reeves, 1966,1972; Reeves and Reeves, 1971; Schreiber and others, 1972; Wood and Ostercamp, 1987 Central Canada: Cole, 1% Grossman, 1949,19f& Last, IW,1989; Last and Schweyen, 1983,1985;Last and Slwak, 1%, 1987; Rawson and Moore, 1944; Sle7ak and Last, 1Y85 Western North America British Columbia: Cummings, 1940, Nesbitt, 1974 US. Basin and Range: a) Death Valley - Blackwelder, I%$ Butler, 1984, Goudie and Day, 1981; Hunt, 196h; Hunt and Mabey, 1966; Hunt and others, 1% b) Saline Valley - Gale, 1914; Hardie, 1%8, Lombardi, 1963 c) Deep Springs Lake - Clayton and others, 1068; Jones, 1965; Peterson and others, 1%3,1%6; Slack, 1967 d) Searles Lake - Eugster and Smith, 1965; Felmy and Weare, 1%; Gale, 1915; SchoU, 1960, Smith, 1966, 1968,1419;Smith and Haines, IW,Smith and others, 1983; Smith and F'ratt, 1957 e) Bristol Dry Lake - Basset and others, 1959; Handford, 1982a,b;Rosen, lMu( f) Clayton Playa - Kunaw, 197% Mein~er,1917 Moiola and Glover, 1965 g) Panamint Valley - Carran~a,1965; Gale, 1915; Motts and Carpenter, 1968; Smith and Pratt, 1957 h) Great Salt Lake -Amow, 1980, Cl(roni, 1962;Cohenour, 19h6; Currey, 1980, Eardlcy, 1938,1%2a,b,lw, Eardley and Gvodetsky, 1960; Eardley and others, 1973; Eardley and Stringham, 1952; Gilbert, 1885,1890, Cnvynn, 1980; Gymn and Murphy, 1% Hahl, 1968; Hahl and Handy, 1969; Hahl and Langford, 1W, Halley, 1976,19n; Jones, 195% Morrison and Frye, 1965; Sandberg, 1975; Scott and others, 1982; Spencer, 1982, Spencer and others, 1984,1985; Whelan, 1973 i) Pyramid Lake - Born, 1972; Main, 1970; Radbruch, 1957; Smoot and LeTourneau, 1989 j) Akrt Lake - Deike and Jones, 1980; Jones and Weir, 1983 Philips and Van Denburgh, 1971; Van Denburgh, 1975 k) Walker Lake - Benson, 1988, Benson and Thompson, 1987; Bradbury and others, 1989; Cooper and Koch, 1% Link and others, 1985; Morrison, 1%. Morrison and Frye, 1965; Osborne and others, 1982; Russel, 1885; Smoot and LeTourneau, 1989, Spencer, 1977,1985 1) Owens Lake - Alderman, 1985; Friedman and others, 1976; Gale, 1915; Hirschkind, 15%; Matsuo and others, 197% Smith and Friedman, 1986, Smith and Pratt, 1957 m) Salton Sea - Carpelan, 1959; Long and Sharp, 1964, Sharp, 1W,Van de Kamp, 1973 n) Others - Anderson, 1958,1959,Anderson and others, 1985; Bailey, 1904; Basset and others, 1959; Bateman and Hess, 1978 Cannon and others, 1975; Carlisle and others, 1978; Droste, 1959,1961; Dunn, 1953; Eberl and others, 1982, Foshag, 1921; Khoq and others, 1982, Langer and Kerr, 1966; Lattman and Lauf- fenburger, 197% Libbey, 1985, Markgraf and others, lm,Meinzer, 1917; Motts, 1965,1969; Motts and Carpenter, 1968; Neal and Motts, 1967; Neal and others, 1%8; Papke, 1976,1985; Phillips and Van Denburgh, 1971; Rooney and others, 1969; schd and Taft, 1964;Shearman and others, 1989, Sheppard and Gude, 1968; Smith, 1985; Smith and Drever, 1976; Smith and Pratt, 19sI; Snyder, 1962; Stone, 1956, Van Denburgh, 1975; Ver Plank, 1958; Walker and Motts, 1969 South America a) Laguna Salinas - Muessig, 1958,1959, Norman and Santini, 1985

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Table 3. I (continued)

b) "yncalaya Saka - Hurlhut and others, 1973 c) Pastos Grandes - Risacher and Eugster, 1979 d) Salar Uyuni - Rettig and others, 1080 e) Salar de Atacama - Dingman, 1967 f) Others - Badnut and Risacher, 1983; Benavidcs, 1%; Brcdtkorh, 1980; Brdtkorband others, 1982, Conway, 1961; Diar, 1984: Ericksen, 1981; Eriksen and others, 1976; Garrett, 1985; Mucssig, 1959,lW); Norman and Santini, 1985; Stnert7 and Ericksen, 1974 Australia a) Lake Eyre - Baas-Becking and Kaplan, 19% Bonython, 1955,1Y.%, Btmythnnand Mason, 1953; Callen andothers, 19W, Dulhunty, 1975,1977,1982, Habermehl, 1088; Johns and Ludhrook, 1963; King, 1956; Kotwieki, 1%; Mag- and others, 1988, Ruello, l%'6; Twidale, 1972 b) Lake Amadeus - C%en Xiang Yang and Barttm, 19m, (%en Xiang Yang and Bowler, 1988, Jacobson and Janknwski, 1988 c) Lake ljrrell - Bowler and Teller, 19M, Luly and others, 1% Teller and others, 1982 d) Lake Csorge - Singh and others, 1981 e) Lake Frome - Bowler and Magee, 1%; Bowler and others, 1% Draper and Jensen, 1976; Singh, 1981; Ullman and McLeod, 19%; Wasson, 1983 f) Maar Lakes - DeDeckker and Last, lW, 19W, Last and DcDeckker, 1987 g) Others - Arakel, 19%; Arakel and McConchic, 1982, Bettenay, 1962; Bird and others, lW,Bowler, 1%; Bowler and others, 1986, Bume and Fergusnn, 198% Bume and others, 1980, Callen, 19x4, Clivas and others, 19% DeDeckker, 1983,1988, Duty, 1973; Geddesand others, 1981;Johns, 1968, Krinsley and others, 1963;Lock, 1988a,b; Magee, 1988, Mann and Hornwib, 1979; Teller. 19M; Veevers and Rundle, 1979; Wakelin-King and Arakel, 1988. Williams, 1963, 1971, 1973; Williams, 1981 Antarctica Angino and others, 1962; Burton, 1981; Tominaga and Fukui, 1981 Spain Garcia del Cura and others, 1418; heyo Mur and others, 1%; Pueyo Mur and Urpinell, 1% Asia Central Asia: a) Qaidam Basin - Bowler and others, 19M; Chen Kemo and Bmvler. 1986, Lnwenstcin and othcrs, 1989; Momin Gu Shuqi and others, 19x6; Petrnv, 196% Qien Ziqiang and Xuan Zhiquiang, 1985; Strakhnv, 1970; Sun-Dapeng and others, 19W, Valyashko, 1972a; Yuan Jianqi and others, 1983,1985; Zhang Pengxi, 1985 b) Gull of Kara Bogux - Dickey, 1968, Garbell, 196% Straknv, 1970; Tcodornvieh, 1x1;Valyashko, l972b c) Lake lnder - Valyashko, 1972a d) Others - Barker and LcFond, 1985; Petrov, 1907; Pnsokhov, 1949; Smith, 1975; Spasskaya, 1988, Straknv, 1970; Valyashko 1972a,b; Wen Xue7e and Wu Dizong, 1985; Zemljanima, 1973 Southeastern Asia: a) Iran Playas - Alsinawi and Saadallah, 197% Bobek, 1959; Krinsley, 1068,1!Y70,1972,197~Stalder, 1975; Stocklin, 1968 b) Lake Van - Degens and others, 1984 c) Toz Golu - Irion and Miller, 1968; Muller and Irion, 1x9; Muller and others, 1972; Uypun, 1980, Uygun and Sen, 3978 d) Indian Lakes - Khandelwal, 1975; Krishnan, 1968, Singh and others, 1972,1974; Wasson and others, 1983, 19x4 e) Others - Barzanj and Stoops, 197% Desio, 197Q Goudie, lm,Mdler and others, 1972; Uygun, 1980 Africa East Africa: a) Dead Sea and Vicinity - Amiel and Friedman, 197l;Assaf and Nissenbaum, 19R;Begin and others, 1974, 1980, Bentor, 1%1,1%8; Beyth, 1980; Garber, 1980, Gomitz and Schreiber, 1981; Jones and others, 1976;

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Table 3. I. (continued)

Karci, 197% Karc7 and Zak, 1987; Kat7 and others, 1977; Lerman, 1967; Livnat and Kronfeld, 1985; Manspewsr, 1978,1985;New, 1978; New and Emery, 1967 Schulman and Bartov, 1978; Sirkes, 1987; Sneh, 1979, 1W, Steinhom, 1983; Steinhom and Gat, 1983; Weiler and others, 1974 Zak, 1974,lW Zak and Bentor, 1963 h) Lake Magadi - Baker, 19%, Eugster, 1969,1970, 1980; Eugster and Jones, 1%8;Hillaire-Marcel and Casanova, 1987; Hillaire-Marcel and others, I%, Jones and others, 1977 Sheppard and others, 1970; Surdam and Eugster, 1976 c) Lake Km - Stdfers and Fishbeck, 1974 Stdfers and Heeky, 1978; Degens and others, 1973 d) Lake Tanganyika - Degens and others, 1971; Stdfers and Hecky, 1978 e) Lake Turkana - Ahell and others, 1978,1982; Cerling, 1979,19%, Halfman and Johnson, in press; Vondra and Bowen, 1978; Yuretich and Cerling, 1%3 f) Lake Urima - Kelts and Shahrahi, 1986 g) Danakil Depression - Holwerda and Hutchinson, 1% Von Damm and Edmnnd, 19x4 h) Others - Baumann and others, 1975; Beadle, 1974; Beachamp, 19M; Degens and Kulhicki, 1973; Ecles, 1974; Hay, 1W; Hecky and Degens, 1973; Hillaire-Marcel and Casanova, 1987: Hillairc-Marcel and others, 19%, Miiller and Forstner, 1973; Rosendahl and Livingstone, 1983; Vnndra and others, 1971 North Africa: a) Lake Chad - Durand, 19Q, Eugster and Maglime, 1979 Gac and others, 1977 Maglime, 1972,1974,1980; Pedro and others, 1978; Roche, 1977 h) Saudi Arabia - Ahlhrandi and Fryherger, 1981; Fryherger and nthers 1983, 1984; Glennie, 1970; Johnson and others, 1978 c) Others - (iavcn and others, 1981; Glennie, 1970, 1987: McKee and Tihhitts, 1%; Meckelein, 1977; Petit-Maire and Riser, 1981: Tortochaux, 1%; Tricart. 1967; Trichet, 1%3 Watson, 1983a,h, 1985, 1988 South Africa: Flint and Bond, 1W& Foshag. 1913; (ievers. 1030, (kvers and Van der Westhym~,1y31; Lancaster and Teller, 1988; McKee, 1982; Milton and Naeser, 1971; Shaw, 1988; Teller and Lancaster, 19W, Watson, 1983a,h, 1985,1988 basins due to higher inflow at that time (Smith, 1979; Rettig and others, 1980; Cooper and Koch, 1984; Bowler and Teller, 1986; Benson and Thompson, 1987; Benson, 1988). If fluctuating lake levels are common, then a vertical sequence of deposits that reflect varying depths and salinities should be equally common. In the Basin and Range Province of the western U.S., many Pleistocene vertical successions contain freshwater lake deposits representing pluvial lake highstands when water depths were over 100 m, that grade upward to playa mudflat or saline mudflat deposits, reflecting modern arid conditions. Fluctuating lake levels of this type have been inferred from core samples of modern lake basins (Smith and Pratt, 1957; Morrison and Frye, 1965; Motts and Carpenter, 1968; Smith and Street-Perrott, 1983; Smith and others, 1983; Benson, 1988; Bradbury and others, 1989) and from ancient deposits (Van Houten, 1962,1964; Picard and High, 1968; Eugster and Hardie, 1978; Smoot, 1983; Olsen, 1986; Olsen and others, 1989; Southgate and others, 1989). Another feature of closed basin hydrology is the response relationship between

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Fig. 3.2. Distribution of saline lakes in A). western North America, B). west-central South America, C). and Australia. Lakes are shown in black. A) is modified from Eugster and Hardie, 1978; B) is modified from Stoertz and Ericksen, 1974; C) is from Bowler, 1986).

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Fig. 3.3. Examp1e.s of saline lake setting.s. A). Luke Chad - drowned dunefield in the northeast corner of the lake f ornzs restricted pockets of higher salinity. (Photo courtesy of H.P. Eugster). B). Saline Valley - block-fault intermontane basin with up to 1800 m of relief from the basin floor. View is to the Qastf'romthe ,saline pan. C). Saskatchewan saline lakes - satellite pictltre of sinuoiis lakes (blocked river channels - black) in aglncial terrain in south-certtral Saskatchewan, C'unadn. Area o,f view is about 180 km from east to west. (Photo coicrtesv of W.M. 1,a.c.t). surface inflow and lake volume (,Lanebein, 1961). Surface inflow into a closed basin usually sinks into the substrate, adding to the groundwater, or contributes to a standing body of water within the basin. Therefore any closed basin with substantial surface inflow inusf have at least an intermittent lake and the area occupied by that lake will be related to the discharge characteristics of the inflow system and the area of the drainage basin (Hooke, 1968: Smoot, 1985). A result of this relationship is that

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deposition during fluvial flooding events will occur while the lake boundaries are shifting laterally. This lateral shift may be very large, if the basin floor has low relief and if there is an ephemeral or very shallow perennial lake. The volume of discharge needed for a lateral shift of the lake boundary necessarily increases as the area of the lake increases. Such dynamic interplay between surface inflow and ponded lake water is particularly important to the geometry of shoreline deposits and their sedimentary features. Groundwater systems are important in the formation and preservation of non-marine evaporites. Groundwater may contribute surface inflow as springs and seeps. Artesian springs develop: 1) where groundwater aquifers approach the sediment surface, 2) where faults or fractures intersect the aquifer, 3) where a porous aquifer abuts against less permeable basinal sediment, and 4) where dilute groundwater flow intersects stagnant basinal brines (Bowler, 1986). Groundwater may be concentrated near the surface by evapotranspiration (Nesbitt, 1974; Eugster and Hardie, 1978) or by evaporation. Hsu and Siegenthaler (1 969) introduced the term evaporative pumpingfor vertical and lateral groundwater flow caused by evaporation of pore waters. Macumber (1983) argues that evaporation will only cause draw-down of the water table and that regional flow paths are defined by the hydraulic head in the aquifer system. He attributes diurnal vertical fluctuations in the groundwater level in Lake Tyrrell to variations in atmospheric pressure. Dense surface brines concentrated by evaporation may sink through sediments pushing less dense groundwater laterally (Valyashko, 1972a; Bowler, 1986). The brines may also diffusively mix with older groundwater across a slowly migrating interface (Lerman and Jones, 1973; Spencer and others, 1985a).

Chemistry

The type of saline minerals found in non-marine evaporite deposits depends on: 1) the chemical composition of the inflow waters, and 2) the mechanisms whereby these waters become brine. Table 3.2 lists evaporite minerals found in non-marine settings. We briefly summarize the types of inflow waters in non-marine evaporite systems and the modification of these parent waters into brine. The processes that modify parent waters during their evolution into brine include evaporative concentration, mineral precipitation, syndepositional recycling, diagenetic mineral reactions, and exchange reactions with pore fluids (Jones and Vandenburgh, 1966; Hardie and Eugster, 1970; Eugster and Hardie, 1978; Eugster and Jones, 1979; Hardie, 1984; Spencer and others, 1985). Major types of inflow in non-marine basins have been reviewed by Hardie (1984),

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Table 3.2. Chemical f ormulas of some evaporite minerals found in non-marinesettings.

MINERAL FORMULA MINERAL FORMULA

Alunite WdOH)dSO4 k Leonite MgSO, K,SO, 4H,O Anhydrite CaSO, Loewitc ZMgSO, 2Na,SO, 5H,O Antarcticite CaCl, 6Hz0 Magadiite NaSi,O,, 3H,O Aragonite CaCO, Magncsian Calcite (MgxCa, x)CQ Barite BaSO, Magnesite M&O3 Bassanite ZCaSO, H,O Meyerhofferite CazB60,, 7H,O

Bischdite MgCl, 6H,O Mirabilite Na2S0, 9 lOH,O Blocdite Na,SO, MgSO, 4H,O Nahcolite NaHCO, Borax Na2B,07 10H,O Natron Na,CO, 10H,O Burkeitc Na,CO, 2NazS0, Nobleite CaB,OIo 4H,O Calcite CaCO, Northupite Na,MgCI(CO,) C'amallite KCl MgC1, 6H20 Pentahydrite MgSO, 5H,O cclestite SrSO, Pirssonite Na,Ca(CO,k 2H20 Colcmanite CaZB60,, 5H,O Polyhalite K,Ca,Mg(SO,), 2H,O Dolomite CaMg(CO,k Priceite CrhB10019 ~HzO Epsomitc MgSO, 7Hz0 Prokritc NaCaBaB,Og5H,O Gaylussite NazCa(CO,), 5H,O Rinneite FeCI, NaCl 3KQ Glauherite CaSO, Na2S04 Sanderite MgSO, 2H20 Gypsum CaSO, 2H20 Schairerite Na,(SO,)(F,CI) Halitc NaCl Schoenitc MgS0, K,SO, 6H20 Hanksite 9Na2SOa 2Na,CO, KCI Searlesite NaBSi,O, H,O Hexahydrite MgSO, 6H20 Shortite ZCaCO, Na,CO, Howlitc Ca2SiB,09 (OH), Sdfohalite Na,C1F(S04), Hydroboracite CaMgB,O,, 6H20 Sylvite KCI lnderitc MgzB60,, 15H,O Syngenite CaSO, K2SOa H,O Inyoite C?dZB60,, 13H,O Tachyhydrite Caa, ZMgCl, 12H,O Jarosite KF%(W,(SO& Thcnardite NazSO, Kainite 4MgS0, 4KCl SH,O Thermonatrite Na,CO, H,O Kenyaitc NaSiO,,~,,(OHJH,O) Tincalconite NazB40, SH,O Kcmite Na,B,O, 4H,O Trona NaHCO, Na,COh 2H,O Kicscrite MgSO, H,O Tychite N%M&(SO,) (CO,), Kumakovite MgB,O,, 15H@ mxite NaCaB,O, 8H,O Langkinite ZMgSO, K2S04 Van'thoffite MgSO, 3Na,S04 Leonhardtite MgSO, 4H,O and include meteoric water, diagenetic reaction water, hydrothermal reaction water, volcanogenic water, and mixed water with contributions from two or more types of inflow. Meteoric water contains solutes derived from the low-temperature weathering reactions between bedrock and rainwater. They are commonly dilute Na-Ca-HC0,-bearing waters, but their chemical composition is variable depending upon the rocks undergoing weathering (Hardie and Eugster, 1970; Eugster and Hardie, 1978). Hydro- thermal source water is produced by circulation and interaction of hot groundwater with

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sediment or rock at low grade metamorphic conditions. Such waters are commonly rich in Na' and Ca' and contain Cl- as the dominant anion (Hardie, 1984). Diagenetic source water is formed by sediment water interactions at temperatures and pressures below metamorphic conditions. Important reactions include dolomitization of carbonates, albitization of detrital feldspars, and clay mineral reactions (Jones and Bodine, 1987). Evaporative concentration of the inflow waters eventually leads to precipitation of saline minerals. Saline minerals may also form by brine mixing with another brine or dilute water, organochemical reactions, or by changes in temperature. The type of salts formed and the order in which they precipitate is fundamentally controlled by the chemical composition of the source waters and the chemical pathways they follow. Precipitation of salts cause major changes in the composition of the remaining water because each salt removes components from solution (Hardie and Eugster, 1970; Eugster and Hardie, 1978; Eugster and Jones, 1979). Precipitation of the early, relatively insoluble mineral phases, especially alkaline earth carbonate and gypsum, is the most important step for determining the pathways of brine evolution during evaporation. The ratio of equivalents of Ca to HCO, + CO, in the inflow water determines whether Ca or HCO, + CO, is depleted in the remaining water (Fig. 3.4). Either an alkaline brine, poor in Ca, or a neutral brine, poor in carbonate and bicarbonate will result (Hardie and Eugster, 1970; Eugster and Hardie, 1978). Similarly, the ratio of Ca to SO, equivalents in the evolving water at gypsum saturation determines whether the remaining water will be depleted in Ca or in SO, following precipitation of gypsum (Figs. 3.4 and 3.5). The major types of inflow water evolve into brines with distinctive chemical compositions and which precipitate characteristic saline mineral assemblages. Dilute meteoric water with equivalents of Ca < CO, + HCO,, precipitates alkaline earth carbonates and evolves into alkaline Na-K-HC0,-C0,-SO,-C1-rich brine that may form trona, natron, nahcolite, halite, mirabilite, or thenardite. This type of non-marine evaporite is common in modern settings such as Mono Lake, Deep Springs Lake, and Owens Lake in California, Abert Lake in Oregon, and Lake Magadi and Lake Chad in Africa. The thick Pleistocene salt sequence of Searles Lake, California, also falls into this group, as do the vast evaporite deposits of the Eocene Green River Formation of Wyoming, Colorado, and Utah, and the trona-rich lacustrine deposits of the Eocene Wucheng Basin, Hunan Province, China (Zhang Youxun, 1985). Alternatively, after alkaline earth carbonate precipitation, dilute meteoric water with Ca in excess of bicarbonate may evolve into neutral Na-K-Mg-CI-SO, brine from which gypsum, halite, mirabilite, thenardite, glauberite, polyhalite, epsomite, or bloedite precipitate. Modern

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so;

Fig. 3.4. Ternary phase diagram illicstrating how inflow waters evolve into 1irine.c.. TWO chemical divide.r (linesfrom Calcite to SO, and Calcite to Gypsum-Anhydrite) segregccte waterS that will evolve (arrowr), upon evaporation and precipitation of calcite and gypm m, into Ca- CI h rin PS. CI - S 0,- rich h r ines, and N a - H C HCO, + CO, + SO,, these waters have been termed calcium chloride brines with Ca at least partly balanced by C1 (Fig. 3.4; Hardie, 1984; Spencer, 1985). Calcium chloride brines occur in Bristol Dry Lake, California, the Dead Sea, and some saline lakes of the Qaidam Basin, China (Fig. 3.5; Zhang Pengxi, 1987; Lowenstein and others, 1989). Dia- gnostic saline minerals formed from this brine type include halite, sylvite, carnallite, bischofite, tachyhydrite, and antarcticite. The "end-member" brine types discussed above may be modified in composition by inflow of more than one source water, and by changes in the relative volume of different inflows through time. For example, in tectonically active block fault or rift basins, inflow of meteoric water may combine with subsurface-derived hydrothermal, diagenetic, or volcanogenic water (Hardie, 1984). In Great Salt Lake, Utah, NaC1-rich

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z 0

W 2 2 U a0 <> W

Fig. 3.5. Flow diagram for brine evolution. Solid boxes are critical precipitates and dashed boxes are water compositions. Final brine compositions include examples of saline lakes. (from Eugster and Hardie, 1978).

hydrothermal springs contribute the bulk of the Na', K', and CI- to the basin, whereas dilute Ca-HCO,-rich river water supplies the majority of the Mg, Ca, HCO,, SO,, and SO,, but the volume of these inflow waters varied over the past 32,000 years (Spencer and others, 198%). During fresh water periods (high lake levels), inflow was dominated by Ca-HCO,-type river water. During saline stages (low lake levels), NaCI-rich hydrothermal springs were the major solute source. Another process that may influence the composition of non-marine water and, hence, non-marine evaporites is syndepositional recycling. In this process, dilute inflow water redissolves previously deposited salt. The bulk of the solutes in saline-pan ephemeral lakes may be derived from the preferential dissolution of the most soluble salts such as, halite, thenardite, natron, in efflorescent crusts by surface runoff and lake expansion. Such recycling has been tested experimentally by Drever and Smith (1977) and has been documented in Saline Valley, California (Hardie, 1968; Lowenstein and Hardie, 1985) and Lake Magadi, Kenya (Eugster, 1969, 1970; Jones and others, 1977; Eugster and Hardie, 1978; Eugster and Jones, 1979). In Saline Valley, dissolution of halite-rich efflorescent crusts produces nearly pure NaCl brines that precipitate halite

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layers, and in Lake Magadi, dissolution of efflorescent crusts containing thermonatrite produces brines that precipitate layers of trona. Finally, the chemical composition of non-marine waters may be syndepositionally modified by diagenetic reactions within the sediment and by diffusion of brine into the pore fluid. Jones (1986) summarizes the literature on the formation of authigenic clays, illitization of smectite and formation of a Mg-rich trioctahedral mineral, that removes K, Mg, and Si from pore water. Examples of clays enriched in Mg, K, and Si relative to detrital clay minerals have been reported from Lake Abert, Oregon (Jones and Weir, 1983), Lake Chad, Africa (Gac and others, 1977; Pedro and others, 1978), the Eocene Green River Formation, Utah (Dyni, 1976), and the Miocene of the Tajo Basin, Spain (Jones and others, 1986). In Great Salt Lake, Utah, Mg is removed in significant amounts from the lake brine by diffusion into pore fluids and by the precipitation of an authigenic magnesium silicate in the top meter or so of sediment (Spencer and others, 1985a). Other early diagenetic processes include precipitation of saline minerals in the sediment by displacive growth or as cement (i.e. gypsum and halite), formation of diagenetic silicates (i.e. zeolites), and adsorption of species in pore water onto active surfaces (i.e. volcanic glass). Although little is known about the extent to which these processes affect surface brines, they all may change the chemical composition of surface water and shallow ground water by removal of species from solution and, hence, exert some control on the path of brine evolution.

DEPOSITIONAL SUBENVIRONMENTS OF NON-MA RINE EVAPORITES

Hardie and others (1 978) suggested a classification of depositional subenvironments for saline lake deposits that would be applicable to the geologic record. We will update this classification with some modification and the addition of a section on saline soils. The subenvironments considered in this paper are: 1) perennial saline lake, 2) saline pan, 3) saline mudflat, 4) dry mudflat, 5) shoreline, 6) alluvial fan-sandflat, 7) ephemeral stream floodplain, 8) perennial stream floodplain, 9) eolian dunefield and sand sheet, 10) spring, and 11) saline soil. Our presentation groups these subenvironments into Lacustrine Deposits (perennial saline lake, saline pan, saline mudflat, dry mudflat, and shoreline subenvironments), Flu vial Deposits (alluvial fan-sandflat, ephemeral stream floodplain, and perennial stream floodplain subenvironments), and Other Deposits (eolian dunefield and sand sheet, spring, and saline soil subenvironments).

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LACUSTRINE DEPOSITS

Lacustrine evaporites and associated deposits range from accumulations in deep, stable lakes to deposits formed in centimeter-deep lakes that persist for a few days before drying completely for several years.

Perennial Saline Lake Subenvironment

Perennial saline lakes contain standing brines that persist for tens to thousands of years. Perennial saline lakes require substantial inflow, normally as perennial streams and rivers, and a closed drainage in order to allow solutes to concentrate. Evaporation concentrates the inflow water into brine which eventually reaches saturation with respect to saline minerals. However, perennial closed basin lakes can exist for a long time without precipitating evaporites. For example, Yuretich and Cerling (1 983) estimate that Lake Turkana has been closed for 2000 to 3000 years, but only minor alkaline earth carbonates have accumulated. Most perennial saline lakes that are more than a few meters deep are stratified (meromictic) with denser brine at the bottom, such as the Dead Sea (Neev and Emery, 1967) or Great Salt Lake (Spencer, 1982). Density stratification may lead to a heliothermal lake with a solar heated bottom water below a cooler, less dense water (Kirkland and others, 1983; Sonnenfeld and Hudec, 1985). Wave action can destroy density stratification, particularly in shallow lakes, but in deep lakes that remain stratified for long periods of time, a stagnant anoxic lower brine layer may develop. The largest modern perennial saline lakes are Lake Chad in North Africa (l0,OOO to 26,000 km2,4-7 m deep) and Lake Balkhash in the U.S.S.R. (17,4OO km2,26 m deep). Most other modern perennial saline lakes are also shallow, typically less than 10 m deep. The only deep lake that is presently accumulating evaporites is the Dead Sea (approximately 900 km2and maximum depth greater than 300 m). Several deep perennial lakes are moderately saline and may evolve into evaporite precipitating systems, such as Lake Issyk Kul, U.S.S.R. (6200 km2, 702 m deep) and Lake Kiw in East Africa (2060 km2, 485 m deep). The Caspian Sea is a landlocked seawater mass that is now a brackish lake with evaporites presently forming in the partially isolated Gulf of Karabogaz. The Caspian Sea has an area of 371,000 km2and a maximum depth of 1025 m. Perennial saline lakes that are relatively dilute precipitate alkaline earth carbonate minerals such as low Mg-calcite, high Mg-calcite, protodolomite, aragonite, monohydrocalcite, or dolomite (Strakhov, 1970; Miiller and others, 1972; Von der Borch, 1976

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Stoffers and Hecky, 1978; Spencer, 1982). However, shallow areas of such lakes, that are restricted from mixing freely with inflow waters, may reach supersaturation with more soluble minerals, such as trona or halite, while relatively insoluble minerals such as aragonite and gypsum accumulate in deep water. An example is Lake Balkhash in the U.S.S.R. where gypsum and halite are accumulating in shallow lagoons along its isolated eastern arm while alkaline earth carbonates are precipitating in the deeper portions of the lake (Strakhov, 1970). Other examples include the Dead Sea (Neev and Emory, 1967), Lake Chad (Maglione, 1974,1980; Eugster and Maglione, 1979) and the Caspian Sea (Garbell, 1963; Dickey, 1968; Strakhov, 1970; Valyashko, 1972b). Perennial saline lake evaporites consist of 1) cumulus crystals that precipitate at the air-water interface and sink to the bottom of the brine, 2) crusts that precipitate on the brine bottom, 3) detrital evaporites, and 4) intrasediment crystal growths and cements. Cumulus crystals form by evaporative concentration of brine at the air-brine interface. Their mineralogy depends upon the chemistry of the inflow water, the evaporation/inflow ratio, the accumulation history of the solutes, and other factors (see section on chemistry and Eugster and Hardie, 1978; Spencer and others, 198Sa; Bowler, 1986). Evaporite minerals that precipitate at the air-brine interface are typically thin, platy, euhedral crystals (Neev and Emory, 1967; Hardie and others, 1985; Last, 1984) that commonly aggregate into floating crusts. The crystals or crystal aggregates float on the brine surface, held by surface tension, until they grow large enough to sink or until waves disrupt the surface tension. The evaporite accumulates on the lake bottom as a layer of loosely-packed crystals ("cumulates") that drape the topography. They may, however, dissolve while falling through the water column or while resting on the bottom. For example, Neev and Emory (1967) postulate that gypsum precipitated at the surface of the Dead Sea is consumed at the lake bottom by sulphate reducing bacteria producing calcite and iron sulfides, such as mackinawite and greigite. Crystal cumulus layers may be graded if the first crystals formed are large, subhedral skeletal crystals and smaller euhedral crystals precipitate later from highly supersaturated brines (see calcite crystals in Lake Zurich, Kelts and Hsu, 1978). Crystal layers with reverse grading may form if the period of time during which the crystals remain afloat becomes progressively longer as the surface brine becomes denser and more concentrated (Hardie and others, 1 985). Hardie and others (1978) describe the sequence of brine concentration and evaporite crystal precipitation in stratified perennial saline lakes based upon a model first proposed by Schmaltz (1969). Surface brine concentrated by evaporation becomes denser and sinks, mixing with the lower water mass. Further evaporative concentration

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increases the salinity and density of the entire water column. Therefore, the surface brine in the newly stratified lake must be concentrated even more before it will be dense enough to sink. If the evaporation rate exceeds inflow for extended periods of time, the saline minerals accumulated on the bottom will reflect the overall increase in brine salinity. A precipitation sequence involving minerals of increasing solubility, such as calcite - aragonite -gypsum -halite, will result. The sequence of evaporites that precipitate from a saline lake may be altered by changes in the evaporation/inflow ratio. If inflow increases, surface waters become more dilute, which could lead to 1) dissolution of evaporite deposits, if the entire water mass is mixed, or 2) isolation of the saline bottom brine. If inflow ceases, the lake brine will drop rapidly and precipitate a single layer of evaporites. The drought of 1930-1935 in the Great Salt Lake caused a severe lowering of lake level and precipitation of a halite layer up to 1.5 m thick directly over carbonate sediments (Eardley, 1966; Whelan, 1973). The halite layer was redissolved when lake level subsequently rose, but small lenses of halite remain in local depressions overlain by halite-saturated bottom brine (Spencer, 1982; Spencer and others, 1985). Neev and Emory (1967) describe precipitation of aragonite and gypsum from surface waters of the Dead Sea in Israel, deposited on top of a late Pleistocene or Holocene halite layer that formed when the lake level was lower. The dense lower water mass was saturated with halite at the time (196(Ys), but was isolated from evaporation by the more dilute upper water mass. A subsequent drop in lake level resulted in the destruction of the salinity stratification by 1975 and renewed deposition of halite and some gypsum on the lake bottom (Beyth, 1980). A man-made cut-off of inflow to Owens Lake, California resulted in the complete desiccation of the lake and precipitation of a 1-3 m thick trona + burkeite layer (Friedman and others, 1976; Alderman, 1985). Layering in crystal cumulates may be defined by different saline minerals or by variations in the abundances of saline minerals. These variations probably reflect episodes of surface water fresheningfollowed by evaporative concentration. Progressive concentration of the lake brine may produce successive laminae of evaporite crystals containing larger fractions of more soluble minerals (Fig. 3.6). Eugster (1969, 1980) suggested that magadiite layers in the Pleistocene Olorongo Beds of Lake Magadi formed by the mixing of freshwater runoff with alkaline lake brines. Layering may also be defined by intercalation of cumulate evaporites with flood-derived detritus or cumulate layers that punctuate the background accumulation of detrital and biogenic sediments. Evaporite crusts precipitate on the floors of perennial lakes from well-mixed supersaturated brines. Crusts are best developed in very shallow saline lakes or the

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I THE VARVED CLASTIC-ORGANIC-EVAPORITE CYCLE

ANN U AL L AM IN AT IONS STRATI GR AP WC FORMITION APPROXIMATE CLIMATE TIME REOUIRED CYCLES ldiograrnrnolicl SECTION iOH DEPOSITION

2-FOLD CYCLE CIOStIC ond gwsum

3-FOLD CYCLE btumi nous limestone, gypsum onc CIOStIC

4-FOLD CYCLE CIOStIC, orgonic, limestone ond gypsum

3-FOLD CYCLE SIOIIIC, organic, ond hmestone

2-FOLD CYCLE CIaZtIc ond \ organic

Fig. 3.6. Schematic drawing of stratigraphic sequence of cumulus crystal laminites in the Jurassic Todilto Formation, New Mexico. The vertical change upward to more soluble minerals and thicker laminae suggests increasing salinity and decreasing water depth. (from Anderson and Kirkland, 1960).

shallow portions of deep lakes. They form as overgrowths on cumulate layers or on older crusts. Crusts may also form on stable objects such as rocks or submerged vegetation. Crystals in evaporite crusts are typically vertically elongated and may show radial textures or competitive growth boundaries similar to crusts in saline pans. Hemispheroidal mounds or biscuits may result from crystal growth at local nucleation

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centers (Muller and Irion, 1969; Weiler and others, 1974; Warren, 1985). Thick crusts comprised of large crystals with growth bands may form in a meter or two of water (Warren, 1982 and Last, 1984), if growth is not interrupted by a large sediment influx or a major dilution episode. Magee (1988) describes millimeterscale,vertically oriented blades of gypsum (sedentary gypsum) growing from cumulate crystal laminae in cores through Holocene deposits of Lake Prungle in Australia. He determined that the brine was no more than a few meters deep from lateral relationships with shoreline deposits (pers. comm., 1988). Similar gypsum crystal layers are described from cores in Lake Eyre (Magee and others, 1988), Lake Frome (Bowler and Magee, 1988), and Lake Tyrrell (Bowler and Teller, 1986) in Australia. Evaporite crusts near the margins of perennial saline lakes may be partially dissolved by dilute surface brines, if lake level rises, or by rainfall andfloods, if lake level falls. The resulting fabrics may be indistinguishable from those in saline pan crusts. Salinity stratification may protect evaporite crusts from dissolution in the deeper portions of the lake, allowing them to be draped by detrital or chemical layers. The absence of dissolution fabrics and the development of layered sediment drapes over crystal terminations may be indicators of perennial lake saline crusts rather than saline pan crusts (Fig. 3.7A). The maximum water depth in perennial saline lakes in which evaporite crusts may develop is not known. In most modern perennial saline lakes, crusts form at water depths of no more than one or two meters. This may be the maximum depth to which supersaturated surface brines may sink without mixing. Beyth (1980) describes aggregates of halite and gypsum in the Dead Sea collected at depths of up to 260 m that apparently precipitated as bottom crusts. The crusts may have formed from sinking density currents of brine formed in the shallow South Basin or from subsurface brine that entered the Dead Sea along faults through salt of the underlying Lake Lisan Formation. Detrital deposits of evaporite minerals may be an important component of perennial saline lakes. In shallow-water areas of saline lakes, gypsum, mirabilite, and halite, may be reworked by waves into ripples or shoals (Fig. 3.7C; Neev and Emory, 1967; Weiler and others, 1974; Gwynn and Murphy, 1980; Last, 1984; Last and Slezak, 1987; Hardie and others, 1985; Karkz and Zak, 1987). Cumulus crystals, intrasediment crystals, or fragments of crusts are commonly abraded or rounded by such reworking. Evaporites may also be reworked by slumps, debris flows, and turbidity flows in perennial saline lakes, particularly those with steep edges to lower basins. The evaporite crystals in these deposits may be euhedral and dispersed in the sediment matrix. Blocks of evaporites with contorted bedding may also be included in these deposits (Fig. 3.7D).

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Fig. 3.7. A). Perennial lake halite crust comprised of' euhedral chevron cr.vstals (H) draped bv cumulus crystals of sylvite (S).Oligocene Rhine Graben, France. Scale =em. B). Pinch-and-swell lamination indicates intermittent wave deposition and layer selective euhrdral crystals of glauherite pseudomorphed by calcite are similar to intrasediment evaporites formed in a perennial lake sptting. Sample=6 cm thick. C). Sur,face of a gypsum Ia,yrr showing oscillatory ripple marks within Holocene perennial lake laminated clays, Lake Eyre, Autralia. Sample = 25 cnz long. D). Perennial lake laminite comprised of cumulus crvstals of' gypsnm (light) and aragonitr. Limn Formation, Dead Sea. Israel. Thickness = 2.5 m. (Photo courtesy of H. P. Eugster).

lrifrasediment crystal growth and cements composed of saline minerals may precipitate from perennial lake bottom brine, from interstitial brine, or from mixtures of the two (Fig. 3.78). Neev and Emory (1967) and Gornitz and Schreiber (1981) describe large skeletal halite crystals that encase laminated lake sediments in the shallow South Basin of the Dead Sea. They could not determine whether these crystals precipitated in a perennial saline lake or if they formed in a saline mudflat during a low lake stand. Spencer (1982) observed euhedral gypsum crystals near the sediment surface in cores taken from the bottom of the Great Salt Lake. These crystals were interpreted

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as modern intrasediment precipitates based on pore water chemistry. They apparently formed by mixing of surface brines and interstitial pore waters along a diffusion gradient. Spencer (1982; Spencer and others, 1985a) also noted angular deflections in disrupted laminae near the modern sediment surface which are interpreted to have formed by intrasediment growth of mirabilite during the winter that redissolved when the water warmed in the summer. He also observed impressions in laminae that resemble gypsum swallow tail crystals and halite hoppers which he attributed to dissolution of cumulate crystal beds. Saturated bottom brines may also alter and cement evaporite beds and convert them into coarsely crystalline mosaics. This is suggested by halite layers in the Great Salt Lake that are known to have formed from surface precipitation, but are now coarsely crystalline and cemented (see Gwynn and Murphy, 1980,fig. 8). Similar cement overgrowths may account for the crystals of halite in the Dead Sea (Beyth, 1980) and mirabilite crystals in Freefight Lake (Last and Slezak, 1987). Diffusion of lake bottom brines into the sediments may result in the formation of authigenic clay minerals and silicates, including illite, stevensite, kerolite, nontronite, and sepiolite (Gac and others, 1977; Jones and Weir, 1983; Spencer, 1982; Yuretich and Ceding, 1983; Von Damm and Edmond, 1984; Spencer and others, 1985b; Jones and others, 1986). These clay minerals may be the only indication of evaporitic conditions in some perennial lake deposits. Non-evaporite sediments in perennial saline lakes are similar to deposits of freshwater lakes (Fouch and Dean, 1982; Dean and Fouch, 1983; Eugster and Kelts, 1983; Allen and Collinson, 1986). Commonly saline lake deposits are interlayered with freshwater lake deposits at scales as small as a few tens of centimeters, reflecting lake level fluctuations. Saline lake fabrics may be superimposed on the freshwater deposits, which may complicate the interpretation of saline and freshwater episodes in lake deposits. Spencer and others (1984) recognized increases in lake salinity related to lake level fall in cores from Great Salt Lake by changes in the biological component (ostracodes - brine shrimp -algal mats) and in the carbonate mineralogy (calcite - Mgcalcite - aragonite). Sedimentary structures in perennial saline lake deposits reflect settle-out of detrital material and chemical precipitates, gravity flows, wave action, desiccation, and biological activity. Settle-out sedimentation may produce fine, continuous laminae, particularly in the deep portions of stratified lakes where the sediments are protected from wave reworking, bioturbation, and fluidization from slumping or sudden influxes of large amounts of sediment (Kelts and Hsii, 1978). The lamination may reflect seasonal conditions so that couplets or triplets represent annual increments (varves)

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(Kelts and Hsii, 1978; Anderson and others, 1985; Anderson and Dean, 1988) or the laminae may reflect periodic or aperiodic non-annual events such as flash floods or droughts (see Calvert, 1966, explanation for clastic mud-diatom laminae in the Gulf of California, also Neev and Emory, 1967, p. 93). The continuity of layers is a function of the sediment dispersal mechanisms (Sonnenfeld and Hudec, 1985, for influence of chemocline) which may be the same for annual and non-annual layers. In shallow water, settle-out deposits tend to have thicker, more irregular lamination or they are massive as layering is disrupted by the action of waves, organisms, or fluidization. Rapid accumulation of flocculent pelagic sediment may produce syndepositional compaction and dewatering, resulting in a structureless deposit or diffuse thin bedding (thick lamination) defined by variations in organic content or grain size. Fine lamination may be preserved in shallow perennial saline lakes when bound by microbial mats (Bauld, 1981a,b). Spencer (1982) describes shallow water laminites in cores from Great Salt Lake, Utah, which contain layers with abundant organic filaments that are interpreted as former microbial mats. Sediment gravity flows may occur in perennial saline lakes by slumping or collapse of sediment accumulations in shallower water or by density flows induced by rapid sediment influxes from rivers or by wave-mixing of sediment during storms. Turbidity flows from slumping or sudden influxes of sediment produce laminae to decimeter-thick, muddy graded beds with load structures at the base that may be internally structureless or that may show complete "bouma" sequences (Bouma, 1964). Waves rework detrital, biological, and chemical sediments producing characteristic structures that depend on water depth. Wave action in relatively deep water produces thin, continuous silt laminae that pinch-and-swell formed by the weak wave stress (rolling ripples of Harms and others, 1982). In shallower water wave action may produce thicker, more lenticular laminae and thin beds of sand, reflecting accumulations of sand in long crested ripples (Fig. 3.8A). The coarse-grained ripple beds may alternate with settled-out mud layers or may only have muddy lenses in the troughs similar to flaser beds. Storm waves may produce scour and fill structures, intraclast beds of crusts and sediment, and graded sand beds with flat or wavy lamination. Wave-formed deposits in shallow water may be interbedded with evaporite crusts or cemented into pavements by evaporites or carbonates (see aragonite and dolomite crusts in Great Salt Lake, Spencer, 1982; Lake Urmia, Kelts and Shahribi, 1986; and Freefight Lake, Last and Slezak, 1987). Biological activity is more diverse in freshwater lakes than in saline lakes, but saline lakes may be highly productive biologically (Evans and Kirkland, 1988). In fresh or brackish lakes, tests or shells of organisms may be an important source of sediment

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Fig. 3.8. A). Pinch-and-swell lamination with sandy layers (light) comprised of brine shrimp fecal pellets in core from Great Salt Lake, Utah. Sample=7..5 cm wide. (Photo courtesy of B. F. Jones). B). Saline pan surface of Salar Uyuni, Bolivia. Pan surface is halite, note polygonal pattern of ridges. (Photo courtesy of B. F. .lanes). C). Halite crust from saline pan of Saline Valley, California. Note f lat parting surf ace and vertical voids produced by dissolution during f lood stages. Thick clay bed (dark) was deposited during a flood stage. Sample = 10 cm thick. (Photo courtesy of E. Casas), D). Trona crust from saline pan at Lake Magadi, Kenya. Note radial crystal arangement and f latpartings with clay. Sample = 20 cm thick. (Photo courtesy of H.P. Eugster). and a significant component of layering. In several modern saline lakes the fecal pellets of brine shrimp are an important sediment component (Eardley, 1938; Spencer, 1982; Kelts and Shahribi, 1986). Algae may contribute significant quantities of organic material and may also form mats. In some marginally saline lakes in Australia, the calcified oogonia of Charophyte algae are an important sedimentary component and calcified algal mounds (tufa) are common in saline and brackish alkaline lakes (Eardley, 1938; Carozzi, 1962; Halley, 1976; Spencer, 1982). Burrowing by a variety of organisms may disrupt bedding. Burrows are

Data Center ,09126599985,[email protected], For Educational Uses 212 NON-MARINE EVAPORITES tube-shaped disruptions with constant diameters that may be randomly oriented or restricted to vertical or horizontal planes. The tubes may be simple or complexly branching or patterned. Where burrowing is intense it may be indicated by a mottled pattern consisting of circular or ovate cross-sections. We do not know of any literature on the tolerance of burrowing organisms to salinity so we cannot comment on their significance. Plant growth at the margins of a lake or plants that have been inundated by a rising lake may be recorded by root structures (branching, tapering, tubes that have a variety of diameters) or by trunks and stems coated by algal tufa crusts or evaporite crusts. Desiccation of perennial lake sediments may be important in shallow lake deposits, particularly if the lake is in a broad, flat-bottomed basin. Subaerial exposure of fine-grained perennial lake sediments will result in the development of polygonal mudcracks. For instance, Currey (1980) postulates that the large polygonal features visible in the shallow water of Great Salt Lake are desiccation polygons formed during a lake low stand. Relatively simple V-shaped polygonal cracks form during short periods of exposure, whereas prolonged exposure may result in the development of soil features or features similar to those of dry mudflats or saline mudflats (see descriptions under those headings). Deep, broad cracks will form in thick accumulations of fine-grained subaqueously deposited sediment when subaerially exposed. Shorter, narrower cracks will form if desiccation occurs repeatedly with only thin accumulations of sediment. Short periods of desiccation may allow deep, narrow cracks to form while the sediment is still very soft. If lake level rises these cracks may remain soft and later compact producing features similar to fluidization structures. Subaqueously formed polygonal cracks (synaeresis cracks) have been reported (Rooney and others, 1969), but their origin never documented. Perennial saline lake deposits commonly reflect variations in lake depth and salinity. Deep, stratified perennial saline lakes form flat lamination in the center and more lenticular and irregular layering towards the margins. Evaporites in the center are limited to cumulate crystal layers and gravity flow deposits. Towards the edges, evaporite crusts and intrasediment evaporites are more common. In shallow saline lakes, crusts and intrasediment growths may occur across the entire bottom. Evaporites may occur as intrasediment growths and crusts in small lagoons while the bulk of the perennial lake deposit is evaporite free. Desiccation cracks may be common near the margins of shallow saline lakes in broad basins. Changes in lake depth and salinity may result in intrasediment evaporites forming in freshwater deposits and desiccation features in deep water deposits. A change from a shallow, stratified saline lake to a deep, well-mixed brackish lake may be indicated by a laminated sediment grading into

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Saline Pun Subenvironment

A saline pan (Lowenstein and Hardie, 1985) is a normally dry area underlain by layered salt (Fig. 3.8B). Our usage of saline pan is equivalent to the ephemeral saline lake salt pan of Hardie and others (1978). Saline pans form in non-marine and marine settings. Non-marine saline pans characteristically occupy the lowest areas of closed drainages and are surrounded by saline mudflats with efflorescent salt crusts. Saline pans vary in size from tiny pans less than 1 km’ (for instance the Basque Lakes, British Columbia are .008 km’, Nesbitt, 1974) to gigantic features covering thousands of km’ (for instance, the Lake Uyuni salt pan is 9,000 km’, Rettig and others, 1980). Saline pan deposits are stratified and consist of evaporite crusts separated from one another by solution surfaces or by mud layers (Fig. 3.8D). Saline pan crusts are most commonly composed of halite (for instance, Stoertz and Erickson, 1974; Lowenstein and Hardie, 1985; Bowler and others, 1986), but may also be composed of trona (Eugster, 1970), gypsum (Stoertz and Erickson, 1974), mirabilite (Last, 1984), epsomite or bloedite (Nesbitt, 1974; Last, 1984), and thenardite (Jones, 1965; Strakhov, 1970). Mud interbedded with salt crusts commonly contains intrasediment evaporite crystals. Lowenstein and Hardie (1985) describe a saline pan cycle (Fig. 3.9) to interpret the formation of thin-bedded halite crusts and interlayered mud. The cycle consists of: 1) a flooding stage, 2) a saline lake stage, and 3) a desiccation stage. Similar conditions of formation are suggested for gypsum crusts (Warren, 1982; Castens-Seidell, 1984; Bowler and Teller, 1986), trona crusts (Eugster, 1970), and mirabilite crusts (Last, 1984). Flooding Stage. Floodwaters inundating a saline pan may form a shallow temporary brackish lake. The area of flooding is commonly larger than the area of the saline pan. For instance, Lake Eyre expanded to 8000 km’ duringflooding in 1950 and evaporated to dryness by 1952 leaving a salt crust over an area of 800 km’ (Bonython, 1956). Flooding may be seasonal or can be sporadic such as the Lake Eyre saline pan is flooded on the average of every 8 years (Kotwicki, 1986). Fine-grained sediment introduced by flood water is deposited as thin layers that are thickest near the flood source areas. Rapid sedimentation of suspended clay is aided by flocculation (Last, 1984; Sonnenfeld and Hudec, 1985). Most solutes in saline pan ephemeral lakes are derived from 1) dissolution of the most soluble salts of efflorescent crusts (i.e. halite, thenardite, natron, etc.) by dilute surface runoff and brackish lake water, and 2) partial to complete dissolution of the previous saline pan crust. Therefore, the solutes in saline pan brines may be dominated by the most soluble “recycled“salts.

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THE SALINE PAN CYCLE STAGE 1. FLOODING

Brackish Lake Floodwaters+-.. 111~ - -cyFloodwaterS saline mudflat deposits

Processes dissolution of salrne crust mud layer followed by deposition of mud dissolution surface on saline crust STAGE II. EVAPORATIVE CONCENTRATION dissolution vugs and piper

I evaooration evaDoralion ,

\ nucleated rafts and Processes subaqueous crystallization hoppers and growth of salts in the saline lake SYnlaxlal overgrowth on rafts and hoppers settled STAGE 111. DESICCATION at lake bottom Dry Pan evaporation surface crust evaporation 1

ground water brine Puffy efflorescence growing ~n \ iractur& of tne pressure ridges Processes diagenetic growth of Salts \ Of polygons within salirie and mud layers growth of beneath the dry pan surface saline cements and disruption of Saline crust ~n lntetgranuler into polygons wolds and dissolution vugs

Lowenstein and Hardie (1985) describe dissolution features in saline pan halite crusts produced during the flooding stage. These include centimeter-scale, vertically-oriented solution pipes at crystal boundaries, horizontal solution caverns, and rounded crystal faces (Fig. 3.10A). Warren (1982, 1985) describes rounding and planing of gypsum crystals in saline pan crusts in Marion Lake, Australia. Eugster (1980) describes leveling of trona crusts and dissolution of trona crystal layers during flooding of the Lake Magadi saline pan. Last (1984) describes mirabilite crusts from Ceylon Lake, Saskatchewan that, during flooding events, have undergone complete dissolution or partial dissolution leaving hummocky remnants.

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Saline Lake Stage. Afterflooding, the saline pan lake water is concentrated by evaporation. Complete evaporation of the ephemeral lake can occur within days in small, shallow lakes or can take years in large, deeper lakes. The earliest precipitates from the brine are small crystals that nucleate at the water-air interface. The crystals may be flat platy euhedra of halite (Lowenstein and Hardie, 1985), gypsum lathes (Schreiber and Kinsman, 1975; Castens-Seidell, 1984), or radial crystallites of natron (Alderman, 1985). The crystals may sink to the bottom and form a loosely packed cumulate layer. Those crystals that continue to float on the brine, held by surface tension, develop into fluid inclusion-rich, hopper-shaped crystals or floating rafts of crystals that eventually sink to the bottom (Fig. 3.10B). Some of these crystals, along with partially dissolved remnants of the underlying crust, may be reworked by waves into flat layers, ripples, or shoreline bars (Stoertz and Erickson, 1974; Castens-Seidell, 1984; Last, 1984). Continued concentration of the brine by evaporation together with mixing of the shallow lake water eventually produce supersaturated brine from which crystals may grow on the lake bottom. For halite, syntaxial growth on earlier formed crystals and competitive growth of crusts produce vertically oriented chevrons with growth bands defined by fluid inclusions (Fig. 3.1OC; Shearman, 1970; Lowenstein and Hardie, 1985). Radial sprays of vertically-oriented crystals have been observed in trona crusts (Eugster, 1970,1980),gypsum (Arakel and McConchie, 1982;Warren, 1982;Castens-Seidell, 1984; Bowler and Teller, 1986), thenardite (Strakhov, 1970, p. 259), and artificial borax crusts (Bowser and Dickson, 1966). Mirabilite-thenardite crusts in the Ceylon Lake saline pan form dendritic to reticulate masses of radiating bladed crystals (Last, 1989). Monomineralic salt crusts may form if the saline pan lake waters are dominated by solutes derived from dissolution of earlier formed efflorescent crusts and saline pan layers (Hardie, 1968, 1984; Eugster, 1970). Alternatively, several minerals may precipitate as crusts during evaporation of the brine. In this case, as the lake area decreases, a bull's eye pattern may develop with the most soluble minerals in the center of the lake. The crusts formed from the last remaining surface brine in the center of the lake may overlie a cumulate layer of less soluble salts (see descriptions of Arakel, 1980; Castens-Seidell, 1984) that grades outward to a crust of less soluble minerals. Alderman (1985, fig. 3) illustrates that successive floods over the saline pan of Owens Lake, California, produced a concentric pattern of trona-burkeite-halite which also appears as vertical successions. Hunt and others (1966) and Stoertz and Ericksen (1974) envisioned a similar process for the zonation of gypsum and halite crusts in Death Valley and several salars in Chile. Some of the mineral zonation, however, may be due to the development of saline mudflat efflorescent crusts or to saline pan crusts of differing ages. Last (1984, 1989) notes that crusts with acicular crystals of bloedite and epsomite

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. .. . ~

Fig. 3.10. A). Thin-section o,f halite crust from saline pan nt Raja California, hlexico. Flood stage dilute water produces the flat solution surf ace and vertical tithes (open porosity is black). Thickness = 3.5 cm. (Photo courtesy of E. Casa.~).B). Paper-thin rclft,s of mirabilite crystals ,flouting on surfucp o,f Ce,vlnn Lake. Saskutchewun. The.w forni cumulate depositsupon sinking. (Photo courtesy o,f W.M. Last). C). Thin-section0.f' halite crust overlying a cumulus crystal layer from saline pan at Buja Calif'ornin. M erico. Growth bands in halite chevrons are defined b.yf'luid inclusions (open poro.sit,y is black). Sample =S.Scm wide. D). Thin section of clear halite cr.vstalsgrowing into solution cuvity of a saline pan halite crust from Baja California, Mexico (openporosity is white). Sumple is from about 10 cm below the present surface. Scale =5 mm long. (Photo courtesy of E. Casas). occupy the central zone of several saline pans in Saskatchewan. Those crusts form during the final stages of brine evaporation of the ephemeral saline lakes. They overlie an earlier formed mirabilite crust. Jones (1965) illustrates a layered saline pan crust of thenardite, burkeite, and halite in Deep Springs Lake, California. Desiccation Stage. The final stage of the saline pan cycle is the complete desiccation of the ephemeral saline lake. During this stage the surface brines sink below

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the surface salt crust and continue to concentrate by evaporation (evaporative pumping?). Lowenstein and Hardie (1985) and Casas and Lowenstein (1989) describe clear halite overgrowths and euhedral cavity-filling crystals of halite that precipitate during the desiccation stage in porous halite crusts immediately below the surface (Fig. 3.10D; Shearman, 1970; Strakhov, 1970). Eugster (1980) mentions that trona precipitation in Lake Magadi continues after the brine drops below the surface crust during the desiccation stage, and that the porosity is reduced, probably by cementation. Dense concentrated brine that sinks below the crust surface may also precipitate more soluble minerals within the pore spaces, if concentration by evaporation continues or if the brines cool (Strakhov, 1970; Valyashko, 1972a; Last, 1989; Lowenstein and others, 1989; Spencer and others, 1990). Halite crystals precipitate in mud layers between salt crusts during the desiccation stage as isolated euhedra or randomly oriented crystal aggregates. Mud is commonly poikilitically encased in the cores of crystals, whereas the outer edges of crystals are clear and form by displacive growth (Handford, 1982a,b; Lowenstein and Hardie, 1985: Casas and Lowenstein, 1989). Mud near the margins of some saline pans in Saskatchewan contains equant crystals of mirabilite with muddy cores that grew in the sediment (Last, 1984, 1989). Surface brine of the saline pan may sink and mix with groundwater creating a hybrid brine different in composition from that produced by evaporative concentration of either water (Hardie, 1968). Evaporative concentration or cooling of this hybrid groundwater brine may produce evaporite minerals or recrystallization of older evaporite minerals. For instance, in Saline Valley, California, thenardite and glauberite crystals with muddy cores and clear edges together with halite form interlocking aggregates in mud partings and cavities of the saline pan halite crust. Hardie (1968) argued that the glauberite replaced previously formed gypsum by reaction with NaCI-saturated groundwater brine. Gaylussite in sediments of the Lake Magadi saline pan may have a similar origin (Eugster, 1980). Crystal growth in saline pan crusts during the desiccation stage also leads to the development of polygonal fractures with overthrust edges due to the expansion of the crust (Fig. 3.1 1B; Lowenstein and Hardie, 1985). The polygonal fractures range in size from meters to tens of meters in diameter (Hunt and others, 1966; Stoertz and Ericksen, 1974). Expansion polygons are most widely reported from halite crusts but have also been observed in gypsum crusts (Castens-Seidell, 1984), trona crusts (Eugster, 1980),and thenardite-burkeite crusts (Jones, 1965). Evaporation of subsurface brines preferentially along the polygonal fractures results in the precipitation of finelycrystalline efflorescent crusts that further widen the cracks and upturn the polygon margins into a series of

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Fig. 3.11. A). Saline mudflat at Salar tirande, Bolivia, covered with an irregular sediment-rich halite eff lorescent crust. (Photo courtesy o,f G.E. Ericksen). B). Polygonal ridges of efflorescent halite in saline pan halite crust at Saline Valley, California. The.st. ridges formed during a long period of desiccation. (Photo courtesy of L.A. Hardie). C). Laminated gypsum crust from a probable relict saline pan at Saline Valley, California. Irregular humpy layers consist of vertically oriented gypsum crystals and pockets of gypsum cement. Finely crystalline gypsum occurs as clasts in sand,y interbeds. Dark layer is clay from a major flood stage. Exposure = 45 cm thick. (Photo courtesy of B. Castens-Seidell).

dish-shaped lenses separated by contorted ridges (Eugster and Hardie, 1978; Lowenstein and Hardie, 1985). The efflorescent ridges may contain significant amounts of more soluble salts such as carnallite (Sebkha Marada, Libya; Desio, 1970; and Lake Inder, U.S.S.R.; Strakhov, 1970).The efflorescent ridges are dissolved by subsequent flooding (which contributes solutes to the ephemeral saline lake waters), but the cracks may be filled with mud or crystals which preserve the polygonal pattern (Lowenstein and Hardie, 1985). Castens-Seidell (1984) describes precipitation of large gypsum crystals in polygonal fractures of gypsum pan crusts from Baja California, Mexico. Blocky prisms of gypsum grown downward from overthrust gypsum crusts form pendant cements, and

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blocky hemipyramids of gypsum form mounds of loose crystals in the cracks. Saline pans commonly contain groundwater spring seeps which may produce cone-shaped dissolution cavites in saline pan crusts (Jones, 1965; Hunt and others, 1966; Last, 1984, 1989). Evaporites may precipitate in the solution cavities as isolated pods that crosscut the layered saline pan deposits. For instance, halite and carnallite precipitate from thermal springs in the Danakil Depression, Ethiopia (Holwerda and Hutchinson, 1968), and epsomite and bloedite precipitate from bottom seeps in the Basque lakes, British Columbia (Nesbitt, 1974) or Ceylon Lake, Saskatchewan (Last, 1989). Spring inflow mixing with saline pan brines may also produce local pods of evaporites which differ in mineralogy from the saline pan crust. Hardie (1968) reports a zone of thenardite within the halite crust of Saline Valley, California, which he attributes to local spring mixing. Repetition of the saline pan cycle produces thin-bedded salt crust deposits. Evaporite crystal beds are separated by crystal-rich mud partings, thin organic films, and/or dissolution surfaces. The boundaries between crystal layers may be flat or highly irregular depending upon the degree and manner of dissolution during the flood stages. Saline pan crusts are thinner and generally more patchy in distribution near the outer margins of the pan where they typically grade into saline mudflat deposits. Thin-bedded saline pan deposits may be disrupted or obscured by continued growth of evaporite crystals below the surface that displace sediment and fill porosity (Casas and Lowen- stein, 1989). G.E. Ericksen (USGS, pers. comm., 1987) noted giant halite crystals, up to 75 cm on a side, that encase thin-bedded halite in a quarry pit through the saline pan deposit of Salar Grande, Chile. Thin gypsum pan crusts are fragile and are commonly disrupted by desiccation and buckling of the underlying sediments, by wind deflation, and by reworking during the flood stages (Stoertz and Ericksen, 1974; Castens-Seidell, 1984). These fragmented gypsum pan crusts may be deposited as clastic gypsum beds (Fig. 3.11C) that are flat-bedded or cross-laminated and that are interlayered with irregular hummocky crusts composed of vertically oriented gypsum crystals (Stoertz and Ericksen, 1974; Neal and others, 1983; Castens-Seidell, 1984; Bowler and Teller, 1986). Saline pan evaporites may accumulate to appreciable thicknesses. For example, there are 300 m of halite saline pan deposits in Death Valley, California (Hunt and others, 1966); 40 m of saline pan trona in Lake Magadi, Kenya (Eugster, 1980); 25 m of possible saline pan gypsum in Trinity Lake, New Mexico (Neal and others, 1983), and 45 m of possible saline pan mirabilite in Ceylon Lake in Saskatchewan (Ruefel, 1968; Last, 1984). In many modern saline pans, however, the salt crusts are completely dissolved duringfloods and the accumulated sediments consist of structureless mud with euhedral, intrasediment evaporite crystals (Teller and others, 1982; Last, 1984).

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Preservation of saline pan evaporite deposits depends upon the chemical maturity of the pan and the stability of the groundwater table beneath the pan. Immature saline pan crusts completely dissolve during each flood stage. However, the solute concentration in groundwater beneath the saline pan progressively increases during each flood stage as brine derived from dissolution of efflorescent crusts sinks into the water table. The more saline groundwater produces thicker efflorescent crusts over a greater area. The dissolution of these readily soluble crusts during saline lake stages creates progressively more saline brines that supersaturate with respect to the evaporite mineral a little earlier. Eventually the saline pan crust in the lowest areas of the pan will no longer dissolve completely during the flood stage and a new crust layer will precipitate over it. The next flood-stage will increase the area of saline pan crust preservation, and so on through time. This mechanism also forces the area of mixing between the saline pan brine and groundwater to migrate outward parallel to the saline pan boundaries (Eugster and Hardie, 1978), and may ultimately produce brines capable of precipitating more soluble salts, such as potash. The initial potash precipitates apparently occur as pore-filling cements formed beneath the salt pan surface during the desiccation stage (Valyashko, 1972a; Strakhov, 1970; Yuan and others, 1985; Lowenstein and others, 1989). They may later form layered crusts, when surface brines reach saturation with respect to the most soluble salts.

Saline Mudf lat Subenvironment

Saline mudflats are narrow fringes or broad extensive flats of fine-grained sediment commonly marginal to perennial saline lakes and saline pans. Saline mudflat deposits consist of wet plastic clay to firm sandy mud in which intrasediment evaporites and surface efflorescence salt crusts form from groundwater brines (Fig. 3.1 1A). Saline mudflats aggrade by a combination of fluvial, lacustrine, and aeolian processes. They may also develop on older desiccated perennial lake or ephemeral lake deposits. The types of evaporite minerals formed in saline mudflats depend upon 1) the stage of chemical evolution of groundwater brine, 2) sinking of surface brine, and 3) the mixing of groundwater brine and surface-derived brine. The groundwater in saline mudflats evolves in conjunction with the adjacent saline pan or perennial saline lake. For example, gypsum is the only intrasediment mineral reported from the saline mudflats of most Australian playas (Bowler, 1986), and halite crusts of the saline pans are typically very thin. This apparently reflects the chemical "immaturity" of the groundwater systems in these lakes. Many saline mudflats display roughly concentric zones of intrasediment-grown evaporite minerals (Fig. 3.12). More soluble minerals

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Saline Valley Playa I d Carbonates (?) -2 KM T -r -7.- Gypsum -N- . ;,. . , :. .- ,..- Gypsum and Glauberite

Glauberite

Glauberite and Halite

A Halite

0 Halite and Mirabilite

Mirabilite and Glauberite 0 No salts

Fig. 3.12. Distribution of intrasediment evaporite minerals in Saline Valley, California. Concentric pattern of most soluble minerals in the saline pan surrounded by progressively less .soluble minerals in the saline mudf lat reflects progressive concentration of the groundwater. Sulfate-rich zone in the northwest corner is related to an artesian spring outlet. (from Hardie, 1968). occur toward the lowest central portion of the mudflat (Jones, 1965; Hardie, 1968). This zonation has been attributed to lateral concentration gradients of groundwater produced by progressive evaporative concentration and mixing of groundwater with saline surface brine. Lateral zonation of zeolite minerals in tuffaceous lake beds and of authigenic clays in lacustrine deposits also may result from lateral groundwater concentration gradients (Sheppard and Cude, 1968, 1969; Hay, 1966, 1970; Hay and others, 1986). Local sources of groundwater or local spring inflow may produce saline minerals that differ from the "normal" concentrically zoned mineral assemblage (for example the mirabilite/thenardite zone in Saline Valley, Hardie, 1968; or the distribution of borates in Teels Marsh, Nevada, Papke, 1976). Although the chemical controls on evaporite minerals in saline mudflats are fairly well known, the controls on their morphology and dktribution are poorly understood. Saline minerals grow in brine-soaked sediment of saline mudflats by displacive growth, where surrounding sediment is pushed aside by the force of crystallization, and by

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incorporative growth. The evaporite minerals display a variety of crystal habits. Gypsum commonly occurs as discoidal hemipyramids (millet seeds) (Fig. 3.13; Eardley and Stringham, 1952; Gwynn and Murphy, 1980; Arakel and McConchie, 1982; Bowler and Teller, 1986), but may also form bladed hemipyramids (Handford, 1982 a,b). Both gypsum crystal habits may form twinned rosettes. Cody (1979) attributed the discoidal morphology of gypsum crystals to growth inhibition of certain crystal faces by organic material. The confining pressure due to the surrounding mud has also been cited as the cause of parallel crystal growth (Shearman, 1978; Teller and others, 1982). Large gypsum crystals containing bands of sediment inclusions that alternate with clear gypsum (Draper and Jensen, 1976; Arakel, 1980) are thought to indicate alternating periods of solution and precipitation. Similar growth bands observed in halite crystals from the Dead Sea (Neev and Emory, 1967; Cornitz and Schreiber, 1981) were attributed to rapid crystal growth. Handford (1982a,b) describes halite crystals grown in the saline mudflat of Bristol Dry Lake that have muddy cores and depressed hopper-like faces which he attributed to slow, preferential growth of crystal edges over crystal faces. Southgate (1982) produced dendritic halite crystals at the brine-sediment interface experimentally by adding humic acid to the precipitating solutions. Mirabilite crystals precipitate in the saline mudflat of Saline Valley, California in the winter but leave molds when they dehydrate in warmer periods. The molds occur as large delicate dendrites in dense laminated clay layers and as barrel-shaped euhedra in porous silty clays. Thenardite may occur in mirabilite molds as a powdery residue from dehydration of the mirabilite crystals or it may grow in the sediment as barrel-shaped crystals. Glauberite in Saline Valley and Death Valley, California forms flattened wedge-shaped crystals and rosettes. Caylussite crystals in Deep Springs Lake, California occur as flattened wedge-shaped crystals up to 2 mm long within brine saturated carbonate sediments adjacent to the saline pan (Jones, 1965). Anhydrite forms intrasediment nodules of finely crystalline material that apparently form by dehydration of gypsum crystal masses (Moiola and Glover, 1965; Castens-Seidell, 1984). An intrasediment nodular habit is also common for ulexite (cottonballs) (see references in Barker and Lefond, 1985; Smith, 1985; Norman and Santini, 1985) and for magadiite (Fig. 3.13C; Eugster, 1970; Maglione, 1980). Evaporite minerals in saline mudflat sediments occur in layers or as isolated crystals and crystal aggregates. Randomly distributed euhedral crystals with muddy cores form in brine-saturated mud on the margins of saline pans and perennial saline lakes (Neev and Emory, 1967; Gwynn and Murphy, 1980; Gornitz and Schreiber, 1981; Arakel and McConchie, 1982; Dulhunty, 1982; Lowenstein and Hardie, 1985). Localized pods and lenses of evaporites also occur in brine-soaked mud. In places, pods of qstals grow

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Fig. 3.13. Phreatic zone intrnsediment gypsiini in the saline mudflat. A). Brine-saturated smndv mud containing abundant discoidnl gypsum crystals in the saline mudflat at Saline Valley, California. Sample thickness = 2.5 cm. B). Thin .section of discoidalgypsum crystals from the saline miinflat at Lake Tyrrrll, .4ustrmlia. The loss of the tapered ends of some crystals suggests they have been transported. Field of view = 2.5 cm wide. (Photo courtesy of.1.T. Teller). C). Nodules and enterolithic beds of magadiite (white) in the Pleistocene Ooloronga Beds of Lake Magadi. Kenya. Thickness = 70 cm. (Photo courtesy of H.P. Eugster). D). Trench in saline mudflat, Pastos Grnndes, Bolivia, where spring brine seeps into the surface muds. White nodu1e.s near the surface ore ulexite in a surf ur-stained mud. Laminated srdiments below are perennial leke deposits pnrtiallv disrupted by gypsum precipitated in the modern saline mudflat. Pocket knife (dork) = I0 c'tn long. (Photo courtesy of H.P. Eugster). as void-filling cements, such as gypsum or glauberite crystals filling mirabilite crystal molds in Saline Valley, California. In other cases, the reasons for localized growth of

Data Center ,09126599985,[email protected], For Educational Uses 224 NON-MARINE EVAPORITES crystal pods are not known. Pod-like lenses (as much as 30 cm thick) of loosely packed euhedral glauberite crystals are common in the saline mudflats of Saline Valley and Death Valley, California. The glauberite is apparently formed by replacement of gypsum (Hardie, 1968) and its distribution may reflect this process. Where saline mudflat sediment is only intermittently saturated with groundwater brine, crystal growth is localized in porous areas such as sandy layers or mudcracks (Arakel and McConchie, 1982). Crystal growth may extend away from these porous areas disrupting the surrounding sediment. Spring seeps may produce pockets of coarse crystals or localized beds of evaporite minerals that reflect brine mixing (Teller and others, 1982; Qian and Xuan, 1985). Patchy occurrences of ulexite nodules, concentrated in zones 10-30 cm below the saline mudflat surface near the saline pan, have been attributed to local groundwater sources such as hot springs or older borate deposits (Fig. 3.13D; Smith, 1985). A deeper groundwater table on the outer margins of saline mudflats leads to the development of vadose zone evaporites and modification of minerals previously formed in brine-saturated sediment. Vadose zone evaporites are typically finely crystalline and confined to porous layers or fenestrae (Fig. 3.14B). The crystals may exhibit dissolution features from intermittent exposure to dilute water, such as rainfall. The evaporites, including crystals formed in "phreatic"conditions, may become dehydrated (i.e. gypsum to anhydrite) after long periods in the vadose zone. Dry mud with "vadose"evaporites may overlie wet mud with "phreatic"saline minerals (Fig. 3.14A). The thickness of the dry-zone fabrics increases away from the saline pan or perennial saline lake, as the depth to brine-saturated sediment increases. In very dry portions of saline mudflats, percolation of rainfall through efflorescent crusts and subsequent desiccation may produce powdery nodules or thin crusts of saline minerals that are similar to saline soils. The depth to which evaporite minerals can form by evaporation of groundwater brine is unknown, but depends upon the porosity of the sediment, the intensity and variability of evaporation, and the salinity of the groundwater. Hardie (1968) determined that the assemblage of evaporites in equilibrium with groundwater brine in Saline Valley, California extends to a depth of at least 4.8 m but probably no more than ten meters. Saline Valley, however, may have a shallow modern groundwater system that overlies an older equilibrium groundwater and evaporite mineral system. Under such conditions, saline minerals and vertical variations in mineral textures and fabrics may be destroyed or modified by later groundwater brine. Descriptions of cores through modern saline mudflats do not mention vertical variations in crystal habits reflecting vadose or phreatic conditions (Smith and Pratt, 1957; Smith, 1979; Smith and others, 1983; Bowler and Teller, 1986). Therefore, it is difficult to evaluate whether textures

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Lacy halie c~st A Gray mud Brown mud lolauberite cwstals subhedrall Green Clay Bmwn mud Bmwn clay wiih glauberile-filledpolygonal cracks Brown mud Gray clay Three fine layers, fmrn top to bottom, 1 of brown, while, and green clay Brown mud

Brown clay with grayish bands

Brown mud with clay patches near base

a Well defined patch of coarse glauberite (1-2 mm) a Well defined patch of fine glauberile (< lmm) ...... -..p..: Poorty detined patch of fine glauberite

:: Fine glauborile crystals F+ Glauberie msenes (large and small) 00 Mirabilite crystal molds (bkky. large and small)

A Mirabilite crystal molds in clay (fern-likefrom large skeletal crystals)

I Mirabolite crystal mold filled with glauberile cement

Fig. 3.14. A). Distribution of evaporite textures in a trench in the saline mudf lat at Saline Valley, California. Evaporite crystals are coarser and more abundant in the brine-saturated muds in the lower part of the trench, and finer and more layer selective in the drier upper part. Remnants of bedding in the lower part of the trench are separated by patches of crystals. Crystals at the top of the trench are rounded by dissolution. Clay layers formed in sheet deltas or shallow lakes. (Based on f ield notes and photos of B. Castens-Seidell and J.P. Smoot). B). Vadose zone evaporites from the outer part of saline mudflat at Saline Valley, California. Top of picture is the base of a halite efflorescent crust. Note the irregulargradational base. Light layer at the base of the crust is crystalline halite cement formed by rain percolating through the eff lorescent crust and ponding at the mud contact (several small remnants of similar cement layers occur in the underlying sediment). Small blebs of f inely crystalline gypsum and halite (white specks) are restricted to silt lenses (s), mudcracks (m), and roots (r). Silt lenses are windblown sediment that was trapped in efflorescent crust irregularities. Thickness = 2.5 cm. produced by surface variability of groundwater brine are preserved. Bowler and Teller (1986) observed several layers with displacive gypsum crystals separated by layers of

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laminated gypsum or gypsum-free clay in cores through the upper few meters of Lake Tyrrell, Australia. The zones of intrasediment displacive gypsum were interpreted to have precipitated from groundwater during periods of lake desiccation. The spacing between these layers is less than one meter, suggesting that the depth of groundwater influence was small. However, deflation may have removed significant portions of the sediment affected by the groundwater during each desiccation stage. Efflorescent crusts composed of finely crystalline porous aggregates of saline minerals cover most of the saline mudflat surface. They form by complete evaporation of groundwater brine brought to the surface. The mineralogy of efflorescent crusts reflects the chemical composition of the underlying brine but not its overall concentration. Efflorescent crusts are ephemeral features that are readily dissolved by rain or surface floods. Efflorescent crusts are thickest where the groundwater table is near the surface and where they are beyond the reach of perennial saline lake or saline pan lake water. The thickest and most common crusts are composed mostly of halite, although thick efflorescent crusts dominated by gypsum are not uncommon (Stoertz and Ericksen, 1974; Watson, 1983a,b, 1985). Other minerals that dominate efflorescent salt crusts are thenardite (Jones, 1965; Last, 1984), thermonatrite (Eugster, 1970, 1980), trona (Hay, 1968; Papke, 1976), and ulexite or borax (Muessig, 1959, 1966; Bowser and Dickson, 1966; Hunt and others, 1966; Barker and Lefond, 1985; Smith, 1985). Efflorescent salt crusts range from powdery surfaces forming broad humps and depressions to rock hard crystalline crusts over a meter thick, with solution-etched pinnacles and narrow depressions creating 40-50 cm of relief. Thin efflorescent crusts commonly exhibit a polygonal pattern of narrow mounds and broad troughs which may grade into progressively narrower troughs and a more irregular pinnacled surface as the crust thickens basinward. Efflorescent crusts commonly display a variety of dissolution features due to rainfall or dew, including fluted pinnacles, tube-like solution pits, and rubbly brecciation (Hunt and others, 1966; Stoertz and Ericksen, 1974). Powdery efflorescent crusts typically exhibit a delicate granular porosity, whereas crystalline efflorescent crusts are commonly rich in clastic sediment and have a characteristic tubular vertical porosity and a nodular "popcorn" surface. Where efflorescent crusts overlie sandy material, they incorporate the upper surface of the sediment into the crust. This distorts the underlying layers and makes the basal contact of the crust indistinct. Where efflorescent crusts overlie muddy sediment, they are more mud rich near the base but their contact is typically sharp and the underlying layers are commonly not deformed. A striking characteristic of saline mudflat deposits is the complete absence of efflorescent crusts in the underlying sediments. The size and mineralogy of the crystals

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in efflorescent crusts make them readily soluble in undersaturated groundwater, so that even meter-thick crusts are apparently not preserved. Efflorescent crusts composed of relatively insoluble minerals, such as gypsum, may be partially preserved upon burial as irregular, humpy layers of sand-sized crystal fragments. The sedimentary record of efflorescent crusts includes the deformation of bedding and the interplay of the irregular crust surface with eolian, fluvial, and lacustrine sedimentation. Over most of the saline mudflat, the efflorescent crusts produce a massive, porous, poorly-sorted silty mud. This sediment is apparently mostly fine-grained eolian material that adheres to the crust surface. The porosity of the sediment reflects the original porosity of the crust as well as the molds formed from dissolution of evaporite minerals. A vague layering may be defined by variations in the concentrations of different grain sizes, probably reflecting mud drapes from ponded flood waters. Irregular pods, lenses, and distorted layers of silt and sand that commonly float in the muddy matrix are eolian and fluvial material trapped in the surface depressions of the crust (sand patch fabric of Smoot and Castens-Seidell, 1982). The mottled texture of sand, silt, and mud resembles burrow disruption, but is distinguished by: 1) coarse-grained patches that have very irregular shapes including angular edges and internal fractures indicatingcollapse by local solution of the salt crust, 2) grain sizes that vary widely within patches and between adjacent patches, reflecting the local sites of accumulation, and 3) boundaries of the coarse-- grained patches commonly have cuspate embayments against the muddy matrix, reflecting the popcorn surface of the salt crust. Efflorescent crusts deform sandy beds into humps or bowl-shaped polygons (Fig. 3.1 5), disrupting the upper surfaces (Smoot and Castens-Seidell, 1982; Fryburger and others, 1983). Ripple structures are commonly deformed by preferential steepening of the crests, producing a distinctive hump-shaped lens. Some of the structures Glennie (1972, figures 62, 111, 113) interpreted as adhesion ripples may have been formed by efflorescent crusts. Massive porous mud with intrdsediment evaporites and sand patch textures may comprise the bulk of the saline mudflat sedimentary record. Massive muds are generally more important in the central parts of the saline mudflat and sand patches become larger and more abundant towards the margins of the flat (Fig. 3.15). Deformation polygons are more important where the sediment becomes sandier, grading into alluvial fan toe, sandflat, eolian dune, or windflat deposits. Saline mudflat deposits at the margin of perennial saline lakes may consist entirely of perennial lake sediments, even deep water deposits with freshwater fossils, that are disrupted by intrasediment evaporites. Saline mudflats developed on intermittently exposed perennial lake deposits have the sedimentary structures of a

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shallow lake that, in addition to the evaporites, may be disrupted by mudcracks and root structures. Sheetfloods onto saline mudflats produce graded laminae similar to those on dry mudflats, if the efflorescent crusts are thin or powdery. If the efflorescent crusts are thick, sheetfloods fill the crust topography with sand and silt deposits that are deformed by solution collapse features. Sheetfloods may be focused into channel flow where they intersect crusts. The channels may be straight or sinuous, cutting the crust to where it intersects the saline pan or perennial lake margin (Fig. 3.15C,D). The channels produce flat lenses of graded silt and sand that are flat laminated and ripple cross-laminated. This layering is commonly deformed by subsequent efflorescent crust growth or by dissolution of the underlying crust. Overbank silt and clay preferentially accumulate in efflorescent crust pockets, producing irregular flat patches of laminated silt that may be inclined at different angles and that contain angular fractures and irregular breaks in layering. Sheetflood deposits and channel deposits commonly grade into sheet deltas near the intersection with the perennial or ephemeral lake margins. The sedimentary characteristics of these deposits will be described under the shoreline subenvironment.

Dry Mudf lat Subenvironment

Dry mudflats are subaerially exposed plains of fine-grained sediment with abundant features produced by desiccation and soil-like processes. Dry mudflats may border saline mudflats or perennial saline lakes. They also commonly occupy the entire floor of closed basins where they are called playa flats or clay pans. The groundwater table in dry mudflats is too deep for the formation and preservation of significant amounts of evaporites. We distinguish dry mudflats from saline mudflats on the basis of the abundance of desiccation features over intrasediment evaporites. The geomorphological characteristics of dry mudflats have been described by numerous authors (Longwell, 1928a; Stone, 1956; Snyder, 1962; Neal, 1969,1975; Neal and Motts, 1967; Motts, 1969; Reeves, 1966,1968; Flint and Bond, 1968; Schreiber and others, 1972). There are three major types of dry mudflats: 1) subaerially exposed perennial lake bottoms, 2) slowly aggrading mudflats, and 3)rapidly aggrading mudflats. In all three, the most striking feature is the pervasively cracked surface which results from repeated wetting and drying. The cracks define polygons that range in size from centimeter-scale to giant polygons tens of meters in diameter. Centimeter-scale polygons with narrow cracks only a few millimeters deep develop in thin sediment layers deposited on hard dIy substrates. Larger polygons with deeper cracks form where water is ponded on the surface for a time sufficient to saturate the underlying sediments.

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Fig. 3.15. A). Efflorescent salt crust at Saline Valley, California, showing flat-topped sand and siltpatchespartially f illing depressions between humps of eff lorascent crust with ’)popcorn”surf aces. Scale =35 cm long. B). Surface view of a sandysaline mudflat adjacent to an eolian sand sheet, Saline Valley, California. The efflorescent crust forms a polygonal pattern of broad troughs and narrow crests (light). Shovel = 1.4 m long. C). Channel in the efflorescent salt crust of the saline mudflat at Saline Valley, California. The channel f loor (darkest area) is covered by salt encrusted linguoid and straight-crested ripples, lighter areas are salt-encrusted silt-mud leveas. Channel and levee area = I0 m wide. D). Trench in channelsimilar to fig 3.15Cparallel to flow (to the left). There are two f ining upward sequences of flat lamination grading to ripple cross-lamination with efflorescent crust deformation at the top, then muddy sand with sand patches and layers of deformed ripples. Thickness =SO cm.

Giant polygons, defined by cracks that are several meters deep, are apparently formed by surface subsidence in response to water table draw down (Neal and Motts 1967; Neal and others, 1968). Cross-sections of mudcracks on dry mudflats are commonly not simple v-shaped,

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sediment-filled structures. Rather, the mudcracks are complicated, jagged to sinuous features that may bifurcate and may contain multiple sediment fillings (Smoot, 1981). These complicated cracks are produced by: 1) lateral "stepping" of cracks and sheet-crack development at bedding inhomogeneities, such as sand-mud contacts or wet-diy sediment contacts, 2) sediment flowage in cracks, if they fill with water before filling with sediment, 3) repeated cracking of the same crack where only sediment filling the crack is moist enough to crack on drying, 4) sediment partially filling cracks during flooding episodes and 5) superimposed cracks produced by random development of polygons on water-saturated mudflat surfaces (Fig. 3.16). Repeated wetting and drying of sediment on the dry mudflat may also produce slickensided shear planes. The shear planes may develop into a series of bowl-shaped polygons, called gilgae structures (see Harris, 1968), in muds dominated by expanding clays. Perennial lakes may, upon desiccation, develop into dry mudflats. Typically, the surfaces of these lake-bottom mudflats are areas of no net accumulation of sediment or even loss of sediment by deflation (Blackwelder, 1931; Reeves, 1966; Motts, 1965; Bowler, 1973,1986; Young and Evans, 1986). The sediment contains features diagnostic of deposition under lacustrine conditions, but it has subsequently been modified by subaerial processes. Repeated flooding and desiccation of the normally dry surface of the mudflat produces a series of superimposed polygonal cracks (Longwell, 192th; or Motts, 3969) that, combined with sheet cracks, brecciate the sediment. The old lake sediments may be further disrupted by bioturbation, including burrows and roots, and by growth of thin efflorescent evaporite crusts. These dry mudflat surfaces are essentially poorly developed desert soils (Dan, 1973; Birkeland, 1974) superimposed on the old lake deposits. The eastern mudflats of the Bonneville Salt Flat in Utah are an example of this type of deposit. There, laminated lake sediments of Pleistocene Lake Bonneville are disrupted by desiccation features and efflorescent crust growth. Many poorly developed soil horizons reported from cores and from outcrops of Lake Bonneville and Lake Lahonton (Eardley and Gvodetsky, 1960; Morrison, 1964; Morrison and Frye, 1965; Eardley and others, 1973) are probably also this type of mudflat deposit. Slowly aggrading dry mudflats are pavement hard surfaces underlain by massive vesicular mud cut by numerous silt and clay filled cracks (Fig. 3.16B). Vesicles are millimeter-sized spherical to ovate cavities that are commonly connected by thin horizontal and vertical cracks formed during drying. Vesicles form in the sediment by trapping air in a water-mud slurry during floods (Springer, 1958; Shinn, 1968; Deelman, 1972). Sediment transported onto slowly aggrading dry mudflats is typically deposited as millimeter- to sub-millimeter layers of silt and clay that is mostly deflated away by

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Fig. 3.16. A). Surface of slowly aggrading dry mudflat near the center of Mud Lake, Nevada. Several generations of partially filled polygonal desiccation cracks randomly cross cut euch other. Note widening and rounding of older crack edges in contrast to the mo.st recent open cracks. Knife=35 cm long. B). Plastic-impregnated slab from A). Angular blocks of sediment are separated by open vertical and horizontal cracks (black) and by sediment-f illed cracks (c). Circular voids (black) are vesicles. Oriented clay lines some cracks and vesicles. Scale = 7 mm long. C). Slabbed sample of plastic-impregnated mud from the slowly aggradingdry mudf lat at North Panamint Valley, California. Silt bed at top is a sheet-flood deposit that is broken by polygonal cracks. Note upturning of the edgesof the layer near the center. Verticalf eatures are sediment-f illed and open polygonal cracks. Circular to flattened ovate features (black and gray) are open vesicles. Thickness =5cm. wind after drying (Motts, 1965, 1969; Young and Evans, 1986). Most sediment accumulates in open cracks. Later, when the surface dries, a new set of cracks forms and these then become the new locus of sediment accumulation. In this manner, a massive mud is produced by uneven sediment accumulation in thin increments. Discontinuous and heavily disrupted layers may be protected from deflation in local shallow depressions such as partially filled cracks, deflation pockets, phreatophyte

Data Center ,09126599985,[email protected], For Educational Uses 232 NON-MARINE EVAPORITES drawdown sinks, or pits and ruts formed by sediment flow or slumping into cracks. Relatively thick layers of clay or silt and sand may accumulate during floods as sheet flood or sheet delta deposits. These thicker layers are brecciated by superimposed polygonal cracks formed during repeated desiccation and wetting. Silt layers may separate from the underlying mud substrate during desiccation which produces polygonal curl structures (Fig. 3.16C). Development of cyanobacterial mats on silt and clay layers during flood stages may produce over-steepened or even rolled curls as they shrink during desiccation. The concave-up silt curl structures may be greatly modified by desiccation before they are finally buried by slow aggradation. Sand may accumulate as mounds around plants that grow on the dry mudflat surface. These sandy phreatophyte mounds are also commonly surrounded by circular depressions produced by local drawdown of the water table by the plants roots (Motts, 1965; Neal, 1969). The mounds are typically deflated away after the plants die, but thin sand lenses with root structures may be preserved and may be intercalated with clay curls from water ponded in the depression. Rapidly aggrading dry mudflats consist of laminated to thin-bedded sediments with structures diagnostic of deposition from shallow sheetflood flow or from a standing body of water (Fig. 3.17A). Sheetfloods on dry mudflat surfaces deposit thin, discontinuous, graded sand lenses that fill irregular scours. The sand layers may be rich in mud intraclasts derived by erosion and slaking of desiccated mud surfaces. Continuous laminae of silty mud commonly cap the sand layers. Where sheetfloods intersect expanding lake waters during floods, they form graded sand-mud laminae with the mud thickening basinward. On the outer margins of rapidly aggrading dry mudflats, shallow streams or sheetfloods may deposit broad, sheet-like deltaic wedges built of graded thin beds of flat laminated silt or climbing ripples of silt and sand. Such sheetlike deltas are further described under the shoreline subenvironment. Each layer of the rapidly aggrading dry mudflat is disrupted by polygonal desiccation cracks which range from a few millimeters to several centimeters deep. The cracks are filled and buried with each flooding event, so that superimposed crack patterns, complex fillings, and subaqueous slumping of crack walls are not as important as in slowly aggrading dry mudflats. Instead, the cracks are typically simple, sinuous v-shaped cracks or jagged stepcracks. Evaporites in dry mudflats are limited to fluffy efflorescences that grow on or near the surface and euhedral crystals and crystal crusts that form in the sediment by processes similar to those described for soils. Efflorescent crusts grown within cracks and fenestrae break the mud into sand- and silt-sized clumps (pellets) separated by pedestals of salt (Fig. 3.18). The surface texture is called puffy ground or self-rking ground (Snyder, 1962; Motts, 1965; Neal, 1969). The pellets are easily eroded by wind

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B A

(3)

0 SAND MUD a MUD CRACKS

Fig. 3-17. Sedimentaryfeaturesof rapidlyaggradingdrymudflats. A). Schematicdrawing of sequence that produces graded sand-mud couplets in a rapidly aggrading dry mudflat. 1) Sheetflood from left cause.s lake (right) to expand, 2) Expanding jake laps over the sheet flood causing deceleration of flow and deposition of sand then mud, 3)After f lood, lake begins to evaporate exposing layers to desiccation. The area affected may be tens of kilometers wide. Thicker clays and larger mudcracks occur basinward. (from Smoot, 1977). B). Thin-bedded silt and clay layers disrupted by deep polygonal cracks, local depression in the dry mudflat of Big Smoky Playa, Nevada. Thick layers were deposited in a shallow ephemeral lake. Mudcracks extend from each bed boundary and erosion of upper crack boundaries indicates prolonged subaerial exposure. Thickness = 40 cm. deflation (Bowler, 1973, 1986; Young and Evans, 1986) and may be deposited as clay dunes. Puffy ground efflorescent crusts may form directly from a moderately shallow

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DESERT DRY LAKE "MODEL CYCLE"

CLAYEY MUDSTONE

BEDDED GRANULAR

NRSSIVE GRANULAR

BRECCIA

CLAYEY MUDSTONE

Fig. 3.18. Puffy ground crusts in dry mudflats. A). Surface of dry mudflat with a well-developedpuffygroundcrust, Desert Dry Lake, Nevada. Thepolygonal crack pattern is niodif ied by the growth of microcrystalline halite in cracks and vesicles pushing clumps of mud apart. Brunton compass = R cm long. (Photo courtesy of L. V. Benson) B). Sketch from a core in Desert Dry Lake, Nevada, showing transition f rom massive calcareous clay to granule-sized clay clumps f illing polygonal cracks separating irregular blocks of the calcareous clay. This represents part of a repeated pattern over a thickness of 65 m. The granular brecciated textures are similar to the modern dry mudflat surface in (A) superimposed on older perennial lake sediment. Thickness = 2 m. dilute groundwater. Bowler (1986) believes that puffy ground development and pellet formation represents an early stage of saline groundwater evolution in Australian playas. They may also form from recycled salts derived from windblown dust. In this case, salts are dissolved by rainwater and are later precipitated as an efflorescence as the rainwater evaporates. Neal and Motts (1967) and Neal (1969) note that puffy grounds may

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develop on those areas of a dry mudflat where ponding of flood water localizes dissolved salts. Recycling of windblown dust by percolation of rainwater through dry mudflat sediments, followed by evaporation, may also produce intrasediment evaporite crystals. These consist of: 1) powdey aggregates or tiny euhedral crystals that line mudcracks and vesicles, 2) thin irregular horizontal crusts with fibrous crystal growths, and 3) single layers of large crystals that extend up or down from the surface of a perched groundwater table, such as on a dense clay layer or cemented pavement. Slowly aggrading dry mudflats may grade laterally into rapidly aggrading dry mudflats in an area adjacent to a major inflow drainage. A different situation is seen in the dry mudflats of the Mojave region in the southwestern U.S.A. There, erosion of desiccated old lake bottom sediments on the outer margins of the mudflat provide sediment for slowly aggrading mudflats in the central part of the basins (Neal, 1969).

Shoreline Subenvironment

Shoreline deposits on the margins of perennial and ephemeral lakes include deltas, beaches, spits, bars, platforms, and carbonate mounds. There are few published studies of shoreline deposits in association with non-marine evaporites or even for freshwater lakes. The sedimentary structures and shapes of shoreline deposits in saline lakes are commonly identical to marine shoreline deposits. We will emphasize those shoreline features that are characteristic of arid closed basin settings. Deltaic deposits form where streams or sheetfloods intersect a lake margin. Sediment is rapidly deposited when stream or sheetflood flow decelerates in the standing lake water. Clay may be rapidly deposited due to flocculation where fresh water intersects a saline lake (Hyne and others, 1979). In arid climate settings, coarse-grained deltas are common. The term fan delta has been used to describe coarse-grained deltas in a variety of settings (Galloway, 1976; Sykes and Brand, 1976; Nemec and Steel, 1984; Nilsen, 1985; McPherson and others, 1987). It has been argued, however, that the term be restricted to subaqueous deposits of alluvial fans and be distinguished from braid deltas formed by braided rivers (Sykes and Brand, 1976; Postma and Ori, 1984; McPherson and others, 1987). Deltaic deposition in saline lakes is also affected by the characteristic radical changes of lake depth and area. Evaporites are commonly not present in deltaic deposits of saline lakes. Where present, they have characteristics similar to those in shallow perennial saline lakes and in saline mudflats. Deltaic deposits in saline lake settings are here divided into three varieties: 1) birdfoot deltas, 2) "Gilbert-type" deltas, and 3) sheet deltas (Fig. 3.19).

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These include some of the deposits fitting braid delta and fan delta definitions. Fan deltas, as defined by McPherson and others (1987), are discussed after wave-formed deposits. Birdfoot deltas have a digitate geometry reflecting the dominance of stream flow in sediment distribution. This is the dominant delta type where perennial streams enter shallow perennial saline lakes. Birdfoot delta deposits consist of distributary channels bounded by levee deposits, distributary mouth bars where the channels terminate in the lake, and interdistributary bays (Fig. 3.19A). The sedimentary traits of birdfoot deltas are described in Coleman and Wright (1975), Coleman (1976), and Donaldson and others (1970). The distributary channel deposits are low-velocity stream deposits which form fining-upward sequences. Levees are composed of flat laminated silt and clay with some ripple cross-lamination which may be disrupted by burrowing or root structures. The distributary mouth bars consist of beds of climbing ripple cross-lamination separated by silt partings. Interdistributary bays are filled with lacustrine silt and clay that reflect shallow lake conditions. Birdfoot delta deposits are similar to fluvial deposits and may be indistinguishable from them in areas where lake levels frequently drop leaving sediment subaerially exposed. Abundant soft-sediment deformation structures and thick, laterally continuous interbeds of mud suggest deltaic conditions. Rapid drops in lake depth may result in distributary channel deposits cutting directly into deep lake deposits. "Gilbert-type" deltas are fan-shaped wedges of sediment built into lakes by sediment gravity flows. This delta type was first described by Gilbert (1885, 1890) from shoreline deposits of the Pleistocene Lake Bonneville in Utah. These deposits form where perennial rivers enter relatively deep lakes or where ephemeral streams (including alluvial fan channels) enter deep or shallow lakes. "Gilbert-type'' deltas consist mostly of relatively steep foresets of gravity flow deposits overlying relatively flat-lying lacustrine bottomset sediments and overlain by flat, fluvial topset sediments. Postma and Roep (1985) distinguish between "Gilbert-type"deltas where sedimentation is by slump-initiated grain flows and debris flows and "Gilbert-type" deltas where deposition is largely by stream flow entrained beneath the water. Delta foresets produced by grain flows and debris flows are characteristicallysteep (near the angle of repose) and coarse grained. Grain flow layers are internally reverse-graded and coarsen down slope. Subaqueous debris flow layers are matrix-supported muddy conglomerates that may have normal grading and clast imbrication (Enos, 1977; Gloppen and Steel, 1981). Postma and Roep (1985) emphasize the importance of resedimented conglomerates that are deposited by debris flows initiated by the failure of the delta front (Nemec and others, 1980; Nemec and Steel, 1984; Massari, 1984).

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Fig. 3.19. Three rnujor types o,f deltas in saline lakes. A). Birdfoot delta of the Bear River into Great Salt Lake, Utah. Distrihutary channel with levees and crevasse splays is intermittently drowned b.y rising lake levels and exposed when lake level drops. (Photo courtesy of B. F. Jones). B). "Gilbert-type" delta at the mouth of an ephemeral stream entering Lake Natron, Tanzania. Whitefringe at top of ,foreset is an efflorescent salt crust formed during a drop in lake level. (Photo courtesy of H.P. Eugster). C). Sheet delta into the saline mudflat of Lake Frome, Australia, at the mouth 0.f an ephemeral stream. Lake level rises during f Ioods to cover the area of small channelsproducing a sheet of sand with delta front characteristics. White siuface is a thin efflorescent .salt crust.

Other sedimentary features common to steep-front "Gilbert-type" deltas are soft-sediment deformation, dewatering structures, turbidite beds, and thick silty clay interbeds (Sneh, 1979; Postma, 1983, 1984; Nemec and Steel, 1984; Postma and Roep, 1985). Topsets consist of braid stream deposits and debris flow deposits. "Gilbert-type"deltas dominated by stream-flow entrained beneath the surface are best known from glaciolacustrine deltas (Jopling and Walker, 1968; Gustavson and others, 1975; Clemmensen and Houmark-Nielsen, 1981; Stone, 1982). These deltas are characterized by finer-grained sediment and have lower angle foresets than the

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"Gilbert-type" deltas dominated by grain flows and debris flows. Foreset deposits are commonly sand or silt graded beds with internal structures similar to the "Bouma sequences" of turbidites (Bouma, 1964). Planar lamination and climbing ripples form "deceleration of flow" sequences wherein ripples exhibit increasing angles of climb and more preservation of their stoss side both vertically within a single foreset bed and downflow (Fig. 3.20A). Packages of delta foreset beds may be separated by broad, curved scour surfaces as the locus of foreset building shifts. Proximal foresets (i.e. those formed where the river first intersects the lake) tend to be steep and coarse-grained, dominated by planar lamination, and may include grain flow and debris flow deposits. Distal foresets are typically low angle and are dominated by ripple cross-lamination with muddy silt drapes. Bottomset beds may be massive silty muds or flat-laminated silty mud with graded layers composed of silt or fine sand. Soft sediment deformation, including slump faults, load casts, and folding, is common in both the foreset and bottomset beds. Topsets are commonly composed of coarse braid stream deposits, but low-gradient, flash-flooding stream deposits may also form topsets. Rise and fall of lake level affects "Gilbert-type" delta deposition either by flooding the topset surface and reinitiating delta building at a higher elevation, or by causing channels to incise foreset deposits as local base level drops. This process results in deposits consisting of stacks of delta foresets, each commonly 2-4 m thick, that are part of a larger-scale coarsening-upward deltaic deposit (Fig. 3.20B). The delta foresets may be separated by thin lacustrine clays or fluvial deposits with root structures or desiccation cracks, all reflecting the fluctuating lake levels. Channel deposits may be absent in most cross-sections due to their localization in steep-edged gullies. Examples of this type of delta are the Truckee River delta into Pyramid Lake (Born, 1972) and deltas in Lake Hazar, Turkey (Hempton and Dewey, 1983; Dunne and Hempton, 1984). Sheet deltas form where streams or sheetfloods are drowned by a shallow lake expanding in response to floods. Sediments reflecting the intersection of stream flow and standing water (delta front deposits) accumulate over an expanding area as flooding is in progress, so that proximal deltaic deposits are overlain by distal deposits. Sheet deltas are nearly flat sand and silt wedges composed of graded thin beds 1-40 cm thick separated by mud (Fig. 3.21). Each thin bed consists of a "deceleration of flow" sequence of mostly ripple cross-lamination that may overlie planar lamination. The graded beds fine and thin basinward and are separated by thicker mud partings. The sedimentary structures in each graded bed also reflect progressively lower velocity flow basinward. When a number of floods cause a lake to progressively expand over wider areas, a meter-scale fining-upward sequence of progressively thinner sand beds and thicker mud partings may form. Load casts and soft-sediment deformation may be common.

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I IR 0.5 km

Fig. 3.20. A). Climbing-ripple cross-lamination from a low-angle delta f oreset oj the Holocene Lake Cuahilla, Salton Basin, California. Ripples show increase 0.f angle of climb with decrease of grain size indicating deceleration of flow. Flow is to the left. Thickness = 20 cm. B). Schematic sketch of stacked "Gilbert-type"de1taf oresets based on observations at Pyramid Lake, Nevada, and Mono Lake, California. Delta progradation to the left makes coarsening-upward sequence. Low-angle f oresets are comprised of climbing-ripple cross-lamination. Proximal foresetsare steep avalanche sets composed of grainflows. Progradation of individual foresets is punctuated by periods of subaerial exposure as lake levels f all, resulting in root structures.

Sheet deltas are the dominant delta type in dry mudflat and saline mudflat settings. In saline mudflats, deltaic sediments often have intrdsediment evaporite crystal growths and may have thick efflorescent crusts. Collapse structures and growth faults are common in this setting due to the dissolution of efflorescent crusts by the expanding lake. Delta deposits combining the characteristics of sheet deltas and "Gilbert-type'' deltas form where large, flash-flood rivers enter shallow perennial saline lakes. In the Salton trough, California, a large, low-angle "Gilbert-type" delta formed where the Colorado River intermittently flooded into Holocene Lake Cuahilla. Lake Cuahilla expanded over a large area each time the Colorado River flooded, nearly filling the entire basin, then shrank to near dryness in the intervening dry years. This delta deposit consists of sheet-like graded sand and silt beds comprising "deceleration of flow" sequences that range from a few centimeters to over 3 m thick. The graded beds are commonly stacked to form fining-upward sequences 5-10 m thick, consisting of progressively thinner beds with thicker mud partings. The tops of these fining-upward

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LAKE TRANSGRESSION FO LLO W ED BY REG RESSlO N

DELTA1C CHANNEL APRON SYSTEM

Fig. 3.21. Sketch showing plan view and cross-section of sheet delta deposit. The cross-section shows bedsformed by rise of lake level fromf looding, resulting in an upward fining sequence and progradation of the delta front, followed by a desiccation period. Flow is to the leftwith planur laminutiongrading to climbing-ripple cross-laminution then planar silt and clay in the downflow direction. Wedge shapes are mudcracks. Horizontal distance shown varies from 300 m to over 10 kni. sequences exhibit root structures, desiccation cracks, and eolian sand layers, indicating periods of subaerial exposure. Each fining-upward sequence is overlain by a thicker, coarser fining-upward sequence that truncates the lower one at a low angle. This progression produces a large-scale, coarsening-upward sequence, that reflects the basinward progradation of the delta. Wave-formeddeposits in ephemeral and perennial saline lakes include beaches, bars, spits, and platforms (Fig. 3.22A). In large perennial lakes wave reworking and longshore drift produce well-sorted sandy deposits that are sedimentologically

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indistinguishable from marine deposits (Wulf, 1963; Fraser and Hester, 1977; Fouch and Dean, 1982). These deposits commonly include low-angle inclined lamination overlying planar scour surfaces, decimeter-scale tabular foresets with bimodal and polymodal orientations, and oscillatory ripples with mud partings and flasers. Similar features may develop on the margins of large ephemeral saline lakes, if they remain as lakes for a sufficient period of time (for instance, the shoreline of Lake Eyre, Australia; Bonython and Mason, 1953, plate 3; King, 1956, fig. 4; Callen and others, 1986). Spit and bar deposits may form tabular foresets several meters thick composed of sand and gravel (Gilbert, 1885, 1890). Wave-formed shoreline deposits in shallow ephemeral lakes are typically composed of thin sheets of oscillatory rippled sand, isolated bar deposits, or strandline deposits. Oscillatory rippled sands are generally well sorted and may be comprised of locally derived material such as brine shrimp fecal pellets, ooids, or evaporite crystals. Ripple crests in gently shoaling lakes may be oriented parallel to shore and the internal stratification asymmetric towards shore. The rippled sands in deeper water may have mud flasers or laterally continuous mud partings where wave action is more intermittent. Bar deposits consist of steep, shoreward-dippingforesets that are eroded by more gently dipping lakeward foresets. Gravel bars and strandlines commonly contain imbricated clasts, indicating a shoreward flow direction. The gravel may also have bimodal imbrication if the clasts are small and reworked into ridges that are several grains thick. The bar and strandline deposits may consist largely of locally derived materials, such as mud clasts from desiccation cracks, crusts, algal mat fragments, soil caliche nodules, and evaporite minerals. Wave-formed shoreline deposits developed on alluvial fans are coarse-grained bar or platform deposits that may be built over wavecut terraces (Smoot and LeTourneau, 1989). These deposits are composed of debris flow and stream deposits that have been partially winnowed and sorted by wave action. Bar deposits form shoreward-dipping foresets ranging from 2 m thick to about 10 cm thick truncated by more gently dipping lakeward foresets. Their internal bedding is defined by alternation of grain sizes and grain shapes (Bluck, 1967; Dobkins and Folk, 1970; Clifton, 1973; Nemec and Steel, 1984; Smoot and LeTourneau, 1989). The large bar foresets are commonly slightly convex and flare towards the base, whereas smaller bar foresets are more tabular. Large, coarse-grained bars form where waves strike the fan without loss of energy due to shoaling, whereas small bars form where wave energy is lower. Therefore, the larger bars form high on steep alluvial fans, whereas smaller bars form on the lower portions of fans where the dips are lower. Platform deposits consist of flat-topped, wedge-shaped deposits that form steep

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Fig. 3.22. A). Spits and tombolos at margin of saline pan at Lake Tyrrell, Australia. Eolian sands are reworked by waves during f lood stages. (Photo courtesy of J. T. Teller) B). Small wave-formedplatform deposit at Pyramid Lake, Nevada, built over a wave-cut terrace into an alluvial fan. Tabular sets inclined lakeward (left) show radical changes in grain size, good sorting of finer grain sizes, and reverse grading. Planar beds on top contain oscillatory ripple lenses. Small beach bar deposit with tabular sets inclined landward overlies a cobble lag on top of the platform. Knife =30 cm long.

lakeward-dipping foresets (Fig. 3.22B). The platform builds lakeward and parallel to shore by the sweeping of wave-sorted sediment off the flat platform top into the lake as grain flows. The internal structure of these foresets resembles grainflow dominated "Gilbert-type"delta foresets (Swirydczuk and others, 1979,1980),but 1) each foreset bed is composed of moderately well-sorted sand or gravel, 2) adjacent foreset beds may differ greatly in mean grain size, 3) the deposit is linear parallel to shore rather than lobate, 4) the deposit may have a grain composition considerably different from that of adjacent rivers or fans. Where waves rework boulder- and cobble-rich alluvial fan deposits, the resultant sediments may be massive or poorly bedded, particularly if wave action only occurs for a short time. Waves winnow and sort the gravel and sand fraction of the deposits, leaving the coarser grains as randomly scattered lag deposits. The large clasts disrupt wave patterns resulting in the development of local pockets of different grain sizes or small lenses of oscillatory rippled sand. Well-sorted granule or gravel pods and lenses usually have a well-sorted sand, mostly introduced by infiltration, in their intergranular spaces. Many slightly saline perennial lakes and the older, higher shorelines of modern perennial and ephemeral saline lakes have algal tufa deposits (Fig. 3.23). The tufas are typically composed of calcite or aragonite (Russel, 1885; Eardley, 1938; Carozzi, 1962;

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Fig. 3.23. A). Large tufa mound at shore of Pyramid Lake, Nevada. Rounded knobs consist of porous radial tufa with a dense micritic outer crust. The radial pattern may in part reflect former ikaite (CaCO, - hH,O) crystals replaced by calcite (Shearman and others, 1989). Mound is about 10 m high. (Photo courtesy of L.V. Benson). B). Finely laminated stromatolitic tufaf ormingsmall club-shaped heads on a cobble from shoreline of Walker Lake, Nevada. Tufa consists of monohydrocalcite (CaCO, - H,O) which dehydrates on subaerial exposure to calcite with shrinkage f eatures. Scale = cm.

Scholl and Taft, 1964; Halley, 1976; Von der Borch and others, 1977; Abell and others, 1978,1982; Spencer, 1982; Dean and Fouch, 1983; Livnat and Kronfeld, 1985; Hil- larie-Marcel and others, 1986; Hillaire-Marcel and Casanova, 1987) but may also be composed of monohydrocalcite (Spencer, 1977). Shearman and others (1989) interpret porous radial tufa (thinolite of Radbruch, 1957) in Pyramid and Mono lakes as pseudomorphs of ikaite (CaCO, 6H,O) replaced by calcite. The porous texture may be due to volume loss during replacement. Tufa deposits form where algal filaments become calcified by chemical and/or biological supersaturation of the lake waters. The resultant mounds and ridges are initiated on stable surfaces along the shoreline such as boulders, cement crusts, or tree branches and trunks. Tufa mounds are best developed where spring waters intersect saline lakes, even well away from shore. This aspect will be discussed later in the section on springs. Tufa mounds also develop where supersaturation with respect to carbonate minerals is reached by evaporation, wave degassing, photosynthesis by algae, or increase of water temperature. Eugster (1980) describes button like tufa coatings on small clasts along the margin of Lake Magadi that apparently formed as freshwater surface runoff mixed with the alkaline brines of the ephemeral lake. Oolitic sands may be associated with tufa mounds, such as in the Great Salt Lake, Utah (Eardley, 1938; Sandberg, 1975; Halley, 1977). Burne and others (1980) observed carbonate-encrusted charophytes in saline lakes in Australia. In the freshwater Green Lake in New York (Dean and Eggleston, 1975; Eggleston and Dean, 1976) and

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a small marl lake in Michigan (Murphy and Wilkinson, 1980), calcareous build-ups of tufa and charophytes produce platforms that extend into the lakes. Last and Slezdk (1987) mention a carbonate platform with abundant algal material from Freefight Lake in Saskatchewan. Fan deltas, as defined by McPherson and others (1987), include modern or Recent examplesfrom the Dead Sea (Sneh, 1979; Manspeizer, 1985), the Gulf of Aqaba (Hayward, 1985), Lake Hazar in Turkey (Dunne and Hempton, 1984), and Walker Lake (Link and others, 1985). The term of fan delta, however, fails to distinguish between subaqueous deposits built out into the lake by alluvial fan sedimentation and those formed by wave reworking of subaerial fan deposits, a common situation in closed basins. An alluvial fan terminating in a lake may build deposits like a bottom-set modified "Gilbert-type" delta, if the lake resides for a long period or if the fan channels are frequently active. In arid settings, however, perennial lakes may rise and fall too rapidly for appreciable subaqueous sediments to accumulate and the fan channels may remain dormant during the entire period of high lake stand. The record of the subaqueous fan, in this situation, is primarily a wave-formed deposit that includes turbidites formed by storm waves reworking sediment and washing it offshore (Smoot and LeTourneau, 1989).

Ancient Non-marine Evaporites In Lacustrine Deposits

Most ancient non-marine evaporites occur in lacustrine environments (Table 3.3) and most are interpreted as perennial saline lake or saline mudflat deposits. The vague term playa-lake is commonly used for the deposits, without stating whether the evaporites accumulated in a saline lake or in a freshwater lake that alternated with a saline mudflat. Saline pan Occurrences have been noted in some of the more recent literature. Dry mudflat deposits are commonly referred to as playa mudflats with little or no distinction from saline mudflat deposits. Descriptions of shoreline deposits have generally been restricted to perennial lakes. They are typically classed as deltaic, wave-formed, or stromatolitic, with few details presented. The criteria commonly used to identify ancient perennial saline lake deposits are: 1) organic-matter-rich lamination, 2) absence or paucity of fossils, 3) intrasediment evaporites, most commonly pseudomorphs after gypsum or halite, and 4) Occurrence of alkaline minerals such as dolomite, zeolites (particularly analcime), diagenetic feldspars, and Mg-rich clays. Examples of lacustrine deposits of cumulus crystals include interlaminated gypsum and carbonate minerals in the Pleistocene Lisan Formation, Israel (Begin and others,

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1974, 1980) and in the Jurassic Todilto Formation, New Mexico (Tanner, 1970, 1971; Anderson and Kirkland, 1960),Tertiary laminated Na-borate deposits in the southwestern U.S. (Barnard and Kistler, 1966; Barker and Barker, 1985), and nahcolite and halite beds in the Green River Formation of the Piceance Basin, Colorado (Dyni, 1974). Many of the "sugary" trona beds in the Wilkins Peak Member of the Green River Formation in the Green River Basin, Wyoming shown by Birnbaum and Radlick (1982), also resemble cumulus crystal beds. Lenticular beds of fine crystals may represent reworking of cumulus crystals by waves (Fig. 3.24A). Examples of crusts formed in perennial lake deposits include vertically oriented gypsum crystals in the Tertiary Mormoiron Basin, France (Truc, 1978, 1980) and oriented borate crystal beds in the Tertiary Kramer lake beds of the Ricardo Formation, California (Bowser and Dickson, 1966). Some of the Green River Formation trona and halite crystal layers may also be subaqueous crusts. Carter and Pickerill (1985b, fig. 1Oc) interpret layers of mudstone with pseudomorphs of vertically oriented gypsum crystals in the Devonian Albert Formation, New Brunswick as crusts, and Bell (1989) interprets irregular limestone layers interbedded with flat laminated limestone in the Jurassic-Lower Cretaceous red beds of northern Chile as replaced subaqueous gypsum crusts. Bedded halite deposits (Deardorff, 1963;Smith and Haines, 1964; Dingman, 1967; Benavides, 1968; Stocklin, 1968; Smith, 1970; Strakhov, 1970; Culbertson, 1971; Dear- dorff and Mannion, 1971; Glennie, 1972; Holser and others, 1972; Wardlaw, 1972; Ar- thurton, 1973; Eugster and Hardie, 1975; Holser, 1978; Orti Cabo and others, 1979, 1986; Smith and Cosby, 1979; Smith, 1979; Smith and others, 1983; Ingles Urpinell and others, 1986) and bedded trona deposits (Culbertson, 1961, 1971; Deardorff, 1963; Bradley, 1964; Bradley and Eugster, 1969; Deardorff and Mannion, 1971; Eugster and Hardie, 197.5; Birnbaum and Radlick, 1982; Zhang Youxun, 1985; Southgate and others, 1989) are reported from many ancient lacustrine deposits. These beds are commonly associated with mudstones that contain abundant euhedral evaporite crystals with displacive and poikilitic growth features. Most of these deposits were probably formed in saline pans (Fig. 3.25). Orti Cab0 and others (1979) describe layered thenardite and gypsum that also resemble saline pan deposits. Intrasediment evaporites interpreted as having formed in perennial lakes commonly occur in laminated or thin-bedded mudstones. The crystals include skeletal or dendritic halite (Llewellyn, 1968; Parnell, 1983; Southgate, 1982), bedding plane crystals of halite or some other mineral (Demicco and Gierlowski-Kordesch, 1986; McDonald and LeTourneau, 1988), or layer specific growths of gypsum crystals (Greiner, 1974; Smith, 1982; Carter and Pickerill, 1985b). In the Tertiary Sentinel Butte Member of the Fort Union Formation, North Dakota (Boyer, 1981), angular, sand-filled

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Table 3.3. Selected references to ancient evaporite deposits interpreted as non-murinr and their associated f acies.

Tertiary North America a) Green River Formation - Bimhaum and Radlick, 19a; Boyer, 19=, Bradley, I%, 1Y2.9,lW, 1970; Bradley and Eugster, I(%(& Braunagle and Stanley, 1977; Cole and Picard, 1981; Cullxrtson, 1961, 107 I; Deardorff, 1963; Deardorff and Mannion, 1971; Desborough, 1978; Dyni, 1974,1976;Dyniand others, 198% Eugster and Hardie, 1975; Eugster and Surdam, 1973; Fahey, 1962; Fouch and others, 1976; (irmde, 19x0; Iijima and Hay, lW, Lundell and Surdam, 1975; Milton, 1971; Milton and Eugster, 195'); Mott and Drever, 1983; Picard. 1957; Picard and High, 1968,1972a,b; Rwhler, 1965,1069;Ryder and others, IWO;Smith, 1974 Smoot, 1977, 1978,1983; Stanley and Surdam, 1978; Surdam and others, 1972; Surdam and Parker, 1972: Surdam and Stanley, 1979; Surdam and Wolfbauer, 197.5; Surdam and Wray, 1976;Tank, 1972; Trudell and others, 1974; Tuttle, 1988, Williamson and Picard, 1974 Woifbauer and Surdam, 1974 b) Ridge Basin - Crowell, 1974; Crowell and Link, 1982 Link, 1982a,h: Link and Osbome, 1978,1082;Link and others, 1978; Smith, 19x2: Wood and Link, 19%' c) Kramer Borate Deposit - Bamard and Kistler, 1W~Btnvser and Dickson, 1966; Christ and Garrels, 1959: Siefke, 1985; Smith, 1985 d) Muddy Creek Formation - Johnson and others, 1989: Longwell, 1028h. 1936; Mannion, 1%)3;Netherland, SeweU, and Associates, 1977; Pierce, 1974, 1976 e)Barstow Formation - Link, 19x0; Sheppard and Gude, 1%)'); Wotdburne and Tedford, 1982 f) Phoenix Basin - Eaton and others, 1972; Eberly and Stanley, 1978; .Johnson and others, 198% Pierce, 1974, 1W6 g) Others - Ballance, 19M, Barker, 19x0; Barker and Barker, 1985; Bethke and Rye, 197'); Blake, 1982; Bohanncm, 1976; Boles and Surdam, 1Wk Boyer, 1981; Butler, IYM. DCM~and Drewes, 196%Eherly and Stanley, 1978; Fouch, 197%Gude and Sheppard, 19% Hay, 1954, 1%: Hay and other\, 1086, Johnson and others, 198'); Leftmd and Barker, 1985; McAllister and Ross, 1978; Noble. 1%?h, Palmer, 19.57; Pierce, 1972, 1973,1974,1976;Piper, 1985; Sheppard and Gude, 1973a,b, 1974. 1986; Smith. 1985; Spittler and Arthur, 1'9Q Squires and Advocate, 1982; Stanley and C'ollinson, 1979; Steven and Eaton, 1975; Stuart and Willingham, 1984; Surdam and others, 1972; Wells, 1983; Wtmdbume and others, 1982 Europe a) Ebro Basin - Bimbaum and Coleman, 1979; Friend and others, 197'); Williams and Bimhaum, 1975 b) Tajo Basin - Brell and others, in press; Calvo and others, 1986, 198% In& Urpinell and others, 19%: Jones and others, 1986, Orti Cabo and others, 1979, 1986 c) Others - Freytet, 1973; Ingles Urpinell and others, 1986; Nickel, 19=. Plmiat, lW5; Rios. Iohx; Stamatakis. 198'); Tmc, 1978,1980 South America Benavides, lW, Dial, 1984; Dingman, 1%7; Flint, 1985,1986, Harrington, 1961, Newell, 1949. Ponte and others, lYW, Pratt, 1961 Asia Alhayrak and Protopapas, 1985; (hen Ouanmao and Diekinson, 19%; Gokcen and others, 1978; Gu Chengao. 1988, Helvaki, 1984; Stocklin, 1968, Strakhov, I97nO: Sun Shu and others, 198% Zhang Youxun, 1985 Africa Cagle and Craft, 1970 Hay, 197%Hay and Reeder, 1978; Ward, 1988 Australia Ambrose and Flint, 1981; Callen and others, 1986

Mesozoic Atlantic Margin Asmus and Ponte, 1973; Belemonte and others, 1%5; Benavides, IW, Bertani and Caroui. 10x4; Brice and Pardo, 1981; Brink, 197%Camp and others, 1974; Evans, 1978; Ghignone and Andrade, 1970; Jan= and others, 1W, Ponte and others, 19W, Szatmari and others, 1979; Tian and others, 19x3 Tortochaux, 1W Early Mesozoic a) Newark Supergroup - Aryden and Rodolfo. 1% Demicco and Gicrlowsk-Kordcsch, 1986; Glaeser,

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Table 3.3. (continued)

1%; Gore, 1988a,b, 1989; Handy, 197% Hawkins, 1W, Hent;., 1985; Hubert, 1978; Hubert and F'lorenw, 19%; Hubert and Hyde, 1082, Hubert and Merb, 1980,1984; Hubert and others, 1Yi6,1978; Kat;., 1983; Klein, 1962; LeTourneau, 19x5; LeTourneau and Smoot, 1985; Lindholm and others, lY79; Lorenz, 19%; Manspeker and Olsen, 1981; McDonald and LeTourneau, 19x8; McLaughlin, lY39; McLaughlm and Gerhard, 1953; Olsen, 1980,1984,1986,Olsen and others, 198% Olsen and Schlische, 19x8; Parker and others, 19%; Parnell, 1983,198Gb Sanders, IW, Schaller, 19372; Smith and Robison, 19%; Smoot, 1985, in press; Smoot and Katr, l9m, Smoot and LeToumeau, 1989;Smoot and Olsen, 1988,Smoot and Robinson, 1988, Stevens and Hubert, 1980; Turner-Peterson, 1980, 1988; Van Houten, 1962,1964, 1%5a,b, 1969a,b, 1980; Wessel, 1969; Wheeler and Textoris, 1Y78; Wherry, 1916; Ziegler, 1983 b) English Triassic - Arthurton, 1973,1080;Evans, 1970; Holmes and others, 1983; Llewellyn, IW,Tucker, 1977,1978; Tucker, 1981 c) North Africa - Beauchamp, 1988, Brown, 19%) Loren;.,1976,1988a,bMattis, 1Y77; Petit and Beauchamp, 1W) d) Greenland - Clemmensen, IY78a,b, lY79,19M) e) Other - McKee, 1954 Others Anderson and Kirkland, 1W,Bell, 198% Chen Quanmao and Dickinson, lYW, Chyi Clang, 1981; Fouch, 1979; Freytet, 1973; Goldkry, 1982a,b; Gradzinski and Jerqkiewicf, 1974; High and Picard, 1965; Lehman, 1989; Platt, 1989; Rawson, 19FQ Suare7 and Bell, 1987, Sun Shu and others, 1989; Sweet and Donovan, 1988; Tanner, lY70, 1971; Turner-Peterson, 1988, Turner-Peterson and Fishman, 19%

Permian Rotliegendes Clemmenwn and AbrahamFen, 1983; Glennie, lY70,1972,1983a,b, 19M; Holux, 198; Mader and Yardley, 19x5; 'jmith, 1970, 1971; Smith and Croshy, 1979 Karoo System Beukes, 1970; Keyser, 1%; Smith, 1979,1990;Stear, 1983; Van Dijk and others, 1978

Carboniferous Rocky Brook Formation (>all and Hyde, 1989

Devonian Alberl Shale Carter and Pickerill, 1985a,b; Greiner, 197% Howie, 1979; Macauley and others, 1984, Pickerill and others, 1985; Smith and Gibliig, 1987; Webb, 1977 Orcadian Basin Donovan, 1975,1980; Donovan and Foster, 19772; Parnell, 1Y86a; Tunbridge, 1981,19&l Others Hardie, 1984, Holser and others, 19772; Kendd, 1979,1984

Cambrian Officer Basin Pitt and others, 1980, Southgate and others, 1989, White and Youngs, 1W

Proterozoic MatArthur Bash Donneb and Jackson, 1988, Jackson, 1985; Jackson and others, 1986,1987;Muir, 1981,1983; Muir and others, 1980, Neudert and Russel, 1982; Oehler and Logan, 1977 Belt Supergroup Grotzinger, 1980, Winston, 1977,1984;Winston and others, 1984 Others Behr and others, 1983; Cheadle, 1986, Clemmey, 1978; Ellmore, 198%Eugster, 1985; Eugster and Chou, 1973; Petola, 1978; Porada and Behr, 1988, Ross, 1983, Ross and Chiarenzelli, 1985; Rowlands and others, 1980; Southgate, 1982.1486

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Fig. 3.24. A). Laminated, perennial lake dolomitic mudstone withpinch-and-swe1lIayer.s and lenses of rounded orsubhedral shortitepseudomorphed by calcite. Lenses are probably oscillatory ripples. Sample = 10 cm thick. B). Succession of shortite crystal habits (pseudomorphed by calcite) in a dolomitic mudstone suggesting a phreatic to vadose saline mudf lat transition. Randomly distributed rosettes are overlain by irregular ma.xseS o,f euhedral crystals. Top of sample contains spherical masses of anhedral to subhedral crystals restricted to sand layers. Scale = 2 cm long. Figs A and B from Eocene Wilkins Peak Member, Green River Formation, Wyoming.

casts occur at the base of dolomitic laminae that are interpreted as syndepositionally-dissolved evaporite crystals formed at the sediment-water interface. Dolomite is a commonconstituent of ancient lacustrinecarbonates associated with evaporite minerals or their pseudomorphs. The dolomite is typically attributed to elevated salinities in the depositional environment, forming by direct precipitation from an alkaline lake (Bradley and Eugster, 1969), by diagenetic alteration of pre-existing carbonate minerals by saline lake water (Desborough, 1978), by later groundwater alteration of lacustrine carbonates during periods of subaerial exposure (Wolfbauer and Surdam, 1974; Freytet, 1973), or by replacement of high-Mg calcite precipitated on marginal flats which was then transported to the lake by floods (Eugster and Hardie, 1975; Smoot, 1978). Van Houten (1962, 1964) argued that dolomitic layers dnd analcime in lacustrine sequences of the Triassic Lockatong and Brunswick Formations of the Newark basin, New Jersey and Pennsylvania indicate alkaline saline lake conditions. Concentric zones of zeolites and clays (especially trioctahedral smectites and sepiolite) have been reported from a number of Tertiary and Pleistocene deposits, commonly in association with saline minerals (Hay, 1964, 1966 1970; Sheppard and Gude, 1968, 1969, 1973a,b; Surdam and Parker, 1972; Surdam and Sheppard, 1978;

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Fig. 3.25. Probable saline pan deposits in the rock record. A). Saline pan crust of trona (?) pseudomorphed by calcite in the Cambrian Parakeelya Formation, Officer Basin, Australia. Radial crystal pattern is truncated by f lat surf ace with mudstone parting. Light layer in middle of sample may be a cumulus crystal bed. Compare tof ig. 3.80. Sample = 6 cm thick. (Photo courtesy of P. N. Southgate). B). Displacive andpoikilitic halite crystals (dark) in mudstone interbedded with saline pan halite of the Permian Zechstein Upper Halite, England. Scale =cm. (Photo courtesyof D. B. Smith). C).Displacive trona crystals in dolomitic mudstone interbedded with saline pan trona crusts and perennial lake trona cumulus crystal layers, Eocene Wilkins Peak Member, Green River Formation, Wyoming. Crystals may have formed in the perennial lake or saline pan subenvironment. Scale = cm. (Photo courtesy of L.A. Hardie).

Boles and Surdam, 1979; Smith, 1979; Lefond and Barker, 1985; Gude and Sheppard, 1986; Hay and others, 1986; Gall and Hyde, 1989; Stamatakis, 1989; Brell and others, in press). The zeolites commonly replace volcanic glass in tuffs. In some of these examples K-feldspars form the innermost of the concentric zones (Hay, 1970; Surdam and Parker, 1972; Sheppard and Gude, 1973a,b; Gude and Sheppard, 1986; Hay and others, 1986). The zonation is attributed in part to early diagenetic alteration of lake sediments under saline mudflat conditions. Hay (1964, 1966) reports examples of lacustrine beds with zeolites, but no evaporites, which he interprets as saline lake deposits. Turner-Peterson (1988) and Turner-Peterson and Fishman (1988) describe zoned zeolites and clays in volcaniclastic mudstones of the Jurassic Morrison Formation in the San Juan basin, New Mexico including a K-feldspar zone and an inner

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Na-feldspar (albite) zone. They interpret the Morrison Formation mudstone as saline lake or playa deposits, although they obsetved no evaporite minerals. Chert has been used as an indicator of saline conditions in some lacustrine deposits. Chert nodules have been attributed to the replacement of gypsum or anhydrite. Layers and nodules of chert have been interpreted as primary gel-like precipitates or have been attributed to the replacement of magadiite or some other Na-silicate precursor (Fig. 3.26; Hay, 1968, 1970; Eugster, 1969, 1980, 1985; Surdam and others, 1972; Eugster and Surdam, 1973; Sheppard and Gude, 1974, 1986; Eugster and Hardie, 1975; Wheeler and Textoris, 1978; Muir and others, 1980; Houser, 1982; Behr and others, 1983; Cheadle, 1986; Parnell, 1986a; Southgate, 1986; Schubel and Simonson, 1990). Surdam and others (1972) presented criteria for identifying magadiite replaced by chert in the geologic record which include: 1) soft-sediment deformation features of the putty-like magddiite, such as enterolithic folding, lobate protrusions, casts of mudcracks and crystals, and extrusion forms, and 2) contraction features due to the loss of volume in the transition to chert, such as reticulation cracks and polygonal ridges. It has yet to be proven that these features are unique to magadiite. Schubel and Simonson (1990) note that chert formed from magddiite has a characteristic rectilinear pattern of quartz crystal orientations visible in thin section that they attribute to pseudomorphing after the original spherulitic magadiite structure. Descriptions of ancient non-marine evaporites rarely distinguish between saline mudflat deposits and deposits of perennial saline lakes or saline pans. Evaporite traits

Fig. 3.26. Comparison of cherts from the Oligocpne Moonstone Formation, Wyoming (upper lef t) and Eocene Green River Formation, Wyoming (bottom) to chert formedfrom magadiite from Lake Magadi, Kenya (upper right). Irregular convolute surface and reticulatr shrinkage features are considered diagnostic by Srirdam and others (I972). (Photo courtesy of H.P. Eugster).

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suggesting vadose or phreatic conditions of formation have generally not been noted. Smoot (1977, 1983) noted vertical and lateral variations in evaporite characteristics and distributions in the Eocene Wilkins Peak Member of the Green River Formation, Wyoming that he attributed to differences in groundwater conditions in the saline mudflat (Fig. 3.24B). Textures indicating deposition in saline mudflats with efflorescent salt crusts were noted in the Triassic Blomidon Formation in the Fundy basin, Nova Scotia and the Triassic Bigoudine Formation in the Argana basin, Morocco (Smoot and Castens-Seidell, 1982; Smoot and Olsen, 1988). Dry mudflat deposits have been mentioned by many authors, but typical descriptions only note mudcracks and associations with evaporites. Many of the dry mudflat deposits appear to be superimposed on old lake bottoms (Plaziat, 1975; Stanley and Surdam, 1978; Van Dijk and others, 1978; Nickel, 1982; Platt, 1989). Deposits with characteristics suggestive of slowly aggrading dry mudflats (ZFig. 3.27A) are described in the lower Mesozoic Newark Supergroup, eastern North America (Van Houten, 1962, 1964, 1965a, 1980; Smoot and Katz, 1982; Katz, 1983; Demicco and Gierlowski-Kord- esch, 1986; Smoot and Olsen, 1988), the Triassic Kueper Marl, England (Arthurton, 1980), and the Devonian Orcadian basin, Great Britain (Tunbridge, 1984). Deposits with characteristics of rapidly aggrading dry mudflats (Fig. 3.27B) are described from the Wilkins Peak Member of the Green River Formation (Eugster and Hardie, 1975; Smmt, 1977, 1978, 1983), the Middle Proterozoic Belt Supergroup in northwestern U.S.A. (Winston, 1984; Winston and others, 1984), and the Precambrian Officer Basin, Australia (Southgate and others, 1989).

Fig. 3.27. A). Mudstone with cement-f illed vesicles and cracks (white) and sediment-- filled polygonal cracks (vertical features) indicating a slowly aggrading dry mudflat deposit. Siltstone layer (light) is broken by polygonal cracks and the ends are upturned. Sample =4 cm thick. B). Dolomitic mudstone transitional from shallow perennial lake (upper and lower dark beds) and dry mudf lat (middle light laminae) in the Eocene Wilkins Peak Member, Green River Formation, Wyoming. Sample = I0 cm thick.

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Shoreline deposits in the rock record that are associated with non-marine evaporites are typically laterally equivalent to fresher-water perennial lake sequences. Birdfoot deltas interbedded with laminated perennial lake deposits are reported from the Eocene Green River Formation, Colorado, Wyoming, Utah (Bradley, 1926; Fouch and others, 1976; Ryder and others, 1976; Smoot, 1983), the Tertiary Ridge Basin, California (Wood and Link, 1987) the Tertiary Ruby River Basin, Montana (Monroe, 1981), and the Permian Karoo Basin, South Africa (Van Dijk and others, 1978). Ancient "Gilbert-type"deltas have been described by Stanley and Surdam (1978); Link and Osborne (1978, 1982), McDonald and LeTourneau (1988), and Smoot (in press) (Fig. 3.28A). Fan deltas or bottom-modified "Gilbert deltas" have been reported by Link and Osborne (1978, 1982), Nemec and others (1980), Manspeizer and Olsen (1981), Pollard and others (1982), Postma (1983,1984), Schultz (l983), Nemec and Steel (1984), Postma and Roep (1985), and Hentz (1985). Smoot (in press) describes deposits interpreted as sheet deltas from the early Mesozoic Newark basin, Pennsylvania and New Jersey, and from the Jurassic Hartford basin, Connecticut (Fig. 3.28B). Sandstone sheets described by Hubert and Hyde (1982) in the Triassic Blomidon Formation, Nova Scotia and by Hubert and others (1978) in the Jurassic East Berlin Formation, Connecticut are also probably sheet deltas. Large-scale coarsening-upward sequences of mostly fluvial deposits have been interpreted as deltas (Manspeizer and Olsen, 1981; Brown, 1980), but these may be fluvial terminal fans (Friend, 1983). Wave-formed bars and spits have been described from the Tertiary Ridge Basin Group, California (Link and Osborne, 1978, 1982), Tertiary deposits in central Australia (Ambrose and Flint, 1981; Callen and others, 1986), and the lower Mesozoic Newark Supergroup, eastern North America (LeTourneau and Smoot, 1985; Smoot and LeTourneau, 1989). Tucker (1 978) described wave-cut terraces and associated wave-sorted gravel in Triassic deposits in Great Britain. Eugster and Hardie (1975) and Smoot (1983) described mud pebble shoreline ridges in the Eocene Green River Formation, Wyoming. LeTourneau and Smoot (1985) and Smoot and LeTourneau (1989) describe intercalations of wave-formed conglomerate, lacustrine shale, and subaerial alluvial fan conglomerate in the lower Mesozoic Newark Supergroup, eastern North America. Regressive wave-formed sandstone and conglomerate sequences in these cases are similar to those found in cores taken from the toes of modern alluvial fans in Walker and Pyramid lakes in Nevada. Examples of ancient shoreline tufas have been described by Bradley (1929), Wolfbauer and Surdam (l974), Surdam and Wray (1976), Link and others (1978), and Abell and others (1982). Many ancient lacustrine deposits, that are associated with non-marine evaporites, display cyclic patterns a few meters thick of subaqueous deposits alternating with

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Fig. 3.28. A). Two climbing-ripple cross-lamination packages separated by a brown mudstone layer from a low-angle delta foreset in the Jurassic Feltville Formation, New Jersey. Vertical structures are root casts. Compare to fig. 3.20A. Flow is to the left. Scale = 7 mm long. B). Sandstone interpreted as a sheet delta deposit with climbing-ripple cross-lamination and basal sof t-sediment deformation features overlying a massive mudstone with dry mudflat features and a shallow perennial lake shale. Entire sequence is cut by polygonal mudcracks. Jurassic East Berlin Formation, Connecticut. Pencil = I5 cm long.

subaerial deposits (Van Houten, 1962,1964;Picard and High, 1968; Eugster and Hardie, 1975; Hubert and others, 1976, 1978; Smoot, 1977, 1983; Clemmenson, 1978a,b, 1979, 1980; Arthurton, 1980; Behr and others, 1983; Olsen, 1984, 1986; Winston and others, 1984; Demicco and Gierlowski-Kordesch, 1986; Porada and Behr, 1988; Smoot and Olsen, 1988). The subaerial deposits include dry mudflats, saline mudflats, and saline pan deposits, as well as soils (Fig. 3.28). This cyclicity has generally been attributed to climatic fluctuations affecting the inflow-evaporation balance of closed basin lakes. Olsen (1986) related the periodicity of cycles in the lower Mesozoic Newark Supergroup, eastern North America to those of Milankovitch cycles of climatic forcing.

FLUVIAL DEPOSITS

Occurrences of non-marine evaporites in modern fluvial deposits are generally insignificant except where fluvial systems intersect evaporitic lacustrine systems or where saline soils are developed over them. Fluvial deposits, however, are an important part of the sedimentary record of non-marine evaporites and should be considered in the

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environmental interpretation. The arid to semi-arid conditions necessary to produce evaporites also control the types of associated fluvial deposits. Most fluvial sedimentation occurs during short-lived flash floods. These sedimentation events are separated by longer periods of low flow, or in the case of ephemeral streams and most alluvial fans, periods of complete dryness of the channels. Rainfall in poorly vegetated catchment areas causes flash-flood runoff into tributary channels producing one or more pulses of flow (Leopold and Miller, 1956; Sharp and Nobles, 1953; Wertz, 1966; Frostick and others, 1983). Flood waters are typically rich in suspended sediment. Very high sediment concentrations form debris flows (Johnson, 1970; Rodine and Johnson, 1976) or hyperconcentrated flows (Beverage and Culbertson, 1964; Pierson, 1980). Floods may last for several months, for instance the normally dry Cooper Creek in Australia flooded for over 12 months in 1949-1950 (Bonython and Mason, 1953) but, more typically, flows last for a few hours to a few days with large portions of the flows lost by infiltration into underlying dry sediment (Cooke and Warren, 1973). The literature on fluvial depositional systems is extensive and a rigorous overview is beyond the scope of this paper. Three varieties of fluvial deposits are considered: alluvial fans, ephemeral streams, and perennial streams. Each of these may be the dominant style of fluvial deposition in a non-marine evaporite setting or any combination of the three may exist together.

Alluvial Fan Subenvironment

Alluvial fans are coarse-grained, cone-shaped wedges of sediment that extend radially from canyon drainages in mountain fronts. Alluvial fans are commonly associated with non-marine evaporites due to the tectonic settings that produce closed basin hydrology. There is voluminous literature on the geomorphologic and sedimentary characteristics of alluvial fans (Bull, 1968, 1972, 1977; Cooke and Warren, 1973; Spearing, 1974; Rust, 1979; Nilsen, 1982; Nilsen and Moore, 1984, and Nemec and Steel, 1984). Alluvial fans become finer grained as they radiate from the mountain front changing from boulder-dominated to sand-dominated deposits over distances typically less than tens of kilometers. Alluvial fans are divided into three parts: the fan apex with deeply incised channels that come from the canyon mouth; the mid-fan where flows break out from incised channels (the intersection point of Hooke, 1967) and radiate over the fan surface; and the fan toe where the fan slope decreases and channels become shallow and complexly braided. Beyond the toe of some fans, the channels may disappear into a sandy plain called the sand flat (Hardie, 1973; Hardie and others, 1978).

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Alluvial fan deposits consist mostly of debris flows and channel-fill deposits of shallow braided streams. The sedimentary characteristics of debris flows have been described by Sharp (1942), Beatty (1963, 1968, 1970), Hooke (1967), Johnson (1970), Bull (1972), Rodine and Johnson (1976), Nilsen (1982), and Nemec and Steel (1984), among many others. Debris flow deposits occur as narrow levees bordering channels or as lobate bodies that may be very broad and sheetlike. They are typically poorly-sorted, matrix-supported conglomerates that drape underlying deposits. The largest clasts, commonly concentrated at the top of flows, create a "coarse tail" reverse grading and at the margins of flows make tightly packed ridges of boulders or cobbles. High-viscosity debris flows tend to form thick deposits on the steep (proximal) portions of fans while low-viscositydebris flows form thin, sheet-like deposits on the low-angle (distal) portions of fans. Debris flows with very low viscosities are gradational to hyperconcentrated stream flow deposits that have a poorly-developed horizontal lamination. The characteristics of stream deposits on alluvial fans depend on the fan slopes. Channels on steep fans are often erosional with thin veneers of sediment accumulating behind step-like ridges of boulders and cobbles (Bowman, 1977). Flood water seepage into the underlying sediment produces poorly-sorted mound-shaped deposits of cobbles and boulders, as much as a few meters thick, behind which finer-grained sediment accumulates. These mounds are termed sieve lobes (Hooke, 1967; Bull, 1972). Decimeter-scale sieve lobes occur within incised channels on steep fans and are surrounded by imbricated gravel and thin lenses of sand. The most common stream deposits on alluvial fans are braided channels consisting of coarse-grained longitudinal bars surrounded by finer-grained interbar sediment. Longitudinal bars form during peak flood stages; surrounding finer-grained sediment is deposited as flow wanes. Bars typically are finer-grained in the downstream direction and vary from well-defined ellipses in plan view to broad diffuse sheets (Williams and Rust, 1969; Rust, 1979; Bluck, 1982). Bars comprised of boulders and cobbles are commonly poorly-sorted with sand and silt filling interstices of the framework-supported clasts. The clasts may exhibit flow shadows, imbricate clusters (Dal Cin, 1968; Gustavson, 1974; Brayshaw and others, 1983; Brayshaw, 1984), and transverse ribs (Foley, 1977; Shaw and Kellerhaus, 1977; Koster, 1978). Longitudinal bars comprised of pebbles and sand commonly exhibit imbrication, horizontal bedding, and low-angle cross-bedding defined by alternations in grain size or orientation of elongate clasts. Interbar sediments are commonly sandy with horizontal to gently inclined lamination. In general, the shallow depths of flow, coarse grain size, and short duration of stream flow on arid alluvial fans do not allow for the development of lower flow regime bedforms (ripples and dunes) except for distal fan toe areas where ripples may cap

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channel-fill sequences. Stream flow during peak flooding often extends beyond the channels of the fan toe producing a sheet-like surfxe of water. These sheets of water flow over unchan- neled surfaces depositing thin tabular sand sheets that comprise the sandflat. Each flood produces centimeter- to decimeter-thick sand beds with erosional bases and thin, mudcracked, mud drapes. The thin beds may have internal horizontal lamination, low-angle inclined lamination from antidune flow, or poorly-defined lamination or coarse-tail grading from hyperconcentrated flow as flood water seeps into underlying sediment. Some thicker beds may exhibit "powering down" sequences of flat lamination to ripple cross-lamination to flat-laminated fine sand and silt. Debris flow deposits may plug channel systems causing streams to change flow directions, particularly below the intersection point. Subsequent debris flows follow the new stream channels and plug them causing areas of aggradation to continuously shift. A lenticular mosaic of debris flow and stream deposits thus forms. Thin debris flow deposits may be washed over by streams transporting the finer grained portions and leaving the coarse clasts as lags. Narrow ridges of tightly-packed boulders and cobbles or scattered outsized boulders surrounded by sandy and pebbly stream deposits reflect the reworking of debris flow levees and lobes. Other features of alluvial fans include: caliche crusts, cements, or nodules (Bull, 1972; Lattman, 1973;Glennie, 1970; Hardie and others, 1978);soil evaporites (see saline soil deposits); coarse deposits formed by deflation of fine sand and silt; eolian dunes and wind flats composed of reworked sand and granules (see eolian deposits); wave reworked fan deposits and deltaic deposits where fans intersect lakes (see shoreline deposits); desert varnish (Hunt and others, 1966; Glennie, 1970; Potter and Rossman, 1977; Dorn and Oberlander, 1981,1982; Dorn 1988); and travertine crusts and cements in spring-fed stream channels (Slack, 1967; Barnes and ONeil, 1971). The size and slope of alluvial fans depend upon the size of the drainage area, the type of sediment in the source area, and the amount of tectonic activity along the mountain front (Hooke, 1972; Heward, 1978; Hooke and Rohrer, 1979). Tectonic activity contemporaneous with fan deposition can help produce very thick deposits (Bull, 1972; Crowell, 1974; Steel, 1976) and can change channel gradients during deposition thus halting scarp retreat. Large-scale (10s to 100s of meters thick) cyclic sequences of coarsening and fining fan deposits have been attributed to tectonic activity during fan accumulation (Steel and others, 1977; Heward, 1978; Nilsen, 1982; Nemec and Steel, 1984; Blair, 1987). Evaporite deposition on alluvial fans is largely restricted to the fan toe and sandflat, except for saline soil development which may occur anywhere on the surface

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(see Saline Soil Subenvironment). Evaporites formed in distal fan deposits are mainly thin efflorescent crusts and tiny anhedral blebs of microcrystalline minerals. Eff lorescent crust growth deforms sedimentary structures (particularly the crests of ripples) and traps sediment on surface irregularities. Polygonal crust deformation is common where sandflats intersect saline mudflats and thin crusts warp silty mud drapes into irregular folds. Efflorescent salt growth in pebbly deposits, particularly debris flows, may disintegrate some clasts by evaporite growth along grain boundaries or small fractures.

Ephemeral Stream Subenvironment

Ephemeral streams are dry river channel systems that flood intermittently. The river systems are gradational to sandy low-angle alluvial fans, but differ in that they do not issue from a mountain front and may even pass through several highland ranges. Examples of major ephemeral stream systems associated with non-marine evaporites are the river systems entering Lake Eyre, Australia (Bonython, 1955; Williams, 1968, 1970a, 1971; Veevers and Rundle, 1979; Rust, 1981; Rust and Nanson, 1986; Nanson and others, 1986) and the wadis of North Africa (Glennie, 1970; Williams, 1970b;Abdullatif, 1989). Smaller ephemeral stream systems associated with non-marine evaporites include, for example, the axial drainages entering Lake Magadi, Kenya (Eugster, 1980) and Death Valley, California (Hunt and others, 1966). The sedimentary characteristics of ephemeral streams have been described by Leopold and Miller (1956), Williams (1968, 1970b, 1971), Karcz (1972), Frostick and Reid (1977), Rust (1981), Sneh (1983), Rust and Nanson (1986), Abdullatif (1989) and summarized by Glennie (1970), Picard and High (1973), and Tunbridge (1981). Most ephemeral stream systems consist of broad flat-bottomed channels with braid bars. The bar and interbar deposits are basically the same as for alluvial fans, but lower flow regime bedforms may be more important and overbank deposits may form adjacent to channels. Ephemeral stream deposits commonly consist of coarse-grained longitudinal bars that exhibit clast imbrication and horizontally laminated interbar sands. Large, flat-topped bars may also exhibit low-angle inclined lamination to planar tabular foresets. These foresets may dip downstream or may be oriented at right angles to flow across the channel. Mud drapes, up to several centimeters thick, are typically mudcracked and rip-up clasts of these may be incorporated into an overlying flood unit. Shepperd (1987) describes a meandering ephemeral stream in New Mexico that produces lateral accretion beds about 5 m thick. Williams (l968,1970a, 1971) described dune-scale bedforms 7-75 cm high and 1-8 m chords in a number of the ephemeral streams feeding Lake Eyre and Lake Frome.

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The stream sections that contained the dune-scale bedforms were 1-4 m deep during peak flooding that lasted up to 3 months. The stream deposits, up to 3 m thick and presumably deposited by a single flood, consist of trough cross-beds 5-25 cm thick and 40-120 cm wide overlain by and laterally gradational to tabular foresets of a wide variety of thickness and orientations reflecting deposition by longitudinal and transverse bars. Horizontal stratification is limited to the tops of bars and a few river stretches near the Lake Eyre basin. The Cooper Creek river system in central Australia consists of a mixture of braid patterns and dendritic sinuous channels (Veevers and Rundle, 1979; Rust, 1981; Rust and Legun, 1983; Rust and Nanson, 1986; Nanson and others, 1986). The braid bars are a mixture of sand and mud peloids (Rust and Nanson, 1989) that are largely structureless, except for deep, complex desiccation cracks (Rust and Nanson, 1986; Nanson and others, 1986). The sinuous channels are interpreted as anastomosing streams that form during falling flood stages. These are comprised of side channel bars forming tabular foresets that grade laterally into relatively small scale dune trough cross-beds. There is no single criterion by which ephemeral stream deposits can be distinguished from perennial stream deposits. The distinction can be made on the abundance of sedimentary features indicating dry conditions. Some of the sedimentary features include: mudcracked mud drapes; early cementation or case hardening reflected in steep scours and sand intraclasts; poor sorting; intercalations of wind deflation gravels or eolian sand beds or dunes (Glennie, 1970; Shepperd, 1987); and abundant mud clasts and peloids. Thick caliche or silcrete crusts and saline soils may develop on the overbank deposits. Root penetration and growth may disrupt entire channel deposits since flood water or rainfall is best stored in the channel sands making them preferred sites for plant colonization.

Perennial Stream Subenvironment

Perennial streams associated with non-marine evaporites are almost exclusively allogenic (Cooke and Warren, 1973) meaning that the source waters lie outside of the arid zone. The perennial rivers may flow from adjacent mountainous areas into basins in a fan-like manner such as the Bear, Weber, and Jordan rivers which flow out of the Wasatch Mountains into the Great Salt Lake (Eardley, 1938; Hahl and Langford, 1964) or the Golmud River which flows out of the Kunlun Mountains into the Qaidam basin (Petrov, 1967; Chen Kezao and Bowler, 1986). Perennial rivers may flow for long distances feeding several non-saline lakes before terminating into evaporitic deposits, such as the Jordan River that flows 105 km into the Dead Sea (Neev and Emery, 1967)

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or the Tarim River which flows over 600 km before losing its character in the Lop Nor Depression (Petrov, 1967). Most perennial rivers are low gradient and meandering as they enter the areas of evaporite formation although some rivers are braided or anastomosing. Rivers draining mountainous source areas have steep gradient braided channels or narrow incised channels in their proximal localities, but are meandering on the basin floor. We know of no studies of the sedimentary characteristics of perennial rivers in saline basins. There is abundant literature on the sedimentology of perennial rivers which will not be summarized here (Allen, 1965; Miall, 1977,1985;Jackson, 1978; Rust, 1978, 1979; Cant, 1982; Friend, 1983). Proximal deposits of perennial rivers initiating in mountains would probably resemble coarse-grained braided rivers (Williams and Rust, 1969; Bluck, 1979) or small flash flooding streams (Laronne and Carson, 1976; Martini, 1977). The distal perennial rivers would most likely be tightly meandering, high suspension load rivers (Taylor and Woodyer, 1978; Woodyer and others, 1979; Nanson, 1980; Nanson and Page, 1983) or anastomosing streams (Smith and Smith, 1980; Smith and Putnam, 1980; Smith, 1983). There is no mention in the literature of evaporites precipitating in perennial river deposits, but there is the possibility of evaporites forming in abandoned meanders (Wasson, 1983, and Callen and others, 1986) or by soil-forming processes on floodplains.

Ancient Fliivial Deposits

There are many papers on ancient alluvial fan deposits (Nilsen and Moore, 1984). Most descriptions emphasize poor sorting and coarse grain size using the catch-all term 'Tanglomerate". Megacycles of grain size variation in ancient alluvial fan deposits are commonly attributed to tectonic processes (see Heward, 1978). LeTourneau (1985) interpreted changes in the sedimentary character of alluvial fan conglomerates in the Jurassic Portland Formation, Connecticut as the result of climatic fluctuations, apparently coincident with transgression and regression of perennial lakes. Smoot (1983) described extensive sheet sandstones interpreted as sandflat deposits, in the Eocene Wilkins Peak Member of the Green River Formation, Wyoming. Hubert and Hyde (1982) interpreted sheet sandstones in the Triassic Blomidon Formation in the Fundy basin, Nova Scotia as sandflat deposits, but these are probably sheet delta deposits. Sheets of ephemeral braid stream deposits have been described by Hubert and others (1978), Veevers and Rundle (1979), Turner-Peterson (1980), Tunbridge (1981, 1984), Rust and Legun (1983), and Petit and Beauchamp (1986). Friend and others

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(1979), Stear (1983), and Flint (1985) describe alternating sheet and ribbon stream morphologies surrounded by mudcrdcked mudstones that contain evaporites. These stream deposits include braided sheets, lateral accretion sets of meandering patterns, and single channels with overbank "wings" that may be anastomosed channels. The streams may have been perennial or ephemeral and at least some of the interbedded mudstone appears to be lacustrine. Hubert and Florenza (1988) describe thick channel-fill sequences with large-scale trough cross-bedding in the Triassic Wolfville Formation in the Fundy basin, Nova Scotia which mostly underlie saline mudflat mudstone. However, there may be a vertical change from perennial to ephemeral rivers in that sequence. Stuart and Willingham (1984) describe the successive changes in fluvial style in Pliocene to Holocene deposits in two basins in Texas, which they attribute to the evolution from closed basins to the present open drainages. Butler (1984) noted variation in the fluvial character of the ancestral Amargosa River in Death Valley, California, that he attributed to responses to faulting. Perennial river deposits have generally not been identified as such in the rock record, but river deposits described in the Eocene Green River Formation in Co~orddo, Utah, and Wyoming by Ryder and others (1976), Surdam and Stanley (1978), and Smoot (1983) are probably perennial braided rivers. Smoot and Robinson (1988) and Smoot (in press) describe muddy river deposits, probably perennial meandering streams, that are associated with evaporite crystal casts in the lower Mesozoic Newark Supergroup in the eastern U.S.A., and Smith (1990) describes similar deposits in the Permian Karoo Basin, South Africa. Friend (1983) defined terminal fans as fluvial deposits that end within a basin in a manner similar to an alluvial fan. Ancient deposits with characteristics fitting this definition include those described by Friend and others (1979) in the Tertiary Ebro basin, Spain; by Winston (1977) in the Middle Proterozoic Belt Supergroup, northwestern U.S.A.; and by Parker and others (1988) in the Triassic Passaic Formation in the Newark basin, New York and New Jersey. Similar deposits described by Brown (1980) in the lower Mesozoic Argana Basin, Morocco were interpreted as deltaic, because they formed coarsening-upward sequences several tens of meters thick.

OTHER DEPOSITS

Although fluvial and lacustrine deposits typically dominate non-marine evaporite environments, significant accumulations of eolian or spring deposits may also occur. In some cases, eolian deposits may dominate the sedimentary record whereas spring deposits are usually less distinct as a separate subenvironment. Saline soils have

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received little attention from geologists but may represent the most extensive deposits in most settings. Since they are superimposed on other deposits, they are less easily separated as a mappable subenvironment and are included here mainly to emphasize that they exist and have been under-utilized by geologists.

Eolian Dunef ield and Sand Sheet Subenvironment

The eolian dunefield and sand sheet subenvironment (Fig. 3.29) is comprised mostly of sand derived from exposed fluvial deposits or mudflats. Large dunefields, referred to as sand seas or ergs, completely dominate the landscape in broad areas of relatively low relief such as North Africa, South Africa, and central Australia. In those areas, sediment deflation is an important basin-forming process, and lakes and playa flats commonly form in interdune areas. In mountainous areas, dunefields are smaller and restricted to the toes of alluvial fans. They are typically located near the ends of basins because of wind circulation patterns through narrow valleys. Eolian sand sheets (Fig. 3.29B) are sandy areas of low relief commonly adjacent to dunefields (Fryberger and others, 1979, Kocurek and Nielson, 1986). Eolian sand sheets may also cover considerable areas. The literature on eolian deposits is expanding rapidly. There are several volumes describing their distribution, origin, and sedimentary features (Bagnold, 1941; Glennie, 1970; Bigarella, 1972; Wilson, 1972; McKee, 1979; Ahlbrandt and Fryburger, 1982;

Fig. 3.29. A). Sand sea of barchan dunes at White Sands National Monument, New Mexico. Dune sands are mostly rounded gypsum derived from adjacent Lake Lucero. Deflation f lat in the f oreground is water-saturated gypsum sand. B). Eolian sand sheet at Salar de Coipasa, Bolivia. Small mounds trapped by vegetation are mostly gypsum sand. Areas between mounds are capped by an armor of deflation gravel. (Photo courtesy of G.E. Ericksen).

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Brookfield and Ahlbrandt, 1983; Brookfield, 1984; Hesp and Fryberger, 1988). Dunes are recognized by their cross-bedding which may attain thicknesses of several tens of meters. The tabular and trough cross-bedding and the distribution of soft-sediment deformation features have been noted by many (McKee, 1966, 1979,1982; McKee and Tibbitts, 1964; Bigarella, 1972; McKee and Moiola, 1975; Hunter, 1977; Ahlbrandt and Fryberger, 1982; Rubin and Hunter, 1985; Fryberger and Schenk, 1981,1988;Fryberger and others, 1983). Eolian dune foresets are commonly separated by large-scale planar bounding surfaces believed to represent deflation planes controlled by groundwater tables (Loope, 1984; Tdlbot, 1985) or, at least in some cases, the bounding surfaces of climbingdunes (McKee and Moiola, 1975; Brookfield, 1977;Kocurek, 1981,1984;Rubin and Hunter, 1982, 1984) including large scale complex dunes (drdas of Wilson, 1972). Interdune areas are sites of erosion by deflation or areas of sediment accumulation under generally lower wind stress than on the dunes. Interdune areas may have ponded water or a near surface groundwater table. Detailed descriptions of interdune deposits are included in McKee and ‘T‘ibbitts (1964), Glennie (1970), Ahlbrandt and Fryberger (1981), Kocurek (1981), Fryberger and others (1983,1984) and Lancaster and Teller (1988). Within large sand seas, interdune areas may cover considerable areas. Interdune deflation deposits commonly consist of pebbly lags overlying planar to trough-shaped erosional surfaces (McKee and Tibbitts, 1964; Cooke and Warren, 1973). Clasts may be polished by saltating sand and may have leeside sand shadows (Twidale, 1972; Lancaster and Teller, 1988). The clasts may be from streams, local bedrock, or fragments of caliche, cemented roots (dikaka), or silcrete. Ahlbrandt and Fryberger (1981) divide interdune deposits into dry, wet, and evaporitic varieties. Dry interdune deposits are similar to sand sheet deposits (see below), but are less widespread. Wet interdune deposits may form in relatively deep lakes, intermittent ponds, or moist areas near the water table (Fig. 3.30). If salinities are low, moist interdune areas become sites of abundant plant growth and the deposits are heavily bioturbated (Glennie and Evamy, 1968; Ahlbrandt and Fryberger, 1981). Otherwise, evaporites develop in the interdune areas as saline crusts and intrasediment crystals. These deposits include euhedral crystals formed from groundwater brine, most commonly as poikilitic growths encasing sand grains. Efflorescent salt crusts also encase sand grains forming polygonal ridge patterns that deform layering and act as sediment traps (see Fryberger and others, 1983, figures 25 and 26). Evaporitic interdune deposits have been noted in the Jaffura Sea in North Africa (Ahlbrandt and Fryberger, 1981; Fryberger and others, 1983,1984;Johnson and others, 1978) the margins of Lake Chad (Maglione, 1974,1980; Roche, 1977; Eugster and Maglione, 1979), the Killpecker Hills and the Nebraska Sand Hills (Ahlbrandt and Fryberger, 1981,1982), and the Kalahari

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WIND 5TEEP DUNE - /1 CRO 55BE DDI NO

Fig. 3.30. Relationship between eolian dune.s, stirid skeets, and suline lakes (subkhn)in the Jaj:fiira Sund Sea, Saudi Arabia. (A)is a cross section of' (B). (from Fryherger and others, 1983).

(Shaw, 1988). Adhesion ripples may be important structures in moist interdune deposits. These form irregular flat to inclined lamination that may be over one meter thick (Hunter, 1973,1980;Kocurek, 1981; Ahlbrandt and Fryberger, 1981; Kocurek and Fielder, 1982). Interdune lacustrine deposits form in isolated pools between dunes that fill during episodic floods (see Petrov, 1967, for the Tarim basin in central Asia) or from major lakes that drown the dunefield, as in Lake Chad. They typically consist of clay, silty clay, carbonate, and calcareous sand and silt (Ahlbrandt and Fryberger, 1981; Kocurek, 1981; Teller and Lancaster, 1986; Lancaster and Teller, 1988). Interdune lacustrine deposits characteristically range from massive to laminated and may contain lacustrine faunal remains. They may be disrupted by sand-filled polygonal cracks or they may consist of thin clay layers broken by desiccation into convex-upward lenses (Glennie, 1970). Sand sheets are areas of eolian sand accumulation without obvious dunes. Fryberger and others (1979) and Kocurek and Nielson (1986) describe the characteristics of these deposits. In general, they are dominated by flat to low-angle inclined strata comprised of interlayered coarse and fine sand. A diagnostic sedimentary feature associated with these layers is isolated ripple forms (Sharp, 1963) with long wavelengths

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and low amplitudes (Fryberger and Schenk, 1981,1988), that may be very coarse grained (Sakamoto-Arnold, 1981). Small scour and fill structures, root casts, burrows, normal and reverse graded laminae, and intercalated fluvial deposits are common. Many areas of the sand sheet appear structureless presumably due to the breakup of wind patterns by plants. Kocurek and Nielson (1986) argue that eolian sand sheets form in areas of wind transport where dunes are not produced due to a lack of fine-grained sediment or to the inhibition of sediment transport by high water table, plant growth, or surface armoring by coarse grains. Zibars are coarse-grained, low amplitude dune structures that are transitional to sand sheets (Nielson and Kocurek, 1986). Broad expanses of stony pavements comprised of pebbles and cobbles, such as the gibber plains of Australia (Mabbutt, 1968; Dury, 1968) and reg of North Africa (Cooke and Warren, 1973), have been attributed to wind deflation (Symmons and Hemming, 1968), but they may be coarse-grained eolian sand sheet deposits. These deposits may also form through matrix removal by flash flooding or rain splash or by upward migration of coarse particles in response to expansion and contraction of sediments during freeze-thaw or wetting and drying (Cooke and Warren, 1973). Eolian dune deposits may consist almost entirely of sediment deflated from adjacent mudflats. Gypsum-rich dunes, as much as 98% gypsum, have been noted from several localities (Jones, 1953; Bettenay, 1962; Tricart, 1967; Trichet, 1963; Eardley, 1962b; Fryberger and others, 1983; Glennie, 1970), most notably at White Sands National Monument, New Mexico (Talmage, 1932; McKee, 1966; McKee and Douglass, 1971; McKee and Moiola, 1975; Allmendinger, 1971; Simpson and Loope, 1985; Schenk and Fryberger, 1988). The White Sands dunes are fed by deflation of saline mudflat and saline pan gypsum from Lake Lucero. The grains are commonly rounded, but those in the interdune areas show evidence of euhedrdl overgrowths subsequent to deposition. Descriptions of other gypsum dunes also allude to roundness of grains and to early cementation. The internal structures of these dunes are generally the same as those of siliciclastic dunes. Clay dunes (parna dunes) are composed of fine sand-sized clay aggregates (Fig. 3.31). These are best known from Australia where they are a ubiquitous component of lunettes that border many playa flats (Butler, 1956, 1974; Bowler, 1973, 1986), but have also been described from other places (Roth, 1960; Price, 1963; Bowler, 1973). Clay pellets are apparently derived from mudflats as salt efflorescence disrupts the surface into puffy ground. Bowler (1986) argues that the formation of clay pellets is part of the chemical and morphological cycle of all Australian playas. The pellets may also comprise a significant part of evaporitic lacustrine deposits. Bowler (1973) noted that clay dunes differ from other eolian dunes in that internal layering is restricted to

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CONTROLLING WINDS

4 A' 0

Fig. 3.31. A). Comparison of clay lunettes (left)to barchan dunes. Clayddurzes are convex downwind and are typically dominated by low-angle cross-strata. Cross-section at base shows a barchan dune draped by a clay dune. (from Bowler, 1973). B). Thin-sectionof an eoliansand composed of clay pellets (dark) and roundedgypsum grains (white)from Lake Tyrrell, Australia. Field of view =3 mm wide. (Photo courtesy of J.T. Teller). low-angle inclined beds that occur on the lee side of the dune and that the convex part of lunettes is in the upwind direction (Fig. 3.31A). Clay dunes do not migrate once established and may consist of several layers separated by soil deposits (see Wasson, 1983). Fine grained eolian sediment (silt and clay) or loess is prominently mentioned in the Chinese literature (see Tungsheng, 1988) but barely mentioned elsewhere (Cooke and Warren, 1973; Glennie, 1970). Loess is apparently an important component of arid soils (Dan, 1973; Dan and others, 1982) and, as mentioned in the saline mudflat subenvironment, eolian dust may be an important contribution to efflorescent crusts. The recent surge in studies of modern eolian deposits has resulted in the

Data Center ,09126599985,[email protected], For Educational Uses 266 NON-MARINE EVAPORITES recognition of many ancient deposits (McKee, 1979). Glennie (1972, 1983a,b, 1987) describes dune deposits of the Permian Rotliegendes in the North Sea area that he interprets as an extensive sand sea formed adjacent to a large inland saline pan. Cores of the sandstones contain beds dominated by distorted fabrics which he attributed to adhesion ripples, but that are probably due to disruption by efflorescent salt crusts (see also the Permian Minnelusa Formation, Powder River Basin, Wyoming in Fryberger and others, 1983, figs 27 and 28). Kocurek and Hunter (1986) describe polygonal fractures with small anhydrite crystals within silica cement in the Triassic eolian Navajo and Page Sandstones, Arizona that they attribute to cementation by evaporite crusts. These polygonal fractures may be similar to root disruption and pedogenic gypsum in dune sands on the margin of Lake Eyre, Australia (upper part of the section shown in Callen and others, 1986). Ross (1983) and Ross and Chiarenzelli (1985) interpret Proterozoic sandstones of the Hornby Bay Group, Northwest Territories, Canada as erg deposits with associated lacustrine evaporite deposits. Other examples of ancient eolian deposits associated with lacustrine, fluvial, and/or alluvial fan deposits include the Proterozoic Masterton Sandstone, McArthur Basin, Australia (Jackson and others, 1987) the Devonian Caherbla Group, Ireland (Horne, 1975), Permian of the Arran Basin, Scotland (Clemmenson and Abrahamsen, 1983), lower Mesozoic of the Fundy basin, Nova Scotia including deflation pebble lags with ventifacts (Hubert and Mertz, 1980, 1984), the Triassic Gipsdalen Formation, Greenland (Clemmensen, 1978a, 1980), and the Tertiary of the Namib Desert in South Africa (Ward, 1988). There have been no reported Occurrence of gypsum dunes or clay dunes in the geologic record (although Glennie, 1970, states that McKee, 1954, observed possibly eolian cross-bedded gypsum in the Triassic Moenkopi Formation, southwestern U.S.A.). Ancient eolian sand sheet deposits have only recently been noted. Kocurek and Nielson (1986) cite ancient eolian sand sheet examples from the Triassic Dolores Formation, Colorado (Blodgett, 1988) and the Pennsylvanian-Permian Rico Formation, Utah (Loope, 1984). Eolian sand sheet deposits have also been reported from the Triassic Chinle Formation, New Mexico and Arizona (Dubiel, 1989; Dubiel and others, 1989). Muddy siltstones interpreted as loess deposits have been reported by Johansen (1988) and Johnson (1989).

Spring Suberi vironm en t

Springs are nearly ubiquitous in basins with lacustrine evaporites. As discussed in the section on hydrology, springs are local surface outlets of groundwater where the water table intersects the surface. Springs also form where groundwater flows upward under hydrostatic pressure along faults, along the margin of a brine lens, or along the

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boundaries between sediments or rocks with different permeabilities. Springs may be a major source of solutes for lacustrine brine, as well as sites of local mineral deposition (Lowenstein and others, 1989). The chemistry of springs reflect their flow path, for instance, cold dilute springs commonly emerge near the apices of an alluvial fans (Slack, 1967), hot acid springs or brines that rise along faults may be derived from deep hydrothermal groundwater reacting with bedrock or dissolving older evaporites (Degens and Kulbicki, 1973), or cold saline springs surfacing along the margin of a saline mudflat form by the mixing of dilute groundwater and stagnant basinal brine (Bowler, 1986, fig. 4). Springs may make local pools or marshes, feed small streams or sheets of water, seep directly into the surface sediments, or enter directly into a lake. In each case, the spring deposits will differ depending on whether spring water cools or warms, degasses, mixes with surface water, evaporates, or reacts with sediment. The most common and easily recognized spring deposits are tufa and travertine consisting of alkaline earth carbonates composed of mostly low-Mg calcite (Fig. 3.32; Slack, 1967; Irion and Muller, 1968; Julia, 1983; Chafetz and Folk, 1984, Shearman and others, 1989). Tufa is porous carbonate deposit that coats algal sheaths, moss thalli, or plant stems. Travertine is more compact carbonate consisting of bands of botryoidal fibrous crystals or micrite commonly with color or fluid inclusion banding. Tufa fabrics and travertine fabrics may be interlayered on a variety of scales producing stromatolitic mounds. Where springs emerge on alluvial fans or in stream channels the travertine and tufa may coat boulders, line channels, or form pisolitic sands (McGannon, 1975). Springs may form conical mounds where waters spill out from orifices and flow as sheets. Habermehl(l988, pers. comm.) describes spring mounds as much as 45 m high along the southern margin of Lake Eyre, Australia made of micritic marl with gastropods and other invertebrates, micritecemented eolian sands, terraced algal tufa, and pisolitic limestone. Risacher and Eugster (1979) describe broad spring flats in Pastos Grande, Bolivia, with tufa crusts, cemented pavements with tepee structures, and pisolites as much as 30 cm in diameter. In some cases, the pisolites developed on sandy deposits exhibit reverse grading. Tufa pinnacles may develop where springs enter lakes (Scholl, 1960; Scholl and Taft, 1964) and may reach heights of several tens of meters. Localized patches of coarse thinolite crystals (Radbruch, 1957) may also be lake bottom spring deposits, but formed as ikaite (CaCO, 6H2’O)during much colder climatic periods (Shearman and others, 1989). Siliceous tufa and sinter commonly form from hot springs (Weed, 1889; Walter, 1976; Walter and others, 1976). Saline springs may deposit a variety of minerals (Fig. 3.33). Eugster and Jones (1%8) describe siliceous gels formed where hot springs enter the Lake Magadi basin, Kenya. Muessig (1958, 1959,1966) described Recent mounds and sheets of borate

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Fig. 3.32. Spring travertine deposits associated with saline lakes. A). Stroniutolitic travertine-tufacomposed of calcite from an alluvial f an streambed f ed by a spring, Saline Valley, California. The carbonatesheet drapesthe channel and coats boulders. Pencil = 25 em long. B). Calcareous tufa towers up to I5 rn high, equivalent to Pleistocene deep-water perennial lake deposits, Searles Lake, Calif ornia. Tufas are linear along a fault in the floor of the basin and probably f ormed by mixing of spring water f rom the f ault with the overlying lake water. C). Spring f lat with travertine pisolites and tufa mounds composed of calcite at Pastos Crandes, Bolivia. Thisflat is adjacent to asaline pan. (Photo courtesy of H.P. Eugster).

minerals (primarily ulexite) forming spring aprons that are in places associated with carbonate tufas (Norman and Santini, 1985). Hot springs in the Danakil Depression in Ethiopia (Holwerda and Hutchinson, 1968) deposit halite, sylvite, carnallite, and bischofite. Hunt and others (1966) describe crystalline deposits of gypsum and crusts of glauberite, thenardite, halite, and trona formed in spring-fed marshes at the toes of alluvial fans in Death Valley, California. Reeves and Reeves (1971) describe spring seeps forming efflorescent crusts of halite and an unnamed alkali mineral. Lock (198821) describes a number of lakes in South Australia that are fed exclusively by acidic springs.

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Sediments precipitated in these lakes include alunite [KAI,(OH),(SO,),], jarosite [KFe,(OH),(SO,),], opaline amorphous silica, and gypsum. Hydrothermal springs entering the bottom of Lakes Tanganyika and Malawi in East Africa have made the lakes stratified and caused precipitation of Fe-silicate minerals and base-metal sulfides (Degens and others, 1971; Degens and Kulbicki, 1973; Miiller and Forstner, 1973; Stoffers and Hecky, 1978). Spring-fed ponds and marshes also may act as sites of high organic productivity including rich floral assemblages (Hunt, 1966), invertebrates, and vertebrates (Jones, 1965; Hunt and others, 1966; Beadle, 1974). Artesian springs whose outlets are in saline pans or saline mudflats commonly form cone-shaped solution pockets that may be filled with brine or with newly precipitated minerals (Fig. 3.33). The spring water may be undersaturated with respect to surface mineral assemblages causing dissolution, but mixing of the spring water with the local groundwater may cause precipitation of new minerals unrelated to the regional groundwater chemistry. Solution pockets 1 to 30 feet in diameter and filled with brine have been noted in saline pans of Death Valley, California (Hunt and others, 1966), salars of northern Chile (Stoertz and Ericksen, 1974), and Ceylon Lake, Canada (Last, 1989). These pockets commonly fill with evaporite crystal cumulates and crusts precipitated on the walls. These deposits have not been described in cross-section. Closely spaced circular depressions through which groundwater brines reach the surface have been noted in saline pans of the Basque Lakes, British Columbia, Canada (Nesbitt, 1974), Alkali Lake, Oregon (Rooney and others, 1969), and several lakes in southern Russia in the Kulunda Steppe (Strakhov, 1970). In the Basque Lakes, epsomite and bloedite precipitate in the depressions but gypsum forms in the surrounding mud. In Alkali Lake, thermonatrite and magadiite form only in the depressions and trona precipitates outside of them. In Malinvoe Lake (U.S.S.R.), pisolitic calcite forms in the depressions and dolomite dominates the surrounding mud, and in Kochkovoe Lake (U.S.S.R.), the depressions are filled with thenardite and the surrounding deposits are encrusted with halite. Qian Ziqiang and Xuan Zhiquiang (1985) describe borate ore bodiesfrom the Qaidam Basin of western China as nests with irregular roots surrounded by gypsiferous clays. They attribute the geometry of these bodies to local freshwater upwellings that mix with surface brine. There are few descriptions of spring deposits associated with ancient non-marine evaporites. Bradley and Eugster (1969) mention mounds in the Wilkins Peak Member of the Green River Formation, Wyoming as much as 15 m thick that are made of brecciated marlstone with silica veins. They interpreted these mounds as spring deposits. Smoot (1977, 1978) described sheet- and mound-shaped stromatolitic structures from the Eocene Wilkins Peak Member of the Green River Formation in

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Fig. 3.33. A). Saline springs at Salar Sun Marlin. Chile. forming small pools where groundwater comes up along the ,faulted busin rnurgin. Ridges around the edges of the pools are composed of coarse g.vpsum crystals. (Photo courtesy of G.E. Ericksen). B). Spring mound o,f' halite and svlvite in the Dunukil Depression. Ethiopia. (Photo courtesy of E. Bonatti). C). Circular depressions in the saline pan at Alkali Luke, Oregon, where artesian springs reach thesurface. The mineral assemblage in the depressions differsfrom that of lhe surrounding saline pan. Each depre.ssion is 3-5 m in diameter. (Photo courtesy oj B. F. Jones). D). Solution pipe in saline pan halite at Death Vu1le.y. California, where an undersaturated urtesian spring reaches the surface. The sides o,f the pipe are now lined with inward growing halite crystals. Tape measure is extended 25 cm.

Wyoming that he interpreted as spring travertine and tufa. These features coat boulders in conglomerate, line channel-form scours in sandstone, and cap polygonal cracks in mudstone. Preferential Occurrencesof sheet-like stromatolites underlying perennial lake laminites may reflect lateral migration of spring outlets as the groundwater table rose in response to lake transgression. Reverse-graded pisolites associated with finger-like stromatolitic tufas were observed in the Laney Member of the Green River Formation

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overlying the Wilkins Peak Member (Surdam and Stanley, 1979). These may be analogous to the pisolitic spring deposits of Risacher and Eugster (1979). Southgate and others (1989) observed laminated carbonates, similar to the Wilkins Peak stromatolites, in cores of the Cambrian Pardkeelya Formation in South Australia which they also interpret as spring deposits. Some ancient examples of tufa mounds mentioned in shoreline deposits may also have formed from springs (for instance, the boulder-coating stromatolites of Elmore, 1983). The organic-rich "algal swamp" facies of Carter and Pickerill (1985b) may also be a spring deposit. Hay and others (1986) interpret brecciated and fenestral carbonate mounds in Pliocene deposits in the Amargosa Desert, Nevada and California as spring deposits. Hot spring activity has been called upon for either the direct precipitation of Tertiary borates or at least as the source of solutes for lacustrine precipitates (Smith, 1985; Norman and Santini, 1985; Albayrak and Protopapas, 1985).

Saline Soil Subenvironment

Soils are primarily subaerial deposits produced by physical and chemical processes operating on rocks or sediments immediately below the geomorphic surface. Soil-forming processes include: shrinking and swelling of sediment due to wetting and drying; translocation of clays and solutes by percolating water; fracturing and disruption of materials by freezing and thawing, crystal growth, or the actions of plants or animals; and dissolution, precipitation, or recrystallization of minerals by pore fluids. Many deposits that we classify as dry mudflat, saline mudflat, and saline pan would also fall under the broad definition of soils. Further complications arise from the concept of aggrading soils, that are floodplain, eolian, or lacustrine deposits partially modified by soil processes. The purpose of this section is to introduce some basic information on saline soils and to highlight the evaporite mineral features produced by soil-forming processes. Several useful background references include Eriksson (1958), Dan (1973), Birkeland (1974), Dan and Yaalon (1982), and Dan and others (1982). Soil scientists have their own terminology for classification of soil profiles (Table 3.4) and for describing textures within soils (Brewer, 1964; Fitzpatrick, 1984). The important soil features examined in this section are: 1) soil churning and slickenside formation by wetting and drying, 2) precipitation of carbonate and silica nodules and cements, 3) solute movement and precipitation of saline minerals in soils, and 4) formation of thick saline mineral crusts. Other soil features noted in arid regions are the fracturing and crumbling of rocks by salt crystallization (Wellman and Wilson, 1965; Evans, 1970; Goudie, 1977,1983b; Goudie and others, 1979; Goudie and Day, 1981),

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TabIe 3.4. Comparison of nomenclature for saline soils (from Dan, 1973).

Common Terms FA0 Classification U.S. Classification

Coarse desert alluvium or Eutric or cdcaric fluviosols Tompsamments 01 gravelly desert alluvium todwents

Sand and sandy soils Psamments, mainly xeropsamments, tompsamments, quartzipsamments, and ustipsamments

Silty alluvium or silty calcaric flwisols Toduvents desert alluvial soils

Alluvial soils Fluvisols Flwents (mainly xercslwcnts and ustiflwents; and torrifluvents)

Lithosols Lithosols Lithic tomorthents, lithic xerorthents: and lithic ustorthents

Yermosols (mainly gypsic Camborthids or calciorthids; yermosols) or orthic solonchaks and some haplargids

(key, greyish brown, and red Yermosols: and some xerosols Argids and orthids (paleargids desert soils and paleorthids)

Saline soils. solonchaks Solonchaks Salorthids

Takyrs Takyric solonchaks Solarthids, natrargids, calciorthids, camborthids

Solonetz Solonetz Natrargids, natrixeratfs, natrustalfs, natraqualfs; and some natraquolls, natrustolls, and nattixerolls

Sdcdized soloneb Natrixeralfs, natrustalfs, natraqualfs

Solods Solodic planosols Natrixeralfs, natrustalfs, natraquatfs

Eutric or Calcic cambisols Eutric or calcic cambisols Ustochrepts, xerochrepts, ustropepts

Brawn and reddish-brown Xerosols, and poorly developed Argids or orthids; and poorly soils castanozems developed ustolls or xerolls

Rendzinas Rendzinas Rendolls

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Table 3.4. (continued).

Common Terms FA0 Classification U.S. Classification

Chestnut soils Castanoxms Ustolls, mainly calciustolls; sometimes xerolls (calcbterolls)

Clemozems C7lemoisems Ustolls, or sometimes xerolls and udolls

Brumllems Phaeo7~ms Udolis, mainly argiudolls; sometimes xerolls

Planosols Planosols Albaqualfs and glossaquafs

Brown forest soils Haplic phaeo7ems, haplic Haplustolls, calciustolls, castanomns, haplic haploxerolls, calcixerolls chernozems

Red and brown Chomic luvisols; hrunic Ustalfs or xeralfs Mediterranean soils luvisols

Vertisols Vertisols Vertisols

Brown vertisols Ckomic vertisols Chromustcrts or chromoxererts

Grey and black Pellic vertisols Pellusterts or pelloxererts hydromorphic vertisols formation of desert varnish (Whalley, 1983), and growth of plants. In arid and semiarid regions, sediment may be saturated with water and then subjected to periods of complete desiccation. Shrinkage of wet muddy sediment during desiccation results in polygonal cracks similar to those in dry mudflats (Sleeman, 1963). If the cracks are filled or partially filled with sediment and then rewetted, a compressional stress is exerted on the surrounding sediment (Dan, 1973; Cooke and Warren, 1973; Wilding and Tessier, 1988; contra Yaalon and Kalman, 1978). If the muds contain sufficient quantities of expanding clays (>30%, Yaalon and Kalman, 1978) and the depth of temporary water saturation is sufficient, the compressional stress will cause the formation of clay slickensides. The slickensides are apparently best developed some distance below the surface where confining pressure is sufficient to retard vertical expansion but not so great as to compact the sediment. If this process is repeated many times without burial, surface mounds and gullies called gilgae may develop (Hallsworth

Data Center ,09126599985,[email protected], For Educational Uses 274 NON-MARINE EVAPORITES and others, 1955; Harris, 1959, 1968; Cooke and Warren, 1973; Dudal and Eswaran, 1988). Slickensides in gilgae are characteristically long and arcuate, intersecting to form bowl shapes (Knight, 1980; Ahmed, 1983). This process may also push coarser particles upward, resulting in the formation of stony pavements (Springer, 1958; Mabbutt, 1968; Cooke, 1970; Yaalon and Kalman, 1978). Soils where this type of disruption dominates are termed vertisols. The maximum depth of disruption is generally believed to be around 3 meters but may be deeper, with the optimal wetting and drying depth of 1.5 m, depending on the sediment (Yaalon and Kalman, 1978). Slickensides may occur on dry mudflats, particularly those developed on perennial lake sediment. Two soil features commonly linked to arid environments are silcrete and caliche. Silcrete occurs as large tabular bodies of cemented sediment or carbonate replacements often in association with evaporite and eolian deposits (Cooke and Warren, 1973, p. 108-109; Smale, 1973; Goudie, 1973, 1983a; Summerfield, 1983a,b), but may also occur in humid and tropical climates. Although commonly assumed to be pedogenic, the origin of many silcretes is in doubt and some may be deposits of groundwater mixing with brine or spring deposits. Caliche or calcrete is a carbonate deposit that occurs as disseminated microcrystalline aggregates, nodular to pisolitic concretions, and massive to laminated crusts (Fig. 3.34; Gile and others, 1966; Krumbein, 1968; Aristarain, 1970; Goudie, 1973, 1983a; Lattman, 1973; Reeves, 1976; Esteban, 1976; Carlisle and others, 1978; Watts, 1977, 1978; Krumbein and Giele, 1979; Mann and Horowitz, 1979; Hay and Wiggins, 1980; Chafetz and Butler, 1980; Klappa, 1980; Esteban and Klappa, 1983; Arakel and McConchie, 1982; Mermut and Dasog, 1986; among many others). Like silcrete, calcrete is not restricted to arid or evaporitic settings and appears to have many origins, not all of which are pedogenic. Gile and others (1966) described caliche profiles in loose sediment that represent progressive stages of development from disseminated patchy micritic cements to small isolated nodules to tightly packed large nodules culminating with a cap of massive to laminated pavements. Tightly packed nodular masses may form polygonal tepee structures (Watts, 1977; 1978). Caliche nodules may also fill or line root cavities (Esteban and Klappa, 1983) or occur randomly scattered (Mermut and Dasog, 1986). Caliche may also form brecciated masses on carbonate rocks or older nodular to massive profiles. These deposits are commonly pisolitic or laminated (Esteban, 1976; Chafetz and Butler, 1980). Thick (up to 10 m), massive to laminated carbonate deposits that cement channel sand and conglomerate and form sheetlike linings may be groundwater deposits that are not pedogenic (Gevers, 1930; Lattman, 1973; Carlisle and others, 1978; Mann and Horowitz, 1979; Arakel and McConchie, 1982; Arakel, 1986). This type of caliche is common on alluvial fans and

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Fig. 3.34. Examples of caliche in deposits associated with saline lakes. A). Carbonate caliche cementing an alluvial f an conglomerate in Saline Valley, California. Notegreater abundance of carbonate (light matrix) in the upperpart of section and thicker carbonate rinds (arrows) on undersides of some boulders. Thickness shown is about 50 cm. B). Coarse gypsum crystals filling root casts (white) in a saline soil developed on a low-angle alluvial fan adjacent to Lake Frome, Australia. Note irregular boundaries of lighter calcareous blocky zones and darker clay rich zones and area of gypsum veins (arrow). Lens cap is about 3.5 cm in diameter. ephemeral stream channels entering saline basins and may form from groundwater, restricted to porous deposits, that becomes more concentrated downflow. Carlisle and others (1978) suggest that some thick gypcrete deposits may also form in this manner. Large quantities of sediment and solutes (in the form of grains and aerosols) are introduced to desert soils by the wind. These deposits may be transported from the surface into the soil by downward percolating rainwater or water from temporary lakes. Clay particles transported in this manner coat pores, grains, and fractures and are called cutans (Brewer, 1964). Accumulation of these clays leads to the formation of clay-rich or argillic soil horizons. Solutes transported by downward percolating water may be sufficiently concentrated to form saline minerals as the pore water evaporates. The most soluble minerals are stripped from the surface first and concentrated at lower levels, while less soluble minerals may reach saturation in the pore water at shallower depths (Drever and Smith, 1977). If airborne solutes are abundant and rainfall infrequent, a vertical profile of evaporites may form with the most soluble salts forming at the greatest depth (Fig. 3.35; Dan, 1973; Dan and Yaalon, 1982; Reheis, 1984). This is opposite the sequence predicted from groundwater concentration in a saline mudflat. A wide variety of salts may precipitate under these conditions (Yaalon, pers. comm.), but gypsum and halite are the most common. Crystals formed under these conditions

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EUIVlAL A HORIZON A

821 ILLWlAl 8 HORlZON

822 ca )I I) J 1) PRYSMATIC OR STRUCTURE Bcw(y SXWTURE CHARACTERISTICS

Cvca

HASSNE LAYERS c12CS c HORUON J LIME -I C13M ACCUHULAnON 10O']O01 GVWH ff SOLUABLE COHWUNOS

SALINE LAYER

PARENT MAERIAL c2

Fig. 3.35. Vertical profiles in saline soils. Schematic profile of a typical saline soil. Calcite and gypsum overlie halite in the vertical sequence. Thickness shown = I m. (from Dan, 1973).

include: 1) powdery microcrystalline aggregates, 2) thin fracture lining crystalline cements (particularly fibrous vein-filling fabrics), 3) large euhedral crystals that grow in sediment displacively; fill openings such as cracks, burrows, or root casts; or poikilitically encase sediments (Barzanj and Stoops, 1974; Stoops, 1978; Dan and Yaalon, 1982; Dan and others, 1982; Reheis, 1984; Murphy and others, 1985; Callen and others, 1986). Soil evaporites may also be distinguished from saline mudflat or saline pan evaporites by: 1) evaporite minerals may not show vertical variations in size or shape, 2) evaporite minerals may occur with root structures and minerals reflecting a less saline groundwater, and 3) evaporite beds may cut across stratigraphic units and occur within fluvial facies because their distribution is related to surface topography and not the groundwater table. Extensive crusts of gypsum (Carlisle and others, 1978; Watson, 1983a,b, 1985; Callen and others, 1986), halite (Gevers and Westhuyzen, 1931; Watson, 1983b), and nitrate salts (Ericksen, 1981; Garrett, 1985), or mixtures of these are attributed to accumulation of airborn solutes (Fig. 3.36). These crusts drape topography, independent of the groundwater table, and overlie a variety of sediments and bedrock. The solutes apparently initially adhere to surfaces as particles and recrystallize into a porous framework by re-solution from rainfall or fog. The surface deposits are typically fine grained, but buried deposits may be coarsely crystalline and develop secondary diage-

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A B C

Fig. 3.36. Vertical profiles of soil gypsum crusts in North Africa and the Namib Desert. Croute de nappe is coarse intrasediment gypsum. (from Watson, 1985).

netic crystals (Watson, 1983a,b, 1985). Sulfate salts may also form by oxidation of sulfides during weathering. Large gypsum crystals are common in weathered outcrops of marine black shales in arid and semiarid conditions. These crystals presumably form as oxidized surface water interacts with sulfides in the rock. Epsomite, bloedite, and hexahydrite also develop by this process (Mazor and Mantel, 1966; Tien, 1971; Murata, 1977; Lock, 1988b). Lock (1988b) described gypsum and epsomite forming in the upper 2 m of weathered Creta- ceous shale covering an area of over 5100 km2 within the Lake Eyre drainage in Australia. Ancient soils have been recognized mostly from carbonate nodules and root casts (Steel, 1974; Blodgett, 1988). Hubert and others (1978) described vertical sequences of carbonate nodules in pebbly sandstone that are comparable to the caliche sequences of Gile and others (1966). Curved synsedimentary slickenside structures have been found associated with carbonate nodules and root structures (Goldbery, 1982a,b; Smoot and Olsen, 1988; Gray and Nickelsen, 1989) and with complex cracking similar to that observed on dry mudflats (Van Houten, 1962, 1964; Smoot and Olsen, 1988). Chert layers interpreted as silcrete associated with evaporite deposits are reported by Ross and Chiaranzelli (1985), Ambrose and Flint (1981), Callen and others (1986), and Jackson and others (1987). Carbonate layers interpreted as calcrete or caliche are reported by Tucker (1978), Freytet (1973), Smoot (1978), Hay and Reeder (1978), Nickel (1982),

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Hay and others (1986), among many others. Goldbery (1982a) and Lehman (1989) are the only examples of ancient evaporites that are directly attributed to pedogenic conditions, but there are other examples of evaporite pseudomorphs associated with root structures and fluvial deposits that may be pedogenic (Flint, 1985; Smoot and Olsen, 1988; Gore, 1988a,b; Smith, 1990). Evaporite pseudomorphs in the upper Triassic Newark Supergroup rocks of the Culpeper and Newark basins in the eastern U.S.A., that occur discontinuously in a variety of sedimentary facies, exhibit no vertical changes in size or shape, and appear to form zones that are not parallel to bedding, are probably pedogenic.

DIAGENESIS

Evaporite deposits may undergo major changes in mineralogy, texture, and porosity following deposition. These diagenetic changes are: 1) syndepositional and controlled by processes operating in the depositional environment, or 2) post-burial and produced in the subsurface burial environment (Hardie and others, 1985). Early diagenetic features are well documented through studies of modern evaporite environments, but much less is known about the processes involved in the burial diagenesis of evaporites. The distinction between syndepositional and burial diagenetic features in ancient evaporites is not always possible because the timing of formation of such features may not be unambiguously documented. Syndepositionaldiageneticfeatures reported from modern non-marine evaporites include: 1) dissolution textures and fabrics, 2) cements, such as crystal overgrowths or cavity fillings, and 3) intrasediment growth of saline minerals (as euhedra, nodules, and fine-grained precipitates), alkaline earth carbonates, and authigenic zeolites, clays, and silicates. Syndepositional dissolution textures and fabrics are common in saline pan and saline mudflat deposits because dilute floodwater may dissolve surface saline crusts. The textural features produced in saline pans during flood stage include: 1) horizontal surfaces that sharply truncate crystal frameworks, 2) rounding of single crystals, and 3) horizontal and vertical dissolution cavities between vertically oriented crystals and along bedding planes (Fig. 3.10A). Such features have been described for modern halite (Lowenstein and Hardie, 1985), trona (Eugster, 1980), and gypsum (Warren, 1982, 1985). Dissolution of efflorescent crusts in saline mudflats during floods may produce small-scale faults and irregular lateral changes in the thickness of channel and sheet delta deposits. Syndepositional cements have been reported in modern evaporites for the

Data Center ,09126599985,[email protected], For Educational Uses DIAGENESIS 2 79

minerals halite (Shearman, 1970; Strakhov, 1970; Lowenstein and Hardie, 1985; Casas and Lowenstein, 1989), trona (Eugster, 1980), and potash salts, such as carnallite (Strakhov, 1970; Valyashko, 1972a; Yuan and others, 1985; Lowenstein and others, 1989). Halite forms inward-growing euhedrally terminated crystals lining solution cavities and intercrystalline voids (Fig. 3.10D) in saline pan crusts (Lowenstein and Hardie, 1985). Trona crusts at Lake Magadi, Kenya (Eugster, 1980) and potash salts, such as carnallite, form anhedral pore-filling cements in halite crystal frameworks which precipitate from groundwater brines beneath the saline pan surface during the desiccation stage (Strakhov, 1970; Valyashko, 1972a; Yuan Jianqi and others, 1985; Lowenstein and others, 1989). In Great Salt Lake, Utah, clear halite, apparently precipitated from bottom brines, has cemented layers of cumulate halite (Gwynn and Murphy, 1980). Halite crusts from modern saline pans show a marked decrease in porosity with depth (Casas and Lowenstein, 1989). In contrast to surface crusts with porosities commonly in excess of 50%, halite layers at about 10 m depth from Saline Valley and Searles Dry Lake, California, Lake Uyuni, Bolivia, and Qarhan Salt Plain, western China, have porosities of less than 10%. Halite layers from Saline Valley are tightly crystallized and contain no visible porosity at all depths below about 45 m. In halite layers at these depths, pore spaces are filled by clear halite cements that are most easily identified where they fill cavities (Fig. 3.37). Syndepositional diagenetic growth of saline minerals commonly occurs in muds interlayered with halite crusts in modern saline pans as well as in surrounding brine-soaked sediments of saline mudflats (Gornitz and Schreiber, 1981; Handford, 1982a,b; Lowenstein and Hardie, 1985). Euhedral crystals and nodular masses of evaporite minerals are common syndepositional features and euhedrdl overgrowths on intrasediment evaporites near the surface have been noted (Arakel and McConchie, 1982). Irion and Miiller (1968) report the formation of finely crystalline polyhalite, huntite, dolomite, and magnesite in mud of the saline mudflats in Tuz Golu, Turkey. Zeolites, clays, and other silicate minerals have also been widely reported as precipitates in saline lake mud below the surface (Hay, 1964, 1970; Millot, 1964; Sheppard and Gude, 1968, 1969, 1973a,b, 1974; Deike and Jones, 1980; Spencer, 1982; Jones, 1986; Jones and Galan, 1988). Mechanisms for producing diagenetic alteration of evaporite assemblages in non-marine settings are poorly understood. Saline brine formed by evaporation at the surface may affect older deposits by sinking due to its high density. These brines may mix with deeper groundwater and cool, both of which could lead to supersaturation with respect to saline minerals. Such a mechanism appears to be particularly important for

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Fig. 3.37. Large halite crystals (30-50 cm on the side) encasing saline pan layers about 10 rn below the modern surface in a quarry at Salar Grande, Bolivia. (Photo courtesy of G.E. Ericksen).

the formation of potash salt cements (i.e., carnallite) whose solubilities are strongly controlled by temperature. Spencer (1982) interprets intrasediment gypsum of the modern Great Salt Lake, Utah, as formed by mixing of surface brine and interstitial pore water along a diffusion gradient. The diffusion of lake bottom brine into underlying sediment may also result in the formation of authigenic clay minerals and silicates, including illite, stevensite, kerolite, nontronite, and sepiolite (Jones and VanDenburgh, 1966; Jones and others, 1972; Gac and others, 1977; Jones and Weir, 1983; Spencer, 1977,1985; Yuretich and Ceding, 1983; Von Damm and Edmond, 1984; Spencer and others, 1985; Jones, 1986; Jones and others, 1986). The criteria for recognizing post-burial alteration of evaporites are summarized in Hardie and others (1985) and include features that disrupt, deform, or completely destroy depositional and syndepositional diagenetic textures and fabrics. Evaporites composed of massive crystalline mosaics which lack primary sedimentary structures and which crosscut bedding may be produced by recrystallization during burial, particularly at elevated temperatures and pressures. Individual crystals in such massive mosaics may display sutured, interpenetrative grain boundaries (Shearman, 1985, fig. 8) or polygonal mosaic textures (Hardie and others, 1985). Polygonal mosaic textures in halite consist of equigranular anhedral crystals that meet at triple junctions that approach 120" angles. This recrystallized halite lacks all primary features such as vertical chevron fabrics and

Data Center ,09126599985,[email protected], For Educational Uses DIAGENESIS 281

fluid inclusion banding. Additional burial diagenetic features in evaporites include deformation structures, such as flattening of grains, leading to the formation of foliation, folds, flow banding, and pressure solution surfaces (Balk, 1949; Schwerdtner, 1966; Atwater, 1968; Dutton and others, 1982; Schreiber and others, 1982). Saline minerals crystallized in fractures that crosscut layering and primary sedimentary structures are commonly interpreted as burial diagenetic features, although similar features are found in saline soils. Gypsum and fibrous aggregates of ulexite fill fractures in shale in the Pliocene Terry borate deposit of the Death Valley region, California (Barker, 1980) and borax fills fractures in the Tertiary borate deposits of Kramer, California (Bowser and Dickson, 1966). Crosscutting zones of different evaporite mineralogies may reflect late diagenetic processes, as in the Kramer borate deposit (Siefke, 1985), but they could also be due to the local influence of springs. Dissolution of buried saline lake evaporites may occur thousands of years after deposition when dilute pluvial lakes fill the basin. Undersaturated water flowing through buried deposits may cause solution collapse breccias (Dyni, 1974; Olsen and others, 1989) and the resulting brine, produced by the dissolution, may form new evaporites. The replacement of evaporites by saline minerals or non-evaporite minerals is a commonly noted synsedimentary and post-burial diagenetic feature. The new mineral may or may not mimic the form of the original mineral. In modern saline mudflats, evolving groundwater brine reacts with previously formed minerals, precipitating new minerals (for instance, gypsum to glauberite, Hardie, 1968). The new minerals do not form pseudomorphs, but they may follow the distribution of the original mineral. Distinguishing between synsedimentary and post-burial replacement of saline minerals in ancient deposits is difficult. For instance, shortite in the Wilkins Peak Member of the Eocene Green River Formation, Wyoming is probably a post-burial diagenetic mineral (Eugster, 1971) but its texture and distribution suggest formation by syndepositional growth in a saline mudflat (Smoot, 1983). Dehydration of hydrous minerals such as gypsum to anhydrite, may or may not produce crystals that resemble the original mineral. Synsedimentarydissolution of a mineral may form a void that may subsequently be filled with another mineral such as glauberite filling mirabilite molds in the saline mudflat. Neev and Emery (1967) suggest that calcite associated with aragonite in bottom mud of the Dead Sea, Israel is a synsedimentary replacement of gypsum originally precipitated at the lake surface and then converted to calcite by chemical reactions initiated by sulfate-reducing bacteria on the lake bottom. Many ancient non-marine evaporites are known only by their molds or by pseudomorphs after them. The minerals are commonly recognized by their crystal habit,

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their gross morphology, or their distribution, although the criteria for distinguishing between mineral types are often ambiguous or not clearly stated. Common replacement minerals include calcite, dolomite, and chert (Fig. 3.38). The pseudomorphs may consist of single crystals with the original saline mineral morphology or may consist of cements filling a former void (Fig. 3.38D,E). Some replacements are thought to reflect specific mineral reactions in response to changes in the chemical environment. For instance, Eugster (1969, 1970) described the conversion of magadiite to chert by freshwater leaching of sodium from the magadiite. A similar mechanism has been suggested for the conversion of sodium carbonate minerals to calcite (Bradley and Eugster, 1969; Southgate and others, 1989). Folk and Pittman (1971) suggested that chert preferentially replaced gypsum in the form of length-slow chalcedony. Chert nodules in sedimentary rocks have often been interpreted as former anhydrite nodules (after Chowns and Elkins, 1974). Small gypsum or anhydrite laths within the nodules are considered strong evidence in support of this interpretation. Many authors have suggested that ancient chert deposits reflect alkaline lake conditions, as former magadiite (Surdam and others, 1972; Eugster and Chou, 1973; Sheppard and Gude, 1974; Parnell, 1986a) or as precipitates of amorphous silica (Wheeler and Textoris, 1978; Cheadle, 1986; Southgate and others, 1989). Even in heavily metamorphosed sequences, non-marine evaporites are inferred from abundant chert or minerals with high sodium or boron (Behr and others, 1983; Eugster, 1985; Porada and Behr, 1988).

DISTRIBUTION OF SUBENVIRONMENTS

Every non-marine evaporite deposit forms under unique physical and chemical conditions. Combinations of subenvironments, and the size and vertical and lateral distribution of subenvironments are dependent on tectonics, climate, hydrology, surrounding bedrock, and many other factors. There are certain associations of subenvironments, however, that result from specific basin geometries, hydrologic constraints, and chemical variables. Several modern and ancient examples of these associations are presented as illustrations.

Alluvial Fan - Saline Pan (Dry Mudflat)

Deep basins bounded by mountains form by faulting in a variety of tectonic settings including rifts (i.e. East Africa), thrust belts (i.e. the Andes), and complex strike-slip/extensional settings (i.e. Basin and Range, western U.S.A.). The central floor of these basins may be occupied by saline pan, saline mudflat, or dry mudflat deposits,

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Fig. 3.38. A). Thin section of bladed crystals of gypsurn(?) pseudomorphed by ferroun dolomite, Proterozoic Amelia Dolomite, McArthur Basin, Australia. Width = 3 mm. (Photo courtesyof M.J. Jackson). B). Rosettes of gypsum(?)pseudomorphed by chert, Proterozoic Woologorong Formation, McArthur Basin, Austrulia. Scale = cni. (Photo courtesy of M.J. Jackson). C). Discoidal gypsum crystals pseudornorphcd by f erroan dolomite overlain by chert nodules interpreted as pseudornorphs after unhydrite, Amelia Dolomite, McArthur Basin, Australia. Scale =cm. (Photo courtesy o,f M.J. Jnckson). D). Calcite crystals (light) pseudomorphinggypsum in mudstone of the Triassic Passnic Formution, New Jersey, Field of view =2 mm wide. E). Gluuberite(?) crystal moldfilled with the cement sequence of dolomite (I), unulaime (2), albite (3), and calcite (4) in mudstone of the Triassic Lockatong Formation, New Jersey. Field of view = 2.5 wzin wide.

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and is surrounded by alluvial fans extending from the mountain fronts. The tectonically subsiding basins act as sinks for sediment and may accumulate great thicknesses over a relatively short time. The mountains act as traps for moisture and provide a hydraulic head to promote the flow of groundwater to the arid center of the basin. This may produce a tremendously thick evaporite deposit. Dry mudflats develop in this setting when the groundwater level is not near the surface. Saline Valley, California is an example of a saline pan-saline mudflat in a block-fault graben (Fig. 3.39). The alluvial fans are dominated by debris flow deposits, so they are short and steep with narrow or absent sandflats. The alluvial fans extending along the basin axis are less steep with more stream deposits and their sandflats are mostly covered by eolian dunes. Small eolian dunes and eolian sand sheets occur beyond the toes of most of the fans. The saline mudflat covers a large part of the basin floor. It is dissected by sinuous channels that feed small sheet deltas at the margins of the saline pan. Springs form travertine and tufa deposits along the toes of alluvial fans and artesian springs in the saline mudflat form solution pockets filled with coarse crystals and windblown sand. A spring outlet along a fault on the western side of the basin produces a freshwater marsh with abundant plants almost adjacent to the saline pan. The central saline pan is comprised of halite beds separated by mud with abundant intrasediment crystals of halite, glauberite, and thenardite. This grades out to the saline mudflat in which glauberite crystals and molds of mirabiiite filled with glauberite are common. A zone of layered gypsum, probably a former gypsum-dominated pan, parallels the southern border of the saline pan. The outer portion of the saline mudflat is dominated by gypsum that is increasingly finer crystalline towards the basin margin. The crystals in the vadose portion of the saline mudflat are mostly subhedral and microcrystalline, except within artesian spring depressions. Sand patches from the efflorescent crusts are larger towards the basin margin and polygonal deformation is more obvious. Deformation by efflorescent crusts is common in fan toe sediments and in the eolian sand sheets. A core taken on the eastern side of the basin with abundant saline pan crusts in the subsurface suggests that the saline pan has shifted west due to faulting. Comparable alluvial fan to saline pan or dry mudflat deposits are very common and include: most of the basins in the Basin and Range and in the Rio Grande Rift of North America, most salars in the Andean Ranges of South America, and many of the basins in the East African Rift, southwestern Asia, and the Tibetan Plateau (see Table 3.1). Alluvial fans are restricted to the edges of these basins because streams that move coarse detritus into the basin center also cause ponding of water in the basin which decreases the streams sediment-transport capabilities. Ephemeral streams typically

Data Center ,09126599985,[email protected], For Educational Uses ALLUVIAL FAN - SALINE PAN (DRY MUDFLAT) 2 85

Saline Valley, California

36"45'

-N

* Spring seeps [=3 Bedrock Spring pond-marsh = Gypsum saline pan Alluvial fan Saline mudflat #Channels Sandflat Halite saline pan Division between Eolian Sandsheet Spring travertine /" / p vadose and phreatic @ Dry mudflat Dune field Sheet delta

Fig. 3.39. Schematic map of sedimentary subenvironments in Saline Valley, California. Cross-section is exaggerated vertically to show lateral relationships. Mineral zonation is shown in fig. 3.12. Dashed line shows approximate boundary between modern phreatic conditions at the surf ace and vadose conditions. (modified from Hardie and others, 1978).

Data Center ,09126599985,[email protected], For Educational Uses 286 NON-MARINE EVAPORITES enter the basins along basin axes or gaps between mountain ranges. The basinal extent of coarse-grained alluvial fan or stream sedimentary wedges may be considerably greater in broad basins than in narrow basins, since more water is needed to produce a lake or to change the lake area (Langbein, 1961; Smoot, 1985). Eolian deposits may also be more extensive in broad basins because wind patterns are not as constricted by mountains. The frequency and extent of surface flooding dictates the relative size of saline pans and saline mudflats in this setting. For instance, floods that frequently cover the entire basin floor produce broad saline pans and narrow saline mudflats (i.e. Lake Uyuni, Bolivia).

Alluvial Fan - Perennial Stream - Perennial Saline Lake

This association is actually a variation of the intermontane basins with saline pan and mudflat subenvironments. The important difference is that a perennial river entering a closed basin produces a permanent lake. Great Salt Lake, Utah is a broad, shallow lake fed by three perennial streams (Fig. 3.40). Alluvial fans extend from the Wasatch Mountains to the east, and to the west, talus cones form around smaller bedrock hills that separate the modern lake from broad exposed flats composed of old lake deposits. The perennial streams are high-gradient braided streams as they come out of the mountains, but they rapidly change to highly sinuous, high suspension-load meandering streams on the basin floor. These meandering streams form birdfoot deltas where they intersect the lake. Frequent changes in lake level probably make the birdfoot delta deposits look more like muddy stream deposits with fine-grained floodplains, since desiccation and soil development disrupt the lacustrine clays during lake lowstands. Shoreline deposits are dominated by oscillatory ripple flats, spits, and bars comprised of ooids and brine shrimp pellets. These are cemented into crusts by carbonate minerals and are commonly broken into flat clasts. Tufa mounds also occur in shallow water, commonly coating clasts or partially eroded pavements. In deeper water, rippled sand alternates with mud layers and many beds are disrupted by dissolved evaporites. Bedded halite occurs in the deepest water pockets, preserved by the lake stratification. Cores from Great Salt Lake (Fig. 3.40B; Spencer, 1982) have layers of intrasediment gypsum and scattered euhedral gypsum crystals in the upper meter or two of sediment (Unit I). This sediment also exhibits evaporite disruption features and pinch-and-swell bedding. Finely laminated aragonitic clay with algal mat partings is apparently laterally equivalent to mirabilite beds and salt disruption beds (Unit 11). Laminated calcitic clay and ostracode-rich calcitic clay with little bedding indicates deeper water conditions (Unit 111). Brown clay

Data Center ,09126599985,[email protected], For Educational Uses ALLUVIAL FAN - PERENNIAL STREAM AND SALINE LAKE 287 Bonneville basin and Great Salt Lake ,

a Bedrock

Alluvial fan

Saline mudflat aSaline Pan a Eolian dunefield

Birdfoot deta

Dry mudflat Perennial lake I ,A Perennial river

- 0%fled pellets and Qolds I Ripple laminae Unsorted parallel fiat laminae 0Oisrupted laminae

111 Ostracodal mud Iv I=] Ripple laminae Unsorled parallel llat laminae v Carbonate tree and aragonite peloidal laminae

I Bedded salt M Mazama ash CS Carson Sink ash

Fig. 3.40. A). Sedimentary subenvironments in the Bonneville Basin. Great Salt Lake is a shallow perennial lake fed by perennial streams forming birdf oot deltas (fig, 3. I9A). The saline pans are halite dominated. The dry mudf lat and saline mudf lat deposits are old lake bottom deposits, The eolian dunes arepredominantlyreworkedgypsum. B). Sequence of sores across Great Salt Lake, Utah. Units I, II, IV, and V areshallow perennial saline lake deposits (fig. 3.8A). Unit I11 was deposited in a deeper lake and contains numerous fresh water fossils. Disrupted laminae in units I and IV are interpreted as evaporite disruption f eatures and contain some cumulus crystals. Thicknesses are given with respect to elevation above sea level. (from Spencer and others, 1984).

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at the base of one core (Unit V) may be a subaerial mudflat deposit. High lake stands are indicated by shoreline deposits preserved on alluvial fans. These deposits include "Gilbert-type"deltaic deposits and wave built deposits (Gilbert, 1885, 1890). The flats west of Great Salt Lake are exposed lake bottom saline mudflats with abundant intrasediment discoidal gypsum and dry mudflats with puffy ground efflorescence. A saline pan formed on part of these flats. Other examples of perennial saline lakes associated with alluvial fans and perennial rivers are known from the same general areas as the saline pan deposits noted above. Lake Urmia in Iran (Kelts and Shahrabi, 1986) is similar to the Great Salt Lake, Utah. The Dead Sea, Israel (Fig. 3.41) is a deep bake in a narrow basin with a shallow saline pan at its southern end. Laminated aragonite and calcite are forming in the deepest part of the perennial lake and gypsum-aragonite laminae, intrasediment growth of gypsum and halite, and gypsum crusts form in the shallower parts (Neev and Emery, 1967). Halite crusts and abundant intrasediment halite and gypsum are forming in the saline pan at the south end of the basin. Other examples of narrow, deep perennial lakes in intermontane settings are generally less saline, with carbonates forming cumulus deposits and tufa crusts. Lakes Tanganyika, Kivu, and Turkana in East Africa and Lakes Baikal and Issyk Kul in the U.S.S.R., and Pyramid Lake in the U.S.A. are examples of this type. Stacked "Gilbert-type"deltas occur where the Truckee River enters Pyramid Lake, Nevada (Born, 1972). Another variation on this subenvironment association is in the Qaidam Basin in China where perennial and ephemeral saline lakes occupy only a small part of the basin floor because perennial river inflow is not sufficient to fill the basin. The shallow perennial saline lakes are bounded by older layered saline lake evaporites. A large delta has formed where the perennial Golmud River enters perennial Dabsur Lake.

Ephemeral Stream - Saline Pan (Dry Mudf lat) - Eolian Dunef ield

This association of subenvironments is common in relatively stable, low-relief, intracratonic areas where sediment accumulates in local depressions such as structural warps, deflation hollows, solution collapse pockets. The area of deposition can be very large, and is only limited by the drainage area of ephemeral streams. Eolian deposits may completely dominate the setting and deflation of sediments may be very important. The thickness of sediment in these basins is generally much less than in intermontane basins. The Lake Eyre Basin, Australia (Fig. 3.42) is a broad flat expanse dominated by longitudinal eolian dunes, ephemeral stream "channel lands", and a large saline

Data Center ,09126599985,[email protected], For Educational Uses EPHEMERAL STREAM-SALINE PAN-EOLIAN DUNEFIELD 289

Dead Sea

I:.:.] Bedrock

l"p0DI Alluvial fan

'Gilbert-type' delta

Deep perennial saline lake

Shallow perennial saline lake

Fan delta

lylri Saline pan "V

Saline mudflat

Perennial river

Fig. 3.4I. Schematic map of sedimentarysubenvironmentsin the Dead Sea Basin, Israel. The Jordan River forms a fluvial-deltaic complex in the northern part of the basin. The areas shown in white are fluvial and deltaic muds. Active faulting exposed old perennial lake deposits at higher elevations (fig. 3.70) and produced a narrow, steep alluvial fan apron with fan deltas into the present lake. The saline pan to the south is dominated by halite. pan-saline mudflat. The ephemeral streams include sheet-like braided deposits, muddy anastomosing channels, and older meandering stream deposits (Fig. 3.43). The channel systems are mostly dry, but flooding may cover hundreds of square miles with water up to 5 m deep for periods of weeks (Kotwicki, 1986). Decimeter-scale trough cross-beds (dune-scale bedforms) are common in sandy channel areas. Fluvial floodplains are desiccated and caliche and gypcrete formation is common. Eolian dunes are mostly

Data Center ,09126599985,[email protected], For Educational Uses 290 NON-MARINE EVAPORITES

4 Ephemeral saline lake

Saline soil i....

0.. Dry mudflat

Saline pan "V

Saline mudflat

Sheet delta

* Spring mounds

Dune field and gibber plain

,," Drainage divide 1

Ephemeral stream floodplain

Lake Eyre (North)

Fig. 3.42. Sedimentary subenvironments of' the Lake Eyre region, Australia. Saline soils to the west overlie Cretaceous shales. ?'he dune field and gibber plain is dotted with small dry mudflats and Pleistocene meandering river channels (fig. 3.43). Sheet delta areas are not well documented. Spring mounds along the southern margin o,f the basin rest on Cretaceous deposits. The margins of Lake Evre (inset) have a narrow band of wave-- formed shorelines against erosional bluffs of Tertiary lacustrine deposits. (modified from Hardie and others, 1978). older deposits with abundant clay pellets. Small eolian dune crests oriented parallel to the edge of the lacustrine area have a large clastic gypsum component. Interdune areas include pebbly "gibber plains", small clay pans with parna dunes, and old fluvial and lacustrine deposits with soil overprints. Where larger stream systems enter the lake, there are broad sheet deltas dominated by silt and fine sand. The sheet deltas are cut by channels (grooves) where floods drain toward the saline pans before the basin fills with water. The basin fills rapidly during floods so the surface is covered with a body of standing water as much as 4 m deep. The saline pan has a relatively thin accumulation of bedded halite surrounded by wet clay with abundant gypsum crystals. The margins of the saline pans are lined with oscillatory ripple sheets, spits, and bars,

Data Center ,09126599985,[email protected], For Educational Uses EPHEMERAL STREAM-SALINE PAN-EOLIAN DUNEFIELD 291

Flow direction of Playas ephemeral stream drainages ..:::: : ,: ..: Eolian dunes .j ;; ...I: : . P P(G c .:: ::::.. ., Lunet t es Meander belt traces <-...... : ,! 9.

Fig. 3.43. Paleochannels of Pleistocene meandering river belts and small dry mudflats (pluvus) in interdune depressions in the Strezlecki dunef ield adjacent to Lake Eyre. Large channel area in the upper part is Cooper Creek drainage. (modified from Callen and others, 1986). reflecting wave reworking of mostly eolian sand during high lake stands. There are no descriptions of the sheet deltas or the surrounding mudflats, other than that there are no intrasediment evaporites and that the sediments have oxidized colors (Dulhunty, 1982). Small cliff faces that expose Quaternary and older lacustrine deposits border the southern saline mudflats. This erosion has been attributed to wind deflation aided by groundwater sapping (Bowler, 1986). Many of the older deposits, including strandline ridges are kilometers from the present shoreline, are capped by silcrete or gypcrete surfaces. Carbonate spring mounds along the southern margin of the basin are built on deflationary pedestals of Cretaceous bedrock (Habermahl, 1988, pers. comm.). Cores and outcrops of the sedimentary fill in Lake Eyre (Callen and others, 1986; Magee and others, 1988) indicate a relatively thin accumulation of Quaternary to Holocene

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sediment disconformably overlying Miocene dolomitic clay. Modern deposits of gypsum-rich clays, halite crusts, and flaser-bedded sands dominate the upper few meters of the saline pan and shoreline deposits. Laminated clays with ostracods and fish fossils occur just above the Miocene unconformity. Gypsum crystals occur as scattered blades and discrete layers within the laminated clays, presumably formed from pore fluids during more recent saline conditions. Other examples of ephemeral streams, saline pans or dry mudflats, and eolian dunes are reported from the Kalahari Desert in South Africa, the Sahara Desert in North Africa, and the Tarim and Gobi Deserts in southern Asia (Table 3.1). In the Jaf- fura Sand Sea, Saudi Arabia (Johnson and others, 1978;Ahlbrandt and Fryberger, 1981; Fryberger and others, 1983, 1984), sandy interdune areas are salt encrusted where deflation intersects groundwater tables. The Qattara Depression in Egypt, a saline flat that covers an area over 18,000 km', is apparently a deflation pocket in an eolian dunefield (Glennie, 1970). The distribution of alluvial fan wadis, inland sabkhas, and eolian dunes and sandsheets adjacent to the Oman Mountains, Libya are provided by Glennie (1970, enclosures 1 and 3).

Perennial Stream - Perennial Lake - Eolian Dunef ield

In geomorphic settings similar to the saline pan - eolian dunefield association, perennial rivers may produce perennial lakes. The perennial lakes may be very large and shallow, and the shoreline may be very irregular with multiple embayments and ponds formed by lake water around eolian dunes. Lake Chad in north-central Africa (Fig. 3.44) is a good example. Lake Chad is fed by two major perennial rivers draining mountains to the south. The rivers produce extensive meander and flood plain deposits terminating as birdfoot deltas. An ancient eolian dunefield (the Kanem erg) to the northeast is partly flooded by the lake. Lake Chad is mostly freshwater, precipitating calcite and aragonite. It is more saline to the northeast, but still does not precipitate saline minerals. Between the dunes on the northeast side of the lake, there are a number of small saline pans and saline mudflats fed by lake water seepage or by flooding. Na-carbonate minerals and Na-silicates form as crusts and intrasediment crystals in these pans. Small saline pans also occur on islands in the lake, formed by dunes surrounded by water. Some of these have Na-sulfate salts precipitating in them.

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Lake Chad basin

...... :. Perennial river floodplains

inactive or ephemeral braided and meandering river .. floodplains and deflation plains

Eolian dunefields

Old lake bottom Elsaline mudflat and dry mudflat

Perennial saline lake

Fig. 3.44. Schematic map o,f Lake Chad basin, North Africa. Perennial rivers from the south are braided near the mountains and meandering near the lake where they form birdf oot deltas, Dunef ields in the southern two-thirds of the basin are mostly inactive. Dry river channels include deposits equivalent to the Pleistocene pluvial lake stage. (modified from Cooke and Warren, 1973 and Roche, 1977).

Springs - Saline Pan (Perennial Lake)

In the central continent of North America and Asia, there are a number of small saline pans and saline lakes forming in a variety of depressions primarily fed by groundwater seeps. Many of the saline lakes occur within old channel systems or within karst depressions. The lakes are almost all Na-sulfate dominated and their deposits are primarily chemical precipitates with mixtures of windblown loess or coarser sediments reworked from older deposits by wave action during flood stages. The surrounding areas are generally heavily vegetated and little surface sediment transport occurs. The source of solutes appears to be dissolution of older evaporites or possibly oxidation of sulfides. Some of these are perennial lakes with perennial stream inflow, such as Freefight Lake in Saskatchewan. The groundwater-fed lakes in Australia with alunite

Data Center ,09126599985,[email protected], For Educational Uses 294 NON-MARINE EVAPORITES and jarosite (Lock, 1988a; Bird and others, 1989) are also in this category. The Chott el Djerid in North Africa (Meckelin, 1977) is a spring-fed, gypsum and halite dominated saline pan in a deflation hollow surrounded by eolian dunes. Ceylon Lake (Fig. 3.45) is a saline pan with fringing saline mudflat. Shoreline features include wave reworked mirabilite crystals and some sandy spits reworked from the surrounding glacial deposits. The central pan is mostly coarsely crystalline mirabilite with thenardite, bloedite and some mud partings. Intrasediment mirabilite occurs in muddy and sandy sediment near the shore. Spring upwelling within the pan forms solution chimneys and mounds of mirabilite or bloedite up to 2 m high. These mounds do not survive subsequent surface flooding. Linear dikes of mud and salt frequently extrude at the surface of the saline pan during desiccation stages. A core from the central pan of Ceylon Lake (Fig. 3.45B; Last, 1989) includes about 3 m of sulfate salts with scattered black mud partings that have abundant mirabilite and gypsum crystals. The amount of mirabilite in the salt layer increases from 55% at the top to 80% near the base. The lower part of core contains black clay with abundant mirabilite and gypsum crystals; calcareous clay with ostracods, gastropods, and charophytes; and irregularly spaced hard carbonate crusts and gypsum laminae. These deposits sharply overlie sandy to pebbly fluvial deposits and glacial till.

Ancient Examples

Comparison between modern subenvironments and ancient deposits is difficult because most descriptions of ancient deposits do not give information on the lateral variability of evaporites or the surrounding sediments. A few examples of ancient deposits illustrating the vertical and lateral distributions of facies are provided below. Wilkins Peak Member, (Eocene) Green River Formation, U.S.A.. The Wilkins Peak Member is a dolomitic lacustrine deposit in the Green River Basin in southwest Wyoming, U.S.A. The Green River Basin is a broad fault block basin of the Laramide thrust belt system (Fig. 3.46A,B). There are three members in the Green River Basin, the oil shale-dominated Tipton Shale and Laney Members and the evaporite-rich Wilkins Peak Member. The Tipton Shale and Laney Members are aerially more extensive than the Wilkins Peak Member, and were deposited for the most part in large perennial freshwater lakes. The 300 m thick Wilkins Peak Member has over 40 organic-rich laminite beds (oil shale) averaging 1-2 m thick intercalated with dolomitic mudstone and sandstone with abundant desiccation features (Fig. 3.47). This intercalation probably reflects repeated rise and fall of a perennial lake. Unlike the oil shale of the Tipton Shale and Laney Members, the oil shale of the Wilkins Peak

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> SEDIMENTARY FEATURES 9 U LAMIN_*TLON 00 co

HARD

+IUD INCLUSlONS

MODERN FACIES COLLUVIUM 'l..""" MUD FLATISAND FLAT ---1 INTRASEDIMENT~Y '1..MIRABILITE CRYSTALS SALT PAN .A GYPSUM LAMINAE MIRABOLITE SHOAL :ALCAREOUS HARDGROUND a CRUSTS FIRM SPRING ORIFICE 4EUNDANT CALCAREOUS : BEACH SHELLS a ORGANIC DEBRiS

OTHER FEATURES

0 SALT KARST CHIMNEYS TIFF - MUD DYKES * DlAPlRS

Fig. 3.45. Depo.7itional environments of Ceylon Lake, Saskatchewan. A). Schematic map view o,f' Ceylon Lake, Saskatchewun; the shupe of' the lake reflects a ,former stream channel. Mirabolite shoals are wave-reworked mirabilite forming oolitic coatings. Solution chimneys. d.ykes, and diapirs in the saline pan (saltpan) result f rom spring seeps. (fromLast, 1989). B). Coresection through thesaline pan lit Ceylon Lake, Saskatchewan. Saline pan mirnbilite overlies aperennial .saline lake deposit with intrasediment mirabilite crystals and gypsum cumulus laminae. Lower clay is a fresh water glacial lake overlying a till deposit. (from Last, 1989).

Member is mostly devoid of fossils, except for insects and, less commonly, ostracodes. Conglomerates and sandstones along the southern margin of the basin resemble stream-dominated alluvial fan deposits grading basinward into quartzose intraclastic sandstone sheets with sedimentary features of sandflats. The sandstone sheets wedge

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I Wyoming

1 km 0 100 200

Fig. 3.46. Map views of'the Green River Formation in Wyoming Utah, and Colorado. A). Distribution of depositional basins of the Eocene Green River Formation (stippled).Each basin contains a separate stratigraphic sequence. Bedded evaporites occur in the subsurface of the Green River and Piceance Basins (lined). (from Eugster and Kelts. 1983). B). Green River Basin showing the inferred area of the hydrographic basin and extent of the Laney Member lacustrine beds (stippled) and the Wilkins Peak Member (shaded). Wilkins Peak deposits east of the Rock Springs "lsland" are predominately f luvial. (from Bradley and Eugster, 1969).

out basinward and intertongue with dolomitic siltstone, mudstone, and oil shale. The extensive sandflat deposits may be due to the sandy nature of the alluvial fans and the broadness of the basin floor. Aggrading dry mudflat fabrics are also abundant probably because of the broad flat basin floor which allowed rapid expansion of very shallow ponded water in response to surface flooding. Sheets and mounds of stromatolitic travertine and tufa, interpreted as spring deposits, are abundant within the sandstone sheets and in mudcracked dolomitic mudstone. The arkosic sandstone deposits are believed to represent episodic influxes from the southeast that resulted in rapid expansion of the central lake. Repeated 4-5 m thick alternations of oil shale and bedded evaporites in the basin center and of oil shale and mudcracked dolomitic mudstone further towards the basin margin probably reflect climatic cycles. Arid stages change a large perennial saline lake into a shallow saline lake and finally a saline pan. Intrasediment evaporites associated with bedded evaporites do not show variations in texture coincident with changes in sedimentary structures. There is a decrease in the abundance of soluble salts away from the bedded evaporites and a

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A Tollgate 6 119 8 5 4 2 rock 16

Unto Mlns Alluv~ol Sand Dry Perennial Saline B I Ion 1101 mudllal mrdfl0l

I/ 9 8 52 I

Fig. 3.47. A). Lateral distribution of sedimentary subf acies in the Wilkins Peak Member along a transect from the Wyoming-Utah border south of the Rock Springs "Island" northeast into the evaporite zone. Numbered lines are measured sections. Vertical and lateral facies changes reflect rise and fall of lake level. The sequences shown here are repeated numerous times in the 300 m thick section and are intercalated with 10-20 m thick siliciclastic f luvio-deltaic sequences. Sedimentary units are 1) alluvial fan conglomerates and sandstones, 2) sandf lat sandstones, 3) dry mudf lat mudstone, 4) intermittent shallow perennial lake, 5)perennial saline lake (vertical lines show ashf all tuff beds), 6)saline mudf lat with vadose evaporite textures (a),phreatic evaporite textures (b), and possible perennial lake intrasediment evaporites (c), 7)bedded trona and halite with saline pan textures and possible perennial lake cumulus beds, and 8) ash-f all tuffs. Note the vertical exaggeration. (modified from Smoot, 1983). B). Schematic representation of lateral facies relationships based on the above cross-section. (from Smoot, 1978).

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general change to more layer selective distribution and anhedral forms. The abundance of intrasediment evaporites in the central facies probably does not indicate uniform saline conditions, but more likely diagenetic precipitation of saline minerals from brines that sank into fresher water deposits during saline stages. Intrasediment evaporites in the outer portions of the saline facies were probably not influenced by the same brines and thus maintained their original distribution. This interpretation is uncertain because the most abundant mineral, shortite, is probably itself a diagenetic replacement of another mineral, such as pirsonnite or gaylussite, and because calcite pseudomorphs are evaporite minerals in outcrops. Some of these crystal shapes may have formerly been sulphate or carbonate-sulphate minerals. No features indicating efflorescent salt crusts were recognized in these deposits. The evaporite deposits of the Wilkins Peak Member are very different from probable time equivalent evaporite deposits in the Piceance Basin (Fig. 3.48; Dyni, 1974; Surdam and Stanley, 1979; Cole and Picard, 1981; Dyni and others, 1985). These evaporite deposits appear to have accumulated exclusively in a perennial saline lake subenvironment as cumulus beds of nahcolite and halite and intrasediment nahcolite (Dyni 1974). Tajo Basin (Miocene), Spain. The Miocene Tajo Basin is a fault bounded depression filled with over 1500 m of siliciclastic sediment and evaporites (Fig. 3.49; Orti Cabo and others, 1979; Calvo and others, 1989). Conglomeratic arkose reflecting shallow flash-flooding streams is interpreted as alluvial fan deposits. These grade laterally into green shale, and then evaporite deposits. Clay minerals in mudstones show a lateral transition from dioctahedral smectite to trioctahedral smectite to illite from the basin margin to the center (Calvo and others, 1986; Jones and others, 1986), apparently reflecting a saline lake system. Sepiolite in the arkosic facies is interpreted as brackish spring pond deposits at alluvial fan toes. Bedded halite deposits are interpreted as saline pan deposits. Glauberite beds include nodular beds and scattered euhedral crystals both within the halite and adjacent to it. These may also represent saline pan deposits, perhaps former gypsum pan deposits altered by halite-saturated brines (Orti Caband others, 1979). Nodular anhydrite in the marginal facies may also have been originally deposited as gypsum which dehydrated either in the depositional environment or after burial. The vertical transition of halite to thenardite, gypsum, and then carbonate (Fig. 3.49A) suggests a change in water chemistry and a generally less saline environment. The gypsum appears to be at least in part a saline pan or shallow saline lake deposit, as indicated by vertically oriented crystal sprays. The thenardite may be a diagenetic mineral formed from concentrated brines derived from the gypsum-producing environment and accumulating at the confining layer formed by the underlying

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sw NI Horizontal scale 10 UINTA FORMATION

Fig. 3.48. Schematic cross-section through the Eocene Green River Formation in the Piceunce Basin. Colorado. The Douglas Creek Member and oil shale facies of the Parachute Creek Member are perennial lake deposits. The saline mineral zone with abundant nahcolite and some halite isdominated by perennial saline lake cumulus crystal beds and intrasedimeiit crystals with some possible crusts. The Garden Gulch Member and the sandstone facies of the Parachute Creek Member are most1.v birdfoot delta deposit,r and some wave-f ormed shoreline deposits. The stromatoliticfacies is shoreline tufa arid the marlstone is shallow perennial lake or saline mudf lat deposits. (fromCole and Picard, 1981). halite. Newark Basin (Triassic-Jurassic), U.S.A..The Newark basin is a half graben structure extending across parts of New York, New Jersey, and Pennsylvania. It is part of a string of fault basins, whose strata are collectively called the Newark Supergroup formed by the extensional processes that ultimately led to the opening of the Atlantic Ocean. The fluvial and lacustrine strata and interbedded basalt flows are several thousand meters thick, and represent about 25 million years of continuous deposition (Fig. 3.50). The sedimentary sequence can be roughly divided into four parts: 1) an initial fluvial stage consisting of braided and meandering river deposits that extend across the basin floor, 2) the first lacustrine stage consisting of fossiliferous organic-rich

Data Center ,09126599985,[email protected], For Educational Uses 300 NON-MARINE EVAPORITES

A Tajo Basin

meters Lacustrine carbonate Halite facies

Gypsum facies

Thenardite facies Anhydrite facies 1",1<.>

B - W sw

Basement Arkose Evaporites M Calcareous shale Carbonate Shale with evaporites

Fig. 3.49. Cross-sections of the Miocene Tajo Basin, Spain. A). Halite facies exhibits saline pan textures with intrasediment glauberite and halite crystals in mudstone partings. Glauberite f orms crystalline beds and intrasedirnentgrowths (Glauberitef acies) that may represent altered gypsum pan and saline mudf lat deposits. Anhydrite occurs as intrasediment nodules and may indicate more vadose influence. Thenardite displays saline pan textures and the overlying gypsum appears to be perennial lake crusts and cumulus crystal layers. The total sequence suggests a gradual decrease in salinity and more perennial lake conditions. (redrafted from Orti Cab0 and others, 1979). B). Conglomeratic arkoses grade laterally into mudstones (shale) with intrasediment gypsum crystals and zones of smectite clays and sepiolite. The central evaporite zone is mostly gypsum, but includes halite and glauberite to the south. The evaporitic sequence is underlain and gradationally overlain by carbonate lacustrine deposits representing less saline conditions.

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Fig. 3.50. Lower Mesozoic deposits in the Newark basin of the Newark Supergroup, New Y ork, NewJersey, and Pennsylvania. Schematicmapview of depositionalsubenvironmentsy in the Passaic Formation and the adjacent Gettysburg basin; S - Stockton Formation, L - Lockatong Formation, P - Passaic Formation. Outline of preserved basin boundaries are shown by thin lines. I) Alluvial fans, 2) Axial braided stream complex, 3) Meandering

stream deposits and stacked "Gilbert-type" deltas and sheet deltas, 4) Alternations of perennial lake deposits and dry mudflat deposits, 5) Ephemeral stream or low-angle alluvial f an deposits. Sheet delta deposits are equivalent to dry mudf lat deposits. Euhedral intrasediment evaporite crystals (gypsum or glauberite) are randomly scattered across f acies 3 and 4, suggesting saline soil development.

laminites alternating with shallow lake to subaerial mudstone, 3) a middle lacustrine stage dominated by shallow lake deposits and subaerial mudstone, and 4) a volcanic stage with thick basalt flows alternating with lacustrine deposits that include fossiliferous black shale. Olsen and Schlische (1988) and Olsen and others (1989) interpret these transitions as a function of the basin gradually widening so that sediments are dispersed over increasingly larger areas. The lacustrine sediments form 3-7 m thick cycles of perennial lake deposits to subaerial deposits, interpreted as rise and fall of lakes in response to climate. The subaerial portion of these cycles in the upper Lockatong

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Formation and the lower Passaic Formation is dominated by slowly aggrading playa mudflat deposits. Root structures may be abundant in the overlying unbedded mudstone below the next lacustrine shale, and these become more abundant near fluvial deposits and also towards the top of the Passaic Formation. Lacustrine deposits grade into alluvial fan conglomerate near the border fault and into braided river and meandering stream deposits towards the narrow ends of the basin. Stacked "Gilbert-- type" and birdfoot delta deposits occur where perennial lake laminites intertongue with the stream deposits. Sheet delta deposits occur where alluvial fan sandstone is transitional to shallow lake and dry mudflat deposits. Wave sorted conglomerates overlie perennial lake laminites where they intertongue with alluvial fan conglomerate. Evaporite minerals are represented by cement-filled molds, probably after gypsum or glauberite. Most of the evaporites occur as euhedral crystals randomly distributed in all of the subaerial facies including fluvial deposits. In the middle part of the Passaic Formation, they commonly occur with root structures and carbonate nodules, suggesting that they are soil evaporites. This is consistent with the common occurrence of dry mudflat deposits suggesting a deep groundwater table. In the lower Passaic Formation, a few occurrences of crystals in the regressive portion of lake deposits exhibit characteristics of bottom growth and layer selectivity, suggesting they precipitated from a saline lake. The organic-rich laminites underlying the lithologies containing the saline mineral crystal molds are sulfide-rich and devoid of fossils. In the Jurassic Boonton Formation, layer selective crystals are common in shallow lake deposits and structures resembling sand-patch fabric were observed in some horizons. This may indicate a shallow saline groundwater table at that time. Fundy Basin (Triassic-Jurassic), Canada. The Fundy basin, also part of the Newark Supergroup crops out in Nova Scotia and extends beneath the Bay of Fundy. Exposed sections are less than lo00 m thick, but equivalent strata in the subsurface are several thousand meters thick. Although stratigraphic relationships are still in doubt, there appears to be a basal fluvial sequence overlain by a cyclic lacustrine sequence like in the Newark basin. Lacustrine cycles in the Fundy basin consist mostly of alternations of clay-rich shale, sometimes with pinch-and-swell lamination, and sandy mudstone with sand-patch fabric (called adhesion ripples by Hubert and Hyde, 1982, by comparison to Glennie, 1970) and abundant anhedral gypsum ciystals. These are consistent with alternation of shallow lake deposits and efflorescent salt encrusted saline mudflat deposits. Massive breccias with abundant sandstone-filled dikes are interpreted as synsedimentary solution collapse of salt beds resulting from freshwater lake expansion. Sheet-like sandstones with climbing ripple cross-lamination that intertongue with lacustrine cycles are probably sheet delta deposits (equivalent to the sandflat of Hubert

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and Hyde, 1982). Eolian dune deposits occur in fluvial sandstones immediately beneath the cyclic lacustrine deposits and at several other stratigraphic levels. The oldest fluvial sandstones are interbedded with eolian sandstones and contain numerous carbonate nodules and pavements, suggesting that they were deposited by ephemeral streams, similar to the modern Australian channel land (Veevers and Rundle, 1979). These are overlain by a thick sequence of fluvial sandstone and conglomerate that reflects perennial braid stream conditions. European Rolliegendesand Zechstein (Permian).The Rotliegendes, as interpreted by Glennie (1972), consists of a large eolian sand sea in which lacustrine carbonates and evaporites accumulated within a local structurally controlled basin. Cross-bedded, pebbly sandstone within the eolian dune deposits suggest shallow ephemeral streams. The evaporite sequence suggests saline pan halite and saline mudflat gypsum deposits. Core photographs show sand-patch structures and deformed layering indicating efflorescent salt crusts. The carbonate layers may represent perennial lake stages similar to the present Lake Chad. Evaporites of the Upper Zechstein in England have also been interpreted as lacustrine deposits (Smith, 1970,1971;Smith and Crosby, 1979). These deposits include layered potash salts associated with halite that has saline pan characteristics. Other Examples. The Miocene Ridge Basin Group in California (Link and Osborne, 1982) contains alternations of deep-water to shallow-water perennial lake and shoreline deposits including wave and deltaic varieties. Smith (1982) reports intrasediment gypsum crystal aggregates from organic-rich perennial lake shale in the shallow water phase. The Tertiary Mormoiron basin in France (Truc, 1978, 1980) contains a thick sequence of gypsum with morphologies suggesting deposition as crusts in a perennial saline lake. The Middle Proterozoic Belt Supergroup in northwestern U.S.A. (Winston, 1977, 1984; Winston and others, 1984) contains cyclic sequences of sedimentary structures reminiscent of the Eocene Wilkins Peak Member deposits described above. The Belt Supergroup cyclic deposits are laterally equivalent to fluvial deposits. The lateral extent of the deposits and their stratigraphic association with deposits more characteristic of a marine setting suggests that these are marine deposits or that the basin is an abandoned marine embayment like the Caspian Sea. The Precambrian McArthur Basin in Australia (Muir, 1983; Jackson and others, 1986, 1987;

Donnelly and Jackson, ' 1988) also contains deposits suggestive of both marine and non-marine deposition. It may have contained small lakes in a coastal setting. The Proterozoic Duruchaus Formation in the Damara Belt of Namibia (Behr and others, 1983; Porada and Behr, 1988) is a metamorphic sequence interpreted as a cyclic Na-carbonate and borate playa deposit. The saline mineral pseudomorphs are mostly

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albite or microcline and the alleged borates are now tourmaline.

RECOGNITION OF ANCIENT NON-MARINE DEPOSITS

As mentioned in the introduction, the distinction between ancient marine and non-marine evaporites is not easy. It has frequently been stated that lacustrine deposits older than Tertiary are rare (Feth, 1964; Picard and High, 1972a, 1981) and their paucity has been attributed to their poor preservation potential. It is probably no coincidence, however, that identification of non-marine organisms or determination of the continental setting becomes progressively less certain in strata older than Tertiary. Lately, the number of reports of lacustrine deposits has increased dramatically and most of these are associated with features interpreted as evaporites. Hardie (1984) reviewed the problems of distinguishing between marine and non-marine evaporites. He suggested three categories of criteria for distinguishing between the different types of evaporites: sedimentological,mineralogical, and chemical. Each of these are reviewed below, but the sedimentaq characteristics of non-marine evaporites are emphasized. Mineralogical criteria are based on the types of evaporites and their abundances that reflect brine chemistries consistent with seawater or incompatible with seawater. The "ideal" seawater evaporite sequence is not easily applied because: 1) the classical "evaporating pan" sequence of salts from seawater (i.e. Braitsch, 1971) is wrong (Harvie and others, 1980; Hardie, 1984); 2) the sequence and abundance of saline minerals may change because seawater is added continuously or episodically to the starting brine, because crystals remain in contact with the brine allowing further reactions (syndepositional reaction minerals), because detrital minerals (i.e. biogenic carbonate or clay minerals) may react with the brine or pore water (Spencer, 1982), and because concentrated brine may move out of the system before complete desiccation; 3) the brine may be chemically altered by mixing with other waters, such as surface runoff or spring inflow (mixed seawater/non-marine water), especially since evaporites typically precipitate from seawater isolated from the oceanic reservoir; and 4) seawater may not have always had the same composition. Any combination of minerals produced by the evaporation of seawater can be duplicated in a non-marine setting. The most common evaporite minerals formed from seawater, gypsum and halite, are also the most common non-marine evaporites. There are some mineral assemblages, however, that cannot be produced from seawater without major modification by non-marine inflow: 1) Nacar- bonate minerals, such as trona, nahcolite, and shortite, 2) Na-silicate minerals, such as magadiite and kenyaite, and 3) abundant Na- or Ca-borate minerals.

Data Center ,09126599985,[email protected], For Educational Uses RECOGNITION OF ANCIENT NON-MARINE DEPOSITS 305

The problem of distinguishing between marine and non-marine evaporites is further complicated by diagenesis which may alter the mineral assemblage and, particularly in older rocks, may leave only cement-filled pseudomorphs as the record of the minerals. The obvious pitfalls in determining mineralogy from pseudomorphs are that many different saline minerals are in the same crystal classes and exhibit similar crystal habits (i.e., trona and gypsum are both monoclinic and form radial sprays), anhedral crystal forms are indistinguishable (i.e., anhydrite or borax nodules), and early diagenetic versus late diagenetic replacement is difficult to distinguish (i-e., chert layers replacing magadiite or opaline silica). The mineralogy of evaporite pseudomorphs is rarely determined by measurement of crystal angles. The identification typically appears to be based on a superficial comparison of crystal habit. Geochemical criteria for distinguishing between marine and non-marine evaporites and for determining their diagenetic histories include trace element, isotope, and fluid inclusion studies. The two most widely cited measurements for identifying marine evaporites are the CI/Br ratio (Holser, 1970,1978;Holser and others, 1972) and 34 S/32Sisotopic ratio. These two ratios depend upon the constancy (or near constancy) of seawater composition and assume non-marine evaporites will have a freshwater signature. Hardie (1984) demonstrated that the range of CI/Br ratios in natural Occurrences of halite from modern non-marine basins and indisputable ancient non-marine basins completely overlap those predicted for seawater halite (Fig. 3.5 1). A similar overlap in range may also exist for marine and non-marine 34Sf% ratios (see data of Tuttle, 1988). The significance of values measured in ancient deposits may also be suspect if care was not taken to separate the primary crystals from diagenetic overgrowths or if fluid inclusions were present. Recent studies on boron isotopes suggest that the 'lB/'OB ratio is higher for borate minerals precipitated from seawater than for those precipitated from non-marine water (Swihart and others, 1986; Oi and others, 1989) although this is based on a small number of samples. Fluid inclusions have potential value as indicators of brine chemistry and temperature at the time of crystal formation (Hardie and others, 1985). Sedimentological criteria for distinguishing between marine and non-marine evaporites include the geographic setting, associated fossils, and sedimentary structures and packaging. If the deposit is known to be separated from seawater by a considerable distance or by an area of high relief, if the sedimentary rocks surrounding the deposits all represent non-marine depositional environments, or if the intercalated sedimentary rocks contain non-marine fossils, a non-marine origin for the evaporites can be ascertained with certainty. None of the depositional subenvironments described in this paper, however, is absolutely diagnostic of a non-marine evaporite setting. Alluvial fans,

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BROMINE IN HALITE 50 75 100 ppm Br

Saline Valley, CA Holocene

(0-3 m) Searles Lake, CA Holocene-Pleistocene (100-200 m) zrPPrn(210-270m) Green River Formation Wyoming Eocene Colorado Predicted for first halite from seawater Basal halite of Zechstein 2 l==!l 50 75 100

Fig. 3.51. Comparison of bromine contrtitpredictecif’or halite precipitatedf’romseawater to observed ranges of bromine in halite in non-marine deposits (black). The non-marine bromine values completely overlap those of seawater. The valuesfor the ”classic”marine rvaporite sequence of the Zechstein 2 in Germany are shown for comparison. (redrafted from Hardie, 1984).

streams, eolian dunes, and springs may all form deposits adjacent to an evaporitic marine embayment or tidal flat (for instance, Baja California or the Persian Gulf). Lacustrine subenvironments are also similar to those of restricted marine lagoons or sabkha flats. The only marine process that would not be expected in lakes is large tides that could produce complex bimodal cross-bedding or tidal bundles (as in Boersma, 1969). The alternation of perennial lake and subaerial deposits, that reflects the instability of inflow and evaporation in a closed basin, may be a useful indicator, but similar alternations of subaqueous and subaerial deposits may occur in a marine lagoon with restricted seawater input or in tidal flats (see for example Lowenstein, 1988, for an explanation of cyclic evaporites in the Permian Salado Formation, New Mexico, and James, 1977, for shoaling upward carbonate tidal flat cycles). The distribution of evaporite fabrics appears to be more complicated in non-marine basins than in marine tidal flats. More detailed studies of the vertical and lateral variations of evaporite minerals are needed in both the non-marine and marine settings, however, to determine if the variability is unique to their setting.

Data Center ,09126599985,[email protected], For Educational Uses ECONOMIC ASPECTS 307

Non-marine evaporites may accumulate on top of marine sediments. This occurs when sea level drops leaving saline lakes and broad saline flats behind. The remaining marine waters can be modified by continental waters, as in the Caspian Sea, or marine salts may be recycled by floods into saline pans, as in the Grand Kavir, Iran. It would not be surprising to find that many of the thick ancient marine evaporites have thin lenses of recycled non-marine evaporites representing drops in sea level. Saline soils form on marine deposits with the source of solutes being seawater. The soil is not a marine deposit even though it is formed on marine sediments and owes it chemistry to seawater. The distinction between marine and non-marine evaporites is not trivial and many ancient evaporites that are literally assumed to be marine may be, at least in part, non-marine in origin (for instance, the thick Paleozoic evaporite sequence of the Elk Point and Paradox Basins, western North America). Furthermore, some of the deposits interpreted as non-marine may actually be marine (for instance, the evaporites of the Precambrian McArthur Basin, Australia). Part of the solution to unraveling the origin of these deposits is to stop interpreting hundreds to thousands of meters of section as one depositional environment, and try to establish the relationship of each part of the deposit to the surrounding facies. Many of the ambiguous deposits will probably have both marine with non-marine influences or alternations between them.

ECONOMIC ASPECTS

Many modern and ancient non-marine evaporites have been mined or have been considered for mining, including halite and gypsum (Ver Plank, 1958; Reeves, 1978), borate minerals (Muessig, 1959; Christ and Garrels, 1959; Barker and Lefond, 1985; Libbey, 1985; Smith, 1985; Qian and Xuan, 1985), sodium and potassium sulfate minerals (Condra, 1918; Hicks, 1921; Hirschkind, 1931; Cummings, 1940; Smith and Crosby, 1979; Slezak and Last, 1985; Lock, 1988), potassium chloride minerals (Holwerda and Hutchinson, 1968; Desio, 1970; Smith, 1970; Smith and Crosby, 1979; Momin Gu Shaqi and others, 1986; Lowenstein and others, 1989), Nitrate minerals (Ericksen, 1981; Garrett, 1985), and sodium carbonate minerals (Culbertson, 1971; Deardorff and Mannion, 1971; Dyni, 1974; Zhang Yoaxun, 1985). Brines have also been pumped for economic elements such as lithium (Kunasz, 1974; Cannon and others, 1975; Bohannon, 1976; Ericksen and others, 1976; Vine, 1976). Saline perennial lakes have been recognized as sinks for base metals, such as copper, zinc, and lead (Degens and others, 1971, 1973; Degens and Kulbicki, 1973; Miiller and Forstner, 1973; Robbins, 1983). Many ancient base-metal deposits, including

Data Center ,09126599985,[email protected], For Educational Uses 308 NON-MARINE EVAPORITES

some of the larger active mines, are interpreted as saline lake deposits (Muir and Ridgeway, 1975; Steven and Eaton, 1975; Rowlands and others, 1980; Muir, 1981, 1983; Neudart and Russel, 1982; Behr and others, 1983; Eugster, 1985; Jackson and others, 1986, 1987; Parnell, 1986b). Some of the sandstone-hosted base-metal deposits, including fluvial sandstones, have been attributed to saline lake brines moving metals (Davidson, 1965; Muir and Ridgeway, 1975; Rose, 1976; Holmes and others, 1983; Eugster, 1985; Flint, 1986; Smoot and Robinson, 1988). Brine migration from saline lakes has also been cited as the cause of some sandstone-hosted uranium deposits (Turner-Peterson, 1980, 1988). Mixing of fresh water with gypsum-saturated brines in association with caliche has been suggested for some uranium deposits (Carlisle and others, 1978; Arakel and McConchie, 1982; Arakel, 1986). The high organic productivity of saline lakes has been recognized by many (for instance, Beadle, 1974; Larson, 1980; Bauld, 1981a,b; Borowitzka, 1981; Burne and Ferguson, 1983; Powell, 1986; Warren, 1986, 1989; Evans and Kirkland, 1988). Petroliferous perennial lake deposits have been long noted in association with evaporite-bearing formations such as the Eocene Green River Formation, western U.S.A. (Bradley, 1964,1970)and the Devonian Albert Shale, New Brunswick (Griener, 1974; Howie, 1979). The realization that many of China’s large hydrocarbon reserves are lacustrine-sourced (for instance, Chiyi Chang, 1981; Tian and others, 1983; Chen Quanmao and Dickinson, 1986; Yinan Qiu and others, 1987) has led to considerable interest in ancient lacustrine deposits associated with evaporites (Ziegler, 1983; Powell, 1986; Warren, 1986, 1989; Smith and Robison, 1988). Many of the petroleum occurrences along the Atlantic rift margin may have source rocks that are saline lake deposits (Belmonte and others, 1965; Wardlaw, 1972; Brink, 1974; Campos and others, 1974; Szatmari and others, 1979; Ponte and others, 1980; Brice and Pardo, 1981). Reservoir rocks in the Chinese lacustrine petroleum occurrences are reportedly alluvial fan conglomerates (Chiyi Chang, 1981). Other reservoir rocks may include wave-reworked shoreline deposits or birdfoot delta distributary channels (Fouch and others, 1976) and fluvial channel deposits (Yinan Qiu and others, 1987). The emphasis on lacustrine exploration should greatly increase our knowledge of these deposits in the next several years.

ACKNOWLEDGMENTS

We would like to thank the following people for contributing photographs and unpublished data for this paper: Larry Benson, Enrique Casas, Barbara Cdstens-Seidell, George E. Ericksen, Hans P. Eugster, Lawrence A. Hardie, Michael J. Jackson, Blair F.

Data Center ,09126599985,[email protected], For Educational Uses REFERENCES

Jones, William M. Last, Peter LeTourneau, F. Orti Cab, Marith C. Reheis, Brian R. Rust, Denys B. Smith, Peter N. Southgate, Ron Spencer, James T. Teller, Michelle L. Tuttle, Don Winston, and Daniel H. Yaalon. Blair Jones, Charles Barker, and Ed Hasser reviewed the manuscript and greatly improved its presentation. We gratefully acknowledge their assistance in preparing this manuscript, but we are solely responsible for any inaccuracies of description, presentation, or interpretation of their data. Linda Gundersen drafted Figures 3.12,3.14A, 3.39,3.40A, 3.41,3.42,3.43,3.44A, 3.47B, 3.49, and 3.50, and assisted in the photography and organization of the manuscript. J. Smoot would particularly like to thank her for her continued support through the ordeal of preparing this paper. We especially thank the editor of this volume for her phenomenal patience.

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Data Center ,09126599985,[email protected], For Educational Uses Chapter 4

PETROLEUM AND EVAPORITES

Robert J. Hite and Donald E. Anders

INTRODUCTION

Asphalt Occurrences in and around the Dead Sea have been known and used since Biblical times (Nissenbaum, 1977) and probably represent man's first documented observation of the affinity between salts and hydrocarbons. In a more scientific vein, Ochsenius (1888) first called the attention of the geologic profession to the association of evaporites and petroleum. Woolnough (1937) in his classical paper on barred basins went so far as to discuss the relationship in terms of petroleum source rocks. Despite the work of these early visionaries, the evaporite environment was, for a long time, generally considered as a sub-facies of red beds and devoid of organic matter. The late fifties and sixties brought about a resurgence of interest in the association. Drawing on the earlier work of Borchert (1940) and Lotze (1957), Borchert and Muir (1964) devoted a chapter of their book to "the relationship between evaporites and oil". Indeed, by this time, it had become obvious that so much of the world's petroleum was at least spatially associated with evaporites (Moody, 1959; Weeks, 1961) that a genetic relationship was sought (Hedberg, 1964). In an attempt to more closely define the relationship, AAPG convened a symposium "Evaporites and Petroleum" in 1968. This resulted in an entire issue of the 1969 Bulletin being devoted to papers presented at the Symposium. Thus, it appeared that at last the profession had moved beyond admission of a genetic relationship and had begun to search for the cause. Unfortunately, and in the words of special editor of the issue (Buzzalini, 1969), "----some of the papers included here say very little about petroleum generation". Editor Buzzalini was more than generous, for a Scan of the 13 papers involved reveals that perhaps only one came close to addressing the subject of the Symposium. Only in this decade has the profession begun to seriously study the evaporite-petroleum association. The reasonsfor this turning point are probably two-fold. First, the paper "Source-rock potential of evaporitic environment," by Kirkland and Evans (1981), which was also the subject of an AAPG distinguished lecture tour by Evans, sparked considerable interest. These authors began the necessary evaluation of high-biomass production within the hypersaline environment. A second important reason for the resurgence is due to the fact that now the biology disciplines are providing much of the badly needed data on the quantity and quality of biomass from salterns and Holocene evaporite environments,

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particularly the algal mats of coastal sabkhas and salinas. Our history regarding the early benchmark studies of the evaporite - petroleum association would not be complete without noting a significant, but perhaps not widely circulated, paper by Szatmari (1980) which provided some of the first important observations and ideas on the evaporite source rock concept in terms of the actual rocks. For this chapter, we were principally concerned with comparing the quantity and quality of organic matter and hydrocarbons produced in a variety of modern evaporitic environments with hydrocarbons and kerogen in rocks deposited under similar conditions. In particular, we attempted to concentrate our study on samples of organic matter and hydrocarbons which were truly indigenous to a range of evaporite rocks. From this approach, our observations may help answer certain questions such as which evaporite lithology can truly be considered a petroleum source rock, are lacustrine evaporites more favorable source rocks than marine, and does petroleum derived from evaporitic source rocks truly have a unique geochemicalfingerprint. In addition, we have explored other parameters such as paleoproductivity, mobile source materials, and reservoir development, all of which provide some insight to the evaporite petroleum relationship.

DEFINITION OF TERMS

There are two groups of terms which are used frequently in this chapter that unfortunately have somewhat imprecise usage in the literature on evaporites. These terms refer to the depth of an evaporite brine body and to the salinities of brines. For a number of years, a controversy has raged over deep vs. shallow origin of certain evaporite deposits. Some of the combatants are remarkably confident and well prepared to argue their particular stand on this issue but can be brought to earth again by the question; what is deep? In fact, it is quite likely that in many cases no argument exists were the bathymetric conditions defined. For purpose of this chapter, we define deep as anything below wave base and the photic zone. It is well known that heavy concentrated brines drastically reduce effective wave base; however, the actual depths under these conditions are poorly known (Sonnenfeld, 1984). Brine transparency is also effected by salinity but not in a consistent manner because of variable concentrations of the standing crop of biomass. For these reasons, there will be a range in values for what is deep using the effective wave base - photic zone criteria. We, therefore, consider a deep water evaporite brine as anything below 50 m. Note, this is in close agreement with Warren (1986) who defined deep water as below 75 m. Recently, however, Evans and Kirkland (1988) have stated that depths below 10 to 20 m constitute a deep brine.

Data Center ,09126599985,[email protected], For Educational Uses DEFINITION OF TERMS 35 1

Evaporite environments involve a wide range of salinities. Terms have been applied to certain ranges of salinities which produce specific rock types. Lang (1937) proposed a classification which in part still has wide usage. Recently Kirkland and Evans (1981) proposed a "mesosaline"evaporite environment for the salinities between normal marine and the point at which CaSO, begins to precipitate. Their "mesosaline" zone would correspond with the "vitasaline"zone of Lang. These same authors used the term "hypersaline" to describe in a general way any brine of greater salinity than seawater. Kirkland and Evans are to be complimented for defining the terms used in their paper; however, they did not explain why a new name was necessary for Lang's "vitasajine". They stressed the importance of their "mesosaline" environment in terms of organic productivity which to a degree was also recognized by Lang who used the Latin "vita" to express this. The term "hypersaline" is well entrenched in the literature but rarely defined. Most commonly it is applied to environments which are precipitating evaporitic salts. We have retained the word to be applied in a general sense to all environments with salinities greater than seawater (Lang, 1937). We would recommend that all those

StaEe Principal Rock Type Salinity Range %o Vitasaline Carbonates 35-142 Penesaline Anhydrite 142-250 Hypersaline Saline Anhydrite + halite 250-350 Supersaline Halite + potash 350 + working in evaporite deposits adapt some common terminology if not that used here. Because they are so important to biomass production in hypersaline environments, it seems appropriate to also include an update on the classification of a group of primitive organisms formerly called "blue green algae".These organisms are prokaryotes and unlike other photosynthetic bacteria contain chlorophyll-a and normally evoke plant type oxygenic photosynthesis. However, in sulfide-rich environments, they are capable of switching to anoxygenic photosynthesis (Cohen, 1984). But, because their cellular organization more closely resembles bacteria, they are no longer classified as algae (Bold, 1973). Most commonly these organisms are called cyanobacteria although there are still some objections to the use of this name. Significant in the case of biomass production, some cyanobacteria contain heterocysts and can fix atmospheric nitrogen. However, under microaerophilic and anaerobic conditions, certain other groups of cyanobacteria fix nitrogen even more effectively than the heterocystous genera (Gerdes and Krumbein, 1984). Thus, phosphorus is probably the limiting nutrient in a cyanobacterial system. Cyanobacteria occur as unicellular, colonial and filamentous species. The term microbial mat is used extensively for a complex ecosystem which is composed principally of cyanobacteria.

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PRODUCTION OF ORGANIC MA'ITER IN EVAPORITE ENVIRONMENTS

Less than two decades ago, organic-richsediments associated with evaporites were generally thought to be the product of a pre-evaporite phase (Schmalz, 1969, and Brongersma-Sanders, 1971). It is true that many of the "saline giants" are immediately underlain by a dark organic-rich facies (Castile, Paradox, Salina, Zechstein, etc.); however, it is now well documented that euxinic conditions prevailed throughout the deposition of many evaporite sequences, resulting in accumulation of organic matter (OM) in all of the facies. The source of some of the organic debris, at least in barred marine basins, is marine plankton swept in by influx currents (Peterson and Hite, 1969). More recently, however, (Kirkland and Evans, 1981) our attention has been drawn to the accelerated rate of biomass production within the actual evaporite environment. Data on biomass production in modern evaporite environments comes from a broad range of natural marine and lacustrine settings as well as salterns constructed for the purpose of sea salt production. Because comparisons are made between source rocks of lacustrine vs. marine derivation in later sections of this paper, a division of setting data was made on the same basis. Under marine we include all settings marginal to or directly connected with the ocean. Most salterns are also included in this setting. For the purpose of our discussion, we have summarized conditions of some of the localities which may best represent geologic deposits. For a very complete compilation of measured organic productivity in evaporite settings, we refer the reader to an excellent paper by Warren (1986). The principal sources of primaq biomass production in hypersaline environments are primitive unicellular organisms. Species of cyanobacteria, such as Aphanothece halophyta, form dense bottom mats which consist of cellular material and large amounts of mucilage. Green algae (Chlorophyta) also contribute large amounts of OM from planktonic forms such as Dunaliella salina and Dunaliella viridis. Dunaliella salina contributes to the red coloration of brines due to the production of carotenoid at high temperatures and salt concentrations. Diatoms (Chrysophyta) can occur in large numbers, especially in the algal mats. Felex and Rushforth (1980) have reported over 27 different species of diatoms from the Great Salt Lake. At salinities, greater than 208/00,the algae are replaced by extreme halophilicbacteria such as the genus Halobac- terium. These bacteria are pigmented by bacteriorhodopsin and bacterioruberins. The former is used to convert sunlight into chemical energy that the bacteria require for metabolism (Stoeckenius, 1976). They are also responsible for the red coloration of concentrated brines.

Marine Settings

Our knowledge concerning the principal sources and rates of biomass production in marine hypersaline environments comes from salterns, a few small ponds marginal to the ocean, marine sabkas, and restricted estuaries. All of these are shallow water environments with maximum depths of no more than 5 meters. Ecologic data from the

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salterns is most complete because of the recognition by industry that the quality and quantity of NaCl production is closely related to the ecology of the pond system. Marine salterns are especially good sources of biologic data for hypersaline systems because they generally represent the full range of salinities achieved by evaporation of seawater. Furthermore, each pond in the chain is semi-segregated so that the ecology of a specific salinity can be precisely determined. Alviso Salterns. Perhaps one of the most intensively studied saltern systems in the world is the Alviso in San Francisco Bay (Figure 4.1). The early work of Carpelan (1957) described the ecology of this system. In the Alviso ponds, as well as all marine salterns, it has been noted that as salinity increases the number of life forms decreases but the overall biomass production increases (Copeland and Jones, 1965). At lower salt concentrations, which correspond to the vitasaline stage, the floor of the ponds are covered by dense microbial mats consisting of the cells and mucilage secretions of Apanothece halophytica. Experimental work by Sammy (1985) shows that Apano- thece grew at approximately the same rate in brines rangingfrom 62 to 224%0. In saltern systems, however, the mats cease to exist in ponds which are precipitating gypsum (Javor, 1983). Assumably, the disappearance of the mats is related more to smothering by rapid sedimentation of gypsum than to intolerance of the higher salinities. At least part of the mucilage secreted by Aphanothece is released to the brine and can create syrup-like viscosities. Little is known about the composition of the mucilage other than it is a complex polysaccharide. The major planktonic contributors to primary biomass in almost all salterns are the unicellular green algae Dunaliella salina and Dunaliella viridis. One notable exception is in the Exportadora de Sal., SA., salterns in Baja California Sur, Mexico, where Javor (1985) has found DunaZiella to be entirely absent through the pond system. This has been explained as probably due to the grazing activities of the brine shrimp Arternia. Above a salinity of about 2000/00, saltern brines support dense populations of halo-

Fig. 4.1. Air view of the Alviso saltern in San Francisco Bay, California. Dark colored ponds are colored blood-red by a dense population of halophilic bacteria. Photograph courtesy of Walter Dean.

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philic photosynthetic bacteria but algae are absent. The bacteria, which impart a pink or red color to the brines, thrive until the supersaline stage at which point the bitterns are devoid of life (Javor, 1983). Halophilic bacteria are not only unique in their ability to stand high salt concentrations, but also show an interesting range of cell morphologies (Figure 4.2). The cubic form must have some utilitarian purpose relating to the crystallization of associated minerals which at high salinities are also dominantly cubic in habit. One of the important characteristics of marine saltern brines is the development of high contents of dissolved organic matter (DOM). Nixon (1970) reported as much as .lo0 mg/liter of DOM in saltern brine. Three brines from the Alviso ponds were analyzed for dissolved organic carbon (DOC) by Hite et a1 (1984). These brines showed an increase in DOC as salinity increased (Table 4.1). The 1.30 brine showed the most drastic increase, in fact much higher than could be achieved by simple evaporative concentration of the 1.21 brine and its DOC content. The concentration of 1.21 brine to 1.31 involves a 62% loss of H,O (Holser, 1979) or a concentration factor of X 5.6 for the DOM. This might suggest an explosion of biomass production in the late stage brines, however, other investigators have found such bitter brines to be devoid of life (Javor, 1983). The 1.30 Alviso brine is quite viscous and is colored a pale yellow-brown quite similar to the color seen in many ancient halite deposits which were deposited from brines near the potash saturation field. Much of the DOC is probably the result of death and autolyzation of algae and bacteria cells (Larsen, 1980) as well as breakdown of the mucilage produced by cyanobacteria. From a 300 g sample of 1.30 Alviso brine, Hite et a1 (1984) were able to extract 10.62 mg of bitumen. When fractionated on a silica gel column, the extract was found to consist of 19.1% saturated hydrocarbons, 12.4% aromatic hydrocarbons, and 64.2% resins, asphaltics and high molecular weight compounds containing N, S, and 0. This brine was studied further by heating to 300°C for 92 hours and analyzing the extracted hydrocarbonsby gas chromatography and gas chromatography-massspectrometry. Some analytical results from this experiment are shown in Table 4.2.

ILE

Fig. 4.2. The various cell morphologies exhibited by extreme halophilic bacteria. Figure courtesy of R. Vreeland.

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Table 4.1. Dissolved organic carbon (DOC) from 3 brines from the Alviso salt ponds (from Hite et al. 1984).

Brine density Saturation field TOC (PPm)

1.16 RvPsum 223 1.21 halite 249 1 ..w halite-potash 2.417 'I'he Alviso saltern is one of the few marine evaporite settings where the amount of organic matter accumulating in the evaporitic sediments has been established. In cores which penetrated up to 50 cm of pond sediment, Klug et a1 (1985) measured 'FOC content and volatile fatty acids (VE'A) from five different salinity sites within the saltern system. These investigators found that TOC values increased up to a Yoo/oo salinity. A maximum of 23.1% '1OC was found in the sediments of the pond with 9oo/o salinity. 'TOCvalues decreased in the sediments of the 150%0 and 3W/oo ponds. 'l'his decrease at the higher salinities is probably due to dilution by rapidly precipitating gypsum and halite. Klug and his associates also analyzed volatile fatty acids (VFA) in these sediments ('I'ables 4.3,4.4,4.5,4.6and 4.7).The amount and spectra of VFA was found to increase with increasing salinity. Acetic acid, which was the principal VFA, was found to increase with increasing salinity in all pond sediments reaching a measured value of 802.33 p mol/liter in the upper 1 cm of sediment in the 3Whpond. Another significant finding was the decrease in bacterial sulfate reduction in pond sediments with increasing salinity. Klug et al (19855) suggested that the increased accumulation of organic matter was due to decreased activities of fermentative bacteria. 'l'he findings of Klug, et a1 (1985) are strongly supported by the conclusions of Koyama et a1 (1973) who showed that the production of VFA was not inhibited by high salinities but that high salinities do retard destruction of OM by methanogenic bacteria. The Solar Lake. Located in the Gulf of Elat, along the Sinai coast, "Solar Lake" is a small heliothermal brine pond. 'The pond covers an area of about 9,100 m2 and has a maximum depth of about 5 m. Salinities in the epilimnion range from 7V)ho to 1Who and may reach up to 185Y00 in the hypolimnion. Overturn and mixing during the period of May to September destroys stratification; and at that time, the brine salinity throughout the lake is about 1W/oo to 185Yoo (Cierdeset al, 1985). When the lake is strongly stratified,

Prislane/phytane <1 Hopdnes/steranes h>s Steranes G>G Diasteranes minimal Tricyclic maximum C2, C2, tetracyclic moderate Extended hopanes moderate Gammacerane minimal

Table 4.2. Organic chemical characteristics o,f a 1.30 density brine from the Alviso salt ponds after heating at 300°C ,for 92 hours.

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Table 4.3. Wt. % TOC on a dry sediment basis, wt. 70 porewater (PW.), and p mol/liter fatty acids in porewaters from various depths from a salt marsh adjacent to the Alviso pond system. Salinity of water overlying sediment is 33%. From Klug et a1 (1985).

Depth (cm) T.0.C. pw. (a) (b) (c) (4 (4 (f) (9)

0.1 11.7 76.0 19.70 -- .. __ -- 0.12 1-2 9.9 70.4 20.59 0.14 __ -- 0.12 2-3 9.8 68.8 19.62 0.83 __ _. -- 0.51 55 9.7 67.9 4-5 ._ -_ 9.62 -- 5-7 9.9 63.8 __ 7-8 __ 10.44 -- 9-11 11.4 69.6 10.11 10.44 -- 11-13 9.6 67.7 __ u-15 9.5 65.9 14-15 ._ -_ 8.69 0.68 17-19 9.0 66.6 ._ 18-19 82) 0.58 21-23 8.3 65.5 __ __ 23-24 5.47 0.03 2525 8.1 64.1 __ 26-28 8.2 62.1 .. 28-30 8.1 61.5 33-32 8.6 60.5

(a) acetic acid @) isobutyric acid (c) propionic acid (d) n-butyric acid (e) iwaleric acid (f) unknown acid retention time 7.35 min. (g) unknown acid retention time 5.25 min.

Table 4.4. Wt. % TOC on a dry sediment basis, wt. %porewater (PW.), and p mol/liter fatty acids in porewaters from various depths from an Alviso salt pond. Salinity of water overlying sediment is 42%0. From Klug et a1 (I 985). See above f or column definitions.

0-3 14.9 74.9 43.64 0.69 2.27 -- -- 2.73 3-6 11.1 75.3 59.30 -- 2.62 -- -- 2.49 6-9 10.8 62.4 19.17 -- 322 -- -- 323 9-12 14.1 69.1 13.M -- 2.24 -- -- 2.14 12-15 16.3 72.7 15.86 -- 1.57 -- -- 1.60 15.18 16.3 72.7 14.51 1.02 0.99 -- 18-21 19.4 74.5 9.34 -- 0.82 -- -- O.% 21-24 19.8 73.8 30.64 -- 0.71 -- -- 0.68 27-30 172 72.3 40.73 2.U 1.90 -- -- 1.69 30-36 13.4 67.9 -- ._ .. 3336 __ -- 39.78 0.71 121 -- -- 0.90 39-42 15.8 71.8 26.09 0.95 0.67 -- -- 052 45-48 15.4 67.8 -- ______51-54 15.9 69.4 34.38 -- 0.07 -- -- 0.28 57-60 15.6 68.9 20.72 -- 0.55 -- -- 0.57 63-66 15.9 70.4 29.47 -- 157 -- -- 1.50 69-72 15.0 70.8 13.82 -- 0.31 -- -- 0.31 72-75 14.6 69.4 l3.97 -- 0.40 -- -- 0.60 78-81 135 63.8 11.75 -- 0.30 -- -- 0.42

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Table 4.5. Wt. % TOC on a dry sediment basis, wt. % porewater (PW.) and p mol/liter fatty acids in porewaters from various depths from an Alviso salt pond. Salinity of water overlyingsediment is 9Mo. From Klug et a1 (1985). See Table 4.6 for column definitions.

Depth (cm) T.0.C. PW. (a) (b) (4 (4 (4 (9 (9)

C-1 20.9 82.1 5.67 -- -- 0.13 -- 1-2 15.1 76.8 16.911 0.03 -- __ -- 021 -- 2-3 172 76.9 21.56 2.0 -- -_ -- 2.50 -- 3-5 17.9 775 -- __ _- __ -_ -- __ 4-5 __ -- 3738 6.45 -- __ -- 8.25 0.17 5-7 23.1 80.8 -- ______7-8 -- 20.76 2.51 -- __ -- 2.91 -- 7-9 22.5 81.0 -- ______- 9-11 17.3 772 -- __ -______10-11 __ -- 31.76 4.59 ------5.14 0.15 11-13 18.1 76.5 -- ______-- -- _. 13-14 __ -- 37.7.64 3.43 -- __ -- 5.01 0.04 16-17 -- 42.62 9.70 -- __ -- 9.73 0.24 19-20 -- 45.62 10.67 -- __ -- 9.09 -- 22-23 __ -- 16.79 5.67 -- -- 4.73 0.18 2526 __ -- 3234 8.68 -- __ -- 8.33 --

Table 4.6. Wt. % TOC on a dry sediment basis, wt. % porewter (PW.), and p mol/liter f arty acids in porewaters f rom various depths f rom an Alviso salt pond. Salinity of water overlying sediment is 150%. From Klug et a1 (1985).

Depth(cm) T.O.C. PW.

-~ 0-1 __ 0.09 1-2 162 72.6 223 3.82 2-3 17.6 73.6 %[email protected] 34 __ _- 3-5 182 71.7 __ __ 4-5 __ 20.63 4.73 5-7 15.0 63.4 __ __ 7-8 -- __ 35.87 735 7-9 12.4 61.1 __ 9-11 12.1 57.8 __ __ 10-11 __ 40.70 8.90 11-13 14.7 66.8 __ -_ 13-14 __ __ 19.80 1.22 15-17 9.7 59.8 __ _- 16-11 __ 18.65 I27 17-19 8.3 55.3 __ __ 19-20 __ 45.76 3.81 19-21 9.4 56.1 __ __ 22-23 __ -_ 30.63 1.U 25-26 __ __ 35.30 1.a

(a) acetic acid (d) n-butyric acid @) isohutyric acid (e) isavaleric acid (c) propionic acid (f) uaknown acid retention time 735 min. (g) unknown acid retention time 5.25 min.

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Table 4.7. Wt. % TOC on a dry sediment basis, wt. % porewater (PW.), und p mol/liter fatty acids in porewutersfrom various depths from an Alviso salt pond. Salinity of water overlying sediment is 3oo%o. From Klug et a1 (1985).

Depth (cm) T.O.C. PW. (a) (b) (c) (d) (e) (f) (g)

0-1 12.1 51.4 802.33 1w.76 9.95 5.88 592 85.92 4.70 1-2 14.0 56.3 736.97 150.40 4.66 2.46 3.85 114.06 2.88 2-3 13.2 55.0 N9.66 34.53 -- 1.07 -- 26.43 1.77 34 13.5 54.0 -- ._ ._ 4-5 -- 288.94 65.05 0.73 2.65 2.30 32.96 2.00 4-6 15.1 55.7 -- ______5-6 __ -- 248.94 65.05 0.73 2.65 2.M 32.96 2.00 6-8 152 55.4 -- _. 8-10 16.0 54.4 -- __ .. 10-12 15.0 54.8 -- __ .. 12-13 __ -- 45.31 1468 -- __ 6.78 0.23 12-14 14.5 59.9 14-16 11.0 59.7 -- .. 18-20 10.9 60.6 -- __ .. .. 20-21 -- 27.41 7.05 -- 3.15 -- 24-25 -- 39.25 -- 6.84 -- 3.03 -- 28-29 -- 57% 7.3 -- 3.71 --

(a) acetic acid (d) n-butyric acid @) isobutyric acid (e)iswaleric acid (c) propionic acid (f) unknown acid retention time 7.35 min. (g) unknown acid retention time 5.25 min.

the bottom brine can reach temperatures of 60°C due to the "Green House" effect (Hirsch, 1980). "Solar Lake" is separated from the Red Sea by a low barrier, about 60 m wide composed of porous sand and gravel through which seawater passes to replace evaporation losses from the lake. The margins of the pond are covered by microbial mats down to a depth of about 2 m. An extremely high biomass production, 1,810 mgC/m'/day, was reported for this penesaline environment by Cohen et al, (1977b). Sediments of the "Solar Lake", which have been described by Friedman et al, (1973) and Gavish (1980), consist of finely laminated microbial mats (carbonate laminae) overlying carbonate mud along the shallow shelves of the lake and gypsum laminated by carbonate and underlain by carbonate on the slopes and bottom of the lake. A total of 70 cm of sediments has been deposited in a period of about 2,400 years. In recent years, the microbial mat sequence in Solar Lake has been under intensive study by organic geochemists because the organic rich sediments are autochthonous, exclusively marine and as a result, can serve as a model to relate to the biogeochemical processes of ancient microbial mats (Boon, 1984). An important observation by Friedman et a1 (1973) was that the finely laminated microbial mat of Solar Lake was probably preserved because high salinities excluded burrowing and grazing animals. They suggested that the mats developed only when a salinity of 60-80%0 was exceeded. Similar salinities in the Gavish Sdbkha, which is also located in the Gulf of Elat, also limit the activities of grazing animals (Gerdes and

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Krumbein, 1984). The first author of this report has also observed the effects of destructive grazing on microbial mats in salt ponds in the Gulf of Suez (Figures 4.3A and 4.3B). The sediments of the "Solar Lake" are rich in organic carbon. Selected samples analyzed by Boon et al (1983) show (Table 4.8) that, except for the first few millimeters of sediment, TOC remains relatively constant down to a depth of 658 mm (Boon et a1 (1983). Shidlowski et al (1985) obtained even higher TOC values from "Solar Lake" in a 1 meter deep core. These higher values (Table 4.9) may be the result of slightly different analytical techniques or different sample locations. In addition to TOC in the "Solar Lake" sediments, the dissolved organic carbon (DOC) in the sediment porewater has also been determined. Lyons et a1 (1982) found significant concentrations of DOC at sampling depths rangingfrom 5 to 81 cm. The high DOCvalues (1 19-818 mg/l), according to Lyons et al(1982), are probably not imparted by "photosynthetic exudutes". These authors also concluded that the high temperatures sometimes reached in the brines (600C) may bring about the initial stage of maturation of the OM in the sediments creating a buildup of low molecular weight organic acids. On the basis of their work in "Solar Lake" and other hypersaline environments, Lyons et a1 (1984) concluded that only where qanobacterial mats are found are the rates of bacterial reduction of sulfate high. This is probably due to the fact that the mat would provide the ma,jor source of carbon. Furthermore, because most species of cyanobacteria fix nitrogen, the principle growth limiting nutrient could be phosphorus. But, as they have shown, phosphorus is rapidly recycled from the upper part of the mat as the result of mat degradation by bacterial sulfate reduction. Since the "Solar Lake" is receiving a constant supply of phosphorus via infiltrating seawater, it would seem likely that the supply of phosphorus would steadily increase with time. It would seem that in hypersaline environments, such as "Solar Lake", conditions are ideal for accelerated biomass production. Ultimately, the phosphorus concentration might reach a point where inorganic precipitation of phosphate minerals would occur. This relation-

Fig. 4.3. (A) Microbial mat in a snzall brine pool on El Heniiet Island. Gulf of Suez, Eg-vpt. Salinity of the brine is about 160%0. (B) Total destruction of microbial mat in the El Herniet lsland brine pool bvgrazing gastropods. Note trails left on bottom sediments.

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Table 4.8. Wt. 70TOC on a dry weight basis from select sediment .samples from “Solor Lake”. From Boon et a1 (1983).

Depth (cm) T.O.C.

0 - 0.3 9.7 0.3 - 1.0 4.6 1.0- 2.0 3.4 11.5 - 14.5 2.1 26.5 - 29.5 2.9 38.5 - 41.5 3.4 62.8 - 6.5.8 3.0

ship might explain the frequent association in ancient sediments of laminated carbonates and phosphorites. Another interesting observation made by Lyons et a1 (1984) is that not all hypersaline environments produce microbial mats. In Bonaire, for example, they noted that hypersaline lakes underlain by Carbonates produce mats but that in lakes underlain by volcanic rocks no mats are present. The Orca Basin. In the Gulf of Mexico on the Louisiana continental slope is a brine filled depression known as the Orca Basin. The closed portion of the basin covers an area of about 400 km2and is depressed roughly 500 m below the sea floor (Demaison and Moore, 1980). The depression is thought to be related to subsidence caused by dissolution of a near surface salt body (Addy and Behrens, 1980). The basin contains a 200 m thick layer of brine which has a salinity of 25Ph0,and thus the basin qualifies as a saline environment. This basin, and others like it, such as the Red Sea brine pools and the recently discovered (ten Haven et al, 1985) Tyro basin in the eastern Mediterranean, constitutes an unusual but effective environment for trapping and preserving OM. It is unusual in the sense that unlike other saline environments, it does not generate OM. Instead, the organic rich sediments in this basin are the product of slow sedimentation of planktonic OM from the overlying Atlantic water of normal salinity. The high TOC levels in these sediments have been explained as the result of high productivity in the overlying Atlantic water rather than any special preservation qualities of the brine pool (Sheu, 1983). We suggest that the minimized activities of anaerobic bacteria, due to the high salinities in the brine pool, is also a factor. An important consideration in the Orca Basin model is that if the density of the brine is too high the OM cannot sink (Holser, 1979). However, this problem may be circumvented by mineral crystals nucleated by, and attached to, the OM making combined densities higher than the brine and allowing settling. Some in situ decomposition of the OM in the brines does occur, however, as indicated by the concentration of methane in the brine (750 mM/L) (Wisenburg et al, 1985). In addition to methane, the brine also contains ethane (1,300 mM/L). Wisen- burg et a1 (1985) believes the methane is the product of in situ decomposition of OM by methanogenic bacteria. Along with the gases, the brine contains as much as 3.5 mg/L of DOC and 70 pg/L of particulate OC (Sackett et al, 1979). Cored sediments beneath the brine pool contain as much as 2-3% organic carbon (Northam et al, 1981). Saline basins, such as the Orca, could accumulate sediments with important

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Table 4.9. Wt. % TOC on a dry weight basis from sediment samplesf'rorn Solar Lake. From Schidlowski et a1 (1985).

wth(cm) TOC

0 4.6 n- 3.4 14.9 3.0 4.6 5.0 4.6 3.4 - 7.5 16.1 12.8 - 16.4 10.7 21.7 - x.n 9.6 40.0 5.4 3.0- 42.0 18.9 .W.O 5.4 3.0- 58.0 20.6 62.0 - 66.0 15.6 M.0 - 72.0 42 90.0 - 100.0 1.3 petroleum source rock characteristics. Discovey of such sediments in the stratigraphic column would present an enigma to organic geochemists because their organic constituents would be made up primarily of marine plankton from a normal marine environment. To make matters even worse, there would be little, if any, mineralogical evidence that the sediments had been preserved in a saline environment.

Lacustrine Settings

The Lacustrine evaporite environments described here can be divided into four types based on brine chemist7 and morphology of the evaporite setting. Alkaline brines of the sodium/bicarbonate/carbonate type seem to dominate with sodium chloride brines playing a subordinate role. Examples of the latter are the Great Salt Lake and the Dead Sea. Morphologically, there is the deep basin type generally with precipitous margins commonly found in rift zones. The Dead Sea is a perfect example of this type. The other setting involves broad flat floored depressions with thin brine layers, some which may at times completely desiccate. Examples of this type are the Tuz Golu in Turkey and the Great Salt Lake of Utah. Lake Van-Turkey. Lake Van in eastern Turkey is the largest soda lake in the world covering an area of 3,574 km'. Maximum measured depth is 457 m and the volume of the lake is 607 km3 which makes it, in terms of volume, the world's 4th largest closed lake (Kempe et al, 1978). The lake sits in a relatively steep-walled basin whose morphology is fault controlled (Wong and Finckh; 1978). The salinity of the lake is 2%0 with a pH of 9.55. The chemical composition of the lake water (Table 4.10) is relatively uniform due to annual overturn which keeps the water well mixed and also oxygenated. The DOC content of the lake water is apparently low although only a few measurements have been made (Kempe and Degens, 1978). Because the phosphate content of the lake water is quite high (Table 4.10), it is somewhat surprising that the lake is not more productive. Gessner (1957) suggests that the biolimiting factor is nitrogen which,

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Table 4-10. Major ions in brine from Lake Van, Turkey in m,q/liter. From Kempe et al, 3978.

Cations Anions

Stdium 7,747.00 Chloride5,450.00 Potassium 508.00 Carbonate 3,331.00 Magnesium 94.80 Sulfate 2,344.00 Calcium 5-10.00 Bicarbonate 2,19l.(X) Lithium 1.50 Phosphate 0.52 Strontium 0.70 in turn, would suggest low activities of nitrogen-fixing cyanobacteria. The low salinity of the lake no doubt is also a factor. The phytoplankton of the lake consists principally of diatoms whose maximum population is at a depth of 20 to 30 m. Below that depth, the siliceous frustules of the diatoms tend to dissolve in the high pH water. Because the entire water column of the lake is oxygenated, much of the diatom biomass may be destroyed by aerobic bacteria before reaching the lake bottom. The sediments of Lake Van have been cored to a depth of about 9.5 m. These sediments are laminated and varve counts show that they represent 10,420yearsof continuous sedimentation (Degens et al, 1978). During this period, it is known that the lake level fluctuated as much as 400 m. At the lowest levels, the lake brine could have been concentrated to the threshold of sodium carbonate precipitation. TOC distribution in the lake sediments show that, other than an interval deposited between 8,500 and 9,500 years B.P., there is a near linear upward increase in TOC. Maximum TOC value was 7.08 wt. 96 at a depth of 2.2 m. The well preserved laminae in the sediments suggests limited activity by burrowing animals which is something of an enigma considering the oxygenated water column and the high TOC values present. The bicarbonate-carbonate enriched water of Lake Van is stated by Degens et a1 (1978) to be the result of the chemistry of the river waters which flow into this closed basin. If this is true, it would suggest that the activities of sulfate-reducing bacteria, which return by-product bicarbonate to the brine, have been minimal. The apparent low productivity of the lake suggests that much of the OM in the lake sediments is carried in by rivers. Unfortunately, the OM in the lake sediments has not been characterized so its origin can only be speculated on. Because of its large size and alkaline water, the organic geochemistry of the lake should be studied in greater detail as a possible analog to ancient associations of oil shale and sodium carbonates. Wadi Natrun - Egypt. About 80 km northwest of Cairo is a subsea depression (-23 m) which contains several small shallow lakes. These lakes, some of which are completely dry during the summer, receive a limited supply of groundwater which seeps into the depression from the nearby Nile River. Because the depression has no outlet and the evaporation rates are high, the water in the lakes has a high salt concentration. In studies of 6 of these lakes, Imhoff et a1 (1979) found salinities ranging from 91.9 to 393.9%0 which makes the concentration of these lakes brines an order of magnitude higher than Lake Van in Turkey. The alkaline brines of these lakes all have a pH of about 1 1 .O and contain large amounts of DOM (Table 4.1 1). The variation in DOM

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Table 4.1 I. Chemical composition of brines from 6 Wudi Natrun Lakes, Egypt. From Imhoff et al. IY79.

Laxcation Gaar Gabara Hamara Muluk Ruunia Zugm

Ions in g/l K 1 ..Y) 0.21 1.cn 0.m 2% 2..w

Na 1.%.99 31.N 79.N 54.(W) 13s.5.50 142.M

c1 173.70 20.92 86.85 66.20 155.98 154.56

SO, 4x.m 17.29 w.m IXSK 31.70 22.57

co, 6.60 19.w 23.40 4.40 49.m 67.21

TDS (dI) 374.20 91 .()o 237.W 159.20 .m.m 393.m

Organic Matter (mM as CO*) 115.70 32.60 47.70 54.70 12930 I13) between the 6 lakes might be a function of different rates of productivity which relates to salinity (highest salinity = highest DOM), but each lake has a somewhat different set of environmental conditions. For example, Muluk frequently is totally desiccated, Zuym has a solid floor of salt, and in others, the lake bottoms are disturbed by mining operations. The latter would have a tendency to release dissolved organic acids in the sediment porewater to the overlying brine. The lakes in Wadi Natrun have high productivities and are populated by dense colonies of cyanobacteria, green algae, halobacteria, and phototrophic sulfur-bacteria. Conspicuously absent are brine shrimp. Apparently sulfate reducing bacteria are active in the lake's sediments and the high carbonate reported in the brines is a metabolic product. Methanogenic bacteria are not present and other than a small population of zooflagellates, there are no higher organisms feeding on the bacteria and algae (Imhoff et al, 1979). Hammer (1981) has estimated the primary production of organic matter to be 2,601 gC/m2/yr in Lake Mariute. Conditions in the lakes of Wadi Natrun are apparently ideal for large biomass production as well as preservation. Unfortunately, little, if any, work has been done on the organic geochemistry of the sediments. Great Salt Lake, U.S.A.. Although it is but a small (6,200 km2) desiccated remnant of the huge ancestral Lake Bonneville, the Great Salt Lake of Utah is the largest saline lake in North America. At its highest level, Lake Bonneville maximum water depth was about 350 m (Currey, 1980); but during historical times, the maximum

Data Center ,09126599985,[email protected], For Educational Uses 364 PETROLEUM AND EVAPORITES depths have ranged from 14.7 m to 21.7 m. In 1959, the lake was divided by a railroad causeway into a north and south arm (Figure 4.4). Because the north arm received no freshwater inflow to replace evaporation losses, and circulation between it and the south arm was restricted to small openings in the causeway, the salinity of the north arm became much greater than the south arm. In 1979 deeper brine of the south arm had a total dissolved solid (TDS) content of 258.82%0 while the bottom brine of the north arm had a TDS of 341.44Yoo (Sturm, 1980). The composition of the Great Salt brine, in general, resembles concentrated sea water (Table 4.12). The saline environment of the Great Salt Lake is not only divided into two distinct salinity regimes but each arm of the lake also represents a different ecologic community. The south arm of the lake is a blue-green color (Figure 4.4) due to a dense population of green algae, cyanobacteria, diatoms, and brine shrimp. Prior to 1984 the north arm was a wine red color due to a massive population of bacteria whose cells contain carotenoid pigments and bacteriorhodopsin (Post, 1980). During 1984, the railroad causeway was breached by a 82.3 m opening to allow salinities to equilibrate between the north and south arms. Since that time, the red color of the north arm has disappeared. Besides bacteria, the next important contributor of biomass in the north arm are algae which consists of two planktonic species: Dunaliella viridis and Dunaliella salina. The latter species is red pigmented and more abundant than the green pigmented D. viridis. During 1973, Stephens and Gillespie (1976) found the average annual phytoplankton production in the south arm to be 145 gC/m’. This production comes principally in March and April. In the north arm, the brine contains as much as 44.1 mg/l of DOC (Post, 1980). In one of several cores taken from the sediments of the south arm of the Great Salt Lake, Spencer et al(1984) measured TOC to a depth of about 530 cm. The TOC content through this interval ranged from 0.4% to 4.9% on a dry weight basis. The highest concentrations of TOC correlate with periods of high salinity and a low level

Fig. 4.4. Air view of brineflowing through an opening in the railroad causeway separating the north arm (upper part of photograph) and the south arm (lower part of photograph) of the Great Salt Lake, Utah. Theflow direction is from south to north. The brine of the south arm (very dark inphotograph) is darkgreen in color. The north arm brine (lightgray in photograph) is pinkish-gray. Photograph courtesy of Lloyd Austin.

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Table 4.12. Chemical composition in 1976 of north arm brine of the Great Salt Lake, Utah. From Sturm, 1980.

Ion 811

K 8.48 Na 105.U Mg 10.42 ca 0.42 a 181.73 so4 21.73 Li .05 Br .l2 B .04 pH - 7.3 TDS - 328.1 1 g/l stage of the lake. At least one interval of high TOC content appears to be time equivalent to a mirabilite (Na2S0,*10H20)bed in the north arm of the lake. The Dead Sea. The Dead Sea occupies a deep subsea rift valley on the Israel-Jordan border. Its surface area covers about 940 km2and it has a maximum depth of about 400 m (Neev and Emery, 1967). When studied by Neev and Emery (1967), the Dead Sea was a stratified lake with an upper water mass (-3W?o), a transitional zone (-320%0), and a lower water mass (-33Woo). Since that time, due largely to the intervention of man, the upper water mass salinity has steadily increased and in 1979 the lake overturned. During the time of their investigation, Neev and Emery (1967) found significant differences in the composition of the upper and lower water masses (Table 4.13). Brine samples from depths of 0 to 130 m averaged 6,625 mg/L dissolved organic carbon (Neev and Emery, 1967). Similar to other saline environments, the high salt concentrations of the Dead Sea have limited the diversity of biota. The principle organisms in the water column are red halophilic bacteria whose population can reach 8.9 x lo" cells per cm3 and the alga Dunaliella viridis whose population has been measured at 4 x 104 cells per cm3 (Kap- Ian and Friedman, 1970). Thus, the biologic community and standing biomass are very similar to the Great Salt Lake of Utah. Oren (1983) sampled surface brine in the au-

Table 4.13. Chemical composition of the upper and lower water masses of the Dead Sea g/l. From Neev and Emergy, 1967. Upper Lower Ion Water Mass Water Mass

ca 16.38 17.18 Mg 36.15 42.43 Na 38.51 39.70 K 6.50 7.59 cl 196.94 21925 Br 4.60 521 SO, 0.58 0.42

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tumn of 1981 and found a very low population of Dunaliella and a dense concentration of halobacteria (4-5 x 106 cells per ml-'). According to Oren (1983), these bacteria may be a dominant factor in photoassimilation of CO, in the Dead Sea as well as in other saline environments. Although Neev and Emery (1976) found as much as 2% TOC in Dead Sea sediments, previous investigations by Nissenbaum et a1 (1972) showed a range of only 0.23 to 0.40%. The latter authors characterized the organic compounds found in sediments of both deep and shallow parts of the Dead Sea. They found that between 40 and 50% of the TOC in the sediments in the deep and strongly anoxic part of the sea were humic and fulvic acids but that these acids were reduced to only 1.5 to 3.2% of TOC in the shallow oxygenated margins. Various plant pigments were isolated from the sediments, the most abundant being a -carotene. Surprisingly, a significant amount of chlorophyll was also discovered in the deep basin sediments. The isoprenoid hydrocarbons phytane and pristane were present with phytane slightly dominant. Nissenbaum et a1 (1972) proposed that the phytane and pristane may well have been derived from an ether-linked lipid in the halophilic bacteria rather than chlorophyll, because the rate of chlorophyll degradation in this environment seems very slow. A similar proposal was made by Kates (1978). It is noteworthy, however, that ether bonds are much more stable than the ester linkage of phytol-chlorophyll. Other data on Dead Sea HC showed that 80% of the HC from deep basin sediments were n-alkanes C,, and C,, but that these compounds were only 4 to 16% of the HC fractions in shallow water sediments (Nissenbaum et al, 1972). All samples had a strong odd carbon preference. Fatty acids were present in the sediments in amounts rangingfrom 407 to 1,503 mg/100 g and were equally divided between free and hydrolyzable ("bound') acids; most of the free fatty acids were C,, and C,, while the hydrolyzable fatty acids showed a broader but unspecified range. Even carbon numbered acids predominated in both groups.

THE SPATIOTEMPORAL RELATIONSHIPS OF PETROLEUM AND EVAPORITES

Recently, Parparova et a1 (1981) plotted the intensity of accumulation of coal, bitumen, oil, gas and total organic carbon through the Phanerozoic. Similarly, Zharkov (1981) has produced the most complete calculations of evaporite volumes in the stratigraphic column for the Paleozoic. We made additional estimates of evaporite volumes for the remainder of geologic time. Although the latter are less than precise and will no doubt be changed by future workers, they allow us to make a direct comparison of oil volumes to evaporite volumes throughout geologic time (Figure 4.5). In examining the plot, one finds correlation between large volumes of evaporites and large volumes of oil only in the Upper Devonian, Lower Cretaceous, and the Lower Neogene. This might suggest that there really isn't a significant relationship between evaporites and petroleum after all. Szatmari (1980) has pointed out that "within each salt basin, petroleum deposits are contained in an envelope that underlies, overlies and surrounds partially dissolved

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I ------N;' r. ------U I \ Pa M I L I I I 100

200

300

400

500

600 I I I I I I 1,400,000 1,000,000 600.000 200,000 1.200.000 800,000 400,000 VOLUME OF EVAPORITES IN Km3

Fig. 4.5. Stratigraphic distribution of the World's oil volumes plotted against evaporite volumes. Oil volumes from Parparova, et al, (1981). Paleozoic evaporite volumes from Zharkov (1981). Mesozoic and Tertiary evaporite volumes estimated by the authors (dashed lin?). or shattered source evaporites". Szatmari acknowledges that to some extent the petroleum envelope may suffer considerable localized stratigraphic dislocation where salt diapirs, erosional hiati, or tensional faulting are present. Some actual numbers on the relationship between oil and gas horizons in evaporite basins have been reported by Kalinko (1974) who states that 46% of the basins have oil and gas horizons below the evaporites, in 41% they are above, and only 13% are in the evaporite facies. Considering the envelope concept, one might expect to see a major period of petroleum accumulation either immediately above or below the Lower Cambrian, Lower Permian, Lower Cretaceous and Lower Neogene. This does hold true for the Lower Cretaceous and to some extent the Lower Neogene, but is not evident for the other evaporite maxima. The lack of correlation between periods of maximum accumulation of evaporites and stratigraphic distribution of petroleum could be due to a number of reasons, such as: 1) not all evaporite sequences were conducive to the accumulation of organic matter

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and the generation of hydrocarbons, 2) the evaporite-petroleum envelope commonly has a larger stratigraphic displacement than suggested by Szatmari ( 1980), 3) appropriate reservoirs for evaporitic petroleum may not have existed, 4) a large amount of petroleum may be associated with evaporite deposits which are too small to effect the volume plots, and 5) many petroleum Occurrences may be directly associated with an evaporite environment (vitasaline) that only produced carbonate sediment and thus would not be included in the evaporite volume data. A resolution of the problem will not come until a worldwide documentation of specific petroleum-evaporite associations is made.

EVAPORITES AS SOURCE ROCKS OF PETROLEUM

There is ample evidence to show that the evaporite environment is capable of generating abundant amounts of organic matter and then preserving it. But what is the evidence that evaporites are actually important source rocks'? The question is difficult to answer because of the somewhat meager amount of data concerning the actual contents of TOC or extractable hydrocarbons (HC) in these rocks. Part of the problem seems to be recognition of the full suite of rocks produced by the evaporite environment.

Evaporitic Carbonates

Kirkland and Evans (1981) have emphasized the importance of the carbonates produced during the vitasaline stage (their mesosaline stage) as source rocks. As we have seen from measurements made on modern evaporite environments, this is not the time of maximum organic productivity. However, the relatively low productivity of the vitasaline stage seems to be more than made up for by the fact that this stage represents the slowest sedimentation rate of an evaporite depositional cycle. For example, in a marine evaporite basin we might expect the depositional rate of carbonate rock compared to halite to have ratios as high as 1:400 or greater. Thus, even though the accumulation rate of OM might be greater during higher salinities, it will be subjected to greater dilution by sediment. The organic geochemistry of the vitasaline carbonates is known much better than rocks of the other evaporite phases. This is due in part by a recent symposium on Carbonate Source Rocks (Ed. J. G. Palacas, 1984) in which several papers present detailed organic geochemistry of carbonates that are evaporitic in nature. One important paper by Palacas et a1 (1984) describes the source rock characteristics of organic rich carbonates associated with anhydrites in the Cretaceous of Florida. These limestones, some of which consist of up to 50% argillaceous matter, have an average TOC content of 1.8%. Extractable organic matter (EOM) ranged from 509 to more than 8,170 ppm. Some of the dark gray laminated micrites of this sequence have HC/TOC as high as 15%, and Palacas et a1 believe this to be the result of in-situ HC generation from oil-prone (hydrogen rich) OM. On the basis of analyses of kerogen concentrates by

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microscopic examination and HC/C ratios (avg. 1.15), Palacas et a1 (1984), suggested that "most of the solid organic matter in the Sunniland and associated carbonate rocks belongs to the type I and type I1 kerogen classes (Tissot and Welte, 1978) derived chiefly from algal, planktonic and microbial residues that accumulated in saline to hypersaline, moderately to highly reducing environments". The study by Palacas and his co-workers was designed principally to show that the organic-rich carbonates were the source rocks for the petroleum accumulations in porous carbonate reservoirs in the same formation. For this purpose, they made an extensive examination using gas chromatography-mass spectrometry of biological marker compounds (steranes, tricyclic and pentacyclic terpanes) which they felt were much more significant in correlation than the more conventional properties such as molecular distribution of n-alkanes, isoprenoid ratios, etc. If we can assume that the rocks they studied represent the vitasaline stage of the evaporite environment, then their biomarker data are valuable for making comparisons to similar data from other evaporite source rocks. One of the interesting biomarker relationships in the Sunniland is that the C,, sterane concentration is equal to or exceeds the c7sterane concentration. According to Palacas et a1 (1984), this relationship might be the result of "...intense microbiologic activity and highly preservative saline conditions...". In contrast, Mello et a1 (1988) found the c7-C;,sterane ratio to range from 1.0 to 2.2 for oils and rocks from Brazilian evaporites. Unfortunately, Mello et a1 (1988) did not give the lithologies of the rocks they sampled. In summary, the literature suggests this ratio is unreliable for the determination of OM source. Two other papers in the symposium on Petroleum Geochemistry and Source Rock Potential of Carbonate Rocks discuss the organic geochemistry of rocks whose facies relationships suggest they are evaporitic carbonates. Oehler (1984) describes carbonates in the lower part of the Jurassic Smackover Formation as algal laminated micrites which were deposited in a hypersaline to shallow intertidal environment. These micrites have average TOC and EOM values of 0.48% and 259 ppm, respectively. The average EOM/TOC is 7% and the HC/TOC average is 3.6%. Oehler stated that these carbonates were the source rocks for the prolific Smackover trend in Mississippi, Alabama and Florida. Organic-rich rocks, generally referred to as black shales, are elements of a thick sequence of evaporitecarbonate cycles in the Paradox basin of southeast Utah and southwest Colorado. These Pennsylvanian age "shales"are actually argillaceous-siliceous carbonates whose TOC values range from 0.5 to 13.0%. These rocks were deposited in a vitasaline environment and apparently represent sedimentational tongues from a deltaic sequence marginal to the evaporite basin (Hite et al, 1984). The kerogen in these "shales" is a mixture of types I1 and I11 and seems to have an increasing component of terrestrial plant contribution towards the eastern margin of the basin. In addition to sedimentological evidence and kerogen type, isotopic signatures of the carbonates show they were derived from meteroic water influx (Magaritz, 1987). Another factor that has been overlooked concerning the genesis of the Paradox black "shales", as well as similar organic-rich facies in other evaporite basins, is the possible contribution of humate

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precipitates. If meteoric waters carried in particulate debris of terrestrial plants, it seems quite possible that some of this water might represent coal swamp drainage and as a result, could cany appreciable amounts of humic acids in solution. When these waters mixed with brines of the evaporite basin, the acids would precipitate in the manner described by Swanson and Palacas (1965). Thin laminae of black coal-like material can occasionally be found in the Paradox black "shales". These laminae may be the result of such a mechanism. Despite all the evidence for a major component of terrestrial plant material in their kerogen, the Paradox "shales"are oil prone source rocks and have been responsible for the production of about 53.6 million metric tons of oil (Baars and Stevenson, 1982). Extracts of some of the organic-rich Paradox "shales"show EOM as high as 13,119 ppm with a ratio of EOM/TOC of 32.8%. Saturated to aromatic hydrocarbon ratios (S/A) range from 3.5 to 14.7. Extracts examined by GC and GC-MS have high pristane/phytane ratios (1.1 to 1.4) and n-C,,/pristane ratios ranging from 1.5 to 7.1 (Figure 4.6). The relatively high pristane/phytane in the Paradox "shales" probably reflect a strong terrigenous input of OM. In all the rock extracts analyzed, no pronounced odd or even n-alkane predominances were observed. The distribution of sterane and tri-, tetra-, and pentacyclic terpane biomarkers in the Paradox black "shales" (Figures 4.7A and 4.7B) are typical of carbonates in the marine evaporitic sequence described by Palacas (1984, p. 85 and 92) for the lower Sunniland Formation. The regular G7/C9sterane ratios average 1.4, the C& tricyclic/GA tetracyclic terpane ratios average 1.8, the C;, to q, extended hopanes decline in relative concentration with increasing carbon numbers, and the integral of the m/z 191 trace for the C,, to C,, hopanes is greater than the integral of the m/z 217 trace for the C7to G, steranes. This difference in concentration of hopanes relative to steranes may suggest extensive reworking of the algal and terrigenous OM by anaerobic bacteria. In Sicily, the Upper Miocene contains a marine evaporite sequence which includes

TIME -

Fig. 4.6. Gas chromatogram of the saturate hydrocarbon fraction from sample of Paradox shale (3847.7-49 ft depth), GD-I Well, Paradox Basin, Utah. For compound assignments, see Appendix 4.1.

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31 I

A TIME

'6 '[

TIME,

Fig. 4.7. (A) Mass chromatogram of m/z 21 7 in the steroid region of the 7'1 C-trace from u sample of Paradox shale, (3847.7-49 ft depth). GD- I Well, Paradox shale, Paradox Basin. Utah. (B) Mass chromatogram of m/z 191 in the hopanoid region ofthe TIC-trace of the GD-I Well, 3847.7-4Yft depth, Paradox shale, Paradox Basin, Utah. For compound assignments, see appendix 4. I. carbonates, anhydrites or gypsum, halite and potash. Some of the carbonates or marlstones in this sequence have TOC values as high as 13.1 % and are considered petroleum source rocks. (Palmer and Zumberge, 1981). The organic geochemistry of these evaporitic carbonates, as well as oils from seeps in the sequence, were studied in detail by Palmer and Zumberge (1981). They found that kerogen concentrates from these rocks were predominantly pale-yellow amorphous material which they considered to be thermally immature. Two of their samples (E-13 and E-14) contained abundant vitrinite, had a high siliciclastic content, and probably represent a fluvial contribution to the evaporite basin. The kerogens had H/C ratios ranging from 1.1 to 1.68 and were considered by Palmer and Zumberge (1981) to fall between a type I and type I1 kerogen. Several of

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the Miocene carbonates had HC/TOC ratios of over 20%, which Palmer and Zumberge (1981) acknowledge may represent migrated hydrocarbons (staining). However, they also offered an alternative explanation that the large amount of soluble organic matter present in these samples may be due to conversion of insoluble kerogen into soluble kerogen by nonthermogenic processes. Gas chromatograms of C,, + saturated hydrocarbons from their rock extracts showed phytane strongly dominant over pristane in most samples. Secondary in abundance to phytane was an unknown component eluting between n-G, and n-C& in chromatograms of the saturated hydrocarbon extracts. In a recent study of these same rocks, Sinninghe Damstk et al(1986) identified the unknown compound of Palmer and Zumberge as an isoprenoid thiophene. This compound is one of a series of n-alkyl and isoprenoid thiophenes identified by Sinninghe Damsti: et al (1986). These organic sulfur compounds (OSC) were suggested to represent early diagenetic incorporation of OSC into bacterial and algal lipids in a hypersaline environment (Sinninghe Damste et al, 1986). Many of the Sicilian samples contained terpane, sterane and n-alkane distributions indicative of immature OM (Palmer and Zumberge (1981). However, the vitrinite reflectance data on these samples is somewhat contradictory since they show R, values ranging between 0.74-1.01%. Analyses of the sterane composition of the rock extracts by GC-MS showed a dominant cholestane (G,) peak (Palmer and Zumberge, 1981). This suggests that the most likely source of the OM in the Sicilian evaporites is algae and cyanobacteria. The extended pentacyclic terpanes of the samples indicate bacterial sources. Gardner and Bray (1984) considered that the Silurian A-1 carbonates and the "Brown Niagdrdn" carbonate in the Michigan Basin are products of a vitasaline (mesosaline) environment and are the principal source rocks for the oil in Niagaran reefs. It is noteworthy, however, that they found the A-1 carbonate averaged only 0.28% TOC and the "Brown Niagaran", on the flanks and in the upper parts of the reefs, averaged only 0.27% TOC. Biomarker components of the A-1 and "Brown Niagaran" carbonates have been examined by Anders (work in progress). These extracts had low pristane/phytane (Pr/Ph) ratios and high relative amounts of n-alkanes. Figure 4.8

1 3 I I41

TIME' Fig. 4.8. Gas chromatogram o,f the saturate hydrocarbon fraction of a Brown Niagaran samplefrom the Michigan Basin, Michigan. For compound assignments, see Appendix 4.1.

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shows a gas chromatogram of the saturated HCs for the "Brown Niagaran." Similar to most carbonate source rocks, the n-alkane distributions (C,,,-q5region) in the Salina A-1 and "Brown Niagaran" of the Michigan Basin show a slight even carbon preference (Figure 4.8). Like the basinal carbonate facies of the lower Sunniland Formation of south Florida (Palacas, 1984), the relative concentrations of cyclic biomarkers in the Salina A-1 and "Brown Niagaran" carbonate facies are: pentacyclics > tricyclics > steranes (Figure 4.9). The G, to C,, steranes are dominated by the C& steranes. In the CI9to C& region of the m/z 191 mass chromatograms, the G3tricyclic and C, tetracyclic terpanes are the dominate peaks and in the extended hopane region (C,, to G5), gammacerane is the dominate peak. Some average valuesfor geochemical characteristics of marine evaporite carbonate rock OM are given in Table 4.14.

An hydr i t es

Anhydrites are nearly impermeable dense rocks that are usually barriers to petroleum migration. Exceptions to this sometimes occur due to faulting and fracturing or to textures developed where anhydrite commingles with other minerals. It is also possible that under special circumstances the development of biogenic calcite bodies within anhydrite deposits by bacterial reduction of sulfate or methanogenisis (see a following discussion in the evaporites and reservoirs section) could release hydrocarbons from anhydrite rocks. However, even if some anhydrites are rich in OM, it seems unlikely in most cases that significant migration would accompany hydrocarbon generation.

t 18 t t v) 2 +Lu z

'p 13.14

Fig. 4.9. Mass chromatograms of m/z 191 and m/z 217 in the terpanoid and steroid regions of the TI C-traceof a Brown Niagaran sample f rom the Michigan Basin, Michigan. For compound assignments, see Appendix 4.1.

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Table 4.14. Geochemical characteristics of bitumen extracts from evaporite rocks.

Clemical Carbonate Carbonate Halite/ Parameter Marine' Lacustrine' Anhydrite.' Potash4 Trod

maturity range immat.-mat. immat.-mat. mature mature

pristane/phytane 0.6-1 .5 0.1-2.0 0.3-0.6 0.9-1.1

pristane/n-C,, 0.1-1.6 0.5-2.0 0.1-0.7 0.8-1.6

phytane/n-C,, 0.2-2.5 0.4- 16.5 0.2-1.25 1.2-1.7

(:PI (C,-C,) il >I

non-HC/HC 0.1 -1.2 < 6 0.5-2.0 0.15.25

naphthencs mod.-abund." abund. abund. mod.-abund.

hopancs/steranes h>s h~~\ h>\~~

~~steranes c,> c2q c,i cz0 ('272 C2q ND~~

~~C;, diasteranes/ C, reg. steranes < 0.1-0.4 <0.1 0.1-2.0 0.5-0.9~~

~~tricyclic terpane maximum (123 C2".C*, ('23 c,~~

~~C, tetracyclic/ C2,tricyclics c'?,> c;, c*4< c*, C24 > C2h CQ4> cz~~

~~extended hopanes abund. min. ahund. mod. (distribution) Ckci4 < c;, c, > C;, G < Gd> G, G, > (14 > G,~~

~~MDBT v-pattern ND v-paltern ND~~

~~gammacerane/ hopane 0.1-0.4 0.2 < 11.1 -0.5 < 0.2~~

~~p -carotane min. mtd.-abund.' min. min.~~

I. Includes rocks from S. Florida Basin (Palacas, el al. 10x4); Michigan Basin and Paradox Basin, Utah. 2. Includes rocks from Uinta Basin, Utah: Piceance Basin, Colo.: and Jianghan Basin. CKna (Philp and Zharan, 1987). 3. Includcs rocks from Permian Basin, Texas; Paradox Basin. Utah; S. Florida Basin (Palacas et al, 1984) and anhydrites from Guatemala (Connan et al, IYXS). 4. includes rocks from Paradox Basin. Utah and Khorat Basin, Thailand. 5. Green River Basin, Wyoming. 6. Symbols: > = greater than; < = less than; min. = minimal, mod. = mcderate. abund. = ahundant, oba. = obsetved: ND = not determined; MDBT = C4-,C;+G-,C,- methyldiben7othiophcnes hopancs/steranes = x&-qqhopanes (m/i 191)/Cg-C& (m/z 217); CPI = carbon preference index; He = hydrocarbon; S = sulphur. 7. p-carotane represcnts as little as one percent to as much as 25 percent d the saturated hydrtmrhon fraction in Green River Fm. mark and carbonate mudstone.

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If we can accept that most anhydrite deposits are the result of dewatering of precursor gypsum, then it is possible that hydrocarbons or bitumen could leave the deposits during diagenesis (Warren, 1986). Since gypsum is 21% water, an enormous amount of water plus bitumen or hydrocarbons should migrate upward as the critical thickness of overburden and temperature is reached. Timing of the fluid loss would determine whether mobile OM is squeezed out before the generation of hydrocarbons or whether expulsion occurs before existence of any overburden capable of sealing in the fluids. Although organic geochemical data are limited for anhydrites, what data are available show that these rocks are rather lean in TOC. Connan et a1 (1986) found that anhydrites from a sabka facies in Guatemala averaged less than 0.08% TOC and had EOM lower than 400 ppm. Somewhat higher TOC values were reported by Palmer and Zumberge (1981) for Messinian Sicilian anhydrites and gypsums. Average for these rocks was 0.27% TOC with one anhydrite reaching 0.55% TOC. In the South Florida Basin, Palacas et al (1984) analyzed five anhydrites and found TOC ranging from 0.05 to 0.1 1%. In addition to these, another sample from a thick downdip basinal anhydrite had an unusually high TOC of 1.14%. Fifteen anhydrite and halite samples from the Michigan Basin were also characterized by a low average (0.04%) TOC (Gardner and Bray, 1984). In our studies, we analyzed laminated anhydrites from the Castile Formation of New Mexico and the Paradox Member of Utah and found an average ‘TOC for the Castile of 0.10% and 0.15% for the Paradox. EOM from these anhydrites was consistently very low (<300 ppm). For other geochemical properties of anhydrites, only the studies by Connan et a1 (1986), Connan (1981), and ten Haven et a1 (1985) are relatively complete. Connan’s work on Messinian and Cretaceous samples, the studies on Messinian by ten Haven et a1 (1985), plus our own data obtained on Permian and Pennsylvanian samples (Table 4.14), consistantly showed low pristane/phytane ratios (0.32-0.88). According to ten Haven et a1 (1985), this may be used as an indicator of hypersalinity as well as anoxic conditions. Normal alkane distributions in the Castile EOM show a strong even carbon preference (Figure 4.10). Relative concentrations of cyclic biomarkers based on m/z 191 and 217 ion intensities are: pentacyclics > steranes > tricyclics. The C,, tetracyclic is the dominate peak in the C,, to G6tricyclic region of the m/z 191 mass fragmento- gram (Figure 4.1 1A). The overall distribution of tri- and pentacyclic biomarkers in the EOM from Castile anhydrite is nearly identical to their distributions in the anhydrite facies of the Upper Sunniland (compare Figures 4.11A and 4.12) and the C;, steranes are the dominate steranes (Figure 4.11B).

~~Halite and Potash~~

Our knowledge of the organic geochemistry of evaporite rocks decreases with increasing salinity of the environment that produced them. Although halite + potash rocks are more important volumetrically than the evaporitic carbonates and the anhydrites, very few studies have been made regarding their organic geochemistry. This

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~~TIME 6 I~~

~~Fig. 4.10. Gas chromatogram of' the saturate hydrocarbon f'raction oj a laminated anhvdritesample f rom the Castile Formation, Permian Basin, New Mexico. For compound assignments, see Appendix 4.1.~~

is not particularly surprising since they are generally considered as unlikely petroleum source rocks. A notable exception to this has been made by Szatmari (1980) who suggests most of the worlds largest oil deposits could have been derived from evaporitic source rocks, especially halites. Szatmari's theory calls for release of hydrocarbons from organic rich halite deposits as the result of dissolution at both the top and base of the deposit as well as from "shattered" halite. The hydrocarbons and associated connate brines are sealed in the inclusions in the halite rocks and are overpressured because the plastic flow of salt allows direct lithostatic pressuring rather than hydrostatic. Due to this overpressuring, the fluids, when released by dissolution of the hosting halite, can migrate upward or downward into reservoirs of lower (hydrostatic) pressures. A test of Szatmari's theory requires data on the amounts of TOC or EOM in halite deposits; and as we have previously stated, such data is meager. Despite this, we feel that some of the data offered here suggests a problem with Szatmari's theory. Our principal source of organic geochemical data comes from halite samples from the Paradox Member in southeastern Utah. We have carefully analyzed 147 samples of halite ranging in color from light gray, pale yellow, and dark gray and have found 70 of these had TOCvalues below our limits of detection (0.001 %). The other 77 samples had TOC contents ranging from 0.001 to 0.28%. In addition to these samples, we examined halite from the A-1 Salt in the Michigan Basin, and halite from Cretaceous evaporites in Thailand, and found similar low values. Treesh and Friedman (1974) reported TOC values ranging from 0.44 to 2.14% for Silurian halite in the Michigan Basin, but we suggest that these high values are not representative of these halite deposits as a whole and may include other rock types. Furthermore, the EOM of many samples was so low that it was difficult to get enough material for chemical characterization. In addition, two potash samples from a potash deposit in the Paradox Member were analyzed and found to have TOC and EOM contents similar to the halites. Previously, a sample from this

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~~TIME~~

~~TIME -c~~

Fig. 4.11. (A) Mass chromatogram of m/z 191 in the terpanoid region of the TIC-trace o,f a laminated anhydrite sample from the Castile Formation, Permian Basin, New Mexico. (B) Mass chromatogram of m/z 217 in the steroid region of the TIC-trace of a laminated anhydritesample f rom the Castile Formation, Permian Basin, New Mexico. For compound assignments, see Appendix 4.1. potash deposit had been reported to contain 3,000 ppm EOM (Peterson and Hite, 1969), but we could not duplicate that high value. Liquid inclusions of hydrocarbons are occasionally found in halite deposits (Figure 4.13), but they are quantitatively unimpor-

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~~PENTACYCLICS I I 19~~

~~>t t v) TRlCYCLlCS z :w w5 U~~

~~M/z 191~~

TIME - Fig. 4.12. Mass chromatogram of m/z 191 in the terpanoid region of the TIC-trace of a core sample ,from an upper Sunniland anhydrite f acies, South Florida Basin, Florida (Palacas, et al, 1985). For compound assignments, see Appendix 4.1.

tant. If we can assume that values for TOC and EOM presented here are representative for halite deposits that are generally considered organic-rich (Paradox, Zechstein, Salina, etc.), then it is difficult to see how the dissolution of halite could be an important source of petroleum (Szatmari, 1980). For example, the dissolution of a 10 m thickness of halite with an EOM content of 0.001% over a km2 area could potentially release about 217 tons of petroleum. Considering the volume of brine generated by the dissolution of this amount of halite, the dilution factor would lower the EOM content of the brine to 0.00033%. Greater volumes of halite than used in the previous exercise might be expected to dissolve where salt tectonics, especially diapirism, is involved. Despite this, the

~~Fig. 4.13. Liquid petroleum inclusions in halite rock from Paradox Member, Paradox Basin, Utah.~~

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tremendous dilution of EOM from the halite by the brine generated by its release, in our opinion, suggests that halites should still be considered as unlikely source rocks. Combustible mixtures of gases causing explosions and fires have been encountered in salt and potash mines of Europe (Sonnenfeld, 1984). The gases of halite and potash deposits most commonly occur in micro-inclusions and consist predominantly of nitrogen with minor amounts of methane, higher alkanes, hydrogen, helium, neon, and argon (Bol'shakov, 1972). In the Solikamsk Basin of the USSR, the greatest concentration of occluded gases (210 ml/kg) was found in milky-white sylvites, whereas carnallites contained <21 ml/kg (Bol'shakov, 1972). Similar relationships can be seen in strongly recrystallized white sylvites from Thailand and Utah. In general, halite rock contains less gas than potash; and in particular, the gas has a smaller percentage of hydrocarbons (Table 4.15). The micro-inclusions of gas in halite and potash rocksfrequently have high internal pressures, and these inclusions have the tendency to break out with explosive force when confining walls are weakened mechanically or by dissolution. The higher concentration of OM in evaporite brines of supersaline salinities suggests that a higher HC occlusion content, as well as higher internal pressures in the occlusions, should be expected in potash deposits. The high internal pressures could simply be the result of maturation of OM in the original fluid filled pores of a rock which allows little or no loss of products or pressure. We have collected and analyzed inclusions of hydrocarbons from some of the Paradox halites. These inclusions (Figure 4.13), which are predominantly light hydrocarbons (C,-C,J, were analyzed by dissolving the halite in water under vacuum and passing the released hydrocarbons directly into a GC. The distribution of n-alkanes in these samples shows a strong odd number carbon preference (Figure 4.14). Using the

~~ccGAS/kg SALT~~

~~Sample # Paradox Member Depth Total N, CH4 C02 O2 Salt Bed (feet)~~

GD 1-100 6 3123.5 30.7 0.25 0.15 28.2 T GD 1-101 6 3231.9 13.9 0.52 0.48 74.5 027 GD 1-lUZ 6 3m.7 142.5 1.10 027 U5.9 0.84 GD 1-103 6 3m.4 39.1 0.71 0.48 37.5 0.23 GD 1-104 6 3321.6 4.5 0.13 0.14 4.3 GD 1-105 6 3340.8 19.9 0.40 0.77 192 0.12 GD 1-106 7 3381.6 9.1 0.40 0.10 7.4 0.05 OD 1-107 7 3421.9 582 1.04 0.64 56.1 0.11 GD 1-10 10 3721.5 57.9 1.04 0.16 58.5 0.18 GD 1-109 13 3914.8 192.0 1.60 1.14 187.0 0.46 GD 1-110 18 4417.7 34.0 0.55 0.24 30.0 GD 1-111 18 4143.6 23.0 0.30 0.06 21.0 GD 1-112 21 5090.7 32 0.32 T 3.4 T GD 1-115 26 5504.7 14.9 0.04 8.40 15.6

~~Table 4.15. Occluded gas in halite rock, Gibson Dome No. I Corehole, San Juan County, Utah (Hite, 1983). Salt bed nomenclature from Hite, 1960.~~

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~~20 1 80~~

~~5 10 15 20 CARBON ATOMS PER MOLECULE~~

Fig. 4.14. Distribution of low molecular weight n-alkanes in petroleum inclusions in halite rock from the Paradox Basin (solid line) plotted against the distribution of C,-C,, fatty acids from 300%0 salinity sediments, Alviso salt pond, San Francisco Bay, California (dotted 1ine) .

data of Klug et a1 (1985), we have plotted the distribution of C,-C, fatty acids from the 3Wh0salinity sediments of an Alviso salt pond (Table 4.6) against the Paradox inclusions (Figure 4.14). It can be noted that there is a strong correlation between the amounts of even carbon acids and odd carbon hydrocarbons. We suggest that this correlation may be the result of thermal decarboxylation of dissolved fatty acids in primary brine trapped as inclusions in the Paradox halite. Hydrogen is a common gas in halite and potash deposits; and if we consider observations made in modern evaporite environments, this is predictable. Lupton et a1 (1984) found that increasing salinities in sediments of the Great Salt Lake brought about increased production of hydrogen by fermentative bacteria; and at the same time, consumption of hydrogen by sulfate reducing and methanogenic bacteria decreased. In the north arm of the lake at a salinity of 290%0 the sediments contained as much as 300 pMof dissolved H, and only 0.4pM of dissolved Cw.Applying these findings to ancient evaporite deposits, we should expect highest H, contents in potash deposits and associated CH, may be largely derived from thermal decarboxylation of acetate rather than being a metabolic product of methanogenic bacteria. The organic geochemistry of halites has received little attention, and this is

Data Center ,09126599985,[email protected], For Educational Uses HALITE AND POTASH 38 1 perhaps unfortunate because these rocks have hermetically sealed in OM and associated HC that wds generated in situ. Furthermore, the OM has only been in contact with NaCl so that catalytic effects of clay minerals, etc. have been prevented. Considering what we know about the modern evaporite environment, it is safe to assume that all of the OM is either algal, bacterial, or a combination of the two. Small terrestrial contributions by wind blown spores or pollen and even an occasional larger wood fragments are known, but these are in most cases quantitatively unimportant. Thus, halites afford the opportunity to precisely study the genetic and diagenetic histories and mass balance between the OM and its derived hydrocarbons. Normal alkanes in the Paradox halites have an even carbon preference (Figure 4.15) the relative intensities of hopanes > steranes, the C& tricyclic is the dominate tricyclic and the C, tetracyclic is present in moderate abundance. It is commonly accepted that anoxic and hypersaline conditions are indicated when pristane/phytane in oils and rock extracts are <1 (ten Haven, et al, 1985). Pristane/phytane ratios for OM associated with halites and potash in this report range from 0.9 to 1.1. Other biomarker values for these rocks may be seen in Table 4.14. Halite deposits also offer an unusual opportunity to study maturation of woody plant material in a totally sealed anoxic system. Normally, one does not associate evaporites with coal; but on occasion, woody plant remains are delivered to evaporite basins in small quantities. Dzhinoridze et a1 (1974) have described coalified land plants collected from potash-bearing rocks of Moisten age in the Ciscarphathian region. This material consisted entirely of vitrinite with a reflectance (R,,) of 0.6 to 0.65%. We have examined a small (3.0 cm diameter) fragment of coalified plant material that was collected from a potash mine in the Paradox Member in Utah. This sample was totally encapsulated in halite. At the contact between coal and halite, the halite was discolored yellow-brown by maturation products expelled from the original wood. The coal was predominately vitrinite with an R,, of 0.81%. Elemental analyses of the coal gave the following results: S = 0.17%, H/C = 0.65, C/O = 8.65, atomic (H/C = 0.78, O/C =

~~4 21~~

~~TIME -~~

~~Fig. 4.15. Gas chromatogram of the saturate hydrocarbon fraction from halite rock, Texas-Gulf Cane Creek potash mine, Paradox Basin, Utah. For compound assignments, see Appendix 4. I.~~

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0.049). Rock-Eva1 pyrolysis of this coal gave 9,100 ppm for S, and 97,000 ppm for S,. Extractable OM relative to TOC amounted to 9.2 mg/g. Gas chromatographic results (Figure 4.16) are: pristane/phytane 0.91, pristane/n-C,, 0.33, phytane/n-C,, 0.43, n-alkane maxima at C,, and C,, in the n-C,, to n-C,, region, no odd or even preference in the higher homologs, and except for n-C,, and n-C,,, and the n-alkanes decrease in smooth stairstep fashion from n-C,, to n-G. Hopane was the only terpane biomarker present in sufficient concentrations to be detected by GC/MS. The Cretaceous halite deposits of northeast Thailand were also sampled from a cored interval consisting of halite containing abundant macerated woody plant remains. This coaly material had an value of 0.74% which places it at a similar maturation level as the previously described Paradox sample. The geochemical characteristics of the OM in the Thailand core were also very similar. In this sample, the pristane/phytane was 1.05, pristane/n-G, 0.34, phytane/n-C,, 0.41, there was no odd or even predominance in the n-alkanes, a high abundance of medium range n-alkanes (C,5-C20)relative to higher homolog (>Go)and smooth stairstep decline in relative abundance of n-alkanes from n-C,, to n-G, (Figure 4.17). Biomarker concentrations were so low that only hopane was detected.

~~Lacustrine Evaporites~~

The various saline facies of the Eocene Green River Formation of Colorado, Utah, and Wyoming typify the organic-rich rocks produced from a lacustrine evaporite environment. In three sub-basins (Green River, Piceance Creek, and Uinta) kerogenous marlstones (oil shales) are associated with sodium carbonate minerals (trona, nahcolite, etc.) and halite (Dyni, 1981). A vertical succession of 1) oil shale, 2) trona or nahcolite, and 3) dolomitic marlstone is repeated many times especially in the Green River and

Fig. 4.1 6. Gas chromatogram of the saturate hydrocarbon fraction from a halite encased coalif ied wood fragment, Texas-Gulf Cane Creekpotash mine, Paradox Basin, Utah. For compound assignments, see Appendix 4.1.

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~~TIME-~~

~~Fig. 4.17. Gas chromatogram of' the saturate hydrocarbon fraction from halite encased codified woody plant debris, Khorat Basin, northeast Thailand.~~

Piceance Creek Basins. It is evident that it was during periods of highest salinities that maximum accumulation of kerogen occurred. The OM in these rocks is hydrogen-rich type I kerogen consisting largely of algal and bacterial remains (Tissot and Welte, 1984). The association of sodium carbonate minerals such as trona and/or nahcolite with kerogenous marlstones can be explained as the result of large-scale bacterial reduction of the sulfate ion in sodium rich lake brines which enriches the brine in bicarbonate and eventually sodium carbonate minerals are precipitated (Dyni, 1981). Prerequisites for this process are eutrophic lakes where organic productivity is so high that there is more than enough OM to satisfy the needs of the anaerobes and or salinities are so high that activities of the anaerobes are limited, and a surplus of OM is left behind which on burial becomes kerogen. The most organic-rich facies of the Green River Formation is present in the Piceance Creek Basin where TOC of select intervals is over 35%, and these rocks are considered commercial oil shales. In the Uinta Basin the marlstones have lower TOC values but have been buried more deeply and have generated significant amounts of petroleum. Because of these petroleum accumulations, the geochemistry of the Uinta Basin rocks has been studied in detail by many workers. Some of the richest rocks in this sequence have TOC contents of over 21% (Anders and Gerrild, 1984). In the basin center, the lacustrine sequence reflects a progressive increase in salinity from base to top. Similarly, the organic geochemistry of the sequence shows progressive changes (Robinson and Cook, 1975; Anders and Robinson, 1973). A study by Anders (in progress) of biomarker compounds in the immature (R., < 0.5%) lacustrine carbonate rich marls and mudstones of the Upper Green River Formation shows many dissim- ilarities to the marine carbonate deposits previously discussed (Table 4.14). Through a 305 m thick interval of the Upper Green River Formation in the southeastern Uinta

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Basin, the pristane/phytane ratios average 0.7, G7/G9sterane ratios average 0.41, odd-carbon n-alkanes predominate over evencarbon n-alkanes in the n-C,, to n-C& region, extended pentacyclic hopanes (GIto C&) decrease rapidly in concentration with increasing carbon number, and gammacerane and p -carotane are abundantly present. Some distinguishing characteristics of the immature Uinta Basin Green River basinal facies are (Figures 4.18 and 4.19): 1) n-C,, is the principal normal paraffin in the n-alkane fraction (averaging 20% of total n-alkanes) suggesting strong algal and bacterial input; 2) perhydro- p carotane is the predominate alkane in the branched and cyclic alkane fraction averaging 23% of the fraction; and 3) in the tricyclic region, the Goand GItricyclics dominate over the G3tricyclic and G4tetracyclic. Organic carbon content in the basinal facies averages 6.3 wt. %, carbonate carbon averages 5.6 wt. %, and the average EOM/TOC ratio is 27.3%. The ratio of n-alkanes/branched and cyclic alkanes/aromatic HC/resins and asphaltenes is 4/26/3/67, respectfully. In stratigraphically lower rocks there is a decrease in relative amounts of pristane and phytane, an increase in waxy n-alkanes, and an increase in saturated HCs (particularly the higher molecular weight n-alkanes) with respect to non-HCs. These changes were tentatively explained by Tissot et a1 (1978) as due to increased bacterial reworking rather than thermal maturation. The degree of degradation of the OM may relate to salinity change because part of the changes may have occurred under mildly oxidizing conditions (Powell, 1986). With progressive increase of salinity in the Uinta Lake, oxygen levels would have become lower so that aerobic bacterial degradation would also have pro-

~~I 5~~

~~w a UJ U~~

Fig. 4.18. Gas chromatogram of saturate hydrocarbon fraction from thermally immature Upper Green River Formation carbonate-rich marl, Uinta Basin, Utah (obtained on a 6 x 0.25"SE-3Opacked column). Insert isgas chromatogram of mole seivefraction of same sample. For compound assignments, see Appendix 4.1.

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~~TIME-~~

Fig. 4.19. Mass chromatograms of m/z I91 and m/z 217 in the terpenoid and steroid regions of the TIC-trace of an immature Upper Green River Formation carbonate rich marl, Uinta Basin, Utah. For compound assignments, see Appendix 4. I.

gressively decreased. That the salinity of the ancestral Uinta Lake did progressively increase is verified by the presence of evaporite minerals (sodium carbonates). Furthermore, the higher salinities, and a corresponding increase in pH would have favored and increased production of OM from cyanobacteria, algae, and halophilic bacteria. Associated with the Green River rocks of the Uinta subbasin are vertical veins or dikes of solid bitumen such as gilsonite. The composition of gilsonite is quite similar to rock extracts from kerogen-rich rocks in the upper part of the Green River sequence (Figure 4.20). The steranes (C&-G9)are particularly abundant with a C& dominance. Also abundant are the C, pentacyclics and pcarotane. The origin of these veins has never been completely resolved, but we have some new data that may shed some light on the problem. In the Eocene rocks of the Green River Basin of Wyoming large deposits of trona (Na2(C0,).Na(HC0,).2H20) occur associated with kerogenaceous marlstones. Also found in this sequence is a sodium carbonate brine which locally may contain in solution as much as 10 wt. % organic acids. The organic acids can be precipitated from the alkaline brine by mineral acids or divalent ions. The precipitated organic material has an elemental composition that is very similar to kerogen in the oil shales of the Eocene sequences. A comparison of the gas chromatograms of the saturated hydrocarbons of the trona brine precipitate (Figure 4.21), gilsonite (Figure 4.20), and rock extracts from the Mahogany Ledge oil shale (Figure 4.18) from the Uinta subbasin show striking similaritiessuggesting a common genetic origin. We suggest that the black trona brine of Wyoming represents fossil brine from the original lake which produced trona and also oil shales. Brines of similar composition were probably

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~~TIME-~~

~~Fig. 4.20. Gas chromatogram of the saturate hydrocarbon fraction from gilsonitc, Uintu Basin, Utah. For compound assignments, see Appendix 4.1.~~

present during the saline stages of all of the Green River basins and it seems likely that at least part of the abundant OM produced at that time would be solubilized in the highly alkaline lake brines. With an influx of calcium and magnesium ions via seasonal runoff, we suggest it would be possible to precipitate organic acids in the lakes. The end result would be the typical laminated oil shale (dolomite and kerogen) where a component of the kerogen is an organic acid precipitate. The gilsonite veins might be explained by expulsion of the still fluid-like precipitates into vertical fractures which were forced open by high flow pressures. Eocene rocks in the Qianjiang Depression of the Jianghan Basin of China were deposited under extremely anoxic, reducing, and hypersaline conditions (Jiang and Zhang, 1982). These rocks and associated oils have been characterized in terms of n-alkanes, isoprenoids, steranes, terpanes, aromatics, and sulfur containing compounds by Jiamo et a1 (1986) and Philp and Zhaoan (1987). Of seven potential source rocks examined by Philp and Zhaoan (1986) only one sample seemed to have a geochemical correlation with the associated oils. This sample had a TOC value of 0.78% and an EOM of 509.0 mg/g OC (Philp and Zhaoan, 1987). The Pr/Ph ratio was 0.68, and the Ph/n-C,, ratio was 3.34. The n-alkanes in this sample had a slight predominance of even-carbon-numbers which according to Tissot and Welte (1984) is characteristic of evaporite environments. The biomarker data of Philp and Zhaoan (1987) shows C, steranes are predominant over G, steranes which they point out could be due to a large input of algal material rather than terrestrial input. They also found various sulfur compounds in the suspected source rock which were predominantly thiophenes, benzo- thiophenes, dibenzothiophenes and their alkylated analogues. Additional biomarker data from these sediments include establishingthe presence of gammacerane, hopanoids G& excluding C,,, pregnane and homopregnane, steranes G,-C,, excluding G5,and finding a tendency for steranes to dominate hopanes (Fu Jiamo et al, 1986). Fu Jiamo et a1 (1986) suggested that the absence of sterenes, diasterenes and near absence of diasteranes in these sediments may be the result of the lack of any clay minerals to bring

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Fig. 4.21. Gas chromatogram of the saturate hydrocarbon f’raction from a tronu brine, Wilkins Peak Member of the Green River Formation, Green River, Wyoming. For compound assignments, see Appendix 4.1.

about catalytic steroid rearrangements. The Officer Basin of Australia contains a Cambrian age carbonate-evaporite facies which is believed to have been deposited in an alkaline playa lake. These rocks contain TOC values ranging from 1.30 to 0.09% with an average of about 0.50% (McKirdy, et al, 1984). The kerogen of this facies is type I and is hydrogen-rich (H/C ratios 1.19-1.34). The source of this OM was probably lipids from cyanobacteria, halophilic, and methanogenic bacteria. Biomarker geochemistry shows that these rocks and their associated oils have high concentrations of sesterterpanes and squalene (McKirdy et al, 1984).

~~EVAPORITE OILS~~

The original geochemical character of evaporite oils is difficult to determine because, in many cases, they have migrated into non-evaporitic reservoirs where biodegradation may have taken place as well as catalyzed in-reservoir reactions due to contact with a different mineral suite. Furthermore, the migration of oils may bring about changes in biomarker ratios due to differential rates of migration between certain biomarker groups (Philp, 1986). We have attempted to summarize our data plus data of other authors on oils that most likely originated from evaporite sources. Most of this data comes from carbonates deposited in the vitasaline stage. Both lacustrine and marine carbonates provided oils for this data base. Our data on oils derived from anhydrites or a penesaline environment are limited to samples from the Castile Formation of west Texas where oil is associated with limestone reservoirs created in the encapsulating Castile anhydrite by bacterial reduction of sulfate. Some of our samples that may be most representative of evaporite oils come from seeps in potash and salt mines as well as cores from these deposits. These samples are somewhat unique because

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~~they have been hermetically sealed in salt since their origin and, therefore, are not effected by catalytic change or biodegradation.~~

~~Marine Carbonate Oils~~

Oils associated with carbonate-rich rocks from marine environments covering a range of maturities have the following general geochemical characteristics (Table 4.16, Figures 4.22A-C): API gravity, 19-50"; sulfur, 0.2 to 4.0%; pristane/phytane, < 1; pristane/n-C,,, 0.1 to 1.0; phytane/n-C,,, 0.3 to 1.5; CPI (G7-G9),< 1.1 (generally < 1where there is little or no siliciclastic and associated OM input); non-HC/HC, 0.1 to 0.5; hopanes > steranes; abundant ring napthenes relative to n-alkanes; G7steranes relative to G9steranes variable (Fig. 4.22B); minimal to moderate amounts of G,diasteranes; moderate amounts of tricyclic terpanes (maxima G,); moderate amount of C, tetracyclic terpane; abundant extended hopanes (GI to C&); v-shaped pattern for the distribution of the methyl dibenzothiophenes; minimal to moderate amounts of gammacerane (C,pentacyclic triterpane); and little or no f3-carotane.

~~Anhydrite Oils~~

The geochemical properties of oil associated with anhydrite rocks is based on a very limited number of oil stained cores from the Permian Basin. Generally, the various HC ratios are in the range for those calculated for oils associated with carbonate rocks (Table 4.16, Figure 4.23A-C): pristane/phytane, < 1; pristane/n-C,,, < 1; phytane/n-C,,, > 1; CPI, < 1; nonHC/HC 0.3; hopanes > steranes; abundant ring napthenes relative to normals; G7steranes sometimes < and sometimes > G9 steranes; minimal tricyclic terpanes (G, maxima); moderate amount of C, tetracyclic terpane, abundant hopanes (C2,-G5);minimal to moderate amount of gdmmacerane; and little or no (3-carotane. We are hesitant to claim that all of these oils are truly indigenous to an anhydrite facies (Castile Formation) because there is some geologic evidence that they may have migrated via faults and fractures from the underlying Bell Canyon Formation. However, the gas chromatogram of our sample does not resemble the Bell Canyon analyzed by Palacas.

~~Halite-Potash Oils~~

We have based the geochemical characteristics of halite-potash oils on samples from seeps in potash mines in the Paradox basin (Pennsylvanian) Utah, and Permian Basin (Permian) New Mexico, and a seep in the Pugwash salt mine (Mississippian) New Brunswick. Additional samples from the Paradox basin included halite with oil inclusions and oil-stained halite cores. The characteristics of these oils are summarized as follows: API -36"; Sulfur <0.1%; 6%, -18.61%0; 6l3Csat., -30300;6I3C arom. -29.@/00;pristane/phytane, 1.0-1.6; pristane/n-C,,, 0.4-0.5; phytane/n-C,,, 0.4-0.6; CPI (G7-G9)= 1; non- HC/HC, 0.2-0.3;n-alkane maxima at q5and q7;hopanes (G7-G)>steranes (%-G9);

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~~Table 4.16. Geochemical characteristics of evaporite oils.~~

~~Chemical Carbonate Carbonate Anhydrite Halite-Potash Parameter Marine' Lacustrine* ~arine~ Marine3~~

~~maturity mod.-mat. mod. mod. mod.-mat.~~

~~API 19-50 22-32 ND 36~~

~~S% 024.0~~

~~pristane/ phytane~~

~~pristane/ n-Ct7 0.1-1.0 >1 <1 0.4-05~~

~~Phytand n-c,, 0.3-1.5 >1 >1 0.4-0.6~~

~~CPI (GG9) < 1.1 > 0.9 <1 1 non-HC/HC 0.1-0.5 02-1.7 0.3 02-0.3~~

~~naphthenes abundant abundant abundant minimal~~

~~hopanes/ h>s h~~s h>s steranes~~~~

~~~~steranes GVSCZ9 GSG9 cnG9 variable C, diasteranes/ C, reg. steranes 02-0.4 minimal 0.1-0.3 0.50.9~~~~

~~~~tricyclic maximum G? G?Gl c, c,~~~~

~~~~C,tetracyclic/ c, tricycfics G>G6 G6'G c24>G6 G>G6~~~~

~~~~extended abundant minimal abundant moderate hopanes G3&4 c,> GS G3~~~~

~~~~MDBT v-pattern stairstep ND ND pattern Gammacerane/ hope 0.1-0.4 02-0.4 024.3 0.1~~~~

~~~~p-carotane minimal abundant minimal minimal~~~~

(1). Based on oils from S. Florida Basin, Michigan Basin and Paradox Basin, Utah. (2). Based on oils from Uinta Basin, Green River Formation, Utah, and Jianghan Basin, China. (3). Based on oil stained cores, mine seeps, or petroleum inclusions in anhydrite, halite, and potash from Paradox, Permian, and Windsor Basins. (4). Symbols: > = greater than, < = less than; mod. = moderate; mat. = mature; ND = not determined; MDBT = C,-,G+G-,C,-methyldikmthiophenes; S = sulphur; he panes/steranes = CC&-GS(M/Z 19l)/cG-C, (M/Z 217); HC = hydrocarbon.

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~~~~/jl TIME-~~~~

~~~~TIME-~~~~

~~~~Silurian NiagaranCarbonate,~~~~

TIME- Fig. 4.22. (A) Gas chromatogram of the saturate hydrocarbon fraction from a Silurian oil (A-2 carbonate), Dorr Field, Michigan Basin, Michigan. (B) Mass chromatogram of' m/z 191 in the terpanoid region of the TIC-trace from a Silurian oil (A-2 carbonate), Dorr Field, Michigan Basin, Michigan. (C) Mass chromutogram of m/z 21 7 in the steroid region of the TIC-tracefrom a Silurian oil (A-2 carbonate), Dorr Field, Michigan Basin, Michigan. Top insert, mass chromatogram of m/z 217 in the steroid region of' the TIC-tracefrom a Silurian oil (A-1carbonate), Hessen Field, Michigan Basin, Michigan. Lower insert, mass chromatogram of m/z 217 in the steroid region of the TIC-trace form a Silurian oil (Niagaran carbonate), Marine City Field, Michigan Basin, Michigun. For compound assignments, see Appendix 4.1.

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~~~~TIME-~~~~

~~~~TIME -w C 36 t~~~~

~~~~TIME -w~~~~

Fig. 4.23. (A) Gas chromatogram of the saturate hydrocarbon fraction from an oil-stained anhydrite core, Castile Formation, Permian Basin, New Mexico. (B) Mass chromatogram of m/z 191 in the hopanoid region of the TIC-trace (ion trap) from an oil-stained anhydrite core, Castile Formation, Permian Basin, New Mexico. (C) Mass chromatogram of m/z 217 in the steroid region of the TIC-trace (ion trap) from an oil-stained anhydrite core, Castile Formation, Permian Basin, New Mexico. For compound assignments, see Appendix 4.1. minimal ring napthenes relative to n-alkanes; G7steranes > G9steranes; moderate to abundant G7 diasterdnes; moderate amounts tricyclic terpanes (maximum G3); moderate amounts C, tetracyclic terpane; moderate to abundant extended hopanes (C&-C3& minimal to moderate amounts of gammacerane (C,pentacyclic triterpane); and little or no p-carotane.

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In spite of the differences in geologic age and location of the halite-potash oils used for the above geochemical profile, striking similarities are observed. Both have nearly identical hydrocarbon distributions (Figures 4.24 and 4.25), both have a dominance of G, steranes over C,, steranes and abundant diasteranes (Figures 4.26B and 4.27B), both have nearly identical tricyclic and tetracyclic distributions and with the exception of the TS/TM ratio, and both have very similar hopane distributions (Figures 4.26A and 4.27A). These chemical similarities suggest that the marine saline to supersaline environment imparts a distinctive geochemical signature to its derived oils and the data presented here can be considered representative of all oils produced from this setting.

~~~~Lacustrine Carbonate Oils~~~~

The chemical features of oils derived from mature lacustrine carbonate rich mark and mudstones are (Table 4.16, Figure 4.28A-B): API gravity 22-32O; sulfur, < 1%; pristane/phytane 0.4 to 1.7; pristane/n-C,, > 1; phytane/n-C,, > 1; CPI (C2,-C2,), 0.9 (generally > 1 because of siliciclastic and associated terrestrial OM input); non-HC/HC 0.2 to 1.7; hopanes < steranes; abundant ring napthenes relative to n-alkanes; G7steranes < C,, steranes; minimal amount of diasteranes relative to regular steranes; moderate to abundant amounts of tricyclic terpanes (maximum is generally C,, or GIbut rarely CZ3); moderate amounts of C2, tetracyclic; minimal amounts of extended hopanes (C31to C&), moderate to abundant amounts of gammacerane (C,pentacyclic triterpane); stairstep

~~~~I I3~~~~

I TIME - Fig. 4.24. Gas chromatogram of' the saturate hvdrocarbon fraction from an oil seep in the Texas-Gulf, Cane Creek potash mine, Paradox Basin, Utah. For compound assign- ment, see Appendix 4. I.

~~~~Data Center ,09126599985,[email protected], For Educational Uses LACUSTRINE CARBONATE OILS 393~~~~

~~~~TIME-~~~~

~~~~Fig. 4.25. Gas chromatogram of the saturate hydrocarbon fraction from an oil seep in halite in the Pugwash salt mine, New Brunswick, Canada. For compound assignments, see Appendix 4.1.~~~~

distribution pattern for the methyl dibenzothiophenes; and abundant amounts of p carotane. Oil derived from lacustrine evaporitic environments is remarkably different than oil from marine evaporites. Some of these observed differences are summarized as follows: pristane/phytane generally <1 in marine, >1 in lacustrine; CPI (G7to (&) generally < 1 in marine, > 1 in lacustrine; non-HC content much lower in marine than in lacustrine; relative abundance of steranes (G,-G9)variable in marine but generally <1 in lacustrine; dominate tricyclic terpanes are C& in marine but C;, or GI in lacustrine; extended hopane (GIto G5)abundant in marine but minimal in lacustrine; p carotane content negligible in marine but abundant in lacustrine.

~~~~PALEOPRODUCTIVITY AND PRESERVATION OF ORGANIC MAlTER IN THE EVAPORITE ENVIRONMENT~~~~

In previous sections, we have given data on rates of organic carbon (OC) productivity measured in modem evaporite environments. Attempts at calculating the paleoproductivity of certain geologic environments have been less than satisfxtory due to insufficient data regarding the effect of oxygen levels on OC degradation and losses due to maturation of OC (Arthur et al, 1987). We suggest that evaporites may be ideally suited for studies of paleo-productivities. Much has been written about the unique preservation parameters of the evaporite environment, such as strong anoxia, limited destruction by anaerobes, etc. Thus, it would seem that a large proportion of the OC produced by the environment should be preserved in the evaporite rocks. The other fortuitous feature of many evaporite rocks is the presence of laminae or banding which we believe represents seasonal depositional cycles. Many geologists disagree with a seasonal interpretation for these features, but the data presented here help support the

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~~~~TIME- TIME -+~~~~

Fig. 4.26. (A) Mass chromatogram of m/z 191 in the terpanoid region of the TIC-trace from an oil seep in the Texas-Gulf, Cane Creek potash mine, Paradox Basin, Utah. (B) Mass chromatogram of m/z 191 in the steroid region of the TIC-trace from an oil seep in the Texas-Gulf, Cane Creek potash mine, Paradox Basin, Utah. For compound assignments, see Appendix 4.1. seasonal concept. Using this built in time scale, we can calculate the yearly amount of OM delivered to and preserved by the environments. In our study, we used four sample sets which included 1) banded halite from the Paradox Member of the Paradox Basin, 2) laminated anhydrite from the Paradox Mem- ber, 3) laminated anhydrite from the Castile Formation of the Permian Basin, and 4) varved or laminated oil shale from the Green River Formation of the Piceance Creek Basin. The latter rock type is a fine-grained kerogenous marlstone which was deposited in a saline lacustrine environment so it qualifies as an evaporite rock.

~~~~A B 31 ii v)a z Lu w~~~~

~~~~TIME- TIME-~~~~

Fig. 4.27. (A) Mass chromatogram of’ m/z 191 in the terpanoid region of the TIC-trace from an oil seep in halite in the Pugwash salt mine, New Brunswick, Canada. (B) Mass chromatogram of m/z 21 7 in the steorid region of the TIC-trace from an oil seep in halite in the Pugwash salt mine, New Brunswick, Canada. For compound assignments, see Appendix 4.1.

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~~~~TIME~~~~

~~~~B 1~~~~

~~~~18 \t 8 rnc z w I- 22 z I w kz 17 i w n: 13 14 MIZ ,x TIME~~~~

Fig. 4.28. (A) Gas chromatograin of thesaturate hydrocarbonfractionj’ronz an oil of the Altamont-Bluebell Fieldderivedfrom carbonate-richrocksojthe Green River Formation, Uinta Basin, Utah. (B) Mass chromatograms of m/z I91 and m/z 217 in the terpanoid and steroid regions of the TIC-trace from an oil of the Altamont-Bluebell Field derived from carbonate-rich rocks of the Green River Formation, Uinta Basin, Utah. For compound assignments, see Appendix 4. I.

Eleven samples of halite rock from the Paradox Member were taken from a single halite bed which was distinctly banded by anhydrite laminae. These samples totaled an aggregate thickness of 31.398 m; and according to numbers of halite-anhydrite couplets present, the interval represents 941 years of deposition. TOC values through this interval ranged from 0.28 to 0.05% and averaged 0.17%. Using a density of 2.168 for the halite rock 1 m3 would contain 3,686 g of OC so that:

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~~~~(3,686 g OC/m3 x 31.398 m (thickness)) + 941 yrs = 123 g OC/m’/yr~~~~

From a suite of 13 samples of laminated anhydrite rock from the Paradox Member, we determined that through a thickness of 0.762 m, the TOC averaged 0.1 12%. On the basis of laminae count, this thickness represents 168years of deposition. Using a density of 2.96 for the anhydrite rock, 1 m3 would contain 3,315 g of OC so that:

~~~~(3,315 g OC/m3 x 0.762 m (thickness)) A 168 yrs. = 15.0 g OC/m’/yr~~~~

Four samples of finely laminated anhydrite rock from a continuous interval of the Castile Formation were analyzed. Analyses for TOC through a thickness of 0.366 m ranged from 0.08 to 0.1 1% and averaged 0.095%. Counting laminae through this interval gave a depositional age of 185 years. Using a density of 2.960, the OC content of 1 m3 of this rock would be 2,812 g so that:

~~~~(2,812 g OC/m3 x 0.366 m (thickness)) + 185 yrs. = 5.6 g OC/mz/yr~~~~

The Green River Formation and its extensively laminated (varved) organic-rich marlstones were studied in great detail by Bradley (1930). By counting varves, he estimated that for oil shale yielding between 15-35 gallons per ton (gpt) of shale, one foot of shale (0.305 m) would accumulate in 4,700 years or 15,410 years/m. Using Bradley’s data, we can make a comparative study with the other samples from marine environments. The density of oil shale yielding 34.7 gpt is 2.08 (Smith, 1976). Gpt is converted to TOC by dividing by 2.14 giving a value of 16.21%. Using these values 1 m3 of this grade of oil shale would contain 337,168 g of OC so that:

~~~~(337,168 g OC/m3 x 0.305 m (thickness)) + 8,200 yrs. = 12.5 g OC/m’/yr~~~~

Even richer oils shales are present in the Green River Formation, so the above figure does not represent maximum accumulation for the ancient lake. For example, using richer oil shale (1.978 density = 42.5 gpt) Bradley (1963) calculated the amount of OC preserved to be 28 g OC/m2/yr for the Green River Formation in Wyoming. A comparison of the numbers arrived at here with the known productivities of modern evaporite environments allows a rough estimate of the preservation efficiency of this unique environment. Using the TOC content of sediments from the Great Salt Lake (Spencer et al, 1984), Eugster (1985) estimated the annual sedimentation and preservation rates of OC at the present sediment/brine interface to be 5-7.5 g OC/m2/yr. Deeper in the sediments (270-320 m) the rate was 10-15 g OC/m’/yr. The latter depths encompass Unit 11, as defined by Spencer et al (1984), which is a time-stratigraphic equivalent of a mirabilite deposit and thus represents a period of high salinity in the lake. Comparing these accrual rates to the current primary production, 145 g OC/mz/yr in the south arm of the lake (Stephens and Gillespie, 1976), Eugster (1985)

Data Center ,09126599985,[email protected], For Educational Uses MOBILE PETROLEUM SYSTEMS 397 estimated that only 510% of the OC produced in the lake is preserved in the sediments. Bradley (1963) estimated by two different methods the productivity of ancient Lake Gosiute. These values, 27% OC/m2/yr and 657-876 g OC/m2/yr, compared to his OC accrual rate of 28 g OC/m2/yr suggest preservation rates ranging from 3.2% to 10.4%. More efficient rates of preservation are suggested for the saline environment by our accrual rate of 123 g OC/m2/yr for Paradox Member halites. If this rate is compared to productivity rates of modern evaporite environments which probably provide a range that covers the paleoproductivity of the Paradox environment, such as 145 g OC/m2/yr (Great Salt Lake) and 1,810 g OC/m2/yr (Solar Lake), the preservation efficiency of the Paradox system was 7% to 85%. All of these exercises achieve a rough measurement of the degree of carbon loss due to fermentative processes by anaerobic bacteria. However, it should be noted that fermentative processes are inhibited by high salinities but deamination and formation of fatty acids is not. If there were large concentrations of fatty acids in the sediment, thermal decarboxylation could result in the loss of carbon as methane. One significant problem which we have been unable to evaluate is how the OM in brine, both particulate and dissolved, can be incorporated into the sediments. Obviously, the settling velocity of dead bacterial cells in a brine of 1.25 density must be extremely slow if not zero. Does this mean then that for the total OM product of the system to reach the sediments, the brine column must totally desiccate? Our data from halite samples show these rocks to have the highest carbon accrual rates of the entire evaporite suite. This we suggest is the result of higher productivity plus limited reworking by anaerobes due to high salinities. Although anoxia is frequently cited as one of the major reasons for effective preservation in hypersaline environments, Calvert (1987) makes a convincing argument that primary production and bulk sedimentation rates are more important. If we had more data concerning occluded methane and carbon dioxide in the Paradox halites (Table 4.15), it would be tempting to attempt a mass balance calculation and arrive at a close estimate of paleoproductivity for the Paradox environment. We may be overly optimistic about the representative accuracy of such estimates but our confidence is boosted by the fact that even though the OM in our samples, plus those from the Green River Formation, occur in a wide range of dispersal (0.095-19.9% TOC); when this OM is equated with time, the range in values becomes quite small (4.1 g OC/m2/yr to 28g OC/m2/yr.

~~~~MOBLLE PETROLEUM SYSTEMS~~~~

Convincing arguments have been made against oil and water migrating together from source rock to reservoir rock (McAuliffe, 1979). The low solubilities of hydrocarbons in water with the exception of low carbon number alkanes and aromatics, which decreases even more with increasing salinity (Price, 1976), rules against this (McAuliffe, 1979). Baker (1959) proposed that petroleum is solubilized by surfactants thus increasing the amounts of hydrocarbons moved with formation waters. The lack of sufficient amounts of these surfactants makes Baker‘s theory improbable (McAuliffe, 1979, and

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Szatmari, 1980). It is beyond the scope of this paper to discuss all of the theories and controversies regarding the primary migration of oil and gas; however, there is one peculiar feature of the evaporite environment which may play an overlooked role in migration. This involves the generation of organic-rich brines which could transport petroleum source materials directly into reservoirs prior to catagenesis. An early suggestion was made by Hunt (1968) that primary migration may involve precursor soluble organic compounds rather than hydrocarbons. A similar theory by Hodgson (1971) called for the generation of petroleum under moderate thermal conditions from organic material in solution in migrating formation water. The large volumes of EOM from evaporitic marlstones in Sicily suggested to Palmer and Zumberge (1981) the possibility of the involvement of a soluble low-molecular weight kerogen. The evaporite environment may offer the potential for development of a mobile petroleum precursor. On the basis of high concentrations of dissolved organic compounds found in brines from the Alviso salt ponds on San Francisco Bay (Table 4.1), Hite et al (1984) suggested that brines of this nature might have carried petroleum precursors into overlying reservoirs when the originally porous salt beds compacted. The abundant veins of halite found in carbonate interbeds in thick evaporite sequences attest to this, especially where the sequence has been "milked" by halokinesis. Heavy organic-rich brines might also be responsible for charging adjacent carbonate reservoirs with petroleum source materials where refluxion takes place. In his paper on the "cool shallow origin of petroleum", Yi-gang (1981) suggested that primary migration was the result of migration of source materials and not petroleum. Yi-gang called upon bacterial remains and other organic residues in aqueous suspension rather than dissolved OM. Sonnenfeld (1985) also pointed out the potential for generating hydrocarbons in mobile evaporite brines where they are more mobile than those generated in the impervious shales and micrites of an evaporite sequence. Organic-rich brines are also generated in lacustrine environments where perhaps even higher contents of dissolved OM can be expected. We have previously discussed alkaline brines associated with the Green River Formation where dissolved OC can be 10% or more. Evaporite brines may also be a factor in primary migration of natural gas. In this context, we refer back to the abundance of aliphatic (fatty) acids in the pore waters of the Alviso pond sediments (Tables 4.3-4.7). Furthermore, it is now recognized that many connate brines associated with evaporite deposits have an abundance of the acetate ion. Interestingly, Kharaka et al (1983) concluded that thermal decarboxylation of these acids, after they migrate into reservoir rocks, could produce a large portion of natural gas accumulations.

~~~~VITRINITE REFLECTANCE SUPPRESSION~~~~

The measurement of the intensity of reflected light from polished sections of the maceral vitrinite have wide usage in determining the degree of maturation for certain petroleum source rocks. Kerogen associated with evaporites is generally type I or I1 and thus particles of vitrinite may be scarce. Where vitrinite is present, the percent R,,

~~~~Data Center ,09126599985,[email protected], For Educational Uses EVAPORITES AND RESERVOIRS 399~~~~

(reflectance in oil) frequently shows much lower values than the level of maturation indicated by other parameters. For example, in the Paradox Basin, samples of Cretaceous coal, stratigraphicallyabout 3,000 m above the Paradox Member evaporites, have an average R, of 0.55% over much of the Basin. However, vitrinite reflectance measured in thin black "shales" in the evaporite sequences show similar values (R, = 0.51%) even though these "shales" have generated abundant hydrocarbons and are considered mature (Hite et al, 1984). In sharp contrast to the black shales, a fragment of coal collected from a potash mine level in the upper one-fourth of the Paradox Member had an Roof 0.81 %. The tendency for suppression of vitrinite index in evaporite facies has also been mentioned by Sonnenfeld (1985). The suppression of R, values from facies other than evaporites is also a problem where high concentrations of exinite (hydrogen-rich maceral family of type I kerogen) are associated with vitrinite (Price and Barker, 1985). A common explanation for retardation of vitrinite reflectance is migration of bitumen into the vitrinite during an early stage of diagenesis. An investigation of the problem by Sitler (1979) led him to suggest that the reflectance of vitrinite is directly related to the hydrogen content of the vitrinite so that losses sustained during coalification would lower R,. Newman and Newman (1982) called on strong anaerobic conditions to bring about hydrogen enrichment in vitrinite. In a detailed study of vitrinite suppression, Price and Barker (1985) agreed with Sitler (1979) and Newman and Newman (1982) and suggested that hydrogen enrichment of vitrinite is probably due to bacterial reworking in an anaerobic environment. Our data shows that, at least in the Paradox Basin where vitrinite is intimately associated with exinite in black shales in the evaporite sequence, severe R, suppression occurs. In halite beds where occasional fragments of coal are found, the R, values seem to be normal according to linear plots (Figure 4.29). In the halite, vitrinite is totally isolated from other macerals such as exinite, but it was exposed to limited reworking by anaerobic bacteria. This leads us to conclude that in carbonate (black shale)-evaporite environments that hydrogen enrichment, leading to suppression of vitrinite reflectance, is brought about by some unknown inorganic transfer mechanism from exinite. One cannot ignore that the associated carbonate minerals may be involved in some catalytic capacity.

~~~~EVAPORITFS AND RESERVOIRS~~~~

Evaporite rocks, and this includes the fine-grained carbonates that are part of the evaporite depositional cycle, are noted for their extremely low permeabilities; and, in general, are considered effective seals to fluid and gas migration. On the other hand, evaporites are frequently responsible for post sedimentary creation of reservoirs conducive to petroleum accumulation. The formation of these reservoirs can have many diverse origins. Perhaps the most unique development of a reservoir system, where evaporites are present, involves anaerobic sulfate-reducing bacteria. In this process, the anaerobes reduce the sulfate ion, which comes from solubilizinggypsum or anhydrite, and produce S2-and HCO;. The ubiquitous association of sedimentary sulfur deposits and petroleum

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has led most geologists to conclude that hydrocarbons are the carbon source for cell growth of these bacteria (Davis and Y arbrough, 1Y66; Davis and Kirkland, 1970; Davis and Kirkland, 1979; Kuckmick et al, 1Y79). This view is not widely supported by microbiologists and according to Postgate (1984), the balance of evidence is that hydrocarbons are not utilized by any of these bacteria. This process has acted on a gigantic scale in many thick deposits of anhydrite and gypsum as well as in the caprock of salt domes. A well known example of the latter is the famous Spindletop salt dome of Texas where the caprock reservoir has produced over 55 million barrels of oil (Halbouty, 1967). In addition, in west ‘I’exasand southeast New Mexico, there are many extensive bodies of limestone in the anhydrite and gypsum of the Permian Castile Formation which are bacterial in origin (Davis and Kirkland, 1970). ‘These limestones

~~~~MEAN VlTRlNITE 03 04 05 06 07 08 09~~~~

~~~~x CUTLER \~~~~

~~~~OO x\ a~~~~

~~~~0 0 0 \ \~~~~

Fig. 4.29. Vitrinite reflectance gradient f or Paradox Basin based on samples f rom one well (Permian thru Pennsylvanian) and adjacent outcrops (Cretaceous thru Triassic). 0 samples f rom argillaceous carbonates X coals coal sample completely encased in halite collected from nearby potash mine.

Data Center ,09126599985,[email protected], For Educational Uses EVAPORITES AND RESERVOIRS 40 1 are lenticular and locally are totally enclosed by the Castile anhydrite (Zimmerman and Thomas, 1969, McNeal and Hemenway, 1972, and Kirkland and Evans, 1976). Frequently, the vuggy porosity of these bioepigenetic limestones host commercial deposits of sulfur (Figure 4.30). The significance here is that thick and extensive beds of anhydrite and gypsum, which normally would not be thought of in terms of potential petroleum reservoirs, can under the right conditions undergo large scale replacement by bioepigenetic carbonates with excellent porosities. Because of their high solubilities, evaporite layers are subject to dissolution especially around basin margins. The removal of even a thin layer of halite can cause collapse of overlying strata forming blanket-like breccias with good reservoir characteristics. In addition, dissolution at the base of a thick evaporite layer may stope upward and form breccia pipes with large vertical dimensions. Halokinesis, especially diapirism, can create extensive fault and fracture systems in sediments surrounding salt domes as well as causing abundant stratigraphic pinchouts along dome flanks due to concurrent sedimentation and rise of salt. Well known examples include the salt domes of the Gulf Coast region of the U.S.A. and Mexico (Hal- bouty, 1967). In addition, many evaporite sequences include thin interbeds of dolomites and sometimes black shales which ordinarily have negligible permeabilities. However, the intense fracturing of these brittle beds that results as they are moved along by the plastic flow of the rest of the sequence can form extensively connected fracture systems that allow accumulation of petroleum and brines. Furthermore, these types of fractures are generally propped open by a filling of halite crystals. The halite fillings of the fractures form fibrous crystalline masses with the long axes of the crystals oriented at 90 degrees to the wall of the fractures. In many cases the original fractures seem to have been enlarged by the force of crystal growth, and the fibrous habit of the halite crystals forms a permeable conduit for fluids and gases. The thick evaporite sequence of Penn- sylvanian age in the Paradox basin of Utah and Colorado contains reservoirs of this type-

Fig. 4.30. Core ,from Castile Formation West Texas showing suy'ur mineralization, bioepigenetic calcite and vuggy porosity. The laminated f abric of the original anhydrite is still evident. Photograph courtesy of E.S. Montgomery.

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~~~~CONCLUSIONS~~~~

This paper has touched on a broad number of subjects which are part of the petroleum-evaporite relationship. Some of our conclusions, as well as others drawn from the literature, that we suggest are important to the understanding of this relationship, are as follows: 1) Primary production of organic matter (OM) in evaporite environments increases with increasing salinity until the supersaline stage is reached. At lower salinities (vitasaline and penesaline) the dominant contributors of biomass are green algae and cyanobacteria.At higher salinities ( > 20CFho) halophilic bacteria dominate the ecosystem. 2) Studies of the brines and porewaters of marine salterns have established the presence of large amounts of dissolved organic carbon (DOC). This DOC, which increases with salinity, includes important amounts of volatile fatty acids (VFA) especially acetic acid. High salinities ( > 2000/60) do not impede the prcduction of VFA, but they do inhibit or prevent destruction of OM and VFA by methanogenic and sulfate reducing bacteria. 3) In addition to protecting OM from destruction by bacteria, high salinities also exclude burrowing and grazing animals which can totally destroy the microbial mats of hypersaline ecosystems. 4) Periods of maximum volume of evaporite facies correlate with major petroleum accumulation only in the Upper Devonian, Lower Cretaceous, and Lower Neogene. Is it possible then that there is no significant relationship between evaporites and petroleum? 5) The maxima for accumulation of OM in the evaporite environment is not in the vitasaline (mesosaline) phase but in the higher salinity phases. However, because of much more rapid depositional rates at the higher salinities, the OM content of the rock will be lower due to dilution by evaporite minerals. 6) Many darkcolored, organic-rich carbonates, which are frequently called "shales", are important source rocks. Many of these carbonates were deposited in hypersaline environments. 7) Terrestrial OM may make a more important contribution to evaporite basins, both marine and lacustrine, than previously recognized. An important part of this contribution may be in the form of humic acids which would precipitate when the transporting meteoric water mixed with high density evaporite brines. 8) A major problem in establishing the true geochemical characteristics of evaporite rocks and oils is the frequent lack of adequate descriptions of the lithologies or geologic framework from which samples were obtained. Many times a sample is simply described as representative of a hypersaline environment and the basis for that judgment is unexplained. 9) Extracts of evaporitic carbonates generally show pristane/phytane < 1.0 which is interpreted as the result of phytol or phytanic acid transformation to phytane in a highly reducing environment. Anomalous exceptions include Paradox black "shales"

~~~~Data Center ,09126599985,[email protected], For Educational Uses CONCLUSIONS 403~~~~

which consistently show pristane/phytane > 1.0. Since we know by other parameters that these rocks were deposited under strongly reducing conditions, we suggest the predominance of pristane is due to oxidation of phytol from land plants outside the basin and then transported into the evaporite basin by fluviatile means. A higher concentration of C&-G,hopanes relative to C,,-C,, steranes suggests extensive reworking of the terrigenous and algal OM in the vitasaline stage by anaerobic bacteria. 10) Anhydrites are characterized by low TOC and low EOM values. However, because of the tremendous loss of H,O during the diagenetic change of the original gypsum to anhydrite (21%), bitumen or hydrocarbons could be expelled along with this water into overlying reservoirs. However, the ratios of H,O to either hydrocarbons or bitumen would seem too high to create significant accumulations of oil and gas. 11) In general, the organic characteristics of immature to moderately mature marine evaporitic rocks are: pristane/phytane < 1 (largest ratios are generally associated with halites and smallest ratios with anhydrites); CPI <1; moderate to abundant concentrations of cyclic biomarkers relative to n-alkanes; abundance of hopanes > steranes; abundance of G, steranes/ G9steranes variable; tricyclic terpane maximum generally CZ3;moderate abundance of C, tetracyclic; abundant extended hopanes; and v-shaped distributions of methyl dibenzothiophenes in immature to moderately mature rocks (stairstep distributions in mature rocks). 12) Szatmari (1980) has proposed most of the world's largest oil deposits as having been derived from evaporitic source rocks, especially halites. Our data suggest that the EHC and EOM of most halites is very low; and since effective release of these materials would involve dissolution of halite, the enormous volume of brine generated would lower the EOM of the brine to about 0.00033%. This concentration of EOM seems unlikely to be the basis of large accumulations of oil. 13) Hydrogen is commonly found in halites and potash deposits. This is predictable because the high salinities required to form these rocks bring about increased production of hydrogen by fermentative bacteria; and at the same time, decreased consumption of hydrogen by sulfate reducing and methanogenic bacteria. 14) In the anoxic saline to supersaline evaporite environment, the coalification process of woody plant material proceeds with minimal influence of bacterial reworking. Extracted HC from two coal samples of different geologic age encapsulated in halite shows no n-alkane carbon preference except at n-C,, and n-C,,, and little to no cyclic terpanes or steranes. The abundance of the C,, and C,, alkanes suggests that thermal decarboxylation of fatty acids and alcohols in halite favors loss of only one carbon rather than the expected two carbon loss reported by Shimoyama and Johns (1972). Decarboxylation is further supported by the evidence in conclusion 15. Pristane/phytane in the two EOM samples was 0.9 and 1.1. In this regard, it has been noted that halophilic bacteria common to hypersaline environments possess abundant lipids with a phytanyl moiety (ten Haven, et al, 1987) that lead to pristane/phytane ratios <1. Therefore, the low pristane/phytane ratios for coaly organic matter in our samples probably reflects some combination of both microbial and coaly signatures. 15) Petroleum inclusions in Paradox halites are predominantly light hydrocarbons

~~~~Data Center ,09126599985,[email protected], For Educational Uses 404 PETROLEUM AND EVAPORITES~~~~

(C,-CJ which show a strong odd carbon preference through C,-C,. This distribution pattern correlates with even carbons from fatty acids extracted from sediments of a modern saltern. This correlation suggests the light hydrocarbons in the inclusions may result from thermal decarboxylation of fatty acids, alpha to the carboxyl group in organic-rich brines occluded in the halite. 16) Hydrocarbon extracts from organic-rich marlstones of the Eocene Green River Formation, associated veins of gilsonite, and precipitates of organic acids and occluded OM from associated sodium carbonate brines, all have strikingly similar geochemical characteristics. This suggests that trona acid brines are capable of transporting H@s (including the large cyclic biomarkers) as well as fatty acids and that a component of the OM associated with the Green River oil shales may be derived from precipitation of the organic acids. 17) The organic geochemical properties of immature to marginally mature lacustrine carbonate rich mark and mudstones are: pristane/phytane < 1; CPI > 1; abundant non-H@s relative to H@s; abundant cyclic biomarkers relative to n-alkanes; abundance of steranes/hopanes > 1 in thermally immature rocks, < 1 in mature rocks; abundance of C& steranes > C2, steranes; tricyclic terpane maximum generally CZl; moderate abundance of C,tetracyclic terpane; minimal abundance of extended hopanes (GI to G5);minimal to moderate abundance of gammacerane; and moderate to abundant amounts of p carotane. 18) The major organic chemical differences between lacustrine and marine evaporites are: CPI < 1 in marine and > 1 in lacustrine; abundance of non-HCs much higher in lacustrine samples than in marine samples; tricyclic terpane maximum, Goor C,, in lacustrine samples and G3in marine samples; abundance of extended hopanes (GIto C&) greater in marine than lacustrine rocks; and p -carotane much more prevalent in lacustrine rocks than in marine rocks. 19) The sulfur content of evaporite oils is probably directly a function of salinities of the depositional environment of the source rock. As salinities increase the activities of sulfate reducing bacteria decrease, thus reducing the chance for incorporation of sulfur into organic matter during diagenesis. Many of the evaporitic carbonates associated with high sulfur oils were deposited under optimum conditions for bacterial reduction of sulfate. Other evaporitic carbonates, such as those in the Paradox basin, were deposited in hypersaline conditions with reduced activities of the anaerobes and this resulted in oils with sulfur contents as low as 0.02%. 20) Geochemical characteristics of immature to moderately mature oils from marine evaporitic sequences are: pristane/phytane < 1in carbonates and anhydrites and 2 1 in potash and halites; CPI < 1; abundant cyclic biomarkers relative to n-alkanes; abundance of hopanes > steranes; relative abundance of G7-G9steranes variable; tricyclic terpane maximum generally G3;moderate abundance of C;, tetracyclic terpane; abundant extended hopanes; and a v-shaped distribution of the methyl dibenzothiophenes. These geochemical features of the oils match those of the rocks very well. 21) The organic geochemical properties of oils derived from marginally mature to mature lacustrine carbonate rich rocks are: pristane/phytane variable (range 0.4to 1.7);

~~~~Data Center ,09126599985,[email protected], For Educational Uses CONCLUSIONS 405~~~~

CPI > 0.9; abundance of non-HC relative HC generally high; abundance of steranes < hopanes; abundance of C& steranes > G7steranes; tricyclic terpane maximum generally or G1;moderate abundance of GA tetracyclic; abundance of extended hopanes minimal; gammacerane generally more abundant than Gl hopanes; and abundant p carotane. 22) The major chemical differences between oils derived from lacustrine and marine evaporites are: pristane/phytane ratios almost always <1 in marine but frequently >1 in lacustrine; CPI generally <1 in marine but >l in lacustrine; non-HC/HC ratio higher in lacustrine than in marine; abundance of C;, steranes relative to C, steranes variable in marine but G74&steranes in lacustrine; C& tricyclic terpane dominate tricyclic in marine but C, or C;, dominate tricyclic in lucustrine; extended hopanes (GIto (&) abundant in marine but minimal in lacustrine; and the abundance of pcarotane is minimal in marine but abundant in lacustrine oils. 23) Laminated or varved evaporite rocks allow the calculation of the yearly amount of OM delivered to and preserved in the sediment. Laminated anhydrites gave a range of 5.6 to 15.0 g C/m'/yr, carbonates (lacustrine) 12.5 to 23.0 g C/mz/yr, and halites 123.0 g C/m'/yr. Thus, the accrual rate of organic carbon seems to be greatest at highest salinities. Considering some of the high productivities of modern day hypersaline environments (145-1,810g C/m'/yr) the preservation efficienciesof a saline environment are high; 7 to 85%. 24) High concentrations of dissolved organic matter (DOM) are common in both marine and lacustrine brines. The DOM can, in a manner of speaking, be considered a mobile petroleum source. It is possible that the classic problem of primary migration of petroleum may be partly solved by moving the highly mobile DOM into reservoirs before it converts to oil. 25) Coalified wood in halite beds of evaporite sequences shows no suppression of vitrinite reflectance values, however, the vitrhite of organic-rich carbonates ("shales") intercalated with the halites often shows strong suppression. Reflectance suppression in these cases appears to be related to vitrinite association with hydrogen rich OM. In addition, the associated carbonate mineral matrix may act as catalyst for hydrogen transfer from the hydrogen rich OM to the vitrinite, slowing the coalification process. 26) Geologic processes associated with evaporite deposit can create high quality reservoir rocks. Particularly noteworthy are porous carbonates which originate as the result of bacterial sulfate reduction in deposits of gypsum and anhydrite. Breccias produced as the result of evaporite dissolution, as well as extensive fault and fracture systems originating from halokinesis, are known to form excellent reservoirs. In this paper, we have attempted to touch on some of the factors pertaining to the complex relationship of evaporites and petroleum. Excellent progress in this field has been made through studies of the organic geochemistry of modern evaporite settings. However, with the exception of evaporitic carbonates, scarcity of good data from ancient evaporite sediments is a definite handicap. Future productive work will need a well orchestrated joint research effort between teams of diverse disciplines. Hopefully then, we can add more to the story than did Lot's wife.

~~~~Data Center ,09126599985,[email protected], For Educational Uses 406 PETROLEUM AND EVAPORITES~~~~

~~~~Appendix4.1. Conipound ussignmentsf’orfigures4.10, 4.11, 4.12, 4.15, 4.16, 4.18, 4.19, 4.20, 4.21, 4.22, 4.23, 4.24, 4.25, 4.26, 4.27, and 4.28.~~~~

1. (heptadecane) 2. 2.h.10,14-tetramethylpcntadecane(pristane) 3. (trtadccane) 4. 2,6,10,14-lelramethylh~xa~~nc(phytane) S. p arotane 6. tricyclic terpane 7. tricyclic terpane 8. tricyclic terpane 9. tricyclic terpane 10. tricyclic terpanc 11. tricyclic terpane 12. tricyclic terpane 13. tetracyclic terpane 14. tricyclic terpane 15. tricyclic terpane 16. 1Xa (H)-trisnomrohopane (Ts) 17. 17a (H)-trisnorhopane (Tm) 18. 17a(H),21p (H)-norhopane 19. 17a (H)JlP (H)-hopane 20. 17a (H),21P (H)-homohopanc (22s) 21. 17a (H),21P (I-1)-homohopane (22R) 22. (?ammacerane 23. 17a (H)Jlp (H)-hiahomohopane (22s) 24. 17a (H),21P (H)-bishomohopane (22R) 25. 17a (H),21p (H)-trishomohopane (22s) 26. 17a (H),21p (H)-trishomohopane (22R) 27. 17a (H),21 p (H)-tetrakishomohopane (22s) 28. 17a (H),21 p (H)-tetrakishomohopane (22R) 29. 17a (H)JlP (H)-pentakishomohopdne (22s) 30. 17a (H),21p (H)-pentakishomohopne (22R) 31. 13p ,17a-diacholestane (20s) 32. 13p ,17a -diacholeatane (20R) 33. 14p .17p -cholestane (20R) t C, 24-ethyl-13P.17a -diacholestane(WS) 34. 14a ,17a -cholestane (20R) t C, 24-ethyl-13p,17a-diacholestane(~R) 35. 24-ethyl- t4a ,17a -cholestane (20s) .%. 24-ethyl-14P ,17p -cholestane (2QR) 37. 24-ethyl-140 ,17P-choIestane (20s) 38. 24-ethyl-14a,17a -cholestane (20R)

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southern Rocky Mountains. Bull. Am. Asstr. Pet. Geol., 53 884-908. Philp, R. P. and Fan Zhaoan, 1987. Geochemical investigation of oils and source rocks from Qian,pang Depression of Jianghan Basin, a terrigenous sahe basin, China. Organic Geochem.. 11: 549-562. Philp, R. P., 1986. Geochemistry in the mrch for oil. Clem. and Eng. News, Feb. 10,1986, pp. 2843. Philp, R. P. and Zhaoan, F., 1987. Geochemical investigation of oils and source rocks from Qianjiang depression tf Jianghan Basin, a terrigenous saline basin, China. Organic Geochem.. 11: 549-562. Post, F. J., 1980. Biology of the north arm. In: J. W. GW~M(Editor), Great Salt Lake: a Scientific, Historical and Economic Ovewiew. pp. 313-321. Postgate, J. R., 1984. The Sulfate Reducing Bacteria. Second Edition. Cambridge University Press, Cambridge, 208 pp. Powell, T. G., 1986. Petroleum geochemistry and depositional setting of lacustrine source rocks. Marine and Pet. Geol., 3: 200-219. Price, L. C., 1976. Aqueous solubility of petroleum as applied to its origin and primary migration. Bull. Am. Assoc. Pet. Geol., 60: 213-244. Price, L. C. and Barker, C. E., 1985. Suppression of vitrinite reflectance in amorphous rich kerogen - a major unrecogruzed problem. J. of Pet. Geol., 8: 59-84. Robinson, W. E. and Cook, G.L., 1975. Compositional Variations (f Organic Material from Green River Oil Shale - WOSCO Ex-I core (Utah). U.S. Bureau of Mines, Rept. Invest., 8017: 40 pp. Ruckmick, J. C., Wimberly, B. H. and Edwards, A. F.. 1979. Classification and genesis of biogenic sulfur deposits. Eon. Geol. Bull., 74 469474. Sackett, W. M., Brooks, J. M.. Bernard, B. B., Schwab, C. R., Clung, H. and Parker, R. A., 1979. A carbon inventory for Orca Basin brines and sediments. Earth Planetary Science Letters, 44: 73-81. Sammy, N., 1985. Biological systems in northwestern Australian solar salt and fields. In: B. C. Schreiber (Editor), Proc. Sixth International Symposium on Salt: Salt Institute, Alexandria, VA, 1: 207-215. Schidlowski, M., Matzigkeit, U., Mook, W. G. and Krumbein, W. E., 1985. Carbon isotope geochemistry and 14C ages CS microbial mats from the Gavish Sabkha and the Solar Lake. In: G. M. Friedman and W. E. Krumbem (Editors), Hypersaline Ecosystems. The Gavish Sabkha. Springer-Verlag, N.Y., pp. 38141. Schmalz, R. F., 1969. Deep-water evaporite deposition: a genetic model. Bull. Am. Assoc. Pet. Geol., 53: 798-823. Shimoyama, A. and Johns, W. D., 1972. Formation of alkanesfrom acids in the prcsence d CaCg. Geochim. Cosmochim. Acta, 36: 81-91. Sinniie Damstt, J. S., ten Haven, H. L., de Leeuw, J. W. and Schenck, P. A,, 1986. Organic geochemical studies of a Messinian evaporitic basin, northern Apennines (Italy). 11. isoprenoid and n-alkyl thiophenes and thiolanes. Advances in Organic Geochemistry, 10: 791-WI5. Sitler, J. A,, 1979. Effects of Source Material on Vitrinite Reflectance. Master’s Thesis, West Virginia University, Morganttnvn, 108 PP. Sonnenfeld, P., 1984. Brines and Evaporites. Academic Press, Inc., Orlando. 613 pp. Sonnenfeld, P., 1985. Evaporites as oil and gas source rocks. J. Pet. Geol., 8: 253-271. Spencer, R. J., Baedecker, M. J., Eugster, H. P., Forester, R. M., Goldhaber, M. B., Jones, B. F., Kelts, K., McKewie, J.. Madsen, D. B., Rettig, S. L., Rubin, M. and Bowser, C. J., 1984. Great Salt Lake and precursors, Utah the last X),(XX) years. Contrib. Mineral. Petrol., 86: 321-334. Stephens, D. W. and Gilespie, D. M., 1976. Phytoplankton production in the Great Salt Lake, and a laboratory study of algal response to enrichment. Limnology and Oceanography, 21: 74-87. Stoeckenius, W., 1976. The purple membrane d salt-loving bacteria. Sci. Am., 234: 38-46. Sturm, P. A,, 1980. The Great Salt Lake brine system. In: J. W. Gwynn (Editor), Great Salt Lake: A Scientific, Historical and Economic Ovewiew. Bull. Utah Geol. and Mineral. Surv., 116: 147-162. Swanson, V. E. and Palacas, J. G., 1965. Humate in Coastal Sands of Northwest Florida. US. Geol. Surv., Bull., 1214-B. Szatmari, P., 1980. The origin of oil deposits: a model based on evaporites. Anais do 31 Congress0 Brasileino de Geologia, 1: 455499. ten Haven, H. L., de Leeuw,J. W., Rullkotter and SieDamst&, J. S., 1987. Restricted utility d the pristane/phytane ratio as a paleoenvironmental indicator. Nature, 33: 641643. ten Haven, H. L., Baas, M., de Leeuw, J. W., Maassen, J. M. and Schenck, P. A,, 1987. Organic geochemical characteristics (f sedimentsfrom the anoxic brine-fiied Tyro basin (eastern Mediterranean). Organic Geochem., 11: 605-611. ten Haven, H. L., de Leeuw, J. W. and Schenck, P. A., 1985. Organic geochemid studies of a Messmian evaporitic basin, northemAppennines(1taly). I. Hydrocarbon biological marker sfor a hypersaline environment. Geochem. Cosmtrhim. A&, 49 2181-2191. Tissot, B. P. and Welte, D. H., 1984. Petroleum Formation and Occurrence. Second Edition. Elsevier, N.Y., 69Y pp. Treesh, M. L. and Friedman, G. M., 1974. Sabkha deposition d the Salina Group (Upper Silurian) of New York State. In. A. H. Coogan (Editor), Fourth Symp. on Salt, Northern Ohio Geological Society, Cleveland, pp. 354.

~~~~Data Center ,09126599985,[email protected], For Educational Uses REFERENCES 41 1~~~~

Warren, J. K., 1986. Shallow-water evaporitic environments and their source rock potential. J. Sed. Petrol., 56: 442454. Weeks, L. G., 1%1. Origin migration, and (Kcurrence of petroleum. In: G. B. Moody (Editor). Petroleum Exploration Handbook. McGraw-Hill Book Co., pp. 5-50, Wiesenburg, D. A., Brooks, J. M. and Bernard, B. B., 1985. Biogenic hydrocarbon gases and sulfate reduction in the Orca Basin brine. Geochim. Cosmochim. Acta, 49 2069-2080. Wong, H. K., wens, E. T. and Fmckh, P., 1978. Structures in modern Lake Van sediments as revealed by 3.5 KHz high resolution prding. In: E. T. Degens and F. Kurtman (Editors), Geology of Lake Van. M.TA. Press, Ankara, Turkey, pp. 11-19, Woolnough, W. G., 1937. Sedimentation in barred basins and source rocks ofoil. Bull. Am. Assoc. Pet. Geol., 21: 1101-117. Yi-gang, Z., 1981. Cool shallow origin of petroleum -microbial genesis and subsequent degradation. J. d Pet. Geol., 3 4274l4. Zharkov, M. A,, 1981. History of Paleozoic Salt Accumulation. Springer-Verlag, N.Y., 308 pp. Zimmerman, J. B. and Thomas, E., 1969. Sulfur in west Texas: Its Geology and Economics. Texas Bur. Econ. Geol., Univ. d Texas, (~eol0gicaJCircular, 69-2 35 pp.

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~~~~Data Center ,09126599985,[email protected], For Educational Uses Chapter 5~~~~

~~~~HALOKINESIS, CAP ROCK DEVELOPMENT, AND SALT DOME MINERAL RESOURCES~~~~

~~~~J. Richard Kyle and Harry H. Posey~~~~

~~~~INTRODUCTION~~~~

Salt domes, their cap rocks, the associated chemical environments, and the adjacent sedimentary rocks and structures form one of the most economically viable evaporite-related geologic settings. The economic resources and utilizations of the salt dome setting are remarkably diverse. The major economic products of the salt dome environment are oil and gas that occur on the margins of salt stocks, halite and potash salts, and cap rock-hosted native sulfur deposits. Some cap rocks are sources of base metals, limestone, gypsum, or anhydrite, and some are potential sources of celestite or barite. Uranium is recovered from units above or adjacent to diapirs. Caverns, excavated within the salt stocks, serve as material storage for diverse products including liquefied petroleum gas and hazardous waste. In addition to their importance as mineral and hydrocarbon resources, salt diapirs are important for their controls on basin depositional architecture, formation water evolution, sediment diagenesis, and hydrodynamic environments. For instance, the margins of diapirs serve as restricted fluid-escape routes during compaction and diagenesis of basin sediments. In addition to being key components of oil and gas migration, these routes are volumetrically important zones of fluid mixing leading to fluid and mineral diagenesis and thus exert important controls over the fate and compositions of basin formation waters. Dissolution of halite in diapirs provides electrolytesthat affect mineral dissolution and alteration in the sedimentary column and, thus, the composition of both the sediments and formation fluids. Man-made salt caverns,which generally are excavated by injectingfresh water into the salt mass and dissolving the halite, are currently being used for several types of material storage; this use appears to be a growth market for salt diapirs. Commodities as diverse in composition and form as crude and refined oil products, liquefied petroleum gas (LPG), and radioactive nuclear waste are currently being stored or are being considered for salt cavern storage. Such ventures, however, are not without risk and are worthy of considerable study prior to such broad applications. Cavern collapse, surface subsidence, ground water contamination, and petroleum combustion are among the risks involved with salt dome cavern storage. The maximum useful life of salt caverns

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is not yet known, and the tectonic stability of salt diapirs is still being evaluated. The low cost of building these storage caverns indicate that construction and use are likely to expand. U.S. Gulf Coast salt diapirs were recently rejected as potential sites for the disposal of high level nuclear wastes, but salt diapirs in Europe remain strong contenders for storage sites for nuclear and other wastes. With the exception of halite, none of the salt dome resources would exist except for complex interactions between halokinesis, basinal fluids, and the considerable temperature differentials that are present in the salt dome environment. Whereas halokinesis creates structures that facilitate mineralization in diapirs, fluid interactions within the diapirs may drive diapirism, and fluid interactions in the proper temperature environment can lead to mineralization. The general mineralization type appears to be a common feature of salt diapirs, although economic metal concentrations seem to be uncommon and known deposits are small. Where salt dome-related metal deposits are mined, economic feasibility is defined more by local than by global economic conditions. One of the focal points of salt dome research is the cap rock. A typical cap rock sequence consists of an upper calcite zone which is underlain by a gypsum-bearingzone that is transitional with a lower anhydrite zone directly above the halite diapir. The upper surface of the halite is generally very flat, having structural amplitudes generally less than a meter over the crest of the diapir. While halite dissolution is active, there is a salt dissolution zone between the diapir and the overlying anhydrite where anhydrite residues left from halite dissolution accumulate and are underplated to the pre-existing compacted anhydrite by the upward pressure exerted by the rising diapir. In relatively shallow and cool ( < looOC) environments, hydrocarbons and associatedformation waters in contact with the anhydrite cap rock will cause part of the anhydrite to dissolve, sulfate in solution will be bacterially reduced, and calcite will form as a by-product. Hydrogen sulfide which forms by this reaction, if trapped and oxidized, can accumulate as sulfur deposits. Gypsum, an apparently late mineral to form, generally fills voids within the lower part of calcite cap rock and replaces anhydrite and fills voids in the upper part of the anhydrite zone. Because gypsum commonly occurs with sulfur, we believe that their formation is genetically related. Oxygen in groundwater serves to oxidize H,S, whereas the low salinity groundwater serves to hydrate the anhydrite and precipitate the excess, unreduced dissolved sulfate as gypsum. Salt dome mineralization, which is a focus of this paper, has been classified broadly as either a sub-set of the Mississippi Valley type (MVT) deposits or a "hybrid of MVT and sedimentary exhalative (Sedex) deposits (Kyle and Price, 1986). The distinct isotopic signatures of associated carbonates in salt domes, the presence of sulfur or celestite in some domes and their general absence in MVT and Sedex deposits, the

~~~~Data Center ,09126599985,[email protected], For Educational Uses GULF COAST OF SOUTHERN NORTH AMERICA 415~~~~

association of salt domes with large oil and gas fields coupled with the general lack of this direct association in MVT and Sedex deposits, and differences in the mechanisms of ore fluid migration lead us to propose that salt dome mineral deposits are a distinctly different mineralization type. We also propose that the sulfide minerals form as a result of local sulfate reduction, at or near the site of mineralization, with reduction being dominated by biochemical processes. The probable differences in the sulfate-reduction mechanisms leading to mineral precipitation--biogenic in salt domes versus thermogenic in MVT deposits--underscore the real differences in these two deposit classes. Halokinesis occurs by three general processes: (1) gravity flow, similar to glacial movement, (2) thrusting, associated with collision tectonics, or (3) diapirism, most commonly attributed to the complementary effects of sediment loading and the differences in buoyancy between a low-density halite-dominated evaporite sequence and the overlying higher-density, fluid-saturated sedimentary sequence. We focus on the diapiric style of halokinesis in this review. The diapiric environment comprises a complex series of interactions between diapirism, depositional systems, basin fluid evolution and migration, hydrocarbon maturation and migration, mineralization, basement tectonics, and fluid-rock reactions, perhaps including metamorphic reactions. By evaluating information provided by each of these major disciplines, we come closer to understanding the characteristics of individual diapirs from the broad vantage, and to making reliable predictions of the safety and integrity of individual domes for human uses. There are many economically important halokinetic basins, including the Gulf Coast region of southern North America, the Paradox Basin of the southwestern United States, the Maritimes region of eastern Canada, the Zechstein Basin of western Europe and the North Sea, the Maghreb region of northwestern Africa, the Congo Basin of west-central Africa, and the Serigipe Basin of eastern Brazil. In this brief review, we will consider aspects of halokinesis, fluid-mineral diagenesis, oil and gas formation and migration, and genesis of economic mineral concentrations using mostly examples from the Gulf Coast region of southern North America.

~~~~GEOLOGIC SETTING OF MAJOR SALT DOME PROVINCES~~~~

~~~~Gulf Coast of Southern North America~~~~

The development of halokinetic structures in the Gulf Coast sedimentary basin is affected by the basin geometry and depositional systems, basin hydrodynamics, and sediment compaction and diagenesis. The Gulf Coast basin hosts a thick succession of

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Mesozoic and Cenozoic sedimentary strata that rest above a Precambrian and Paleozoic basement (Figs. 5.1 and 5.2; Martin, 1978). Basin development began as a passive continental margin in late Triassic, following rifting of the north Atlantic Ocean and apparently just prior to the rifting of South America away from western Africa (Buffler and Sawyer, 1985). The basin floor has been subsiding throughout the Mesozoic and Cenozoic, and the depositional center has shifted progressively toward the present Gulf of Mexico (Fig. 5.1). Thus, most Mesozoic and Cenozoic formations regionally thicken and dip gently toward the present Gulf of Mexico (Fig. 5.2). During late Triassic and perhaps earlier, local rift-related basins formed in response to extension within the incipient Gulf of Mexico basin. Although most of the

~~~~EXPLANATION~~~~

~~~~Bo Boling BU Butler DM Damon Mound GH GypHill HO Hockley LP Long Point OK Oakwood RA Rayburns W Richton ST Spindletop SU Sulphur Mines Taturn VH Vacherie WH West Hackberry WF Winnfield~~~~

~~~~f Salt dome e Salt massif~~~~

~~~~0 200 km 0- 100mi~~~~

~~~~Fig. 5.1. General geologic setting of the Gulf Coast showing location of selected salt domes. Modified after Martin (1978) and Posey and Kyle (1988).~~~~

Data Center ,09126599985,[email protected], For Educational Uses GULF COAST OF SOUTHERN NORTH AMERICA 417 extensional basins are too deeply buried to have been reached by drilling, the composition of the initial Triassic sediments within them is inferred from the shallower basins (the North Louisiana, East Texas and Mississippi salt basins) and from age equivalent Triassic basins in the eastern United States. These consist of relatively coarse first-cycle sediments and red beds along with intermittent shales and local basaltic intrusions of the Eagle Mills Formation (Benyhill et al., 1968; Martin, 1978; Salvador and Buffler, 1982; Salvador, 1987). These basins became sites for the accumulation of the thick evaporite deposits of the Jurassic Louann Formation. The four main basins (East Texas, North Louisiana, Mississippi, and Gulf Coastal) are now defined by the abundant and diverse halokinetic structures (Fig. 5.1; Martin, 1978; Salvador, 1987). Diapirism in the interior basins (East Texas, North Louisiana, and Mississippi) apparently reached a peak earlier than in the Coastal Basin (Seni and Jackson, 1983a; Labao and Pilger, 1985). Halokinesis is currently active in the deeper Gulf (e.g. Humphris, 1978), whereas diapirism along the coast or further inland is dormant to weakly active. These age variations are due to the halokinetic response of the mother Louann evaporites to sediment loading; sediments progmded from the continent toward the present Gulf, thus diapirism began earlier in the interior basins.

~~~~N S East Texas Basin Present Coast Line~~~~

~~~~Continental Crust~~~~

~~~~Oceanic Crust~~~~

~~~~Transitional Crust 0 100 200 300 400km~~~~

Q Quaternary UK Upper Cretaceous Middle Jurassic salt UT Upper Tertiary LK Lower Cretaceous diapirs and pillows LT Lower Tertiary J Middle ' -Upper Jurassic Upper Triassic? red beds Fig. 5.2. Northwest-southeast geologic cross section of the Gulf Coast Basin depicting a thick succession oj Mesozoic and Cenozoic sedimentary strata that rest above a Precambrian and Paleozoic basement. Approximate line of section shown on Fig. 5.1. Modified af ter Salvador and Bujjler (1982). Reproduced with permission from Poseyand Kyle (1988).

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The Louann Formation is overlain by a late Jurassic Smackover Formation shelf carbonate sequence which is in turn covered by the regionally extensive Cretaceous limestone of the Edwards Group. Late Jurassic and Cretaceous strata in the Gulf Coast region consist principally of shallow-marine carbonate rocks with an increased fine-grained clastic component in the eastern section. Slow subsidence of the carbonate shelf and limited clastic supply during the Early Cretaceous (Aptian to Cenomanian) promoted the growth of a shelf margin reef complex that is particularly well developed in the subsurface of south Texas where it is an important petroleum reservoir known as the Stuart City Trend (Bebout and Loucks, 1974). Toward the craton, the Edwards Limestone thins and crops out to provide an important aquifer unit. Basinward, the Edwards changes from a shallow platformal facies into a major reef at the paleoshelf edge, then into a deep-water fine-grained clastic facies. The Gulf Coast Cenozoic section is dominated by Tertiary and Quaternary fluvial-deltaicclastic strata. Bulk differences in detrital components are functions of both their specific depositional environment and their provenances. The source areas and fluvial depocenters for these units changed with time, giving rise to significant variations in the detrital mineralogy at both the facies scale and at the scale of formations. Several Tertiary units were particularly high in detrital Ca-plagioclase at deposition; however, most of the Ca-plagioclase now has either dissolved or has been albitized (Boles and Franks, 1979; Boles, 1982; Land and Milliken, 1981; Milliken, 1985). These differences affected the composition of fluids that passed through the section during burial diagenesis, as reviewed by Sharp et al. (1988). Contemporaneous growth faults influenced the nature of individual fluvial units throughout deposition (Gregory et al., 1979; Salvador and Buffler, 1982). These down-to-the-basin faults are the result of shelf margin instability as a consequence of rapid depositional loading. Growth faults are major structural features of the Gulf Coast and are extremely important in controlling the upward migration of diagenetic fluids, including hydrocarbons. These fault trends commonly exert controls on local depositional facies, also an important factor in petroleum reservoir development. The Gulf of Mexico is a relatively young sedimentary basin, the deeper parts of which are undergoing significant fluid evulsion from overpressured sediments (Sharp et al., 1988). Fluids from the geopressured zone escape dominantly through growth faults but also up the margins of salt diapirs. The mechanism of fluid ewlsion in parts of the basin evolved in the basin history from gravity-driven flow, which operated during early basin history, to overpressure-induced fluid flow.

~~~~Data Center ,09126599985,[email protected], For Educational Uses MAGHREB REGION OF NORTHWESTERN AFRICA 419~~~~

~~~~Maghreb Region of Northwestern Africa~~~~

The mother salt of diapirs around the western Mediterranean is a Triassic halite-dominant sequence that is at least 1,OOO m thick. Triassic evaporites in northwest Africa and southern Spain became diapiric shortly after deposition, as is shown by their truncation of beds as old as Jurassic (Rouvier et al., 1985). In northern Tunisia, these diapirs occur most commonly as elongate structures that have breached the surface (Fig. 5.3). Because of weathering, halite generally has been dissolved, leaving a surface accumulation of mostly anhydrite (which has generally converted to gypsum at the surface), shale, and dolostone. Evaporites at the surface have been complexly folded into tight enteroiithically folded gypsum and shale units that envelope broader more regularly folded dolostone layers. The enterolithic folds are caused by material solution (probably of halite) and hydration of anhydrite. However, the broader folds appear to have two origins: diapirism and compression associated with Alpine-age tectonism. The elongate structures may have been influenced by pre-existing Precambrian structures (Rouvier et al., 1985). The composition of the Triassic evaporites is not well documented, although it

~~~~Fig. 5.3. Generalized geologic setting of the Triassic diapirs of the Maghreb region of northwestern Africa showing location of Pb-Zn deposits. Modified after Rouvier et al. ( 1 Y 85).~~~~

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appears to bear less halite, overall, than the Gulf Coast. Jurassic marine carbonates and Cretaceous mark with only minor fine-grained siliciclastics cover Triassic rocks (Bishop, 1988). Mark dominate the Tertiary section up to the Miocene which consists of fluvial sandstones above an angular unconformity. Most of post-Miocene time has been a period of erosion. A few north African diapirs bear bioepigenetic carbonate cap rocks that contain metal sulfides, several of which have been or are being mined. In Tunisia, these include the Bou Grine and Fedj el Adoum deposits (Fig. 5.3). The Bou Grine deposit is the second largest Zn-Pb deposit in northwest Africa containing proven reserves of 7.3 million tons of ore at 2.4% Pb and 9.7% Zn (Orgeval et al., 1989). In addition to the sulfides of zinc, lead, and iron, these deposits contain significant concentrations of barite and celestite with lesser Occurrences of fluorite, magnesite, siderite, and chalcopyrite. Oil and gas are recovered from several localities in north Africa, but large production has not come from the diapir areas (Bishop, 1988).

~~~~THE MECHANICS OF DIAPIRISM~~~~

~~~~General~~~~

Hundreds of salt tectonic structures, including anticlines, pillows, rollers, walls, stocks, nappes, massifs, or namakiers (Fig. 5.4; Jackson and Talbot, 1986) have been identified in the onshore and offshore Gulf Coast region (Martin, 1978; Halbouty, 1979). In general, these salt structures become younger toward the center of local depositional basins. This younging effect is due principally to loading from edge to center but may be due also to the invasion of fresh water in the prograding sediment wedge. Jackson and Talbot (1986) identified six mechanisms for salt diapirism: buoyancy, differential loading, gravity spreading, thermal convection, contraction, and extension (Fig. 5.5). Although buoyancy generally has been regarded as the dominant mechanism for diapirism, Jackson and Talbot demonstrated that density contrast is an ineffective diapiric mechanism because it requires substantial pre-existing relief beneath denser overburden. Differential loading is an effective and geologically realistic mechanism for the initiation and early stages of diapirism and commonly results in asymmetric salt structures. This coupled relationship between sediment loading and halokinesis forms the geologic basis for determination of the timing and extent of salt movement (Halbouty, 1979; Seni and Jackson, 1983a,b). In some regions (Iran and North Africa, for example) salt masses extrude onto the land surface or sea floor (Fig. 5.6; Jackson et al., in press). This situation appears

~~~~Data Center ,09126599985,[email protected], For Educational Uses MECHANICS OF DIAPIRISM 42 1~~~~

Fig. 5.4. Principal types of large salt structures. Structure contours are in arbitrary units. Salt nappes of the Sigsbee Scarp type and irregular salt massifs are an order of magnitude larger and have been omitted. Reproduced with permission from M.P.A. Jackson and C. Talbot (1986)." Geological Society of America, v.97, p. 306, Boulder, Colorado.

to be rare in most halokinetic basins. Most shallow diapirs lift their cover materials like shallow laccoliths. Growth of the salt structure ceases when it (I) penetrates a layer of lower density, (2) encounters resistant overburden, (3) becomes detached from the root, or (4) exhausts the source layer (Jackson and Talbot, 1986). Diapiric growth can be reactivated by further burial and compaction of the cover or by fluid softening. In theory, a water-softened thick salt sequence exposed to a relatively high geothermal gradient typical of the early stages of basin formation could become unstable without any overburden; later, a heat-induced density inversion could lead to convective circulation within a tabular salt mass (Jackson and Talbot, 1986). Halokinetic movements can be initiated, accelerated, or retarded by regional tectonic forces. The ultimate fate of bedded and diapiric salt in most sedimentary basins is dissolution and return to surface or formation waters to complete the geochemical cycle. Although the margins of salt diapirs are noted for their degree of brecciation and fluid flow, unequivocal examples of fossil diapirs, that is, diapirs wherein the evaporite minerals have been removed by solution leaving only vertical breccia pipes, have not been documented. Although most older literature indicates that diapirs intrude overlying rocks in response only to sediment loading, there may be additional causes of diapirism. Posey et al. (1987a) and Posey and Kyle (1988) speculated that diapirism is aided, if not induced, by fluids that pass into the halite during either the seminal stages of diapirism

~~~~Data Center ,09126599985,[email protected], For Educational Uses 422 SALT DOME MINERALIZATION~~~~

~~~~A. BUOYANCY HALOKINESIS B. DIFFERENTIAL LOADING HALOKINESIS~~~~

~~~~c. GRAVITY SPREADING HALOKINESIS D. THERMAL CONVECTIVE HALOKINESIS~~~~

Fig. 5.5. Six principal mechanisms of salt tectonics. A). Buoyancy halokinesis; B). Differential halokinesis; C). Gravity-spreading halokinesis; D). Thermal-convection halokinesis; E). Contraction halotectonics; F). Extension halotectonics. Reproduced with permission from M.P.A. Jackson and C. Talbot, 1986, Geol. SOC.Amer. Bull., v. 97, p.

312. @ Geological Society of America, Boulder, Colo. or during the pillow stage. These fluids may induce diapirism in several ways. For instance, intragranular fluids lower the shear strength of halite, thus making a halite body more plastic. Also, the dissolution of halite at the top or along the margins of a

~~~~Data Center ,09126599985,[email protected], For Educational Uses MECHANICS OF DIAPIRJSM 423~~~~

Fig. 5.6. Subaerial salt diapirs of Iran. Photo by A. Gansser, supplied by M. P. A. .Jackson. Reproduced with permission from M.P.A. Jackson et al. (inpress). Geological Society of America, Boulder, Colo. diapir should allow halite lying beneath the fluid-bearing units to flow into the space created by halite dissolution. Likewise, because the solubilities of several minerals increase in the presence of brines, the dissolution of halite within a diapir may promote dissolution of marginal sediments, which would also induce flow of underlying halite into the open space created by that dissolution. In this section we consider the concepts of sediment loading versus fluid softening as mechanisms that promote diapirism. Fluids that move and interact within and adjacent to salt diapirs are critical to the formation of most of the mineral resources that occur in the salt dome environment. Hydrocarbons are perhaps the best known example. Oil and gas are found commonly in structural traps adjacent to and above salt diapirs, and it is our belief that diapirism and hydrocarbon migration are co-dependent, at least in some cases. Sulfur, the salt dome commodity second only to petroleum in economic value, depends on hydrocdr- bons for its existence, Sulfide deposits in salt dome cap rocks, although small by commercial standards, are significant enough to be mineable in some regions, and are apparently the result of the interaction of at least two different fluids in the cap rock environment. Concentrations of celestite, uranium, and halogen-enriched brines may

~~~~Data Center ,09126599985,[email protected], For Educational Uses 424 SALT DOME MINERALIZATION also be associated with diapirs and are locally of economic importance.~~~~

~~~~Fluid Infiltration vs. Sediment Loading~~~~

Salt domes grow upward in pulses, a fact generally attributed to the effects of differential sediment loading (see Seni and Jackson, 1983a,b; Jackson and Talbot, 1986). According to this theory, as sediment deposition causes evaporites in the mother salt to evulse from the mother salt unit closest to the diapir, clastic sediments fill the area around a diapir forming the rim syncline (Trusheim, 1960; Seni and Jackson, 1983a). This activity causes the rim syncline to fill with proportionally thicker sediments than the adjacent area away from the diapir. By knowing the age of the rim syncline-thickened formations, it is possible to determine the time of diapirism; by knowing the amount of overthickening, it is possible to determine the amount of salt withdrawal from the mother salt. By combining the age of salt withdrawal and diapirism with the abundance of overthickening or salt withdrawal, the rates of diapirism have been determined (see Seni and Jackson, 1983b for a review). These rates are on the order of to 10-l6/sec (strain rate = e/t, where elongation e = change in length/original length and t = duration of strain in seconds). This strain rate is slower than average rates of orogeny, 10-14/sec(see Jackson, 1985, for a discussion). This method of calculating the rates of diapirism assumes that all of the salt accounted for within the overthickened rim synclines has evulsed into the neck of the diapir, and that none has been lost by solution. Although it is recognized that diapirism induces brecciation alongside the diapirs, the amount is not known, and the effect is generally regarded as insignificant. Halite and anhydrite within the diapirs have experienced a considerable amount of dissolution and reprecipitation (Land et al., 1988). Posey (1986) and Posey et al. (1987a) showed that salt-hosted anhydrite, which accounts for about 5% of the material within Gulf Coast diapirs, has a considerable range in strontium isotope ratios. The lowest strontium isotope ratios are identical to those of midJurassic seawater, suggesting that the evaporites are marine in origin. All higher values must be explained by the introduction of radiogenic strontium accompanying later fluid events. Land et al. (1989) reported similar high values for a more extensive suite of salt dome samples and for bedded anhydrites of the same age. The conclusion forwarded by the strontium isotope studies is that anhydrite-strontium within the salt domes underwent partial isotope exchange either before or during diapirism. In addition, the fact that all cap rocks inherited most of their strontium from the salt-hosted anhydrite and have similar strontium isotope ranges indicates that this

~~~~Data Center ,09126599985,[email protected], For Educational Uses FLUID INFILTRATION VS. SEDIMENT LOADING 425~~~~

strontium exchange took place prior to cap rock formation (Posey, 1986). Light et al. (1987) and Light and Posey (in press) speculate that fluids of possible metamorphic origin from beneath the evaporites participated in this event. Similar conclusions regarding halite were reached by Land et al. (1988). Their study found that bromide concentrations in Middle Jurassic halite are high in bedded evaporite or undeformed evaporite units but uniformly low in diapiric halite. The high values in the undeformed halite are typical of halite precipitated from marine water. However, the low values in the diapirs indicate that bromide was lost from the halite after deposition from seawater. The fact that bedded halites have higher, marine-like values, whereas the diapiric halites have lower values indicates that bromide loss occurred during and as a consequence of recrystallization accompanying diapirism. The loss of bromide in diapiric halite and the development of a broad range of strontium isotope ratios in diapiric anhydrite indicate that both the halite and the anhydrite have undergone at least partial recrystallization. Land et al. (1988) noted that such reactions can take place only under aqueous conditions and must indicate mineral dissolution and reprecipitation. Thus, it appears that all of the diapirs studied to date have interacted with fluids some time within their diapiric ascent. The introduction of even small volumes of fluid into salt diapirs would have profound effects on halokinesis as such fluids would lower the strength of the halite, thereby increasing the strain rate (Jackson, 1985). For example, the addition of only 0.1% water to the evaporite mineral bischoffite (MgCI, 6H,O) decreases the flow stress by five times (Urai, 1983). Apparently, water along the grain boundaries promotes dynamic recrystallization by movement of high angle grain boundaries and possibly by increasing intracrystalline plasticity (Jackson, 1985). The addition of fluid may actually have induced diapirism, rather than just decreased the flow stress (Posey et al., 1987a; Posey and Kyle, 1988). Thus, we introduce a debate: is diapiric halokinesis caused by sediment loading or is halite-softening induced by fluid invasion? The answer to this debate has particular importance to the selection of inland didpirs as sites for material disposal or storage on the assumption that such domes will remain free of further diapirism. Indeed, this may be a relatively safe assumption provided sediment loading is the cause of diapirism. However, without a clear understanding of the causes and timing of fluid softening, these diapirs could present some greater, albeit uncertain, amount of risk of movement. A point to consider about a sediment-loading mechanism as a source of diapirism is that a substantial amount of diapirism takes place during deposition of carbonate above the evaporite sequence. This situation has been shown especially for the North Louisiana Basin and for northern Tunisia. In the Winnfield dome, northern Louisiana,

~~~~Data Center ,09126599985,[email protected], For Educational Uses 426 SALT DOME MINERALIZATION~~~~

anhydrite cap rock-hosted sulfide mineralization (Ulrich et al., 1984; Kyle et al., 1987) began during late Jurassic, at a time when the only cover above the Louann Evaporites was the Smackover Limestone. By comparison, in northwest Africa, practically all of the material covering the Triassic evaporites is carbonate (Bishop, 1988), and diapirism was initiated during carbonate sedimentation very soon after deposition of the evaporites. Thus, rapid loading, like that of the Gulf Coast, which would be expected of clastic-dominated sedimentary terranes, is not necessary for diapirism. Diapiric halokinesis can occur due to loading by a variety of sediment types. We believe that fluid intrusion into the salt mass during early deformation, perhaps as early as the pillow stage of development, is a viable mechanism for inducing diapirism, and is a certain mechanism for promoting diapirism. Subsequent sediment loading, typically accompanied by fluids, would further halokinesis.

~~~~Timing of Diapirism~~~~

The early basin-fill sequence of the East Texas Basin, one of the rift basins that developed at the cratonic margin of the Gulf Coast basin, is shallow enough to have been penetrated by drilling. Diapirism within this basin was compared with centrifuge models to illustrate the elements critical to the development of the major salt dome provinces in the Gulf Coast (Seni and Jackson, 1983a,b; Jackson and Tdlbot, 1986). The original Louann Salt thickness in the East Texas Basin, prior to halokinesis, is estimated at 1500 m and presumably was even greater in the main Gulf Basin (Martin, 1978). The salt was covered by sediments deposited from the west, north, and east, and diapirism was apparently induced by the uneven sediment loads supplied at these input points. Most of the diapirs in the East Texas Basin lie within a N-S to NE-SW axial zone of the basin. Studies of depositional patterns within the East Texas Basin indicate that the growth of each of these diapirs is slightly different in time (Seni and Jackson, 1983b). Their key finding, that diapirs near the borders of the basin formed early, whereas those nearer the center formed later, is critical to the interpretations of the mechanisms that activate diapirism and that affect fluid compositions and fluid flow. Centrifuge modeling by Jackson and Cornelius (1987) shows that salt units resting on a slightly inclined surface exhibit gravity flow, and halokinetic structures grow younger in the direction of that flow. Where salt is loaded from more than one side of a basin, diapirs grow younger toward the centers of basins, whereas salt that is loaded from just one side of a basin produces diapirs that become younger leeward of the sediment source.

~~~~Data Center ,09126599985,[email protected], For Educational Uses FLUID MIGRATION AROUND SALT DIAPIRS 42 7~~~~

In the East Texas Basin halokinesis was initiated in Late Jurassic during the loading of the Louann Salt by a prograding carbonate wedge and about 90% of the diapirism took place in the first 30 million years of the 160 million years since salt deposition (Fig. 5.7; Seni and Jackson 1983a.b, 1984). Rapid progradation of the Schuler-Hosston fluvial deltaic sediments over the carbonate shelf in Late Jurassic and Early Cretaceous promoted diapir development which continued throughout the Creta- ceous. Although dome growth has been slow in the East Texas Basin since Cretaceous (about 1.5 x 10I5/sec), sedimentation effects continued into the Tertiary (Fig. 5.7; Seni and Jackson, 1983b). Salt structures of the Gulf Coast range generally from late Jurassic structures in the interior salt basins (Seni and Jackson, 1983a,b; Kyle et a]., 1987) to the currently active salt structures of the deep Gulf (Humphris, 1978). Diapirism began at least as early as Late Jurassic in the North Louisiana Basin. Based on stratigraphic relationships determined from seismic studies, the principal period of diapirism at the Winnfield Dome took place from 147 to 144 Ma (Labao and Pilger, 1985). An inverted stratigraphy--oldest at the top, becoming progressively younger toward the base--has been documented within the Winnfield anhydrite cap (Ulrich et al., 1984; Kyle et al., 1987). Paleomagnetic studies suggest that cap rock formation (and diapirism?) began as early as 157 Ma (Gose et al., 1985, 1989). This difference could be significant concerning interpretations of the mechanism of diapirism and the source of mineralizing fluids, as will be discussed in the following sections.

~~~~Fluid Migration Around Salt Diapirs~~~~

The infusion of fluids into diapirs and the paleomagnetic evidence regarding the timing of diapirism and cap rock formation raise important questions about the mechanisms of intrabasinal fluid flow. The Gulf Coast Basin again provides the geological setting for discussion. Presently, in the more seaward parts of the Gulf Coast basin, fluids below about 3 to 4 kilometers are generally under higher pressures than could be expected if the pressures were established by a normal hydrostatic gradient (see Bethke, 1986, and Hunt, 1990, for expanded discussions of overpressuring). These abnormal formation pressures, which can be detected on drill stem tests, form an irregular but mappable surface, the top of geopressure. Pore fluids in the deeper zone are not interconnected with those higher in the section (above the zone of geopressure), so the fluid pressure within the formation is a combination both of the weight of the fluid column and the rock density (Fig. 5.8). Measured formation pressures, therefore, fall somewhere between normal hydrostatic pressures (the pressure that would be measured at any point

~~~~Data Center ,09126599985,[email protected], For Educational Uses 428 SALT DOME MINERALIZATION~~~~

~~~~Bossier-Gilmer- carbonate build~~~~

~~~~her).~~~~

~~~~TOPOGRAPHIC SW~~~~

~~~~EXPLANATION 0 Primary peripherol sink 0 Secondary peripheral sink !3 Terliory peripheral sink INFERRED BASAL~~~~

Fig. 5.7. East Texas Basin during the last 150 million years. A). Initiation of salt flow in Late Jurassic, 150 to I37 Ma; B). Initiation of salt diapirism in Late Jurassic-Early Cretaceous, 137 to 115 Ma; C). Continued development of salt diapirs in Early Creta- ceous, 115 to 98 Ma; D). Deceleration of diapirism in Early Tertiary, 56 to 48 Ma. Reproduced with permission from M.P.A. Jackson and S.J. Seni, 1981, Geology, v.11, p. 134; Geological Society of America, Boulder, Colo.

~~~~Data Center ,09126599985,[email protected], For Educational Uses FLUID MIGRATION AROUND SALT DIAPIRS 429~~~~

~~~~FLUID PRESSURE~~~~

~~~~I a LL~~~~

Fig. 5.8. Schematic representation of overpressuring in a sedimentary sequence modeled after the geologic setting of the Gulf Coast. Structural by-pass represents a conduit f or upward fluid flow which could be a growth fault or salt dome margin. Modified after Hanor (1987a). in a column of fluid) and normal lithostatic pressures (the pressure that would be measured at any point in a column of rock). Both pressure measurements vary, of course, with density of both rock and sediment. Geopressuring occurs under several circumstances.When water-rich sediments are buried too rapidly for the pore fluids to be expelled or to reach hydrologic equilibrium with fluids in the water column above, the trapped fluids will become overpressured by the weight of the sediment column above (Fig. 5.8; Hanor, 1987a). Likewise, when relatively impermeable rocks, such as anhydrite, form a seal--even a leaky seal--above water-bearing sediments, the upward flow of fluid that would otherwise move as a consequence of sediment loading is blocked (Hunt, 1990). Thus, it is the pressure of the overlying rocks plus pore waters in the affected unit, rather than the pressure from the water column alone, that compress the pore fluids. Hunt (1990) proposed that a more common type of geopressuring forms through the precipitation of diagenetic calcite above organic source rocks in subsiding basins. This "carbonate curtain" is found most typically at pressure/depth conditions equivalent to a vitrinite reflectance (RJ range of 0.4 to 0.5. CO,, which evolves mainly from the breakdown of kerogen in the & range of about 0.9, first dissolves calcite which is reprecipitated at lower pressure higher in the sedimentary sequence. This secondary

~~~~Data Center ,09126599985,[email protected], For Educational Uses 430 SALT DOME MINERALIZATION~~~~

~~~~calcite generally forms at temperatures between about 90-100"C, which is equiv,~~~~

~~~~Data Center ,09126599985,[email protected], For Educational Uses CAP ROCK FORMATION 43 1~~~~

overpressuring caused by the evaporite seal. Gravity-driven fluid movement from the continent may likewise have caused fluid flow. Consequently, we may conclude that metallic mineralization, at least in the interior basins, formed as a result either of gravity-driven flow, with fluids being driven off the North American craton, or by overpressuring of pre-Louann-hosted fluids by the Louann itself. However, the southern Gulf, a geopressured basin, may have been mineralized via geopressure-driven flow. Likewise, in the Mediterranean area where the Triassic evaporites generally are covered by mark (Bishop, 1988), geopressured flow may have never occurred. Gravity-driven flow during early basin history probably came from the craton to the south. Perhaps the more significant flow occurred during Alpine collision and, subsequently, by gravity flow.

~~~~CAP ROCK FORMATION~~~~

~~~~General~~~~

Cap rock formation processes have been reviewed many times, most recently by Posey and Kyle (1989). A brief review of cap rock nomenclature, the stratigraphy, and recent concepts of cap rock formation will be presented to provide background for later discussions. Although diapirs are, in fact, intrusive, the actual thickness of the stratigraphic section that they intrude has been exaggerated. Diapirs actually behave like vertical columns of evaporites (mostly halite) around which younger successions of sediments slide like a sleeve. Once formed, the top of a diapir remains at a relatively constant elevation with respect to the sediment/water interface, while younger sediments that are deposited above and alongside the diapir subside in response to sediment loading, basin compaction, and basin floor subsidence. The upper surfaces of halite diapirs remain relatively close to the surface throughout diapirism or become buried permanently (Posey and Kyle, 1988). This shaft of halite is continually fed from beneath by the mother salt and grows upward until it is either cut off from the mother salt or until the top of the diapir is buried by more sediment than the diapir can effectively penetrate (Trusheim, 1960). Typically, onshore salt diapirs in the Gulf Coast are mantled by cap rock consisting of a basal anhydrite zone directly overlying the halite stock and an upper complex calcite zone (Fig. 5.9). The calcite and anhydrite are commonly separated by an irregular gypsum-bearing zone. The lower part of the calcite zone and the upper part of the anhydrite zone near the contact with gypsum may contain native sulfur

Data Center ,09126599985,[email protected], For Educational Uses 432 SALT DOME MINERALIZATION concentrations. The Winnfield dome in the North Louisiana Basin has provided an unequaled situation in the Gulf Coast to study a salt dome cap rock because active quarrying provides access to three-dimensional exposures of the cap rock zones. The nature and origin of the cap rock features are discussed and illustrated by Ulrich et al. (1984) and Kyle et al. (1987). The salt diapir is roughly circular in outline with a diameter of about 1.1 km at the -180 m structural elevation. The stratigraphic position of the underlying "mother" midJurassic Louann Salt is at a depth of over 5 km in this region (Eversull, 1984). Halite from the Winnfield salt dome is typical of Gulf Coast salt diapirs and includes white to dark gray, coarsecrystalline bands. The color generally is indicative of the amount of non-halite materials in the salt (Fig. 5.1OA-C). Banding in the salt is generally vertical in the interior of the diapirs, but the margins of salt stocks have been complexly deformed to produce inclined banding. Salt produced during mining operations at Winnfield averaged about 3% impurities, dominated by fine- to medium-- grained, idio-blastic to subidioblasticcrystals of anhydrite. Trace quantities of dolomite, siderite, quartz, barite, and pyrite are also found in the salt. Steeply dipping, internally folded anhydrite bands, generally 2-10 cm in width, parallel the foliation and contain 15- 90% anhydrite. High pressure brine and gas inclusions are common throughout the salt, particularly within anhydrite-rich bands (Belchic, 1960). Larger, gas-filled pockets occur in several salt domes in the Gulf Coast. These pockets can cause rock bursts underground when the removal of rock salt at the mine face allows gas pressure in the salt to exceed the rock strength (Iannacchione et al., 1982, 1984; Schatzel and Hyman, 1984; Molinda, 1988).The composition of the gases is domi-

~~~~EXPLAN AT ION FALSE CALCITE CAP IMARlNE CALCITE CAP VARIEGATED CALCITE CAP =BANDED CALCITE CAP GYPSUM (transitional) CAP ANHYDRITE CAP BSALT DISSOLUTION ZONE aSALT STOCK~~~~

~~~~Fig. 5.9. Schematic cross-section of the upper part of a Gulf Coast salt diapir showing an idealized sequence of cap rock lithotypes. Modified after Posev and Kyle (1988).~~~~

~~~~Data Center ,09126599985,[email protected], For Educational Uses CAP ROCK FORMATION 433~~~~

nantly CH, although lesser volumes of C&, GH,, C,H,,, and C,H,, also occur. The source of the methane is not completely known, but its association with higher hydrocarbons (Molinda, 1988) indicates that it is probably thermogenic instead of biogenic. 'I'he Winnfield salt stock is mantled by a typical cap rock sequence of calcite, gypsum, and anhydrite. The top of the salt is 76-80 m below sea level, is horizontal, and is directly overlain by anhydrite cap rock. The salt/anhydrite cap rock contact is marked by a brine zone as much as a few meters thick. Salt dissolution zones are atypical of stable salt domes within the interior salt basins in the Gulf Coast in which the salt/cap rock contact generally is tightly fused. 'I'he anhydrite zone reaches a maximum thickness of 94 m over the center of the dome; because of the dip of the anhydrite zone away from the crest of the dome, the total amount of anhydrite stratigraphy present along the cap rock perimeter is perhaps twice this amount. A thinner anhydrite zone is present along the flanks of the dome to a depth of at least 1500 m. The anhydrite cap rock is light to medium gray and generally layered. Banding and domal structures within the anhydrite zone represents repeated cycles of halite dissolution and anhydrite accretion at the salt/anhydrite interface (Kyle et al., 1987). 'I'he anhydrite consists of tightly interlocking, predominantly xenoblastic crystals with no visible preferred orientation. Porosity and permeability of the anhydrite cap rock are very low. Deformed, banded anhydrite clasts are commonly contained within the layers (Fig 5.10D). These clasts are derived from primary anhydrite beds within the Louann Salt which have been deformed during diapiric emplacement. Sulfides occur as laminar zones parallel to the tightly intergrown, layered anhydrite grains or as coarse euhedral crystals growing in open space within the anhydrite cap rock. Sulfide layers generally are less than one millimeter to several centimeters thick and can be traced laterally for tens of meters. The principal sulfide is monoclinic pyrrhotite, but sphalerite, galena, pyrite, and marcasite are abundant locally. 'The anhydrite/calcite interface at Winnfield is highly irregular and a gypsum zone as much as 12 m thick occurs locally. The calcite zone varies in areal extent and in thickness from about a meter to over 90 m. The calcite zone consists of at least three major types of limestone, ranging from a fossiliferous marine limestone to a highly solution-brecciated,multi-generation detritus-rich limestone to horizontally-layered dark and white bands of brecciated vuggy limestone (Posey et al., 1987b; Prikryl et al., 1988). 'lhe upper portion of the calcite cap rock is a fossiliferous, glauconitic, pelleted marine limestone, the "marine false cap rock" of Posey et al. (1987b). "his unit also contains abundant euhedral quartz and carbonate crystals that are believed to have been derived from the salt mass (Kyle et al., 1987). 'I'he underlying "true" calcite cap rock is distinctly banded with layers of dark Fay, finecrystalline calcite separated by layers of white,

~~~~Data Center ,09126599985,[email protected], For Educational Uses 434 SALT DOME MINERALIZATION~~~~

Fig. 5.10. Geologic features of the upper part of Gulf Coast salt diapirs. A). Tight vertical isoclinal folds in halite showing light and dark salt layers in the interior of the Belle Island, Louisiana, salt dome. B). Sawed core of the steeply inclined salt/anhydrite contact from the periphery of the Boling, Texas, dome. C). Transmitted light photomicrograph of the above contact zone showing disseminated anhydrite in halite overlain by the accretionary zone of anhydrite residue in a halite matrix. Crossed nicols with gypsum plate. Field of view = 20 mm. D). Anhydrite cap rock from the Winnfield, Louisiana dome showing layers of breccia and homogeneous residue anhydrite separated by sulfide laminae. E). Irregularly banded calcite cap rock from the Winnfield dome showing multiple generations of light and dark calcite. coarsecrystalline calcite (Fig 5.10E). Several roughly laminated massive sulfide lenses occur along the calcite/anhydrite transition zone. The largest lens has a maximum thickness of about 5 m and extends for more than 30 m laterally. The massive sulfide

~~~~Data Center ,09126599985,[email protected], For Educational Uses FLUID VOLUMES REQUIRED FOR CAP ROCK 435~~~~

lens consists of marcasite and pyrite with local concentrations of pyrrhotite, colloform sphalerite, galena, and barite. Other cap rocks in the Gulf Coast are generally similar. Gypsum or calcite may be absent, but anhydrite is always present if there is any cap rock at all. Cap rocks of any lithology are not known to occur below about 2 kilometers (Halbouty, 1979). The absence of calcite can be attributed to both the relatively high temperatures which inhibit bacterial sulfate reduction and a lack of meteoric water at those depths. However, the absence of anhydrite cap rocks is less easily explained. Given that the formation water compositions, even in some of the deepest parts of the basin, are below halite saturation, there must be some other chemical or physical mechanism that inhibits halite solution and consequent anhydrite accumulation, or that promotes anhydrite solution as well. The apparent absence of deep anhydrite cap rock remains enigmatic.

~~~~Fluid Volumes Required f or Cap Rock Formation~~~~

Various isotopic and chemical indicators show that fluids have interacted with the halite mass, some time during diapirism, a phenomenon that should have a significant effect on the nature and timing of diapirism. In addition, the very presence of cap rock, not to mention oil, sulfide minerals, or sulfur, indicates that diapiric margins are a major structural conduit through which basin fluids migrate into shallower parts of the basins. Of the many criteria indicating that fluids have interacted within the environments of salt diapirs, the presence of cap rocks is the most indicative. Mass balance calculations indicate that the volume of fluid that accompanies cap rock development and mineralization is enormous. For instance, a one-meter thickness of anhydrite cap rock (anhydrite averages about 5% of the salt in the Gulf Coast diapirs) is indication that at least 20 meters of halite have been removed from the top of the diapir (see Taylor, 1937; Price et al., 1983; Dix and Jackson, 1982; Seni and Jackson, 1983a; Ulrich et al., 1984). Under these conditions, a 2-km diameter dome that has a 300-m thick anhydrite cap rock represents the loss of about 18.8 cubic km of halite from the upper part of the diapir. Likewise, a single volume of halite dissolves in about six volumes of pure water. Correspondingly, the solution of this mass of halite requires about 113 cubic kilometers of water (saline water volumes would be even greater). The presence of carbonate cap rocks indicates the involvement of additional volumes of fluid. Carbonate cap rocks form in a two-stage process: first, fluids dissolve CaSO,, thus releasing Ca2' and SO,= to solution; then, as sulfate is reduced and hydrocarbons are oxidized, carbonate precipitates. The volume of carbonate produced is roughly equal to the volume of anhydrite dissolved. Thus, for a carbonate cap rock

~~~~Data Center ,09126599985,[email protected], For Educational Uses 436 S A L‘T D 0ME MI N E K A L I2AT ION~~~~

that is 30 m thick and present in the dome described above (a total cap rock thickness of 330 m), the amount of fluid required would have been greater than 124 cubic kilometers. ‘This calculation accounts neither for the volume loss that accompanies the transformation of anhydrite to calcite, nor does it consider anything other than perfect conservation of Ca in the system. If the volume loss is considered (which is required) and part of the Ca is lost from the system (which we presume is reasonable), then the volume of fluid required to dissolve the anhydrite that ultimately became carbonate would have been higher. On the other hand, this calculation is balanced somewhat by the fact that the solubility of anhydrite increases in saline solutions, thus indicating that the required fluid volume could be lower. Overall, it seems conservative to assume that at least 124 cubic kilometers of water are required to form 30 m of calcite and 300 m of anhydrite over a 2 km diameter salt dome. For a rock with 5% porosity, this represents a rock volume roughly 2,500 square kilometers by one kilometer thick. Hence, the fluid volumes involved in cap rock generation are substantial in comparison to the volume of cap rock that forms. The f1uid:rock ratio in this case is nearly 120:l and does not consider the volume of fluid that transported the hydrocarbons for the formation of calcite. Thus, the margins of salt domes play a major role in transporting basinal fluids into the shallower section.

~~~~Conditions of Cap Rock Formation~~~~

‘I’hefollowing general sequence of events, modified from Murray (1966), results in cap rock formation. The upper part of a rising salt diapir, which contains a few percent of anhydrite crystals, dissolves as it penetrates a zone of halite-undersaturated water resulting in the accumulation of anhydrite and other minor insoluble components. Compaction and cementation of the insolubles accompanying diapiric movement results in a dense anhydrite rock. Further cycles of dissolution and residue underplating result in the formation of a banded anhydrite zone that becomes progressively younger downward (Fig. 5.11; Ulrich et al., 1984; Kyle et al., 1987). Calcite cap rock and sulfur form as a product of a complex series of bacterially mediated chemical reactions that, in the presence of a hydrocarbon food supply, alter the sulfate zones (Feely and Kulp, 1957; Sassen, 1980; Sassen et al., 1988; Prikryl et al., 1988). These hydrocarbons are oxidized, ultimately, to HCO;, and the aqueous sulfate derived from anhydrite or gypsum is reduced. These reactions result in precipitation of biogenic calcite and production of hydrogen sulfide which may be subsequently oxidized to elemental sulfur (Fig. 5.12; Feely and Kulp, 1957). Alternatively, reduced sulfur may encounter metals supplied by heated formation waters and result in precipitation of

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~~~~A B~~~~

~~~~LATE DIAPIRIC STAGE OLD~~~~

~~~~@ \ MATURE DlAPlRlC STAGE~~~~

~~~~YOUNG ......~~~~

Fig. 5.11. Model f or the origin of stratiform sulfide layers within anhydrite cap rocks showing the development of an inverted stratigraphy by underplating. A). Accumulation of anhydrite crystal residue at the anhydrite cap rock/salt diapir contact as halite is dissolved. Introduction of mineralizing solutions into this zone will form stratif orm sulfide laminae. B). Petrographic character of a stratif orm surf ide laminae after many subsequent stages of anhydrite cap rock accretion. Euhedral anhydrite crystals are preserved within the surf ide zone, whereas anhydrite crystals typically are tightly intergrown, i.e. crystal boundaries are sutured. metallic sulfide minerals (Kyle and Price, 1986; Kyle and Agee, 1988). Posey (1986), Posey et al. (1987b), Pnkryl et al. (1988), and Prkyl(1989) have shown that calcite cap rocks are composed of several genetically distinct zones and have proposed that the calcite zone forms from top to base (Fig. 5.9). The majority of the cap rock gypsum probably forms after calcite from hydration of anhydrite along the anhydrite/calcite contact by low-temperature, low-salinitywaters. Given the relatively high salinities and high temperatures that exist prior to and during cap rock formation, gypsum probably forms after all other cap rock materials including native sulfur. Anhydrite accumulates when halite-undersaturated waters reach the diapiric surface or edge. Most studies document cap rocks above the salt diapir, but it is clear that several cap rocks including calcite also occur along the margins of diapirs (see studies on Hockley dome by Price et al., 1983; Boling and Moss Bluff dome by Seni, 1987; and the general compilation by Halbouty, 1979). There is remarkably little information available on the nature of the diapir flank contacts with the enveloping sediments. Considering the high density of calcite and anhydrite and their high shear

~~~~Data Center ,09126599985,[email protected], For Educational Uses 43 8 SA L1’ D 0ME MINEK A L I Z A‘1‘I 0N~~~~

Fig. 5.12. Flow chart f or the f ormation and mineralization of calcite cap rocks. OXD = Oxidation: RDT = Reduction: SOLN = Solution: PPTN = Precipitation: HC = Hydro- carbons: Me+ = Metal cations: MeS = Metal sulfide minerals: MeSO, = Alkaline earth sulfate minerals (e.g. barite).

strengths relative to halite, diapir margin cap rocks may be more common than documented. Early formed dome-crest cap rocks commonly are carried into younger stratigraphic positions by the “rising”diapir. For example, dome-crest cap rock that began forming in late Jurassic at Winnfield dome in north Louisiana is in contact with overlying sediments as young as Eocene (Gose, et al., 1Y85; Kyle et al., 1Y87). This cap rock is clearly diapiric along with the parent salt stock. Average sulfur isotope values of anhydrite are heavier than the midJurassic mother evaporite anhydrite (Fig. 5.13; Feely and Kulp, 1957;Kyle and Price, 1986; Posey et al., lY87a). The majority of anhydrite within the halite diapirs have 6”s values ranging from + 15 to + 17%0 (CDI‘) that are identical to midJurassic seawater (as defined by Claypool et al., 1Y80). However, anhydrite cap rocks range up to 6Yoo heavier. Posey et al. (1Y87a) interpreted this to mean that a minor amount of sulfate reduction occurred during the accumulation of anhydrite cap rock. ‘I‘he unreduced sulfate, which was ”S enriched by partial sulfate reduction, then re-precipitated along with the cap rock,

~~~~Data Center ,09126599985,[email protected], For Educational Uses CONDITIONS OF CAP ROCK FORMATION 439~~~~

~~~~I I I I ANHYDRITE, I (HJ) SALT RESIDUE m (B,II)~~~~

~~~~ANHYDRITE, CAP ROCK~~~~

~~~~GYPSUM, CAP ROCK I~~~~

~~~~MARCASITE / PYRITE (H,17)~~~~

~~~~SPHALERITE (H,II)~~~~

~~~~GALENA~~~~

~~~~SULFUR~~~~

~~~~BARITE (H,5)- (S,2 SELENITE W (H,3) I~~~~

Fig. 5.13. Generalized sulfur isotopic compositions of Gulf Coast salt dome cap rock materials. H = Hockley dome (HS = sulfidesfrom overlying andflanking sediments); B = Boling dome; M = Moss Bluff dome; S = Spindletop dome. Number of analyses are in parenthesis; vertical bur is the mean. Modified after Kyle and Price (I 986). resulting in cap rocks with a higher 6% ratio. Geochemically, the formation of calcite cap rocks is much more complex than anhydrite. Carbon and oxygen isotope profiles have been used to show that the fluids responsible for calcite cap rock formation consist of both formation fluids and meteoric water and, possibly, seawater (Posey, 1986; Posey et al., 198%; Prikryl et al., 1988; Prlkryl, 1989). Formation fluids containing oil (Feely and Kulp, 1957; Sassen, 1980; Sassen et al., 1988) and, locally methane (Posey, 1986; Posey et al., 198%) supply the bulk of the carbon, whereas meteoric water supplies the bulk of the dissolved oxygen.

~~~~Data Center ,09126599985,[email protected], For Educational Uses 440 SALT DOME MINERALIZATION~~~~

The carbon isotopic composition of calcite from several Gulf Coast salt dome cap rocks shows a range of depleted 613Cvalues that generally are interpreted to indicate hydrocarbon sources (Feely and Kulp, 1957; Posey, 1986). Calcite cap rocks typically have lighter with depth carbon isotope profiles ranging in most domes from about -10 to -30%0 (PDB) or, in a few domes, from about -10 to about -5@?o (Posey, 1986). Calcite cap rocks form in a general downward-younging sequence, and the carbon isotope profile owes either to evolution of a single (petroleum) source (Feely and Kulp, 1957; Sassen, 1980; Kreitler and Dutton, 1983) or, more probably, mixing between two end members (Posey, 1986; Posey et al. 1987b; Prlkryl et al., 1988; Sassen et al., 1988). Carbon isotope values lighter than about -30%0 indicate derivation from methane; those between -30 and about -15%0 indicate a liquid petroleum source, and values heavier than about -5Yoo indicate a direct marine source. Although the source of methane is generally regarded to be derived from deeply buried hydrocarbon source rocks, a biogenic methane source has not been ruled out. Likewise, methane found within salt diapirs (see previous discusssion) rather than in diapir flank sediments, has not been considered as a source of cap rock carbon. Oxygen isotope values range from about -4to -1 l%o (PDB) and profiles in calcite cap rocks are mixed (Posey, 1986); some grow lighter with depth, whereas others grow heavier with depth, and some are mixed, showing excursions within the profile (Prikryl, 1989). The range of values is due, in a slight degree to variations in temperature of formation, but more importantly to the oxygen fluid source. Meteoric fluids supply light oxygen, whereas formation fluids supply heavy oxygen. The strontium isotopic compositions of cap rock components from several Gulf Coast domes suggest the presence of fluids from several sources during cap rock formation, including fluids from the Louann Formation, at least two different brines, and meteoric waters (Posey, 1986; Posey et al., 1987~).Cap rocks in the East Texas and North Louisiana basins have 87Sr/86Srratios similar to Jurassic seawater (0.7068-0.7076). However, some salt dome cap rocks from the Gulf Coast province, including Hockley dome, have very high values. Similarly, barite and associated calcite from the Winnfield cap rock have high 87Srp,6Srratios (Posey, 1986). Strontium in salt dome cap rocks is buffered by strontium from the anhydrite. Anhydrite, calcite, and gypsum cap rocks, all of which form after the salt-hosted anhydrite, have the same range of strontium isotope values as found in the salt-hosted anhydrite itself (Posey, 1986). Although the salt and associated anhydrite are considered to be of marine origin, and even though the lowest strontium isotope values match that of mid-Jurassic seawater, higher values, which are found in virtually all domes, indicate interactions with radiogenic strontium of non-marine origin (Posey, 1986; Posey et al.,

~~~~Data Center ,09126599985,[email protected], For Educational Uses ENERGY RESOURCES 44 1~~~~

1987a; Land et al., 1988). Apparently, strontium-bearing fluids invade the salt prior to cap rock formation. This information coupled with studies of bromide concentrations in bedded and diapiric salt (Land et al., 1988) indicates that the salt recrystallized in the presence of a radiogenic strontium-bearing fluid during diapirjsm, leaving the salt-hosted anhydrite with strontium isotope values ranging from midJurassic seawater (0.7068) to very high values (0.7100). Cap rock materials that formed subsequently have either the same or, possibly, slightly higher 87Sr/86Srranges (Light et al., 1987).

~~~~SALT DOME MINERAL RESOURCES~~~~

~~~~General~~~~

The most economically significant mineral resource of the salt dome environment is petroleum. Various industrial rocks and minerals have been produced from salt domes, including salt, potash, limestone, gypsum, anhydrite, and sulfur. Zinc, lead, silver, and manganese sulfides, barium and strontium sulfates, and many other minor minerals occur in many domes along with the iron-sulfides pyrrhotite, pyrite, and marcasite (Kyle and Price, 1986). Zn and Pb are mined from cap rocks and associated sediments in northwestern Africa. In this section we describe selected salt dome mineral resources and focus on processes responsible for their concentration.

~~~~Energy Resources~~~~

The porous calcite cap rock zone not only requires petroleum for its formation, but also provided a reservoir for spectacular, but short-lived, oil production in the early history of the Gulf Coast, including the famous Spindletop gusher that issued in the "age of liquid fuel" in 1901 (Tyler et al., 1985). However, the relationship between sediment loading and halokinesis that results in complex sedimentary f acies and related structural effects is more significant to the development of the major oil reservoirs in the salt dome environment. Variations in thickness and syndepositional facies characterize the near-dome strata that form during salt flow. Seni and Jackson (1984) summarized these relationships as a framework to understand the development of known and potential hydrocarbon traps during the halokinetic evolution of the East Texas Basin (Fig. 5.14). The initial stage of pillow growth caused broad crestal uplift so that syndepositionally and postdepositionally thinned strata overlie the pillow crest. Fluvial-deltaic strata deposited over the crest of salt pillows are sand poor but are likely to be flanked by stratigraphic pinch-outs of sandy reservoirs. Sand-rich fluvialchannel systems

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~~~~EXPLANATION SlLlClCLASTK FLUVIAL, DELTAIC, WSTDIAPIR AND SLOPE SYSTEMS STAGE a Mudslone sond.1-~~~~

~~~~CARBONATE AN0 SlLlClCLASTlC SHELF SYSTEMS 63 E2~,2:OTIMi Corbonoles rllh enh0honc.d~~~~

~~~~SlLlClCLASTlC FLUVIAL, DELTAIC, AND SLOPE SYSTEMS I CARBON4TE AND SlLlClCLASTlC SHELF SYSTEMS~~~~

~~~~DEPOSITIONAL SYSTEMS m991~~~~

Fig. 5.14. Schematic cross section through a mature salt diapir showing typical f’acies variations and potential petroleum traps (numbered) in siliciclasticf’luvial, deltaic, and slope depositional systems and in carbonate and siliciclastic shelf’ depositional systems. From Seni and Jackson (1984). bypassed pillow crests and occupied adjacent primary peripheral sinks. Under marine conditions, topographic swells over pillows were sites favorable for reef growth, high-energy grainstone deposition, and sand concentration by winnowing; these depositional environments form potential hydrocarbon reservoirs. Primary peripheral sinks formed preferentially updip of the salt pillows because of greater salt flow in the pillow from the updip side. Structural reversal during diapirism generally transforms a primary peripheral sink into a turtle-structure anticline (Fig. 5.14). Thus, the location of a primary peripheral sink establishes the core of the subsequent turtle-structure anticline, generally 5 to 20 km updip from the dome crest in the East Texas Basin. This relationship is a valuable exploration guide for one of the most important salt-related structural traps, especially for the deeper, less explored horizons in this basin (Seni and Jackson, 1984). Large secondary peripheral sinks enclosed or flanked the rising diapir. In East Texas, marine strata filled these secondary peripheral sinks and represent thickened, but otherwise normal, low-energy sequences. Because secondary peripheral sinks represented local sites of greater subsidence and hence were depressions, they were more likely to preserve marine sand bodies formed during transgressive reworking. These pinch-outs of marine sand bodies can subsequently act as subtle hydrocarbon traps (Fig. 5.14). Seafloor mounds over diapirs may become petroleum reservoirs because, as with pillows, they were sites of reef growth, grainstone deposition, and sand

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concentration by winnowing. However, these supradomal mounds were much smaller than analogous suprapillow swells. Furthermore, they commonly are destroyed by further uplift, erosion, and salt emplacement. Raised saddles between secondary peripheral sinks may provide another favorable site for reef growth in younger environments; this type of Early Cretaceous depositional setting provided the stratigraphic trap for the giant Fairway Field in the East Texas Basin (Seni and Jackson, 1984). Postdiapir growth typically has only minor effects on surrounding strata. Mounds over domes undergoing postdiapir growth may deflect fluvial-channel systems around supradome areas, so that mud-rich sediments are deposited over the diapirs. Differential subsidence may cause these fluvial-channel sandstones to stack vertically in tertiary peripheral sinks. Subtle petroleum traps formed during this stage generally are much smaller than those formed during earlier stages of diapirism. Salt diapirs may also be responsible for the formation of critical genetic factors for the concentration of non-hydrocarbon types of energy resources. Uranium, produced from open-pit mining and in situ leaching operations in the South Texas Uranium District, is a major mineral resource of the Gulf Coast. Genetic concepts for these "roll front" deposits include sulfidization of detrital iron and titanium oxides by H,S which was supplied along growth faults from deeper hydrocarbon reservoirs (Goldhaber et al., 1978,1983). This alteration zone provides a reducing environment for the precipitation of uranium carried in oxidizing ground waters. The common association of H,S with salt dome hydrocarbon reservoirs has resulted in widespread sulfidization of detrital oxides in the contiguous sedimentary units; this feature has resulted in a widely tested exploration model for roll front uranium deposits surrounding salt domes. Significant uranium has been produced from superdome sedimentary units at the Palangana dome in south Texas (Weeks and Eargle, 1960). Thus, salt dome-associated iron sulfide concentrations provide a variation of the critical reductant for roll front uranium deposits in the Gulf Coast and elsewhere.

~~~~Sulfur Deposits~~~~

Despite the economic importance of salt dome native sulfur (SDNS) deposits, little has been written about their geologic character. Thode et al. (1954) and Feely and Kulp (1957) described the isotopic characteristics of SDNS and developed a theory for their formation. Ruckmick et al. (1979) classified native sulfur deposits and speculated on the various chemical routes taken to produce native sulfur from H,S. Subsequently, Price et al. (1983) incorporated several sulfur isotope values of sulfur into their study of

Data Center ,09126599985,[email protected], For Educational Uses 444 SALT DOME MINERALIZATION the Hockley dome sulfide mineralization. Although much data exist on the isotopic composition of SDNS, most of the geological observations deal with the host rocks and associated sediments. Even the paragenetic position of sulfur, particularly with respect to gypsum, has not been established. A systematic petrographic and geological investigation of SDNS deposits has never been reported. The following discussion, which is based on our personal, albeit limited, studies of SDNS geology and the available literature, is an initial attempt to synthesize the geological and geochemical information and to underscore the information deficiency regarding these important deposits. Native sulfur deposits in salt dome cap rocks seem to occur most commonly within the lower part of the calcite cap rock and in the upper part of the anhydrite cap rock, although some native sulfur also occurs within shallower parts of the calcite cap. This lithologic relationship has been quantified for a few domes, notably Boling dome and Moss Bluff dome, Texas (Fig. 5.15; Seni, 1987), and we have observed the same situation at the New Gulf and Damon Mound domes in Texas. Where present, gypsum occurs most commonly at or near the calcite/anhydrite contact and in close association with the sulfur. However, the major cap rock-hosted sulfur deposit at Boling Dome, located a few kilometers from New Gulf and with a similar cap rock at a similar depth, lacks appreciable gypsum (Seni, 1987). Because gypsum generally occurs as a cement in both the calcite cap rock and anhydrite cap rock rather than as a discrete gypsum layer, this unit commonly is called the transition zone (see Seni, 1988, for an example). Salt dome cores from the Long Point dome in Texas have sulfur that is paragenetically later than gypsum. Sassen et al. (1988) documented native sulfur later than calcite at Damon Mound dome, but the relation to gypsum is not known. Price et al. (1983) placed sulfur as the latest mineral of their general paragenesis, but this placement was clearly based on observation of minor sulfur occurrences in a sulfur-poor cap rock. We also generalize that sulfur is the latest mineral to form in a SDNS deposit. Although gypsum may form later than sulfur, the gypsum-precipitating solutions will probably dissolve sulfur. Most authors conclude that native sulfur forms as a result of the oxidation of hydrogen sulfide (Thode et al., 1954; Feely and Kulp, 1957; Ivanov, 1968; Ruckmick et al., 1979). Feely and Kulp (1957) and Ruckmick et al. (1979) assumed that H,S was oxidized by excess SO,=.However, Davis et al. (1970) showed that this reaction would proceed only under acidic conditions. Davis et al. (1970) showed further that oxygen, probably dissolved in groundwater, was a thermodynamically more viable oxidizer of H2S, and evaluated the reaction rates under lab conditions. This reaction has credence in that the reaction is charge balanced and that there is no reported evidence of acid attack of the substrate minerals on which the sulfur forms, although we hasten to add

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~~~~a Surface a'~~~~

~~~~Miocene and Younger Siliciclostics - 100~~~~

~~~~v) -200- aW I- W I -300-~~~~

~~~~++++++++++++++ Moss Bluff Dome' + Cop Rock Lithofacies -400- 0 zoom -VE:2 Sulfur-Wring zone OSolt stock - 500 1 ,-- I'. ,__---~~~~

~~~~0 loo0 rn - 400 u~~~~

Fig. 5.15. Geologic features of sulfur deposits in Gulf Coast salt dome cap rocks. A). Structure contour map in meters on the Boling dome cap rock showing location of sulfur and petroleum reservoirs. B). Generalized geologic cross section through the Moss Bluff dome cap rock showing location of the sulfur-producing zone relative to cap rock lithofacies. Modified after Seni (1987).

~~~~Data Center ,09126599985,[email protected], For Educational Uses 446 SALT DOME MINERALIZATION~~~~

that a systematic investigation has yet to be carried out. We feel that oxidation of H,S by groundwater oxygen is the most plausible mechanism for precipitating native sulfur, but it requires several physical and chemical constraints. There must be a physical barrier to gas migration in order to trap the H,S. Without a physical trap, the gas would eventually escape and a significant sulfur concentration would not form. Groundwater must be able to enter this environment as well as exit. Oxygen in the groundwater must contact the H,S, but a significant supply of oxygen must be available in order to form a sulfur deposit of appreciable size. At some point, the groundwater must become saturated with S, or the native sulfur crystals would redissolve. Native sulfur can dissolve freely in groundwater, as can be seen from historical examples. For example, native sulfur that was present in limestone exposures in West Texas in the early part of this century has now been dissolved by rainwater (G. Wessel, pers. comm., 1989). This requirement further suggests that groundwater flow into the deposit must cease at some point, either as a result of reservoir sealing by the sulfur, or perhaps by another mineral, most probably gypsum. The relationship has not been documented and should be tested. Overall, the general Occurrence of SDNS deposits in the lower parts of calcite cap rocks lends credence to the model of hydrogen sulfide oxidation by groundwater and accumulation beneath a gas seal, as discussed above. The upper parts of most salt dome calcite cap rocks consist of more tightly cemented calcite lithotype than the lower parts (Posey, 1986; Posey et al., 198%). This denser calcite zone could behave as a seal, allowing H,S to accumulate in the porous zone below. At some point, calcite cap rock growth is arrested, most probably when the hydrocarbon supply is exhausted, and groundwaters invade along the most permeable routes. The fluid migration routes, in this case, just happen to be the lower part of the calcite zone, which is also where H,S has accumulated. Thus, native sulfur accumulates in that part of the cap rock, or nearby. It is probably no coincidence that gypsum also forms here, because this porous zone close to the anhydrite is the most likely place for gypsum to begin filling open voids.

~~~~Metallif erous Deposits~~~~

General. Cap rocks are best known for their great native sulfur concentrations, but a few in northwestern Africa have commercial concentrations of metal sulfides. Recent exploration results suggest potential for commercial metal concentrations in the Gulf Coast salt domes (Price et al., 1983; Kyle and Price, 1986). At least two Zn-Pb deposits associated with Triassic age diapirs are mined in Tunisia (Fig. 5.3), and Zn-Pb sulfide Occurrences associated with diapirs in northern Africa, Spain, and France suggest

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that the metallic mineral potential is very widespread (Rouvier et al., 1985; Charef and Sheppard, 1987; Montacer et al., 1988; Orgeval et al., 1989). Although some of these cap rock-hosted Zn-Pb-Ag-Ba concentrations are thought to be genetically related to the Mississippi Valley-type (MVT) Zn-Pb-Ba and sedimentary exhalative Zn-Pb-Ag-Ba deposits that typically occur in older strata (Price et al., 1983; Charef and Sheppard, 1985; Kyle and Price, 1986), and even though similar mineralization in associated sedimentary strata at Bou Grine, Tunisia, (Montacer et al., 1988; Orgeval et al., 1989) support this affiliation, there are enough differences in the mode of origin and fluid drive mechanisms to allow the classification of salt dome mineral deposits as a distinct group. The cap rock-associated minerals, notably native sulfur, celestite, and strontianite, further support consideration of the salt dome deposits separately from the MVT or Sedex deposits. Metalliferous oil field brines of the type that have been well documented in the Gulf Coast (Carpenter et al., 1974; Land and Prezbindowski, 1981; Kharaka et al., 1977, 1987; Macpherson, 1989) have been widely accepted as modern analogs of the ore-forming fluids for MVT and Sedex deposits (Anderson and Macqueen, 1982; Svejensky, 1984; Gustafson and Williams, 1981; Sawkins, 1984), and existing data indicate that the ore solutions responsible for salt dome cap rock mineralization are generally the same. However, the sulfate reduction mechanisms, the temperature of formation and the fluid flow mechanisms may be different. Additionally, there is ample evidence that salt dome cap rock mineralization is more closely related to large scale oil and gas migration than either the MVT or Sedex deposits. The presence of metalliferous formation brines in the Gulf Coast suggests the potential for commercial metallic concentrations originatingfrom these fluids or similar geologically young brines. Metallic minerals have long been identified in the cap rocks of Gulf Coast salt domes (Table 5.1; Hanna and Wolf, 1934), but typically had been regarded as mineralogic curiosities. Sphalerite and galena have now been identified at 17 domes; numerous other domes contain locally abundant iron and manganese sulfides, barite, celestite, sulfur, and other "exotic" minerals (Kyle and Price, 1986). Therefore, mineralization in the salt dome environment appears to be a common process and provides an unusual example of the relationship of evaporites with concentrations of metallic, nonmetallic, and energy resources (see chapter 6, this volume, for a general review of this topic). The only cap rock sulfide deposits that have been systematically evaluated to date are the Hockley dome within the Houston diapir province in south central Texas (Price et al., 1983; Kyle and Price, 1985, 1986; Agee, 1990), the Fedj el Adoum deposit (Rouvier et al., 1985; Charef and Sheppard, 1987) and the Bou Grine deposit (Orgeval

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~~~~Table 5. I. Geologic characteristics and commercial significance of selected Gulf Coast salt domes. CAPROCK DEPMTO: COMMERCIAL PROWCTION SALT DOME BASIN CAPROCK SALT SULFUR OTHER~~~~

~~~~~ . -ImL-ImL~~~~

~~~~Barbers Hill, Gc 73 105 305 17 0 107 NR ER,SF Texas~~~~

~~~~Belle Isle. Gc 36 30 10 22 20.524 NR HL Louisiana~~~~

~~~~Boggy Creek, ET 17 2 NR 560 09 3 NR Texas Eohng. Gc 33 7 115 290 47 NR 80 7 SF Texas~~~~

~~~~Cailbu Isle, Gc 40 760 835 71 9 16.908 NR Louisiana~~~~

~~~~Cinco Presidenles. IS NR NR 2,000 28 8 9 NR Tabasco~~~~

~~~~Damon Mound, Gc 54 0 115 29 9 0.1 LS Texas~~~~

~~~~Encutla. MS NR 3,595 3,600 55 NR w Mississippi~~~~

~~~~Grand Saline, ET 54 50 65~~~~

~~~~High Island. Gc 44 45 370 17 5 17 02 Texas Hockley, Gc 85 20 300 01 NR NR K Texas~~~~

~~~~Humble, Gc 14 3 210 370 22 2 NR NR Texas~~~~

~~~~Jalpan. IS 33 4 100 300 NA NR 30.2 Veracruz Lake Washington, Gc 10 0 335 480 27 5 9.578 41.0 Louisiana~~~~

~~~~Palangana. Gc 34 115 150 41 22 0.3 ER,U Texas~~~~

~~~~Swlh Pass, Gc N) NR 2.860 36 2 6,964 NR Louisiana~~~~

~~~~Spindlelop. Gc 18 210 360 20 7 17 10.0 BR Texas~~~~

~~~~Sulphur Mines, Gc 05 120 440 39 32 9.6 ERSF Louisiana~~~~

~~~~Weeks Island, a: 79 0 15 27 6 5,788 NR ER.HL Louisiana~~~~

~~~~Wnnlield. NL 11 0 105 NR N) NR LS.AN.HL Louisiana EXPl ANATDN~~~~

~~~~Data Center ,09126599985,[email protected], For Educational Uses METALLIFEROUS DEPOSITS 449~~~~

et al., 1989; Montacer et al., 1988) both in Tunisia. At Hockley, exploration drilling has defined a multi-million ton deposit of Zn-Pb-Ag sulfides. Initial economic evaluation, which was based on composite core samples averaging 3% Zn, 1% Pb, and 20 ppm Ag, and which is believed to be representative of the massive sulfide concentrations in the calcite cap rock, indicate that there are 5.5 million tons at 7.1% Zn + Pb or 12 million tons at 4.2% Zn + Pb (Kyle and Price, 1986). This deposit compares in size and overall mineralogy with the Bou Grine deposit in northern Tunisia which contains 7.3 million tons averaging 2.4% Pb and 9.7%Zn (Orgeval, et al., 1989). A brief discussion of each of these deposits follows. Hockley Dome, Texas, U.S.A. Hockley dome is elongate northwest-southeast with dimensions of 4.4 km by 3 km at a depth of 200 m (Fig. 5.16A). Although the cap rock is within 12 m of the surface, there is no topographic expression of the dome. The cap rock consists of a typical sequence of calcite, gypsum, and anhydrite zones totalling as much as 285 m (Fig. 5.16B). The calcite zone is highly fractured and brecciated with vugs containing calcite, barite, and sulfur crystals; petroleum is present locally in the calcite zone and within a small oil field in an overhang trap on the southwestern side of the diapir (Fig. 5.16B). The highest sulfide concentrations occur around the periphery of the Hockley cap rock at depths ranging from 120 to 365 m (Price et al., 1983; Kyle and Price, 1985,1986; Agee, 1990). Anomalous amounts of sulfides occur in both calcite and anhydrite hosts, and sulfide-rich zones appear to be continous between the two lithologies in some localities (Fig. 5.16B). Marcasite, pyrite, sphalerite, and galena range from trace components to local massive sulfide zones; significant zinc + lead drill intercepts include 5.5 m of 7.1% and 12 m of 4.2%. Sphalerite is the dominant economic mineral with an overall Zn/Pb ratio of about 3:l. Metal distribution studies suggest that there is a trend of increasing zinc/lead downward in the anhydrite cap rock (Kyle and Agee, 1988). Silver is associated with major base metal zones in amounts generally less than 20 ppm. The most significant silver concentrations occur stratigraphicallylow within the anhydrite cap rock zone and are independent of major base metal concentrations. A maximum silver analysis of 550 ppm over 0.3 m occurs within a 2-m intercept averaging over 225 PPm. The zonal distribution and relative abundance of identified minerals at Hockley dome were discussed by Kyle and Price (1985,1986). Marcasite and pyrite are the most common accessory minerals in the Hockley dome cap rock, ranging from thin laminae typical of mineralization at the salt/anhydrite contact to local massive sulfide concentrations (Fig. 5.17A-C).The iron sulfides locally form the major portion of the cap rock section over as much as a few meters; high sulfide intercepts in the calcite cap

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Fig. 5.16. A). Structure contour map showing depth to cap rock in meters at Hockley dome. Texas. Modified from Kyle and Price (1985). B). Geologic cross section and histograms of major Zn + Pb intercepts along the H drill fence at Hockley dome. rock typically are brecciaform (Fig. 5.17B,C). Aggregates of fine-crystalline sphalerite may be intergrown with other sulfides or form monominerallic colloform crusts (Fig. 5.17E,F). Hockley sphalerites have unusual trace element compositions, generally with low concentrations of iron, but with variable enrichment of cadmium and/or silver. Microprobe analysis suggests that the relative proportions of Fe-Cd-Ag substituting for Zn controls the luminescent properties of sphalerite (Kyle and Price, 1985). Some of the sphalerite types from Hockley and other domes have brilliant yellow to reddish-orange fluorescence (Fig. 5.17F). Galena typically forms relatively large (0.1 to 2 mm) crystals that may be intergrown with sphalerite. Acanthite is locally associated with galena and appears to be responsible for the higher grade silver assay intervals. Hauerite, a relatively rare manganese sulfide mineral, occurs as small isolated octahedral crystals. Larger crystals of hauerite and alabandite have been reported from other Gulf Coast cap rocks (Table 5.1; Hanna and Wolf, 1934). Barite is a locally abundant component

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Fig. 5.17. Geologic features of mineralized cap rock core from the Hockley, Texas, dome. A). Anhydrite cap rock showing closely spaced sulfide laminae typical of mineralization at the salt/anhydrite contact. B). Brecciated iron sulfides and sphalerite from the interior calcite zone on the southern flank of the dome (Fig. 5.168). Note subhorizontal sulfide laminae and sulfide clasts inherited from the stratiform sulfide laminae in precursor anhydrite cap rock (e.g. Fig. 5. I7A). C). Brecciaform massive iron sulfides with interfragmental sphalerite frotn the interior calcite zone. D). Massive barite with irregular iron sulfide laminaefrom the interior calcitezone. E). Sphalerite-cemented anhydrite cap rock overlain by tan collof’orm sphalerite and rainose rnnrcasite. Field of view = S cm. F). Broad band fluorescence photomicrograph of yellow-f luorescing colloform sphalerite cementing anhydrite residue which represents stratiform mineralization at the salt/anhydrite contact. Field of view = 2 mm.

~~~~Data Center ,09126599985,[email protected], For Educational Uses 45 2 SALT DOME MINERALIZATION~~~~

of the calcite cap rock (Fig. 5.17D) and also forms concretions in unconsolidated superdome sediments. Although barite distribution has not been systematically evaluated, the average barite content of several 5-m intervals associated with sulfide concentrations is about 30 wt %. Sulfur crystals occur locally in late fractures and vugs in the calcite cap rock. The paragenetic relationships among the metallic minerals and host rocks at Hockley and other Gulf Coast salt domes are complex. Generally the iron sulfide minerals appear to have been the first sulfides to have precipitated, followed by sphalerite and galena, with the more oxidized sulfur species, such as elemental sulfur, manganese sulfides, and the sulfate minerals, forming later. However, it is readily apparent that there are many repetitions of this general sequence. The stratiform sulfide layers within the anhydrite cap rock are believed to represent mineralization at the salt/cap rock interface (Ulrich et al., 1984) and are therefore syngeneticwith regard to the enclosing host rocks. Euhedral crystals may form simple open-space fillings of late fractures or vugs within the calcite cap rock in some cases; in other cases, earlier generations are brecciated and cemented by later sulfides or calcite (Kyle and Price, 1985). Calcite may also be intimately intergrown with sulfide minerals. Therefore, calcite precipitation, i.e. calcite cap rock formation, can pre-date, be contemporaneous with, or post-date local sulfide precipitation. The complex textural relationships among sulfide minerals and their host rocks are believed to be a function of an evolving geochemical system in which earlier mineral products are overprinted as the result of later halokinetic and fluid events. Fedj el Adoum and Bou Grine, Tunisia. The Fedj el Adoum diapir is a northeast trending elongate diapir located in the approximate center of the Tunisia-Algeria diapir belt, about 45 km northeast of the city of Le Kef, and about 105 km west-southwest of Tunis (Fig. 5.3; Rouvier et al., 1985). The Bou Grine deposit, located along the Lorbeus diapir, occurs about 20 km southeast of Le Kef (Rouvier et al., 1985). The diapir belt is a well-defined northeast-trending zone of folded Mesozoic to Neogene sedimentary rocks containing Triassic evaporite diapirs. The diapir belt is flanked to the north by the Atlas Mountains thrust belt which contains Triassic evaporites in nappes and to the southeast by less intensely folded, diapir-poor, mostly Mesozoic sedimentary units. Like the Triassic evaporite diapirs and nappes in general, the diapirs at Fedj el Adoum and Lorbeus consist mostly of gypsum, dolostone, and variegated, marly clay with minor accompanying sandstone, anhydrite, and mafic volcanic segregations (Pervinquiere, 1903; Perthuisot, 1977; Orgeval et al., 1989). The majority of the identified diapirs have breached the surface. Although these diapirs are considered to

~~~~Data Center ,09126599985,[email protected], For Educational Uses METALLIFEROUS DEPOSITS 45 3~~~~

be saliferous (see Rouvier et al., 1985), they do not contain halite at the surface. Surface halite is preserved at only one diapir in central Tunisia where rainfall is extremely slight (A. Abbes, S. Bouhlel, 1988, pers. comm.). Mineralization in these diapirs consists of sphalerite, pyrite, galena, barite, and celestite, although only Zn and Pb are recovered. Mineralization occurs as replacement and open space filling in both dolostone and limestone at Fedj el Adoum and in these lithologies plus dark brown shale at Bou Grine (Orgeval et al., 1989; Rouvier et al., 1985). Between the Triassic evaporites and Cretaceous marls and shales, especially at Fedj el Adoum, is a dolomitic limestone unit of uncertain age referred to locally as 'Transition Zone" or "X-Carbonate" (Laatar, 1980; Rouvier et al., 1985). This unit appears, from rock descriptions and stable isotope studies, to be similar to carbonate cap rock described commonly in the Gulf Coast. A black calcite in this zone contains segregations or bands of fine- and coarse-crystallinewhite calcite and celestite, and the rock is brecciated near the diapir contact (Rouvier et al., 1985). Carbon and oxygen isotope values from the 'Transition Zone" calcite are identical in range to those of the "detritus-bearing false calcite" cap rocks in the Gulf Coast compiled by Posey (1986). Detritus-bearing false calcite cap rock is the same as variegated calcite cap rock of Pnkryl et ai. (1988), and consists generally of those calcite cap rocks in the upper part of the calcite cap rock section that contain little to no banded "zebra-rock" textures. Carbon isotope values range from about +8 to -26960 (PDB), and oxygen isotope values range from about -3 to -@!o (PDB) (Rouvier et al., 1985). The oxygen isotope compositions suggest at least two carbon sources, petroleum and marine bicarbonate, and the oxygen isotope values suggest a mixture of formation water and/or marine water with meteoric water. Hydrogen isotope values from calcite and ore mineral fluid inclusions range from about -20 to -1 18/00 (SMOW); oxygen isotope values from fluid inclusions range from about + 10 to +Woo (SMOW). The lighter values are associated generally with the lighter carbon isotopes and heavy oxygen isotopes from calcite (Charef and Sheppard, 1987). Charaf and Sheppard (1987) interpreted these relations to mean that both light hydrogen and light carbon were derived from organic matter (petroleum) during bacterial sulfate reactions. The light carbon would have a composition of liquid petroleum (ca. -25%0) and the oxygen which is supplied by formation waters, would have a heavy value, but the generalvalue of + 1Woo which is associated with the lighter carbon and hydrogen isotope values, is the lightest of the oxygen isotope range. This indicates that at the time these minerals precipitated, there was a more significant component of meteoric water than at other times. The temperature range over which calcite of this

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composition formed is cool enough to be compatible with bacteriogenic sulfate reduction is 70 to 90°C. The hydrogen isotope composition of a single gypsum fluid inclusion (-43%0) is identical to modern meteoric water in the area (Rouvier et al., 1985), indicating, as would be expected, that the gypsum formed in a cool setting. Fluid inclusion analyses of sphalerite indicate that sulfides formed over a range of temperatures and salinities from an early, warm (> 170°C) saline (ca. 23 eq. wt. % NaCl) through intermediate temperatures and salinities to late, cool (~70°C)low salinity (12 to 15 eq. wt. % NaCl) (Rouvier et a]., 1985). Overall, this range of values trend from warmer, more saline values having heavier carbon isotope numbers near the diapir contact to cooler, less saline and lighter carbon isotope numbers higher in the section. The isotope, salinity, and temperature trends at Fedj el Adoum are similar to that shown by Prikryl et al. (1988) for the Damon Mound calcite cap rock where carbon isotope values decrease with depth (toward the diapir) and oxygen isotope values increase. In both settings, the cap rocks appear to have formed from a mixture of meteoric and formation waters. However, the mineral concentrations do not follow this trend at Fedj el Adoum. Instead, it shows the progressive influence of formation fluid away from the diapir contact rather than toward it. Mineralization at Bou Grine is similar in some ways to that at Fedj el Adoum, with the major exception that most of the Bou Grine sphalerite is contained in an organic-rich shale adjacent to the diapir, the Cretaceous (Cenomanian-Turonian) Bahloul Formation. Transition zone mineralization is minor (Orgeval et al., 1989; Montacer et al., 1988). According to Orgeval et al. (1989), the Bahloul Formation may have served as both the petroleum source rock, the metals source rock, and the petroleum and metals conduit for mineralization in the Bahloul as well as in the adjacent mineralized Cenomanian and Transition Zone (i.e. cap rock) deposits. No other deposit is known to have diapir-related mineralization of this magnitude in the sedimentary units adjacent to the diapir. If units like the Bahloul serve as both source and conduit, this factor limits the volume of metalliferous fluid that can flow into the formation. Fluids that form the more classic MVT deposits, which are of more regional extent and much longer lived than fluids within a single compacting shale formation, supply practically unlimited fluid and, therefore, metal volumes as they probably rely on topographically driven fluid flow. The amount of metal that may be derived from an organic source rock, if traveling in the same fluid as petroleum, is limited to the amount of metal within the individual petroleum source rock unit. However, regional flow systems, which are topographically driven, may inherit metals from any of the formations through which they travel. Thus,

~~~~Data Center ,09126599985,[email protected], For Educational Uses ORIGIN OF CAP ROCK METALLIFEROUS DEPOSITS 455~~~~

it seems reasonable that the longer lived, regional, topographically-drivenflow systems have greater opportunity to produce large metal deposits that the more local, compaction-driven flow systems through organic source rocks.

~~~~Origin of Cap Rock Metallif erous Deposits~~~~

Our model for salt dome metals mineralization is biased toward the geologic situation at the Hockley, Damon Mound and Winnfield domes in the Gulf Coast (Price et al., 1983; Kyle and Price, 1985, 1986; Ulrich et al., 1984; Gose et al., 1985; Kyle et al., 1987; Posey et al., 1987; Kyle and Agee, 1988; Prikryl et al., 1988), and toward the Fedj el Adoum and Bou Grine deposits in northwestern Africa (Charef and Shepherd, 1987; Rouvier et al., 1985; Orgeval et al., 1989; Montacer et al., 1988). This model incorporates carbon, oxygen, strontium, and sulfur isotope data and metal concentration data from numerous sources mentioned throughout the text. The ores and mineral Occurrences appear in several components of the salt diapir system, including calcite cap rock (both detritus-rich and true calcite cap), anhydrite cap rock, gypsum (rarely), and in coarse- and fine-grained siliciclastics alongside the domes. Mineralization associated with epigenetic calcite having a salt dome cap rock type geochemistry extends as far as 1500 m from the West Hackbeny Dome in southwest Louisiana (McManus and Hanor, 1988) and is probably characteristic of the extent of circulating fluids around the domes. The major metals include Fe, Ba, Zn, Pb, and Sr, with locally significant Ag. Sr mineralization is very common in north African salt domes but, with a few exceptions (see Saunders et al., 1988), Sr minerals occur only in trace abundances in Gulf Coast domes. On the other hand, Ag is present in several Gulf Coast domes (Kyle and Price, 1986), but has not been reported from the domes in northwestern Africa. Petrographic studies indicate that some sulfide concentrations were emplaced into fracture and intergranular pores within a consolidated and compacted, dominantly calcite cap rock (Price et al., 1983; Kyle and Price, 1985). However, sulfides also Occur as cement around undeformed anhydrite grains and represent mineralization at the salt/anhydrite contact before local consolidation of the anhydrite zone (Ulrich et al., 1984; Kyle et al., 1987). Sulfide mineralization at the Bou Grine deposit occurs in three dominant modes (Orgeval et al., 1989; Montecer et al., 1988): (1) as replacement sulfides in a carbonate 'Transition Zone" between the diapiric Triassic evaporites and adjacent sediments; (2) as stratiform replacement fine disseminationswithin the Cenomanian-Turonian Bahloul Formation, a black, organic-rich shale and; (3) as a multiply-brecciated and mineralized

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carbonate and marly unit that cuts across other mineralized units. It is our opinion that the 'Transition Zone" is actually cap rock and that mineralization within the Bahloul Formation very nearly approximates a true sedimentary exhalative deposit. In the Gulf Coast deposits, sphalerite is typically very finecrystalline, whereas it appears to be more coarsecrystalline in the Tunisian deposits. The finecrystalline nature of the cap rock sphalerite in the Gulf Coast deposits has prevented direct fluid-inclusion investigations of the nature of the ore-forming fluids for the main stage of sulfide mineralization. Barite and celestite crystals in late vugs are the only materials that have provided information somewhat relevant to the main stage of mineralization; limited data for these fluid inclusions suggest precipitation over a range of 110-140°C from solutions ranging from 3-12% NaCl equivalent (Kyle and Price, 1986). Some late calcites also contain low-salinity, single-phasefluid inclusions which suggest precipitation at less than 60°C (Prkryl et al., 1988). Hydrocarbon fluid inclusions are locally present in calcite and barite. Although Hockley sulfur isotope values for galena-sphalerite pairs are compatible with sulfide precipitation between 110 and 180°C (Price et al., 1983), the samples are not clearly co-precipitated. Thus, the broad range of sulfur isotope values may indicate a broad temperature spectrum, but it might indicate different sulfur isotope reservoirs. Therefore, the sulfur isotope values are considered more useful for determining sulfur sources than temperatures of formation. The presence of metalliferous formation waters in several areas of the Gulf Coast is permissive evidence for a genetic association with the cap rock deposits (Agee, 1988, 1990). Furthermore, the ratios of trace metallic components in Gulf Coast oil-field brines to those present in the cap rock deposits suggests an affiliation (Kyle and Price, 1986). Galenas from the Hockley sulfide deposit are non-radiogenic with 206Pb/204Pb ranging from 18.73 to 18.79 (Price et al., 1983). The lead isotopic composition of a modern Pb-rich brine and minor galena from a Cretaceous carbonate reservoir in southcentral Texas are similar, further suggesting metal supply to cap rock deposits from saline formation waters. Mineralized domes of the Houston diapir province occur within the geopressured- geothermal "fairways" of various Tertiary formations (Fig. 5.18; Morton et al., 1983). Formation waters in the lower Frio Formation below 4.5 km at the Pleasant Bayou geopressured-geothermal test well in Brazoria County are highly saline and enriched in metals, methane, carbon dioxide, and aromatic hydrocarbons (Kharaka et al., 1980). Svejensky (1984) modeled the Pleasant Bayou formation waters as potential ore-forming fluids for Mississippi Valley type deposits. These fluids may also represent the type of formation waters that resulted in cap rock mineralization (Agee, 1988, 1990). Although it appears that the metals now present in the cap rock deposits were

~~~~Data Center ,09126599985,[email protected], For Educational Uses ORIGIN OF CAP ROCK METALLIFEROUS DEPOSITS 457~~~~

~~~~STATES~~~~

~~GEOTHERMAL FAIRWAYS 0 10 20 30 40 5Okm aFRIO-VICKSBURG TREND 0WlLCOX TREND YEGUI TREND 0 10 20 ,om, @ PLEASANT BAYOU #2 -GEOTHERMAL TEST WELL a SALT DOME P NORMAL FAULTS hochures on dawnthrown side~~

Fig. 5.18. Geologic setting of the salt domes of the Houston diapir province relative to growth f aults andgeothermal fairways in Tertiaryreservoirs. Modified after Morton et al. (1983). supplied by saline formation waters, the ultimate metal source is more difficult to define. In fact, the mechanisms by which formation waters achieve their high major ion contents are not totally understood (Hanor, 1983). Light et al. (1987) developed a model that integrates salt diapirism, diagenetic mineral transformations, hydrocarbon maturation, geopressured brine generation, and metallic mineralization in the Gulf Coast. They concluded that several phases of fluid migration occurred during the last 25 Ma within the Houston diapir province. Rock-water interactions including the conversion of smectite to illite, albitization of detrital feldspars, and feldspar dissolution, could account for the release of significant quantities of metals to formation fluids (Burst, 1969; Hower et al., 1976; Light et al., 1987; Light and Posey, in press). Carbonate dissolution, siltstone stripping, and reduction of detrital iron oxides (red beds) could also supply metals

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(Carpenter et al., 1974; Sverjensky, 1984). Several of these sources can be evaluated where the age of mineralization is known, and some are incompatible with the timing of fluid migration. The margins of salt diapirs represent significant stratigraphic bypasses, and their fractured aeroles typically provide conduits for the migration of deep heated formation waters to the cap rock environment (Bennett and Hanor, 1987; Hanor and Bailey, 1983; Hanor and Workman, 1986; Hanor, 1987a, 1987b; Kyle and Price, 1986; Price et al., 1983; Ranganathan and Hanor, 1988). With regard to the composition of fluids and diagenetic minerals in a basin, and with regard to the mechanisms and pathways of mass transport, the importance of this physical and fluid-rock environment in diapiric basins cannot be overstated. Diapiric margins in diapiric basins are like fault conduits in other sedimentary basins and are capable of focusing major fluid volumes. Hanor and Workman (1986) and Bennett and Hanor (1987) have shown that plumes of saline fluid that occur in the southern Gulf Coast are caused by the dissolution of halite in salt diapirs. Whereas any brine will eventually reach equilibrium with the host sediments by dissolving and/or precipitating minerals, it is virtually certain that any brine, when introduced into a different PTX regime, will be out of equilibrium with the new regime. The composition of the brine will change through fluid-rock reactions until it again reaches equilibrium. Therefore, because the margins of diapirs are typically the upward escape routes for numerous formation waters, it is clear why the diapiric environment is so diverse in mineralogy and geochemistry. The dissolution of minerals as a result of halite dissolution may have profound implications on the creation of porosity for oil reservoirs, for fluid migration along diapir margins, and for the metallization of fluids. These same pathways also may undergo abnormal cementation relative to the adjacent sediments simply because they are the avenues of travel for a great number of compositionally different fluids.

~~~~Source of Sulfur~~~~

Most salt dome studies have determined or assumed that the ultimate source of sulfur in all sulfur-bearing species is anhydrite from the salt stock. These anhydrite crystals have a narrow 6% range of 16 f l%o (CDT), which is indicative of a Jurassic seawater sulfate parent (Fig. 5.13; Claypool et al., 1980; Posey et al., 1987a). However, massive anhydrite and gypsum cap rocks have slightly enriched average 6”s values and show a wider 6%Srange than anhydrite from the salt diapirs (Kyle and Price, 1986; Posey et al., 1987a). Late cap rock sulfates are enriched in 34Swith 6% ranging from 22 to 78%0 for barite and selenite (Kyle and Price, 1986).

~~~~Data Center ,09126599985,[email protected], For Educational Uses SOURCE OF SULFUR 45 9~~~~

Native sulfur samples show highlyvariable 6%vaIues which are believed to be the result of varying fractionation during bacterial metabolism (Feely and Kulp, 1957; Orr, 1982). 6%Svalues for cap rock sulfides range from -23 to 18%0, with a mean near @/oo (Fig. 5.13). Sulfates that are extremely enriched in 34Sprobably reflect the residual sulfate enriched in ?S following late-stage bacterial reduction to isotopically light, reduced sulfur (Posey et al., 1987a). The source of reduced sulfur for metallic sulfide precipitation was assumed to be bacterial reduction of cap rock sulfate (Thode et al., 1954; Feeley and Kulp, 1957). However, although the bacterial reduction mechanism is viable for some cap rock sulfide mineralization, it appears to have two major problems regarding reduced sulfur production: (1) the mineralization temperatures are higher than those generally accepted for effective bacterial metabolism (On, 1974), and (2) the post-sulfide origin for some of the calcite cap rock is a timing problem for the bacterial mechanism, especially for the early-formed anhydrite-hosted sulfides (Ulrich et al., 1984; Kyle and Price, 1985). However, it is possible to account for these apparent discrepan- cies if the mineralization system involves spatially or temporally distinct fluid pulses that result in episodic cooler conditions in the cap rock environment. The sulfur isotopic composition of H,S from some deep south Texas gas fields is similar to the Hockley sulfides, thus suggesting that externally generated reduced sulfur might have migrated to the structural trap provided by the salt diapir (Kyle and Agee, 1988). Trapping of reduced sulfur in the cap rock environment would provide a viable precipitation mechanism for metals carried in formation waters (Anderson, 1983). Kyle and Agee (1988) demonstrated from a profile of cap rock sulfides at Hockley dome (Fig. 5.19) that fluid mixing involving two distinct sulfur sources is a possible cause of the range of sulfide sulfur isotope values. This profile consists of a regular downward trend in 6% values from an initial light value of about -38ho at the top of the section to a heavy value of about +5%0 followed by an excursion back to heavy values of about -2@/00 (Fig. 5.19). By analogy to studies of the Winnfield Dome (Ulrich et al., 1983; Kyle et al., 1987), the majority of these Hockley sulfide laminae formed at the anhydrite salt interface while anhydrite accumulated by underplating. Accordingly, each layer is older than any successively deposited layer beneath because the anhydrite cap rock grew from top to base and the sulfides formed during that accumulation. As Kyle and Agee (1988) noted, the heavy sulfide values correlate with high metal abundances and are close, spatially, to a lens of calcite within the anhydrite cap rock. Their interpretations of these associations are that metals traveled along with reduced sulfur to the site of sulfide deposition but that carbon (presumably as hydrocarbons) reached the local site of sulfide precipitation somewhat later and formed calcite which

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~~~~METALS (pprn/l03) 0 100 200 300~~~~

*+ 0 Fe +++ + Pb+Zn *+ ++ + 500 k .H

~~~~a ' 700 + + + 800 + + + + + 1000~~~~

~~~~-40 -30 -20 -10 0 10~~~~

~~~~634~o/m pyrite~~~~

Fig. 5.19. Vertical profile of metal composition and pyrite 634Scomposition of drill hole HH2 through the cap rock on the southern margin of Hockley dome. Compare with Fig. 5.16. Modified after Kyle and Agee (I9x8).

is superimposed over the sulfide laminae (Fig. 5.17B). Sulfur is suggested to have been supplied from two sources: (1) a relatively shallow, isotopically light sulfur derived from the local biogenic reduction of cap-rock derived sulfate, and (2) an isotopically heavy sulfur derived from the external thermochemical reduction of sulfate present in deep-sourced formation waters. Evidence of these two end member reduced sulfur sources may be provided by the diagenetic pyrite in the sediments surrounding the cap rock (Fig. 5.13). Thus, this model interprets that salt dome sulfides incorporate reduced sulfur from both thermochemical sulfate reduction and biochemical sulfate reduction. Furthermore, these two fluids, each bearing reduced sulfur, mix at the site of sulfide deposition which principally was the anhydrite/salt contact. Although Kyle and Agee (1988) preferred a mineralization model in which the heavy sulfide sulfur came from a deeper source, there are other possible explanations

Data Center ,09126599985,[email protected], For Educational Uses SOURCE OF SULFUR 46 1 for the sulfur isotopic trends. The amount of sulfur isotope fractionation is controlled by several factors including (1) the rate of sulfate reduction, (2) pH, and (3) oxygen fugacity (Kaplan and Rittenberg, 1964; Kaplan, 1983; Sakai, 1968; Ohmoto, 1972). Rapid sulfate reduction, such as could occur when there is a great abundance of available reductant, leads to a low degree of fractionation, whereas slow reduction leads to something approaching equilibrium or maximum fractionation. For the case at hand, it is conceiveable that the variation in the Hockley dome sulfide sulfur isotope profile indicates variations in the amount of biogenic sulfate reduction: by this process, the light values represent slow rates of reduction when the availability of reductant, presumably hydrocarbons, was low, and the heavy values represent rapid rates of reduction when the amount of reductant was high. Consequently, the rapidly reduced sulfates, which are associated with the greatest metal sulfide concentrations, are also associated with calcite which formed due to the abundance of hydrocarbons. By this mechanism, it is not necessary to explain a mechanism for moving reduced sulfur and metals in the same solution without precipitating sulfides enroute and also provides a simplier mechanism for precipitating sulfides via sulfate reduction at the site of deposition. A pH increase induces an increase in the 6%Sof precipitated sulfides becaue H,S concentrates ”S relative to the dissolved sulfide ion as pH increases (see Hoefs, 1987). A pH increase, as might accompany precipitation of calcite in the interior part of the anhydrite cap rock, should be associated with heavier 6”s values in co-genetic sulfides. Because of the large fractionation between sulfate and sulfide, minerals precipitated under low sulfide/sulfdte conditions will have lighter 6%Svalues than minerals precipitated under high ratios (Rye and Ohmoto, 1974). Like the range of sulfur isotope values in sulfur found in MVT deposits, the 69 values of sulfur in salt domes cover a broad range. However, the two ranges are different. The range of sulfur 6% values of SDNS deposits is about -15 to + 15%0. In contrast, 6”s values of native sulfur from Pine Point, a large MVT deposit in the Northwest Territories of Canada cluster around + 23%0 (Powell and Macqueen, 1984). Although these variations in average sulfur isotope values are due, in part, to differences in sulfur sources and to mixing of sulfur from various sources, significant controls on the sulfur isotope values are also 1) the temperature of formation, 2) the interaction of sulfur species with bacteria, and 3) the mechanism of sulfate reduction. Salt dome sulfur formation appears to be a low-temperature, bacterially mediated process, whereas MVT sulfur formation appears to be unrelated to bacterial activity, largely because the temperature of formation is too high for bacteria to metabolize sulfur. Thus, the difference in temperature of formation, and the difference in sulfate reduction mechanism, both set apart the SDNS deposits from the MVT sulfur occurrences.

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~~~~Timing of Mineralization~~~~

Although the timing of salt diapirism can be well constrained by the responses of contemporaneous sedimentation (Seni and Jackson, 1983a,b), timing of cap rock formation and mineralization is less well defined. By analogy with the geologic setting of some Holocene salt diapirs that actually breach the land surface or sea floor (e.g. Kent, 1979; Jackson et al., in press), the initial products of salt dissolution may be removed from the site, and anhydrite cap rock formation may not take place until the diapir is mantled by subsequent sediments. In some cases, initial calcite cap rock development, particularly of detritus-rich Yalse” cap rock, could take place on the sea floor or within shallow sediments prior to or contemporaneous with anhydrite cap rock formation (Kyle et al., 1987). Similar biogenic calcite zones have been documented as elevated carbonate mounds overlying petroleum seeps associated with the flanks of shallow salt diapirs in the present Gulf of Mexico (Roberts et al., 1987). The timing of cap rock mineralization is a complex topic that is tied to the overall patterns of halokinesis, cap rock development, sedimentation, and diagenesis in the Gulf Coast. It can be readily proposed that mineralization was a relatively long-lived process and was initiated at different periods in the growth history of individual domes. Conceptually, mineralization, like the salt structures, may become younger toward the interior of individual basins, and indeed it is possible that mineralization systems are active in the main Gulf Basin. It has been demonstrated that sulfide mineralization associated with many Gulf Coast salt domes is intimately related to cap rock formation (Kyle and Price, 1986). Textural relationships and chemical composition of the sulfide layers in the anhydrite cap rock at Winnfield dome suggest that distinct episodic pulses of metalliferous brines entered the cap rock environment (Ulrich et al., 1984; Kyle et al., 1987). Anhydrite grains that are completely surrounded by sulfides are euhedral and undeformed, similar to the anhydrite disseminated throughout the salt diapir (Figs. 5.10C, 5.11,5.17F). Anhydrite grains outside of the mineralized layers are deformed and tightly intergrown. These textures suggest that mineralizing fluids entered the cap rock along the salt/anhydrite interface at the zone of halite dissolution, before that portion of the anhydrite zone was accreted to the overlying anhydrite cap rock. As halite dissolution and accompanying anhydrite accumulation and compaction continued, sulfide layers were successively precipitated in the anhydrite cap rock. Therefore, the oldest stratiform sulfide layers originating by this mechanism occur at the top of the present anhydrite cap rock zone while the youngest are found at the base (Fig. 5.11). The sequential accretion of the anhydrite cap rock and pyrrhotite within the stratiform layers at Winnfield dome is conducive to magnetostratigraphic investigation

~~~~Data Center ,09126599985,[email protected], For Educational Uses CLASSIFICATION OF SALT DOME MINERAL DEPOSITS 463~~~~

of the timing and rate of anhydrite cap rock formation (Kyle et al., 1987). Data provided by both the paleomagnetic pole position and the correlation with established sea floor magnetic reversal pattern indicates that the sampled portion of the Winnfield cap rock formed between 150 and 157 Ma. Furthermore, an average accretion rate for the measured anhydrite interval of 6.7 m/m.y. is indicated (Gose et al., 1989). These data for the timing and rate of cap rock formation are compatible with independent geologic evidence for timing and rates of diapir emplacement for interior basin salt domes.

~~~~Classification of Salt Dome Mineral Deposits~~~~

Cap rock metallic sulfide deposits share features in common with, but are distinctly different from, several classic ore deposit types. The genetic affiliations with the Mississippi Valley-type Zn-Pb deposits have been developed in the earlier discussions; even a portion of the cap rock sulfide concentrations is "carbonate-hosted". Therefore, salt dome associated sulfide deposits could be classified within the "mttlange" of Mississippi Valiey-type deposits (Sangster, 1983a).The unusual geologic setting of the salt dome deposits dictates many obvious differences from the typical Mississippi Valley-type deposit in Paleozoic shelf carbonates. Atypical features of the cap rock sulfide deposits also include the intrabasinal setting of mineralization, the unique structural features of the diapirs, the relatively early timing of mineralization, and the presence of significant amounts of silver. Sedimentary exhalative stratiform Zn-Pb-Ag sulfide deposits typically occur in fine-grained siliciclastic sequences (Sangster, 1983b). Some workers have proposed that the mineralizing fluids for these deposits are saline formation waters that were tapped by normal faults and discharged to the sea floor to form stratiform sulfide concentrations in topographic depressions (Gustafson and Williams, 1981; Badham, 1981 ; Lydon, 1983; Sawkins, 1984). The overall tectonic environment for these deposits is similar to the early rift phases of the Gulf Coast basins, e.g. the East Texas Basin (Jackson and Seni, 1983). Ore deposit classification is useful to highlight both similarities and differences between mineral deposits. Without debating the merits of classification schemes in general, we shall assume at this point that it is sufficient to note that classifications which have placed salt dome mineral deposits in the same class as MVT deposits or Sedex deposits leave out important differences in the composition and probable origins of these deposits. In addition to the major differences in the host rocks, salt dome mineral deposits have different mineral assemblages, geochemical and isotopic characteristics, origins, and, in some cases, f luid-drive mechanisms.

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Compositionally, the salt dome metal sulfide deposits have associated ore grade sulfur deposits and local extremely high grade celestite and strontianite deposits that the MVT and Sedex type deposits generally lack. Likewise, the extremely light carbon isotopes found in salt domes do not occur in either MVT or Sedex deposits. In addition, light sulfur isotopes found in salt dome sulfur deposits appear nowhere in either of the other deposit types, and the extremely heavy sulfur isotopes found in barite of salt domes likewise are not common in the other deposit types. Sulfur sources and mechanisms of sulfate reduction also indicate significant differences between these deposit types. Evaporites, which are ultimately of marine origin, are the probable sulfur source for cap rock and MVT deposits. However, bacterially reduced sulfate dissolved in seawater is the probable source for some Sedex deposits (e.g. Goodfellow, 1987). Biogenic processes are the dominant sulfate-reducing mechanism for salt dome sulfur, sulfides, and barite (Feely and Kulp, 1957; Sassen, 1980, 1988), but thermochemical sulfate reduction is the probable sulfate-reduction mechanism in some (most?) MVT deposits. According to Orr (1974,1977), native sulfur (or H,S) with extremely low values (ca. <-15%0) are indicative of biogenic sulfate reduction; higher values characterize thermochemical sulfate reduction. Marine sulfate sulfur ranges from about + 10 to +30%0 (see Claypool et al., 1980, for a discussion), and native sulfur or H,S values that are close to these marine values are generally construed to indicate thermochemical sulfate reduction. Additionally, the heavier the sulfur or H2S 6 %Svalue, the greater the likelihood of higher temperature sulfate reduction. Extremely light native sulfur is characteristic of salt domes, whereas heavier values are found in MVT native sulfur. The dominant source of all sulfur in salt dome deposits is most probably anhydrite in the cap rocks, as described above, although thermochemically reduced sulfate from deeper sources may play a role in the formation of sulfides (Kyle and Agee, 1988). The sulfur isotope ratios of native sulfur are low, thus a biogenic sulfate reduction mechanism is favored. A component of biogenic sulfur (native sulfur with extremely light values) occurs near the Pine Point Zinc District, NWT, Canada, but this sulfur is removed from the deposits and appears to be forming now rather than having been formed at the time of lead-zinc mineralization (Sasaki and Krouse, 1969; Powell and Macqueen, 1984). Overall, the source of sulfur in the Pine Point deposits and other MVT deposits, as deduced from the sulfur isotope values of sulfides, is thermochemically reduced marine sulfate. This sulfate-reduction mechanism sets the genesis of MVT sulfides apart from salt dome sulfides, and justifies a tentative classification of these as separate deposit types. The mechanism of fluid transport provides another criterion for considering the

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origin of MVT sulfides different from most salt dome deposits. Most North American MVT deposits apparently formed when fluids were expelled from sedimentary basins by gravity-driven flow (Bethke, 1985; Garven and Freeze; 1984). For most of these deposits, the topographic (gravity-driven) flow systems were activated by orogeny. However, some salt dome flow systems include components of both topgraphic drive and overpressure drive. Most salt domes in the nearshore and offshore Gulf Coast are undergoing basin dewatering or fluid mixing because fluids from geopressured shales are being expelled upward. For many, this fluid drive mechanism is a feasible one for supplying metals, hydrocarbons and, perhaps, reduced sulfur to the cap rocks. However, this has not always been the case, and fluid flow systems around many salt domes, notably those in the interior salt dome basins are topographic-drive systems rather than geopressured. In northern Africa, where the stratigraphic section is composed mostly of mark rather than siliciclastics, overpressuring probably never developed. Whereas the hydrodynamic mechanism criterion for segregating the deposit types is not as precise as some of the mechanisms described previously, it is worth noting that none of the MVT deposits have been clearly demonstrated to have formed due to overpressure fluid drive systems.

~~~~GENERAL MODEL FOR DLAPIRIC HALOKINESIS, CAP ROCK DEVELOPMENT, AND MINERALIZATION~~~~

Shortly after deposition of the mother evaporite sediments, just prior to cap rock formation, the evaporites were intruded by fluids most probably from beneath the evaporite section. These fluids may have been driven by overpressuring caused by the evaporites or by topographic drive. These early invading fluids recrystallized anhydrite (or, perhaps, gypsum) and in the process of recrystallization,strontium in the fluid mixed with strontium from the minerals. During this process, halite also recrystallized. By this process, anhydrite inherited strontium isotope values covering a range from their original marine value to the higher values carried by the fluids. This also left the halite with lower bromide concentrations than it inherited during deposition from seawater. The intrusion of these fluids probably induced or, at least, helped to induce diapirism. Anhydrite cap rock formation began soon after deposition of the mother evaporites. Using the known age of the Winnfield dome cap rocks and the apparent age of cap rocks around Tunisian diaph, it appears that cap rock formation begins at the onset of diapirism, or very shortly thereafter. The oldest paleomagnetic ages of pyrrhotite-bearing anhydrite cap rocks at Winnfield are only slightly younger than Mid-Jurassic, which is the commonly accepted age of the mother Louann salt. In

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Tunisia, calcite cap rocks that are interlayered with marine carbonate rocks just slightly younger than the Triassic, and which have been deformed and truncated by events only slightly younger than the marine carbonates, also indicate the cap rocks formed very soon after deposition of the mother evaporites. Anhydrite cap rock formed in a reverse stratigraphic sequence by salt dissolution and anhydrite residue accumulation as cap rock by underplating over a significant period of time. Sulfide mineralization commonly accompanied the formation of anhydrite. As the mother evaporite beds were buried, a neck of evaporites in the diapirs was continuously fed from the mother salt. The top of the diapir stayed at a relatively constant position relative to the land surface or sediment-water interface, while newly deposited sediments slid like a sleeve down the diapir margin. This process appears to moved in pulses closely associated with times of major sediment imput. The margins of the diapir became fluid conduits, either because of brecciation, or because saline fluids, which formed around the diapir as a result of minor halite dissolution, continually dissolved minerals to create secondary porosity. As organic-rich sediments were buried alongside the diapirs, they matured, and at unspecified times, in company with associated formation fluids and fluids from nearby formations, migrated upward along the margins of the diapirs. Metals accompanied the hydrocarbons and other formation fluids, depending on the mineral composition of the formations from which the fluids came (Fig. 5.20). The drive for fluid migration may have been topographic or overpressuring. Topographically driven flow probably affected the region of the proto-Atlas Mountains as meteoric recharge of Alpine highlands to the north drove fluids deep into the basins and then along permeable exit points such as growth faults and the margins of diapirs. Overpressuring is currently operating in the Gulf Coast such that fluids from beneath the top of the geopressured zone seep upward along growth faults and diapir margins. Petroleum that reached shallow depths above or alongside domes, where there was abundant sulfate in solution, where temperatures were lower than about 75 to 85C, and where meteoric water was available, at least, periodically, provided energy for bacteria that reduced sulfate and released HCO;, H,S, and, possibly, CO, or CH, (see Machel, 1987, for a discussion of the major reactions). During this process, calcite, dolomite, native sulfur, sulfides and metal sulfates may have been precipitated, leaving by-product bitumen and biodegraded hydrocarbons. Throughout at least the early stages of calcite cap rock formation, excess CO, led to considerable dissolution of previously-formed carbonate. With time, as anhydrite continued to be replaced by calcite to form calcite cap rock, acid abundance diminished and the calcite zone became stable. Throughout calcite cap rock formation, dissolution

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of anhydrite beneath the calcite created voids that led to collapse brecciation. Cracks within these breccias were healed by later generations of calcite. Thus, in most cap rocks, two types of breccia were created: 1) a polymictic carbonate breccia, which is more common in the upper part of the calcite cap rock, and which contains partly acid-- leached and thereby rounded clasts having multiple generations of breccia development, and 2) a crackle breccia, most generally found in the lower part of the calcite cap rock, that has alternating bands of replacement (of anhydrite) calcite and vein or open-cavity fill calcite. The formation of metal sulfides may have accompanied cap rock formation (both calcite and anhydrite), or may have post-dated it. Metals that accompanied hydrocarbons formed deposits limited in size to the amount of metal in the hydrocarbon-bearing formation waters. These fluids were inherently self-limiting in abundance because they led to the formation of secondary calcite, both above the domes, as cap rocks, and alongside the domes as cements. Metals that were carried in formation fluids alone, which accompanied thermally-reduced sulfate, and which did not form secondary seals that limit fluid flow may have formed larger metal sulfide concentrations either in the cap rock environment or in the Mississippi Valley type settings. Native sulfur formed after calcite cap rock which formed both a reservoir and a physical seal that trapped H,S. When oxygenated groundwater invaded the cap rock, it oxidized H,S to native sulfur. During this process, some of the anhydrite was converted to gypsum. The formation of gypsum may have served to seal the system to further water penetration, thus preserving the native sulfur from re-dissolving in groundwater.

~~~~CONCLUSIONS~~~~

The salt dome environment provides a focus for interrelated geologic processes that result in the development of major mineral resource concentrations. The coupled relationship between sediment loading and salt diapirism creates a wide variety of hydrocarbon traps. The margins of the salt diapirs provide major conduits for vertical fluid transfer within sedimentary basins. The salt mass undergoes substantial dissolution upon entering strata with low-salinity pore fluids, thus initiating cap rock formation. The resulting high-salinity fluids promote diagenetic reactions that further contribute to changes in pore fluid chemistry (Fig. 5.20). These reactions appear to control the evolution of formation fluids that ultimately create valuable mineral concentrations and that create porous zones and permeability barriers to form major oil and gas traps. The recently Zn-Pb-Ag-Baconcentrations in salt-domecaprocksare an interesting

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Fig. 5.20. Schematic hydrodynamic model for the interaction of cool shallow meteoric water and relatively hot, deep-sourced metal-bearing saline formation waters resulting in cap rock mineralization at Hockley dome. See text for discussion oj alternative models for geochemical and isotopic patterns for cap rock mineralization. Reproduced with permiasion from Kyle and Agee (I 988).

type of mineralization in sedimentary environments. These salt-dome metal deposits appear to result when metalliferous formation waters move up the flanks of salt diapirs to cap-rock environments (Fig. 5.20). Episodic interaction between low temperature, oxidizing pore water and warm, saline formation water probably initiated sulfide precipitation. Hydrocarbons provide an essential food source for the bacterial production of CO, and H,S for calcite cap-rock formation and for sulfur-bearing mineral metallic precipitation, respectively. Sulfate-reducing bacteria probably are responsible for the generation of reduced sulfur for the main stage of calcite cap-rock mineralization, but other mechanisms may have contributed sulfur, particularly for the early stages

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of anhydrite cap-rock mineralization. Salt dome mineral deposits have mineralogic, hydrologic, and genetic characteristics that distinguish them from the MVT and Sedex ore deposits with which they are most commonly compared. Additional study is required to refine the bacterogenic model for the cap rock hosted sulfur deposits. The commercial significance of salt domes is expanding beyond their traditional role as a source of energy and industrial minerals. Salt domes currently are being converted into storage vessels for several types of solid and liquid products. Expanded use of salt dome storage presents special environmental concerns and requires more understanding of the tectonic stability of salt diapirs.

~~~~ACKNOWLEDGMENTS~~~~

We are pleased to acknowledge many research colleagues who have collaborated with us in the study of salt domes and on the geology of the Gulf Coast, in particular W. Agee, W.A. Gose, M.P.A. Jackson, T. Jackson. L.S. Land, P.E. Price, J. Pnkryl, S.J. Seni, J. Sharp, M.R. Ulrich and G.R. Wessel. We are also indebted to the WinnRock Quarry, Damon Mound Quarry, Texasgulf Sulphur Company and Freeport Sulphur Company and staff for numerous collecting excursions in their facilities. HHP gratefully acknowledges discussions on the geology of the salt diapir mineral Occurrences in Tunisia with D. Graveson (Freeport-McMoRan), A. Abbes, M. Chickhaoui, A. Mansouri and B. Said (Office National des Mines, Tunisia), S. Bouhlel (University of Tunis), and J. Orgeval (Bureau de Recherches Geologiques at Minieres, France). We are grateful for salt dome research support provided by National Science Foundation grants EAR-8407736 and EAR-8709319 and by the Petroleum Research Fund grant 19624-AC2 of the American Chemical Society. Jeff Horowitz and David Stephens provided skilled drafting and photographic support, respectively.

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Jackson, M.PA. and SeN, SJ., 1983. Geometry and evolution of salt structures in a marghal rift basin of the Gulf of Mexico, East Texas. Geol., 11: 131-135. Jackson, M.PA. and Talbot, CJ., 1986. External shapes, strain rates, and dynamics of salt structures. Geol. Soc. Am. Bull., Y7: 305-323. Kaplan, I.R., 1983. Stable isotopes of sulfur, nitrogen and deuterium in recent marine environments. In: MA. Arthur (Organizer), Stable Isotopes in Sedimentary Geology. Soc. Eon. Paleon. Mineral., Short Course No. 10 2-1 - 2-108. Kaplan, I.R. and Rittenburg, S.C., 1964. Microbiological fractionation of sulfur isotopes. J. Gen. Microbiol., 34: 195-212. Kent, P.E., 1979. The emergent Homw salt plugs of southern Iran. J. Pet. Geol., 2 117-144. Kharaka, Y.K., Callender, E. and Carothers, W.W., 1977. Geochemistry of geopressured-geothermal waters from Texas Gulf Coast. In: J. Meriwether (Editor), Proc. Third Cong. Geopressured Geothermal Energy, pp. 121-165. Kharaka, Y.K., Lico, M.S. and Carothers,W.W., 1980. Predicted corrosion and scale-formation properties ofgeopressured-geothermal waters from the northern Gulf of Mexico Basin. J. Sed. Petrol., 32 319-324. Kharaka, Y.K., Lico, M.S., Wrrght, A. and Carothers,W.W., 1979. Geochemistry offormationwaters from Pleasant Bayou No. 2 well and adjacenty areas in coastal Texas. In: M.H. Dolfman and W.L. Fisher (Editors), Proc. Fourth Cong. Geopressured Geothermal Energy, Austin, pp. 178-193. Kharaka, Y.K., Maest,A.S., Carothers,W.W., Law, L.M., Lamothe, PJ. and Fries, T.L., 1987. Geochemistry of metal-rich hrincs from central Mississippi Salt Dome Basin. Appl. Ceochem., 2 543-561. Kreitler, C.W. and Dutton, S.P., 1983. Origin and Diagenesis of Cap Rock, Gyp Hill and Oakwood Salt Domes, Texas. University of Texas, Austin, Bur. of Econ. Geol., Rept. Invest., 131: 58 pp. Kyle, J.R. and Price, P.E., 1985. Mineralogical investigation of sulfide concentrations in the Hockley salt dome cap rock, Texas. In: W.C. Park, D.M. Hausen and R.D. Hagni (Editors), Applied Mineralogy. Metall. Soc. of AIME, New York, pp. 1065-1082. Kyle, J.R. and Price, P.E., 1986. Metallic sulphide mineralization in saltdome cap rocks, Gulf Coast, USA. Trans. Instn. Min. Metall., 95 B6-Bl6. Kyle, J.R., Ulrich, M.R. and Gose,WA., 1987. Textural and paleomagnetic evidencefor the mechanism and timimg of anhydrite cap rock formation,Winnfield Salt Dome, Louisiana. In: I. Lerche and JJ. OBrien (Editors), Dynamical Geology d Salt and Related Structures. Academic Press, Orlando, pp. 497-542. Kyle, J.R. and Age,W.N., 1988. Evolution of metal ratios and d34S composition of sulfide minerahtion during anhydrite cap rcxk formation, Hockley Dome, Texas, USA. Clem. Geol., 74 37-56. Laatar, E., 1980. Gisements de Plombzinc et Diapirisme du Trias Salifere en Tunisie Septentrionale: Les concentrations peridiapiriques du district minier de Nefate-Fedj-el-Adoum (repion de Teboursouk). Unpub. Thesis, P. and M. Curie Univ., Pans. 280 pp. Land, L.S. and Milliken, K.L., 1981. Feldspar diagenesis in the Fno Formation, Brazoria County, Texas Gulf Coast. Geol., 9 314-318. Land, L.S. and Prezbmdwski, D.R., 1981. The origin and evolution of saline formation water, lower Cretaceous carbonates, south-central Texas, USA.. J. Hydrol., 41: 51-74. Land, L.S., Kupecz, JA. and Mack, L.E., 1988. Louann salt Geochemistry (Gulf of Mexico sedimentary basin, U.SA.): A preliminary synthesis. Chem. Geol., 74: 25-35. Light, M.P.R., Posey, H.H., Kyle, J.R. and Price, P.E., 1987. Model for the origins of geopressured brines, hydrocarbons, cap-rocks, and metallic mineral deposits, Gulf Coast, USA. In: I. Lerche and JJ. OBrien (Editors), Dynamical Geology of Salt and Related Structures. Academic Press, Orlando, pp. 787-830. Light, M.P.R. and Posey, H.H. (in press). Diagenesis and its relation to mineralization and hydrocarbon reservoir development: Gulf Coast and North Sea Basins. In. K.H. Wolf and G.V. Chhgar (Editors), Diagenesis 111, Elsevier, New York. Lobao, JJ. and Pilger, R.H., 1985. Early evolution of salt structuresin the North Louisiana Salt Basin. Trans. Gulf Coast Assoc. CiOl. Soc., 35 189-197. Lydon, J.W., 1983. Chemical parameterscontrollug the origin and deposition of sediment-hosted stratiform lead-zinc deposits. In: D.F. Sangster (Editor), Short course in Sediment-hosted StratiformLead-zinc Deposits. Min. Assoc. Can., Toronto, Handbook 9 175250. Machel, H.G., 1987. Some aspects of diagenetic sulphate-hydrocarbon redox reactions. In: J.D. Marshall (Editor), Diagenesis

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cf Sedimentaly Sequences. Geol. Soc. Spec. Publ. No. 36, Blackwell, London, pp. 15-28. Macpherson, G.L., 1989. Lithium, Boron, and Barium in Formation Waters and sediments, northwest Gulf d Mexico Sedimentary Basin. Ph.D. dissertation, University cf Texas, Austin, 314 pp. Martin, R.G., 1978. Northern and eastern Gulf d Mexico continental margin: stratigraphic and structural framework. In: A. H. Bourn, G.T. Moore and J.M. Coleman (Editors), Framework, Facies, and Oilil-trappingCharacteristics cf the Upper Continental Margin. Am. Assoc. Pet. Geol., Stud. in Geol. No. 7 2142. McManus, K.M. and Hanor, J.S., 1988. Calcite and iron sulfide cementation d Miocene sedimentsflanking the West Hackberry salt dome, southwest Louisiana, USA.. Chem. Geol., 74 99-112. Milliken, K.L., 1985. Petrology and Burial Dugenesis d fio-Pleistocene sediments, northern Gulf cf Mexico. Ph.D. dissertation, University d Texas, Austin, 112 pp. Molinda, G.M., 1988. Investigation d Methane Occurrence and Outbursts in the Cote Blanche Domal Salt Mine, Louisiana. US. Bur. Mines, Rept. Invest., 9186: 21 pp. Montacer, M., Disnar, J.R., Orgeval, J J. and Trichet, J., 1988. Relationship between Zn-Pb ore and oil accumulation processes: Ezample of the Bou Grine deposit (Tunisia). Org. Geochem., 13 423431. Morton, RA., Ewing, T.E. and Tyler, N., 1983. Continuity and Internal Properties d Gulf Coast Sandstones and Their Implicationsfor Geopressured Fluid Production. University of Texas, Austin, Bur. d Econ. Geol., Rept. Invest. 132, 70 PP. Murray, G.E., 1966. Salt structures d the Gulf d Mexiw Basin - a review. Bull. Am. Assoc. Pet. Geol., 50: 439478. Ohmoto, H., 1972. Systematics cf sulfur and carbon isotopes in hydrothermd ore deposits. &on. Geol., 67 551-578. Orgeval, J-J, Giot, D., Karoui, J, Monthel, J. and Sahli, R., 1989. The discovery and investigation of the Bou Grine Pb-Zn deposit (Tunisian Atlas). Chron. Res. Min, Special Issue, pp. 53-68. Orr, W.L., 194.Changes in sulfur content and isotopic ratios of sulfur during petroleum maturation - study cf Big Horn Basin Paleozoic oils. Bull. Am. Assoc. Pet. Geol., 58: 2295-2318. Orr, W.L., 1977. Geol0g;c and geochemical controls on the distribution d hydrcgen sulfide in natural gas. In: R. Campos and J. Goni (Editors), Advances in Organic Geochemistry. E~di~ma,Madrid, pp. 571-597. Orr, W.L., 1982. Rate and mechanism d non-microbial sulfate reduction. Geol. Soc. Am., 14 580 (Abstracts with Programs). Perthuisot, V., 1977. Dynamique et PetrcgenM des Ektrusions Triasiques en Tunisie Spetentrionale. Travaus du Labo. Geol., &ole Normale Sup., Paris, No. 12: 312 pp. Pervinquiere, L., 1903. Etude Geologique de la Tunisie Centrale. F.R. de Rudeval, Park, 359 pp. Posey, H.H., 1986. Regional Characteristics cf Strontium, Carbon, and Oxygen Isotopes in Salt Dome Cap Rocks cf the Western Gulf Coast. W.D. dissertation, University d North Carolina, Chapel W, 235 pp. Posey, H.H.,Kyle, J.R., Jackson, TJ., Hurst, S.D. and Price, P.E., 1987a. Multiplefluid componentsd salt diapirs and salt dome cap rocks, Gulf Coast, USA. Appl. Geochem., 2 523-534. Posey, H.H., Price, P.E. and Kyle, J.R., 1987b. Mixed carbon sources for calcite cap rocks cf Gulf Coast salt domes. In: I. Lerche and JJ. O'Brien (Editors), Dynamical Geology cf Salt and Related Structures. Academic Press, Orlando, pp. 593-630. Posey, H.H. and Kyle, J.R., 1988. Fluid-rock interactions in the salt dome environment: An introduction and review. Chem Geol., 74 1-24. Powell, T.G. and Maqueen, R.W., 1984. Precipitation d sulfide ores and organic matter: sulfate reactions at Pme Point, Canada. Sci., 224: 63-66. Price, P.E., Kyle, J.R. and Wessel, GR., 1983. Salt dome related zinc-lead deposits. In: G. Kisvarsanyi et. al. (Editors), Pr- International Gmfemce on Mississippi Valley-typelead-&c deposits, Univ. Missouri-Rda, pp. 558-571. F'rikryl, J.D., 1W. Origin cf Salt Dome Limestone Cap Rocks in the US. Gulf Coast. MA. Thesis, University cf Texas, Austin, m-5 PP. Pnkryl, J.D., Posey, H.H. and Kyle, J.R., 1988. A petrographic and geochemical model for the origin d calcite cap rock at Damon Mound salt dome, Texas, USA. Chem Geol., 74 67-97. Ranganathan, V. and Hanor, J.S., 1988. DensityitriVen groundwater flow near salt domes. Chem. Geol., 74 173188. Robem H.H., Sassen, R. and Aharon, P., 1987. Carbonates d the Louisiana continental slope. In: Proc. Nineteenth Offshore Tech. Gmf.,Houston, pp. 373-382. Rowier, H., Perthuisot, V. and Mamnui, A., 1985. Pb-Zn deposits and salt-bringdiapirs in southern Europe and North

~~~~Data Center ,09126599985,[email protected], For Educational Uses 474 SALT DOME MINERALIZATION~~~~

Africa. Eon. Geol., 80: 666-fXl. Ruckmick, J.C., Winkrly, B.H. and Edwards, AF., 199. Classification and genesis of bicgenic sulfur deposits. Eon. Geol., 74 460474. Rye, R.O. and Ohmoto, H., 194. Sulfur and carbon isotopes and ore genesis: a review. Econ. Geol., 69: 826-842. Sakai, H., 1968. Isotopic properties of sulfur compounds in hydrothermal prmsses. Ckochem. J., 2 2949. Salvador, A., 1987. Late Triassic-Jurassicpaleogeography and origin cf Gulf cf Mexico Basin. Bull. Am. Assoc. Pet. Geol., 71: 419451. Salvador, A. and Buffler, R.T., 1982. The Gulf of Mexico basin. In: A.R. Palmer (Editor), Perspectives in Regional Geological Synthesis: Plannhg for the Geology of North America. Geol. Soc. Am., Decade of North American Geology, pp. 157-162. Sangster, D.F. (Editor), 1983. Short course in Sediment-hosted Stratifom Lead-zinc Deposits. Min. Assoc. Canada, Toronto, Handbook 9: 309 pp. Sassen, R., 1980. Biodegradation of crude oil and mineral deposition in a shallow Gulf Coast salt dome. Oqanic Geochem., 2 153-166. Sassen, R., Ch,E.W. and McCak, C., 1988. Recent hydrocarbon alteration, sulfate reduction and formation of elemental sulfur and metal sulfides in salt dome cap rock. Clem. Geol., 74 57-66. Sasaki, A. and Krouse, H.R. 1969. Sulfur isotopes and the Pine Point lead-zinc mineralization. Econ. Geol., 64:718-730. Saunders, JA., Prikryl, J.D. and Posey, H.H., 19%. Mineralogic and isotopic constraints on the origin cf strontium-rich cap rock, Tatum Dome, Mississippi, USA. Chem. Geol., 74 137-152. Sawkins,FJ.. 1984. Ore genesis byepisodic dewatering of sedimentary basins: application to giant Proterozoic lead-inc deposits. Geol., 12: 451454. Schatzel, S.J. and Hyman, D.M., 1984. Methane Content of GuP Coast Domal Rock Salt. US. Bur. Mmes, Rept. Invest., 8889: 18 PP. Seni, SJ. and Jackson, M.PA., 1983a. Evolution of salt structures, east Texas dmpu province, part 1, xdimentary record of halokincsis. Bull. Am. Assoc. Pet. Geol., 67: 1219-12114. Seni, SJ. and Jackson, M.PA., 1983b. Evolution of salt structures, east Texas diapu province, part 2, patterns and rates of halokinesis. Bull. Am. Assoc. Pet. Geol., 67: 1245-1274. Seni, SJ. and Jackson, M.PA., 19M. Sedimentary record of Cretaceous and Tertiary Salt Movement, East Texas Basin. University of Texas, Austin, Bur. of Eon. Geol., Rept. Invest., 139 89 pp. Sharp,Jr., J.M., Galloway, W.E., Land, L.S., McBridc, E.F.,Blanchard, P.E., Btdnar, D.P., Dutton, S.P.,Farr, M.R.,Gold, P.B., Jackson. TJ., Lundegard, P.D., Macpherson. G.L. and Milliken, K.L., 198x. Diagenetic processes in northwest Gulf d Mexico sediments. In: G.V. C'ingarian and K.H. Woif (Editors), Diagenesis 11, Elsevier, New York, pp. 43-113. Sverjensky, DA., 1984. Oil field brines as ore-forming solutions. Econ. Geol., 79 23-37. Taylor. R.E., 1938. Origin of Cap Rock of Louisiana Salt Domes. Louisiana Geol. SUN. Bull., 11: 191 pp. Thode. H.G., Wanless, R.K. and Wallouch, R., 1954. The om of native sulphur deposits from isotope fractionation studies. Geochim. Cosmochim. Acta, 5: 286-289. Trusheim, F., 1960. Mechanisms of salt migration in northern Germany. Bull. Am. Assoc. Pet. Geol., 44: 1519-1410. Tyler, N., Ewing, T.E., Fisher, W.L. and Galloway, WE., 1985. Oil exploration and prtduction plays in the Texas Gulf Coast basin. Proc. Fourth Res. Conf., Gulf Coast Sect., Soc. Eon. Paleon. Mineral., pp. 81-99. Ulrich, M.R., Kyle, J.R. and Price, P.E., 1984. Metallic sulfide deposits in the Winnfield salt dome, huisknd: evidence for episodic introduction of metalliferous brines during cap rock formation. Trans. Gulf Coast Assoc. Geol. Socs., 34: 435-442. Urai, J.L., 1983.Water assisted dynamic recrystallization and weakening in polyctystalline bischofite. Tectonophysics, % 125-157. Weast, R.C. (Editor-in-Chef), 1989. CRC Handbook of Clemistry and Physics. Siwty-nineth Edition. CRC Press, Boca Raton, Fla., 2A67 pp. Weeks, A.D. and Eargle, D.H., 1060. Uranium at Palangana Salt Dome, Duval County, Texas. US. Geol. SUN.,Prof. Paper 400-8: B48-852.

~~~~Data Center ,09126599985,[email protected], For Educational Uses This Page Intentionally Left Blank~~~~

~~~~Data Center ,09126599985,[email protected], For Educational Uses Chapter 6~~~~

~~~~EVAPORITES, EVAPORITIC PROCESSES AND MINERAL RESOURCES~~~~

~~~~J. Richard Kyle~~~~

~~~~INTRODUCTION~~~~

The role of evaporites and evaporitic processes in the formation of commercial concentrations of mineral resources is diverse indeed. Many evaporite deposits of both marine and lacustrine types are valuable industrial rocks that supply widely used materials for the construction, chemical, and agricultural industries. Evaporites also serve as the feedstock for natural processes that either have formed valuable materials or have provided a favorable environment for the chemical trapping of externally supplied elements. In other cases, intrastratal dissolution of evaporites has resulted in the creation of permeability trends that have been important controls governing ore-forming fluid movement or have provided ore-hosting porosity. Some "liquid orebodies" are saline formation waters of evaporative parentage that are unusually enriched in valuable ions. In addition, metal-rich formation waters have been widely accepted as modern analogs of ore-forming fluids for a variety of classic ore deposit types. Although the mechanism by which these fluids become metal rich is difficult to establish, it appears that rock-water interactions involving saline formation waters are generally important. Evaporation also has been proposed as a hydrologic agent responsible for focused ground water flow that controls several major geologic processes including ore deposition. It is quite common for several valuable commodities to occur within the same evaporitic sequence or basin. For example, the Maritimes Basin in eastern Canada produces major amounts of gypsum, salt, and potash from Mississippian evaporites of the Windsor Group. The shelf carbonate strata of the Windsor Group host major concentrations of lead, zinc, barite, and celestite, although the precise genetic relationship between the evaporites and the valuable mineral concentrations is more complex. This paper briefly summarizes the varied role of evaporites in the formation of commercial concentrations of valuable materials with emphasis on geologic environments that are not covered elsewhere in this volume. There is a biased coverage of evaporitic sequences in North America with which the author is more familiar. Although evaporitic geologic environments are important in the formation of petroleum and oil shales, these types of energy resources are not considered in this review. Other

~~~~Data Center ,09126599985,[email protected], For Educational Uses 478 EVAPORITIC PROCESSES AND MINERAL RESOURCES~~~~

~~~~reviews of selected aspects of this complex topic are provided by Davidson (1966), Sonnenfeld (1984), and Eugster (1985).~~~~

~~~~EVAPORITES AS MINERAL RESOURCES~~~~

~~~~Introduction to Evaporite Formation~~~~

Evaporitic strata are major sources of important industrial rock and mineral commodities, including gypsum, halite, potash, sodium carbonate, sodium sulfate, and borax (Table 6.1A). These materials are consumed in large quantities for many uses that are essential to the modern construction, chemicals, and agricultural industries. Many of these evaporitic materials are widely distributed geographically (and geologically). Because of the relatively low unit value of most industrial rocks, the location of the deposit relative to consuming areas may have dominant economic significance. The reader is referred to Lefond (1983) and Harben and Bates (1984) for reviews of the geologic nature, processing, and uses of these commodities. Much of the significant industrial rock and mineral production is from strata that generally are interpreted to represent marine environments. However, many of the hosting basins owe their origin to rifting within cratons and therefore include sedimentary sequences that span the transition from terrestrial to open marine environments (Hardie, 1984). The evaporite sequences can reach great thicknesses and are common in many parts of the world and in basins of many diverse geologic periods. Major North American evaporitic sequences that are significant in terms of production of commercial products include the Silurian and Devonian strata of the Michigan and Appalachian basins, Devonian evaporites of the Western Canada Basin, Mississippian strata of the Maritimes Basin, and Permian evaporites of the Permian Basin (Fig. 6.1). Evaporite sequences of Paleozoic age also are important sources of industrial rocks and minerals in Europe, Asia, and South America. The composition and origin of natural saline fluids, particularly seawater, have been the subject of innumerable geochemical investigations for well over a century. Although origin of these solutions and evaporite deposits is a complex subject that is far beyond the scope of this review, a few introductory remarks are in order. Because major evaporitic sequences generally are interpreted to represent marine depositional environments, there have been numerous studies comparing predicted mineral sequences resulting from the evaporation of seawater with natural evaporite deposits (see Borchert and Muir, 1964; Braitsch, 1971; Holser, 1979; Sonnenfeld, 1984, and Sonnenfeld and Perthuisot, 1989, for major reviews). It has long been recognized that

~~~~Data Center ,09126599985,[email protected], For Educational Uses INTRODUCTION TO EVAPORITE FORMATION 479~~~~

~~~~Table 6.1. Relation of Evaporites and Mineral Resources.~~~~

~~~~COMMODITY EXAMPLE AGE REFERENCE~~~~

~~~~A. Evaporites as mineral rcsources~~~~

1. Gypsum Appalachian Basin Silurian Applcyard (198.3a,b) Michigan Basin Devonian, Mississippian Mantimes Basin Mkissippian Anadarko Basin Permian Gulf d C'alifomia Pliocene Salton Trough Miocene Paris Basin Etrene

2. Halite Appalachian Basin Silurian Zharkov (1984) Michigan Basin Silurian, Ldond and Jacoby (1983) Devonian Maritimcs Basin Mississippian Permian Basin Permian Zechstein Basin Permian Gulf Coast (diapirs) Jurassic

3. Potash Elk Point Basin Devonian Zharkov (1984) Moncton Basin Mississippian Adams and Hite (1983) Paradox Basin Pennsylvanian Lowenstein and Spencer (1990) Delaware Ba.sin Permian Hardie (1993) Zechstein Basin Permian Congo Basin Cretaceous

~~~~4. Sodium Carbonate Green River Basin Eocenc Mannion (1983) (Trona) Lake Magadi Holocene Parker and Mannion (1971) Searles Lake Modem brine Eugster (198Ob, 1986)~~~~

~~~~5. Sodium Sulfate Searles Lake Modem brine Weisman and McIlveen (1983) Ingebrigt Lake Holocene Broughlon (1984) C7aplin Lake Modern brine~~~~

~~~~6. Borax Boron Miocene Kistler and Smith (1983) Death Vdey Pliocene Smith (1966, lm9) Kirka Pliocene Searles Lake Modem brine~~~~

~~~~I. Bromine c;Uif Coast Basin Modern brine Jensen et al. (1983) Michigan Bash Modem brine Carpenter and Trout (lg8) Searles Lake Modern brine Dead Sea Modern brine~~~~

~~~~8. Iodine Anadarko Basin Modern brine Jan and Roe (1983) Michigan Basin Modern brine~~~~

~~~~9. Lithium Clayton Valley Modern brine Kunasz (1983) Salar de Atacama Modem brine~~~~

~~~~10. Nitrate Atacama Modern caliche Ericksen (1981)~~~~

~~~~Data Center ,09126599985,[email protected], For Educational Uses 480 EVAPORITIC PROCESSES AND MINERAL RESOURCES~~~~

~~~~Table 6.1 (continued) Relation of Evaporites and Mineral Resources.~~~~

~~~~COMMODITY EXAMPLE AGE REFERENCE~~~~

~~~~B. Evaporite solution or reaction contributes component(s) to ore-forming fluids~~~~

~~~~1. Pb-Zn-Ba Gulf Coast Modem brines Carpenter et al. (1974)~~~~

~~~~2. Zn-Cu-Ag Red Sea Modern brines Degens and Ross (1969)~~~~

~~~~3. Zn-Pb-Cu Salton Sea Modem brines McKibbzn and Elders (1%)~~~~

~~~~4. Zn-Pb-Ag-Au San Juan Volcanic Field Oligocene Berger and Bethke (1985)~~~~

~~~~5. Zeolites San Simon Valley Pliocene Mumpton (1981)~~~~

~~~~C. Evaporite alteration to form or precipitate valuable commodities~~~~

~~~~1. Sulfur Delaware Basin Permian Ruckmick et al. (1979) Gdf Coast (diapirs) Jurassic Feely and Kulp (1957) Fergana Basin Paleocene Mesopotamian Basin Miocene~~~~

~~~~2. Zn-Pb-Ag Gdf Coast (diapirs) Jurassic Kyle and Price (1986) Northern Africa (diapirs) Triassic Kyle and Posey (this volume)~~~~

~~~~3. Pb-Zn Pine Point Devonian Powell and Macqueen (1984) Gays River Mississippian Akande and Zentilli (1984)~~~~

~~~~4. Pb Bunter Basin Triassic Bjbrlykke and Sangster (1981)~~~~

~~~~5. Cu-Ag Zambia-Zaiie Proterozoic Bartholomti (1974) Belt Series Proterozoic Lange et al. (1987) Kupferschiefer Permian Sverjensky (1987) Lisbon Valley Permian Morrison and Pany (1986)~~~~

~~~~6. Pb-Zn-Cu-Ag McArthur River Proterozoic Sangster (1983) Dugald River Proterozoic Muir (1983) Mt. Isa Proterozoic Neudert (1986)~~~~

~~~~7. Iron Formations Hamersley Basin Proterozoic Trendall and Moms (1983) Lake Superior Proterozoic Eugster and Chou (1973)~~~~

~~~~8. Barite Walton-Cheverie Mississippian Mossman and Brown (1986) Delaware Basin Permian Kyle (in press) Gulf Coast (diapirs) Jurassic Kyle and Price (1986)~~~~

~~~~9. Celestite Loch Lomond Mississippian Forgeron (1%) East Greenland Permian Scholle et al. (1990)~~~~

~~~~10. Cu-Ni-Platinoids Noril'sk-TalnakhComplex Devonian Naldrett and Macdonald (1980)~~~~

~~~~Data Center ,09126599985,[email protected], For Educational Uses INTRODUCTION TO EVAPORITE FORMATION 48 1~~~~

~~~~Table 6.1 (continued) Relation of Evaporites and Mineral Resources.~~~~

~~~~COMMODITY EXAMPLE AGE REFERENCE~~~~

~~~~D. Evaporite solution to create orelosting porosity~~~~

~~~~1. PbZn GapRiver Proterozoic Hewton (1982) pine Point Devonian Beales and Hardy (1980)~~~~

~~~~2. Barite Central Texas Cretaceous Kyle (m press)~~~~

~~~~3. Celestite Coahuila Cretaceous Kesler and Jones (1981)~~~~

~~~~E. Evaporation as a hydrologic agent in ore formation~~~~

~~~~1. Reflux(PbZn) Mississippi Valley Paleozoic Lange and Murray (1977)~~~~

~~~~2. Pump(Cu-Ag) Zambia-Zaire Proterozoic Renfro (1974)~~~~

~~~~3. Calcrete (U-V) Western Australia Modem Mann and Deutscher (1978)~~~~

Note: In order to keep the bibliography to a reasonable length, many references cited here and in the text are secondary sources to which the reader is referred for primary information sources. major evaporite strata are not formed as the result of simple evaporation of seawater. Among other geochemical evidence is the unrealistic volume of seawater that would be required for single-stage evaporation to account for evaporite sequences hundreds of meters thick. Eugster et al. (1980) and Harvie et al. (1980,1984) have modeled mineral equilibria for initial solutions approximating the seawater system at 25°C. Braitsch (1971) provided data on phase relations of evaporite minerals to higher temperatures. Many major evaporite sequences represent the lateral and vertical facies relationships resulting from complex precipitation and reworking processes within marine basins. The initial precipitates generally are modified by clastic reworking or solution-reprecipitation reactions, particularly involving potash minerals, accompanying diagenesis at somewhat elevated temperatures. Major recent studies of the depositional environments of evaporite deposits are included in Dean and Schreiber (1978), Handford et al. (1982), Kendall (1984), Schreiber (1988), Warren (1989), and this volume. The present mineral assemblages within evaporite deposits typically consist of complex intergrowths of several minerals that are the result of these primary and secondary depositional processes. Evaporation of a closed body of seawater can proceed to within 90% of completion with the precipitation of only calcite (or dolomite) and gypsum. The residual brine is highly concentrated in sodium, magnesium, and

~~~~Data Center ,09126599985,[email protected], For Educational Uses 482 EVAPOKI'TIC PKOCESSES AND MINEKAL KESOUKCES~~~~

potassium salts. Further evaporation causes the precipitation first of halite, followed by complex salts of potassium and magnesium under conditions of extreme concentration. 'Thus, potash-rich "bittern"evaporite strata commonly occur stratigraphicallyhigh within evaporite sequences and within the interior part of basins, whereas halite, and particularly gypsum, generally mcur stratigraphically lower and toward the basin margins. Potassium may be initially precipitated as carnallite, and secondary changes, such as influx of fresh seawater that dissolves carnallite, bring about the formation of beds of sylvinite (a mixture of halite and sylvite). Further influx of seawater will cause this cycle to be repeated, with the result that potash minerals are found in relatively thin beds within thick sequences of other evaporites. 'The saline minerals have undergone varying degrees of recrystallization, dissolution, and reprecipitation during burial diagenesis that have affected the nature of the evaporite strata. For example, fluid inclusion and isotopic evidence support multiple episodes of recrystallization of the potash-bearing Devonian evaporites of the Elk Point Basin at temperatures of 3570°C from waters originating in overlying forma-

~~~~Fig. 6.1. Major commercial evaporite basins of North America. Areas containing potash-bearing evaporites are shown in the stippled pattern.~~~~

~~~~Data Center ,09126599985,[email protected], For Educational Uses GYPSUM AND ANHYDRI'I'E 483~~~~

tions (Chipley and Kyser, 1989). Lowenstein and Spencer (1990) have defended the syndepositional origin of potash evaporites using petrographic and fluid inclusion evidence to support formation from warm surface brines. Hardie (lY90) suggested that relatively high temperature, deep sourced CaCI,-rich brines actually formed the major deposits of potassium salts.

~~~~Gypsum and Anhydrite~~~~

Gypsum ( CaS04*2H,0) provides an excellent example of an evaporite product that is consumed in large quantities worldwide, principally for use in the construction industries; over 70 countries commonly report annual production (Appleyard, 1Y83a,b). 'The commercial uses of gypsum are tied to the mineral's hydrated character. When gypsum is calcined at 160"C, it loses 1.5 moles of combined water to form the calcium sulfate hemihydrate (CaS04*1/zH20),commonly known as Plaster of Paris. When mixed with water, this compound can be spread, cast, or molded into a desired form which will then set to a rocklike hardness. Its common use as a filling between sheets of heavy paper produces the plasterboard widely used in "drywall"construction; more than 40% of all gypsum consumption worldwide is in this form (Harben and Bates, 1Y84). Uncalcined gypsum is used as a retarding agent in cement, as a soil fertilizer and conditioner, and as a filler in such diverse products as paper, paint, and toothpaste. Anhydrite- and gypsum-bearing strata are widespread in time and space. For example, the commercial gypsum deposits of the United States occur in nineteen states and represent a variety of Phanerozoic evaporitic depositional environments. Proximity to market may compensate for less than ideal geologic character or mining conditions of a gypsum deposit. 'lhus, I-m thick gypsum beds (Fig. 6.2A,B) have been mined underground and processed economically near major population centers in western New York and southern Ontario (Haynes et al.,lY8Y), whereas much thicker and areally extensive gypsum strata exposed on the Gypsum Plain of west l'exas are not currently exploited. Most gypsum deposits, and all commercial ones, are relatively close to the present land surface and commonly are the result of hydration of anhydrite by shallow circulating, low-salinity waters. The Maritimes region of eastern Canada is the world's largest producer of gypsum, generally mining over 6 million metric tons annually from Mississippian strata of the Windsor Group (Cameron, 1Y84; Howie, lY88). Furthermore, these deposits have supplied much of the gypsum for the eastern seaboard trade since the late 170's. Most of the gypsum production has come from deposits within a 200 km' zone along the southwestern edge of the Maritimes Basin in Nova Scotia and New Brunswick (Fig. 6.3).

~~~~Data Center ,09126599985,[email protected], For Educational Uses 484 EVAPORITIC PROCESSES AND 'MINERAL RESOURCES~~~~

Fig. 6.2. Geologic features of commercial evaporite deposits from Silurian strata of the Appalachian Basin, southern Ontario. (A)Photograph of mine f aceshowingstromatolitic dolomite overlying nodulargypsum. (B) Handspecimen showinggypsum nodulesseparated by thin stringers of c1a.v and dolomite. (C)Photograph of salt mine face showing bedded halite. (D) Hand specimen showing salt layer bound by dolomite and clay laminae.

Some current production is also from similar deposits in southwestern Newfoundland. Much of the strata overlying the evaporitic sequence has been removed by erosion, thus exposing the units to hydration. The gentle dip of the shallow basement surface underlying the Windsor Group, combined with faulting and fracturing have aided ground water circulation which has formed mineable gypsum deposits as much as 60 m thick (Cameron, 1984). Commercial gypsum deposits generally occur within evaporitic strata of marine origin, but several major deposits were probably formed in non-marine playa environments. These deposits include the largest U.S. gypsum producer within Miocene evaporites in the Salton Trough at Plaster City, California, and the major producer within Pliocene evaporites on San Marcos Island in the Gulf of California. The gypsum

~~~~Data Center ,09126599985,[email protected], For Educational Uses SALT AND POTASH 485~~~~

~~~~DEPOSITS~~~~

~~~~0 GYPSUM Zn-Pb SULFIDES 0 BARITE ''Ohm CELESTITE~~~~

~~~~Fig. 6.3. Generalized geologic setting of the Maritimes Basin showing locations of metallic and nonmetallic mineral deposits. Ruled pattern shows the distribution of Carboniferous exposures.~~~~

strata on San Marcos Island occur between non-marine sandstone units and range from 10 to 30 m thick. Gypsum grades laterally into anhydrite-bearing strata and has not been found at depths greater than 35 m below the present surface. Although gypsum is the calcium sulfate form that is valuable for construction materials, an innovative use of anhydrite from the Winnfield salt dome cap rock in northern Louisiana is as aggregrate for temporary and secondary roads. This use is ideal in this humid climate with a near-surface water table, where the hydration of the clast margins and matrix fines welds the aggregate into a tough, flexible roadway that is remarkably durable.

~~~~Salt and Potash~~~~

Halite has many important industrial uses and is produced from a variety of geologic occurrences including direct evaporation of seawater and saline lake water. The majority of salt production is used in the chemicals industry to produce sodium and chlorine chemicals which are used in the manufacture of many thousand products ranging from soaps to insecticides (Lefond and Jacoby, 1983). The principal basic chemicals, chlorine and sodium hydroxide (caustic soda), are produced from brine by electrolysis. Sodium chloride is reacted with limestone in the Solvay process to produce

~~~~Data Center ,09126599985,[email protected], For Educational Uses 486 EVAPORITIC PROCESSES AND MINERAL RESOURCES~~~~

sodium carbonate (soda ash) and with sulfuric acid to produce hydrochloric acid and sodium sulfate (salt cake). Salt is used directly as a food preselvative and seasoning, as well as in a variety of other applications ranging from dye manufacture to textile processing. A major use of salt in northern climates is for deicing roads during the winter. Man-made caverns in bedded salt units and in salt domes are also used for the storage of a variety of chemical substances. These openings may be those left from conventional or solution mining of salt, but recently storage caverns have been specifically designed and excavated to provide the most stable configuration for long-term storage. Solution-mined caverns are utilized for storage of several types of unrefined and refined petroleum products. Recent investigations have evaluated the suitability of bedded salt units and salt domes for nuclear and toxic-chemical waste storage facilities. Major North American salt-producing units (Fig. 6.1) includes Silurian and Devonian strata of the Michigan and Appalachian basins (Fig. 6.2C,D), Devonian evaporites of the Western Canada Basin, Mississippian strata of the Maritimes Basin, and Permian evaporites of the Permian Basin. Evaporite sequences of Permian age also are important sources of salt in Europe and Asia (Harben and Bates, 1984). Intensely deformed and recrystallized halite is also produced from salt diapirs derived from the Jurassic Louann Salt in the Gulf Coast, as well as from the Permian Zechstein evaporites in western Europe. Salt is recovered either by hard rock mining, generally by room-and-pillar methods, or by solution mining, i.e. by injecting water down wells to dissolve the halite, then pumping the artificial brine back to the surface where the salt is recovered by evaporation under pressure. Solar evaporation of seawater and saline lake water in shallow ponds is an important local source of salt. Some modern saline lake waters, relict lacustrine brines, and deep subsurface brines are chemically processed, but compounds other than salt generally are the principal economic products. Potassium-bearing minerals, ores, and processed products are generally referred to as "potash", a term which owes its origin to early processing techniques in which potassium carbonate was produced by leaching hardwood ash in an iron pot. Potassium is widely distributed in igneous, sedimentary, and metamorphic rocks and forms 2.6% of the earth's crust (seventh most abundant element). Furthermore, potassium, along with phosphorous and nitrogen, are the three primary plant nutrients. Not surprisingly, about 95% of all current potash production is used in the manufacture of commercial fertilizers. The remainder is used in soaps, glass and ceramics, drugs, and chemicals. Most commercial potash concentrations are predominantly in salt beds that formed during many geological periods by evaporation of shallow seas; the remainder are

~~~~Data Center ,09126599985,[email protected], For Educational Uses SALT AND POTASH 48 7~~~~

largely in modern salt lakes and natural brines commonly developed from the evaporation of Pleistocene freshwater lakes. These brines are variable in their ionic contents and ratios and may contain as much as 3.2% K,O. A variety of potash minerals occur in natural evaporite deposits (Holser, 1979). Carnallite (KCI*MgC1,*6H2O)generally is the primary mineral in potash deposits, and sylvite (KCI) probably is derived through the alteration of carnallite in most cases. Langbeinite (K2S0,*2MgS04)is an important ore mineral in some areas, including the Carlsbad district, and is believed to result from the alteration of halite and kainite into an assemblage of langbeinite, sylvite, and kieserite, or of sylvite and kieserite into langbeinite and halite (Sonnenfeld, 1984). Kainite (4KC1*4MgS04*11 H,O) occurs as an accessory mineral in potash ores, but is not nearly as abundant in evaporite sequences as would be predicted by equilibrium precipitation accompanying evaporation of seawater (Eugster et al., 1980). As many as 70 other potash minerals occur naturally or appear in brine treatment, but their economic significance is negligible. Common salts, such as halite, gypsum, anhydrite, and kieserite, are associated with potash evaporites. Commercial potash products are sold in a variety of mineralogical and physical forms. Potassium chloride, which as sylvite is the chief naturally occurring economic mineral, is known as muriate of potash in the trade. North American muriate concentrates prepared by flotation are pink in color, due to the presence of occluded hematite. Similar concentrates made by dissolution and recrystallization or by evaporation of brines are white. Potassium sulfate products include langbeinite, the naturally occurring double sulfate (K,S0,*2MgS04) and synthetically produced potassium sulfate (K2S04).Potassium nitrate (KNO,) is produced by reacting potassium chloride with nitric acid. About 5% of the U.S. consumption of potash is as potash sulfate, which is preferred for tobacco, potato, sugarbeet, and citrus crops to prevent chloride burning. The vast majority of fertilizer consumed is muriate of potash blended with nitrogen and phosphorus compounds. The average potash application to crops in the United States is about 30 pounds of K,O per acre (3.3 kg/1000 m’). About 3% of domestically consumed muriate of potash is converted by aqueous electrolysis into potassium hydroxide for use in the manufacture of chemicals for use in the chemicals and ceramics industries. Major end uses are soaps and detergents, glass and ceramics, chemical dyes and drugs, and liquid fertilizer (Adams and Hite, 1983). Large potash reserves occur in Devonian strata in Saskatchewan and in Mississippian strata in the Maritime provinces of Canada. Total recoverable resources of potash in Saskatchewan are conservatively estimated at 67 billion metric tons of K,O of which 4.5 billion metric tons can be extracted using conventional mining techniques,

~~~~Data Center ,09126599985,[email protected], For Educational Uses 488 kVAPOKI'l'IC PKOCESSES AND MINERAL RESOUKCES~~~~

and the remainder can be extracted using conventional solution mining techniques (Harben and Bates, 1984). 1 he Middle Devonian evaporite-rich strata of the Elk Point Basin occupy an area about 325 km wide and 1,300 km long, extending from North Dakota through Saskatchewan to Alberta (Holter, 1969; Fuzesy, 1984; Meijer Drees, 1986; Fig. 6.4). Significant potash deposition appears to have been confined to the Saskatchewan subbasin where the Middle Devonian Prairie Evaporite is as much as 200 m thick. The Prairie Evaporite is the youngest of three major Devonian evaporite cycles and is composed of halite and anhydrite in the lower half of the formation, overlain by three potash-rich members and intervening halite beds that total 60 m in thickness. 'I'he potash-bearing units are, in ascending order, the ksterhazy, Belle Plaine, and Patience Lake members. Each member contains potash beds up to 6 m thick, separated by barren or potash-poor halite beds. Carnallite is more abundant in the northern part of each member, and sylvinite increases to the south and reaches a maximum of 30% sylvite. The upper limit of the carnallite zone is a distinct contact with the overlying

Fig. 6.4. Distribution of the Middle Devonian Elk Point Group (hachured outline) and isopach of the Prairie Evaporite in the Elk Point Basin. western Canada and contiguous United States. Area of potash-hearing evaporites within the Saskatchewan sub-basin is shown in ruled pattern: contours in meters. Modified after Holter (1969).

~~~~Data Center ,09126599985,[email protected], For Educational Uses SALT AND POTASH 489~~~~

sylvinite and halite units that can be traced regionally and is independent of bedding. Potash is recovered by solution mining from the Belle Plaine member and by conventional mining from the Esterhazy and Patience Lake members near Saskatoon. The Prairie Evaporite ranges in depth from 600 m along its northern edge to 2,750 m at the international boundary. The potash-bearing Middle Devonian strata extend into the United States and underlie about 36,000 km2of the North Dakota-Montana section of the Williston basin at depths of 1,800 to 3,650 m. Potash occurs in several beds up to 6 m thick within a thick halite sequence (Anderson and Swinehart, 1979). Most of the United States’ potash production, reserves, and resources occur within Permian evaporitic strata of the Delaware Basin that covers some 100,OOO km2 in southeastern New Mexico and western Texas (Jones, 1972; Fig. 6.5A). Potash-bearing evaporites occur within the Ochoan Salado Formation, an evaporite sequence that is as much as 700 m thick. The area’s most abundant potash mineral is polyhalite, but minable sylvinite occurs in a region near Carlsbad, New Mexico, that is only about 10% of the polyhalite-bearing area (Fig. 6.5A). At least 60 potash layers have been encountered of which 11 ore zones within a 100-m section in the middle Salado are significant (Fig. 6.5B); four of these have provided most of New Mexico’s production. The productive zones are 250 to 550 m below the surface and are between 1 and 5 m thick in the mining areas (Fig. 6.6). Zone 1 has accounted for about 80% of the district production but is essentially depleted. Langbeinite ore is taken from zone 4, and mixed sylvite-langbeinite ore is mined from zone 5. Sylvinite ore is also being extracted from zones 3, 7, and 10. Lowenstein (1988) documented that the present potash-bearing Salado evaporites are products of primary (including early diagenetic) and secondary processes that result in complex relationships of minerals, textures, and fabrics. Vertical cycles represent the evolution of the primary chemical environment from initial relatively dilute waters to halite-saturated brines in a shallow, marginal marine basin. The primary sedimentary environment represents a shallowing-upward sequence that began with a shallow perennial lagoon or lake at alkaline earth carbonate saturation and evolved to gypsum saturation. The cycle culminated in a dry salt pan and saline mudflat at halite saturation. Muddy halite zones represent the effects of meteoric inflow. These primary features and minerals have been extensively modified by the precipitation of secondary minerals, including polyhalite, kieserite, carnallite, sylvite, langbeinite, kainite, and leonite. The timing and conditions for the replacement of the original halite-dominated mineral assemblage by secondary potash-bearing minerals has not been established (Lowenstein, 1988). Potash-bearing evaporites occur within the Paradox Member of the Middle Penn-

~~~~Data Center ,09126599985,[email protected], For Educational Uses 490 EVAPORITIC PROCESSES AND MINERAL RESOURCES~~~~

~~~~NW SE - Norlhweslern Shelf 1 Reef ZoneDewey I Lake redbedrDelaware Basin~~~~

~~~~--. Rustler Fm Solado Fm -i'-----.--.._.._._..~~~~____ ,Bedded deposits ',~~~~

~~~~Capif,~~~~

~~~~100 Bell Canyon Fm , 0 20 km ...... z,, Polyhalile - - - Soluble K %Its I I~~~~

Fig. 6.5. Permian geology of the Delaware Basin, New Mexico and Texas. (A) lsopach map o,f Ochoan strata; contours in meters. Polyhalite-bearing strata shown in the stippled pattern; sylvite- and langbeinite-bearing strata in black. Modified after McKee et al. (1967). (B) Generalizedgeologic cross-section ofthe Permianstrata in the Carlsbad area, southeastern New Mexico, showing the distribution of potash-bearing strata in the Salado Formation. Modified after McKee et al. (1967). sylvanian Hermosa Formation which underlies about 8,000 km' of the Paradox Basin in southeastern Utah (Fig. 6.1). Twenty-nine evaporite cycles have been identified in the central part of the basin of which 19 contain potash (Hite, 1968). These cycles include limestone, dolomite, black shale, anhydrite, and halite, with or without sylvite. Much of the Paradox basin has undergone major folding which has severely disturbed the potash beds. The original thickness of the Paradox Member reaches 2,300 m in the central part of the basin, but the evaporite zone has been thickened to as much as 4,700 m in the diapiric anticlines. Exploration in four areas has delineated a potential total resource of 1.8 billion metric tons of K'O. Potash is solution mined from a 3.5-m thick sylvite zone within salt unit 5 in the upper part of the Paradox Member at a depth of 1,000 to 1,200 m in the Cane Creek area near Moab. A potentially commercial sylvinite deposit occurs within salt unit 13 on the northwestern flank of the basin. This deposit is essentially horizontal and undeformed and occurs at a depth of about 1,800 m; it represents a significant potash resource that would most likely be extracted by solution mining (Adams and Hite, 1983). High-grade sylvinite deposits were discovered in Mississippian Windsor Group strata in New Brunswick in 1971 (Figs. 6.3,6.7). The Carboniferous geology of the Mari-

~~~~Data Center ,09126599985,[email protected], For Educational Uses SALT AND POTASH 49 1~~~~

~~~~ORE ZONE GAMMA RAY NEUTRON~~~~

~~~~11 I~~~~

~~~~lo I~~~~

~~~~71 6.~~~~

~~~~31 PI~~~~

~~~~El POLYHALITE ond HALITE ARGILLACEOUS ANHYDRITE HALITE~~~~

Fig. 6.6. Stratigraphic column of the McNuttpotash zone of the Carlsbad district showing gatnma ray and neutron logs. Modified after Jones (1972). times is characterized by a series of northeast-trending rift basins filled with sedimentary sequences up to 10 km thick (Howie, 1988). Three areas, the Moncton, Cumberland, and Central basins, appear to have significant potash potential (Webb, 1984). Current potash production is restricted to the Moncton Basin where three potash deposits have been defined. Two mines are in production at depths of 600 to 1,OOO m. ‘The potash-bearing beds are as much as 21 m thick with grades of 26-28% K,O (Raulston and Waugh, 1983). The Penobsquis (Plumweseep) deposit occurs within a large salt pillow structure in which the potash-bearing strata dip at about 45” away from a core

~~~~Data Center ,09126599985,[email protected], For Educational Uses 492 EVAPORITIC PROCESSES AND MINERAL RESOURCES~~~~

zone of older salt (Fig. 6.7). This atypical geometry permits the separate extraction of both potash and salt beds using a relatively simple mining plan. The evaporite sequence is overlain locally by a complex solution collapse breccia zone (Fig. 6.7). This material is believed to be the residue resultingfrom the dissolution of the evaporite sequence by groundwater. The Windsor Group strata are atypical for marine evaporite deposits in that they contain avaried local borate mineral assemblage (Raulston and Waugh, 1983). The association with continental environments is further suggested by the "red bed" composition of the pre- and post-evaporite depositional sequence. In addition to these ancient deposits, saline commodities are produced from a variety of modern continental lakes and their Tertiary analogs. The evolution of these closed basins and their contained brines was reviewed by Hardie and Eugster (1970), Eugster and Hardie (1978), Eugster (1980a), and Hardie (1984). Subsurface transport of solutes in ground water and discharge within the depositional basin via warm springs may be particularly important to commercial evaporite mineral deposition in this setting as these processes allow the accumulation of relevant ions without the sediment and fluid dilution effects that would accompany the discharge of permanent streams into the basin. Probably the best known of the modern saline lakes is the Great Salt Lake in western Utah which covers a variable area between 2,500 to 5,000 km'with an average depth of about 10 m. The present lake is a remnant of ancient Lake Bonneville, a Pleis-

~~~~EXPLANATION REDBEDS =POTASH SOLUTION COLLAPSE BRECCIA 0HALITE ANHYDRITE BASAL CARBONATE / SULFATE~~~~

~~~~Fig. 6.7. Geologic cross-section of the Penobsquis (Plumseweep)potash deposit of the Moncton Basin, New Brunswick. After Kingyton and Dickie (1979).~~~~

~~~~Data Center ,09126599985,[email protected], For Educational Uses BORATES 493~~~~

tocene freshwater lake that is believed to have covered over 50,ooO km2with a maximum depth of 300 m (Spencer et al., 1984). The past 30,000 year history of the Great Salt Lake provides a well documented example of the sensitive balance between inflow and evaporation in closed basins. The ancestoral lake at 30,000 years was a relatively small, shallow saline body with similar features to the modern lake (Spencer et al., 1985). In response to climatic changes, the basin began to fill and in several stages reached its maximum areal extent, depth, and freshest water lake conditions about 16,OOO years ago. Outflow and evaporation caused the lake to shrink to near its present level about 15,000 years ago. Even within the historic time frame from 1847 when the Mormon settlers arrived, the lake level and extent has fluctuated markedly from an historic high in 1873 at an elevation of 1,283.7 m with an area of about 6,500 km2 to a low in 1963 at 1277 m with an area of only about 2,300 km2. The lake has been rising at a rapid rate in response to increased inflow and is predicted to exceed the historic high and perhaps attain the 1,285-m elevation that will cause natural outflow on the western side toward the Bonneville Salt Flat. The commercial activities in the Great Salt Lake area are markedly affected by the variability in the lake level and chemistry. The construction of a railroad causeway in 1959 caused the lake to be divided into northern and southern segments, each with its own physical and chemical characteristics. Salinity has markedly increased in the northern arm at the expense of the southern segment. Several operations recover salt via solar evaporation ponds, and one major operation also recovers potassium and magnesium chlorides and sodium sulfate. Extraction of bromine and iodine are technically feasible, but economically unattractive at present. The Bonneville Salt Flats are another remnant of Lake Bonneville and consist of salt or mud flats impregnated with a saturated or nearly saturated brine. A 1-m top layer of hard-packed salt is underlain by a layer of fissured clay several meters thick. The brine below the salt contains about 1% K20which is consistently higher than surface brines in surrounding areas. Muriate of potash is produced by a combination of solar evaporation and flotation (Adams and Hite, 1983).

~~~~Borates~~~~

Borates have a long history of commercial use, dating from the use of the mineral in the mummifying processes in Egyptian and other ancient cultures. Commercial concentrations of boron are rare, and current production is dominated by the United States and Turkey. Borates have a wide variety, and ever increasing number, of industrial uses, but the major end uses are in the glass and ceramic industries. The

Data Center ,09126599985,[email protected], For Educational Uses 494 EVAPOKI'I'IC PROCESSES AND MINERAL RESOURCES major borate minerals contain calcium or sodium with variable amounts of bound water molecules. Borax (Na2B40, 10H20), ulexite (NaCaB,O, *8H20), and colemanite (Ca,B,O,, *5H,O) generally are the most important ore minerals, although the mineral assemblage varies greatly among deposits. The general genetic model for borate mineralization involves a combination of geologic factors including the initial supply of boron through local volcanic activity, the accumulation of boron within a lacustrine body of water, evaporation of the lake to the point of borate mineral precipitation, and preservation of the highly soluble borate minerals by a protective layer of sediments. The inital borate mineral precipitates, e.g. borax, are likely to be converted to borates with less water content, e.g. kernite, during burial diagenesis, but these are highly susceptible to rehydration during weathering and exposure to shallow circulating meteoric fluids. Most of the commercial borate deposits are mid-Tertiary or younger in age, with the alkaline lakes of the southwestern United States and northwestern Turkey providing many notable examples. Searles Lake is one of several alkaline lakes in southeastern California that are remnants of much larger freshwater lakes that existed in the Pleistocene (Figs. 6.8,6.9). Today Searles Lake is a mud and sand flat underlain by two superimposed crystal bodies of mixed salts saturated with concentrated brine (Fig. 6.10). The modern formation waters of Searles Lake are important sources of several industrial minerals including potash, sodium carbonate, sodium sulfate, and borates. The upper crystal/brine body, from which potash is obtained, averages 21 m thick and has a total area of about 100 km'; the brine also contains 1% B,O,. The lower crystal/brine body, which is 10 m thick, is separated from the upper by 3 to 4.5 m of impervious clays; the brine from this level is similar to the upper brine but contains less potash and more sodium carbonate and borax (1.2% B,O,). Searles Lake borate reserves are estimated to be more than 30 million metric tons of recoverable B203(Kistler and Smith, 1983). The evaporites and brines are thought to be the result of desiccation and concentration of salts supplied by thermal springs to a series of late Quaternary lakes that extended along the east front of the Sierra Nevada (Fig. 6.8). The ultimate source of boron and associated elements is believed to be the felsic volcanic rocks associated with the Long Valley caldera to the north. Stratification of the denser Bcontaining brines occurred during several periods in the 140,OOO year history of Searles Lake (Fig. 6.10), while the fresher waters flowed into Death Valley, the last lake in the chain (Figs. 6.8,6.11A; Smith, 1979). Felmy and Weare (1986) modeled the borate and associated mineral assemblage of the upper brine zone as an end product of the evaporation of present-day Owens River water. Tertiary analogs of the depositional environment at Searles Lake are major sources

~~~~Data Center ,09126599985,[email protected], For Educational Uses UO K A'I'kS 495~~~~

~~~~I I~~~~

~~~~SIERRA NEVADA 1000 (m)~~~~

~~~~Level T)EATH~~~~

Fig. 6.8. Pleistocene and modern intermontane lakes of southeastern Caltfornia. (A) Pleistocene and modern drainage pattern. (B) Topographic profile from Owens Lake to Death Valley. Modlfied after Smith (1966. 1979). of bordtes in California and 'I'urkey (Fig. 6.11; Kistler and Smith, 1983). 'I'he elevated boron content of these deposits is believed to have been supplied from contemporary volcanic rocks by hot springs activity. The Boron deposit near Kramer, California, consists of a lenticular central mass of borax (Na,B4O,*l0H,O) and interbedded clays that is about 1600 m long, 800 m wide, and 100 m thick (Fig. 6.11C,D). Afacies change to ulexite ( NaCaB50,*8H,0) and clays takes place laterally andvertically away from the borax core. 'I'he deposit is believed to have formed in a shallow Middle Miocene lake fed by Na- and B-rich thermal springs that were associated with the late stages of local volcanism. 'I'he importance of hot springs in the formation of the borate deposit is fur-

~~~~Data Center ,09126599985,[email protected], For Educational Uses 496 EVAPOKI'I'LC PKOCESSES AND MINERAL KESOUKCES~~~~

Fig. 6.9. Aerial view of the geologic setting of modern alkaline lakes in the Owens Valley-PanamintValley-Death Valley area of southeastern California. Approximate area of' photo shown by dashed outline on Fig. h.XA. Note typical Basin-and-Range topography with closed basins including Searles Lake (S). Shuttle (Mission 3)photograph provided by William R. Muehlberger.

ther suggested by local mineral concentrations containing arsenic, antimony, and other elements that are commonly associated with "epithermal" metallic mineralization. 'l'he lake basin was controlled by faulting contemporaneous with borate deposition; continued movement along these faults resulted in the lake sediments being buried under at least 300 m of late Miocene-Pliocene arkosic sediments (Fig. 6.1 1C). Plio-Pleistocene uplift resulted in erosion and hydration of part of the deposit (Kistler and Smith, 1983). 'I'urkey is the worlds largest producer of borates, principally from several deposits in the northwestern part of the country. 'I'he deposits at Uigddic, Emet, and Kirka each contain tens of millions of metric tons of borate reserves. 'l'hese deposits occur in Upper Pliocene or Middle Oligocene lacustrine sediments which are interlayered with volcanic sediments and flows. Capping clay-rich sequences are important in the

Data Center ,09126599985,[email protected], For Educational Uses SODlUM CAKBONA'I'E AND SODIUM SULFA'l'E 49 7 preservation of the soluble borate minerals. Dehydration accompany burial diagenesis of the original borate minerals has resulted in the development of colemanite and kernite concentrations (Kistler and Smith, 1983).

~~~~Sodium Carbonate and Sodium Sulfate~~~~

Although the bulk of the world's soda ash is produced synthetically by the Solvay process which utilizes salt and limestone as the raw materials with an ammonia catalyst, production from trona deposits and evaporitive brines is becoming increasingly important. Soda ash (Na,CO, ) is produced by combining calcined trona (Na,CO,*Na- HC03*2H,0) with water, removing the solids, and evaporating the solution. Soda ash is used extensively in glass manufacture, chemicals, pulp and paper production, and many other uses.

~~~~DEPTH WATER DEPTH (rn)~~~~

~~~~SALINE UNITS Halite, trono, honkrite. borax, goylurite, pirssonite, 0 northupile, nohcolile, burkeila CLAY-RICH UNITS 0 f;;;: o;~~;;~~muds containing inkgrown~~~~

~~Fig. 6.10. Stratigraphic column of the upper 60-m of Recent and Pleistocene sediments below the surface of Searles Lake relative to depositional water depths. Modified after Smith (1966, 1979).~~

~~~~Data Center ,09126599985,[email protected], For Educational Uses 498 EVAPOKITIC PROCESSES AND MINERAL KESOUKCES~~~~

Fig. 6.11. Geologic features of modern and f ossil borate deposits. (A) View looking west of the alkaline flat within the Death Valley graben. Note the large alluvial fans sourced from the upthrown Panamint Range. (B) Active borate-depositing spring. Antuco. Argentina. Photograph provided by Robert B. Kistler. (C) Open pit showing Miocene arkosic sandstone overlying the Kramer borate orebody. California. (D) Upper Kramer ore zone overlain by ulexite-bearing claystone unit.

Lake Magadi and other lakes in the East African Kift contain substantial sodium carbonate resources as brines and trona beds whose origin is related to alkalic volcanism along the Rift (Figs. 6.12, 6.13). The lake is located within block-faulted Pleistocene trachyte flows and was preceded by two earlier lakes which were fresher and formed lake beds of volcanic sediments and Magadi-type chert beds (Fig. 6.12). Lake Magadi consists of a large ephermeral salt pan, saline and dry mudflats, ephemeral stream deposits, and small alluvial fans (Figs. 6.13A-D). Water enters the lake from ephemeral streams and through perennial alkaline hot springs which are driven by the high geothermal gradient and the deep ground-water reservoir (Fig. 6.13B). Efflorescent salt

Data Center ,09126599985,[email protected], For Educational Uses SODIUM CAKBONA'I'E AND SODIUM SULEA'I'E 499 crusts form wherever the groundwater is close enough to the surface to result in capillary draw in porous sediments related to the wetting and drying cycles in this arid climate (Fig. 6.13C). l'he precipitation of alkaline earth carbonates and silica and the dissolution of efflorescent crusts are believed to be the dominant mechanisms controlling brine evolution that results in trona deposition (Eugster, 1970, 198Ob, 1986). Lake Magadi contains billions of metric tons of trona which are believed to have formed within the last 9,ooO years. Trona beds up to 40 m thick have been intersected in shallow drill holes (Eugster, 198ob). Because the trona is presently forming, the geologic setting and chemistry of Lake Magadi have markedly affected the development of general genetic models for trona deposition (bugster, 1986). 'I'rona saturation is a- chieved when the Lake Magadi lagoonal brine is only a few centimeters deep. As evaporation proceeds, trona grows as long blades upwards from the substrate and as a thin film floating on the surface. 'lrona precipitation results in bicarbonate depletion

Fig. 6.12. Generalized geologic map and schematic geologic cross-section of the Luke Magadi area within the East African Rift. Both saline lagoons and salt pans are shown in black on the cross-section. Modified after Baker (I 986) and Eugster (1986).

~~~~Data Center ,09126599985,[email protected], For Educational Uses 500 EVAPOKITIC PROCESSES AND MINEKAL, KESOUKCES~~~~

Fig. 6.13. Geologic features of modern and ancient trona deposits. (A) 7'hr trona deposits of Lake Magadi. (B) Streani ,fed hy hot springs that is activelv depositing tronu. Lake Magadi. (C) Efflorescent trona crirst. Luke Magadi. (U) C'htjrt ufter magadiitr. Luke Magudi. (E) Homogenous mirdstonr with eiihedrnl shortite crystals. Green River Formation. Wyoming. (F) Clear crvstalline columnar vein frona introdircrd throirgh hydraulic fracturing. Green River Formation. Wyoming. (G) Nodiilar crvstalline trona overlain by sugary textured tronu. Green River Formation. Wyoming. Lake Muyadi photographs provided tiv Duniel S. Barker: Green River surnple.s provided by Stuart Birnbaunz.

~~~~Data Center ,09126599985,[email protected], For Educational Uses SODIUM CARBONATE AND SODIUM SULFATE 501~~~~

and a rise in pH above 11; temperatures of the brine below the surface crust may rise to 60°C. The resulting trona deposit is banded on a 2-5 cm scale with layers composed of tightly spaced rosettes of upward-pointed trona blades separated by thin dark bands of trona containing windblown dust. During the dry season, the level of interstitial brine within the trona layers falls several meters below the surface. During the rainy season, flooding of the trona surface results in partial dissolution of the previouslyformed trona. Only minor chemical additions are made to the lake brines at any point through the perennial alkaline hot springs and the majority of components in any trona-depositing event is the result of elemental recycling. The renewed lake waters evaporate after a few months and the surface trona layer is precipitated. Net trona accumulation rate has been estimated at 2-3 mm/year (Eugster, l980b, 1986). Magadiite (NaSi,0,,(OH),*3H20)is the most common of several sodium silicates that forms beds in the lake area; Na-Al-silicate gels are also present locally. Conversion of magadiite to bedded chert takes place close to the surface and is related to the leaching of sodium by dilute waters in shallow environments or spontaneous conversion to silica in deeper brine-saturated zones (Fig. 6.13L)). Magadiite beds are white, soft, and plastic as long as they are soaked by the alkaline lake brines. Drying and conversion of magadiite to chert is accompanied by a considerable reduction in volume. 'l'hese "magadi-type"cherts show abundant effects of soft sediment deformation and shrinkage cracks (Eugster, 1970, l980b, 1986). 'l'hese types of cherts have been identified in older alkaline lake environments. 'I'he major North American source of sodium carbonate is the extensive lacustrine trona beds of the Eocene Green Kiver Basin in southwestern Wyoming (Fig. 6.14; Mannion, 1983). 'I'he Green Kiver Basin is an intermontane basin that originated during the Paleocene; the basin was occupied by an extensive Miocene body of water (Lake Gosiute) that varied considerably in size and chemical character during its 4 million year history (Bradley and Eugster, 1969; Fig. 6.14A). The Wilkins Peak Member of the Green River Formation includes thin persistent beds of oil shale, trona, and halite, interbedded with siliciclastic sediments (Figs. 6.13E-G, 6.14B). 'I'he oil shales represent deposition from relatively dilute water during periods of expanded volume of Lake Gosiute, whereas the trona beds were deposited in the interior part of the basin during periods of restricted circulation and volume (Surdam and Wolfbauer, 1975; Fig. 6.14B). At least 42 trona beds are present at depths ranging from 120 to over 1,ooO m. Individual beds are as much as 11 m thick and underlie as much as 2,200 km'. Total trona reserves are estimated at 90 billion metric tons. Volcanic ash is believed to have been responsible for the sodium enrichment of the lake waters through thermal springs and surface runoff (Fig. 6.15; Parker and Mannion, 1971; Eugster and Hardie, 1Y75;

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~~~~WILKINS PEAK STAGE (moxbmum extent)~~~~

~~~~__ -WYOMING--- __ COLORAOO /~~~~

~~~~I I~~~~

~~~~6. LAKE EXPANSION LAKE CONTRACTION PERIPHERY INTERIOR~~~~

~~~~MUDCRACKED I I- c LAMINATED 1 1 1 LIME MUDSTONE E I 0 B OIL SHALE 1 2 1 p, FLAT PEBBLE CONGLOMERATE~~~~

Fig. 6.14. Green River trona deposits, Wyoming. (A) Inferred drainage basin and maximum extent of Lake Gosiute including the area of trona deposition in the Eocene Green River Formation. Modified after Bradley (1964). (B) Cyclic units within the Wilkins Peak Member. I. Transgressive cycle representing expansion of the lake margin. 2. Regressive cycle representing contraction within the basin interior culminating in trona deposition. Modified after Eugster and Hardie (1975).

Surdam and Wolfbauer, 1975). The supply of sodium sulfate, commercially known as salt cake or Glaubefs salt, from natural and synthetic sources essentially repeats the case of sodium carbonate. Natural sodium sulfate is produced largely in North America, generally from crystalline beds of mirabilite (Na2SOp10H20)or as brines from saline lakes. Significant sodium sulfate production from brines includes Chaplin Lake, Saskatechewan, as well as the Great Salt Lake and Searles Lake in the United States. Sulfate is produced at Ingebrigt

~~~~Data Center ,09126599985,[email protected], For Educational Uses BROMINE, IODINE, LI'I'HIUM AND NI'I'KOGEN 5 03~~~~

Lake from Holocene mirabilite beds up to 40 m thick, either by direct mining of the permanent beds or by solution mining of deeper zones (Broughton, 1984). The Great Plains of western Canada and northern United States contain hundreds of saline and hypersaline lakes with surface areas up to 300 km'. I he Saskatchewan deposits occur in shallow, closed lakes developed on glacial sediments in this semi-arid climate (Last and Schweyen, 1983; LdSt and Slezak, 1987; Last, 1989). Ground water leaches elementsfrom the soils and carries them in solution to the basins. The process may be largely SUbSUrfdCe in which the basins are fed by springs associated with buried, pre-glacial river valleys. Evaporation during the summer concentrates salts in the lake waters, and cooling of the water during the winter results in deposition of mirabilite crystals on lake bottoms. Salt concentration and mirabilite precipitation may also be effected by ice formation. Daily or seasonal warming of the lake brines will cause mirabilite to go into solution, hence the importance of clay-rich sediment layers in preservation of permanent sodium sulfate beds (Broughton, 1984). Similar deposits occur in contiguous Alberta and North Dakota.

~~~~INTENSE EVAPORATION~~~~

~~~~GRAVELS MUDS verticol scale'~~~~

~~~~Fig. 6.15. Block diagram representing the general depositional environment of the Wilkins Peak Member. Modified after Eugster and Hardie (1975).~~~~

~~~~Bromine, Iodine, Lithium and Nitrogen~~~~

Bromine, iodine, and lithium are other valuable elements for a variety of specialized industrial applications whose concentrations reach economic amounts in some modern lake and formation waters. Bromine is recovered from Searles Lake and the Dead Sea, as well as modern formation waters from the Gulf Coast and Michigan

~~~~Data Center ,09126599985,[email protected], For Educational Uses 504 EVAPORI'lIC PROCESSES AND MINERAL RESOURCES~~~~

basins (Jensen et al., 1983). A brine averaging about 5,000 ppm Br is recovered from the upper Reynolds Member of the Upper Jurassic Smackover Formation in southern Arkansas. The Smackover is a highly permeable ooid grainstone unit that IS an extremely prolific oil reservoir in the Gulf Coast. The brine, including its elevated bromine content, is believed to have been derived from the underlying Louann Salt (Carpenter and 'I'rout, 1978). 'I'he Dead Sea occupies the deepest part of the Dead Sea Basin that occurs along a major continental rift zone. 'I'he lake has an average salinity of 25.7% and is at or near saturation with regard to calcium and magnesium carbonate, calcium sulfate, and halite; carnallite crystallizes when the water volume is halved. Several commodities are recovered through solar evaporation in salt pans. 'I'he lake waters contain 5,200 pprn Br which is recovered from carnallite-producing brines (Carpenter and 'I'rout, 1978). Iodine is recovered, generally as a co-product with other brine components, from formation waters from Paleozoic clastic strata in the Anadarko and Michigan basins (Jan and Roe, 1983). The iodine content of these brines ranges from less than 50 to as much as 1,500 ppm. Iodine is also recovered from brines containing about 150 ppm iodine that are associated with natural gas reservoirs within Miocene and younger clastic sediments in Japan. Lithium attains economic concentrations averaging 200 ppm in formation waters within lacustrine sediments of a modern playa in Clayton Valley, Nevada, and is believed to have been supplied from 'I ertiaq felsic volcanic materials (Kunasz, 1983). 'I'he brines of the Great Salt Lake, Searles Lake, the Imperial Valley geothermal field, and the Smackover Formation contain lithium in non-commercial amounts under present economic constraints, but constitute important resourcesfor the future. Lithium also occurs in lithium-bearing clays, principally the mineral hectorite, which typically result from the alteration of volcanic materials by alkaline brines in the playa lake setting. The interstitial brines of the Salar de Atacama, one of about 75 playa lakes in the Atacama region in northern Chile, is a major source of lithium and other salts. 'I'he brines have an average lithium content of 1500 ppm with total resources of 4.3 million tons of lithium (Kunasz, 1983). A major solar evaporation production facility for these brines is scheduled to be completed by 1992 to recover potassium chloride, potassium sulfate, lithium salts, and boric acid (Anonymous, 1989). Nitrogen is abundant on and near the earths surface, forming 78% of the atmosphere and forming an essential aspect of the biosphere. It, along with potassium and phosphorous, is one of the three primary plant nutrients. Natural deposits of nitrogen-bearing minerals are rare because of the extreme solubility of nitrates in water. Increased use of nitrogen products in the early part of this century in the manufacture

~~~~Data Center ,09126599985,[email protected], For Educational Uses METALLIFEROUS SALINE FOKMA'I'ION WA'I'EKS 5 0s~~~~

of explosives, nylon, plastics, and resins, in addition to agricultural uses, resulted in the development of processes for the production of synthetic nitrogen compounds. Although natural sources of nitrogen are far surpassed by synthetic material primarily produced by the Haber-Bosch process, natural nitrate deposits in northern Chile are still mined. The ores have an average composition of 7-10% NaNO,, 4-10% NaCI, 10-3OC7o Na,SO,, and 2-7% Mg, Ca, K, Br, and I. These unique deposits owe their origin to the climate of the host Atacama Desert, one of the driest regions of the world. 'l'his exceptional climate has prevailed since the middle Miocene and has permitted the concentration of various saline materials from several sources, including Ocean sprav, volcanic emanations, rock leaching, and airborne volcanic and playa dust. Ericksen (1981) further suggested that nitrogenous materials were transported and deposited from the atmosphere as dry fallout and as fog-condensate solutions, and were concentrated by dissolution in rainwater prior to redeposition in soil or salt pans. Commercial concentration and preservation of these highly soluble nitrate deposits was a function of the consistently dry climate for a long period of time.

~~~~EVAPORITES AND THE ORIGIN OF ORE-FORMING SOLU'I'IONS~~~~

~~~~Metallif erous Saline Formation Waters~~~~

Metal-rich saline formation waters have been identified in a number of areas worldwide and have been accepted as modern analogs of ore-forming fluids for a variety of classic ore deposits types ('l'able 6.1B). Although the detailed mechanisms by which these fluids achieve their present composition is difficult to establish, it would appear that rock-water interactions involving formation waters that have elevated salinities are generally important (Land, 1987). Therefore, evaporite dissolution in the subsurface may have a marked influence on the development of ore-forming solutions in a variety of geologic settings. Metal-rich "oil field brines" have been identified in a number of areas of the Gulf Coast of southern North America (Fig. 6.16; Carpenter et al., 1974; Carpenter and Trout, 1978; Land and Prezbindowski, 1981: Hanor, 1987). It is readily apparent that not all saline formation waters are enriched in metals, nor is there a direct relationship with temperature. The most metal-rich formation waters are Ca-rich and generally occur within the Mesozoic sedimentary section. The contained trace metal suite, dominantly zinc, lead, iron, and barium, is compatible with the composition of ore concentrations that typically occur within sedimentary environments, particularly the carbonate-hosted "Mississippi Valley type" and shale-hosted "sedimentary exhalative"

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Fig. 6-16. Generalgeologic setting of the Gulf Coast sedimentary basin showing location of metal-bearing oil field brines (ruledpattern). Geologic setting modified after Martin (1978). (I) CentralMississippi(Carpentereta1..1974): (2) SouthernArkansas(Carpenter and Trout. 1978): (3)South Texas (Land and Presbindowski, 1981). deposits to be discussed in following sections. Furthermore, the present temperature and major ion composition ranges of these formation waters generally correspond to those of the ore-forming fluids for these types of deposits as determined by fluid inclusion studies (Roedder, 1976). Zinc-lead-silver sulfide concentrations have recently been identified in the Jurassic Smackover Formation in southern Arkansas within the area of Br-rich brine production (Kyle et al., 1989). The source of the metals in oil field brines and in ore deposits is difficult to define. Models range from metals stripped from detrital iron oxides in red beds, released during clay mineral transformation, or released during albitization of detrital feldspars during deep burial diagenesis (Land, 1987). The metal-rich formation waters in the Gulf Coast Basin will be discussed further in a separate article on the relationship of mineralization to salt dome development.

~~~~Geothermal Systems~~~~

~~~~Evaporites may also play a significant role in metal concentrations in geologic~~~~

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terranes which at best could be considered to be transitional to typical sedimentary environments. For example, the zinc-copper-silver concentrations in the Atlantis I1 Deep within the axial rift zone of the Ked Sea have precipitated from a density stratified brine pool (Degens and Koss, 1Y6Y; Shanks, 1983). Dissolution of Miocene evaporites on the flanks of the Ked Sea by deeply circulating seawater has been documented as the source of the elevated salinities of the brine pool and perhaps is related to the elevated metal content of the mineralizing fluid (Shanks and Bischoff, 1977). It has been well established that meteoric water is the dominant source of fluids for most geothermal systems in tectonically active continental settings; predictably, most modern geothermal fluids are relatively dilute, generally containing less than 1 wt total dissolved solids (White, 1Y8l). However, the involvement of evaporitic strata with these deeply circulating dilute fluids may have a substantial effect on the chemistry of the final geothermal reservoir fluid. For example, the fluids of the Salton Sea geothermal field have much higher salinities than the fluids associated with most modern or ancient geothermal fields. Dissolution of evaporites in the Salton Trough is believed to be the source of the dissolved ions and has been suggested as the cause of the high metal content of the fluids (McKibben and Elders, lY8S). "Epithermal" mineral deposits in many areas of the world are the product of "fossil"geothermal systems (Berger and Bethke, 1985). Commonly these deposits are associated with explosive felsic volcanism and caldera development in Tertiary continental environments. 'I'hese deposits are particularly well developed in the western United States and Mexico, but occur worldwide typically in association with young volcanic terranes. The deposits classically consist of veins that represent mineralizing fluid conduits along pre-existing normal faults, but the total mineralization system may include surface precipitates and shallow subsurface stratabound concentrations that are the result of hot springs activity. Gold and silver are the dominant economic products of these veins, and they are generally enriched in precious metals, particularly gold, toward the top and become increasing base metal dominant downward. Fluid inclusion data indicate that mineralization took place from low-salinity fluids at temperatures ranging from 200"- 300" C. Stable isotope evidence suggests that these fluids were largely of meteoric water composition. These geologic and chemical characteristics confirm the genetic affiliation with modern geothermal systems. Some classic epithermal veins, e.g. in the San Juans of southwestern Colorado, are associated with late structural events in caldera formation, perhaps related to resurgence. The San Juan volcanic field in southwestern Colorado is the erosional remnant of a huge volcanic field that covered the southern Rocky Mountains in the

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Oligocene. Initial Tertiary volcanic activity from about 35-30 Ma was dominated by construction of andesitic stratovolcanoes. From about 30-26 Ma, the nature of volcanism changed to the eruption of felsic ash flow tuffs derived from at least 18 collapsed calderas. Some of the calderas have been resurgently domed, which was accompanied by intrusion of felsic magmas around the caldera margins. Most of the mineral production has been derived from radial fractures and graben faults along resurgent domes; however, mineralization was considerably younger than local caldera development in most cases and may have been related to emplacement of felsic plutons at depth which resulted in deep convective circulation of meteoric water. Discharge of the heated and chemically modified mineralizing fluid was focused along structural zones, and mineral precipitation may have been effected by fluid boiling at relatively high levels. The Creede district in the central San Juans includes some of the most highly studied epithermal vein deposits (Barton et al., 1Y77; Fig. 6.17A). The district has been a major silver producer, but recent exploration has defined a gold-rich region in the northern part of the district. ‘I’he major veins occur along apical fracture zones of the elongate resurgent dome of the Bachelor caldera (Fig. 6.17A). However, mineralization is believed to be related in time to the younger Creede caldera to the south and to be genetically associated with a buried pluton underlying the Bachelor caldera that provided the heat “engine”for a large convective cell (Fig. 6.17B). Fluid inclusion data indicate that the mineralizing fluids for the Creede district veins were atypical in comparison with most epithermal vein deposits and modern geothermal systems in that they range from 4 to 12 wt % NaCl equivalent (Barton et al., 1Y77). ‘I‘hishighly saline character is compatible with the zinc-lead-silver composition of the deposits. ’I’his anomalous character is believed to be due to the involvement of pore waters from the lacustrine sediments (Creede Formation j of the Creede caldera moat in the mineralizing fluid system. The Creede Formation is predominantly thin-bedded volcaniclastic sediments that accumulated in a playa environment. The lake and pore waters evolved toward an alkaline brine through the combined processes of evaporation and diagenesis of volcanic glasses. Discharge of thermal springs into the moat sediments is indicated by local tufa deposits. The moat sediments also host local disseminated silver concentrations near the caldera margins. Although evaporites have not been reported, some of the tuff beds have been altered to clinoptilolite, and calcite pseudomorphs after gaylusite are common on some bedding surfaces. Hydrologic, geochemical, and stable isotope evidence confirms that Creede caldera lake was the source of relatively saline fluids, whereas dilute meteoric water was supplied from the northern part of the region (Barton et al., 1977; Bethke and Rye, 1979; Fig. 6.17B). Mixing of the fluids within the

~~~~Data Center ,09126599985,[email protected], For Educational Uses ZEOLI’I’IZA‘I’IONIN ALKALINE LAKES 509~~~~

~~~~B. NORTH SOUTH Son LUlP Caldero~~~~

Fig. 6.17. Geology of the epithermal Pb-Zn-Ag vein deposits in Tertiary volcanic rocks of theCreede miningdistrict. southwestern Colorado. ModifiedafterBartonet al. (1977). (A) Generalized geologic map showing the mineralized structures in the Creede graben. (B) Conceptual model of Creede hydrothermal system. 200°C isotherm is shown by the dashed line: ore zone is in the ruledpattern. Fluid-flow patterns for saline fluids sourced from the Creede Caldera moat are shown by heavy arrows: dilute meteoric solution.\ are shown by the dashed arrows. conduits provided by the Bachelor structural zones resulted in the formation of the zinc-lead-silverores,

~~~~Zeolitization in Alkaline Lakes~~~~

~~~~The formation of zeolites through diagenetic alteration of tuffs and volcaniclastic~~~~

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sediments by pore fluids within the alkaline lake environment is a widespread geologic process. In appropriate geochemical and hydrologic settings, these processes may result in the development of economic concentrations of valuable zeolite minerals. 'l'he commercial uses of zeolites are tied to the unusual zeolite structure which consists of three components--the aluminosilicate framework, interconnected void spaces within the framework, and water molecules present as an occluded phase. The zeolite structure results in certain characteristic properties, particularly ion-exchange and adsorption and related molecular sieve phenomena (see Mumpton, 1981, for reviews of the geology and commercial aspects of natural zeolites). For example, the ready exchange of loosely bound ions in the zeolite structure results in the use of zeolites for removing harmful ions from various types of chemical waste products. 'l'he basis for the molecular sieve use is tied to the dehydration of zeolites which results in an interlocking network of micropores in the 2 to 7 A range. Small ions can be preferentially adsorbed into the dehydrated zeolite structure, whereas larger molecules are excluded. 'l'he efficiency of zeolites as molecular sieves is indicated by the surface area available for adsorption which can be as much as several hundred square meters per gram (Harben and Bates, 1984). 'I'hese and many other commercial uses make zeolites increasing valuable industrial minerals. Zeolites are formed by the reaction of pore water with a variety of materials including volcanic glass, clay, feldspars, and silica. 'l'he rate of zeolitization depends on a wide variety of chemical and hydrologic factors, but commonly begins early in the history of the volcanic material and can proceed at extremely rapid rates. 'l'he resulting zeolite mineral assemblage depends on the composition of the starting material and the temperature and chemical character of the pore fluids. Zeolites typically continue to react with evolving pore fluids, and new minerals are formed by the replacement of the initially formed zeolites. Not surprisingly, the greatest quantity and variety of zeolites occur in geologically young deposits which have never been deeply buried. More than 30 different zeolite minerals have been identified in Cenozoic rocks; however, only 7 occur in Mesozoic strata, and even fewer in older geologic ages (Harben and Bates, 1984). Numerous zeolite deposits have been identified and evaluated in the western United States, a function of the abundance of Cenozoic volcanic activity and a tectonic and climatic environment that has produced many closed basins containing alkaline lakes and pore fluids. The deposits in the San Simon Valley in eastern Arizona serve as a commercially significant example (Sheppard et al., 1978). The zeolite deposits occur within an intermontane basin in the Basin and Kange Province. Several zeolite-bearing units crop out within the thick Pliocene to Holocene fluvial and

~~~~Data Center ,09126599985,[email protected], For Educational Uses BIOEPIGENE'I'IC SU LFUK DEYOSI'T'S 511~~~~

lacustrine basin fill, but the most significant is an altered tuff bed that locally is a meter or more in thickness. Chabazite, Ca?l(Alo,),(Si02)8J*13H20,forms up to 80% of the unit with lesser amounts of erionite, (Ca,Mg,K,,Na,),,[(A10,)o(SiOz)z,J*27H20,and clinoptiloltte, Na,1(AI(>,),(SiO2),,J*24H,O.'lhe deposit is believed to have formed from the alteration ot volcanic ash within an alkaline saline lake (Sheppard et al., 1978).

~~~~EVAPOKI 1'E AUI'EKA'I'ION '1'0 PKECIPI'I'AI t: VALUABLE COMMODI'I'IES~~~~

~~~~Bioepigenetic Sulfur Deposits~~~~

A readily demonstrable example of evaporites serving as source beds for the formation of other valuable materials is the "bioepigenetic"sulfur deposits ('l'able 6.1C). Although gypsum and anhydrite deposits represent vast resources of sulfur, the cost of industrial extraction of sulfur from sulfates is economically prohibitive. However, gypsum and anhydrite either in bedded deposits or in salt dome cap rocks have provided sources of sulfate ions for the natural bacterial production of reduced sulfur which may be subsequently oxidized to elemental sulfur (Fig. 6.18). Sulfate-reducing bacteria, generally Desulfovibrio sp., utilize hydrocarbons as the energy source to metabolize aqueous sulfate and produce hydrogen sulfide and carbon dioxide as waste products (Feely and Kulp, 1957; Kuckmick et al., 1979). 'I'hese reactions result in the formation of bacterogenic calcite zones with distinctive light carbon isotope signatures that generally are the hosts for the elemental sulfur concentrations (Fig. 6.19). Frasch production of sulfur from these types of deposits typically accounts for about 30% of the worlds annual production (Harben and Bates, 1984). 'l'he salt dome cap rock deposits in the Gulf Coast have yielded over 300 million metric tons of sulfur. Boling dome is the largest producer and has accounted for over 80 million metric tons of sulfur since production began in 1928 (Fig. 6.19C). 'I'he commercial sulfur concentrations in Gulf Coast salt dome cap rocks typically occur within the calcite zone, although lesser concentrations may occur in the gypsum zone. Oil may be produced from cap rock and supradome sediments reservoirs, but the majority of production comes from reservoirs in the flanking sedimentary units. 'I'he bioepigenetic calcite zones created by these processes also may be valuable sources of limestone for industrial purposes if the dome is in a favorable mining and marketing location.

~~~~Carbon ate - h ost ed Zinc - Lead Deposits~~~~

~~~~As is discussed in more detail in a separate section of this book, bacterially~~~~

~~~~Data Center ,09126599985,[email protected], For Educational Uses *.**.I METEORIC WAT ***I*.~~~~

~~~~H2S SPRINGS~~~~

~~~~B DELAWARE BASIN \FAULT~~~~

~~~~0ANHYDRITE f GYPSUM SULFUR DOLOMITE =PETROLEUM r3SANDSTONE SALT (DIAPIRIC) MUDSTONE~~~~

Fig. 6.18. Conceptual modelsfor the origin of commercial sulfur deposits in the United Stutrs. Modified a,fter Ruckmick et al. (1979). (A) Genetic model f or the sulfur deposits iti salt dome cap rocks in the Gulf Coast region. (6) Genetic model for the strata-bound .sulfur deposits derived ,from bedded sulf ates in the Delaware Basin. produced reduced sulfur may also serve to fix metals as metallic sulfide minerals in this environment. 'I'his mechanism has been developed for the Fe-Zn-Pb-Ag sulfide deposits that occur within salt dome cap rocks in the Gulf Coast (Kyle and Price, 1986). Metals are supplied to the cap rock environment by saline formation waters of the type that have been well documented in the Gulf Coast and other sedimentary basins (Carpenter et al., 1974; Kharaka et al., 1980). Dissolution of and rock-water reactions within evaporitic sequences have contributed markedly to the major ion contents of saline formation waters, including the enrichment of valuable metals (Hanor, 1979,1987; Land, 1987). 'I'hese metalliferous formation waters have been accepted as modern analogs of the ore-forming solutions for several classic ore deposit types including the "Mississippi

~~~~Data Center ,09126599985,[email protected], For Educational Uses CARBONATE-HOSTED ZINC-LEAD DEPOSITS 513~~~~

Fig. 6.19. Geologic f eatures of bioepigenetic sulfur ores. (A) Sulfur crystals within an iron sulfide breccia from the Long Point dome calcite cap rock. (B) Sulfur with barite from a bioepigenetic calcite alteration zone derived from Permian bedded evaporites in the Delaware Basin of west Texas. (C) Storage vat of Frasch-produced sulfur at Boling dome representing about 100,000 metric tons of sulfur. (D) Zebraic calcite with laminar vugs filled with sulfur representing a bioepigenetic calcite zone derived from Permian bedded evaporites in the Delaware Basin of west Texas.

Valley type" Pb-Zn deposits and sedimentary exhalative Pb-Zn-Ag deposits (Anderson and Macqueen, 1982; Lydon, 1983). Mississippi Valley type deposits generally occur in dolomitized shelf carbonate sequences, some of which have associated evaporitic facies, e.g. Pine Point, Northwest Territories (Middle Devonian), and Gays River, Nova Scotia (Mississippian). Ore minerals were precipitated in secondary porosity in the carbonate strata from highly saline brines at temperatures ranging from 75 to 150°C (Roedder, 1976). Associated sulfate strata have been proposed as the source of reduced sulfur for sulfide precipitation (Beales and Jackson, 1966; Sasaki and Krouse, 1969; Akande and Zentilli, 1984).

~~~~Data Center ,09126599985,[email protected], For Educational Uses 514 EVAPOKI'I'IC PROCESSES AND MINEKAL KESOUKCES~~~~

Although the bacterial sulfate-reduction mechanism is viable at the low temperature end of this environment, Powell and Macqueen (1984) used organic geochemical data to support the kinetically unfavorable mechanism of thermochemical reduction of evaporite sulfate by organic matter ('l'rudinger et al., 1985). 'I'hiede and Cameron (1978) proposed that the extensive Devonian evaporite strata of the Elk Point Basin were the source of the metalliferous brines that formed the Pb-Zn orebodies of the Pine Point district, although a source from the evaporite-poor MacKenzie Basin has been favored by most workers (Beales and Jackson, 1966; Kyle, 1981; Garven, 1985). 'I'he Gays River carbonate-hosted zinc-lead deposit in Nova Scotia shows the most direct association with evaporites (Akande and Zentilli, 1984). The orebody occurs in a dolomitized Mississippian carbonate bank (basal Windsor Formation) that developed during marine transgression over a local erosional ridge on the lower Paleozoic metasedimentary basement. 'I'he basement ridge formed a barrier separating two subbasins within the Carboniferous Maritimes Basin, the Musquodoboit basin to the south and the larger Shubenacadie basin to the northwest. 'I'he Windsor Group carbonate strata host important barite and celestite deposits; the carbonate bank is overlain by marine evaporites that regionally comprise the economically important gypsum, halite, and potash deposits (Fig. 6.3). 'I'he Gays Kiver orebody consists of fault-controlled high-grade massive sphalerite and galena veins and lower grade stratiform sulfides interpreted to represent filling of primary and secondary pores in the dolomite (Akande and Zentilli, 1984). Although the most significant sulfide concentrations are restricted to the carbonate bank, local massive sulfide lenses occur at the carbonate-evaporite contact and disseminated sulfides occur within the contiguous evaporite strata. Earlier genetic studies principally of the stratiform sulfide zones at Gays Kiver suggested that mineralization was temporally and genetically related to sabkha evaporitic processes (MacEachern and Hannon, 1974). However, subsequent fluid inclusion investigations revealed that the mineralization occurred at higher temperatures (14U-215"C) than are tenable for this environment (Akande and Zentilli, 1984). Current genetic models involve the deep circulation of high temperature brines during the late Paleozoic and probable leaching of metals from the lower Paleozoic metasedimentary basement. The capping evaporitic strata provided an aquitard that controlled the local circulation of the metal-bearing fluids and probably served as the source of reduced sulfur for sulfide precipitation, possibly by thermochemical reduction of sulfate in the presence of hydrocarbons (Akande and Zentilli, 1984). Subaerial exposure with attendent karstification during the Cretaceous modified and partially obliterated the Gays Kiver orebody.

~~~~Data Center ,09126599985,[email protected], For Educational Uses SEDIMENTARY COPPER-SILVER DEPOSITS 515~~~~

Some Mississippi Valley type districts,e.g. Southeast Missouri, contain appreciable quantities of metallic sulfides in sandstone units that underlie the main carbonate host. Thus, these "sandstone-hosted'' lead deposits appear to represent metal concentrations that were precipitated in the sandstone aquifer from metalliferous formational brines enroute to the carbonate host. Similar deposits elsewhere, e.g. Laisvall, Sweden, are major metal concentrations in continental to shallow marine quartzarenites without associated carbonate-hosted ores (Bjbrlykke and Sangster, 1981). Rickard et al. (1979) provided evidence for the origin of the Laisvall sulfides from concentrated formation waters at temperatures of about 150°C. Bjbrlykke and Sangster (1981), on the other hand, proposed a mineralization model for sandstone-hosted lead deposits involving a gravity-driven hydrologic system for metal transport in dilute, low temperature meteoric waters from adjacent basement topographic highs; sulfides would be precipitated in sandstone aquifers that contain concentrations of reduced sulfur, perhaps derived from evaporitic sulfates.

~~~~Sedimentary Copper-Silver Deposits~~~~

Other types of stratabound sulfide concentrations in sedimentary terranes provide a substantial portion of the world's total production of copper, silver, cobalt, and other metals. 'I'hese include the great orebodies of the Copperbelt of Zambia-Zaire, the Belt Series of the northwestern United States, and the Kupferschiefer of eastern Europe (see Bartholome, 1974). 'Ihe metallic concentrations are hosted by reduced mudstones, siltstones, and sandstones or evaporitic carbonate units that are commonly underlain by an oxidized continental clastic sequence. The copper deposits typically are contained within a relatively narrow stratigraphic interval but have great lateral extent. The metal concentrations are discordant locally and regionally and contain textural evidence for replacement of pre-existing components such as evaporites, organic material, and early diagenetic iron sulfides (Brown, 1978; Lange et al., 1987). One mineralization mechanism that has been proposed for some of these copper-bearing sequences has been termed the sabkha model (Renfro, 1Y74). Coastal sabkhas are nourished by subsurface flow of landward migrating, low Eh-high pH sea water and by seaward-migrating, high Eh-low pH meteoric water. Commonly they are bordered on the seaward side by intertidal mudflats and lagoons that are covered by sediment-binding, blue-green algae. As the sabkha migrates basinward, formation water eventually must flow upward through buried, strongly reducing algal mats in order to reach the surface of evaporation. This continental water initially would have a low pH and high Eh and thus could mobilize and transport trace amounts of metals. As the

~~~~Data Center ,09126599985,[email protected], For Educational Uses 516 EVAPORI'TIC PROCESSES AND MINEKAL KESOUKCES~~~~

metalliferous fluid passes through the hydrogen sulfide-charged algal mat, sulfide minerals are precipitated interstitially. Kesulting metal deposits generally are conformable to the geometry of hydrogen sulfide-bearing host strata. 1 he extent of mineralization is dependent upon quantity of available reductant, duration of the sabkha process, and chemistry of the metal-bearing meteoric water (Kenfro, 1974). However, current genetic models stress that regional fluid flow systems are necessary to account for the magnitude of the metal concentrations in the major districts (e.g. Morganti, 1981; Gustafson and Williams, 1981; Eugster, 1985; Jowett, 1986; Sverjensky, 1987). The iron-bearing minerals in the associated red-beds commonly are proposed as the source of metals for the sulfide deposits. Evaporites generally have been regarded as important sources of chlorinity that contribute to metal transport, but Sverjensky (1987) has proposed that sulfate-rich evaporites are critical to the formation of these deposits by oxidizing basinal brines to facilitate transport of metals and sulfate. 'I'he associated red-beds permit the preservation of the oxidizing brines, and the sulfide- and organic-rich shales serve as precipitants (Sverjensky, 1987). Copper-silver vein mineralization is associated with anticlinal halokinetic structures cored in the Pennsylvanian Hermosa Formation within the Paradox Basin (Morrison and Parry, 1986). 'I'he copper deposits of the Lisbon Valley occur as veins, as sandstone pore fillings, and as direct replacements of plant matter in coal-bearing clastic strata of the Cretaceous Dakota Formation in close association with major faulting along the crest of the Lisbon Valley anticline. Fluid inclusion, mineral chemistry, and stable isotope evidence suggests that warm, metal-bearing, saline basinal fluids migrated upward along fault zones where mixing with shallow-circulating cooler, reduced ground water caused mineral precipitation (Morrison and Parry, 1986).

~~~~Sedimentary Exhalative Zinc-Lead-Copper-SilverDeposits~~~~

In addition to the sedimentary copper deposits, many stratiform lead-zinc-silver sulfide deposits occur within thick sequences of Proterozoic and Paleozoic siliciclastic sedimentary rocks (see Sangster, 1983, and Turner and Einaudi, 1986, for reviews). These orebodies commonly consist of stacked stratiform sulfide lenses separated by weakly mineralized strata (Fig. 6.20A). The sulfide deposits are often associated with rapid sedimentary facies changes and/or an abrupt change in thickness of the hosting sedimentary sequence. l'he deposits are generally associated with a major regional fracture system. These relationships and the local conglomerates or breccias and slumped strata which underlie many of the massive sulfide deposits suggest contemporaneous faulting was the major factor controlling the local sedimentation environment

~~~~Data Center ,09126599985,[email protected], For Educational Uses 1jX H A LA'I'I V E Z1 NC-L E A D -COPPE K-SI LVEK DEPOS 1'1's 5 1 7~~~~

Fig. 6.20. Sedimentary exhalative lead-zinc-silver-barite deposits in f ine-grained siliciclastic strata. (A) Conceptual model f or sedimentary exhalative lead-zinc-silver-- barite deposits. Sedimentary brines are discharged along fracture systems onto the contemporaneous basin f loor or into sediments below the sediment-water interface. The mineralization environment can be either within marine or continental euxenic basins. (B) Plan view of the McArthur River (H.Y.C.) deposit, Australia, showing the inferred direction of fluid flow suggested by metal distribution. Discharge of metal-bearing fluids along the normal faults into the depositional basin resulted in cross-cutting sulfide concentrations in shelf carbonates and stratif orm sulfides in the basinal siliciclastic strata. Modified after Morganti (1981). including the discharge of metalliferous fluids onto the basin floor or into permeable sediments (Figs. 6.20A,B). It has been demonstrated that some of these "sedimentary exhalative" deposits occur in intracratonic troughs which developed as the result of tensional rifting and attendant slow subsidence. Restricted euxenic conditions generally have been proposed for the thinly laminated carbonaceous sulfide-hosting mudstones (Large, 1983), but some sulfide-hosting sequences appear to have been deposited in relatively shallow water as suggested by associated evaporitic carbonate strata.

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Furthermore, the Proterozoic sulfide-hosting strata in Australia contain pseudomorphs of continental evaporite minerals (,Muir, 1979, 1983,1987;Connor et al., 1982; Neudert, 1986). 'l'hese features suggest that the metalliferous sediments were deposited in alkaline lakes in rift valleys, perhaps similar to those of the East African Kift (Kobbins, 1983; Eugster, 1985). Along those lines, the anomalous concentrations of zinc and other metals in Lake Kivu are of interest (Degens, 1973). Eugster (1985) developed the analogy of the Proterozoic sulfide-hosting sequences in Australia with the alkaline lake setting, including the recognition of magadi-type cherts. A detailed sedimentological study by Neudert (1986') indicates that the Middle Proterozoic sequence that hosts the Mount Isa lead-zinc-silver stratiform lenses and the "silica-dolomite" copper ores accumulated in a faulted bounded basin following the deposition of shallow water sediments and continental volcanics (Fig. 6.21 ). 'l'he sedi-

~~~~0STROMATOLITES HALITE CASTS &LARGE CROSS BEDS A SMALL w 4 GYPSUM PS/M a&'&'FLAT- PEBBLES 0 NODULES &TEEPEES~~~~

~~~~Pb-Zn OREBODY - 00o-0 PlSOLlTES~~~~

~~Fig. 6.21. Geologic cross-section through the Mt. Isa lead-zinc-copper-silver deposit showing the sedimentary and diagenetic structures in the ore-hosting sequence. Modified after Neudert (I986).~~

mentary sequence is interpreted to represent a saline-lake complex that generally formed in flood-plain and subaerial to subaqueous saline mudflat environments. 'I'he lead-zinc lenses are hosted by the Upper Urquhart Shale slope to basinfacies sediments that formed in a perennial saline lake (Neudert, 1986). The intrabasinal sediment compositionevolved fromcarbonate, to carbonate-sulfate, to carbonate-sulfate-Na-evaporites, with the ore-hosting sediments forming during the main sulfate-precipitating stage. Current-deposited sulfate-richsediments provided the sulfur source for formation of diagenetic pyrite that formed as the result of microbial sulfate-reduction. Primary

~~~~Data Center ,09126599985,[email protected], For Educational Uses SUPERIOR-'L'Y PE IRON FOKMATIONS 5 19~~~~

lead-zinc mineralization occurred when metal-rich chloride brines inorganically reduced the remaining sulfate; subsequent high-temperature mineralization overprinting caused extensive sulfide replacement (Neudert, 1986).

~~~~Superior-type Iron Formations~~~~

Although many geologic types of iron deposits supply raw materialsfor the worlds steel industries, the vast majority of commercial iron ore is derived from enormous iron concentrations in sedimentary sequences, the "Superior-type'' iron formations (see 'Trendall and Morris, 1983, for a recent review of these deposits). These deposits occupy large intracratonic sedimentary basins of about l.Y to 2.2 Ga. 'I'he deposits occur on virtually all continents and include the major ore-producing districts of Lake Superior region in the United States, the Labrador 'I'rough of Canada, the Hamersley Basin of Australia, the 'lransvaal Basin of South Africa, the Krivoy Rog of the U.S.S.K., and Minas Gerais and Serra dos Carajas of Brazil. Many different iron-bearing minerals occur in the units, a function of original depositional environments superimposed by mineralogic changes accompanying diagenesis and metamorphism. 'I'hese mineral assemblages generally can be subdivided into oxide, carbonate, sulfide, and silicate mineral groups which have been used to define depositional "facies" for the iron formations (James, 1954). 'These facies are commonly seen as vertical sequences in deformed terranes, but their transitions have been traced laterally in some relatively undeformed basins. The term "banded iron formation" is commonly used to refer to the typical appearance of a thinly layered rock in which the iron minerals and chert form alternating laminae. The contents of iron and SiO, typically are in the range of 20-35% and 40-50%, respectively. Much has been written about the geology and origin of the banded iron formations. James (1954) and Lepp and Goldich (1Y64) provided classic summaries of the facies relationships and their proposed depositional environments in relation to water depth and chemical conditions. Iron oxide mineral deposition was believed to take place in shallow, oxygenated waters, perhaps on a shelf environment. lron carbonates were formed in slightly deeper environments, and iron sulfides represented the most restricted (and deeper?) environments. Iron silicates have a wide stability range, but their major development appears to be near the oxide/carbonate transition. Most workers have indicated marine conditions, and evaporation has been suggested by some workers as a means of concentrating iron; Govett (1966) suggested that a lacustrine environment was more compatible with the geological characteristics of the depositional basins. Among the currently unresolved problems about these unusual

Data Center ,09126599985,[email protected], For Educational Uses 520 EVAPOKI'I'IC PKOCESSES AND MINERAL KESOUKCES sedimentary rocks is the iron source, with discussion typically involving relative contributionsfrom continental runoff, upwelling marine currents, andvolcanic exhalative contributions for particular basins. 'I'he Hamersley Basin in western Australia provides a well-documented example of an unmetamorphosed sequence of Proterozoic strata that includes several units of banded iron formation (Trendall, 1973). The total commercial iron resources of the basin is indicated to be over 36 billion metric tons. The basin is believed to have been ovoid in shape with an area of over 100,ooO km' during deposition of the iron formations. It appears to have been enclosed on all but the northwestern side, where there was at least a partial connection to the open ocean. The banded iron formations occur within the Hamersley Group and form over 40% of its total thickness of 2,500 m. 'The iron formations are particularly well banded on several scales, and some of these laminae have been correlated over the entire basin, i.e. on the scale of hundreds of kilometers (I'rendall, 1973). 'I'his amazing lateral continuity led 'Trendall (1973) to suggest that the Hamersley Group banded iron formations were early Proterozoic "ferruginous evaporites" with reference to the barred basin depositional model commonly proposed for saline evaporites. The microbanded iron oxide/chert couplets are interpreted to represent seasonal precipitation episodes essentially affecting the entire depositional basin. The banded cherts provided additional evidence to further suggest that the Hamersley and perhaps other Superior-type iron formations were deposited in the evaporative setting of an alkaline playa-lake complex (Eugster, lY69; Eugster and Chou, 1973). This playa-lake complex is envisaged to have been similar to the depositional and hydrologic setting suggested for the Green River Formation with a major chemical difference provided by a primitive atmosphere with a low oxygen content. Magadiite or sodium silicate gels are considered to be the precursors of the bedded cherts; most magadiite was converted to chert during early diagenesis, but some reacted with iron to form riebeckite, the sodium-iron amphibole, which is common in the Hamersley and some other iron formations. In additon to the rhythmic compositional banding, this model further accounts for the scarcity of clastic sediments in association with the banded iron formations. As demonstrated for playa lakes in younger settings, most of the material transport into the basin would be through springs with ephemeral precipitation and brine evolution taking place on the playa flats fringing the alkaline lake. Magadiite is suggested to have been deposited during wet periods when the lake was permanently stratified. Precipitation of iron minerals is principally effected by changes in pH and fcoz. During dry periods, silica would accumulate in the high pH brines by evaporative concentration, while iron would be precipitated as hydroxides or

~~~~Data Center ,09126599985,[email protected], For Educational Uses ASSOCIATED BARITE AND CELES’I’ITE DEPOSITS 52 1~~~~

silicates, perhaps triggered by fresh-water influxes; iron carbonates may form on the playa margins. If early Proterozoic seawater is assumed to be near saturation with respect to amorphous silica, this model could be adapted to a marine environment, most likely a barred lagoon in a climate arid enough to excludelarge perennial rivers (Eugster and Chou, 1973).

~~~~Evaporite-Associated Barite and Celestite Deposits~~~~

Formation of valuable sulfate minerals, e.g. celestite or barite, would be a likely result of the interaction of a heated Sr- or Ba-rich formation water with sedimentary calcium sulfate beds. This replacement process would be facilitated by the retrograde solubility of gypsum and anhydrite in the diagenetic temperature range (Holland and Malinin, 1979), i.e. cooling of pore waters would tend to take calcium sulfates into solution while precipitating celestite or barite. Barite replacement of early diagenetic calcium sulfate nodules in sabkha sediments was proposed by Mossman and Brown (1986) for the Carboniferous Windsor Group sediments in the Walton-Cheverie area of Nova Scotia, although they acknowledged more complex processes for major sulfide-baritedeposits in the same stratigraphic unit (Fig. 6.3). The Windsor Group also hosts the Loch Lomond celestite deposits that have been interpreted to result from sabkha brine processes (Forgeron, 1984). Major concentrations of celestite, albeit as yet noncommercial, also occur within karst breccias in Upper Permian evaporitic limestones at Karstryggen, central East Greenland (Scholle et al., 1990). The mineralization model proposed by Scholle et al. evokes the leaching of Sr from underlying arkosic redbeds and the precipitation of celestite in the karsted limestones with the sulfate source provided by the associated gypsum beds. The formation of barite and celestite concentrations in salt dome calcite cap rocks is probably related to alteration of anhydrite, although sulfur isotope data indicates that it is not a direct replacement process (Kyle and Price, 1986; Posey et al., 1987; Kyle and Posey, this volume). Similar barite concentrations are associated with the bioepigenetic calcite and sulfur deposits that formed as the result of alteration of Permian evaporites in the Delaware Basin. Several barite deposits occur within Permian carbonate-evaporite strata of the Delaware Basin in west Texas and contiguous New Mexico (Kyle, in press). The most significant barite deposit is at Seven Heart Gap in the southern part of the basin along the northern flank of the Apache Mountains, where the restricted basin deposits of the Ochoan Castile and Rustler Formations onlap the Capitan reef carbonates that form the Apache Mountains core. The barite deposits occur within the basal Castile Formation overlying the transition zone from the Capitan off-reef talus to

Data Center ,09126599985,[email protected], For Educational Uses 522 EVAPOKI‘I’IC PROCESSES AND MINERAL KESOUKCES basinal deposits of the Bell Canyon Formation. The dominant barite-hosting structures at Seven Heart Gap are brecciated and slumped zones that have been interpreted to be the result of evaporite dissolution (McAnulty, 1980). Barite typically occurs as intergrown radiating crystals that precipitated within open space in dilatent zones and slump breccias in the basal Castile Formation. The Delaware Basin also contains elemental sulfur deposits in altered carbonate and evaporite strata of Ochoan age, including the Culberson deposit, one of the largest commercial sulfur concentrations in the world which is believed to contain in excess of 60 million tons of recoverable sulfur (Kuckmick et al., 1979). The ore zone consists of sulfur within vuggy bioepigenetic limestone and is overlain by a clay-rich unit that probably acted as an important seal for the sulfur-generating system. ‘I’he sulfur-bearing zone commonly contains barite and celestite including local concentrations up to several feet thick (Fig. 6.19B). At least 8 smaller sulfur deposits occur at or near the base of the Castile Formation in the region. Requirements for the formation of the sulfur deposits of the Delaware Basin include the presence of petroleum-bearing strata underlying a sulfate-rich environment in which structural and hydrologic conditions permit the interaction of water, hydrocarbons, and bacteria (Kuckmick et al., 1979; Fig. 6.18B). ’I’he age of mineralization is a particularly significant aspect concerning genetic relationships between the Seven Heart Gap barite deposits and the elemental sulfur deposits of the Delaware Basin. Initiation of the sulfur-depositing process clearly postdates the development of ore-controlling structures and the establishment of the hydrologic framework for the interaction of bacteria, hydrocarbons, and meteoric water. Sulfur deposition appears to be currently active in the Delaware Basin, suggesting that the commercial sulfur and associated barite concentrations formed relatively recently. The Seven Heart Gap barite concentrations are clearly epigenetic, but mineralization was relatively early in the diagenetic history, perhaps beginning in late Ochoan. Stable isotope data also suggest major differences between the barite deposit types (Kyle, in press). All of the barites are variably enriched over sea water sulfate of Ochoan age as documented for evaporites of the Permian Basin and elsewhere (Claypool and others, 1980). ‘These data further indicate that the barite concentrations are not the direct result of replacement of Permian evaporites, i.e., enrichment of aqueous sulfate in ”S must have taken place prior to the precipitation of barite. Barite in the elemental sulfur deposits has extremely heavy sulfur isotope values up to 73%0 (CUT), suggesting the bacterial reduction mechanism as has been proposed for barite in the Gulf Coast salt dome cap rock sulfur and metal sulfide deposits (Kyle and Price, 1986). The Seven Heart Gap barite deposits have slightly to moderately heavy sulfur isotope compositions (Kyle, in press). The mixing of heavier sea water sulfate of early Mesozoic age, perhaps

~~~~Data Center ,09126599985,[email protected], For Educational Uses EVAPOKITES AND IGNEOUS-HOS'I'ED OKE DEPOSI'I'S 523~~~~

somewhat enriched in 34S by bacterial processes, with aqueous sulfate derived from Ochoan evaporites could account for the range of sulfur isotope values of the Seven Heart Gap barites. The barite concentrations at Seven Heart Gap and at the C'ulberson sulfur deposit share many features that result from common geologic terrane, including hosting porosity that developed as the result of solution and/or alteration of evaporites. However, there is sufficient geologic and geochemical evidence to indicate that they are not directly related genetically (Kyle, in press).

~~~~b'vaporites and Igneous-hosted Ore Deposits~~~~

Evaporites also may be of considerable importance in the formation of ore deposits in non-sedimentary environments. An excellent example is provided by the extensive magmatic nickel-copper-platinoid sulfide ores of the Noril'sk-'l'alnakh districts at the western edge of the Siberian platform. 'I'he region is underlain by a thick sedimentary sequence of Paleozoic age that includes major Devonian evaporites. Extensional tectonics resulted in the eruption of a large volume of Late Permian and 'I'riassic flood basalts that covered the emergent Paleozoic sedimentary sequence. Sill-liketholeiitic intrusions were emplaced contemporaneously with, and are feeders to, the extrusive volcanics. 'I'he nickel-copper-platinoid sulfide ores form relatively persistent zones in the lower parts of the intrusions. Sulfur isotopic data show that the sulfides are anomalously heavy, a feature that suggests that sulfur was supplied from sedimentary sulfates rather than directly-- by the mantle-derived magma. Geological and geochemical evidence indicates that the basaltic magma assimilated sulfur from the Devonian evaporites as it rose through the sedimentary cover along the rift faults. 'The evaporitic sulfate was reduced to sulfide which reacted with iron from the magma. 'I'he resulting immiscible iron sulfide acted as collectors for the economic metals now concentrated in the basal part of the intrusions (Naldrett and Macdonald, lY80). Considering the association of continental rifts, mafic magmatism, and evaporites, it might be anticipated that this type of deposit would be more common than is apparent from known examples. 'I'he Viscaria copper deposit in the Kiruna district of Sweden also has been interpreted to represent mineralization in a playa lake environment within a mafic volcanic terrane of Proterozoic age (Godin, 1986; Godin and Lager, 1986). 'I'he deposit consists of four mineralized horizons within a 700-m section in a dominantly tholeiitic basalt sequence containing units of both subaqueous and subaerial character. Lake level changes, perhaps related to climatic variations, are proposed to be reflected in the cyclic lithologic sequence of graphitic units, limestones, and chert-albite units. 'I'he highest

~~~~Data Center ,09126599985,[email protected], For Educational Uses 524 EVAPOKI'I'IC PKOCESSES AND MINEKAL KESOUKCES~~~~

concentrations of sulfide minerals occur within the graphitic schist and limestone units. 'I'he graphitic units probably formed in deeper anoxic parts of a fault-bounded alkaline lake, whereas the limestones represent shallow nearshore deposition which may have been resedimented in the lake environment. Evaporite mineral pseudomorphs suggest the chert-albite units originated in a saline lake environment with precipitation of sodium silicate precursors which are subsequently replaced by magadi-type chert. Areas of scapolite and albite alteration may represent authigenic reactions in volcanic sediments around the alkaline lake (Godin and Lager, 1986). 'I'he sulfide mineral concentrations appear to have replaced the lacustrine sediments relatively early in the diagenetic history of the strata.

~~~~EVAPOKI'I'E SOLU'I'ION '1'0CKEA'I'E ORE-HOS'IING POKOSI'I'Y~~~~

'I'he selective removal of more soluble materials within a sedimentary section, commonly accompanied by collapse brecciation of associated strata, can be effective in the creation of permeability trends that govern local movement of mineralizing solutions and may provide ore-hosting porosity. bvaporites have the lowest average residence times of any common geologic materials because of their high solubilities in aqueous solutions, and their former presence in sedimentary basins may be difficult to document in many cases where solution has led to total removal. It is difficult to provide examples in which simple dissolution of evaporites has been instrumental in the development of porosity zones and permeability trends that are important in localization of mineral resources ('I'able 6.1 D). 'I'hat is, the diagenetic conditions that result in evaporite dissolution commonly are associated with geochemical environments that directly involve the dissolved components in mineral precipitation. It has been proposed that the ore zones for several of the Mississippi Valley type districts result from the selective removal of evaporitic strata and the collapse of associated dolomites to form ore-hosting breccias (Jackson and Beales, 1967; Beales and Hardy, 1980; Beales, written commun., 1985). 'I'he presence of baroque (saddle) dolomite filling pores, cementing breccia fragments, or forming breccia-moldic textures commonly has been cited as evidence for the former presence of evaporites. Indeed, the precipitation of saddle dolomite may favor sulfate-rich solutions (Radke and Mathis, 15)80),but its presence does not necessarily signify local evaporites. The ore-hosting collapse breccias for several of the classic Mississippi Valley type districts including those that occur in evaporitic settings, e.g. Pine Point, have been shown to be the effect of karstification and meteoric diagenesis on carbonate strata in the upper part of a subaerially exposed carbonate platform (Kyle, 1983; Rhodes et al., 1984). However,

~~~~Data Center ,09126599985,[email protected], For Educational Uses AS A HYDROLOGIC AGENT IN ORE FORMATION 525~~~~

evaporite removal and attendent collapse of associated carbonate strata has been suggested to be important to the development of ore-hosting breccias in other districts, as well as in the supply of reduced sulfur for precipitation of metals from formational brines (Beales and Hardy, 1980). Major stratabound celestite and barite deposits occur within Lower Cretaceous (Albian) carbonate strata in northern Mexico. Fluid inclusion and isotopic data indicate that these deposits originated by mixing of low temperature formational fluids and that the sulfate was supplied from gypsum deposits within the sequence (Kesler and Jones, 1981). Geologically similar deposits occur in central Texas and consists of barite and celestite of several textural types within secondary porosity in the Albian Edwards Limestone. The mineralization interval is a widespread brecciated zone in central Texas that formed as a result of dissolution of evaporitic beds of the Kirschberg Member and collapse of associated carbonate strata. A commercial gypsum zone up to 12 m thick is locally present. Some of these barites have sulfur isotope values reflecting Albian seawater values, whereas others are substantially heavier indicating that simple replacement of local gypsum did not occur (Kyle, in press). This "S enrichment of the residual aqueous sulfate relative to Albian evaporite values suggests the involvement of sulfate-reducing bacteria in the shallow meteoric environment of evaporite dissolution.

~~~~EVAPORATION AS A HYDROLOGIC AGENT IN ORE FORMAlION~~~~

Evaporation has been proposed as a hydrologic agent in a variety of geologic terranes to drive fluid circulation that resulted in the concentration of valuable commodities (Table 6.1E). Many of the models that have been postulated have not had rigorous geochemical and hydrologic evaluations that would allow a realistic appraisal of the importance of evaporation in governing fluid flow. Following models developed for dolomitization in evaporitic settings (e.g. Deffeyes et al., 1%5), Lange and Murray (1977) proposed that Mississippi Valley type deposits resulted from the deep "reflux" circulation of hypersaline brines forming in evaporite pans. However, there are a number of geologic, hydrologic, and geochemical arguments against this mechanism for mineralization in most classic districts. Furthermore, there is considerable debate as to the viability of the reflux mechanism for regional dolomitization in evaporitic settings (see Land, 1985; Machel and Mountjoy, 1986; and Hardie, 1987, for current reviews of the "dolomite problem"). Certainly the mechanism for formation of regional dolomites is a major concern for Mississippi Valley type mineralization, as these typically are the host rocks. That is, the presence of

~~~~Data Center ,09126599985,[email protected], For Educational Uses 526 EVAPOKI'TIC PKOCESSES AND MINEKAL KESOUKCES~~~~

permeable dolomites in the evaporitic carbonate shelf environment has governed regional fluid flow systems and has provided ore-hosting porosity. As discussed earlier, the sabkha model for the formation of sedimentary copper deposits is intimately tied to evaporation as the driving mechanism for mineralizingfluid flow (Kenfro, 1974). Although this mechanism could operate on a local scale, there is considerable doubt that it could function on the scale required to form many of the great orebodies of this type (Eugster, 1985). l'he importance of evaporitive pumping to solute enrichment in the continental playa environment has been discussed in earlier sections. One type of metal concentration in which evaporitive discharge zones appear to be critical are the calcrete uranium-vanadium deposits of western Australia (Mann and Deutscher, 1978). Weathering of a granitic basement rock is the source of the metals which are carried as oxidized metal complexes in ground water. Partial evaporation of the slow-moving ground water mass concentrates the uranyl ions by decomplexing the soluble uranyl carbonate complexes and promotes the formation of carnotite-rich calcrete zones (Mann and Deutscher, 1978).

~~~~CONCLUSIONS~~~~

Some examples of the diverse relationships that evaporites and evaporitic fluids play in the concentration of valuable commodities have been presented. Some "ores" are saline formation waters of evaporative parentage that are unusually enriched in valuable ions. Evaporitic brine-associated alteration may be important in the formation of commercial concentrations of valuable industrial minerals such as zeolites. In addition to being sources of important industrial rocks and minerals, evaporite deposits also serve as the feedstock for geologic processes that either formed valuable materials or provided a favorable environment for the chemical trapping of externally supplied elements. Metal-bearing saline formation waters have been widely accepted as modern analogs of ore-forming fluids for a variety of classic ore deposit types. 'I'he bacterial or thermochemical reduction of evaporite-sourced aqueous sulfate over the wide range of temperatures appropriate for sedimentary and diagenetic environments may be of critical importance for the precipitation of metal sulfide minerals for a number of ore deposit types. Furthermore, removal of soluble evaporites by intrastratal dissolution may create permeability trends that can govern ore-forming fluid movement or provide ore-hosting porosity. Clearly much more could be written on any topic presented in this review, and many more examples of these processes and products could have been included. Many

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of the processes that have been subdivided for the purposes of organization are related in actual geologic environments. However, it is hoped that this brief discussion will serve to stimulate evaluation of the importance of evaporitic processes in a wider range of geologic environments than is commonly considered. ’The recognition of the former presence of evaporitic brines or minerals in the rock record requires thorough evaluation of the geologic, geochemical, and hydrologic aspects of the local and regional environment.

~~~~ACKNOWLEDGM EN13~~~~

1 am grateful to numerous individuals who contributed photographs or samples to illustrate this review, as credited in the figure captions. Other field and mine photographs are by the author; the excellent hand specimen photographs were taken by David Stephens. 1 am grateful for the skilled drafting support of Jeff Horowitz. Manuscript preparation was supported by the Own Coates Fund of the Geology Foundation of ‘I’he University of ‘lexas at Austin.

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328 pp. Naldrctt. A.K. and Macdonald. A.J.. 19%).Tectonic settings of Ni-Cu sulphide ores: their importance in genesis and exploration. In: D. W. Strangway (Editor). The Continental Crusl and Its Mineral Deposits. Geol. Assoc. Can.. Sp.Pap.. 20: 031 -657. Neudert. M. 1986. A Depositional Mtdel for the Upper Mount Isa Group and Implications for Orc Formation. Unpubl. Ph.D. Thesis. Research School of Earth Sciences. Australian National Univ.. Canherra 1983. Second Edition. Publ. by SEDCON. Brishane. 19%. 324 pp. Parker. R. B. and Mannion. L. E.. 1971. (Editors). Trona Issue. Univ. Wyoming, Contrib. Geol.. 10: 72 pp. Posey, H. H.. Kyle. J. R.. .Jackson. T. J.. Hurst. S. D. and Price. P. E.. 1987. Multiple fluid components of salt diapirs and wlt dome cap rtrks. Gulf Coast. U.S.A.. Applied Geochem.. 2: 523-53. Powell. T.G.and Macqumn. R.W.. 19M. Precipitation of suifide oresand organic matter: sulfate reactions at Pine Point. Canada. Science. 2M 63-66, Radke. B.H. and Mathis. R.L.. 19XO. On the formation and Occurrence of saddle dolomite. J. Sed. Petrol.. 50: 1140-116X. Raulslon. B.V. and Waugh. D.C.E.? 1983. Stratigraphic comparison of the Mississippian potash deposits in New Bn.nswick. Canada. Sixth International Symposium on Salt. Salt Institute. 1: 115-129. Renfro, A.R.. 1974. Genesis (fevaporite-asstriated stratform metalliferous deposits: a sahhka process. Eon. Geol.. 09: 345: Discussion: Eon. Geol.. 70: 407409.

Rhodes. D.. Lantos, E.A.. Lantos. JA.. Webb. R.J. and Owens. D.C.. 10M. Pine Point orebodies and their rclationship to th ~ stratgraphy. structure. dolornitbation. and karstification of the Middle Devonian barrier complex. Eon. Geol.. 74: 991-1055. Rickard. D.T.?Willden. M.Y.. Maringer. N.E. and Donnelly. J.H.. 1970,Studies on the genesk of the Laiwall sandstone lead-tinc deposits. Sweden. Eon. Geol.. 74 1255-1285. Robhins. E.I.. 1983. Accumulation of fossil fuels and metallic minerals in active and ancient rift lakes. Tcctonophysics. 'W: 0333-658. Roedder. E.. 1976. Fluid inclusions evidence on the genesis of ores in sedimentary and volcanic rtrks. In: K. H. Wolf (Editor). Handbook tfStrata-bound and Stratiform Ore Deposits. Amsterdam. Elsevier. 2: 67-110. Rouvier. H.. Perthuisot. V. and Mansouri. A,. 1985. Pb-Zn deposits and salt-hearingdiapirs in southern Europe and north Africa. Eon. Geol.. 80: C3XAX7. Ruckmick. J.C.. Wimkrly. B.H. and Edwards. A.F.. 1979. Classification and genesis of biqenic sulfur deposits. Eon. Gcol.. 74 469474. Sangsier. D.F.. 1983. (Editor). Short Course in Sedimeni-hosted Stratiform Lead-zinc Deposits. Min. Asstr. Can.. Handhook No. 9, Toronto. 3(W pp. Sasaki, A. and Krouse. H.R.. 1y69. Sulfur isotopes and the Pine Point Iead-Tinc minerahtion. Eon. Geol.. 64:718-730. Scholle. P. A,. Stemmerik. L. and Harpbth, 0.. 1993. Origin of major karst-associated celestitc mineralkition in Karstrygen. central East Greedand. J. Sed. Petrol., fik 397410. Schreikr, B. C., 19%. (Editor). Evaporites and Hydrocarbons. Columbia Univ. Press, New York. 475 pp. Shanks, W. C., 111 and Bischoff, I. L., 1977. Ore transport and deposition in the Red Sea geothermal system: a sctrhcniical model. Geochim. Cosmochim. Aeta, 41: 1507-1519. Shanks. W. C.. 111: 1983. Economic and exploration significance d Red Sea metalIiferous brine deposits. In: W. C'. Shank\. 111 (Editor). Unconventional Mineral Resources. Cameron Volume. Am. Inst. Min. Metall. Pet, Engn.. New York. pp. 157-171. Sheppard, R. A,, Gude. A. J. and Elson, G. M.. 1Y78. Btnvie mlite deposit. Cochise and Graham Countie\. Arimna. In: L. R. Sand and F. A. Mumpton (Editors). Natural Zeolites-- Occurrence. Properties. Use. Pergamon. pp. 319-328. Smith. GI.. 1%. Seology of Searles Lake--aguide to prospectingfor buried continental salines. In: R. L. Rau (Editor). Sccond Symposium of Salt, Northern Ohio Geol. Soc., pp. 167-180. Smith. GI., 1979. Subsudace Stratigraphy and Geochemistry of Late Quaternary Evaporites. Searles Lake. California. CJ. 4. Geol. Sum., Prof. Pap., 1043 I30 pp. Sonnenfeld, P., 1984. Brines and Evaporites. Academic Press, Orlando, 613 pp. Sonnenfeld, P. and Perthuisot, J. P., 1989. Brines and Evaporites. Am. Geophy. Union, Short Course in Geology, 3: 126 pp. Spencer, RJ., Baedecker, M. J., Eugster. H.P.. Forester, R. M., Goldhaber. M. B., Jones, B.F., Kelts, K., MeKen7ie. J., Madsen,

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D. B., Rettig, S. L., Rubin, M. and Bowser. C. J., 1984. Great Salt Lake and precursors, Utah: the last W,ooO years. Contrib. Mineral. Petrol., 86: 321-334. Spencer. RJ., Eugster, H.P. and Jones, B.F., 1985. Geochemistry of Great Salt Lake, Utah. 11. Pleistocene-Holocene evolution. Geochim. Cosmochm. Acta, 49 739-747. Surdam, R.C. and Wofiauer, CA., 1975. Green River Formation, Wyoming: a playa-lake complex. Geol. Str. Am. Bull., W 335-345. Sverjensky,DA., 1987. The role d migrating oil-field brines in the formation t# sediment-hosted Cu-rich deposits. Econ. Geol., K? 1130-1141. Thidc, D.S. and Cameron, E.N., “8. Concentration of hawmetals in the Elk Point evaporite sequence, Saskatchewan. Eon. Geol., 73 401415. Trendall, A.F., 1973. Iron-formations d the Hamcrsley Group of Western Australia: type examples of valved Precambrian evaporites. In: Genesis d Precambrian Iron and Manganese Deposits. UNESCO, Pans, pp. 257-270. Trendall, A.F. and Moms, R.C., 1983. (Editors), Iron-Formation: Factsand Problems. (Developmentsin Precambrian Geology, 6), Elsevier, 558 pp. Trudinger, PA.. Clamhers, LA. and Smith, J.W., 1985. Ltw-temperature sulphate reduction: bidogicalversus abiological. Can. J. Earth Sci., 2’2 1010-1918. Turner, RJ.W. and Emaudi, M.T., 1986. (Editors), The Genesis of Stratifom Sediment-Hosted Lead and Zinc Deposits. Cod. Proc., Dept. Geol. Sci., Stanford Univ. Publ., 20: 225 pp. Warren. J. K., 1989. Evaporite Sedimentology. Prentice-Hall, Englewd Clffs, 285 pp. Webb. T.C.. 19%. New Brunswick potash deposits. In: G. R. (;diet and W. Martin (Editors), Industrial Minerals of Canada. Can. Instn. Min. Metall.. Spec. Vol. No. 2(? 4147. Weisman. W. 1. and Mcllveen, S.. Jr., 19x3. Sodium sulfate deposits. In: S. J. Ldond (Editor). Industrial Minerals and Rocks. Ffth Edition. Soc. Min. Engnrs., New York, pp. 1207-1223. White, D.E., 1081. Active geothermal systems and hydrothermal ore deposits. In: B. J. Skinner (Editor). Economic Geology, 75th Anniversary Volume. pp. 3Y2-423. Zharkov, MA., 1984. Paleozoic Salt-bring Formations d the World. Springer. Berlin, 427 pp.

~~~~Data Center ,09126599985,[email protected], For Educational Uses This Page Intentionally Left Blank~~~~

~~~~Data Center ,09126599985,[email protected], For Educational Uses INDEX~~~~

a-carotene 366 calcareous 82 A-1 Carbonate 37 Charophyte 211 A-2 Carbonate 390 green 352, 364 Abu Dhabi 37 Algal sabkha 54 bioherm 126 Abu Sir Limestone 96 biscuits 125, 126 Acid flats 100 acetic 355 lamination 84 fatty 380, 397 mat 78, 81, 114 ferricyanide 17 peat 83 fulvic 366 tufa 242 humic 17, 366, 370, 402 tufa crust 212 hydrochloric 486 Alluvial fan 37, 95, 254, 282, hydrolyzable 366 286 organic 359, 386 Alteration sulfuric 486 postdepositional 2 volatile fatty 355 syndepositional 2 Acropora 82 Alviso saltern Africa brine density 355 Afar region 36 dissolved organic Congo Basin 415 compounds 398 Danakil Depression 36, organic chemical 355 43, 44, 45 Andean Altiplano 75 East African Rift 498, 499 Anhydrite Karoo Basin 260 chicken-wire 84 Lake Asal 36 economic uses 483 Lake Kivu 203 nodular 85, 89 Lake Magadi 46, 109, 201, Anhydrite oils 388 498, 499 Apa n o t h e c e h a I oph y t i c a 353 Maghreb region 415, Arabian Gulf sabkha model 419 144 Mediterranean coast 37, 96 Argentina Qattara Depression 292 Antuco 498 Algae Arkansas blue green 351, 515 Joachim Dolomite 59 brown 80, 82 Asphalt 349

~~~~Data Center ,09126599985,[email protected], For Educational Uses 536 INDEX~~~~

Atacama Desert 505 Barred-basin model 70 Australia Barrier Coorong lakes 83, 118-121, Presqu’ile reefs 5 127 tectonic 39 Cygnet Carbonate 108 topographic 5, 41 Hutt Lagoon 36, 47, 114 islands 82 Ibis Gypsite 108 Basement Lake MacLeod 36, 46, Precambrian 416 103, 107, 114 Basin Lake Prungle 207 Amazon 40 Leeman Lagoon 36, 47, Appalachian 478 114 Argana 260 New Lake 48 Canning 176 North Pole 71 Carnarvon 47 Officer Basin 387 closed 189, 191 Shark Bay 54 Congo 415 Southern 103 Cumberland 491 Western coast 103 Delaware 49 p-carotane 384, 391 East Texas 426, 441 Bacteria Elk Point 40, 482, 488 aerobic 362, 384 Fundy 302 anaerobic 360 Gulf Coast 153, 417 fermentative 355 Hamersley 520 halophilic 352, 353, intermontane 191, 286 366, 385 Jianghan 386 in halite 31 Khorat 383 methanogenic 355, 360, Lake Eyre 191 363, 380 MacLeod 36, 46, 103, photosynthetic 354 107, 109, 114 sulfate reducing 363 Maritimes 477, 478 bacterial reduction 359 Michigan 101, 372, 390, Bahamas 478 Great Island 36 Midland 49 Baja California 36, 57 Moncton 491 Banded iron 519, 520 Newark 299, 301 Bar deposits 241 Officer 387 Barite 413, 514, 521 Orca 360

~~~~Data Center ,09126599985,[email protected], For Educational Uses INDEX 537~~~~

Palo Duro 49, 135, 138 Salar Grande 218, 280 Paradox 415, 490, 516 Bonaire 12, 22, 36, 130 Permian 49, 50, 129 Bonneville Perth 114 Lake 493 Piceance 382 Salt Flat 230, 493 Powder River 142, 144 Borates 493 Qaidam 200, 258 borax 494 Serigipe 415 colemanite 494 Solikamsk 379 economic uses 493 Tajo 200, 298, 300 ulexite 494 Tyro 360 Borax 478 Uinta 382, 384 Boron 493 Western Canada 478 Bouma sequences 210 Williston 90, 138, 140, Boxwork limestone 117 153 Brazil Wucheng 199 Serigipe Basin 415 Zechstein 415, 486 Breccia Basin and Range 194 solution collapse 140 Bassanite 84, 173 sulfide 513 Biogenic 415, 433 Brecciat ion Biomarker salt 42 1 compounds 383 Brine evolution cyclic 373, 403 flow diagram 201 data 386 Brine mixing 33 geochemistry 387 Brine reflux Biomass 349 dolomite 100, 131 marine settings 352 dolomitization 129 production 352, 353 Brine reflux model Bischoffite 425 hydrologic model 130 Bitumen extracts 363 Brine shrimp 105, 363 geochemistry 374 Artemia 353 Boling dome Brine/air interface 19 contour map 445 precipitation 10, 13 sulfur production 5 11 Bristol Dry Lake Bolivia vertical succession 48 Pastos Grandes 223 Bromine Salar de Coipasa 261 economic uses 503

~~~~Data Center ,09126599985,[email protected], For Educational Uses 538 INDEX~~~~

Brown Niagaran 372 435 Buckner Formation formation 431 vertical succession 55 hydrocarbons 44 1 Burial depth 34 marine false cap rock 433 Burial diagenesis 150, 497, Ca pi llary 506 action 22 Cabin Creek field 138, 141 brines 52 Calcite crust 107, 225, 126 bioepigenetic 401, 513 evaporation 92 zebraic 513 pressure 69, 88 Calcite cap rock zone 22, 88, 91, 116, mineralization 438 164 Calcium sulfate hemihydrate Carbon 483 even carbon 381, 384 Caliche 258 odd carbon 366 cement 256 organic 359, 360 crust 256 Carboniferous 46 nodules 256, 274 Carnallite 167, 176, 488 California Catagenesis 398 Badwater basin 27 Celestite 514, 521 Bristol Dry Lake 21, 33 stratabound 525 Death Valley 27, 29 Cement Devils Golf Course 28 clear halite 26, 32 Cambrian 16 syntaxial 13, 14 Canada syntaxial overgrowths Great Plains 503 13, 32 Maritimes region 415 Cenozoic 416, 510 Canary Islands 131 Central Basin Platform 49, Cane Creek potash mine 392 156 Cap rock Cerithid gastropods 83 "true" calcite 433, 438 Cerithid grainstone 87 anhydrite 444 Chert 250, 500 basal anhydrite zone Chile 43 1 nitrate deposits 505 calcite 444 Salar de Atacama 504 conditions for 436 Salar Grande 219 fluid volumes required Salar San Martin 270

~~~~Data Center ,09126599985,[email protected], For Educational Uses INDEX 539~~~~

salars 31 Cornets 16 salt pans 27 Cracks China mud 121 Hunan Province 199 synaeresis 212 Jianghan Basin 386 thermal 30 Qaidam Basin 200 Creede district 508 Qianjiang Depression hydrothermal system 386 509 Wucheng Basin 199 Cretaceous 133, 137 Chlorophyll 366 Early 418 Chlorophyll-a 35 1 Crust Chott 189 efflorescent 226, 499, Clastic 500 halite 49 evaporite 205 halite particles 22 puffy ground 234 particles 10 Crystal Clear Fork displacive 22, 167 vertical succession 53 fluid inclusions 9, 15 Climate 118 growth bands 9 Columbia morphology diagram 18 San Andres 131 replacive 167, 208 Contraction sediment inclusions 9, thermal cracking 30 208 Coorong lakes 83, 118 vertically aligned 137 cross section 121 Crystallization dolomite 122 brine-air interface 7, facies schematic 127 10, 13, 19, 204 Halite Lake 119 displacive-incorpora tive Milne Lake 119 7, 167 North Stromatolite 119 dissolution cavities 8 Pellet Lake 119 efflorescent crust 8, regional setting 119 205, 226, 499, 500 stratigraphy 150 fibrous cement 8 vertical section 12 Crystals Copper 515 acicular 31 Copper-silver bottom-nucleated 13, vein mineralization 5 16 204

~~~~Data Center ,09126599985,[email protected], For Educational Uses 5 40 INDEX~~~~

chevron 15, 26, 51 birdfoot 235, 236 cornet 15, 26 braid 236 cuboids 15 description 235 cumulus 10, 204 fan 236, 244 dendritic 21 " G i 1 be r t -type " 235, 2 36 displacive 19, 167 sheet 235, 238 hopper 20 Shatt el Arab 78 incorporative 19 Desert skeletal 21 Atacama 505 Cuboid growth 13 intracontinental 191 Cumulates 49 Namib 277 layered 14 varnish 256, 273 Cyanobacteria 83, 351, 364, Desiccation stage 216 3 72 DesuIf ovibrio 51 1 Aphan o tti ece Detrital deposits 207 halophyta 352 Devonian 40, 478 mats 105, 232, 359 Middle 489 mucilage 354 Upper 366 nitrogen fixing 362 Diagenesis Cygnet Carbonate 108 burial 150, 167, 173, Cyphastrea 82 178, 497, 506 Danakil fluid-mineral 173, 415 Alps 44 halite 10, 33 Depression 36, 43, 44, Diapir 45, 219 cross section 442 Dead Sea see domes bromine 503 subaerial, photo 423 chemistry 365 Diapiric South Basin 208 structures 97, 420 Death Valley grabon 498 Diapirism Deep Sea Drilling Program buoyancy 420 (DSDP) 71 contraction 420 Deep-basin model 71 differential loading 420 Delaware Basin extension 420 Permian cross section fluid infiltration 424 490 fluid migration 427 Delta fluid softening 423

~~~~Data Center ,09126599985,[email protected], For Educational Uses INDEX 541~~~~

gravity spreading 420 Boling 434, 437, 445, halokinesis 415 51 1 mechanics 420 Hockley 437, 449 thermal convection 420 Long Poinit 5 13 timing 426 Moss Bluff 437 Diatoms 362, 364 Palangana 443 Chrysophyta 352 Winnfield 432, 434, 485 marine 352 Drain age Displacive growth 167 surface 190 Dissolution Dry mudflat cavities 58 description 228 events 25 Dunaliella halite 25 salina 352, 353, 364 karst formation 25 viridis 352, 353, 364, Dissolved organic carbon 3 65 (DOC) 402 Dunefield, drowned 196 Dissolved organic content Dunes 359 barchan 2151 Dissolved organic matter clay 264 (DOM) 354 coastal 103 Dolomite drowned 196 basal 120, 123 eolian 42, 47, 56, 87, 264 brine reflux 141 gypsum 26#4 detrital 93 parna 264, 290 penecontemporaneous White Sands 264 90 zibar 264 primary 90, 118, 122 East African Rift sabkha 91 cross section 499 salina 118 economic resources 498 secondary 90, 129 East Texas Basin type A 122 diagram 428 type B 122 Efflorescent Dolomite formation crust 226, 108 schematic 91 halite 56 DOM 354 EgY Pt Domes Mediterranean coast Belle Island 434 37, 96

~~~~Data Center ,09126599985,[email protected], For Educational Uses 5 42 INDEX~~~~

Qattara Depression 197 292 Evaporite oils 389 vertical succession 55 Evaporite volumes Wadi Natrun 362, 363 through time 367 El Hemiet Island 359 Evaporite-petroleum Elk Point Basin 5, 40 envelope 368 Enterolithic folding 419 Evaporitic carbonates Eocene 501 biomass 368 Eolian dunes 42, 47, 56, 87, Evaporitic pond stage 116 264 Even carbon acids 380 Eolian dunefield 261, 288, Even carbon n-alkanes 381, 292 384 Ephemeral Extractable organic matter halite 42 (EOM) 368 halite crust 42 Facies salinas 35 alluvial fan 37, 95, 254, 282, salt pans 41 286 stream 38, 257, 288 bar deposits 241 Epilimnion range 355 bouma sequences 210 Epit hermal deltaic 235, 236, 244 mineral deposits 507 dry mudflat 228 mineralization 496 dunefield 196 vein deposits 508, 509 eolian dunefield 261, 288, Erg 87, 94, 144 292 Ethane 360 fan delta 38, 54 Ethiopia fluvial 253, 259 Danakil Depression interdune 38, 54, 56, 191 219 intertidal 83 Ethiopian Plateau 44 lacustrine 361 Ethiopian rift 44 lakes 124, 383 Eutrophic lakes 383 marginal marine sabkha 52, Evaporative 54, 87 concentration 33 marginal marine salina 37, dolomite 121 38, 100 drawdown 103, 108, middle supratidal 84 110, 131, 151 mudflat 39, 40, 50, 138 Evaporative pumping 59, 92, perennial salina 35, 41, 202,

~~~~Data Center ,09126599985,[email protected], For Educational Uses INDEX 5 43~~~~

203 Hessen 390 perennial lake 292 Yates 144, 146 perennial saline lake 286 Flood recharge 92 perennial stream 258, 286, Flooding stage :213 292 Fluid peritidal zone 91 ore-forming 477 platform evaporites 134, 135 saline 478 playa 74, 189, 508 Fluid inclusions 51 playa lake 189 hydrocarbon 377, 456 sabkha 74, 75, 76, 81, 93, 94, Fluid infiltration 189 diapirism 424 salar 189 Fluid migration saljna 100, 118, 165, 189 around diapirs 427 saline mudflat 220 Fluid softening saline pan 213, 214, 282, 288, diapirism 423 293 Fluid-mineral saltern 136, 150, 353 diagenesis 4 15 sand sea 94, 191 Fluids subtidal 79 basinal 414 supratidal 83 ore-forming 505 upper intertidal 83 pore 197 upper supratidal 84 Fluvial Facies mosaic 136 deposits 2153, 259 Fairway field 152 systems 2.53 Fan delta 38, 54 Formation Fanglomerate 259 Bahloul 454 Faults Capitan 89 contemporaneous Castile 49, 175, 375 growth 418 Clear Fork 40, 135 down-to-the-basin 418 Creede 508 Ferry Lake Anhydrite 137, Diyab 1341 152 Eagle Mills 417 depositional model 155 Glorieta 25 Field Green River 128, 382, Altamont-Bluebell 395 500 Dorr 390 Hanifa 134 Fairway 152, 443 Hermosa 490, 516

~~~~Data Center ,09126599985,[email protected], For Educational Uses 5 44 INDEX~~~~

Louann 417 Gilsonite 385, 386 Minnelusa 135, 142, Gold 507 143 Goniolithon 82 Mission Canyon 171 Great Salt Lake 492 Parakeelya 249 chemistry 365 Queen 149 Green algae Red Cave 55 Chlorophyta 352 Red River 138 unicellular 353 Salado 29, 489 Green house effect 358 San Andres 9, 14, 15, Growth 25, 51, 135, 149, displacive 167 154 replacive 167 Seven Rivers 90, 135, Gulf 144 Arabian 54, 70, 75, 87 Smackover 90, 369 California 484 Srnackover/Buckner Elat 110, 355 138 Mexico 151 Sunniland 373 of California 54 Tansill 33, 89 of Elat 54 Todilto 206 of Suez 54 Frasch mining 511, 513 of Tadjoura 44 Fuerteventura 131 Persian 54, 70, 75 Fundy Basin 302 Suez 94 Gas chromatography-mass Gulf Coast diapirs 415 spectrometry Gulf of California as tools 369 Baja 36, 57 M/Z 191 371, 394, 396 cross section 58 M/Z 217 371, 394, 396 vertical succession 55 Gas inclusions 379 Gulf of Elat Gays River 514 Solar Lake 110 Geopressured Gulf of Suez strata 430 cross section 95 zone 430 Gypsarenite 104 Geothermal discs 85 gradient 42 1 Gypsite 104 systems 506 Gypsum Gilgae structures 230, 273 cement 116

~~~~Data Center ,09126599985,[email protected], For Educational Uses INDEX 5 45~~~~

croute de nappe 277 rafts 12, 13, 28 displacive 80, 84 reticulate iridge 17 economic uses 483 rippled 23 enterolithic folding 419 saturation 6 eolian 105, 106 Halite morphology ghost 152, 153 dendritic 19 hemipyramidal 116 excrescent dendritic 19 mush 81, 83 horizontal chevron 19 pore-filling 84 reticulate iridge 19 secondary 105 Halite saturation level 11 7 sedentary 207 Halite-potash oils 388 Gypsum Plain 483 Halobacterium 352 Gypsum precipitation field Halokinesis 401, 414, 415, 116 425, 44 1 Haber-Bosch process 505 buoyancy 422 Halite collision tectonics 415 "immature" 13 diapirism 415 "mature" 13, 33 differential 422 chevron 33, 58 gravity flow 415 crossbedded 23 gravity-spreading 422 crust 28 t h e r m a 1-convect ion 42 2 crystal growth 6 Halokinetic crystallography 2, 6 response ,417 cuboids 11 structures 415, 417 dendritic 16, 17, 245 Halop h i 1ic bacteria deposition 2 cell morphiology 354 diagenetic 33 cubic form 354 efflorescent 56 €3 a 1o t ect o n i c s ephemeral 84 contraction 422 foundered hoppers 15 extension 422 horizontal chevron 17 Heliothermal 110 morphology 8 brine pond 355 oil seep 394 .Hith Anhydrite 133 ooids 23 Hockley dome pagoda 17 contour map 450 pisoids 23, 24 cross section 450 precipitation 4 Holocene 47, 101

~~~~Data Center ,09126599985,[email protected], For Educational Uses 5 46 INDEX~~~~

dolomite 118 Inclusions salina evaporites 110 fluid 51 Hopane 382 gas 379 Hopanes/steranes ratio 355 petroleum 377, 456 Houston diapir province India 456, 457 Ranns of Kutch 37 FI u t t Lagoon Interdune cross section 115 depression 38, 54, 191 description 47, 114 sabkha 56 vertical succession 48 Intertidal facies Hydraulic head 5, 197 definition 83 Hydrodynamic 164 Intrasediment basin 415 precipitates 10 drive 5 Iodine 503, 504 environment 413 Iran head 103 subaerial salt diapirs model 468 423 structures 22 Iron Hydrology banded 519, 520 burial 173 Superior-type 5 19 closed basin 194 Isotope Coorong Lakes 128 carbon 439, 440, 453, evaporite diagenesis 455, 464 169 hydrogen 453, 454 flowchart 172 light carbon 51 1 modern coastal basins oxygen 439, 440, 453, 101 455 sabkha 86, 89, 92 stable 507 schematic 102 strontium 424, 425, syndepositional 169 440, 455 Hydromagnesite 12 1 sulfur 438, 439, 455 Hydroseal 117 Isotopic Hydrostatic equilibrium 90 gradient 427 signatures 369, 414 pressure 266 Jaffura Sand Sea 262, 292 Hypolimnion 355 Jebels 79 Ibis Gypsite 108 Joachim Dolomite

~~~~Data Center ,09126599985,[email protected], For Educational Uses INDEX 547~~~~

vertical succession 55 Cuahilla 239 Jurassic 133, 138, 369, 417, ephemeral 189 486 Eyre 75 Arab cycles 138, Freefight 209 489 Gosiute 397, 501, 502 Hith Anhydrite 133 Great Salt 363, 397, Smackover 138 492 Karst 293 heliothermal 203 in halite 60 MacDonnell 170 pits 27 MacLeod 103, 107, 114 Kenya Magadi 201, 498, 499 Lake Magadi 201 Maneotis 96 Kerogen 350, 369 Marion 107 type 1 371, 387 Muluk 363 type 2 371 New 48 Kieserite 489 Owens 495 Ku pf e rsc hi ef er perennial 189 copper deposits 515 pluvial 191, 194 Lacustrine carbonate oils Searles 494 392 Solar 110, 355, 358 Lacustrine settings stratified 365 biomass 361 Superior 519 deposits 203, 382 Tyrrell 197 morphology 361 Van 361, 362 biolimiting factor 361 Lake MacLeod Lagoon cross section 46, 109 Bardawil 99 seawater seepage 5 Hutt 36, 47, 114 vertical succession 48 Khor a1 Bazam 80, 85 Lake Magadi Leeman 36, 114 cross section 499 Pekelmeer 130, 131 Lakes Lake eutrophic 383 Abert 202 Type 1 124 Bonneville 363, 492 Type 2 124 Chad 196 Langbeinite 487 Chaplin 502 Law of Uniformitarianism Coorong 118 132

~~~~Data Center ,09126599985,[email protected], For Educational Uses 548 INDEX~~~~

Lead-zinc-silver Coolan Carbonate 108 stratiform 5 18 Esterhazy 488 Lead-zinc-silver-barite Patience Lake 488 sed i men t ary ex ha la t ive Wilkins Peak 294, 502 deposits 517 Meromictic brine 110 Leonite 489 Mesozoic 4 16 Liquid inclusions Metalliferous deposits 446 hydrocarbon 377, 456 cap rock origin 455 Lithium Meteoric water 3, 28, 35, 57, economic uses SO3 198, 507 Great Salt Lake 504 Methane 360 Searles Lake 504 Mexico Lithophyllum 82 Bahia Adair 30 Lit hot h am n ion 82 Baja California 36, 57 Louann salt 427, 432, 486 Carman Island 36 Lower supratidal Gulf of California 37 definition 83 Laguna Mormona 36 MacLeod Ojo de Liebre 36 evaporite basin 46, 107 Salina Grande 36, 42 graben 46 Salina Omotepec 12, Magadiite 500, 501, 520 14, 16, 26, 57 Maghreb Region Michigan Basin diapirism 419 A-1 Carbonate 37 map of Pb-Zn deposits A-2 Carbonate 390 419 bromine 504 Marginal marine Brown Niagaran 372 sabkha 52, 54 Micro-inclusions salina 37, 38 gas 379 Marine carbonate oils 388 Microbial mats 358 Marine false cap rock Middle Devonian 489 diapirs 433 Middle supratidal Marine lagoonal stage 116 definition 84 Maritimes Basin Miliolite 87 geological setting 485 Mineralization timing 462 Mass chromotogram 371, 394, 396 Mining Member Frasch 511, 513 Belle PIaine 488, 489 Haber-Bosch process 505

~~~~Data Center ,09126599985,[email protected], For Educational Uses INDEX 5 49~~~~

room-and-pillar 486 Bonaire 12, 22, 36, 130 solution 486 Goto Meer 169 Miocene 47, 501 Pekelmeer 130, 131 Mirabilite 209, 365, 503 New Lake Mississippi Valley Type vertical succession 48 (MVT) 414, 447, Newark Basin 505, 513, 525 cross section 301 Mississippian 486 description 299 Mobile petroleum systems Nickel-copper-platinoid 397 sulfide ore 523 Models Nile Delta Coast 96 barred-basin 70 Nitrogen 35 1, 487, 504 brine reflux 129, 130, 131 biolimiting factor 361 hydrodynamic 468 economic uses 503 sabkha 99 North Louisiana Basin salina dolomite 118 Belle Island 434 sea-marginal sabkha 87 Boling 434, 437, 445, sea-marginal salina 100, 102 511 Trilogy 79 Hockley 437, 449 Morocco Long Point 513 Argana Basin 260 Moss Bluff 437 Moss Bluff dome Palangana 443 cross section 445 Winnfield 432, 434, 485 Mt. Isa Northern Sinai 98 cross section 518 Bardawil lagoon 99 Mud cracks 121 Northwest Shelf 130 Mudflat Odd carbon preference 366 cross section 58 Odd-carbon n-alkanes 384 dry 39 Oil shale 382 sabkha 138 Mahogany Ledge 385 saline 40, 220 Oil volumes Muriate of potash 487, 493 through time 367 MVT 414, 447, 505, 513, 525 Olorongo Beds 205 n-alkane 366, 369, 370 Oolites Nahcolite 383 halite 23 Neogene 167 grainstone 98 Netherland Antilles Open-marine embayment

~~~~Data Center ,09126599985,[email protected], For Educational Uses 550 INDEX~~~~

stage 114 Persian Gulf Orbitulina 153 vertical succession 55 Ordovician 20, 59, 138 Peru Ostrocod 105 Bocana de Virila 36, Owens Valley 496 169 Padre Island 56 Petroleum envelope 367 Paleocene 501 Phanerozoic 366, 483 Paleomagnetic 427 Phosphorus 351, 359, 487 Paleozoic 3, 504 Photic zone 350 Pangea 46 Photosynthesis Paradox Basin 415, 490, 516 anoxygenic 351 Pekelmeer Lagoon 130, 131 oxygenic 351 Peneroplid 82 Phreatic zone 223, 285 Penesaline 351, 358, 402 Phytane 366 Peninsula Phyt o 1-c h lor ophy 11 366 A1 Zibaya 85 Phytoplankton 362 Sinai 36, 112, 355 Piceance Basin Penobsquis potash deposit cross section 299 cross section 492 Pine Point 513, 524 Perennial Pine Point Zinc District 464 lake 292 Pisoids 23 salina 35 Pisolite 108, 270 saline lake 41, 202, Plankton 203, 286 marine 352, 361 stream 258, 286, 292 organic matter (OM) stream floodplain 202, 360 258, 286, 292 Platform evaporites Peritidal zone depositional settings schematic 91 135 Permeability 6 model 134 Permian Platygyra 82 25, 27, 32, 40, 156, Playa 74, 189, 508 478 Playa-lake 189 Ice Age 156 Pleistocene 44, 47, 494 Texas 21 Pliocene 484 Permian Basin Polygons 29 cross section 50 compressional 8 1

~~~~Data Center ,09126599985,[email protected], For Educational Uses INDEX 55 1~~~~

Polyhalite 489 Peninsula 75 “Popcorn”surf ace 22 Sabkha Faishakh Pore fluids 70, 92 37 Porites 82 Qattara Depression 292 Porosity Quaternary 33 diagram 34 coral reefs 45 ore-hosting 477 Radioactive nuclear waste potash 413, 478, 486 413 economic uses 485 Radioactive waste repository in North America 482 155 kainite 487 Radiocarbon dating 98 kieserite 487 Ras Muhammed Pool langbeinite 487 description 112 muriate of 487 cross section 113 sylvite 487 Red Cave Formation Potassium 482 vertical succession 55 Prairie Evaporite 488 Red River Dolomite 141 Precambrian 95 Reef Precipitation fringing 110 intrasedimentary 17 patch 82 Pressure Replacive growth 167 atmospheric 197 Rock-Eva1 pyrolysis 382 lithostatic 429 Room-and-pillar mining 486 water vapor 6 Rotliegendes 303 Pressure ridges Sabkha 29 Abu Dhabi 75 Pristane 366 A1 Zibaya 80 Pristane/phytane 355, 370, Arabian Gulf 77 382, 384 barrier island-lagoon Pseudomorphs 76, 81, 93 barite 71 classification 75 chert after halite 16 coastal 72, 88, 97, 98 Puffy mat 80 continental interdunal Pyramid a 1 93, 94 hopper 11, 12 definition 74 plates 11 deflation 89 Qatar dolomite 91

~~~~Data Center ,09126599985,[email protected], For Educational Uses 552 INDEX~~~~

dolomitization 90 saline 351 estuarine 76 supersaline 351 Faishakh 37 penesaline 35 1 fan-delta 76, 93, 94 vitasaline 35 1 Cavish 358 Salt inland 189 diapirs 367, 413 interdunal 75, 94 economic uses 485 mainland beach-dune nappes 421 93 Salt dome physiography 79 Belle Island 434 progradation 78 Boling 434, 437, 445, siliciclastic 98 51 1 stacked 88 cavern storage 413 tidal-beach-dune 76 Hockley 437, 449 tidal estuarine 93, 94 hydrodynamic mode i Sabkha model 468 reservoir occurrence 99 Long Point 5 13 Salado Formation mineral resources 441 vertical succession 53 Moss Bluff 437 Salar 189 Palangana 443 Salina 118, 165, 189 Winnfield 432, 434, 485 definition 100 Salt dome native sulfur Salina dolomite model 118 deposits 443 Saline giant 352 Salt flat 74 Saline lake stage 215 Salt pan 189 Saline mudflat 220 Salt saucers 31 Saline pan Saltern description 213, ancient 150 282,288, 293 definition 136 schematic 214 marine 353 Saline soil Salton Trough 484 classification 271, 272 San Andres Formation description 271 cross section 158 vertical profile 276 Northwest Shelf 154 Saline Valley 201 PI 137 Salinity ranges 351 P2 137 hypersaline 35 1 vertical succession 53

~~~~Data Center ,09126599985,[email protected], For Educational Uses INDEX 553~~~~

San Francisco Bay Sediment loading Alviso Salterns 353 diapirism 423, 424 San Simon Valley Sedimentary exhalative 447 zeolite deposits 5 10 Sedimentary exhalative Sand sea 94, 191 deposits 414, 517 Sand sheet description 261 Selenite 105, 117 Saskatchewan 487 Seven Heart Gap 523 Saudi Arabia Seven Rivers Formation Jaffura Sand Sea 263 Yates field 144 Schuler-Hosston delta 427 Shale SDNS 443 black 369, 370, 399, 401 Sea Shallowing upward sequence Australasian 109 105 Black 103 Shamal 78, 89 Dead 20, 22, 70, 191, Shark Bay 107 349, 365, 504 Shoreline description 235 Mediterranean 7 1, 103, Shortite 248, 500 133 Sicily North 415 Calcare de Base 137 Red 44, 103, 507 Miocene evaporites Salton 507 101 Sea-margin sabkha model Sigsbee Scarp 421 definition 87 Silurian 23, 372, 478 Sea-marginal basin Silver 449, 463, 507 hydrology 102 Sinai coast 355 Sea-marginal salina model Sinai Peninsula 100 Northern Sinai 98, 99 Seal Ras Muhammad 36, evaporitic 69 112 stratigraphic 1 Southern Sinai 112 Searles Lake Smackover/Buckner Formations borates 494 138 bromine 503 Soda ash 497 stratigraphic column Sodium carbonate 478, 486 497 economic uses 497 Sedex 447 Sodium sulfate 478, 486, 502 deposits 414 economic uses 497

~~~~Data Center ,09126599985,[email protected], For Educational Uses 554 INDEX~~~~

Soils Steppes 191 caliche 241, 275 Strandline classification 271, 272 definition 135 description 27 1 Stromatolite 12 1 pedogenic 132, 274 domes 104 saline 189 Stuart City Trend 418 vertical profile 276 Stylophora 82 vertisols 274 Subaqueous Solar evaporation 486 bottom precipitates 10 Solar Lake 355, 358 cumulates 10 cross section 111 Subtidal facies Solution mining 486 definition 79 Solvay process 485 Sulfate Source rock 70 bedded 118 South Australia displacive 85, 94 Coorong Lakes 118 magnesium 3 salina 103, 106 reduction 92 Southern Sinai replacive 85 Ras Muhammed Pool Sulfate reduction 112 biochemical 460 Spain thermochemical 460, Tajo Basin 200, 298, 464 300 Sulfide Sphalerite stratiform 437, 462, 5 17 colloform 435 Sulfur Spindletop bioepigenetic 511, 513 salt dome 441 biogenic 464 Spring conceptual model 5 12 artesian 197, 269 deposits 443 cold saline 267 elemental 436, 51 1 deposits 267, 284 native 413, 444, 446, description 266, 293 45 8 hot 500, 501 reduced 511 hot acid 267 sedimentary 399 hydrothermal 201 source of 458 thermal 508 Supersaline 35 1 Stacked platform cycles 133 Supratidal facies

~~~~Data Center ,09126599985,[email protected], For Educational Uses INDEX 555~~~~

definition 83 364 Sylvinite 482, 488 Total organic carbon 366 Sylvite 176 Travertine crusts 256 secondary 167 Triassic 31, 416 Syndepositional processes Trilogy desiccation 24 definition 79 dissolution 24 Trona 383, 385, 497, 499 thermal contraction 24 Trucial Coast 89 Synsedimentary processes Tufa 508 dissolution pits 29 algal 242 Tajo Basin calcareous 268 cross section 300 pinnacles 267 description 298 stromatolitic 243 Telegraph salts Tunisia vertical succession 53 Bou Grine 420, 447, Tepee 57 45 2 diagram 104 diapirs 419 extrusion 125 Fedj el Adoum 420, groundwater indicators 45 2 108 Sabkha el Melah 36 structure 82, 106, 274 Sabkha el Melah de Tertiary 164 Zarzis 171 Texada halite 47, 108 Turkey Thailand 376 borate production 493, halite 382 496 Khorat Basin 383 Lake Hazar 244 sylvite 379 Lake Van 361 Thermal Tuz Golu 279, 361 contraction 31 Turtle-structure 442 expansion 31 Type 1 lake 120 Thermogenic 415, 433 Type 2 lake 120 Thermonatrite 202 United Arab Emirates Tibetan Plateau 284 Abu Dhabi 37 Tidal flats 94 A1 Zibaya peninsula 80 Todi 1to Format ion sabkhas 93 schematic 206 United States Total dissolved solids (TDS) borate production 493

~~~~Data Center ,09126599985,[email protected], For Educational Uses 556 INDEX~~~~

Upper Devonian 138 198, 507 Canada 138 mixed 198 Upper supratidal pore 70 definition 84 volcanogenic 198 Uranium 443 Wave base 350 USSR Wave-formed deposits 240 Lake Balkhash 203 Western Australia Lake Inder 218 Hutt Lagoon 36, 47, Lake Issyk Kul 203 114 Solikamsk Basin 379 Lake MacLeod 107 steppes 191 Leeman Lagoon 114 Vadose Wilkins Peak Member 294 cement 116 cross section 297 zone 8, 27, 28, 86, 224, Williston Basin 225, 285 contour map 139 Venezuela Windsor Group 477, 483 Los Roques 169 barite 521 Viscaria copper deposit 523 Winnfield dome 485 Vitasaline 351, 353, 368, diapirism 427 369, 402 Wireline log 137, 143, 147 Vitrinite reflectance 372, Yates Field 429 cross section 146 definition 398 Zambia-Zaire gradient 400 copperbelt 5 15 index 399 Zechstein Basin 303 suppression 398 Zeolite 202, 509, 221 Volcanic glass 202 aluminosilicate 5 10 Wadi 54, 110, 164 chabazite 5 11 floodwater 39 clinoptilolite 5 11 Wadi Natrun lakes 362, 363 dehydration 5 10 Walther’s Law 43, 60, 63 economic uses 509 Water erionite 51 1 diagenetic reaction 198 Zibar 264 ephemeral 74 Zinc-Lead-Copper-Silver hydrothermal 3, 198 sedimentary exhalative marine, 3 deposits 516 meteoric 3, 28, 35, 57, Zooplankton 70

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